Genesis of Precambrian iron and manganese deposits

Genesis i \ of Precambrian iron

' and manganese deposits

Proceedings of the Kiev Symposium, 20-25 August 1970

Résumés en français

Genèse des formations précambriennes de fer et de manganèse

Actes du colloque de Kiev, 20-25 août 1970

Unesco Paris 1973

Earth sciences Sciences de la terre 9

In this series / Dans cette collection:

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The seismicity of the earth, 1953-1965, by J. P. Rothé / La séismicité du globe, 1953-1965, par J. Rothé. Gondwana stratigraphy. IUGS Symposium, Buenos Aires, 1-15 October 1967 / La estratigrafía del Gondwana. Coloquio de la UICG, Buenos Aires, 1-15 de octubre de 1967. Mineral map o$ Africa. Explanatory note / Carte minérale de l’Afrique. Notice explicative. 1/10 O00 000. Carte tectonique internationale de l’Afrique. Notice explicative / International tectonic map of Africa. Explanatory note. 1/5 O00 000. Notes on geomagnetic observatory and survey practice, by K. A. Wienert. Méthodes d‘observation et de prospection géomagnétiques, par K. A. Wienert. Tectonique de l’Afrique / Tectonics of Africa. Geology of saline deposits. Proceedings of the Hanover Symposium, 15-21 May 1968 / Géologie des dépôts salins. Actes du Colloque de Hanovre, 15-21 mai 1968. The surveillance and prediction of volcanic activity. A review of methods and techniques. Genesis of Precambrian iron and manganese deposits. Proceedings of the Kiev Symposium, 20-25 August 1970 I Genèse des formations précambriennes de fer et de manganèse. Actes du Colloque de Kiev, 20-25 août 1970. Carte géologique internationale de l’Europe et des régions riveraines de la Méditerranée. Notice explicative / International geological map of Europe and the Mediterranean region. Explanatory note. 1/5 O00 O00 (Édition multilingue: français, anglais, allemand, espagnol, italien, russe / Multilingual edition: French, English, German, Spanish, Italian, Russian). Geological map of Asia and the Far East. 1/5 O00 000. Explanatory note. Second edition. Carte géologique de l’Asie et de l’Extrême-Orient. 1/5 O00 000. Notice explicative. Deuxième édition. Geothermal energy. Review of research. Carte tectonique de l’Europe et des régions avoisinantes. 1/2 500 000. Notice explicative / Tectonic map of Europe and adjacent areas. 112 500 000. Explanatory note. (A paraître / To be published.)

Published by the United Nations Educational, Scientific and Cultural Organization, 7 Place de Fontenoy, 75700 Paris Printed by Presses Universitaires de France, Vendôme

Publié par l’organisation des Nations Unies pour l’éducation, la science et la culture, 7, place de Fontenoy, 75700 Paris Imprimerie des Presses Universitaires de France, Vendôme

ISBN 92-3-001107-X (Paper / Broché) ISBN 92-3-001108-8 (Cloth /Relié) L.C. NO 73-79858

The designations employed and the presentation of the material in this publication do not imply the expression of any opinion whatsoever on the part of the Unesco Secretariat concerning the legal status of any country or territory, or of its authorities, or concerning the delimitations of the frontiers of any country or territory.

Les désignations employées et la présentation adoptée ici ne sauraient être interprétées comme exprimant une prise de position du Secrétariat de l’Unesco sur le statut juridique ou le régime d’un pays ou d‘un territoire quelconque, non plus que sur le tracé de ses frontières.

0 Unesco 1973 Printed in France

Foreword Avant-propos

The Precambrian is of very special significance in the evols ution of the Earth’s crust. It represents almo st seven-eighth of the geological history of our planet. During this period of time, lasting approximately 4,000 million years, the basement of continental land masses and the deposits of iron and manganese ore were formed. These latter are of world-wide significance both in quantity and extent. They form part of the natural resources of the geographical environment and their study is important both for developed and developing countries.

The study of Precambrian rocks and ore deposits includes various theoretical and practical aspects-economic, mineralogical, geochemical, tectonic. The research methodology applied to the Precambrian is very specific and fundamentally different from that used for other geological eras. Straightforward time-stratigraphical methods are not applicable here because of the lack of palaeon- tological criteria, destroyed by metamorphism. Successive granitizations form a complex which is very often difficult to bring into conventional order.

In an attempt to throw some light on these complex geological phenomena, Unesco, in collaboration with the International Association of Geochemistry and Cosmo- chemistry of the International Union of Geological Sciences and the Academy of Sciences of the Ukrainian S.S.R., organized a symposium on the geology and genesis of Pre- cambrian iron-manganese formations and ore deposits. At the invitation of the Academy of Sciences, the meeting was held in Kiev from 20 to 25 August 1970. Some sixty special- ists coming from twelve countries met at the Main Con- ference Hall of the Academy of Sciences of the Ukrainian S.S.R. and presented papers at this meeting. The participants were welcomed by R. V. Babiychuk, Minister of Culture of the Ukrainian S.S.R. and Chairman of the Ukrainian National Commission for Unesco, Opening ad- dresses were also given by Academician N. P. Semenenko, Chairman of the symposium, and D r K. Lange of the Natural Resources Research Division of Unesco.

In order to provide a systematic consideration of the problems, the programme was divided into four sections:

L’ère précambrienne a une importance toute particulière dans l’évolution de l’écorce terrestre. Elle couvre les sept huitièmes de l’histoire géologique de notre planète. Pen- dant cette période, qui a duré approximativement quatre milliards d’années, se sont formés le socle des masses conti- nentales et les gisements de fer et de manganèse. Ces gisements précambriens présentent un intérêt mondial à la fois sur le plan de la quantité et sur celui de l’étendue. Ils font partie des ressources naturelles du milieu géographique et leur étude est utile tant aux pays développés qu’à ceux qui sont en voie de développement.

L’étude des roches et gisements de minerais précam- briens revêt divers aspects théoriques et pratiques : écono- miques, minéralogiques, géochimiques, tectoniques. Les méthodes de recherches appliquées, lorsqu’il s’agit du Précambrien, sont très spécifiques et diffèrent fondamen- talement de celles qui sont utilisées pour d‘autres ères géologiques. Les méthodes stratigraphiques de datation ne sont pas applicables ici, faute de critères d‘ordre paléonto- logique, dont l’absence est due au métamorphisme. Les granitisations successives ont donné naissance à un ensemble complexe qu’il est très souvent difficile de classer dans l’ordre conventionnel.

Afin d‘essayer d‘éclairer quelque peu ces phénomènes géologiques complexes, l’Unesco, en collaboration avec l’Association internationale de géochimie et de cosmo- chimie de l’Union internationale des sciences géologiques et l’Académie des sciences de la République socialiste soviétique d’Ukraine, a organisé un colloque sur la géologie et la genèse des formations précambriennes de fer et de manganèse. Sur l’invitation de l’Académie des sciences, la réunion s’est tenue à Kiev du 20 au 25 août 1970. Une soixantaine de spécialistes venus de douze pays se sont réunis dans la grande salle des conférences de l’Académie des sciences de la République socialiste soviétique d‘Ukraine et ont présenté des communications. Les participants ont été accueillis par M. R. V. Babiychuk, ministre de la culture de la RSS d‘Ukraine et par le président de la Com- mission nationale ukrainienne pour l’Unesco ; des discours d‘ouverture ont été prononcés par M . N. Semenenko,

I. Genesis and types of iron-silicate and ferruginous cherty formations, their position in geosynclinal sedimentary or volcanic sequences and the relation between these and analogous manganese-bearing formations.

II. Absolute age dating of iron-silicate and ferruginous formations and their position in the Precambrian stratigraphic sequence. Analogous formations from the Phanerozoic.

III. Differing degrees of metamorphism, the mineral facies and the petrographic nomenclature of ferruginous rocks such as ferruginous quartzites, taconites, jaspilites, itabirites.

IV. Genesis of high-grade secondary iron and manganese ores from iron-silicate and ferruginous formations and ores, metasomatic processes and processes of oxidation in them.

An exhibition of Precambrian/manganese rocks was arranged: consisting of samples from the U.S.S.R., in particular from the Ukrainian S.S.R., as well as samples brought by foreign participants.

Immediately following the meeting, from 25 to 30 August, a field trip was organized to the well-known Krivoyrog deposits. The Ukraine occupies a leading position in industrial mining and exploration of Precambrian iron formations, and this visit enabled participants in the symposium to make comparisons and correlations with rocks of Precambrian iron formations from elsewhere.

The symposium was the first international gathering toprovide an opportunity for a wide exchange of results obtained through studies of rather intricate problems concerning the nature and specific features of the unique iron- bearing metamorphic Precambrian strata of the Earth.

As a result of a broad discussion of the presented papers, it was recommended that further studies be made on basic regularities of occurrence, distribution and genesis of Precambrian iron-manganese ore formations, with special attention to modern geological, mineralogical, geochemical, and geophysical methods and research techniques.

The symposium also drew attention to the need for the determination and detailed investigation of iron- manganese deposits, including studies of interrelations between chert-iron-manganese deposits, including studies of interrelations between chert-iron-manganese and volcanogenic formations.

It was considered that one of the first tasks to be undertaken should be the systematization and classification of the rocks of chert-iron-manganese formations, the correlation of nomenclature of these rocks in different countries, the elaboration of a unified system of nomenclature for iron rocks in different regions of the world, and the study of analogues of these rocks in conditions of different degrees of metamorphism.

A second important task was also recommended: intensification of investigations on the problem of formation of secondary ores, study of characteristic features of these ores in zones of oxidation and supergene alterations along with formation of iron-rich ores related to hypogene processes.

académicien, président du colloque, et par M. K. Lange, de l'Unesco (Division des recherches relatives aux ressources naturelles).

Afin d'assurer l'examen systématique des questions, le programme a été divisé en quatre sections. I. La genèse et les types de formation de silicate de fer et

de chert ferrugineux, leur position dans les séquences sédimentaires ou volcaniques géosynclinales et les relations entre ces dernières et les formations manga- nésifères analogues.

Il. La datation absolue des formations de fer et de silicate de fer et leur position dans la série stratigraphique précambrienne. Les formations analogues phanéro- zoïques .

III. Différents degrés de métamorphisme, faciès des miné- raux et nomenclature pétrographique des roches ferrugineuses telles que les quartzites, taconites, jaspilites et itabirites ferrugineux.

IV. La genèse des minerais de fer et de manganèse secondaires à haute teneur, à partir des formations de minerais de fer et de silicate de fer, les processus méta- somatiques et les processus d'oxydation qui s'y rattachent.

Une exposition de roches manganésées précambriennes a été présentée. Elle était composée d'échantillons provenant de l'URSS, en particulier de la RSS d'Ukraine, ainsi que d'échantillons apportés par des participants étrangers.

Immédiatement après le colloque une visite, qui a duré du 25 au 30 août, a été organisée aux célèbres gisements de Krivoyrog. L'Ukraine occupe une place prépondérante dans l'exploitation minière industrielle et l'exploitation des formations de fer précambriennes et cette visite a permis aux participants de faire des comparaisons et d'établir des corrélations entre les roches des formations ferrugineuses précambriennes.

Ce colloque a été la première rencontre internationale qui ait permis un large échange de résultats d'études consacrées à des questions relativement complexes concer- nant la nature et les caractéristiques des remarquables couches précambriennes métamorphiques ferrugineuses de l'écorce terrestre.

A la suite d'une ample discussion des communications présentées, les participants ont estimé qu'il y avait lieu de procéder à d'autres études sur les constantes fondamentales de la présence, de la répartition et de la genèse des formations précambriennes des minerais de fer et de manganèse, en se préoccupant particulièrement des méthodes et des techniques modernes de recherche géologique, minéra- logique, géochimique et géophysique.

Ils ont en outre souligné la nécessité de délimiter et d'étudier en détail les gisements de fer et de manganèse en recherchant notamment les relations entre les formations de chert-fer-manganèse et les formations volcanogéniques.

L'une des premières tâches devrait être, a-t-on estimé, de systématiser et de classer les roches des formations de chert-fer-manganèse, d'établir la concordance des nomen- clatures de ces roches en vigueur dans différents pays, d'élaborer une nomenclature unifiée des roches ferrugi-

It was agreed that publication of the Proceedings of the symposium would be a valuable contribution to geological and geochemical sciences, and while the Academy of Sciences of the Ukrainian S.S.R. has undertaken to provide such a publication in the Russian language, Unesco was asked to ensure publication in English [with summaries in French).

The undoubted success of this symposium was assured on the one hand by the preparatory work undertaken by the International Association of Geochemistry and Cosmo- chemistry, and in particular its President, Professor E. Ingerson, and on the other hand by the excellent organization of the meeting in Kiev by the Academy of Sci- ences of the Ukrainian S.S.R. Special thanks are due to the Chairman of the Organizing Committee, Professor N. P. Semenenko.

The papers presented at the symposium are reproduced in this ninth volume of the Earth Sciences series. The selection of material, the points of view, and the opinions presented are those of the authors and are not necessarily endorsed by Unesco.

neuses des différentes régions du monde et d'étudier les roches analogues à différents degrés de métamorphisme.

Une autre tâche importante a également été recom- mandée : l'intensification des recherches sur la formation des minerais secondaires, l'étude des traits caractéristiques de ces minerais dans les zones d'oxydation et d'altération supergene ainsi que l'étude de la formation des minerais riches en fer liée aux processus internes.

Les participants ayant estimé que la publication des Actes du colloque constituerait une aide précieuse pour les sciences géologiques et géochimiques, l'Académie des sciences de la RSS d'Ukraine s'est chargée d'assurer cette publication en langue russe, et l'Unesco a été chargée d'en assurer la publication en anglais (avec résumés en français).

Le succes incontestable de ce colloque est attribuable, d'une part, au travail préparatoire accompli par 1'Asso- ciation internationale de géochimie et de cosinochimie, et en particulier par son président, le professeur E. Ingerson, et, d'autre part, à la façon remarquable dont l'Académie des sciences de la RSS d'Ukraine a organisé la réunion à Kiev. Le président du comité d'organisation, le professeur N. P. Semenenko, doit être tout particulièrement remercié.

Le présent ouvrage, qui fait partie de la collection (( Sciences de la terre », reproduit les communications présentées au colloque. Les opinions qui y sont exprimées n'engagent évidemment que leurs auteurs.

Contents Table des matières

Genesis and types of iron-silicate and ferruginous cherty formations, their position in geosynclinal sedimentary or volcanic sequences and the relation between these and analogous manganese bearing formations / Les types de formations de silicate de fer et de chert ferrugineux; leur genèse, leur position dans les séquences sédimentaires ou volcaniques géosynclinales et les relations entre ces dernières et les formations manganésifères analogues

The depositional environment of principal types of Precambrian iron-formations Milieux dans lesquels se sont déposés les principaux types de formations précambriennes de fer [Résumé]

Archaean volcanogenic iron-formation of the Canadian shield La formation de fer volcanogénique archéenne du bouclier canadien [Résumé]

The facial nature of the Krivoyrog iron-formation Les faciès des formations ferrugineuses du Krivoyrog [Résumé]

Jacobsites from the Urandi manganese district, Bahia (Brazil) Jacobsites du district de manganèse d’ Urandi, Bahia (Brésil) [Résumé]

Time-distribu tion and type-distribution of Precambrian iron-formations in Australia Répartition de l’âge et du type des formationsprécambriennes de fer en Austr.alie [Résumé]

The origins of the jaspilitic iron ores of Australia Les origines des minerais de fer jaspilitique d’Australie [Résumé]

Occurrence and origin of the iron ores of India Manifestations et origine des minerais de fer de l’Inde [Résumé]

Precambrian iron ores of sedimentary origin in Sweden Minerais de fer précambriens 21 caractères sédimentaires, en Suède [Résumé]

The ferruginous-siliceous formations of the eastern part of the Baltic shield Les formations de fer siliceux dam la partie orientale du bouclier baltique [Résumé]

Precambrian ferruginous-siliceous formations associated with the Kursk Magnetic Anomaly Les formations de fer siliceux du Précambrien dans la région de l’anomalie magnétique de Koursk [Résumé]

Structural-tectonic environments of iron-ore process in the Baltic shield Precambrian Environnement tectonique et structural des processus de formation de minerai de fer dans le Précambrien du bouclier baltique [Résumé]

G. A. Gross

A. M . Goodwin

A, I. Tugarinov, I. A. Bergman and L. K. Gavrilova

E. Ribeiro Filho

A. F. Trendall

R. T. Brandt

M . S. Krishnan

R. Frietsch

V. M. Chernov

N. A. Plaksenko, I. K . Koval and I. N. Shchogolev

P. M . Goryainov

15 20

23 33

35 39

41 47

49 55

59 66

69 75

77 82

85 86

89 94

95

98

Geology of the Precambrian cherty-iron formations of the Belgorod iron-ore region Géologie des formations précambriennes de fer siliceux dans le gisement de Belgorod [Résumé]

Iron-formation and associated manganese in Brazil Formation de fer et de manganèse en association, au Brésil [Résumé]

The Precambrian iron and manganese deposits of the Anti-Atlas Gisements de minerai de jer et de mangatièse dans le Précambrien de l’Anti-Atlas [Résumé]

Tectonic control of sedimentation and trace-element distribution in iron ores of central Minas Gerais (Brazil) Le contrôle tectonique de la sédimentation et la répartition des éléments-traces clans les minerais de fer de la partie centrale de l’&tut de Minas Gerais, au Brésil [Résumé]

Yu. S. Zaitsev

J. Van N. Dorr II

G. Choubert and A. Faure-Muret

A. L. M. Barbosa and J. H. Grossi Sad

Absolute age dating of iron-silicate and ferruginous formations and their position in the Precambrian stratigraphic sequence. Analogous formations from the Phanerozoic / L a datation absolue des formations de fer et de silicate de fer et leur position dans la série stratigraphique précambrienne. Les formations phanérozoïques analogues

The iron-chert formations of the Ukrainian shield Géologie et genèse des formations de fer siliceux du bouclier cristallin d’ Ukraine [Résumé]

Occurrences of manganese in the Guianas (South America) and their relation with fundamental structures Les indices de manganèse dans les Guyanes (Amérique du Sud) et leurs relations avec les structures fondamentales [Résumé]

Precambrian ferruginous-siliceous formations of Kazakhstan Les formations de fer siliceux dans le Précambrien du Kazakhstan [Résumé]

Geology and genesis of the Devonian banded iron-formation in Altai, western Siberia and eastern Kazakhstan Géologie et genèse de la formation dévonienne de fer rubaiié dans I’Altai, la Sibérie occidentale et le Kazakhstan oriental Désumé]

Genesis of high-grade iron ores of Krivoyrog type Genèse des minerais de fer à haute teneur de Krivoyrog [Résumé]

Effusive iron-silica formations and iron deposits of the Maly Khingan Les formations de fer siliceux eflusif et les gisements de fer du Maly Khingan [Résumé]

Effusive jasper iron-formation and iron ores of the Uda area La formation du minerai de fer à jaspe effusifet les minerais de fer de la région d’Ouda [Résumé]

N. I?. Semenenko

B. Choubert

I. P. Novokhatsky

A. S. Kalugin

Y. N. Belevtsev

E. V. Egorov and M. W. Timofeieva

E. L. Shkolnik

Differing degrees of metamorphism, the mineral facies and the petrographic nomenclature of ferruginous rocks such as ferruginous quartzites, taconites, jaspilites, itabirites / Différents degrés de métamorphisme, faciès des minéraux et nomenclature pétrographique des roches ferrugineuses telles que quartzites ferrugineuses, taconites, jaspilites et itabirites

Mesabi, Gunflint and Cuyuna Ranges, Minnesota (United States of America) Les chaînes de Mesabi, Gunflint et Cuyuna dans le Minnesota, aux ÉLats-Unis d’Amérique [Résumé]

Physico-chemical conditions of the metamorphism of cherty-iron rocks Les conditions physico-chimiques du métamorphisme des formations de fer siliceux [Résumé]

The Serra do Navio manganese deposit (Brazil) Le gisement de manganèse de Serra do Navio, au Brésil [Résumé]

G. B. Morey

Y. P. Melnik and R. I. Siroshtan

W. Scarpelli

1 o1 1 Ó3

105 112

115 123

125

131

135 141

143 150

153 156

159

164

167 177

181 184

187 189

193 206

209 215

217 227

Genetic studies on the Precambrian manganese formations of India with particular reference to the effects of metamorphism 229 Étude génétique des formations de manganèse précambrien en Inde avec références particulières aux efsets du métamorphisme [Résumé] 239

Precambrian ferruginous formations of the Aldan shield 243 Formations ferrifères du Précambrien inférieur du bouclier d’Aldan [Résumé] 246

I. D. Vorona, V. M . Kravchenko, V. A. Pervago and I. M . Frumkin O n the issue of genesis and metamorphism of ferromanganese formations in Kazakhstan 249 Formation et métamorphisme des roches ferrugineuses de diverses époques dans les provinces du Kuzakhstan [Résumé] 253

S. Roy

V. M . Shtsherbak, A. S. Kryukov and Z. T. Tilepov

Genesis of high-grade secondary iron and manganese ores from iron-silicate and ferruginous formations and ores, metasomatic processes and processes of oxidation in them / Genèse des minerais de fer et de manganèse secondaires à haute teneur, à partir des formations de minerais de fer et de silicate de fer; processus métasomatiques et processus d’oxydation qui s’y rattachent

Iron-formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites Formations de fer du groipe de Hamersley, en Australie occidentale : exemples typiques d’évaporites précambriennes en

257

varve [Résumé]

Geology and iron ore deposits of Serra dos Carajás, Pará (Brazil) Géologie et dépôts de minerai de fer de la Serra dos Carajás, Pará, Brésil [Résumé]

Enrichment of banded iron ore, Kedia d’Idjil (Mauritania) Enrichissement des minerais zonés de fer de la Kedia d’ldjil en Mauritanie [Résumé]

Iron ores of the Hamersley Iron Province, Western Australia Les minerais de fer d’Hamersley, en Australie occidentale [Résumé]

Significance of carbon isotope variations in carbonates from the Biwabik Iron Formation, Minnesota Signification des variations des proportions des isotopes du carbone dans les carbonates des gisements de fer de Biwabik, dans le Minnesota [Résumé]

Genesis and supergene evolution of the Precambrian sedimentary manganese deposit at Moanda (Gabon) Genèse et évolution supergène du gisement sédimentaire précambrien de manganèse de Moanda, au Gabon [Résumé]

The Belinga iron ore deposit (Gabon) Les minerais de fer de Bélinga, au Gabon [Résumé]

Itabirite iron ores of the Liberia and Guyana shields Les minerais de fer d’itabirite du Libéria et du bouclier guyanais [Résumé]

Structural control of the localization of rich iron ores of Krivoyrog Détermination structurale de la localisation des minerais de fer à haute teneur de Krivoyrog [Résumé]

Iron deposits of Michigan (United States of America) Gisements de fer du Michigan, aux États-Unis d’Amérique [Résumé]

A. F. Trendall

G . E. Tolbert, J. W. Tremaine, G. C. Melcher and C. B. Gomes

F. G. Percival

W. N. MacLeod

E. C. Perry Jr and F. C. Tan

F. Weber

S. J. Sims

H. Gruss

G. V. Tokhtuev

J. E. Gair

Problems of nomenclature for banded ferruginous-cherty sedimentary rocks and their metamorphic equivalents

268

27 1 279

281 288

291 297

299

304

307 320

323 332

335 3 57

361 3 64

365 374

377

381 List of participants / Liste des participants

Genesis and types of iron-sihcate and ferruginous cherty formations, their position in geosynclinal sedimentary os volcanic sequences and the relation between these and analogous manganese-bearing formations

Les types de formations de silicate de fer et de chert ferrugineux ; leur genèse, position dans les séquences sédimentaires ou volcaniques géosynclinales et les relations entre ces dernières et les formations manganésifères analogues

The depositional environment of principal types of Precambrian iron-formations

G. A. Gross Geological Survey of Canada, Ottawa 4, Ontario (Canada)

Iron-formations composed of thinly bedded chert and iron minerals which contain at least 15 per cent iron are probably the most abundant chemically precipitated sedimentary rocks known. They occur in a wide variety of geological environments and because of the diversity in chemical properties of their elemental constituents are highly sensitive indicators of the depositional environments in which they formed. Much of the geological literature on these rocks has been based on separate iron ranges or formations and interpretations from these specific studies have been applied to the whole group of cherty ferruginous sediments. In- terpretations and extrapolations are frequently made without distinguishing adequately the diversity in depositional, tectonic, chronological and host rock environments iii which the many different lithological varieties of these chemical sediments occur. The purpose of this paper is to distinguish differences between some of the principal geological environments where siliceous iron sediments occur and to recognize variations in the physical and chemical characteristics of banded cherty iron-formations as found in these different environments. It is necessary to distinguish and define the various types of depositional environments of these rocks before concepts and hypotheses pertaining to their origin and genesis can be satisfactorily evaluated and the geological significance of iron-formations fully appreciated.

Of the broad group of iron-rich sediments, only the banded cherty iron-formation sediments are considered in this paper. The oolitic chamosite-siderite-goethite clay-rich rocks commonly referred to as ironstones are recognized as a distinctly separate type of iron sediment, They formed in different environments than the cherty iron sediments and probably have a different origin and source of iron. The separate group of cherty iron sediments which are associated with a wide variety of sedimentary and volcanic rocks indicate pronounced diversity in conditions in their sedimentary environments. The cherty iron-formations are chemically precipitated sediments and the many different sedimentary facies demonstrate the changes in physical and chemical environment during their deposition. The distinct

variations in geological environment and physical and chemical characteristics of the cherty iron-formations are such that it cannot be assumed that all of these rocks have similar sources of iron and silica and similar genetic affinities.

It is highly probable that there are other fundamental factors affecting the origin of these sediments which have still not been recognized. Because there are relatively few examples of cherty iron-formation in rocks of Mesozoic age or younger and apparently no modern examples exist where banded cherty iron sediments are forming today, we have no complete contemporary model or guide to the geological parameters affecting the origin of these special sediments. For these reasons investigations of the depositional environment of cherty iron sediments have to be comprehensive both in scope and in definition of sedimentary features and environment if the mode of origin of these rocks is to be understood. Detailed comparisons of iron ranges throughout the world may provide a composite picture of the complex factors and conditions which con- tribute to the deposition of iron-formations.

It has proved highly instructive to classify or group the cherty iron-formations according to general features and characteristics of their depositional environments and the kinds of sedimentary rocks associated with them. In Canada the name ‘Algoma type’ has been used in recent years to designate cherty iron-formations and their equivalent facies variants that are intimately associated with volcanic rocks and greywacke type sediments in eugeosynclinal belts. The iron-formations associated with quartzites, dolomites and black slates in continental-shelf environments are classified as ‘Lake Superior type’. This broad classification may not be entirely satisfactory for all occurrences of cherty iron- formation, but it serves to distinguish the two main environments in which cherty iron sediments most frequently occur.

The Lake Superior type of iron-formation forms prominent iron ranges of middle to late Precambrian age in nearly all of the shield areas of the world. Most of the geological literature on cherty iron-formations is based on this type of iron sediment and it is the host rock, or protore, for

Unesco, 1973. Genesis of Precambrian iron and niunganese deposits. Proc. Kiev Syrnp., 1970. (Earth sciences, 9.) 15

G. A. Gross

the largest and best known iron ore deposits in cherty iron- formations.

Lake Superior type iron-formations are characteristically thin-banded cherty rocks with iron-rich layers representing various sedimentary facies. Oxide facies are composed of magnetite, hematite or mixtures of these oxide minerals which were deposited mainly as primary iron oxides. Silicate minerals in the silicate facies commonly range from greenalite and minnesotaite to stilpnomelane, cummingtonites and grunerite to hypersthene depending on their rank of metamorphism. Carbonate facies are represented predominantly by siderite associated with magnetite or iron silicates but ankerite and ferrous dolomites are prevalent where carbonate is associated with hematite-rich facies. Sulphide facies of this type of iron-formation usually consist of fine-grained carbon-rich mudstones with interlayered chert or siliceous shale. Characteristic features of the various facies of this type of iron-formation have been described by Gross (1965), James (1954) and others.

Granules and oolites composed of both chert and iron minerals are typical textural features of these sediments and they are practically free of clastic material except in the transitional border zones or in distinct well-defined members within the formation. The alternate or rhythmic banding of iron-rich and iron-poor cherty layers, which normally range in thickness from a few millimetres to 1 metre, is a prominent feature. Individual layers may pinch and swell to give a wavy-banded member or the uniformity of the layering may be disrupted by nodular or stubby lenses of chert and jasper, by rare occurrences of cross- bedding, or by cherty forms resembling in shape and structure ‘Collenia’ or ‘Crystozoan’ growths in limestones formed by algal colonies. Tension, syneresis and desiccation cracks are present in some chert granules and nodules, and styolites are common. Textures and sedimentary features of this type of formation are remarkably alike in detail wherever examined, although certain sedimentary features are more prominent in some formations than in others.

The close associations of this type of formation with quartzite and black carbonaceous shale, and commonly also with conglomerate dolomite, massive chert, chert breccia, and argillite, are recognized throughout the world. Volcanic rocks, either tuffs or flows, are not always directly associated with Superior type iron-formation, but they are nearly always present in some part of the stratigraphic succession. The sequence dolomite, quartzite, red and black ferruginous shale, iron-formation, black shale and argillite, in order from bottom to top, is so common on all continents that some investigators have been led to believe in the past that it is invariable. However, stratigraphic studies have shown that, although there is a persistent association of these sedimentary rocks, the succession may differ in local areas; it does so for example in the Labrador geosyncline. Quartzite and red to black shale lie below the iron- formation and black carbonaceous shale above it, but the presence of other sedimentary rocks and their position in the stratigraphic succession may vary from place to place, even in a single range or sedimentary belt.

Continuous stratigraphic members of Superior type iron-formation commonly extend for hundreds of miles along the margins of ancient continental platforms or geosynclinal basins. The formations may vary in thickness from a few tens of metres to several hundred metres and occasionally up to 1,000 metres, but their persistence is truly remarkable. The rock successions in which the iron- formations occur usually lie unconformably above highly metamorphosed gneisses, granites or amphibolites, and the iron-formations are, as a rule, in the lower part of the succession. In some places they are separated from the basement rocks by only a few metres of quartzite, grit and shale or, as in certain parts of the Gunflint Range, they lie directly on the basement rocks. However, in most areas they occur at least some hundreds of metres above the basement rocks.

The Lake Superior type iron-formations are present in late Precambrian rocks in nearly all parts of the world and possibly in some early Palaeozoic rocks (O’Rourke, 1961). They apparently formed in fairly shallow water on continental shelves or along the margins of continental shelves and miogeosynclinal basins, and consist of sediments derived from the adjacent land mass and also some material from the volcanic belts within the basin. It is still considered uncertain as to whether the iron and silica in this type of iron-formation were derived from the eroding of a land mass or a volcanic source.

This type of siliceous formation is the protore or host rock for the rich hematite-goethite orebodies of the Lake Superior region in the United States, Quebec-Labrador in Canada, north-western Australia, Orissa and Bihar states in India, Krivoyrog and Kursk areas in the U.S.S.R., in Brazil and for many other major iron deposits in the world.

Algoma type iron-formations are present in nearly all of the early Precambrian belts of volcanic and sedimentary rocks in the Canadian shield, in parts of the Australian shield and in belts of similar rock of Palaeozoic and Mesozoic age in many other regions. This type of iron- formation is characteristically thin-banded or laminated with interlayered bands of ferruginous grey or jasper chert and hematite and magnetite. Massive siderite and carbonate beds, iron-silicate mineral facies and iron-sulphide mineral facies are frequently associated in the formation but are less abundant than the oxide facies. In the Michipicoten area of Ontario, massive siderite and pyrite-pyrrhotite beds form part of the formation. Single iron-formation members of this type range from more than a hundred metres to less than 1 metre in thickness and rarely extend more than a few kilometres along strike. A number of these lenticular beds may be linked together or distributed en échelon throughout a belt of volcanic and sedimentary rocks. The Algoma type iron-formations are intimately associated with various volcanic rocks including pillowed andesites, tuffs, pyroclastic rocks, or rhyolitic íìows and with greywacke, grey- green slate, or black carbonaceous slate. Tuff and fine- grained clastic beds or ferruginous cherts are interbedded in the iron-formation and detailed stratigraphic successions

16


show heterogeneous lithological assemblages. These iron- formations have streaked lamination or layering, and oolitic or granular textures are apparently absent, except in rare cases in post-Precambrian rocks. The associated rocks indicate a eugeosynclinal environment for their deposition and a close relationship in time and space to volcanic activity.

The direct association of Algoma type iron-formations with centres of volcanism or volcanic activity is recognized in a number of volcanic belts in the Canadian shield. Rhyolitic and dacitic volcanic rocks are usually thickest and most abundant in the succession of volcanic sedimentary rocks in and around the ancient volcanic centres. In general the Algoma iron-formations overlap the bulk of the acidic volcanic material and are in turn covered by andesitic volcanic rocks and associated greywacke type of sediments. Sulphide and carbonate facies of iron-formation occur at or near the centres of volcanism and the oxide facies are usually distributed farther away, even where they are almost entirely enclosed by clastic sediments. Carbonate and silicate facies occur near the centres of volcanism, but a general zonal relationship from sulphide through carbonate to oxide facies of Algoma type iron-formation is commonly found. This direct relationship between the type of iron- formation facies and the various kinds and distribution of volcanic rocks leaves little doubt about the genetic relationship of these cherty iron sediments and volcanic processes.

Thin beds of graphitic schist or black carbon-rich mudstones are commonly associated with Algoma type iron-formation and occur mainly in parts of the succession where volcanic rocks are more abundant than the greywacke sediments. Much of the fine clastic material in the black mudstone may be derived from tuff and volcanic ash and collected in depressions in the depositional basin. Usually they contain pyrite and pyrrhotite and parts of them have appreciable amounts of lead, zinc and copper. Black mudstones of this type are closely associated with stratiform base metal sulphide deposits and are one of the common host rocks in which the thin-banded and layered sulphide beds occur. The black mudstones may be a facies of the Algoma type iron-formation and occur in the same bed or member as oxide and carbonate facies. They also occur as separate beds or horizons which are closely associated with thicker beds composed of other facies of iron- formation.

Algoma type iron-formations are widely distributed in the volcanic-sedimentary belts in the older parts of the Canadian shield and some of the better known examples of this type of iron-formation occur in the Michipicoten District, near Kirkland Lake, Moose Mountain Mine, Timagami Lake, the Kapico Iron Range, north of Nakina, at Red Lake, Bruce Lake and Lake St Joseph in Ontario. Examples of Ordovician age occur near Bathurst in northern New Brunswick and northern Newfoundland and some of Mesozoic age on Vancouver Island.

Iron-manganese formations of Algoma type are of particular interest but are relatively rare compared with the frequency of occurrence of the iron-rich beds. Iron- manganese formations were deposited under much the same

conditions and in a similar geological environment to those for typical Algoma iron-formation. The manganese content may range from nearly pure cherty manganese sediment to cherty sediments with a low manganese to iron ratio. Examples of this type of cherty sediment are found in the Karazdhal range in the U.S.S.R. and near Woodstock in New Brunswick, Canada. These appear to be formed by volcanic exhalative processes and are classified with Algoma type iron-formation.

Nearly all of the cherty iron-formations can be classified satisfactorily in these two principal environmental types. Many of the iron-formations and their associated rocks are highly metamorphosed and their sedimentary environments can only be interpreted from the relict sedimentary features that are still recognizable. Many other iron-formations are not known in detail and their immediate geological setting or depositional environment has not been studied or reported.

A n interesting iron-formation in an unusual geological setting extends along the Yukon and Mackenzie District border in north-western Canada. The Snake River iron- formation forms a succession of jasper and blue hematite beds more than 150 metres thick which occur near the base of the Rapitan formation; a crudely stratified, poorly sorted conglomerate at least 1,500 metres thick. The Rapitan conglomerate lies between two angular unconformíties. It overlies a thick succession of dolomite, shale, gypsum and shale, shaly carbonate, limestone, and quartzite which may be Lower Cambrian in age but is believed to be Precambrian. The Rapitan conglomerate is overlain by dark shale and silty dolomite of late Cambrian age. The exact age of the Rapitan formation and the enclosed iron- formation is still not known.

The Rapitan formation as a whole is composed of conglomeratic siltstone and shale, siltstone and silty shale with about 10 per cent of its volume made up of rounded to subangular coarse fragments mostly in the 1-5 centimetres size range with isolated boulders up to 5 metres in dimension. The coarse fragments consist of carbonate, basic igneous rocks, sandstone, quartzite and shale in decreasing order of abundance. Much of the conglomeratic siltstone in the lower part of the formation associated with the iron-formation is highly ferruginous and dark red to maroon in colour. Parts of the Rapitan formation some distance from the thicker iron-formation contain a high proportion of coarse fragmental volcanic rocks and considerable tuffaceous material.

The iron-formation has an average iron content of 46 per cent and is composed of interlayered bright red jasper and fine-grained deep blue hematite beds which range in thickness from thin laminae to several centimetres. The jasper and hematite layers are mostly well-segregated, but some hematite beds have conspicuous round nodules of red, grey or buff chert 0.5-1 centimetre in size which may make up 20 per cent of the hematite layer. Granular or oolitic textured beds were not found in the iron-formation. Other common siliceous layers and beds are deep red to maroon in colour and are made up of very fine-grained clastic mud

17

G. A. Gross

in a highly siliceous matrix. There are numerous thin lenticular beds of coarser clastic material distributed throughout the iron-formation. Some of the fine-grained silty material is composed of tuff fragments and coarser fragments are similar in composition to the coarse fragments in the main conglomerate. Two thin but continuous silty sandstone beds, one near the base and one near the top of the iron-formation sequence, have been used as horizon markers for correlation of detailed stratigraphy. Thin laminae and beds of ankeritic and dolomitic carbonate are interlayered with the chert and hematite in some parts of the iron sequence. The iron-formation appears fresh and there is little evidence of metamorphism. Primary sedimentary and diagenetic features are well preserved and much can be determined about the sedimentary environment and nature of these beds.

Differential compaction features, slump and glide structures, intraformational breccias composed of cherty iron-formation fragments, scour and €ill structures, tension and syneresis cracks are all conspicuous throughout the iron-formation. Many of the coarse fragmental beds appear to have been mud flows which spread over beds of partly consolidated iron-formation causing distortion and disturbance of the underlying bedding in the iron-formation. The iron-formation overlying the mud flow is straight, undisturbed, horizontally bedded jasper and hematite. In some places mud flows were observed which had scoured and cut channels in the soft iron-formation 5 metres deep, and tens of metres wide. Large blocks of iron-formation are suspended in the mud flow and iron-formation fragments in the flow are most abundant near the walls of the channel.

The suggestion has been made that some of the large isolated boulders found in the iron-formation, which caused warping and depression of the underlying chert beds, were rafted by ice and dropped in the soft semi-consolidated cherty iron-formation. Most of these lie along thin seams of conglomerate and tuffaceous material and the writer believes that this material is the product of explosive volcanism which took place during the deposition of the iron- formation. No volcanic vents or diatremes have been identified, but the occurrence of tuffaceous layers and volcanic materials in the conglomerate and iron-formation are evidence of volcanic activity during the deposition of these rocks.

The thick lenticular iron-formation described here is exposed over a width of 10 miles (16 kilometres) and extends laterally for more than 30 miles (48 kilometres). It thins towards the east and west, is terminated at the uiicon- formity surface to the north and its extent to the south, where it dips under younger strata, has not been determined. The total dimensions of this iron-formation, either for the thicker lenticular zone or for its complete lateral extent, are not known. Thinner beds of lithologically similar iron- formation, which may be a continuation of this same stratigraphic zone, have been observed in isolated occurrences for more than 200 miles (320 kilometres) to the north-west and also for some considerable distance to the south-east.

The Rapitan formation represents a rapid filling of a deep basin depression with poorly sorted and stratified silty and conglomeratic material. Chemical precipitation of the iron and silica of the cherty iron-formation has taken place at the same time as the inpouring of the silty conglomerate and the two types of sedimentation, clastic and chemical, have been superimposed on one another. The iron- formation is fresh and relatively unmetamorphosed. Primary sedimentary features indicate that alternate chemical precipitation of silica- and hematite-rich layers was interrupted by the influx of mud flows and conglomerate which spread over the partly consolidated chert and hematite, in places scouring channels in the soft iron-formation. The conglomerate and iron-formation are believed to have been deposited in a broad depression or basin on the ocean floor, and slumping and flow of unconsolidated rocks from adjacent fault scarps or basin shelves may have been triggered by movement along bordering faults or by explosive volcanic activity. Some of the fine-grained clastic beds impregnated with hematite appear to be tuff or volcanic ash that settled in soft hematite ooze.

The hematite and silica are believed to have been transported in solution by hot fumarolic waters and precipitated when these solutions were discharged on the sea floor along fault zones (Gross, 1965). The Snake River iron-formation may be the product of exhalative-sedimentary processes and therefore have a very close genetic aíñnity to the main goup of iron-formations classified as Algoma type. The origin of the Snake River iron beds may be closely analogous to the siliceous iron, manganese and base-metal deposits at present being precipitated in the deeps of the Red Sea (James, 1969).

The Snake River iron-formation represents a voluminous influx of chemicalIy precipitated iron and silica into a basin that was being rapidly filled by conglomerate and coarse silt. There is no apparent genetic relationship between the source and manner of derivation of the two types of sediment. In the case of the Algoma type iron-formations, the chemically precipitated iron and chert beds deposited contemporaneously with a great variety of volcanic and sedimentary rock and the specific genetic relationship between the chemical sediment and the various kinds of clastic and volcanic material is subject to conjecture and interpretation. Important empirical relationships of different facies of iron-formation with certain phases of volcanic activity or kinds of volcanic rock and sediments, and the zonal distribution of different iron-formation facies and exhalative deposits around volcanic centres, leave little doubt that deposition of iron-formation and the volcanic rocks are both expressions of a common igneous-volcanic phenomenon.

In the case of the Lake Superior type of iron-formation, very thick successions of chemically precipitated silica and iron sediment have been deposited in sequences of normal and common types of continental shelf sediment. In many of these areas there was contemporaneous volcanic activity and deposition taking place along the outer edge of the shelf or basin. A possible common source for the iron and silica

18


in the iron-formation aiid the quartzites, dolomites and argillaceous sediments has been proposed by postulating deep chemical weathering of a land mass and special erosion and sedimentation conditions to account for the whole assemblage of sedimentary rocks. Geological models based entirely on these concepts of erosion, transportation and deposition of the iron and silica have not provided a satisfactory explanation for the origin of this type of iron- formation. The problem of the source and origin of the iron and silica has not been solved conclusively by appealing to exhalative-sedimentary processes related to volcanic activity in the adjacent volcanic belts. The writer believes, however, that the source of iron and silica most probably lies in the volcanic belt rather than in an eroded land mass. This opinion is based more on comparison of common features and aspects in the environments of Algoma and Lake Superior formations and analogies which may be made between the two types. It is expected that continuing detailed study of the depositional environments of both Algoma and Lake Superior types of iron-formation will provide examples of iron ranges deposited under conditions intermediate between the volcanic eugeosyncline environment of the Algoma type and the stable continental shelf environment of the Lake Superior type. If this proves to be the case, then the two prominent types of environment now recognized can be considered as two depositional models or sedimentary expressions with the iron and silica derived or supplied from a common kind of source and by a common phenomenon.

Recognition of the two principal types of cherty iron- formation and characteristic features of their depositional environment is an important step towards determining the critical or essential geological processes and features that are involved or related to the origin and development of these chemical sediments. Only some of these processes or features are mentioned here in a qualitative way and it is not possible in this short paper to elaborate on their significance or implication with regard to the source of iron and silica and the origin of cherty iron-formations.

There are also many important economic implications related to these typical environments which are being considered in mineral exploration. Distinctive characteristics of the different kinds of iron ore derived from the principal types of iron-formation have been recognized and described in the literature on iron-ore deposits and will not be elaborated here. Recognition of the characteristic features of the types of iron-formation plays an important part in the evaluation of newly discovered or developed iron-ore deposits. Probably one of the most significant factors relating to the type of iron-formation is recognition of the kind of manganese, base-metal or gold deposits that may be associated with it.

Important stratiform base-metal sulphide deposits in the same geological environment as Algoma type iron- formation are recognized as facies variants of sulphide facies or iron-formation and, like the iron-formation, are considered to be exhalative-sedimentary volcanic deposits. There is little doubt about the genetic relationship of these

stratiform base-metal sulphide deposits with sulphide, car- boiiate and oxide facies of iron-formation, and recognition of this fact has fostered new and highly rewarding concepts in mineral exploration in the Canadian shield. The empirical association of gold deposits and Algoma type iron- formation has been recognized for many years. In the past, some have explained this relationship on a structural basis, believing that the brittle cherty iron-formations were a favourable host rock for quartz vein development. Evidence is now being accumulated to show that the carbonate and some of the sulphide facies of Algoma type iron-formation are source beds for gold and probably silver which were later concentrated in veins and stockworks associated with the iron-formation.

It is noted that the composition and physical characteristics of some of the stratiform base-metal sulphide deposits associated with Algoma type iron-formation are very similar and directly comparable with the contemporary layered siliceous sulphide sediment being deposited in the deeps of the Red Sea. It is highly probable that deposits in the Algoma type iron-formation and Red Sea environments are both products of deep-seated magmatic processes centred along major faults or tectonic features in the crust, Fumarolic activities and circulation of water caused by near-surface thermal gradients have probably given rise to the solution and transport of large quantities of silica and metallic ions in both cases. In the Algoma type environment there has been a prominent deposition of volcanic rock contemporaneous with the discharge of these metal bearing solutions and deposition of their salts, while in the Red Sea solutions are being discharged from the deep- seated fault systems without active volcanism.

Referring briefly to the global distribution of cherty iron-formations, it is noted that many of the major Precambrian iron-formations of the world lie close to or parallel to the borders of the continental masses. This is the case for iron-formation near the west coast of the African continent and those in South America along its east coast, and for the distribution of iron-formations in India and Australia. These iron-formations are Precambrian sediments in ancient shield terrain which may have been closely related to, or even parts of, the sanie depositional basins and tectonic belts prior to continental drifting and segmentation of the principal Precambrian land masses. The type of cherty iron-formation, its associated rocks and depositional environment for each of these iron belts, need to be defined and compared in detail to determine whether the iron ranges now on the borders of different continents may at one time have formed parts of the same depositional basins and tectonic belts. This comparison of the type and environment of iron-formation belts is of course dependent on better determination of the age of sedimentation of the iron beds and much inore detailed information on the chronological sequence of events in each of the iron- formation ranges.

The writer believes that many of the Precambrian iron ranges and their depositional basins may be closely related to, if not parts of, the same sedimentary sequences of rocks

19

G. A Gross

which were separated during the segmentation and drifting of the continents. The iron-formations may be closely related to major deep-seated fault and tectonic systems of global dimensions which existed in the Precambrian land mass and have not been recognized because of continental drift. The separation of large volumes of iron and silica and their transportation by fumarolic water or by circulation of water currents caused by thermal gradients along these tectonic zones may be related to deep-seated igneous aiid volcanic processes. If this is the case, we can then appreciate some fundamental reasons and basic causes for finding this large group of cherty sediments in such a diversity of depo-

sitional environments, The fundamental reasons for finding voluminous sequences of silica, iron and other metallic elements on continental shelves, in volcanic-sedimentary rock assemblages in eugeosynclines, or in thick sequences of conglomerate, as in the case of the Snake River iron- formations, will not be found by exclusive studies of the sedimentation in typical iron-formation environments. These answers will most likely be found through study of major tectonic features and the associated deep-seated igneous processes which may have had a common genetic relationship to all of these distinctive sedimentary environments of iron-formation.

Résumé

Milieux dans lesquels se sont déposés les principaux types de formations précambriennes de fer (G. A. Gross)

Les formations de fer, veinées de silex, qui sont réparties très largement dans toutes les régions du bouclier précam- brien se rencontrent dans deux types principaux de milieux ; d'où le nom qui leur a été donné en Amérique du Nord : (( Algoma )) et (( Lac Supérieur )).

Le type (( Algoma 1) est étroitement lié à la fois par sa genèse et par sa localisation aux roches volcaniques. O n pense qu'il a été produit par des processus d'exhalation volcanique dans un milieu eugéosynclinal. I1 consiste dans une grande variété de faciès sédimentaires qui vont de l'oxyde de fer siliceux aux faciès des carbonates, silicates et sulfures. Ce type est largement distribué dans les roches volcaniques archéennes dans tout l'ensemble du bouclier canadien.

Le type (( Lac Supérieur )) s'est déposé sur la plate- forme protérozoïque et dans les environs du plateau continental. Il est associé avec la quartzite, la dolomite et l'argile schisteuse noire, et avec du tuf en moindre quantité et d'autres roches volcaniques. Ce type de formation de fer siliceux atteint des épaisseurs de plusieurs centaines de

pieds et est distribué de façon continue sur des centaines et même des milliers de kilomètres près de la ligne de côte des anciens continents. Un exemple remarquable de ce type de formation de fer est la région du lac Supérieur et le géosynclinal du Labrador dans le bouclier canadien.

Une image précise de la position relative des zones couvertes par le bouclier précambrien avant la dérive des continents est nécessaire pour l'étude des milieux sédimen- taires où s'est formé le fer siliceux. Les zones où l'on trouve le fer dans l'hémisphère nord à l'intérieur des masses conti- nentales actuelles ont pu être préservées dans un milieu phanérozoïque tectonique relativement stable. Les régions où se rencontre le fer près des bordures des masses conti- nentales actuelles dans les régions équatoriales et dans l'hémisphère sud semblent avoir été fragmentées à la suite de la dérive des continents.

Des comparaisons entre des milieux où se sont formés les dépôts des formations de fer devraient permettre de reconstruire les principales zones de dépôts de fer, les vastes plateaux continentaux ainsi que les environs des bassins où les deux types se sont déposés avant la dérive des continents.

Bibliography / Bibliographie

CANADA. GEOLOGICAL SURVEY OF CANADA. 1963. Geology of northern Yukon territory andnorthwestern district of Mackenzie. Ottawa, Geological Survey of Canada. (Map 10-1963 .)

GOODWIN, A. M. 1962. Structure, stratigraphy and origin of iron formations, Michipicoten area, Algoma district, Ontario, Canada. Bull. Geol. Soc. Amer., vol. 73, p. 561-86.

GROSS, G. A. 1965. Iron-formation, Snake River area, Yukon and Northwest territories; Report of activities; Field, 1964. Geol. Surv. Pap. Can., 65-1, p. 143. -_ . 1965-68. Geology of iron deposits i?z Canada. Ottawa, Geological Survey of Canada. (Econoinic Geology Report no. 22.) Vol. I: General geology and evaluation of iron deposits (1965); Vol. II: Iron deposits, Appalachian and Grenville

regions (1967); Vol. III: Iron ranges of the Labrador geosyncline (1968).

JAMES, H. L. 1954. Sedimentary facies of iron-formation. Econ. Geol., no. 49, p. 235. - . 1966. Data of geochemistry, sixth edition, chapter W. Chemistry of the iron-rich sedimentary rocks. Prof. Pap, US. Geol. Surv., 440-W. - . 1969. Comparison between Red Sea deposits and older ironstone and iron-formation; Hot brines and recent heavy metal deposits in the Red Sea. Edited by Egon. T. Degens and David A. Ross. New York, N.Y., Springer.

O'ROURKE, J. E. 1961, Paleozoic banded iron-formation. Econ. Ceol., vol. 56, p. 331-61.

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SAPOZHNIKOV, D. G. 1963. Karadzhal'skoe zhelezo-margant- sevoe rnestorozhdenie [The Karadzhal iron-manganese deposit]. Transactions, Institute of the Geology of Ore Deposits, Petrography, Mineralogy and Geochemistry, no. 89, p. 123- 95. Moscow, U.S.S.R. Academy of Sciences. (In Russian.) (Unpublished translation by the Canada Department of the Secretary of State, Bureau for Translations).

ZELENOV, K. K. 1958. On the discharge of iron in solution into the Okhotsk Sea by thermal springs of the Ebeko volcano (Paramushir Island). C.R. Acad. sci. U.R.S.S., vol. 120, p. 1089-92. (In Russian; English translation published by Consultants Bureau Inc., 1959, p. 497-500.) - . 1970. Survey of world iron ore resources. New York, N.Y., United Nations. (Sales no. E. 69, II. C.4.)

21

Archaean volcanogenic iron-formation of the Canadian shield

A. M. Goodwin Department of Geology, University of Toronto, Canada

Introduction Iron-formation is widely distributed in Archaean (older than 2,500 m.y.) rocks of the Canadian shield. Although individual iron-formations are comparatively small, their wide distribution compensates in total quantity. Thus total estimated iron-ore reserves in Archaean iron-formation amount to 35,000 million tons with an average content of 25-30 per cent Fe. This constitutes 25 per cent of the total estimated iron-ore resource in Canada (Gross, 1968).

This paper demonstrates the genetic relationship of Archaean iron-formation to volcanism by focusing on three increasing levels of relationships: (a) the Helen iron range

where the quantities of silica present in the iron-formation are equivalent to those leached from subjacent footwall volcanic rocks; (b) the Michipicoten area where basin analysis has revealed the genetic relationship of iron facies to basin bathymetry and volcanic centres; and (c) the Canadian shield with exclusive regional relationship of Archaean iron- formation to volcanic-rich greenstone belts.

Helen Iron Range

The Michipicoten area, situated in southern Superior tectonic province (Fig. i), is underlain by Archaean

Michi picoten

FIG. 1. Location of Michipicoten area and HeIen iron range.

Unesco, 1973. Genesis of Precambrian iron and manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 23

A, M. Goodwin

FIG. 2. Geology of Helen iron range.

supracrustal rocks of Michipicoten Group and by younger intrusions. In the central part, which includes the main iron range, the Helen iron-formation forms a part of a thick, varied volcanic succession. Situated at the top of thick felsic pyroclastics and overlain by mafic flows, it occupies a unique stratigraphic position at an abrupt felsic-mafic volcanic interface (Fig. 2).

Structurally, the Helen iron-formation and enclosing volcanic rocks have been overturned to the north about an east-trending fold axis; the rocks dip steeply southward but face to the north. They plunge eastward at 30-45 degrees.

HELEN IRON FORMATION

This iron-formation comprises three distinctive and mutually transitional facies which are in descending stratigraphic succession, banded chert, pyrite and siderite members (Fig. 3). In addition, thin discontinuous chert zones are present within and at the base of the siderite member.

The main banded chert member is from 400 to 1,000 ft (120-300 m) thick. It is typically composed of alternating bands of white to grey chert and pale brown siliceous sider-

Stratigraphic tops

-, i < - < I - \ - ’

FIG. 3. Cross-section of Helen iron-formation.

24


ite. Individual bands range in thickness from barely perceptible lamina to 3 in (7.5 cm) thick. The banding is generally irregular with much pinching, swelling and bifurcation of individual layers. Thin zones of sooty, black carbonaceous, sulphide-bearing chert up to 30 ft (9 m) thick and several hundred feet long are present.

The banded chert member normally grades downwards to the pyrite, or sulphide, member, a lensy, discontinuous unit ranging in thickness from 50 ft (15 m) at the west end (1000 E in Fig. 2) to less than 10 ft (3 m) at the east end of the range (SO00 Ej. This member is composed of pyrite with some pyrrhotite, siderite, magnetite and chert. The pyrite normally has a well-developed granular texture, although veins and bands of dense, massive pyrite occur locally.

The pyrite member grades down in turn to the siderite, or carbonate, member which forms a persistent tabular body ranging in thickness from less than 100 ft (30 m) at the west end of the range to more than 350 ft (107 m) at the centre (3500 E-5000 E). It contains several hundred million tons of iron ore. The member is composed of fine-grained siderite with variable proportions of ankeritic carbonate, calcite, quartz, sulphides, magnetite, chlorite and iron silicates. The siderite is light grey to brown to black in colour and is typically dense and structureless in appearance. Local fragments of volcanic tuff as well as some thin, persistent pyritiferous shaly units are present in the siderite member.

A thin, discontinuous zone of banded chert normally 5-10 ft (1.5-3 mj thick but ranging to 35 ft (9.5 mj thick commonly lies at the base of the siderite member. It conformably overlies the footwall volcanic rocks without discernible evidence of erosion or deformation of subjacent volcanic rocks.

FOOTWALL VOLCANIC COMPLEX

The footwall complex in the vicinity of the iron range is mainly composed of alternating rhyolite-dacite tuff, flow and breccia but also includes substantial masses of carbonatized diorite (metadiorite) and smaller quartz porphyry intrusions (Fig. 2).

The footwall complex within 3,000 ft (914 m) stratigraphically of the iron-formation comprises (a> coarse grained rhyolitic pyroclastics, marking proximity to source vents to the west (1000 W to 3500 E); (b) finer grained dacitic tuffs and flows to the east (4500 E to 8000 E); these two principal volcanic facies which are transitional through (c) a mixed volcanic zone (3500 E to 4500 E) have been traced down and plunge to the east at least 3,000 ft (914 mj below surface.

The footwall volcanic complex is zoned in both lateral and vertical stratigraphic directions. Lateral stratigraphic zoning, described above, is a function of the original compositional zoning of the volcanic pile. Vertical stratigraphic zoning, on the other hand, is a function of progressive upward chemical alteration of footwall volcanic rocks. Alteration occurred while the volcanics were flat-lying and

during the general period of deposition of the Helen iron- formation. The quantity of SiO, leached from the footwall volcanic rocks equals that present in the Helen iron-formation (Goodwin, 1964).

A zone of maximum chemical alteration of footwall volcanic rocks, on average 150 ft (46 ni) thick, immediately underlies the Helen iron-formation (Fig. 3). This zone is uniformly thick regardless of local wall-rock composition. Beneath this zone of maximum alteration, the underlying volcanic rocks have themselves been altered to stratigraphic depths of at least 3,000 ft (914 m). The chemical nature of the wall-rock alteration is the same throughout varying only in degree. On this basis the footwall volcanic complex is divided in a vertical stratigraphic direction into the ‘highly’ altered zone (approximately 150 ft (46 m) thick) and stratigraphically beneath this, the mildly altered- zone (several thousand feet thick).

The chemistry of wall-rock alteration has been investigated by means of chemical analyses of fresh diamond drill core (Goodwin, 1964). Fifty-five equally spaced intersections of footwall volcanics provided 1,375 samples. Each sample was analysed for SiO,, AlzOe, total Fe, Cao, MgO, MnO, S, Tio, and loss on ignition (CO, + H,O). In addition, selected samples were analysed for Na,O, KzO, Fe0 and Fe,O,.

These chemical data reveal that wall-rock alteration resulted in subtraction of SiO, and addition of the carbonate components-Feo, MgO, MnO and CO,. The Alzo, content of volcanic rocks in individual intersections has remained essentially constant despite the degree of chemical alteration (Goodwin, 1964).

Typical oxide trends across footwall rhyolitic and dacitic assemblages respectively are shown in Figure 4. Initial average SiO, contents in rhyolite of 72.3 per cent and in dacite of 59.8 per cent decrease stratigraphically upwards within 150 ft (46 mj of the ore zone (siderite member) to an average of 64.1 and 58.4 per cent SiO, respectively. Carbonate contents increase proportionately (Goodwin, 1964).

A systematic plot of SiO, : CO, + HzO contents in the 55 intersections represented by the 1,375 wall-rock samples reveals that each weight per cent addition of CO, + HzO in excess of 2 per cent has resulted in 2.8 weight per cent loss of SiO, in rhyolite and 2.4 weight per cent loss SiO, in dacite (unpublished data), The corresponding weight per cent carbonate addition is almost equal to the weightper cent SiO, loss.

INTERPRETATION

Widespread chemical alteration of assorted volcanic rocks to stratigraphic depths of several thousand feet together with the presence of a uniformly thick, continuous, highly altered zone within assorted volcanics indicates that chemical alteration occurred as a function of depth in flat-lying volcanic rocks. The chemical data show that alteration was largely independent of local structural features such as faults and fractures as well as lithofacies boundaries. The

25

A. M. Goodwin

Rhyolitic Volconics

1

Docitic Volconics

50 ' 1

3

.C. ._... - - . d.. . . . m. - .4. : .-:.-:.-

IO 51: M g o

Feet 400 300 200 I O0 O

Horizonto/ distance south of the ore contact -

70 [ U-3-49

50' \ J

c - 5 'O!

I L io o> Co0 a 5

O ..

MgO IO 5 o

Feet 400 300 2 O0 I O0 O

Horizontal distance south of the. ore contoct A FIG. 4. Oxide trends across footwall volcanics.

mechanism of alteration apparently involved upward migration of volcanic-derived solutions by way of hot springs, fumaroles and similar migrational systems. This resulted in pervasive alteration of freshly deposited predominantly felsic pyroclastics .

According to the theory, volcanic solutions and gases were released to the volcanic pile during this explosive volcanic phase. Together with available ground waters, they migrated upwards through freshly deposited felsic pyroclastics, replete with glassy, vesicular fragments, by way of available openings including pores, fragment boundaries and fractures as well as conventional passage ways such as conduits, pipes, fissures, necks and other orifices. During their upward migration, the primary, bicarbonate-sulphur- charged hydrous solutions promoted pervasive devitri- fication of volcanic glass and breakdown of primary rock silicates to secondary mineral assemblages together with release of surplus soluble SiO,. This material there-

upon joined the upward migrating iron bicarbonate-charged solutions.

High confining pressures in the lower levels of the volcanic pile presumably restricted the degree of chemical alteration (mildly altered zone). Within 150 ft (46 m) of the top of the volcanic pile, however, the expanding internal pressures in the solutions gradually exceeded the confining pressures exerted by the overlying rock and water columns. As a result chemical alteration in the upper 150 ft (46 m) attained a maximum ('highly' altered zone).

The volcanic solutions, charged with primary, igneous- derived iron bicarbonate and sulphide components and aug- mented with leached Sioz, thereupon entered the overlying aqueous environment whether sea or lake. Sequential precipitation of chemical components resulted in accumulation of, in order, siderite, pyrite and banded chert members. The main environmental controls were probably, in order of effect: (a) release of pressure at the rock-water interface

26


with loss of CO, resulting in rapid, blanket deposition of massive siderite in proximity to centres of discharge and (b) gradual temperature decrease resulting in delayed precipitation of banded chert upon the main siderite mass. The intermediary position of the sulphide member may reflect similar pressure-temperature control in addition to the influence of selective biogenic activities (Table 1).

TABLE 1. Principal environmental controls and chemical products during precipitation of Helen iron-formation

Subaqueous environment

Surñcial Increasing (volcanic) time-

environment Near source distance from source

~

Pressure Temperature

P H Eh Salinity Organic activity Principal chemical products

High Higher

Low Low Nil Nil

Sileceous bicarbonate and sulphide solutions; wallrock alteration

Lower Low High- Low

Intermediate Intermediate Higher

Low Higher Low Higher

High (?) High (?) Siderite Pyrite Chert

In summary the main chemical components of the Helen iron-formation are considered to represent direct volcanic contributions. The characteristic threefold facies construction is attributed to variations in the aqueous environment during the period of precipitation and lithification. Careful evaluation of the chemical data has shown that the SiO, requirements of the Helen iron-formation may be reasonably met by chemical leaching of footwall volcanics (Goodwin, 1964). Taken together with the intimate stratigraphic relations, this constitutes the main quantitative evidence for a direct volcanic origin.

Michipicoten basin

The Michipicoten area, 70 miles (112 km) long by 30 miles (48 km) broad, is underlain by mafic to felsic volcanic rocks, clastic sediments and banded iron-formation, including the Helen formation discussed above, in addition to younger intrusions. Of present concern are the distribution, nature and relationship of the supracrustal rocks, particularly felsic pysoclastics, clastic sediments and iron-formation, to the original basin of deposition (Goodwin and Shklanka, 1967).

Michipicoten rocks have been complexly folded about east-trending and north-west-trending áxes. This has resulted in complex, doubly plunging fold patterns in both longitudinal and cross-section (Fig. 5). In addition, numerous north-trending faults have disrupted the rocks.

As a result of structural deformation the original lithofacies pattern has been distorted to prevailing easterly trends. To reconstruct the original facies pattern it has been necessary to unravel the superimposed structures and to consider the lithofacies within individual rock-stratigraphic units throughout the area. The net result reveals a predominant northerly trend of the various lithofacies within the main rock-stratigraphic units. This northerly trend corresponds to a principal direction of basin configuration during Michipicoten time. Accordingly an east-west stratigraphic section of the Michipicoten area which is oriented normal to this principal direction of basin configuration, transects the original lithofacies trend.

STRATIGRAPHY

Michipicoten stratigraphy exihibits volcanic and sedimentary facies of conventional ‘eugeosynclinal‘ type. Lensoid rather than ‘layer-cake’ stratigraphic relations predominate. Some units blanket the asea but most have limited distribution.

The Michipicoten assemblage contains distinctive volcanic and sedimentary facies (Fig. 6). The lowermost mafic volcanic member (1) is continuous across the area. It is overlain in the west by clastic sediments (3) which thin and become finer grained to the east, and by felsic volcanic masses (2) in the east which are, in turn, overlain by mafic volcanics (5, 6). Iron-formation is transitional from sedimentary association in the west to volcanic association in the east.

Large masses of felsic volcanic rocks are present at Goudreau in the east, Magpie in the centre and Wawa in the south-centre. The felsic piles contain a great assortment of andesite-dacite-rhyolite pyroclastics and flows. The more silicic types, especially rhyolite, are concentrated at the top of the piles. The coarser, more angular pyroclastic fragments as much as three ft (91 cm) in diameter are more common in thicker parts of the piles. Elsewhere tuff and tuff breccia predominate. Some andesite-dacite lava flows are present. Carbonatization of the felsic pyroclastic rocks is a common feature.

The felsic volcanic masses are clearly products of highly explosive eruptions which rapidly produced thick, irregular, high-rising piles. The ‘coarse fragment-thick pile’ spatial relationship described above reflects proximity to eruptive sources probably central vents.

IRON-F OR M A TIO N

Iron-formations are present in nearly all parts of the Michipicoten area. In central and eastern parts, they occur in volcanic rocks; enclosed, they typically occur at prominent felsic eruptive-mafic effusive contacts. In the western part of the area, laterally equivalent iron-formations are enclosed in clastic sediments.

Michipicoten iron-formations of volcanic association

27

A. M. Goodwin

I I West I Idealized section of Michipicoten basin illustrating shore-to-depth relationship of iron facies.

Dore sediments: conglomerote; greywocke, shale. STRUCTURAL SECTIONS

Felsic volcunics.

I/---I Geologhl boundary. [T=j Fuult.

Gradational ond approximoie boundory of iron facies.

FIG. 5. Iron facies distribution in Michipicoten basin

in central and eastern parts of the area contain a threefold arrangement of, in descending order, banded chert, sulphide and carbonate members (Fig. 7, Helen and Goudreau sections). The banded chert member is commonly 100-200 ft (30-60 m) thick; it reaches a maximum thickness of 1,000 ft (304 m) at the Helen range in the central part but is thinner or absent to the east, The sulphide member is commonly 10 to 30 ft (3-9 m) thick but attains a maximum of 120 ft (36 m) at Goudreau in the eastern part of the area. The underlying carbonate member is a similarly lenticular and discontinuous unit composed mainly of iron-bearing carbonate minerals with minor pyrite, magnetite, pyrrhotite and silicate minerals. The carbonate member is commonly less than 200 ft (60 m) thick; it reaches a maximum thickness of 350 ft (106 m) at the Helen iron range near Wawa. The member is predominantly sideritic in the central part of the area but becomes increasingly calcareous to the east.

Michipicoten iron-formations of sedimentary association in the western part of the area (Fig. 7, Kabenung section) comprise thick cherty units composed of thinly interbanded chert, siliceous magnetite and jasper (hematitic chert). Individual zones of iron-formations may reach thicknesses of 600 ft (180 m). This type of iron-formation is typically intercalated with greywacke and shale.

Michipicoten iron facies are gradational across the area (Fig. 5). Oxide facies predominates in the west as at Kabenung Lake, carbonate facies in the centre, as at the Helen and Magpie ranges, and sulphide facies in the eastern part, as at Goudreau. The sinuous yet parallel distributions of the oxide-carbonate and carbonate-sulphide facies boundaries are shown in Figure 5. Examination of structural relations reveals that this sinuous pattern is due mainly to regional folding.

Thus Michipicoten iron-formations are arranged from

28


FIG. 6. Reconstructed east-west section of Michipicoten basin.

west to east in oxide, carbonate and sulphide facies, the three facies being broadly transitional one to the other across the area (Fig. 5). Considered relative to clastic sedimentary patterns, it is apparent that oxide iron facies is largely coincident with shallower water, conglomerate- bearing clastic facies in the western part, the carbonate iron facies with deeper-water, conglomerate-free sediments in the central part, and the sulphide iron facies with still deeper-water, shaly sediments in the eastern part of the area. This relationship points strongly to common environmental controls during iron-silica precipitation. The principal factors were depth of water coupled with a transition from shallow water, and the <oxidizing' environment in the west to deeper water, reducing environment in the east. In this regard it is noted that the environment in the west permitted extensive precipitation of the ferric oxide- bearing minerals, hematite and magnetite.

Considered relative to tectonic and eruptive elements it is apparent that a north-trending line through Wawa and Magpie marks the western limit of felsic eruptive masses. This line also coincides in a general way with the oxide- carbonate iron facies boundary. This relationship, together with the presence of shallower-water clastic facies to the west, implies that an eastward sloping shelf-to-depth aqueous environment was functioning during this period of iron-silica accumulation,

Considered relative to felsic pyroclastic distribution it is apparent that the oxide iron facies was deposited to the west of the main felsic pyroclastic piles, the carbonate iron facies upon coarse-grained felsic pyroclastics, and the sulphide

facies upon finer grained felsic pyroclastic piles in the east. Michipicoten iron-formations are viewed as' direct vol-

canic products because: (a) the major concentrations of iron (siderite and pyrite) overlie the thickest felsic volcanic piles thereby implying maximum iron accumulation in proximity to volcanic vents; (b) iron-formation was deposited rapidly at the end of a mafic-to-felsic volcanic cycle implying an end stage of igneous differentiation as a source of iron; (c) deposition of iron-formation was associated with pervasive leaching of silica from freshly deposited felsic pyroclastics. In brief the relationship of Michipicoten iron-formation to volcanism is direct and impressive.

Although the chemical components of Michipicoten iron-formations were volcanically derived and thus associated with specific exhalative centres, yet the environmental parameters exerted by the existing basin configuration were sufficient to have produced a general shelf-to-depth facies pattern that conforms to the world-wide pattern of iron deposition as presented by James. Thus Michipicoten iron facies, both in vertical and lateral distribution, are attributed to depositional changes and transitions of environment within the Michipicoten basin. It is noteworthy that the predominantly cherty oxide facies iron-formation which is the most commonly recognized Archaean iron-formation is best developed in the sedimentary environment on the slopes of the basin some distance removed from volcanic exhalative sources. Thus within an Archaean basin, the distribution of cherty oxide facies iron-formation may be used to locate the margins of the basin rather than specific volcanic sources (Goodwin and Ridler, 1970).

29

A. M. Goodwin

KABENUNG SECTION (Oxide Facies)

HELEN SECTION (Carbonate Facies)

GOUDREAU SECTION (Sulphide Facies)

~

:IC. 7. Typical sections of oxide, carbonate and sulphide facies.

Pyrite

Banded chert Interbedded chert-magnetite

Canadian shield Archaean iron-formation is broadly distributed across the Canadian shield. Distribution patterns (Figs. S(a) and (b)) are based substantially on a compilation (Geol. Surv. Canada, Maps 13-1963 to 20-1963 inclusive and Map 1187A) prepared by G. A. Gross in 1963, with some additions (Davidson, 1969). Additional bands of Ar- chaean iron-formation will doubtless be revealed particularly in northern parts of the shield. Also certain iron- bearing belts presently classified as Proterozoic in age may in future be reclassified as Archaean. However, considering that all parts of the Canadian shield will have been mapped, at least on reconnaissance scale, by 1972, there is reason for confidence in present distribution patterns.

Carbonate s-siderite c-limestone

1: ,.'L'j Rhyolite-dacite tuff, brôccia,flows

The term ' iron-formation' has been generally applied to layered rocks with 15 per cent or more iron (Gross, 1965). The most commonly recognized type is chert-magnetite or oxide iron facies. Problems of classifying other types (including associated non-ferruginous chemical sediments) e.g. sulphide-bearing aiid calcareous facies, pose certain problems not yet satisfactorily resolved. For example, should the numerous thin, pyritiferous zones common in quartz-mica schist be classified as iron formation? What about massive sulphide ore deposits, e.g. Kidd Creek Cu- Zn-Ag deposit, of presumed exhalative origin? A broader definition of the term iron-formation may be required in future. In this event established distribution patterns of Ar- chaean iron-formation in the Canadian shield would require appropriate revision.

30

100. 92. 84. 76. 56.

\

LEGEND

Other Precambrian rocks ; mainly grmitic Archaean iron-formation Archaean sedimentary rocks Archaean felsic volc~nic rocks Archáaan mafic voicanic rocks

Boundary of shield

I I I I 1 \ I 100. 9 2' 64' 76'

116. 100- 92' 64. 76.

bn'

. Boundary of shield Orher Precambrian rocks ; mainly grmitic - Tcelonic boundary Archaean iron-formation Archaean sedimentary rocks Archaean feisic volcanic rocks o

Scale rn Miles

106. 100. 92' 64' 76'

(b)

FIG. 8. Distribution of Archaean iron-formation in Canadian shield : (a) southern part ; (b) northern part.

A. M. Goodwin

Flyschoid Facies

DISTRIBUTION

Volcanogenic Facies

TECTONIC SETTING

The distribution of recognized iron-formation in Archaean rocks is shown in Figure 8(a) (southern shield) and 8(b) (northern shield), With few exceptions iron-formation is located either in volcanic-rich belts or in nearby sediments. Iron-formation within volcanic-rich belts is closely related to felsic volcanic centres represented by felsic volcanic rocks and cogenetic tuff and greywacke.

The largest concentration of Archaean iron-formation is in Superior Province (Fig. 8(a)). At least thirty volcanic- rich belts contain iron-formation mainly associated with thirty-four established felsic concentrations each a volcanic centre. In Churchill Province iron-formation is present in six greenstone belts; the major iron-bearing belt is in the Rankin-Ennadai area west of Hudson Bay. In Slave Prov- ince a single, small iron-bearing belt has been recognized.

Obviously the abundance of Archaean iron-formation is directly proportional to the number of volcanic-rich greenstone belts especially those containing felsic pyroclastics and related sediments.

Craton

The typical association of Archaean iron-formation is illustrated in a reconstructed section of the Archaean crust (Fig. 9; Goodwin, 1968). According to this model the main lithofacies, arranged in lateral succession outwards from the more stable parts of the primitive crust, are: (a) sialic craton; (b) flyschoid facies; and (c) volcanogenic facies. The predominant orogenic characteristics of both the flyschoid and volcanogenic facies point to their accumulation in mobile belts corresponding to thin-crustal zones such as may be reasonably ascribed to the margins of, and between (i.e. intracratonic), expanding sialic cratons.

The primitive sialic cratons of the Canadian shield are identified only on the basis of widespread clastic detritus of appropriate sialic composition. The exact nature and degree of stability of the presumed cratons are conjectural beyond their function as a source of sialic detritus.

Flyschoid facies, which includes the common quartz- feldspar mica schist assemblage of the Canadian shield, represents orogenic detritus apparently derived by rapid weathering of sialic provenances and deposited in the manner of classical flysch. The facies is moderately thick (10,000-20,000 ft (3,048-6,096 m)) and tends to be lithologically uniform and stratigraphically continuous.

I I Sedimentary. Basin Island Arc I

. . ..

L E G E N D t i Approx. 100 Miles [1113 Greywacke, argillite Basalt

Granodiorite. &@ Gabbro, norite, etc. . . . . . . . . . Dacite, rhyolite Voicaniciastics

Andesite a iron-formation FIG. 9. Idealized crustal section showing Archaean lithofacies.

32


The flyschoid facies is transitional through a subfacies to the volcanogenic facies, the principal Archaean volcanic- bearing facies. The volcanogenic facies, which is commonly

(9,140 m)). It typically comprises basalt-andesite-rhyo-

ments (iron-formation) and numerous intrusions. Evidence of tectonic instability during accumulation is widespread. Abundant pillow lavas indicate a prevailing subaqueous environment of accumulation.

Iron-formation is common in the volcanogenic facies as well as iii transitional sub-facies the latter being dominantly volcaniclastic and located on the flanks of the volcanic piles. Locally some iron- formation is also present in marginal flyschoid zones adjoining volcanic assemblages.

TABLE 2

Weighted average percentage composition called greenstone, is very thick (approximately 30,000 ft

lite assemblages, assorted volcaiiiclastics, chemical sedi- SiO, CO, HZ0

ORIGIN

Weighted average compositions of Archaean volcanic rocks of the Canadian shield have been determined 011 the basis of detailed geochemical studies in four widely separated belts (Baragar and Goodwin, 1969). The results demonstrate that Archaean volcanic rocks have substantial CO, and H,O contents; these are listed in Table 2 with corresponding SiO, contents (Baragar and Goodwin, 1969).

Thus the indicated average volatile content of Archaean volcanic rocks is sufficient to have facilitated leaching of substantial quantities of SiO, from the volcanic rocks. Although the presence of CO, and H,O in the volcanic rocks does not constitute proof of pervasive leaching of SiO, at the time of deposition of iron-formation, yet the demon-

Average Archaean

Average Archaean mafic

Average Archaean felsic

volcanic belts 53.9 1.1 2.4

volcanic fraction 52.6 1.2 2.5

volcanic fraction 66.8 1.1 1.5

strated example of leaching of SO, from felsic volcanic rocks at the Helen iron range and the intimate shield-wide association of iron-formation with volcanic rocks do support a silica-leaching mechanism for obtaining most if not all the large quantities of SiO, present in Archaean iron- formation of the Canadian shield.

Acknowledgements

I gratefully acknowledge the value of discussions with G. A. Gross on the distribution of Archaean iron-formation in the Canadian shield and with A. Davidson on recent investigations in the Rankiii-Ennadai belt of Northwest Territories.

Some of the illustrations were prepared by F. Jurgeneit and all were photographed for reproduction by P. B. O’Do- novan, both of the Department of Geology, University of Toronto.

Résumé

La formation de fer volcanogéniqsre archéenne du bouclier canadien (A. M. Goodwin)

Des formations de fer rubanées datant de l’époque ar- chéenne (c’est-à-dire ayant plus de 2,5 milliards d’années) présentent des relations étroites de localisation et d’origine avec les roches volcaniques dans de nombreuses zones ar- chéennes de (( greenstone >) du bouclier canadien. Dans la région de Michipicoten, dans la partie septentrionale de l’Ontario central, une formation de fer du type (( Algoma )) comprenant, de bas en haut, dans l’ordre stratigraphique, des faciès carbonate, sulfuré, et de silex est située à un contact stratigraphique remarquable entre les pyroclastites- felsiques sous-jacentes et les coulées de lave mafique sus- jacentes. Des études détaillées de l’altération de la roche encaissante portant sur plus de 2 500 analyses complètes de roche couvrant une zone stratigraphique de roches volcaniques sous-jacentes de plus de 5 kilomètres carrés et atteignant des profondeurs stratigraphiques de près d’un

kilomètre ont permis de définir la nature et le degré de son altération chimique, altération attribuée à des sources thermales contemporaines et à une activité fumérolique. Parmi les éléments qui composent la formation de fer on trouve : le fer, le manganèse et le gaz carbonique, qui se sont ajoutés aux roches volcaniques sous-jacentes et ont évidemment une origine volcanique plus profonde. Cependant SiO, a été lixivié des roches volcaniques felsiques immédiatement sous- jacentes en quantités à peu près égales à celles présentes dans la formation de fer sus-jacente. Ainsi les relations de localisation et génétiques de cette formation de fer aux roches volcaniques et aux processus volcaniques sont à la fois directes et remarquables.

Outre la construction du faciès vertical à l’intérieur de l’empilement des roches volcaniques mafiques et felsiques, les formations de fer de Michipicoten présentent des changements de faciès latéraux qui rappellent la construction du bassin original. Ainsi, les formations de fer où prédo- minent les faciès oxydés, carbonatés et sulfurés sont associées

33

A. M. Goodwin

avec des structures volcaniques sédimentaires de plus en plus profondes quand on progresse de l’ouest vers l’est à travers la région de Michipicoten. Ce système de faciès latéraux rappelle le dépôt depuis le littoral jusqu’aux couches profondes de composants chimiques volcaniques, Fe - M n - CO, - S - Sioz, sur des pentes inclinées vers l’est dans un bassin archéen primitif. Le milieu archéen a permis apparemment que se déposent des combinaisons d’oxydes (magnétite-hématite par exemple).

Des relations volcaniques similaires se retrouvent dans plus de trente bandes de (( greenstone )) du bouclier canadien, où les formations de fer sont associées directement

avec des empilements volcaniques épais ou des assemblages clastiques voisins. Dans bien des régions du bouclier, les schémas de distribution : oxyde, carbonate, sulfures, sont des indications utiles sur la construction du bassin original. Bien qu’elles soient abondantes surtout dans la région du lac Supérieur, les formations de fer archéennes sont assez communes dans les provinces de Slave et de Churchill pour mettre en évidence que les conditions qui prévalaient dans la croûte ont favorisé le développement de la formation de fer volcanogénique dans toutes les parties de la croûte primitive précambrienne représentées dans le bouclier canadien.

Bibliography/ Bibliographie

BARAGAR, W. R. A.; GOODWIN, A. M. 1969. Andesite and Archaean volcanisin in the Canadian shield. In: A. R. McBir- ney (ed.), Proceedings of the Andesite Conference. Bull. Ore. Dep. Geol., no. 65.

DAVIDSON, A. 1969. Eskimo Point and Dawson Inlet map- areas. District of Keewatin (55E, north half, 55F, north half). Geol. Surv. Pap. Can. 70-1, Part A, p. 131-33.

GOODWIN, A. M. 1962. Structure, stratigraphy and origin of iron-formations, Michipicoten area, Algoma district, Ontario, Canada. Bull. geol. Soc. Amer., vol. 73, p. 561-86. __ . 1964. Geochemical studies at the Helen iron range. Econ. Geol., vol. 59, no. 4, p. 684-718.

-- . 1968. Preliminary reconnaissance of the Flin Flon volcanic belt, Manitobaandsaskatchewan. Geol. Surv. Pap. Can. 69-1. Part A, p. 165-68. __ ; RIDLER, R. H. 1970. Abitibi orogenic belt. Geol. Surv. Pap. Can. 70-40, p. 1-24. __ ; SHKLANKA, R. 1967. Archaean volcano-tectonic basins: form and pattern. Canad. J. Earth Sci., vol. 4, p. 777-95.

GROSS, G. A. 1965. Geology of iron deposits in Canada. Geologi- cal Survey of Canada (Economic Geology Report no. 22), 181 p.

-. 1968. Detailed survey tabulates billions of tons of iron. North. Min., November 28, Annual Review Number, p. 51.

34

The facial nature of the Krivoyrog . n .. iron-Iorrnation

A. I. Tugarinov, I. A. Bergman and L. K. Gavrilova Institute of Geochemistry and Analytical Chemistry, Academy of Sciences of the U.S.S.R., Moscow (Union of Soviet Socialist Republics)

The Krivoyrog iron-formation includes a complex of metamorphosed sedimentary rocks of magnesium-iron and iron composition. Beginning with a talc horizon, it passes into a rhythmically constructed stratum of successively alternating shale and iron horizons. In the central part of the Krivoyrog Basin (Saksagan region), where the iron-formation is most intensively developed, up to seven rhythms are distinguishable. The thickness of the separate horizons varies over a wide range, averaging from tens to a few hundreds of metres.

The abundance of rocks in the Krivoyrog iron-formation is shown in Table 1 according to data from a study in five sections (Skelevatski-Magnetitovy, Zelenovski district, Inguletz, Zholtaya Reka and Frunze mines).

TABLE 1

Rocks Abundance (ratio, %)

Iron cherts and jaspilites 55 Magnesium-iron schists 19 Ore-free and low in ore cherts 13 Talc-containing rocks 7 Carbonaceous shales and cherts (carbonate- alumosilicate) 3

Alumosilicate shales and cherts 2.5 Others 0.5

Structural peculiarities of the Krivoyrog iron-formation

The second peculiarity is that quartz-chert interlayers are integral elements of all rocks which compose shale and iron horizons. From the point of view of the general thickness of quartz-chert interlayers, three rock groups are distinguishable (Table 2).

TABLE 2

Rocks Contents of quartz-chert

interlayers (in volume %)

Shales Iron cherts and jaspilites Ore-free and low in ore cherts

20-40 40-60 60-80

Between the rock groups shown in Table 2, all types of mutual transition are taking place. In other words, quartz-chert interlayers alternate both with shale and ore interlayers in practically any quantitative ratios. Hence, to a certain degree quartz-chert interlayers are, with respect to shale-ore interlayers, an independent component- formational peculiarity and not a peculiarity connected with a concrete composition of some rocks.

The third peculiarity of the Krivoyrog iron-formation is the dual nature of its structure. Within one formation there are depositions of two facial series, one at present represented by quartz-chert interlayers and the second by slate- ore interlayers. Alumosilicate slates and iron cherts are two extreme subelements of the rhythm into which the depositions of the second series are differentiated.

Facial position of the The first peculiarity of magnesium-iron and iron rocks is the rhvthmic character of their structure-the alternation of

~~i~~~~~~ iron-formation quartz-chert (quartzite) with interlayers which, as it will be shown below, form a successive series connected by mutual transitions: alumosilicate shales-magnesium-iron schists- iron cherts.

No data are available for a direct reconstruction of the region of iron-sediment accumulations and of their relationships to such types of sediments. At present, the alternation

Unesco, 1973. Genesis of Precambrian iron and nranganesr deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 35

A. I. Tugarinov, I. A. Bergman and L. K. Gavrilova

of rocks in the section, especially their composition and structure, are the sole sources of information.

During the determination of the facial position of the Krivoyrog iron-formation, the following observations were taken into account: (a) the iron-formation replaces in time the lower terrigenous formation, its thinnest fractions; (b) in the lower terrigenous formation only the ‘roots’ of the iron- formation can be found (cummingtonite shales,l relicts of iron carbonates in phyllites and some other, less distinct, signs), on the other hand rudaceous sediments are lacking in the iron-formation itself; (c) the persistent character of the iron-formation and of the rocks composing it; (d) the high degree of differentiation of the material; (e) the distribution of alumina in the rocks, etc. All these features bear witness to a facial displacement of the iron-formation into a region of pelagic conditions rather than coastal ones with respect to pelites.

In Palaeo-Cenozoic formations some types of carbonate deposits (pelitomorphic limestones) occupy a facial position of iron sediments.

Compositional peculiarities of the iron-formation

From the point of view of mineral composition, the magnesium-iron shales-iron cherts (including jaspilites) are composed of one and the same suite of minerals, but in different quantitative ratios? magnesium iron carbonates, magnetite, cummingtonite (grunerite), chlorite, garnet, stilpnomelane, hematite, biotite.

Within the Saksagan region a noticeable, and sometimes even a cardinal, role of the ore mineral in the composition of magnesium-iron and iron rocks is fulfilled by carbonates: siderite-sideroplesite (-pistomesite).

Structural interrelations do not reflect the evolution of the mineral composition of rocks (Betekhtin, 1951), but to study them is an indispensable precondition for singling out mineral parageneses of different ages and for the analysis of the paragenetic relationships of minerals.

In iron rocks of the Krivoyrog iron-formation, beyond the oxidation zone, the fractured and ore zones, a quite definite sequence of structural replacement is revealed, shown by the series: siderite -+ garnet i cummingtonite -+ magnetite -> quartz + alkaline amphiboles -+ hematite + chlorite i stilpnomelane +- biotite.

A typical feature of structural interrelations between minerals is their evolution in the process of regional metamorphism: in some cases positional and genetical connexions are lost (siderite-magnetite, siderite-cummingtonite), in other cases reactional relationships arise which are of no genetic significance (garnet-cummingtonite, cummingtonite- magnetite). Contradictory views on the formation of magnetite, cummingtonite and other minerals are explained by the fact that investigators study rocks in which the structural interrelations have ‘stopped’ at different movements of their evolution.

Taking into account the character of structural inter- relationships between minerals, we have singled out the following typical parageneses of the second temperature degree of regional metamorphism in magnesium-iron rocks -‘cummingtonite + magnetite’, in magnesium-iron rocks with pelitic matter-‘garnet + cummingtonite’, in rocks which are intermediate in composition-‘garnet + cummingtonite + magnetite’. Transition to the second degree of progressive metamorphism is associated with the appearance of rocks with different quantitative ratios of minerals of both degrees: ‘siderite + magnetite’, ‘siderite + cummingtonite’, ‘siderite + cummingtonite + magnetite’, etc. Rocks with higher temperature degrees of regional metamorphism have local spreading.

Siderite is the sole mineral which does not replace other iron minerals. In its turn it is replaced by cummingtonite, magnetite, chlorite, stilpnomelane and others. This is why strict proportionality in the content of siderite with each of the enumerated minerals has not been observed in bi- mineral rocks (e.g. siderite-magnetitic rocks); the inversely proportional dependence is frequently disturbed as a consequence of siderite displacement by quartz.

The structural interrelations between siderite and magnetite, as they bear the most direct relation to the problem of the genesis of iron-formations, will now be dealt with in greater detail.

Siderite and magnetite are the chief ore minerals of the Krivoyrog iron-formation outside the oxidation zone in areas with a low degree of progressive metamorphism (e.g. Frunze mine).

In iron rocks where iron silicates and quartz either are absent or are contained in negligible amounts, siderite and magnetite are positionally and genetically closely connected. The following cases occur: siderite is solely replaced by magnetite; siderite is almost simultaneously replaced by quartz and magnetite (late quartz displaces siderite which has not been replaced by magnetite); siderite is solely replaced by quartz. With the increase of the quartz content a separation of magnetite from siderite occurs, chiefly at the expense of siderite, and the earlier existing genetic relationship between them is lost.

Hematite3 characterizes, in the paragenesis with chlorite and/or stilpnomelane, the regressive stage of regional met amorphism.

The mineral-forming processes in rocks of the iron- formation are concluded by biotite crystallization. It seems that its development was the consequence of the superimposed regional migmatization which, during the Post- Krivoyrog time, covered the territory directly to the west of the Krivoyrog structural-facial zone.

Let us first consider the chief regularities of minor- element distribution in iron ores in general from the tectonic and the physico-chemical aspects.

1. Cummingtonite is formed solely according to the reaction mentioned

2. Evaluated for shale ore interlayers. 3. Outside the zone of oxidation, the ore and fractured zones the hematite

below.

content amounts to 5 per cent, rarely more.

36

The facial nature of the Krivoyrog iron-formation

In the tectonic aspect, the regularities of minor-element distribution were studied by Strakhov (1947), according to whom ores formed in regions of just-completed folding are distinguishable by the greatest concentration and the greatest diversity of minor elements. The least accumulations and the least diversity of minor elements are met in geosynclinal ores. In this respect ores of platform regions occupy an intermediate position.

Arkhangelsky and Kopchenova (1934) approached the interpretation of minor-element distribution from other standpoints. They have established that the chemical composition of iron ores depends on their formation conditions. Ores of an oxidizing medium contain substantially more admixtures (e.g. phosphorus, arsenic, vanadium, nickel, cobalt, chromium) than siderites and other ores of a reducing medium, i.e. the primary ore of Precambrian quartzites was siderite.

When comparing both variants we can easily establish that the chief factor controlling the distribution of this group of elements is the physico-chemical medium of ore formation and their tectonic position becomes less significant.

Tables 3 and 4 show the distribution of typical minor elements in the iron Krivoyrog cherts of the iron-formation.

For comparison, data on iron-formations of similar tectonic position and age, iron ores of oxygenous and oxygen-free media and clarkes of minor elements in sedimentary rocks are included.

Analysis of the distribution of these elements permits the following conclusions to be drawn: In iron cherts and jaspilites of the Krivoyrog iron-formation (e.g. Kursk Magnetic Anomaly), elements which have a tendency to accumulate by sorption or by the formation of slightly soluble compounds of the type of arsenates or molybdates of ferric iron (vanadium, chromium, nickel, cobalt) are present in amounts smaller by as much as a half or even a whole order of magnitude than their clarkes.

Iron cherts and jaspilites of different chemical composition and different degree of metamorphism do not differ in general in the character of distribution of this group of elements and their contents in samples from various sections are similar.

In their content of minor elements iron cherts and jaspilites of the Krivoyrog (and Kursk Magnetic Anomaly) markedly differ from the marine iron ores of the oxygenous zone and, on the contrary, have much in common with

TABLE 3. Distribution of sulphur, phosphorus and arsenic (X10-4 per cent)

Region of deposit Author Rocks S P As

Krivoyrog

K, Basic syncline

Krivoyrog Kursk Magnetic Anomaly, Mikhailovskoye deposit ((

((

Kursk Magnetic Anomaly ((

((

(<

Lake Superior Negaunee

(< Gunflint <( Ironwood <( Iron River

Ukrainian S.S.R., Kerch

France, Landres- basin

Amerrnont basin

Gershoig (ed.-in-chief Magnetite cherts Belevtsev, 1962)

(C Martitic and specularite-martitic

(( Magnetitic, martitic and goethite- jaspilites

hematite cherts and jaspilites Fomenko and Chernovsky Amphibole-magnetitic and carbonate- (Belevtsev, ed.-in-chief, magnetitic cherts 1, 2 and 4 of iron- 1962) containing horizons

Fedorchenko (1965) Sideritic rocks Illarionov (1965) Hematitic quartzites

Van Hise, Baley and Smith

Irving and Van Hise (1892) Huber (1959) James (1951) Vinogradov (1 962) Green (1953) Litvinenko (1964)

(1 897)

Specularite-magnetitic quartzites Magnetitic quartzites Ore-free and low ore quartzites Cummiiigtonite-magnetitic quartzites Dolomite-magnetitic quartzites Hematite-magnetitic quartzites

Hematite-magnetititic quartzites

Carbonate cherts

(hematite : magnetite = 10 : 90)

(hematite : magnetite = 40 : 60)

(<

((

((

Clays and shales, average Limestones, average Oolitic ores

Coche et al. (1954, 1955) Minette ores

770

1,100 3,300

310 420 610

3,300 3,000 1,480

670

420

590 525

3,000 1,100 1,850

740

410 no data-5 samples

50-1 sample

170 no data-15 samples

70-13 samples

270 340

160 510 530 700 750 740

710

540

130 570 470

4,000 770 6.6

1,100 9,700 1,020

8,300

37

A. I. Tugarinov, I. A. Bergman and L. K. Gavrilova

TABLE 4. Distribution of the iron-group elements, uranium and molybdenum ( X IO-* per cent

Novo-Yaltinsk deposit

Region of deposit Author Rocks Ti , M n V Cr Ni Co U M o

Tarkhanov (1969)

According to authors

Plaksenko and Koval (1967)

Iron quartzites

Iron cherts and jaspilites

Magnetite quartzites

Specularite-magnetite

Specularite quartzites Silicate-magnetite quartz-

Magnetite quartzites Specularite-magnetite

Magnetite quartzites Specularite-magnetite quartzites

Specularite-magnetite quartzites

Specularite quartzites Iron quartzites

quartzites

ites

quartzites

30 300 < 10 20

13

< 10

5

<10 1.10 1.0

6/ 1.3 340 7

70 100 20

50 45

60 50

13 10

<(

50 1 O0

730 720

38 38

20 20

35 30

620 330

38 13

20 22

340 17 30 20

650 80

2,000

25 15 20

25 25 200

25 25 30 Gorlitski

(1962)

Vinogradov (1962)

Green (1953) Ronov 1956) Baranov et al. (1956)

Gerasimova etal. (1969)

60 10

zone Clays and shales, average

Limestones, average Carbonate rocks Carbonate rocks

4,500 670 130 1 O0 95 20 3.2 2.0

385 < 10 2 O? 7.5

O 1.3

2.1 4.3 Russian platform

Limestones 5,400 200

1,300

80 30 470

10 8 20

1

6 - 1

2 -

- 3 32

Dolomites Siderites Carbonate rocks Volga-Ural, Emba

regions and N.W. Caucasus Ukrainian S.S.R., Kerch basin

Katchenkov (1959)

Litvinenko (i 964) and Kantor (1937)

1,200 9 24 11 Oolitic ores

16,000 590 200 190 1. According to three out of five.

carbonate rocks of the Russian platforni, the Volga-Ural and Emba regions and the North Caucasus.

All these peculiarities and, in the first place, the lower contents than clarkes of minor elements which have a tendency to accumulate by sorption or by the formation of slightly soluble compounds with ferric iron, may be explained solely by the fact that iron was present in the primary sediment in a form in which neither sorption, nor formation of slightly soluble compounds, could manifest themselves. Such a form could only be the carbonate with all the ensuing consequences.

A number of factors (such as the hardly probable 100 per cent reduction of iron oxides, the facial position of iron sediments being peculiar to chemogenic carbonates and not to iron oxides, the scale of the phenomenon, the absence

of typical concretional forms, and composition peculiarities of magnesium-iron carbonates, including pistomesites unknown as diagenetic formations) point to a chemogenic and not to a diagenetic nature of ferric carbonates and, hence, to the formation of iron cherts and jaspilites in the process of metamorphism of only chemogenic carbonates.

Conclusions

The Krivoyrog iron-formation is a binary flyschoid formation which consists of two rhythm elements: quartz-chert and shale-ore interlayers, the second rhythm element in its turn differentiating into two rhythm subelements-pelites and iron carbonates. The first corresponds to alumosilicate

38

The facial nature of the Krivoyrog iron-formation

shales, the second to iron cherts; magnesium-iron shales occupy an intermediate position corresponding to magnesium-iron marls. reducing conditions prevailed.

The chemogenic nature of magnesium-iron carbonates of the Krivoyrog iron-formation, which is in agreement with finds of magnesium-iron carbonates in pebbles of

conglomerates (Molass formation, according to Kalyaev (1965)), permits us to conclude that in the Lower Proterozoic

The motive factor in the evolution of the iron-ore process in the history of the Earth was the progressive increase of oxygen content.

Résumé

Les faciès des formations ferrugineuses du Krivoyrog (A. I. Tugarinov, I. A. Bergman, L. K. Gavrilova)

1. Des recherches dans la zone de strates de Krivoyrog, axées sur la paragenèse minérale, mettent en évidence que le minerai de fer siliceux trouve son origine dans les roches carbonatées dont le dépôt a commencé par des carbonates de manganèse ferrugineux. Tous les minéraux qui sont venus par la suite ont été formés au cours de métamor- phismes progressif et régressif.

2. La comparaison des données recueillies sur les micro-éléments a montré que la région de Krivoyrog, dans son ensemble, a un contenu maximal de ces éléments dans

K, = K, et un contenu minimal accusé dans K,. Le minimum est tout à fait comparable au contenu des mêmes éléments dans les faciès carbonatés à travers le monde.

3. A la lumière de ces observations, on peut suggérer une origine des formations de fer dans les carbonates primaires sédimentaires dont l’accumulation a eu lieu dans les conditions offertes par les mers ouvertes précambriennes en contraste avec les conditions côtières d’accumulation de K et le faciès K, de la plate-forme continentale.

4. La nature spécifiquement précambrienne des dé- pôts K, a été régénérée par l’accumulation de FeCO, (fer ferreux) avec une différenciation de certains micro-éléments typiques du Précambrien.


ARKHANGELSKY, A. D.; KOPCHENOVA, E. V. 1934. O zavisi- mosti khimicheskogo sostava osadochnykh zheleznykh rud ot usloviy ikh olrazovaniya. [Dependence of the chemical composition of sedimentary iron ores on the conditions of their formation]. Bull. Mask. obshch. ispit. prirody, otd. geologii, no. 2, p. 12.

BETEKHTIN, A. G. 1951. Parageneticheskiye sootnosheniya i posledovatel’nost obrazovaniya mineralov [Paragenetic relations and sequence of mineral formation]. Rec. Russ. miner. soc., no. 2.

COCHE, L.; DASTILLON, D.; DEUDON, M.; EMERY, P. 1954. Compléments à l’étude du bassin ferrifère de Lorraine. Le Bassin de Landres-Amermont . Paris, Centre de documen- tation sidérurgique.

FEDORCHENKO, V. S. 1965. Mineralnyy sostav iskhodnykh porod i genezis kraslcovykh i krasko-maritovykh rud krivorozhskogo zhelezorudnogo basseina [Mineral composition of initial rocks andgenesis of krassyk and krassyk-martite ores of the Krivoyrog iron-ore basin]. Vol. 1-2 of dissertation, Krivoyrog.

GLAGOLEV, A. A. 1966. Metamorfizm dokombriyskikh porod KMA. [Metamorphism of Precambrian rocks of the Kursk Magnetic Anomal]. Moscow, Nauka.

GREEN, J. 1953. Geochemical table of the elements for 1953. Bull. Geol. Soc. Amer., vol. 64, no. 9.

HUBER, N. K. 1959. Some aspects of the origin of the Ironwood iron-formation of Michigan and Wisconsin. Econ. Geol.,

IRVING, R. D.; VAN HISE, C. R. 1892. The Penokee iron-bearing series of Michiganand Wisconsin. US. Geol. Surv. Monogr. 19.

vol. 54, p. 82-118.

JAMES, H. L. 1951. Iron-formation and associated rocks in the Iron River district, Michigan. Bull. Geol. Soc. Amer., vol. 62.

KALYAEV, G. I. 1965. Tektonika dokembriya Ukrainskoy zhele- zorudnoy provintsii [Tectonics of the Precambrian of the Ukrainian iron-ore province]. Kiev, Naukova Dumka.

KANTOR, M . I. 1937. Geneziz kerchenskikh zheleznykh rud [Genesis of the Kerch iron ores]. Trudy konjerentsii PO gene- zisu rud zheleza, margantsa i aluminia.

PLAKSENKO, N. A. 1966. Glavneyshiye zakonomernosti zhelezo- rzidnogo osadkonakopleniya v dokembrii [Main regularities of iron-ore sedimentation in the Precambrian]. Voronezhskogo University.

PONOMAREV, M . S. 1964. K voprosu o sootnoshenii ‘zhelezo- rudnoy formacii’ i fillitoarkozovoy tolshchi proterozoya Kri- vorozhskogo basseina. [On the correlation between iron-ore formation and phyllite-arkose Proterozoic series of Krivoyrog basin.] Soviet. Geol., MOSCOW, no. 7.

STRAKHOV, N. M . 1947. Zhelezorudnye facii i ikh analogi v istorii Zemli. [Iron-ore facies and their analogues in the history of the earth]. Trudy inst. geol. nauk A N S S S R , no. 73. (Geological series, no. 22.)

TARKHANOV, A. V. 1969. K voprosu o proiskhozhdenii zhelezistykh kvartsitov [Origin of ferruginous quartzites]. Probbmy obrazovaniya zhelezistykh porod dokembriya [Problems of formation of Precambrian ferruginous rocks]. Kiev, Naukova Dumka.

VAN HISE, C. R.; BAYLEY, W . S. 1897. The Marquette iron- bearing district of Michigan. US. Geol. Surv. Monogr. 28.

39

Jacobsites from the Urandi manganese district, Bahia (Brazil)

E. Ribeiro Filho Instituto de Geociencias, Universidade de São Paulo, São Paulo (Brazil)

Introduction The area where jacobsite associated with manganese minerals occurs is situated in the south-west part of the state of Bahia, Brazil. It is found localized within the crystalline complex that forms part of the major geotectonic unit known as the Espinhaço Geosyncline. The deposits studied are restricted to the southern region of the belt of metamorphic rocks near Urandi (14'49' S, 42O38' W).

The manganese deposits in this region of Brazil, as in most other occurrences, are associated with Precambrian rocks. These form a lower sequence of gneisses, granitic gneisses, schists and amphibolites overlain by an upper sequence of phyllites, green schists, metaconglomerates and quartzites. The manganese deposits are always intercalated between the schists and phyllites.

The age of the rocks in this region is between 463 and 791 m.y., based on K/Ar determinations.

In the manganiferous district of Urandi, there are at least three different types of manganese ore deposits. The first type, which occurs in the Barreiro dos Campos mine, refers to a carbonate protore. The most common mineral of this protore, which could be structurally classified as mangano-dolomite, has a DTA curve very similar to that of kutnahorite. The two other types were both formed by regional metamorphism of syngenetic sediments with primary manganese oxides. The difference between these two types depends on the grade of metamorphism. In the Barnabé mine the manganese assemblage of minerals as well as the associated rocks shows evidence of high-grade metamorphism. In some deposits, such as Pedra Preta, where the associated rocks are green schists and sericito- schists and jacobsite and hausmannite are absent in the ores, the grade of metamorphism is lower.

A common feature of all manganese deposits of Urandi district is the occurrence of ore formed by weathering and supergene enrichment of protore. Thus, the paragenesis of the manganese ore minerals of Urandi district may be exemplified as shown in Table 1.

Brazilian occurrences of jacobsite In Brazil, the first report of the discovery of jacobsite was from the Barnabé mine, which is situated near the town of LicÍnio de Almeida, Bahia (Ribeiro Filho, 1966). With the progress of studies on the manganiferous deposits of this region, it has been possible to identify several jacobsites, whose mineralogical, textural, chemical and etching test behaviour show some variations. However, the mode of association of jacobsite with other manganese minerals is not always identical.

Jacobsite is always a mineral of metamorphic origin, formed in a stratified manganese deposit. It shows a granoblastic texture in polished section, with crystals of variable shape phose proportions range from 0.05 to 3.0 m m . The most common diameters range between 0.2 and 1.0 m m . It is strongly magnetic and under reflected light, exhibits a rose-grey or olive-grey colour; colour is the most useful means of distinguishing jacobsite from magnetite. It is isotropic and inert with all reagents commonly used in etching tests. Nevertheless, some samples, including those found in the Urandi district, give a positive reaction with a 40 per cent aqueous solution of HF. Previously Roy (1959) observed this phenomenon in jacobsite from Andhra Pradesh. H e advanced the hypothesis that this reaction occurred only when jacobsite had a chemical composition approaching the transition to magnetite. However, jacobsites from the Urandi district, despite having a composition very close to magnetite according to Roy's data, give a negative reaction with HCl.

Samples containing jacobsite, collected from various deposits have been examined in polished section by X-rays, and by chemical analysis. The results are given in Tables 2 and 3.

At the Pedra Preta mine, magnetic ore is rare and is situated only at one of theexploratory work faces. It contains crystals up to 3.0 mm in diameter, which show the rose- grey colour in polished section and are isotropic. They give a positive reaction in the etching test using HF, but a negative reaction with all other commonly used reagents,

Unesco, 1973. Genesis of Precambrian iron and manganese deposits. Proc. Kiev Syinp., 1970. (Earth sciences, 9.) 41

E. Ribeiro Filho

TABLE 1. Paragenesis of the manganese-ore minerals of Urandi district

Mine Minerals

~

Period of mineralization

%dim.-Metam. ore Metasomatic ore Supergene ore

Barreiro dos Campos Mangano-dolomite Spessartite Rhodonite Quartz Cryp tomelane Hollandite Pyrolusite Todorokite Ramsdellite

Pedra Preta

Barnabé

Alpha MnO, Mangano-magnetite Cryp tomelane Pyrolusite

Bixbyite Jacobsite Hausmannite Manganite Alpha MnO, Cryp tomelane Todorokite Pyrolusite

-7-7-1 . . .

7-7-7-3- . . . .

?-?

TABLE 2. X-ray powder data for jacobsite. Radiation FeKor, M n filter, Barnabé mine

1 2 3 4 5 6 7 i i -+ i

hkl d&) d&) i d(Å) dí& d(& d&) 1/10 111 4.85 4.84 4.88 4.88 4.88 4.91 4.89 40 220 2.96 2.96 2.98 3 .O0 2.99 3 .O0 3 .o0 40 311 2.53 2.52 2.55 2.55 2.55 2.56 2.56 100 222 2.42 2.41 2.43 2.43 2.44 2.45 2.45 3

2.09 1.71 1.61 1.48 1.41 1.32 1.27 1.26 1.21

2.10 1.71 1.61 1.48 - - 1.28

1.21 -

2.10 1.72 1.62 1.49 1.42 1.33 1.28 1.27 1.22

2.11 1.72 1.62 1.49 1.42

1.28

1.22

-

-

2.11 1.72 1 .63 1.49 1.43 1.33 1.29 1.27 1.22

2.12 1.73 1.63 1.50 1.43

1.29 1.27 1.22

-

400 422 333 440 53 1 620 533 622 444 551 - 1.18 - 1.18 - 642 1.12 1.12 i .13 - 1.13 1.13 553 1 .O9 1 .o9 1.10 1.10 1.10 1.10 800 1 .O4 1 .O5 1 .O5 1 .O5 1 .O5 1 .O6 660 0.98 0.98 - 0.98 0.99 1 .o0 555 0.97 0.97 0.97 0.97 0.97 0.98

-

a,, =8.38 a, = 8.39 a, = 8.43 a, = 8.46 a, = 8.47 a, = 8.49 1. Jacobsite from amphibolite. 2. Jacobsite with exsolved hematite. 3. Jacobsite crystals cut by quartz veins. 4. Granoblastic jacobsite crystals. 5. Granoblastic jacobsite crystals. 6. Granoblastic jacobsite crystals. 7. Granoblastic jacobsite crystals with exsolved hausmannite.

2.12 1.73 1.63 1.50 1.43 1.34 1.29 1.28 1.22 ~~

1.19 3 1.13 5 1.10 40 1 .O6 20 1 .o0 10 0.98 40

ao= 8.49

60 10 60 60 5 3 20 5 10


TABLE 3. X-ray powder data of magnetite-hausmannite series. Radiation FeKcc, M n filter, Urandi district

111 220 311 222 400 422 333 440 531 620 533 622 444 551 642 553 800 660 555

4.83 2.95 2.51 2.40 2 .O9 1.70 1.61 1.47

1.30 1.27

1.21

-

-

- -

1 .O9 1 .O4 - -

4.87 3 .O8 2.54 2.43 2.11 1.72 1.62 1.49 1.42

1.28 1.27 1.22

1.12 1.10 1 .O5 0.99 0.97

-

-

4.87 3 .O9 2.54 2.43 2.11 1.72 1.62 1.49

1.34 1.29 1.27 1.22

1.13 1.10 1 .O5 0.99 0.97

-

-

a,, = 8.37 a,=8.44 a, = 8.46 1. Mangano-magnetite from Pedra Preta manganese mine. 2. Jacobsite from Covão manganese deposit. 3. Vredenburgite from Feixe de Vara manganese deposit. 4. Vredenburgite from Piedade manganese deposit. 5. Jacobsite from Pau de Rego manganese deposit. 6. Vredenburgite from Piedade manganese deposit.

4.87 3 .O0 2.54 2.44 2.11 1.72 1.62 1.49 1.43 1.34 1.29 1.27 1.22 1.18 1.13 1.10 1 .O5 0.99 0.97

a,, = 8.46

4.89 3 .O9 2.55 2.44 2.12 1.73 1.63 1.50 1.43

1.29 1.28 1.22

1.13 1.10 1 .O6 1 .o0 0.98

a, = 8.48

-

-

i

d(A) 4.94 3 .O0 2.55 2.45 2.12 1.73 1.63 1 S O 1.43 1.34 1.29 1.28 1.22 1.18 1.13 1.10 1 .O6 1 .o0 0.98

a,, =8.49

1/10

40 40

1 O0 3 60 10 60 60 5 3 20 5 10 3 5 40 20 10 40

H,O, +H,SO, and concentrated H,SO,. When the crystals are attacked by a 40 per cent aqueous solution of HF, they darken instantly showing clearly the exsolution texture of hematite in mangano-magnetite (widmanstetten texture).

Thus the percentage of Mn,O, and the results of X-ray studies indicate that this mineral has to be at the beginning of the isomorphous series magnetite-hausmannite, as it has an excess of iron relative to manganese.

At the Barnabé mine typical jacobsites, jacobsites with exsolved hematite, and jacobsites with intergrown hausmannite occur. These have to be classified as vredenburgite.

Jacobsite with intergrowth of hematite occurs in ore with a granobiastic texture having crystals of diameter between 0.05 and 0.5 mm. When the material is submitted to an etching test, the jacobsite has a negative reaction with all common reagents and a positive reaction with a 40 per cent aqueous solution of HF. The HF accentuates the exsolution texture of the hematite in jacobsite (Fig. 1).

Thus the results of X-ray analysis, as much as chemical analysis, and behaviour in the etching tests prove that we are dealing with jacobsite. The unit parameter, a,,= 8.397+ 0.003 A, shows that this mineral is almost the end member of the magnetite-hausmannite series. The chemical analysis, giving Mn,0,=17.36 molecular per cent, is also evidence that we are dealing with jacobsite.

In agreement with the phase diagram of Muan and

Somiya (1962), it is not possible to consider the exsolution mineral formed in an excess of manganese to be hausmannite; it may be bixbyite or hematite. Hematite is indicated by the colour of the exsolution lamellae in reflected light, as well as by the negative reaction to all reagents used in the etching test. In our sample which has exsolution hematite, the percentage of Mn,O, rises to 44.70 molecular per cent and the parameter ao = 8.498 f 0.002 A. There is a sample with Mn,0,=34.64 molecular per cent, in which the exsolution hematite develops as rare and sparsely distributed grains. In other samples of magnetic ore from the Barnabémine the proportion of Mn,O, is higher than54mol- ecular perocent, and the unit-cell parameter is greater than 8.46 A, thus defining the mineral as 'vredenburgite (Table 4).

Magnetic ores that contain jacobsite and vredenburgite also occur at other mines located close to the Barnabé mine, such as Feixe de Vara, Covão, Piedade and Pau de Rego Mines. They are ores having a granoblastic texture with crystals of jacobsite or vredenburgite partially substituted by cryptomelane (Fig. 2), todorokite and pyrolusite.

The minerals here defined as vredenburgite are differentiated on the basis of their chemical composition, because they do not contain an exsolution texture formed by lamellae of hausmannite in jacobsite. Crystals of vredenburgite are granular and generally homogeneous. When

43

E. Ribeiro Filho

FIG. 1. Hematite in jacobsite.

FIG. 2. Cryptomelane veins cutting crystals of jacobsite. The cryptomelane was etched with SnCI, (x 50).

44

Jacobcites from the Urandi manganese district, Bahia (Brazil)

FIG. 3. Corroded crystal of jacobcite associated with MnO, minerals (x 63).

FIG. 4. ‘Cranzon ore’ with jacobcite being encrusted by secondary MnO, minerals.

45

E. Ribeiro Filho

TABLE 4. Comparison of data on unit-cell diinensions of jacobsite and vredenburgite ~

ao& Locality References

8.44 Sintético Mason (1943, 1947) 8.50505 0.0005 8.452rL: 0.004 Kiuragi mine, Japan Hirowatari and Myashisa

8.49 Jakobsberg, Sweden 8.42- 8.52 Tirodi and Kodur, India Mukherjee (1959)

Andhra Pradesh, India Roy (1959) 8.51+ 0.01 Negev, Israel Katz (1960) 8.38-8.43 Buryat, U.S.S.R. Rumyantsev (1965)

Pedra Preta mine Ribeiro Filho 8.373 8.385k0.002 Amphibolite, Barnabé mine Ribeiro Filho

Barnabé mine Ribeiro Filho 8.397+ 0.003 8.436- 8.488 Several deposits, Urandi Ribeiro Filho

Several deposits, Urandi Ribeiro Filho

Weabonga, Australia McAndrew (1952)

(1955) Ramdohr (1956)

8.410 - 8.506

8.465 - 8.493

exsolution occurs, it shows as small grains of hausmaimite diffusely exsolved in jacobsite. The jacobsite crystals have been subject to a rate of cooling that was not sufficiently slow to permit the formation of exsolution lamellae. In some samples, the vredenburgite crystals appear to be corroded predominantly along preferential directions that correspond to the now totally altered exsolution lamellae (Figs. 3 and 4).

Conclusions

Some conclusions may be inferred in the light of results obtained during the present investigation concerning jacobsite and the associated manganese minerals,

The values of the unit parameter together with a comparison of the chemical analysis data of natural jacobsite agree with the graph given by Mason (1943) (Fig. 5), who studied the system Fe,O, - Mn,O,. H e concluded that the dimensions of the unit cell increased with increasing quantities of manganese.

Hematite and hausmannite are exsolved in jacobsite based on the ratios of Mn,O, and Fe,O, found in the ores studied here. The ratios agree with the limits postulated from the phase diagrams of Mason (1943), Van Hook and Keith (1958), and Muan and Somiya (1962).

The behaviour of jacobsite when submitted to the etching test shows that there may be a relationship between a positive reaction to HF and the manganese content, as already suggested by Roy (1959). On the other hand, the results with concentrated HC1 disagree, since Roy (1959) observed a positive reaction, whereas the jacobsite studied here gave a negative reaction.

The variation in the unit-cell parameter of jacobsite based on the ratio Mn/Fe is more or less in agreement with the results obtained by Mason (1943). The discrepancies are probably caused by the fact that in natural jacobsite other chemical elements may enter into the crystalline structure replacing iron and/or manganese (Fig. 5).

The mineralogical association found in the ore minerals

9'5 T

FIG. 5. Variation of lattice dimensions in the system Fe,O,- Mn,O, (o denotes findings of the present investigation; rn signifies those of Mason (1943)).

from the Urandi district is similar to that of other localities where metamorphic ores occur. An exception is a single locality where native copper, cuprite and malachite occur together with manganese minerals.

46


The observed textures associated with the identification of jacobsite gave rise to information useful in the interpretation of the paragenesis of the manganese minerals studied.

Where the magnetic mineral, jacobsite, occurs as a major component in lenticular ore bodies, it provides a sufficiently strong anomaly to register during the use of the magnetometric prospecting method. acknowledged.

Acknowledgements The present study was made possible through the generous co-operation of Urandi S.A. Mining Company. Financial support for field and laboratory research on the ore deposits came from F.A.P.E.S.P. (Fundação de Amparo à Pesquisa do Estado de São Paulo). All this assistance is gratefully

Résumé

Jucobsites c h district de manganèse d’Urundi, Bahia, Brésil (E. Ribeiro Filho)

Le district de manganèse d‘Urandi est situé dans la partie sud-ouest de l’État de Bahia au Brésil. I1 est localisé à l’in- térieur du complexe cristallin qui constitue une partie de l’ensemble géotectonique important connu sous le nom géo- synclinal d‘ (( Espinhaço ». Les gisements étudiés sont limités à la partie méridionale de la ceinture de roches métamorphiques près d’Urandi .

Les gisements de manganèse dans cette région du Brésil, comme dans beaucoup d’autres circonstances, sont associés à des roches précambriennes. Ces dernières forment une séquence inférieure de gneiss, de gneiss granitiques, de schistes et d’amphibolites recouverte par une séquence de phyllites, de schistes verts, de métoconglomérats et de quartzites. Les dépôts de manganèse sont toujours inter- calés entre les schistes et les phyllites.

L’âge des roches précambriennes dans cette région se situe entre 463 et 791 millions d’années d‘après les déler- minations potassium-argon.

Dans le district manganifère d’Urandi il y a au moins trois types différents de gisements de minerai de manganèse d’après leur origine. Le premier type, qu’on rencontre dans la mine de Barreiro dos Campos, se rattache à un protore carbonaté. Le minerai le plus commun de ce protore, qui

pourrait être classé d‘après sa structure comme une mangano-dolomite, a une courbe de dosage à I’EDTA très semblable à celle d‘une kutnahorite.

Les deux autres types ont été formés tous les deux par un métamorphisme régional de sédiments syngénétiques avec des oxydes primaires de manganèse. La différence entre ces deux types dépend du degré de métamorphisme. Dans la mine de Barnabé, les assemblages des minéraux manga- niques aussi bien que les roches associées mettent en évi- dence un métamorphisme très avancé. Dans quelques gisements tels que ceux de Pedra Preta, où les roches associées sont des schistes verts et des schistes séricites et OU la jacobsite, I’hausmannite sont absents dans le minerai, le degré de métamorphisme est moins avancé.

U n caractère commun de tous les gisements de man- ganèse du district d’Urandi est l’existence de minerais for- més par l’altération par les agents atmosphériques et l’enrichissement supergène du protore.

Les jacobsites provenant de gisements manganifères métamorphosés ont été identifiés par la diffraction des rayons X et l’examen de surfaces polies. Les pics de la jacobsite varient de 8,38 à 8,49 angströms.

Les résultats obtenus montrent que les méthodes ma- gnétométriques d’exploration peuvent être utilisées comme des outils précieux dans la région étudiée.


DEER, W. A.; HOWIE, M. A.; ZUSSMAN, M. A. 1962. Rock Forming Minerals, vol. 5, p. 77. New York, Wiley. 317 p.

HIROWATARI, F.; MIASHISA, M. 1955. Jacobsite from manganese deposit of Kiuragi mine. Min. Geol. (Soc. Min. Geol. Japan), vol. 5, no. 16, p. 95-107.

KATZ, G. 1960. Jacobsite from the Weger, Israel. Amer. Min., vol. 45, no. 6, p. 734-39.

MCANDREW, J. 1952. The cell edge of jacobsite. Amer. Min., vol. 37, no. 5-6, p. 453-60.

MASON, B. 1943. Mineralogical aspects of the system FeO- Fe,O,-MnO-Mn,O,. Geol. Fören. Stockh. Förh., vol. 65, UO. 2, p. 97-180.

-. 1947. Mineralogical Aspects of the system Fe,O,Mn,O,-

MONTORO, V. 1938. Miscibilita frai sesquiossidi di ferro e di

MUAN, A.; SOMIYA, S. 1962. The system iron oxide-manganese

MUKHERJEE, B. 1959. An X-ray study of manganese minerals.

RAMDOHR, P. 1956. Die Manganerze. XX Int. geol. Congr.,

RIBEIRO FILHO, E. 1966. Jacobsita de Licinio de Almeida,

ZnMn,O,. Amer. Min., vol. 32, no. 7-8, p. 426-41.

manganese. Gazz. chim. ital., vol, 68, p. 728-33.

oxide in air. Amer. J. Sci., vol. 260, no. 3, p. 230-40.

Miner. Mag., vol. 32, no. 247, p. 332-39.

Mexico, vol. 1 (Symposium del manganeso), p. 19-73.

Bahia. Soc. Bras. Geol., vol. 15, no. 2, p. 43-8.

47

E. Ribeiro Filho

RIBEIRO FILHO, E.; ELLERT, N. 1969. Magnetometria relacionada às jazidas de manganês do sudoeste da Bahia. Mineraç. e Metall., vol. 49, no. 289, p. 11-3.

ROY, S. 1959. Variation in the etch behavior of jacobsite with different cell dimensions. Nature, vol. 183, p. 1256-7. - . 1965. Comparative study of the metamorphosed manganese protore of the world. Econ. Geol., vol. 60, no. 6,

RUMYANTESEV, G. S. 1965. Sostav i svoystva vnov’ obnarugenih mineralov ryada magnetit-jacobsit v mestorogenii Magne- titovoye (Buryatskaya ASSR) [Composition and properties

p. 1238-60.

of minerals of the magnetite-jacobsite series redetected in the Magnetitovoye deposit (Buryat A.S.S.R.)]. C.R. Acad. Sci. U.R.S.S., vol. 164, no. 5, p. 1143-6.

STILLWELL, F. L.; EDWARDS, A. B. 1951. Jacobsite from the Tamworth district of N.S.W. Miner. Mag., vol. 29, no. 212,

VAN HOOK, H. J.; KEITH, M. L. 1958. ThesystemFeso,. Amer. Min., vol. 43, no. 1-2, p. 69-83.

YUN, I. 1958. Experimental studies on magnetic and crystallographic characters of Fe-bearing manganese oxides. Mem. Coll. Sci. Kyoto, series B, vol. 25, p. 125-37.

p. 538-41.

Discussion

S. ROY. W h y have you distinguished between clMnO, and cryptomelane in the paragenesis table? I thought they were the same.

E. RIBEIRO FILHO. In some ores I know that both cryp-

impossible to distinguish them. Because of this I preferred to give a general classification as crMnO,.

W . SCARPELLI. Could you give us an idea of the average M n and Fe content of the Urandi ores?

E. RIBEIRA FILHO.

tomelane and hollandite occur together, but in some it was Mn % Fe ..D SiO, % Ba0 % P %

Pedra Preta mine 40-53 1-2.4 0.60-3.0 0.70-14.0 0.05-0.21

S. ROY. Does cryptomelane coexist with jacobsite and hausmannite as a metamorphic mineral? It is rather unexpected -physico-chemically .

E. RIBEIRO FILHO. Cryptomelane, jacobsite and hausman-

(aMnO,) is a secondary mineral which was formed later by weathering. In the slide of the polished section I showed some crystals of jacobsite which were cut by veins of cryptomelane.

\ nite coexist in the same ore. Probably the cryptomelane

S. Roy. ?ave you chemically analysed the jacobsite with a, = 8.46 A which you call vredenbuyite? McAndrew (1952) described jacobsite with a, = 8.52 A which has been chemically analysed and shown to belong to jacobsite field.

E. &BEIRO FILHO. Yes, I have the results of chemical analyses and I called it vredenburgite, based on the definition which was used by Mason.

Barreiro dos Campos mine

Barnabé mine 45-52 1 .O-57 0.25-4.60 0.05-0.13 0.008-0.12

39-45 10.5-16.5 0.40-13.50 - 0.012-0.31

I. P. NOVOKHATSKY. What is your opinion on the presence of copper-rich minerals?

E. RIBEIRO FILHO. In one of the slides I tried to show the relationship between copper minerals andquartz vein. I think that the native copper was formed later than the Mn oxide.

I. P. NOVOKHATSKY. What is the role of silica in the protores?

E. RIBEIRO FILHO. Most of the M n deposits of the Urandi area are very low in SiO,. Only in the Barreiro dos Campos mine was rhodonite found. There are no silicatic protores in the area.

48

Time-distribution and type-distribution of Precambrian iron-formations in Australia

A. F. Trendall Geological Survey of Western Australia, Perth

Introduction Brief reviews of the Precambrian banded iron-formations of Western Australia, with particular reference to their associated mineral deposits, have been given by Miles (1953) and MacLeod (1965). N o complete account of the Precam- brian iron-formations of the Australian continent exists, and this contribution is intended to fill this gap. Little new information is presented in this paper, which is compiled largely from published sources.

The Australian Precambrian iron-formations are here classified into six divisions. The criteria for this classification are geographic separation, age, lithological type and stratigraphic geometry. With minor exceptions, discussed below, all these four criteria are equally applicable to all six iron- formation divisions. After a summary of the six divisions the following characters of each are discussed sequentially: age and regional geological environment; lithology; stratigraphic geometry.

Australian political divisions

The mainland continent of Australia forms a single political unit administered by a federal or commonwealth government. There are six mainland states, namely N e w South Wales, Northern Territory, Queensland, South Australia, Victoria, and Western Australia (Fig. 1). These, with the island State of Tasmania to the south, constitute the Com- monwealth of Australia. This explanation is given to avoid the confusion sometimes caused by the use of indefinite geographic terms, such as ‘southern Australia’, in parallel with precisely defined political terms, such as ‘South Australia’.

The six categories of iron-formation

The six divisions used in the following descriptive summaries are listed below and their regional locations marked with equivalent numbers in Figure 1.

1. Iron-formations of the Yilgarn Block and Pilbara Block of Western Australia.

2. Iron-formations of the Hamersley Group of Western Australia.

3. Iron-formations of the Cleve Metamorphics of South Australia, best known from the Middleback Range.

4. Hematite-rich sediments of the Yampi Sound area of Western Australia.

5. The Roper Bar and Constance Range iron-formations of the Northern Territory and Queensland.

6. The Holowilena Iron-Formation and Braemar Iron-For- mation of South Australia.

These six divisions exclude some minor occurrences which are strongly metamorphosed and of uncertain age and relationships.

Age and regional geological environment

IRON-FORMATIONS OF THE YILGARN BLOCK A N D PILBARA BLOCK

These two ancient stable blocks are the only extensive areas of rocks older than about 2,400 m.y. in the Australian continent. The northern Pilbara Block (Fig. i), with an area of about 47,000 km2, is bounded by the Indian Ocean to the north and by the basal unconformity of the overlying Mount Bruce Supergroup (see below) to the south. Its eastern margin is poorly known. The much larger Yilgarn Block, to the south, has an area of about 603,000 kmz. It is bounded to the west by a major fault, the Darling Fault, which defined by its movement throughout Phane- rozoic time the sharp eastern margin of the neighbouring Perth Basin. To the north-west, as well as along the southern and south-eastern margins, the Yilgarn Block has boundaries against younger metamorphic rocks. The nature of these boundaries is not well known, neither is it clear whether the adjacent younger rocks are re-metamorphosed parts of the block or subsequently generated crust. The north-western boundary of the block is the

Unesco, 1973. Genesis of Precambrian iron und manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 49

A. F. Trendall

BOU@ ...

:::: . . .

.......

B

t

n

2 Q

R n

<1

50


basal unconformity of the overlying younger Precambrian Bangemall Group, while the eastern boundary is the basal unconformity of the Mesozoic rocks of the Officer Basin.

Within both these blocks the geology is closely comparable with that of the Archaean 'nuclei' of other continents. Sinuous belts of tightly folded metasediments and metavolcanic rocks encircle extensive areas of granitic rocks, often with some foliation defining broad domes. At their margins the granites are usually more strongly foliated and, although on a regional scale they appear to be structurally basal to the adjacent sedimentary sequences, in detail their contacts are discordant. Close to the granites the metasediments have undergone thermal metamorphism, although, with the exception of the south-western part of the Yilgarn Block, the general metamorphic grade of the metasediments and metavolcanics is low.

The stratified sedimentary and volcanic rocks, whose curvilinear outcrops are traditionally known as 'greenstone belts', form almost exactly a quarter of the total area of the two blocks. Over much of the eastern part of the Yilgarn Block the belts have a pronounced north-north-westerly trend, but elsewhere their direction is less clearly aligned. Banded iron-formations occur in the sedimentary successions of most of the belts, but details of local stratigraphy are only now being systematically recorded, so that the proportion of iron-formation present is not known, nor its status in the Stratigraphic sequence. However, in the Kur- nalpi area (3Oa-3l0 S; 121" 30'-123"00' E) Williams (1969) has related iron-formation occurrence to a quiet interval between successive depositional cycles, beginning with a thick basic and intermediate volcanic (extrusive and intrusive) succession with some intercalated sediment and ending with intermediate and acid lavas, breccias, agglomerates and tuffs, with interbedded greywacke, shale, siltstone and sandstone. Each depositional cycle has a likely total thickness of about 12,000 m. It is unlikely, in this area, that banded iron-formation makes up as much as 1 per cent of the total volume of the associated depositional pile.

Compston and Arriens (1968) have summarized the available geochronological data for the age of the depositional sequences, and thus of their included iron-formations, in the Pilbara and Yilgarn Blocks. Granite ages vary from about 3,000 m.y. for the Pilbara Block to a range of 2,900 to 2,600 m.y. for the Yilgarn Block. Some acid volcanics in the Pilbara Block succession must be at least as old as 3,000 m.y. and, although the Yilgarn Block evidence is unsatisfactory, some of the eastern volcanics are at least as old as 2,670&30 m.y. Thus, although all the banded iron-formations of this division appear to be older than this latter figure, it is still uncertain what the total age range may be between various occurrences.

IRON-FORMATIONS OF THE HAMERSLEY GROUP

The Hamersley Group is one of three constituent groups of the Mount Bruce Supergroup; conformably below it lies the Fortescue Group, and above it, with some local discon-

tinuity, the Wyloo Group. All three groups were laid down sequentially in an ovoid depositional basin (the Hamersley Basin) about 500 km long and 250 km wide, with a west- north-westerly elongation.

The Hamersley Basin developed by the steady sinking of a presumed southern extension of what is now the Pilbara Block and may have originally extended over the whole of the presently exposed area of this. However, the present outcrop of the remaining part of the Mount Bruce Super- group is bounded (Fig. 1) to the south and east by sediments of the unconformably overlying younger Precambrian Ban- gemall Group, to the north by its eroded outcrop termination exposing the Pilbara Block, and to the west by the overlying Mesozoic sediments of the Phanerozoic Carnarvon Basin. In the northern part of the Mount Bruce Supergroup outcrop both the Fortescue Group and Hamersley Group have a southerly dip of only a few degrees, and are unmetamorphosed. In the central outcrop area folding is open and dips are gentle, but along the southern edge there is strong folding of all three groups.

The banded iron-formations of the Hamersley Basin are virtually confined to the Hamersley Group. They form some 1,145 m, or over 40 per cent, of its total thickness of 2,500 m, in five main stratigraphic units which are set out in another paper in this volume. Even taking the most generous estimate for the original total volume of the Mount Bruce Supergroup, the Hamersley Group iron-formations could not have constituted less than 7 per cent of the total material of the basin.

The beginning of deposition in the Hamersley Basin was at about 2,200 to 2,250 may., an age obtained by Compston and Arriens (1968) from Fortescue Group material. The Woongarra Volcanics, acid lavas of the Hamers- ley Group, gave an age of 2,000 k 100 m.y., and Wyloo Group acid igneous rocks one of 2,020I 165 m.y.

IRON-FORMATIONS OF THE CLEVE METAMORPHICS

The Gawler Block of South Australia is a massif of metamorphic rocks which formed the stable south-eastern margin of the later Precambrian Adelaide Geosyncline (Fig. l, see also Parkin, 1969). The western parts of the block consist largely of granite, gneiss and migmatite. In the east iron-formations and other metasediments, including quartzites, mica schists and amphibolites are present also. Miles (1955) placed the iron-formations and most of the metasediments in the Middleback Group, and most of the granitic and gneissicrocks in an underlying Gneiss Complex.

It now appears established (Parkin, 1969) ñrstly that granite and gneiss also occur stratigraphically above the Middleback Group, and secondly that the iron-formations of the Middleback Group occur at the base of more than 9,000 m of metasediments to which, with the underlying gneisses, the name Cleve Metamorphics is now applied. The stratigraphic section given by Parkin (1969) shows two main bands of iron-formation, the lower about 300 m and the upper about 150 m thick, separated by a thin layer of

51

A. F. Trendall

schist, Some 9,000 m above these, another iron-formation about 50 m thick occurs, so that iron-formation forms about 5 per cent of the sedimentary material above and including the Middleback Group. Parkin’s (1969) map shows the whole sequence tightly folded about axial planes dipping very steeply east, and striking roughly north-south, but swinging north-westerly and south-westerly in the northern and southern parts of their outcrop respectively. Whitten (1966) has correlated these Middleback Group iron- formations with minor occurrences of iron-formations in the northern part of the Gawler Block (including the Wilgena Hill Jaspilite; see below), and in the pre-Adelaide Geosyncline metamorphic rocks on the other (eastern) side of the geosyncline.

Gneissic granulites from the Gneiss Complex of the southern Gawler Block have an age of 1,780-t- 120 m.y., whereas granites farther north which transect the metasediments of the Cleve Metamorphics give ages of 1,550 1: 70 and 1,590 30 m.y. (Compston and Arriens, 1968). Apparently the older (granulite) age sets a minimum depositional age for the the Middleback Group iron-formations.

HEMATITE-RICH SEDIMENTS OF T H E YAMPI SOUND AREA

The Kimberley Basin is a, relatively undisturbed sedimentary basin in the northernmost part of Western Australia, bounded on its south-western and south-eastern sides by metamorphic belts, which it unconformably overlies (Fig. 1; see also Gellatly, Derrick and Plumb, 1968). Three main sedimentary groups, successively the Speewah, Kimberley and Bastion Groups, were laid down in the basin, with a combined thickness of about 3,000-5,000 m . Sandstone is by far the most abundant constituent, with siltstone, conglomerate, volcanic rocks and carbonates also present.

The uppermost unit of the Kimberley Group, the Pentecost Sandstone, is about 1,000 m thick and has very local developments of hematitic conglomerate, sandstone, quartzite and schist (siltstone). In the Yampi Sound area (Reid, 1965) the hematite content reaches ore proportions, and there are also noteworthy enrichments at scattered localities along the eastern basin edge. At Yampi Sound sediments with over 60 per cent of hematite are interbedded cyclically with normal clastic sediments at a range of scales. The total thickness of iron-rich sediment may be as much as 200 m, representing theoretically about 4-7 per cent of the total sedimentary thickness, but the lateral restriction means that much less than 1 per cent of the total basin volume is hematitic.

Dating of the sediments and volcanics of the Kimberley Basin by Bofinger is summarized by Compston and Arriens (1968). Deposition began at about 1,820 m.y. and appears to have been relatively fast, since a high shale gave a date of 1,790+ 60 m.y. The hematite sediments thus have a depositional age of about 1,800 m.y.

ROPER BAR A N D CONSTANCE RANGE IRONSTONES

The name ‘Carpentaria Province’ was applied by McDou- gall et al. (1965) to the area of deposition of a thick sequence of unmetamorphosed Precambrian sedimentary rocks laid down in three contiguous basins in the northern part of the Northern Territory and Queensland. The large central basin is called the McArthur Basin, its smaller south-eastern neighbour the South Nicholson Basin (Fig. l), and an over- lapping northern part the Wessel Basin. In the McArthur Basin four sedimentary suites, with a combined thickness of about 12,000 m, were deposited in succession throughout the basin. The uppermost of these, the Roper Group, varies between 1,800 and 4,500 min thickness, and consists mainly of siltstone and shale with prominent sandstone units.

Near the top of the Roper Group there are thin beds of oolitic ironstone, known formally as the Sherwin Iron- stone Member of the McMinn Formation, of the Maiwok Sub-Group (5A of Fig. 1). The outcrop of the member is shown by Randal (1963), Dum (1963u, 19636) and by Plumb and Paine (1964), who give a maximum thickness of about 20 m-less than 0.5 per cent of the total basin sedimentation thickness.

In the South Nicholson Basin the Roper Group is believed by McDougall et al. (1965) to be represented by the South Nicholson Group, with a thickness of about 5,800 m; it consists mainly of siltstone and sandstone. Within the Mullera Formation of this group (Carter, Brooks and Walker, 1961) there are up to 10 ferruginous beds, each up to 12 m thick, with a total stratigraphic thickness of about 90 m (5B of Fig. 1). Carter and Öpik (1961) show the outcrop of this ironstone. Harms (1965) places it in a separate Train Range Ironstone Formation.

McDougall et al. (1965) argue a minimum depositional age of 1,390420 m.y. for one formation of the Roper Group. The age of the granitic basament of the basin, at 1,800 ? 50 m.y., and intermediate ages upwards which are consistent with the stratigraphic succession, suggest that this age is approximately correct, and that a depositional age for these ironstones of 1,400 m.y. is likely.

HOLOWILENA A N D BRAEMAR IRON-FORMATIONS

The Adelaide Geosynclide (Fig. 1) is a great belt of relatively unaltered Precambrian sediments extending north- westerly across South Australia. A n excellent summary of these rocks has recently been given by Thomson (in Parkin, 1969). The varied sedimentary succession of the geosyncline attains a maximum thickness of about 25,000 m, and has been divided into four major units : the CaUanna Beds, the Burra Group, the Umberatana Group and the Wilpena Group. The sediments are mainly lagoonal and shallow marine, and accumulated slowly on a gently sinking down- warp.

The Umberatana Group, whose thickness varies from about 3,500 to 6,100 m, has glacial sediments in the upper and lower parts, separated by a non-glacial sequence. The

52


lower glacial sediments-the Yudnamutana Sub-Group- consist largely of an immense thickness (over 5,000 m) of massive or stratified tillite and related rocks, but contain, very locally, two unusual iron-formation units.

The Braemar Iron-Formation (Whitten, 1970) lies at the top of this glacial sequence, with a thickness of about 600 m. Most of this thickness consists of shale and tillite, but about one-fifth of the total thickness consists of as many as 30 beds, up to about 12 m thick, of highly ferruginous sediment, with or without an admixture of dropped glacial erratics. The regional outcrop of the Braemar Iron-For- mation is shown by Mirams (1962).

The Holowilena Iron-Formation, or Ironstone, is described by Dalgarno and Johnson (1965) as a glacigene siltstone close to the base of the Yudnamutana Sub-Group. The occurrences lie about 100 km west of those of the Braemar Iron-Formation (Dalgarno and Johnson, 1966). No information on thickness is published.

The Umberatana Group has not been dated, but Compston and Arriens (1968) give an age range of 850- 600 m.y. for the underlying Burra Group, and by correlation with glacial rocks in the Kimberley area of Western Australia these iron-formations would have a depositional age of about 750 m.y.

Lithology

IRON-FORMATIONS O F THE YILGARN BLOCK AND PILBARA BLOCK

In these rocks bands of red, white, brown, yellow, pale green-grey, or clear chert alternate with dark iron-rich bands consisting largely of iron oxide, usually with some silica. The chert bands consist of crystalline quartz of average grain diameter about 20-30 p, although in more strongly metamorphosed areas the quartz mosaic may have a grain diameter of over 1 m m . Most of the brown and yellow chert is due to weathering, but it is possible that both red and white chert persist in depth. The most common iron oxide is magnetite, especially in deep samples, and Miles (1941) has suggested that magnetite is the only primary oxide. The evidence seems inconclusive, however, and hematite and magnetite may both be primary. Rarely does hematite, other than martite, occur together with magnetite. Goethite is always secondary. In areas of low metamorphism these rocks consist almost entirely of two minerals, but silicate and carbonate appear in areas of higher regional metamorphism (Baxter, 1965).

The individual bands vary from less than 1 mm in thickness up to about 15 m m . Although no systematic measurements exist, the median band thickness of both chert and iron-rich material is probably about 5 m m . In most rock there are, over any stratigraphic thickness involving about 20 or more bands, roughly equal thicknesses of each of the two band types. Since the usual Fe content is about 30 per cent this implies an average silica content in the iron-rich bands of about 15 per cent.

IRON-FORMATIONS O F THE HAMERSLEY GROUP

The lithology of these iron-formations has been described in detail by Trendall and Blockley (1969). Like the older iron-formations of the Pilbara and Yilgarn Blocks they are typically banded, with silica-rich chert bands alternating with iron-rich bands, called ‘chert-matrix’ by Trendall and Blockley (1969).

Taking all five units (see above) as a group, their lithology differs from that of the older iron-formation in the following respects: (a) the chert bands tend to be thicker, with a median thickness in the one measured part of 7.9 mm, and a proportion of much thicker bands (12 per cent above 30 mm); (b) most of the chert mesobands have afine internal lamination (microbanding) defined by iron-bearing minerals, so that the cherts have a significant iron content; (c) carbonate minerals, including major ankerite, siderite and highly magnesian ‘siderites’, as well as minor dolomite and calcite, are important constituents of the iron-formation; (d) at many levels there is an internal cyclicity of chert types; (e) ‘primary’ magnetite and hematite are present and commonly Co-exist; (f) sheet silicates (stilpnomelane and minnesotaite) and riebeckite are significant, but erratically distributed, constituents.

Among the five main stratigraphic units of iron-formation in the Hamersley Group there are characteristic and consistent differences in thickness distribution of the banding, in the degree of development of cyclicity, in the amount and distribution of riebeckite, in the nature of the common sheet silicate, and in the distribution of interbedded shale.

IRON-FORMATIONS OF THE CLEVE METAMORPHICS

The iron-formations of the Middleback Group have been described by Miles (1955) and Edwards (1955), who refer to them as banded hematite quartzites. The following summary by Owen and Whitehead (1965) is worth direct quotation: ‘Typically these are layered rocks which at the surface are composed essentially of quartz and iron oxides with, in some localities, relicts of silicates. Silicification of silicate and carbonate minerals is not uncommon. The bands range in width from microscopic to about 1 cm and the iron content of the rock as a whole is commonly between 24 and 34 per cent.

‘Where encountered at depth by diamond drill holes, these rocks have a much more varied mineral composition. Magnetite is the predominant iron oxide, non-aluminous amphiboles such as grunerite, cummingtonite, actinolite and tremolite are common, and carbonate minerals (excluding siderite) occur in some bands, lenses and zones. Accessory minerals include minor but persistent apatite which tends to occur along some bands, and minor pyrite and pyrrhotite which have migrated and recrystallized.’

The metamorphism of these banded iron-formations make their lithological comparison with either of the two preceding categories difficult. Possibly the least altered representative is the Wilgena Hill Jaspilite described by

53

A. F. Trendall

Whitten (1968). This is a finely laminated rock with 25 per cent hematite, 75 per cent quartz and a trace of magnetite. Whitten believes at least some of the hematite to be primary.

HEMATITE-RICH SEDIMENTS OF THE YAMPI SOUND AREA

Canavan and Edwards (1938) have described the lithology and mineralogy of these rocks. The hematite occurs in quartzites, sandstones, conglomerates and shales, all of which are intergradational varieties of similar clastic sediments. The hematite sandstones and quartzites are generally banded in hematite-rich and quartz-rich bands. This banding defines the stratification, and ripple-mark, cross- bedding and small-scale graded bedding are as well developed as in normal sandstones. Unlike any of the iron- formations of the above three categories both quartz and hematite are in rounded and clearly detrital grains, although both have locally recrystallized into an even mosaic and often hematite acts as an interstitial cement to the quartz grains.

In the more pelitic rocks the hematite tends to take the textural position of the sericite in the interstratified iron- poor siltstones. Conglomerate occurs in which rounded quartz cobbles 5-15 cm in diameter lie in a matrix of hematite; more rarely the cobbles also consist of hematite.

In summary these sediments are closely similar to normal clastic sediments, but have hematite locally occupying the textural position of either the clastic grains, or their matrix, or both.

ROPER BAR A N D CONSTANCE RANGE IRONSTONES

The lithology of these oolitic ironstones has been summarized by Canavan (1965) and Harms (1965) respectively, largely from the work of Edwards (in both cases) and of Whitehead (in the second); Canavan and Harms provide references to this previous work. In both places recent weathering affects the rocks for up to 30 m below the surface, and the following descriptions apply only to unoxidized material.

Of the Constance Range rocks Harms (1965) summarizes: <. . . the ironstones consists of oolites of ochreous or finely crystalline hematite, siderite and/or chamosite and silica grains set in a matrix of siderite, hematite, minor microcrystalline quartz, and carbon. The mineral referred to as chamosite has not been deñnitely identified; it resembles chamosite, greenalite, or glauconite, and may include all these minerals. The oolites vary from 0.2 mm to 3 mm in diameter and the successive shells may be composed of different iron minerals. In the groundmass, siderite frequently forms crystals up to several centimetres across, enclosing oolites and quartz grains in a Fontainebleau tex ture’.

In the Roper Bar rocks (Canavan, 1965) ‘unoxidized

ore consists of quartz grains in a siderite matrix with oolites of red ochreous hematite and occasionally of chamosite. . . , The bottom ironstone bed consists of closely packed oolites of red ochreous hematite up to 3 mm in diameter. Detrital quartz grains occur in the matrix and also form the nuclei of some oolites’.

The average iron content in both areas is about 45 per cent.

HOLOWILENA A N D BRAEMAR IRON-FORMATIONS

Bucknell’s (1970) description of these rocks shows them to resemble siltstones in which magnetite euhedra about 20- 70 p in diameter emphasize, by systematic variations of concentration from about 3 to 80 per cent, the lamination of the rocks. Some of the magnetite is usually altered to martite, and this proportion varies from 3 to 100 per cent. Often, some fine-grained flaky hematite is present in addition to martite. Other minerals present to varying extents are chlorite, stilpnomelane, biotite, quartz, feldspar and carbonates. The magnetite euhedra are believed to represent a primary precipitate, and to have been penecontempor- aneously reworked, and locally concentrated in ripple crests.

Stratigraphic geometry

By ‘stratigraphic geometry’ here is meant the ratio of thickness to lateral extent at various scales. The heading is included because the recently demonstrated fine-scale stratigraphic continuity of the Hamersley Group iron-formations (Trendall and Blockley, 1969) has set a standard against which other iron-formations may usefully be compared.

IRON-FORMATIONS OF THE YILGARN BLOCK A N D PILBARA BLOCK

Many iron-formations (or jaspilites) of these blocks are tightly contorted, and reliable thickness estimates are rare. It seems unlikely that any exceeds a true thickness of 65 m, and most are no thicker than about 10 m. The present state of mapping does not permit an accurate estimate of lateral extent. However, recently published aeromagnetic maps indicate strike lengths of up to 75 km. Within this distance it is uncertain whether the anomaly is due to one or several iron-formations, and no detailed interval subdivision of any of these iron-formations has been carried out to allow investigation of small-scale continuity.

IRON-FORMATIONS OF THE HAMERSLEY GROUP

The present area of the Hamersley Group is about 85,000 km2, and within selected levels of the group the likelihood of outcrop-wide lateral continuity of laminae less

54


than 1 mm thick, within thin chert bands of the iron-formation, has been demonstrated (Trendall and Blockley, 1969). The probable original continuity of such laminae over the full basinal area of about 100,000 km2 is argued, and this degree of lateral correlation is believed to be possible at any selected level within the iron-formations of the Hamersley Group; that is, over a total stratigraphic thickness of about 1,000 m within the 2,500 m total thickness of the group.

ROPER BAR AND CONSTANCE RANGE IRONSTONES

Each of these appears to have been deposited initially over a maximum area of about 10,000 km2, although information is vague since there is a tendency to overlook areas of no potential economic interest. In both places the ferruginous beds are discontinuous laterally; no precise details are published.

HOLOWILENA A N D BRAEMAR IRON-FORMATIONS IRON-FORMATIONS OF THE CLEVE METAMORPHICS

Whitten (1966), in his correlation chart, has implied an initial area of deposition of these iron-formations at least equal to that of the Hamersley Group iron-formations. However, the scattered nature of the occurrences and the lack of any detailed stratigraphy to support the correlation makes it necessary to reserve final judgement on this (not improbable) proposition. Certainly, the original extent of the Middleback Group iron-formations is unlikely to have been less than the 10,000 km2 implied by Parkin (1969). Within the iron-formations no record is available of the lateral continuity of internal subdivisions.

HEMATITE-RICH SEDIMENTS OF THE YAMPI SOUND AREA

These are very local, and it seems unlikely that these sandstones are notably ferruginous over areas greater than a few tens of square kilometres. It is not known, within this order of area, how the small-scale stratigraphy varies laterally.

Thomson (in Parkin, 1969) has shown these formations with a depositional area of about 10,000 km2. For the Braemar Iron-Formation Whitten (1970) has emphasized 'the lenticular nature of the formations as a whole and of individual members of it'. At Razorback Ridge, where the formation is best developed, ironstone bands several metres thick pinch out laterally within a kilometre.

Summary and conclusions

Among the Australian Precambrian sedimentary basins which contains iron-rich sediments the following conclusions seem valid: 1. Banded iron-formation in the generally accepted sense

is apparently confined to sediments with a depositional age of 2,000 m.y. or older.

2. In such sediments banded iron-formation is quantitatively more significant at about 2,000 m.y., although it occurs from the earliest record of sedimentation onwards.

3. There is a progressive decrease in the abundance of all types of iron-rich sediment after about 2,000 m.y.

There is little consensus of opinion about the origin of any of these sediments, and this has not been dealt with in this factual summary.

Résumé

Répartition de l'âge et du type des formations pi4xmbriennes de fer en Australie (A. F. Trendall)

Les formations précambriennes du fer du continent aus- tralien peuvent être classées en six groupes, qu'on peut défi- nir aussi bien par leur localisation géographique que par le type lithologique ou la géométrie stratigraphique, ou encore l'âge. Ces groupes, dans l'ordre des âges décroissant, sont les suivants : (a) les formations de fer dans le massif de Yilgarn Block et dans celui de Pilbara, dans l'ouest de l'Australie : ces massifs sont situés le premier entre les mé- ridiens 116" et 123" est et les parallèles 26" et 34" sud et le second entre les méridiens 118" et 121" est et les parallèles 20" et 22" sud ; (b) les formations de fer des monts Hamers- ley dans l'Australie occidentale, qui se situent dans un

bassin formé sur des roches précambriennes plus anciennes situé approximativement entre les méridiens de 116" et 122" est et les parallèles de 21" et 23" sud entre les deux massifs indiqués en (a) ; (c) les formations de fer des gisements métamorphiques de Cleve en Australie méridionale, dont le principal affleurement linéaire est situé dans la chaîne de Middleback à la longitude de 137" est entre les latitudes de 33" et 34" sud ; (d) les sédiments riches en hématites de la région du Yampi Sound dans l'ouest de l'Australie situés approximativement à la longitude de 124" est et à la latitude de 16" sud ; (e) les formations de Roper Bar et de Cons- tance Range du McArthur Basin, situées respectivement à environ 134" de longitude est et 15" de latitude sud et 138" de longitude est et 18" de latitude sud ; (f) les formations de fer d'Holowilena et de Braemar, dans le sud de l'Australie,

55

A. F. Trendall

qui étaient initialement déposées sur une surface s'étendant approximativement entre les méridiens de 138" et 141" est et les parallèles de 31" et 33" sud. Ces six divisions ignorent quelques apparitions de moindre importance qui sont fortement métamorphosées et d'âge et de relation incertains. D u point de vue lithologique, (a), (b) et (c) sont toutes des formations zonées. Entre (a) et (b) les différences lithob- giques sont les suivantes : les carbonates de (a) sont rares, les bancs de silex sont relativement peu épais, la périodicité n'est pas apparente, le quartz et la magnétite sont de loin les éléments dominants ; les carbonates de (b) sont un constituant important, les silex sont d'épaisseur variable, on observe une périodicité complexe, et thématite, stilpnomelane et riebeckite sont des constituants significatifs en plus du quartz, de la magnétite et des carbonates. Les formations de fer zonées de (c) sont pour la plupart trop fortement métamorphosées pour présenter une similitude approchée avec.(a) ou (b), mais eíles semblent se rapprocher de (b). Celles de (d), (e) et (f) manquent toutes de zonalité ; (d) est un sédiment clastique riche en hématites, (e) est un sédiment

pisolitique d'hématite-sidérite chamosite et (f) est une roche d'association glaciaire dans laquelle magnétite à grain fin et hématite auraient pu, être précipitées simultanément. Quant à la géométrie stratigraphique, le seul contraste signi- ficatif est celui qui existe entre les formations relativement peu épaisses et peu étendues latéralement de (d) et la plus grande puissance et l'extension latérale de celles de (b) ; les formations de fer zonées de (c) sont en général trop éner- giquement plissées pour permettre une comparaison stratigraphique, tandis que la stratigraphie des formations de fer non zonées n'est pas aussi significative. L'âge auquel se sont déposées les diverses formations de fer de (a) se situe entre 3 O00 et 2 700 millions d'années, tandis que celui de (b) est d'environ 2 O00 millions d'années. Les formations de (c) furent métamorphosées il y a environ 1 780 millions d'années et peuvent avoir le même âge que celles de (b). Les sédiments de Yampi (d) ont environ 1 810 millions d'années tandis que ceux de (e) ont probablement un peu plus de 1 500 millions d'années. Ceux de (f) remontent au voisinage de 750 millions d'années.


BAXTER, J. L. 1965. Petrology of a banded iron-formation, Koola- nooka Hills, Western Australia. Honours thesis, University of Western Australia, 130 p. (Unpublished.)

BUCKNELL, M . J. 1970. Petrological reports. In: G. F. Whitten, 1970, p. 91-114 (this reference list).

CANAVAN, F. 1965. Iron ore deposits of Roper Bar. In: J. McAndrew (ed.), Geology of Australian ore deposits, p. 212-15. Melbourne, Australasian Institute of Mining and Metallurgy. (8th Commonw. Min. Metall. Congr. Publ. no. 1.)

CANAVAN, F.; EDWARDS, A. B. 1938. The iron ores of Yampi Sound, Western Australia. Proc. Aust. Inst. Min. Engrs., no. 110, p. 59-101.

CARTER, E. K.; BROOKES, J. H.; WALKER, K. R. 1961. The Precambrian mineral belt of northwestern Queensland. Bull. Bur. Min. Resour. Aust., no. 51, 343 p.

CARTER, E. K.; ÖPIK, A. A. 1961. Lawn Hill. 17 p. (Bureau of Mineral Resources 1 : 250,000 explanatory notes series.)

COMPSTON, W.; ARRIENS, P. A. 1968. The Precambrian geochronology of Australia. Canad. J. Earth Sci., vol. 5,

DALGARNO, C. R.; JOHNSON, J. E. 1965. The Holowilena ironstone, a Sturtian glacigene unit. Quart. Notes Geol. Surv. S. Aust., no. 13, p. 2-4. - . 1966. 1 : 250,000 geological atlas series sheet H 54-13. Parachilna, South Australia Geological Survey.

DUNN, P. R. 1963a. Hodgson Downs, N.T. 16 p. (Bureau of Mineral Resources 1 : 250,000 explanatory notes series.) - . 19636. Urapunga, N.T. 17 p. (Bureau of Mineral Resour- ces 1 : 250,000 explanatory notes series.)

EDWARDS, A. B. 1955. Banded hematite quartzites from the Middleback Range, South Australia. In: K. R. Miles, 1955, p. 206-10 (this reference list).

GELLATLY, D . C.; DERRICK, A. M.; PLUMB, K. A. 1968. Pro- terozoic palaeocurrent directions in the Kimberley region,

p. 561-83.

northwestern Australia. Bureau of Mineral Resources Record no. 141, 10 p. (Unpublished.)

HARMS, J. E. 1965. Iron ore deposits of Constance Range. In: J. McAndrew (ed.), Geology of Australian ore deposits, p. 264- 69. Melbourne, Australasian Institute of Mining and Metal- lurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1.)

MACLEOD, W . N. 1965. Banded iron-formations of Western Australia. In: J. McAndrew (ed.), Geology of Australian ore deposits, p. 113-17. Melbourne, Australasian Institute of Mining and Metallurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1.) MCDOUGALL, I.; DUNN, P. R.; COMPSTON, W.; WEBB, A. W.; RICHARDS, J. R.; BOFINGER, W. M . 1965. Isotopic age determinations on Precambrian rocks of the Carpentarian Region, Northern Territory, Australia. J. Geol. Soc. Aust., vol. 12, no. 1, p. 67-90.

MILES, K. R. 1941. Magnetite-hematite relations in the banded iron-formations of Western Australia. Proc. Aust. Inst. Min. Engrs, no. 124, p. 193-201. - . 1953. Banded iron-formations in Western Australia. In: A. B. Edwards (ed.), Geology ofAustralianore deposits, p. 115- 54. 1st ed., Melbourne, Australasian Institute of Mining and Metallurgy (5th Emp. Min. Metall. Congr.). - . 1955. The geology and iron ore resources of the Middle- back Range area. Bull. S. Aust. Geol. Surv., no. 33, 247 p.

MIRAMS, R. C. 1962. The geology of the Manunda Military Sheet. Rep. Invest. Geol. Suvv. S. Aust., no. 19, 39 p.

OWEN, H. B.; WHITEHEAD, Sylvia, 1965. Iron ore deposits of Iron Knob and the Middlehack Ranges. In: J. McAndrew (ed.), Geology of Australian ore deposits, p. 301-11. Melbourne, Aus- tralasian Institute of Mining and Metallurgy (8th Coinmonw. Min. Metall. Congr. Publ. no. 1.)

PARKIN, L. W . (ed.) 1969. Handbook of South Australiungeology. South Australia Geological Survey. 268 p.

PLUMB, K. A.; PAINE, A. G. L. 1964. Mount Young, N.T., 19 p.

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(Bureau of Mineral Resources 1 : 250,000 explanatory notes series .)

RANDAL, M. A. 1963. Katherine, N . T. 26 p. (Bureau of Mineral Resources 1 : 250,000 explanatory notes series.)

REID, I. W. 1965. Iron ore deposits of Yampi Sound. In: J. McAndrew (ed.), Geology ofAustralian ore deposits, p. 126- 31. Melbourne, Australasian Institute of Mining and Metal- lurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1.)

TRENDALL, A. F.; BLOCKLEY, J. G. 1969. The iron-formations of the Hamersley Group, Western Australia, with special reference to the associated crocidolites. Bull. W. Aust. Geol. Surv., no. 119, 353 p.

WHITTEN, G. F. 1966. Suggested correlation of iron ore deposits within South Australia. Quart. Notes Geol. Surv. S. Aust., no. 18, p. 7-11.

-. 1968. The section of iron-formations, Tarcoola District. Quart. Notes Geol. Surv. S. Aust. no. 26, p. 4-7.

-. 1970. The investigation and exploitation of the Razorback Ridge iron deposit. Rept. Invest. S. Aust. Geol. Surv. no. 33, 165 p. WILLIAMS, I. R. 1969. Explanatory notes on the Kurnalpi 1/250,000 geological sheet, Western Australia. Rec. Geol. Surv. W. Aust., no. 197011.

Discussion

R. P. PETROV. What terminology is used for iron-formation rocks by Australian geologists? What term is preferable: chert, jaspilite, itabirite, taconite, iron quartzite?

A. F. TRENDALL. The following definitions are used by the Geological Survey of Western Australia: 1. Iron-formation is a chemical sedimentary rock with a

high iron content; in addition to ‘banded iron-formation’ (defined below) it includes the high-iron shales, slates, carbonate rocks and mixed oolitic rocks frequently associated with Precambrian banded iron-formations, as well as Phanerozoic chamositic sediments.

2. Banded iron-jormation (often shortened to BIF) is a chemical sedimentary rock consisting in its least metamorphosed state of successive thin layers (mesobands) of fine-grained quartz, iron oxides, carbonates and silicates of various proportions. Some nondefining characteristics are: (a) there is usually a wide range between the relatively low iron content of quartz-rich (chert) mesobands and the high iron content of adjacent magnetite-rich or hematite-rich mesobands; (b) the total iron content of samples taken across the banding over widths about 10 times the average mesoband thickness is often in the range of 15 to 35 per cent; (c) most banded iron-formation is Precambrian.

3. Jaspilite is a rock with alternating thin bands of jasper (red chert) and some other material, usually either black iron oxides or chert of another colour.

With these definitions, most jaspilite is a special type of banded iron-formation, and banded iron-formation is itself

a special type of iron-€ormation. Outside the Geological Survey of Western Australia jaspilite is sometimes used in a broader sense, more or less synonymously with banded iron-formation.

The Brazilian term itabirite and the North American term taconite are not used in Australia. Personally I would regard both as special types of banded iron-formation, although some massive granular taconite is only slightly banded. Banded iron-formation (including jaspilite) is known in India as banded hematite quartzite, but many English-speaking geologists, including those in Australia, do not use this term because quartzite usually means a quartz-rich sandstone which has been either thoroughly cemented or metamorphosed.

All persons working on iron-rich sediments should include exact definitions of their terms in their publications.

R. T. BRANDT. Is it justifiable to extend the term iron-formation to include iron-bearing sediments, such as the oolitic iron-stones of Roper Bar and Constance Range? The latter sediments have quite different characteristics and compositions from the banded iron-cherts and jaspilites which are usually understood to be the essential rock types of banded iron-formation.

A. F. TRENDALL. Yes. A general term is needed to cover all iron-rich sedimentary rocks, and ‘iron-formation’ seems to be the best one available. I would be glad to consider any suggested alternative, but have not yet met one.

The origins of the jaspilitic iron ores of Australia

R. T. Brandt Goldsworthy Mining Ltd, Port Hedland, Western Australia

Introduction Iron-ore deposits in Australia are of several different types and can be broadly classified, in order of economic importance, into four categories, as follows. 1. Deposits within or directly associated with Precambrian

banded iron-formations, generally known in Australia as jaspilites. These deposits are here described as jaspilitic iron ores and are the subject of this paper.

2. Sedimentary iron deposits younger than banded iron- formations and of different character.

3. Massive deposits with igneous affiliations, usually known as magmatic-type deposits.

4. Superficial iron-rich laterites, conglomerates, etc. By far the most important are the ores of the first category, of which there are potential reserves amounting to thousands of millions of tons, mostly in Western Australia. The protores or source rocks of these deposits are banded iron- formations or jaspilites of Archaean and Proterozoic age.

The other types of deposit listed above are outside the scope of this paper and will not be discussed here.

Banded iron-formations

The iron-formations are conspicuously banded iron-silica rocks quite similar to the itabirites of Brazil, the taconites of North America and equivalent formations in other parts of the world. Silica is present as finely granular or cryptocrystalline quartz or chert. The iron is in the form of oxide, carbonate, silicate or occasionally sulphide and is invariably oxidized in the zone of weathering to hematite or limonite, giving rise to the characteristic red and white or brown and white striped rock known in Australia as jaspilite.

The character of the surface rock depends on the composition of the original rock from which it is derived. Some hematite jaspilites may be primary oxide facies iron-formations; others are oxidation products of magnetite-carbonate rocks into which they pass below the zone of weathering. Iron silicate rocks, on oxidation, yield limonite and

clay material and are commonly represented at the surface by banded limonitic cherts or shaly rocks. Pyritic iron-formations are characterized by gossanous outcrops. The typical protores of hematite ore deposits are hematite-rich jaspilites, usually with some magnetite more or less oxidized to martite. Drillholes have shown that magnetite is often the dominant iron mineral in the primary zone.

The presence of siderite or iron silicates in the protore appears to be detrimental to the formation of hematite deposits, though much limonite is commonly associated with this type of iron-formation.

The jaspilites of the Archaean and Proterozoic are similar mineralogically but differ greatly in extent, uniformity and geological associations. Those of the Archaean, in the Yilgarn and Pilbara Blocks, are closely associated with greenstone volcanics, which generally underlie them, and are relatively thin and non-persistent units which finger out or change in facies laterally into cherts, fine clastic sediments or tuffaceous rocks. They were evidently deposited in small and relatively short-lived basins in extensive volcanic terrains. The grade of metamorphism is low to medium in general and high in the vicinity of granitic plutons. Iron-ore deposits occur where the jaspilites are unusually thick and particularly where they have been tectonically thickened by intricate internal folding.

The Proterozoic iron-formations are individually far more extensive than those of the Archaean and constitute stratigraphic units of great persistence and uniformity in conformable sequences of sedimentary and volcanic rocks. Thin jaspilite marker beds in the Median Belt can be traced in outcrop for hundreds of miles with little change in composition or thickness, indicating very extensive depositional basins and a uniformity of sedimentary environmental conditions over considerable areas. The grade of metamorphism in the Median Belt is uniformly low, but in the Middleback Ranges is much higher, the associated rocks being metamorphosed to schists, quartzites, amphibolites, etc.

The association of the iron-formations with volcanic rocks is less evident in the Proterozoic than it is in the Archaean, but thin volcanic and tuffaceous horizons occur


R. T. Brandt

throughout the sedimentary successions and volcanic shards and pyroclastic fragments have been identified by La Berge (1966) in the principal iron-formation of the Median Belt. In the present writer’s opinion the field evidence shows that contemporaneous volcanic activity occurred and was probably a necessary accompaniment to the deposition of the banded iron-formations .

In the Middleback Ranges the most productive iron- formation is underlain by dolomitic rocks, a fact which has also been notedin South Africa (Alberts and Ortlepp, 1961). This is thought to be of possible significance as a pointer to the activity of magnesian solutions in the genesis of the ore deposits. The association with dolomite does not exist in the Archaean because of the absence of carbonate rocks in that era, but magnesian solutions may well have been active in association with the basic igneous rocks that are abundant in the Archaean.

Iron-ore deposits

GENERAL

Iron-ore deposits formed in situ in jaspilites by assumed chemical weathering processes are often referred to as Lake Superior type deposits. This rather unsatisfactory term rests on the assumption, unjustified in the present author’s opinion, that all the major deposits of the North American Lake Superior provinceare of the same type and were formed in the sameway. It is difficult to reconcile the great differences in character between, for example, the extensive sheet-like deposits of the Mesabi Range and the deep, discrete, dyke- controlled orebodies of the Gogebic Range, and to regard both as manifestations of the same genetic process seems illogical.

Concentrations of hematite or limonite in jaspilite could theoretically have been formed in a number of different ways which may be broadly summarized as follows: (a) concentration during sedimentation or diagenesis (syngenetic concentration); (b) concentration by Precambrian supergene weathering processes, shortly after or even during the period of sedimentation; (c) concentration by hypogene (metamorphic and/or igneous) processes either before, during or after folding of the beds; (d) concentration by supergene processes after folding, uplift and exposure of the protore beds to erosion.

One of the difficulties in the genetic interpretation of iron-ore deposits arises from the fact that they are generally well exposed and form high ground, any covering rocks which may once have existed having been removed by erosion. In the absence of any direct evidence of the depth at which a deposit was formed, it may be difficult to decide whether its superficial situation and limited depth are original or are due to geological structure and the fortuitous level of current erosion. Cases in point occur in the Middle- back Ranges, where many prominent hill-forming deposits are shallow because they occupy the cores of deeply eroded synclines, the adjacent anticlinal portions having been

removed. In one such deposit a depth zonal arrangement, from siliceous through talcose to carbonate-rich ore, has been noted and roughly parallels the present land surface. This might be interpreted as evidence of a recent supergene origin, but a more likely explanation is that the zoning is a hypogene post-folding phenomenon and its parallelism with the existing topography is merely a reflection of the high resistance of the uppermost zone to erosion. The Thabazimbi deposit in South Africa exhibits similar features (Alberts and Ortlepp, 1961).

Australian jaspilitic iron ores can be classified into three types, which are here given the following descriptive names: (a) crust type; (b) derived type; (c) lode type. The first two types, which account for by far the largest tonnages of ore, were unquestionably formed by post-folding supergene processes. The third type, which constitutes very important high-grade reserves and includes the Middleback Ranges deposits, is of less-certain origin. The present author advocates a hypogene metamorphic origin similar to that proposed by Dorr (1965) for the high grade hematites of Brazil, but syngenetic or pre-folding supergene origins cannot be excluded from consideration.

CRUST TYPE DEPOSITS

Crust type deposits are essentially surficial crusts of hematite and limonite formed in situ on tilted or folded protore beds in the zone of weathering. They are very numerous and widespread in the Median Belt and the Pilbara Block, much less abundant in the Yilgarn Block and rare in the Gawler Block. They range from shallow limonitic cappings of a lateritic nature to relatively thick concentrations of hematite replacing jaspilites in areas of deep weathering. These deposits are similar in all essential respects to the supergene iron ores of Brazil described by Dorr (1964) and are undoubtedly attributable to the same genetic process, namely downward leaching of the protore beds by surface waters during climatic periods of high rainfall and high temperature.

The process, as described by Dorr, involves two kinds of iron concentration, namely residual enrichment by preferential removal of silica in solution, and secondary cementation by solution of iron and its redeposition as limonite in the spaces vacated by the silica. Residual enrichment is the earlier and deeper process which, under favourable conditions, may reach depths approaching 1,000 ft (304 m), though the average is much less. It produces concentrations of porous and disaggregated hematite in which the hematite or martite laminae of the original jaspilite, after removal of the silica, have usually broken up and slumped together. Some secondary limonite is nearly always present as a cementing material or coating on the hematite fragments, but is subordinate in amount to the original hematite.

Secondary cementation is a shallower process by which the hematite fragments near the surface become hydrated and cemented together by secondary limonite, forming a

60

The origins of the aspilitic iron ores of Australia

hard surface limonitic capping of lateritic character known in Brazil as canga. The hard capping is nearly always underlain by a collapsed zone, characterized by slump structures, cave breccias and open caverns due to leaching. The principal concentrations of residual hematite are found below this zone.

A hard canga capping, according to Dorr (1964), is essential for the preservation of the underlying residual hematite, which would otherwise, because of its largely disaggregated nature, be rapidly removed by mechanical erosion. Part of the iron which forms the capping is derived not from the rocks originally above, but from below, by precipitation from solutions drawn upwards by capillarity during dry periods of the year. By this means the capping renews itself continuously, as fast as its upper surface is eroded.

The controlling factors in the formation of crust type deposits are climate and drainage, the chemical and physical nature of the protore beds, and geological structure. The fact that Australian crust type deposits are virtually confined to the tropical northern region which, though semi- arid at present, had a hot and humid climate in the past, indicates that without the necessary climatic conditions these deposits are not formed, even though suitable protore beds exist. The optimum physiographic conditions would be good sub-surface drainage, a low and steadily falling water-table and a slow rate of mechanical erosion.

The nature of the protore beds is also a powerful controlling factor. The most favourable formations for hematite deposits are hematite jaspilites or those in which primary magnetite has been oxidized to martite. The highest-grade deposits consist of hematite residual from the original jaspilite, with a minimum of secondary limonite. There is evidence in some cases that secondary iron has been precipitated not as limonite but as massive, fine-grained colloform hematite, which cements and further enriches the deposit. Carbonate and silicate iron-formations do not normally contain hematite deposits but yield much limonite and commonly carry extensive canga cappings, underlain by thoroughly limonitized zones which are often collapsed and cavernous due to leaching out of silica.

Due to their mode of formation, crust type deposits are generally shallow in relation to their horizontal dimensions and have very irregular basal profiles, due to differences in the depth of iron enrichment from bed to bed. Sedimentary structures are normally preserved in the ore on a large and a small scale. Compositional differences in the protore beds are reflected by differences in the proportions of hematite, limonite, silica, alumina, etc., in the ore. The laminations of the jaspilite are preserved as a plate-like structure, causing the hematite to split easily into thin slabs.

The principal physical controls of ore emplacement are the bedded structure and state of fracturing of the host rocks. Iron enrichment is deepest and most thorough where the jaspilites are well-jointed or shattered in the vicinity of faults. Where the beds are unfractured and only mildly deformed there may be no enrichment at all. Areas of open folding with fairly steep dips and numerous cross fractures

appear to be the most favourable. The deepest hematite concentrations are often found in the cores of plunging synclinal structures, especially where these are floored by relatively impervious shaly limonitic strata. A typical cross- section of such a deposit is shown in Figure 1. The synclines evidently acted as conduits in which downward-moving silica-leaching solutions became channelled. The same feature has been noted in Venezuela (Ruckmick, 1963).

The grades of crust type hematite ores range up to 66-67 per cent Fe. Material below an arbitrary grade of 55 per cent Fe is classed as iron-enriched jaspilite and not as iron ore. The degree of iron enrichment that can be achieved by silica leaching is limited by the presence of small amounts of combined water and other non-leachable impurities, which are enriched at the same time as the iron. Consequently the ore has a somewhat higher content of alumina, phosphorus and combined water than the parent jaspilite.

The Median Belt of Western Australia has enormous reserves, estimated at about 8,000 million tons, of crust type hematits of grade from 55-66 per cent iron, with a phosphorus content between 0.03 and 0.16 per cent. The exceptionally high grade deposits at Mount Tom Price and Mount Whaleback together contain nearly 1,000 million tons of ore of an average grade higher than 64 per cent Fe, including some which assays up to 69 per cent Fe and is very low in phosphorus. It appears possible that these very high assays, which are said to be above the limit attainable by hematites formed by supergene processes (Dorr, 1965), may be from lode type and not crust type ore.

The Median Belt region is an elevated undulating plateau with deeply weathered convexly rounded hill profiles and flat-floored valleys formed by an old erosion surface which is under dissection by present drainages. The deposits are related to this old surface, known as the Hamersley surface, which is believed to have been perfected some time during the Tertiary and was then upwarped and vigorously incised by rejuvenated drainages. The beginnings of the iron-enrichment process could thus date back at least as far as the Mesozoic and possibly even earlier, and it may even be proceeding at a reduced rate at the present day, though current erosion is chiefly engaged in dissecting and de- stroying the previously formed deposits.

The crust type deposits of the Pilbara Block are of similar composition to those of the Median Belt but of somewhat lower average grade and very much smaller, probably because the thin, closely folded and metamorphosed jaspilites of the Archaean did not lend themselves to supergene iron enrichment to the extent that the thicker Proterozoic formations did. The total reserves are probably somewhere between 100 and 200 million tons in numerous small scattered deposits, many of which are too small to be economically workable. The Pilbara Archaean area is relatively flat and low lying and once carried a cover of Mesozoic sediments. Supergene iron enrichment probably dates from the removal of this cover in the Tertiary, though it could have commenced much earlier, in early or pre- Mesozoic times.

61

R. T. Brandt

E[ . . . . . . . . .

MAfNL Y HEMA TI TE ENR/CtfMENT

MAfNL Y LfMONf TE ENRfCHMENT

JASPfL i TE

Scale of Feet (approx 1 O 500 1000 I I I

FIG. 1. Section of typical crust type iron-ore deposit.

DERIVED TYPE DEPOSITS

Derived type deposits are extensively developed in the Median Belt of Western Australia and are numerous, though very much smaller, in the Pilbara Block. The deposits are surficial accumulations of limonite derived from jaspilite but not formed in situ, the iron having migrated a short or a long distance from its source. In Brazil and Venezuela these deposits are classed as cangas (Dorr, 1964; Ruckmick, 1963), no distinction being made between the in situ canga cappings on crust type deposits and the transported cangas which commonly surround them as aprons on hill slopes and in valleys. There is in fact no real distinction, as the one type passes gradationally into the other, but in Australia the enormous tonnages of transported cangas that exist independently of crust type deposits make it necessary to consider them as a separate type of deposit, namely derived type.

Derived type deposits have clearly originated from both mechanical and chemical weathering of jaspilites, whereby detrital accumulations were formed and were subject to supergene leaching of silica, hydration of original hematite and other iron minerals and the deposition of much limonite cementing and replacing the original rock fragments. Their formation has involved the downslope transportation of enormous amounts of iron as detritus, in solution, or both. The deposits range from limonite-cemented screes on slopes, which are sometimes continuous with in situ canga

cappings on hilltops, to thick valley-floor accumulations of massive and pisolitic limonite without detrital material.

The latter variety is very extensively developed in the catchment area of the Robe river, which drains the north- western part of the Median Belt. The pisolitic limonites form extensive mesa deposits 100 ft (30 m) or more in thickness, which are deeply dissected by the present rivers. The upper surfaces of the mesas have been shown to be continuous with the Tertiary Hamersley erosion surface on the hill-forming jaspilites to the south-east (Fig. 2), from which it can be inferred that the limonites were formed during the period of this extinct erosion cycle and are thus essentially contemporaneous with the crust type deposits in the jaspilites.

While there is no dispute as to the origin of the limonite, which was derived by erosion of large areas of jaspilite within the drainage basin of the ancestral Robe river and its tributaries, some controversy exists regarding the manner in which the iron was transported and deposited. In the headwater areas are marginal canga aprons around the flanks of jaspilite hills which consist largely of cemented jaspilite fragments. This variety passes downstream into mesaform pisolitic limonite, which occupies extensive valley floors eroded on softer rocks. Harms and Morgan (1964) regard the mesa deposits as bog iron ores formed by chemical precipitation of iron in large lakes or swamps. MacLeod (1966) interprets them as sheet-like accumulations of jaspi-

62


2 HAM E RSLEY SUR FA

I CRUST TYPE /RON OR€ DEPOSíT 2 IN srru GANGA

4 PrsoL r nc L rMoNrrE 3 DETR/TAL CANGA

FIG. 2. Illustrating relationship of derived type deposits to Hamersley erosion surface.

litic detritus so thoroughly leached by iron-charged waters and replaced by limonite that their original detrital character is lost.

Certain Lower Cretaceous conglomerates in the lower Robe river valley have been shown to be older than the mesaform limonites, which are therefore of late Mesozoic or Tertiary age. It seems reasonable to regard them as essentially cogenetic with the crust type hematite deposits, both being products of the prolonged chemical and mechanical degradation process which culminated in the attainment of maturity of the Tertiary Hamersley erosion surface. The abundance of limonite suggests that the chief source of the iron may have been large areas of carbonate or silicate iron- formation.

Total reserves of mesaform pisolitic limonite in the Median Belt are estimated at about 6,000 million tons. The grade ranges from 40-60 per cent Fe, with considerable amounts in the range 52-58 per cent, with silica and alumina less than 10 per cent, combined water 10-12 per cent and phosphorus less than 0.05 per cent. This material constitutes excellent beneficiating ore which can be upgraded to 63- 65 per cent by removal of water.

Limonitic deposits, including canga cappings, invariably contain small amounts of hematite and maghemite, which is developed mostly at the surface as thin crusts on outcrops and loose fragments. Evidently hydrated iron oxides are unstable in the present semi-arid climate and dehy- drate slowly when exposed to the atmosphere. This possibly represents the beginnings of the process by which canga-type cappings of hard surface hematite were formed under conditions of extreme aridity, as in Mauritania (Gross and Strangway, 1961; Baldwin and Gross, 1967).

LODE TYPE DEPOSITS

Compared with the huge crust and derived type deposits of the Median Belt, lode type deposits are small, the largest known being less than 100 million tons. They are of sporadic occurrence in the Archaean jaspilites of the Yilgarn

and Pilbara Blocks and in the South Australian Middleback Ranges, but have not been recognized in the Median Belt, though there seems to be no compelling reason why they should not occur there.

The Western Australian Archaean deposits appear to be exact parallels of the Brazilian high grade hematites described by Dorr (1965). They are structurally controlled replacement bodies of massive hematite in jaspilite, which commonly have the form of narrow, steeply dipping lenses similar to certain hydrothermal metalliferous lodes. Some are of saddle-reef type in the cores of folds. The lenses are generally concordant with the bedding of the jaspilite and have sharp contacts with it except at the extremities, where the hematite interfingers with jaspilite and passes gradationally into it. The metasomatic origin of the hematite is clearly shown by the preservation in it of the laminated structure of the jaspilite, as bands of slightly different texture or porosity in massive hematite. Some silica expelled during replacement segregates in scattered blebs, pods and random veinlets of crystalline quartz, which are of common occurrence in the marginal transition zones of the hematite bodies. This type of siliceous hematite is entirely different from that formed by incomplete supergene leaching in crust type deposits, the hematite being very pure and the quartz recrystallized and often coarse-grained.

Lode type deposits have no discernible genetic relationship to present or past land surfaces, their exposure at the surface being due to the fortuitous level of current erosion. Small lenses without surface outcrops have been encountered in drilling and it is possibly only a matter of time before large 'blind' orebodies with no surface expression are discovered by underground or geophysical methods.

The control of ore emplacement is both stratigraphic and structural. The typical host rocks are hematite or martite jaspilites with primary iron contents up to 40 per cent Fe or more, especially the variety containing red hematite- stained jasper in place of chert. As in Brazil, the relationship of the deposits to geological structures is their most conspicuous feature. They are localized in zones of relatively intense deformation, where the host formation is

63

R. T. Brandt

r \

1000 O 1000 2000

E d JASPlL/ JE

CHERT

QUARTZ/JE

LODE TYPE ORE

CRUST JYPE ORE

mmmmn BASIC DYKE

3 No. OF OREBODY

FIG. 3. Plan and section of Mount Goldsworthy orebodies.

tectonically thickened by internal crumpling, due to drag on fold limbs or against faults.

The relationship to faulting is very Weil exemplified at Mount Goldsworthy in the Pilbara Block (Fig. 3). The geology of this area has been described in previous publications (Matheson et al., 1965; Brandt, 1964, 1966). The host rock is an unusually thick Archaean iron-formation with a steep northerly dip, which has been further thickened locally by crumpling against an oblique vertical transcurrent fault with a horizontal displacement of about 2 miles (3 km). The beds on the south side of the fault are thrown into a series of drag folds plunging westwards at about 45", and are also crumpled in the opposite sense about vertical axes.

Lode and crust type hematite deposits occur in three stratigraphic positions marked by iron-rich jaspilites within the formation. The lode type bodies are conformable lenses which are localized against the fault and plunge westwards at 45" in conformity with the drag folds. The largest lens, known as No. 1 orebody (Fig. 3), has an abrupt and steep westerly termination, apparently due to a complex Z-shaped vertical fold in the adjacent beds. This structure is also the locus of a narrow vertical dyke which may have played a significant part in the ore localization. No, 3 orebody occurs stratigraphically below No. 1, outcrops against the fault to the east of it, and its westward-plunging extension has been located underground below No. 1. Part of the lode type ore of No. 3 orebody is directly overlain by a body of younger crust type ore which follows the outcrop of the same jaspi-

SECTION A - B

lite protore bed. The structural controls are thus the two sets of crumplings associated with the fault, namely the westward-plunging drag folds and the vertical Z-shaped folds, which between them furnish a favourable structural environment for the deposition of lode type ore.

Another less-thoroughly explored lode type deposit occurs at Shay Gap, 40 miles (64 km) east of Mount Goldsworthy. Here the orebody is a long narrow conformable lens in jaspilite dipping north-eastwards at about 45" (Fig. 4). The ore is localized near the intersection of two

FIG. 4. Plan of the Shay Gap lode orebody. \'

64


structures. One is a set of drag folds plunging eastwards, obliquely to the strike of the beds. The other is a possible fault or sharp flexure trending northwards, against which the thickest part of the orebody terminates abruptly. The same jaspilite formation extends uninterruptedly for many miles in both directions, but most of it is devoid of suitable structures and therefore devoid of ore.

In South Australia, the deposits of the Middleback Ranges occur in jaspilites which are folded about north- south axes. The folds are themselves gently cross-folded and interrupted by faults, dykes and other transgressive structures, so that the plunge of the main folds undulates, changing in direction and steepness from point to point. The terrain is deeply eroded, the main ranges being formed by the deepest parts of the structure, where fold synclines are intersected by plunge synclines. The hematite deposits are generally conformable with the jaspilite bedding and occur on the limbs and in the troughs of synclines. The apparent shallowness of some of the deposits and their occurrence in basin-shaped synclines once led to the idea that they were of supergene origin. On many other grounds, however, it seems reasonably certain that they are lode type deposits deeply eroded, with the synclines preserved and the adjacent anticlinal portions removed.

As in the Pilbara Block, the Middleback deposits are localized by particular structural features. Where the jaspilites are regularly folded and plunge uniformly there is generally no ore. Deposits which occupy fold-limb positions have been shown to be associated with local drag folds on the limbs. The principal synclinal deposits occur in modified basin-like structures where the fold plunge changes or where the folds are interrupted by cross-faults, dykes or sharp flexures. Some deposits are delimited by amphibo- litized basic dykes (Fig. 5). Most of the dykes have been shown to be older than the ore, but a few are younger (Miles, 1954). An interesting feature of the deposit illustrated in Figure 5 is the occurrence locally of highly manganiferous ore where, it is believed, the hematite has replaced part of a bed of manganiferous dolomite (Miles, 1954).

Lode type ores have grades from 69 per cent Fe downwards. In some deposits the bulk of the ore has a grade of 66 per cent Fe or higher, but the silica content of the thinner marginal portions brings the average grade lower. The only important impurity is silica, the alumina and phosphorus contents being significantly less than in crust type ores of similar grade. Limonite is very scarce in the ore and where it is present it can usually be attributed to hydration by subsequent supergene processes. The ores consist of massive hematite with a little martite or magnetite. The texture is usually granular, but quite often micaceous or schistose and occasionally specular. Most varieties have an inherited banded structure, often more or less contorted and cut by veinlets of later coarse-grained hematite. Sometimes the banding imparts a fissility which causes the ore to split easily into thin slabs, a characteristic which is developed much more strongly in crust type ores.

Most lode type hematite is hard, compact and resistant to weathering, but some ore of identical composition is

CROSS -SECTION

EAST WEST

-_- + + ++++' +

LONGITUDINAL SECTION

NORTH -__---- -- SOUTH .--

Scale of Feet 1000 O 1000 2000

I I I I

u JASPIL /TE SCHIST

GNEISS COMPLEX

AMPHIBOLITE (BASIC INTRUSIVES)

/RON ORE

FAULT - FIG. 5. Section of the Iron Monarch orebody, Middleback Ranges (by courtesy of the Australasian Institute of Mining and Metallurgy).

soft, porous and weathers easily. The soft ores are sometimes hardened by weathering at the surface but more often are leached, limonitic and covered by thin canga cappings. This indicates that supergene solutions must have been instrumental in enlarging the pore spaces by dissolving out siliceous impurities and some hematite, but whether this represents the sole origin of the soft ores, or whether the porosity has a more deep-seated origin, is still uncertain.

Lode type deposits are clearly of different origin from crust type deposits and much older. A Precambrian age for the Middleback deposits is indicated by the presence of water-worn hematite pebbles in nearby Cambrian conglomerates (Miles, 1954). Similar pebbles occur in Proterozoic conglomerates in the Pilbara Block. A syngenetic origin for the hematite, as advocated by Baldwin and Gross (1967)

65

R. T. Brandt

for the deep hematite ores of Mauritania, would seem to be invalidated by the obviously metasomatic character of the ore and its localization by tectonic structures. The same would apply to any hypothesis of supergene iron concentration before folding.

The origin favoured by the present author, at least for the Western Australian lode type deposits, is that proposed by Dorr (1965) for the high-grade hematites of Brazil. This postulates that the deposits were formed at elevated temperatures and pressures by hypogene fluids, possibly acti- vated by and partly derived from igneous bodies at depth. The preservation of folded structures in the ore indicates that the process took place mainly after folding of the beds.

The hematite is chemically similar to that of crust type deposits, though with a smaller content of certain impurities, and so in all probability was derived from the same source, namely the protore beds themselves. The deposits characteristically occur where the jaspilite protores have been thickened by folding, a process which must have involved the splitting apart of bedding planes and the creation of voids suitable for penetration by hypogene solutions. It is significant that deposits do not occur where the jaspilites are tectonically thinned or where they are tightly and plastically folded in a manner which would preclude the existence of open spaces between bedding surfaces.

It could be assumed that the solutions travelled mainly along open bedding planes in an up-dip or up-plunge direction, dissolving iron at depth and re-precipitating it as hematite in place of silica at higher levels in zones of reduced pressure. The jaspilites below a deposit should therefore be impoverished in iron and those above it should be enriched in silica. Field evidence in support of this hypothesis is very scanty and inconclusive, but very iron-poor jaspilites have been intersected by deep drillholes below hematite deposits in the Pilbara Block, and one diamond drill core showed jaspilite with the hematite bands leached out. Hy- drothermally leached and silicified jaspilites have been noted in the Middleback Ranges (Miles, 1954). Evidence of hydrothermal argilic wall rock alteration adjacent to hematite orebodies at Mount Goldsworthy has been mentioned previously (Brandt, 1966).

The Middleback Ranges deposits have various features not possessed by those of the Western Australian Archaean and the writer is at a disadvantage in having no personal

knowledge of these deposits. The genetic hypothesis proposed by Miles (1954) is essentially in accord with that outlined above, but other investigators (Catley, 1963; Owen, 1964) have different ideas on which the writer does not feel qualified to comment.

In conclusion the writer wishes to draw attention to the widespread association of deep hematite deposits with basic dykes, not only in Australia but elsewhere, such as the Gogebic Range in the Lake Superior province. It is felt that these dykes could have played a more substantial role than that of fortuitous structural barriers impounding supergene or hypogene solutions. Talc is a minor accessory constituent of many lode type hematites and is especially abundant in the Iron Duke area of the Middleback Ranges where zones of siliceous, talcose and carbonate-rich hematite and magnetite, in descending succession, have been identified. The formation of talc implies the presence of magnesium in the solutions. Owen and Whitehead (1965) envisage magnesium bicarbonate solutions connected with the intrusion of dolerite dykes as the agents by which the iron, silica and carbonates were mobilized and selectively re-precipitated. The dykes, which commonly occur within or marginal to the deposits, could thus be channels by which iron-charged magnesian solutions arose and entered the beds. The common occurrence of dolomite in association with the deposits also suggests the activity of magnesian solutions.

Conclusion

Australian jaspilitic iron ores are in this paper classified on a genetic basis into three types, which have their equivalents under different names in other parts of the world. Though probably far from comprehensive, it is felt that this classification could, with suitable additions, eventually form a basis for one of world-wide application, but this will have to await further knowledge of the genesis of these deposits and the adoption of a world-wide terminology.

Acknowledgement

The author wishes to thank the management of Golds- worthy Mining Ltd for permission to publish this paper.

Résumé

Les origines des minerais de fer jaspilitique d'Australie (R. T. Brandt) fer jaspilitique.

Les plus importants gisements de minerai de fer en Aus- tralie, avec des réserves potentielles s'élevant à des milliards de tonnes, sont ceux qui sont directement associés aux formations zonées précambriennes de fer ou jaspilites et, de

ce fait, ont été groupés sous la désignation de minerais de

Les formations de fer zonées, généralement désignées en Australie sous le nom de jaspilites, sont des membres remarquables des séries archéennes et protérozoïques du sud et de l'ouest de l'Australie. Elles sont en tout point compa- rables aux formations de fer du même âge dans les autres

66


parties du monde. Elles comprennent les hématites-jaspilites zonées et les quartz-magnétites oxydés superficiellement, les roches ferreuses carbonatées et silicatées. Les minerais primitifs (protores) de la plupart des gisements d'hématite sont constitués d'hématites et de jaspilites contenant de la magné- tite. Quelques gisements limonitiques sont associés au faciès du bicarbonate et du silicate de fer.

Les minerais de fer peuvent être classés en trois caté- gories sous les descriptions suivantes : (a) type (( croûte )) ; (b) type (( dérivé )) ; (c) type (( filon D.

Gisements du type (( cuolite D. On les connaît aussi sous le nom de type supergène ou type (( Lac Supérieur n. Ce sont les plus abondants. Ce sont essentiellement des croûtes superficielles d'hématite secondaire et de limonite formées in situ par lixiviation de silices provenant des lits des minerais primitifs au cours des désagrégations et décompositions présentes et passées. Quoique fréquents, ces dépôts sont discontinus et de dimensions, forme et profondeur très variables, car les facteurs déterminants de leur formation sont le climat et le drainage, la structure géologique et la nature physico-chimique des lits des minerais primitifs.

Gisements du type (( dérivé 1). On les connaît aussi sous le nom de gisements de Canga. Ce sont des accumulations sur les versants et dans les vallées de limonite provenant de jaspilite par décomposition et désagrégation mécanique et chimique, dont le résultat est l'élimination de la silice, l'hydratation de l'hématite originale et des autres minéraux ferreux et le transport le long des pentes de la plus grande partie du fer soit en solution, soit comme détritus rocheux ou les deux. Les dépôts vont des talus de limonite agglo- mérée aux accumulations épaisses dans les fonds de vallée

de limonite pisolitique sans matériel détritique. L'assoda- tion étroite entre les gisements type (< croûte )) et ceux du type (( dérivé 1) laisse à penser qu'ils ont pris naissance dans des régimes climatiques identiques ou similaires.

Gisements du type ((filon ». Ce sont des corps de remplacement à structure contrôlée constitués d'hématite massive dans la jaspilite, généralement en accord avec la stratification et n'ayant en apparence aucune relation avec les surfaces de désagrégation et de décomposition présentes ou passées. Certains dépôts se présentent comme des lentilles abruptes ressemblant à des filons métallifères, d'autres sont du type des gîtes en selle au cœur des plissements. Le minerai d'hématite est chimiquement analogue à celui des dépôts du type (( croûte )) à haute teneur, quoique d'une pureté légèrement plus grande, et, par suite, il provient probablement de la même origine, à savoir les minerais primitifs eux-mêmes, par quelque processus de reconcen- tration secondaire du fer. Il paraît vraisemblable d'en attribuer l'origine à la redistribution métamorphique du fer et de la silice pendant le plissement, sans pour autant exclure la possibilité soit d'une origine syngénétique, soit de processus supergenes au cours des cycles de désagrégation et de décompositions précambriens.

Tandis que les gisements du type (( filon >) sont distri- bués sporadiquement dans les jaspilites du sud et de l'ouest de l'Australie, sans relation avec la géographie, ceux des types (( croûte )) ou K dérivé )) ne se trouvent guère que dans la région tropicale du nord, qui a eu un climat chaud et humide dans le passé géologique récent. La dépendance de ces derniers dépôts à l'égard du climat se trouve ainsi clairement démontrée.


ALBERTS, B. C.; ORTLEPP, J. A. L. 1961. Iron ore mining in South Africa. Proc. 7th Commoniv. Mirz. Metall. Congr., South Afiica. Institution of Mining and Metallurgy in South Africa.

BALDWIN, A. B.; GROSS, W. H. 1967. Possible explanations for the localization of residual hematite ore on a Precambrian iron formation. Econ. Geol., vol. 62, no, 1, p. 95.

BRANDT, R. T. 1964. The iron ore deposits of the Mount Golds- worthy area, Port Hedland district, Western Australia. Proc. Aust. Inst. Min. Engrs., no. 211, p. 157. - . 1966. The genesis of the Mount Goldsworthy iron ore deposits of northwest Australia. Econ. Geol., vol. 61, no. 6, p. 999.

CATLEY, D. E. 1963. Some aspects of the genesis of the Iron Duke iron orebody and associated rocks. Proc. Aust. Inst. Min. Engrs., no. 208, p. 81.

DANIELS, J. L. 1966. The Proterozoic geology of the north-west division of Western Australia. Proc. Aiwt. Inst. Min. Engrs., no. 219, p. 17.

DORR, J. Van N. II 1964. Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., vol. 59, no. 7, p. 1203.

--. 1965. Nature and origin of the high grade hematite ores of Minas Gerais, Brazil. Econ. Geol., vol. 60, no. 1, p. 1.

GROSS, W. H.; STRANGWAY, D. W . 1961. Remanent magnetism and the origin of hard hematites in Precambrian banded iron formation. Econ. Geol., vol. 56, no. 8, p. 1345. HARMS, J. E.; MORGAN, B. D. 1964. Pisoliticlimonite deposits in northwest Australia. Proc. Ausi. Inst. Min. Engrs., no. 212, p. 91.

LABERGE, G. L. 1966.Altered pyroclasticrocks iniron-formation in the Hamersley Range, Western Australia. Econ. Geol., vol. 61, no. 1, p. 147.

LIDDY, J. C. 1968. The jaspilite iron ores of Australia. Econ. Geol., vol. 63, no. 7, p. 815.

MCANDREW, J. (ed.) 1965. Geology of Australasian ore deposits. Melbourne, Australasian Institute of Mining and Metallurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1).

MACLEOD, W . N. 1966. The geology and iron deposits of the Hamersley Range area, Western Australia. Bull. geol. Surv. W. Aust., no. 117.

MATHESON, R. S.; ANDREWS, P. B.; BRANDT, R. T.; LIDDI- COAT, W. K. 1965. Iron ore deposits of the Port Hedland district. In: J. McAndrew (ed.), Geology of Australasian ore deposits, p. 132. Melbourne, Australasian Institute of Mining and Metallurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1).

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MILES, K. R. 1954. The geology and iron ore resources Knob and the Middleback Ranges. In: J. McAndrew (ed.), Geology of Australusiun ore deposits, p. 301. Melbourne, Aus- tralasian Institute of Mining and Metallurgy (8th Commonw. Min. Metall. Congr. Publ. no. 1).

RUCKMICK, J. C. 1963. The iron ores of Cerro Bolivar, Vene- zuela. Econ. Geol., vol. 58, no. 2, p. 218.

of the Middleback Range area. Bull. Geol. Surv. S. Aust., no. 33.

OWEN, H. B. 1964. The geology of the Iron Monarch orebody. Proc. Aust. Inst. Min. Engrs, no. 209, p. 43.

OWEN, H. B.; WHITEHEAD, S. 1965. Iron ore deposits of Iron

Discussion

J. Van N. DORR II. In Brazil at least much confusion has been caused by superimposition of supergene process on hypogene lode deposits. Such superposition obscures contact relations and criteria within ore bodies.

R. T. BRANDT. This phenomenon of imposition of supergene processes on the sites of previous hypogene action is very common in Australia too. Pure lode type hematite bodies are not usually much affected by supergene action, but the adjacent jaspilites are, and may be converted to crust type ore enveloping lode type ores in such a way that it is difficult to tell where one begins and the other ends. It seems quite feasible that the hydrothermal wall-rock alteration associated with lode type ore deposits renders the jaspilites more susceptible to supergene action at a later date. Consequently, the two types of ore often occur together in intimate association.

G. A. GROSS. In the case of the Middleback iron deposits what criteria do you suggest which require that these deposits are of hypogene and not of supergene origin? They appear to be remarkably similar to the Shefferville ore in Labrador, Canada, which are undoubtedly of supergene origin.

R. T. BRANDT. Criteria which are strongly suggestive of a hypogene origin, although not conclusive, are the metamorphosed state of the iron ore and associated rocks, the metasomatic replacement character of the ore and its localization by tectonic structures, and the level of erosion, which has exposed the deepest parts of the folded structures

and demonstrates that the ore bodies now exposed were once buried to great depth. It is difficult to visualize a supergene process giving rise to deposits of this kind, which have been buried, metamorphosed and then exposed by deep erosion.

R. P. PETROV. Is it possible to refer to the third type you have mentioned as a lode type? Was the type formed by the filling up of cavities, or by the replacement of the host jaspilites?

R. T. BRANDT. The lode types ores were definitely formed by metasomatic replacements of jaspilites. Mineralogical and textural evidence indicates that this replacement took place at elevated temperatures under hypogene conditions.

J. E. GAIR. Can you comment on whether deposits related to deep dykes occur at intersections of dykes and iron-formation that open upward toward the surface or downward? Clarification of this relationship might favour a supergene origin if the ore is in upward-opening structural traps, or a hypogene origin if the ore is in downward-opening structural traps.

R. T. BRANDT. The dykes I have observed have all been vertical, transverse to the strike of steeply dipping jaspilites, and hematite ore bodies terminate laterally against them. If these dykes did in fact act as structural traps which impounded mineralizing solutions, they do not provide any clues as to whether the solutions travelled upwards, downwards or sideways.

68

Occurrence and origin of the iron ores of India

M. S. Krishnan Hyderabad (India)

Introduction India's iron deposits fall into three major types: (a) Banded hematite-quartzites or jaspilites of the Lake Superior type and banded magnetite quartz deposits which may be considered a variant of this type. This is the most extensively developed and, at present, the only exploited type. (b) Sedi- mentary beds of siderite intercalated with beds of shales. (c) Magmatic segregations, generally lensoid in shape, consisting of titaniferous (and vanadiferous or chromiferous) magnetite associated with intrusive masses of gabbro or olivene bearing rocks. Figure 1 shows the distribution of the important iron deposits in India.

Banded hematite ores

These constitute the major deposits in India and are being actively exploited at several places, both for internal con- sumption and for export. They are derived from and are closely associated with banded hematite-quartzite (BHQ) also known as banded hematite jasper or jaspilite. The major deposits are those of Singhbhum district of Bihar and adjacent parts of Orissa, Drug, Raipur and Bastar districts of Madhya Pradesh, Goa territory and Mysore.

The BHQ in Orissa forms part of the sedimentary succession of the Iron Ore Series described by Jones (1934), consisting of the following formations: upper shales with volcanics; BHQs; lower shales; pink and purple sandstones with some limestones; sandy and conglomerate beds followed by phyllitic shales and tuffs and basic lavas.

The most important group of this type occurs in South Singhbhum, Bonai, Keonjhar and Mayurbhanj districts of Bihar and Orissa (Fig. 2), the last being separated from others by the large exposure of Singhbhum granite. The Iron Ore Series is intruded by Singhbhum granite, which occupies a large area and contains numerous inclusions of the country rocks which have been metamorphosed and partly assimilated. The Iron Ore Series which is intruded by the granite is believed to have an age exceeding 2,100 m.y.

The total thickness of the succession is of the order of 2,000-2,500 m. The rocks have been isoclinally folded with the axes in an NNE.-SSW. direction with steep WNW. dip. The BHQ is exposed as ridges along the anticlinal axes, while the shales occupy the less elevated synclinal portions. The most prominent fold, which is also marked along its crest by massive hematite deposits, extends from Gua (22"13'N., 85'23'E.) to Chendongra (21"43'N., 85'06'E.) over a distance of some 56 km. Two other parallel anticlines also expose numerous deposits. At the northern end, these turn sharply eastward and follow the southern margin of the Singhbhuin thrust zone, but only incomplete sections are seen in this part. The thickness of the BHQ is rather variable, being of the order of 1,000 m in the Korhadi river section in Bonai district, but only about 350 m in the main range on the border of Keonjhar and Singhbhum districts.

The BHQs consist of thin, parallel alternating layers of hematite and jasper or chert. The individual layers vary in thickness from 1 to 20 mm or more, but the average in typical exposures is from about 3 to 5 m m . In addition to the major folds and fractures seen in the whole formation, there are small-scale structures shown by individual layers over small distances. The layers often show intricate folding, contortion and.faulting on a minute scale. These may be attributed to plastic deformation and readjustments during consolidation of the strata and also during the process of enrichment which involved solution and transport of material. Although the layers are generally uniform in thickness, occasional bulging and thinning may be seen. The ore layers consist mainly of hematite which is massive or lamellar. In some bands small octahedral crystals and grains of magnetite are found. Some of this ore is of the character of martite. The silica bands consist of very fine-grained cherty or chalcedonic material whose colour varies from white through lavender and light red to dark red to brown and black. The colour depends upon the amount and the density of distribution of the iron mineral in each band. The iron in the siliceous bands occurs in the form of thin flakes and dust of red hematite or tiny grains and needles of magnetite. Sometimes these may be distributed in lenticular

Unesco, 1973. Genesis of Precantbrian iron und manganese deposirs. Proc. Kiev Symp., 1970. (Earth sciences, 9.) '69

M. S. Krishnan

68" 72O 76O 80' 84' 8$ 92O 9$ 100" 36; I I I I I

3Z0

2E

24

20

I 6"

12"

9O

FIG. 1. Map of India showing the more important iron ore deposits. Bihar and Orissa: 1. Singhbhum; 2. Bonai &. Keonjhar; 3. Palamau. Bombuy: 4. Goa; 5. Ratnagiri. Central Provinces:

or irregular clots. In the deposits at Noamundi, Jones (1934) has recorded the presence of occasional rhombic shapes in silica which he regards as pseudomorphs of silica after siderite. Some of these crystals were actually found to be colourless or light grey siderite. Such pseudomorphs are, however, not common in the oxide deposit.

In a few deposits in Orissa iron carbonate is prominently developed in the banded iron-formation. For instance Acharya et al. (1968) have described the deposit of Kan- dadhar Hill (21°45'N., 85'5'E.) in which the alternating bands are from 1 to 2 mm thick, being composed of tiny euhedral crystals of colourless to pale grey siderite which are from 0.2 to 0.35 mm across and show light brown colouring along their margin. Quantitative measurements under the microscope indicate that the three constituents of the different bands, namely siderite, chalcedony and hematite, are present in roughly equal amounts. Occasion- ally crystals of magnetite are found amidst the siderite or hematite bands. Similar observations have also been made in a few other deposits in Orissa.

6. Chanda-Drug; 7. Bastar. Hyderabad: 8, Adilabad. Madras: 9. Salem; 10. Kurnool. Mysore: 11. Bababudan. Putialu: 12. Narnaul. Himachal: 13. Mandi.

Under the microscope the siliceous bands are seen to consist of very fine-grained quartz showing undulose extinction. The silica bands contain ñakes of red hematite and grains of dark magnetite and sometimes martite. When hematite ñakes are abundant the silica band assumes a red or brown colour. Occasionally small crystals of siderite are also present.

Spencer and Percival (1952) have recorded the occurrence of micro-spherulitic structure in the hematite bands, which they attribute to the shrinkage of original colloidal hydroxide during consolidation of the BHQ. They have also noted that the BHQ is free from clastic sediments, a fact confirmed by other observers. This characteristic indicates that the BHQ was deposited in quiet and fairly deep waters some distance from the shore.

There is lateral variation within the BHQ. It may pass into the solid thick band of hematite without any silica or into shaly-looking hematite or into lenses of fine-grained dusty nearly blue-black crystalline hematite mixed with a certain amount of martite. Partially enriched masses of

70


- 15'

I

22 oc -

2 4 -

FIG. 2. Iron ore deposits of the Singhbhum-Keonjhar-Bonai region, Bihar and Orissa.

BHQ may also be found amidst rich hematitic ore. The analyses given in Table 1 by Percival (1931) indicate such partial enrichment.

TABLE 1. Analyses of hematite jasper (percentages)

I II III

Fe 20.60 30.50 52.30 SiO, 69.00 54.24 22.30 A1203 2.01 1.47 3.56

BASTAR DISTRICT (MADHYA PRADESH)

According to Crookshank (1938) the deposits of the Baila- dila Range occur along two parallel ridges, separated by a valley. This area is located between 18"35' and 18'45'N. and roughly along 81"13'E. (Fig. 1). The two ridges are synclinal, while the valley is along an eroded anticline. The deposits occur in the BHQ and the immediately underlying ferruginous schists of the Bailadila Series which are of Precambrian age. The BHQ has a thickness of 400-500 m.

71

, I M. S. Krishnan

FIG. 3. Deposits of banded magnetite-quartzites of Salem and neighbouring areas, Madras.

72

. ..

i

"'I i


There are fourteen large hematite deposits located on the two ridges which run practically N.-S. The ore at the surface is generally massive and compact, but in some cases a few feet at the top is composed of porous hematite looking like laterite, but really of high grade. The high grade ore at the outcrop yielded on analysis 66-68 per cent Fe, 0.06- 0.12 per cent P and less than 0.05 per cent S.

Similar deposits also occur at and near Rowghat along two ridges which run N.-S. on the east and west of Kolur (19"55'N., 81'8'E.; Fig. 1).

iron enrichment has been only partial. These ores contain only 40-45 per cent Fe and are therefore not worked at present.

Banded magnetite quartzites

Deposits of these rocks are found mainly in southern India, particularly in southern Mysore, in the district of Salam and Tiruchirapalli of Madras and in the Guntur and Nellore district of Andhra.

DRUG DISTRICT (MADHYA PRADESH) MADRAS

Hematite deposits derived from BHQ occur at four or five places in the Rajhara Hill (2Oo34'N., 81"5'E., Fig. 1) and its neighbourhood along a zig-zag ridge several kilometres long. The average ore at and near the surface contains 65-69 per cent Fe, 0.5-2.0 per cent SO,, 0.10-0.20 per cent Mn, 0.05-0.07 per cent P and 0.05 per cent S.

G O A A N D MYSORE

Similar ores associated with BHQ are found in the Dhar- warian formations of the Dharwar district (now in Mysore State) and in the Goa territory (Fig. 1). In the Dharwar district the BHQ forms part of a succession consisting of chlorite and hornblende schists, phyllites, conglomerates and quartzites. The deposits in this area are small and unimportant.

Dharwarian formations similar to those of the Dhar- war district occur also along the border between Ratnagiri and Goa and at three or four places within the Goa territory (Bicholim, Sirigao, Kosti, etc.). In all these deposits the original BHQ has been converted into hematite. The deposits are, however, not as rich as those of Orissa and Madhya Pradesh, but contain 58-62 per cent Fe. Some of the material may be rather flaky and schistose hematite and lateritic in appearance.

Several deposits of hematite associated with BHQ are known in central and northern Mysore, where they form part of the Dharwarian rocks. In some places the banded rocks also contain magnetite. In southern Mysore particularly the BHQ is found to have been converted to magnetite quartzite because of the metamorphism to which it has been subjected.

A synclinorium of Dharwar formations including BHQ phyllites and amphibolites occurs in the Sandur area of Bellary district. The structure is tightly folded along NNE.-SSE. axes, with a steep dip towards the ENE. Two large groups of hematite deposits have been developed along two parallel ridges. The associated phyllites contain numerous secondary manganese ore deposits.

A few deposits also occur in the Chityal Hills (19"5' N., 78"45'E., Fig. 1) and their neighbourhood in Andhra Pradesh. This group contains both hematite and magnetite in varying proportions and the grade of ore is poor as the

Banded magnetite quartzites occur as part of the sedimentary succession which has suffered regional metamorphism. The various associated rock types are chlorite-and micaceous quartz-schists, quartzites, phyllites etc. There are generally some metamorphosed basic igneous rocks in the older part of the succession, these being now seen as amphibole schists or amphibolites with or without garnet. The rocks in this region are folded along NNE.- SSW. direction and two or three parallel bands are found on the flanks of the folds. The structures are cut across by well-marked faults at the southern end along the Attur valley, where a few ore bodies occur in a sheared and disturbed condition. Several hillocks expose magnetite- quartzites very prominently and in some cases they form perpendicular cliffs sometimes 150 m high, as in Godumalai about 16 m east of Salem town (11"38'N., 78"22'E.; Fig. 3). In the deposits of Perumamalai some 10 km east of Godumalai the ore bands are found to be sheared and disturbed. The individual layers are up to 2 c m in thickness. The magnetite shows alteration to niartite along octahedral planes. According to Gokhale et al. (1961) the associated rocks show sedimentary characters such as current bedding. The hill called Kanjamalai (Fig. 3), 7 km long and 4 km wide, is situated 10 kni WSW. of Salem town. It is a basin- shaped structure in which all the exposures are concentric and show dips towards the centre. The rocks exposed in the hill are amphibolites (which are usually garnetiferous) at the base, overlain by magnetite-quartzites, sericite and chlorite- schists, phyllites and talc-schists, amidst which are found two other bands of magnetite-quartzites in the higher part of the succession. This structure is flanked on either side by the Peninsular Gneisses, whose age is probably around 2,500 m.y. The succession in the Kanjamalai appears to be younger than the gneisses. The three parallel bands of magnetite-quartzites in the hill have thicknesses of 30 m, 1O"m and 10 m respectively.

The magnetite-quartzite formation is conspicuously banded like the BHQ, but the individual bands are more irregular because of movements during metamorphism. The magnetite and quartz grains are medium to coarsely crystalline and the bands show appreciable variation in thickness, even within short distances. A few bands show the presence of grunerite, which has apparently been formed

73

M. S. Krichnan

by the reaction of the magnetite with silica. Such occurrences are more common in southern Mysore and Bastar. The magnetite quartzite constitutes the ore which contains 35- 50 per cent Fe (38 per cent average), 41-56 per cent SiO,, 0.2-2.7 per cent Al,O,, 0.1-1.5 per cent lime, 0.1-2.6 per cent M g O , 0.017-0.193 per cent P and negligible S. The rock can be easily crushed and the magnetite concentrated electromagnetically. The degree of fineness of crushing to free all the magnetite is variable, but grinding to minus sixty mesh size is adequate in most cases. The magnetite concentrate is rich in iron, the content ranging usually between 60 and 65 per cent.

MYSORE

A few deposits of magnetite quartzite occur, especially in southern Mysore, of which the more important are the Kudremukh deposits near the western coast and a few in Tumkur and Mysore.

A N D H R A PRADESH

In the Nellore and Guntur districts of Andhra Pradesh there are several hillocks showing bands of magnetite quartzite. They are found between 15'15' and 15O48'N.; 79"27' and 8O003'E. along a narrow arcuate belt. They were originally described by Foote (1879) and form part of the Precambrian succession, consisting of quartzites and micaceous schists surrounded by granitic gneisses, charnockites and amphibole schists. They occur in two groups of several exposures (a) the Gundlakamma group in the north and (b) the Ongole Group in the south near the town of Ongole (15"30'N., 80O3'E.). The southern group appears to be continuous, although covered by soil between the hillocks. It shows two or three bands of magnetite quartzite of which the middle or the upper one may be the thickest. The rocks form an anticlinorium with its axis trending NNE.-SSW. The folds also plunge to the NNE. and have a general steep easterly dip. The northern group of deposits turns sharply towards the west and apparently forms part of the plunging folds. Sastry (1967) has given a general description of individual exposures and their possible structure.

According to Sastry et al. (1968), who described the southern group, the fold axes trend N.-S. with cross folding movement along NNW.-SSE. axes. The final folding was imposed on the strata with NNE.-SSW. axes.

The magnetite quartzites are interbedded with garnetiferous quartzites, hypersthene quartzites and ferruginous schists. The magnetite quartzite passes along the strike into the garnetiferous quartzite or pyroxene-bearing magnetite rocks. The pyroxene is variable in composition and may be ferro-hypersthene salite or jeffersonite. In the hypersthene-bearing rocks the magnetite is occasionally bordered by thin rims of garnet. Green spinel is also seen in a few thin sections.

The mineral assemblage as well as the textural charac-

ters show that the rocks have been subjected to medium to high grade metamorphism resulting in the production of hypersthene and garnet. The average rock contains about 35 per cent Fe, the rest being mainly silica. It is possible to concentrate the magnetite by electromagnetic means, bringing the iron content to about 60 per cent.

BASTAR DISTRICT (MADHYA PRADESH)

The Bailadila series, associated with large hematite deposits, lie to the west of, and apparently superimposed upon, the Bengal series. These consist of amphibolite, quartzite, crystalline limestone and a variety of schists. The youngest formations lying on top of the hills east of Baila- dila are banded magnetite quartzites.

According to Chatterjee (1968) the bands consist of alternating layers of quartz and magnetite. Cumingtonite and an amphibole, identified as ferro-hastingsite, occur amidst the magnetite. Grunerite is fairly common and it may often contain lamellae of martite. The minerals in the bands show marked parallelism to the banding. The amphiboles are elongated and are sometimes poikilitic with inclusions of magnetite. The amphiboles have resulted from the interaction between magnetite and silica. Occasionally almandite and hastingsite are found, indicating that the original rock contained some alumina and calcareous constituents.

Bands of magnetite quartzite are also found intruded by granite in this area. Riebeckite and aegirine are developed in the magnetite quartzite as a result of metasomatism during the intrusion of the granite. It is of interest to note that the magnetite-quartzites do not show any hypersthene in this region, leading to the inference that the temperature of metamorphism was not high enough for its formation.

Sedimentary siderite deposits

Clay ironstone derived from sedimentary siderite beds occurs mainly in the Ranigunj coalfield of West Bengal and the Auranga coalfield of Bihar, some distance to the west. Early surveys showed that the siderite beds occurred in the stratigraphic unit named the 'ironstone shales', which form the middle part of the Lower Gondwana group. The ironstone shales, which are of Middle Permian age, lie above the Barakar series and below the Raniganj series, both of which are coal-bearing. The estimated thickness of the Ironstone shales was approximately 420-450 m. Data from outcrops and from a shaft 15 m deep, showed the ironstone (limonite and goethite) to form thin layers 5-25 cm thick, the proportion of the material in the formation being about 6 per cent. Later examination of this formation by Hughes (1874) and by Walker (1914) established that: the Ironstone shale formation has a total thickness in the Ranigunj coalfield of 300 m and can be traced over a length of 53-55 km; the ironstone bands are intercalated with shales; the layers of ironstone do not persist for long dis-

14


tances either along the strike or dip, but when one layer disappears another begins to appear slightly above or below. From numerous analyses of the material which was used as iron ore in blast furnaces at Kulti in the same coalfield the range and average composition is given in Table 2.

TABLE 2. Analyses of ironstone, Ranigunj coalfield

Constituent Range (per cent) Average (per cent)

Iron 39 .OO-47.70 45.20 Manganese 0.57- 3.62 1.85 Silica 16.00-21.81 18 .O5 Phosphorus 0.23- 1.37 0.72 Moisture 1.00- 5.10 1.77

As these analyses are old they do not show the content of alumina, sulphur and other constituents. These ores were replaced by hematite from the newly discovered banded ferruginous formations in Orissa from about 1914.

At a depth of 15-20 m the limonitic ore changes into siderite which is generally granular in texture. This shows that the material as deposited was siderite and that the carbonate was converted into hydroxide by meteoric waters. The siderite beds are generally fairly pure, but in some cases are mixed with a small amount of detrital clay and sand. The repeated occurrence of layers of siderite in the Ironstone shale formation indicates that deposition was in shallow waters under reducing conditions, but subject to periodic inundations of clastic material. Similar beds of Ironstone shales have also beennoticed to occur in the Aurangacoalfield in southern Bihar, along the same tectonic trough in which the Ranigunj and a few other coalfields are now located.

The total area covered by this formation in the Rani- gunj coalñelds has been estimated as a little over 100 km2.

Titaniferous magnetite

Numerous large lensoid and vein-like bodies of titaniferous magnetite occur near the border of the Singhbhum and Mayurbhanj district of Bihar and Orissa. Though their occurrence has been known since 1908, they were described in some detail much later by Dunn (1937). Several occurrences are found between 22'16' and 22'29'N.; 86"15' and 86"20' E. They are found traversing gabbros and serpentinized olivine bearing rocks which have in some cases been

converted into steatitic material by post-magmatic changes. The augite in the gabbro is usually uralitized. As the area is thickly forested the exposures are not good, but the distribution of magnetite debris gives an idea of the extent of the deposits.

The magnetite ore bodies are composed of magnetite with subordinate hematite. Polished sections of material viewed under the reflecting microscope show the presence of much magnetite enclosing lamellae of ilmenite, coulsonite (Fe-V-oxide), hematite, rutile and goethite and a little apatite. The occurrence of ilmenite along the octahedral planes in the magnetite suggests that the two minerals originally formed solid solutions and that the ilmenite was exsolved on cooling. Coulsonite occurs as minute grains or needles closely associated with the ilmenite. Sometimes magnetite and ilmenite show graphic intergrowths. Dunn and Dey (1937) postulated that part of the hematite may also have been originally present in solid solution in the magnetite.

The magnetite ore contains both titanium and vanadium. The vanadium oxide content generally ranges between 0.6 and 4.84 per cent. The titanium oxide content is much more, ranging from 10 to 25 per cent.

Similar deposits are also present in the Simlipal Hills of the Mayurbhanj district, but they have not been investigated in detail. The deposits near the Singhbhum-Mayur- bhanj border are believed to be large enough to yield a few million tons of magnetite ore. The close association of the magnetite with gabbroid and olivine rocks, and its titanium and vanadium content clearly indicate that it is a product of segregation from ultramafic magmas. These ores are now receiving attention for the extraction of their vanadium content for the manufacture of ferro-vanadium.

MYSORE

Several lensoid bodies of magnetite occur in or along the borders of ultramafic rocks such as gabbros, peridotites, saxonites, etc., in parts of Mysore. They contain approximately 60 per cent Fe, 1 per cent SiO,, less than 2 per cent Alzo, and very low S and P. The titanium oxide content ranges up to a maximum of 12 per cent and there is always a small quantity (less than 3 per cent) of chromic oxide. A little vanadium may also be present. These ore bodies are small in dimension and may have only a limited industrial importance.

Résumé

Manifestations et origine des minerais de fer de l'Inde (M. S. Krishnan)

On ne rencontre pratiquement de gisements de minerai de fer que dans la partie péninsulaire de l'Inde. Ils sont de

quatre types : (a) jaspe-hématite rubané ; (b) quartzite-ma- gnétite rubanée ; c) magnétite titanifère ; d) carbonate de fer sédimentaire décomposé en limonite. Les trois premiers datent du Précambrien tandis que le dernier est de l'âge du Gondwana inférieur (Permien).

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M. S. Krishnan

Les strates de jaspe-hématite rubané ont été déposées dans la succession sédimentaire du système de Dharwar de Mysore, dans les séries de minerai de fer d'Orissa et dans les séries de Bailadila de Madhya Pradesh, c'est-à-dire à une époque comprise entre 2,3 et 2,5 milliards d'années. Dans- tous les cas, la succession contient des laves basiques d'origine sous-marine en même temps que des grès, des argiles schisteuses et des roches ferrugineuses. Les strati- cules dans les roches ferrugineuses sont alternativement de l'hématite et du jaspe, dont l'épaisseur varie de 3 à 5 mètres, mais avec certaines variations. Ces strates et les minerais de fer qui en dérivent sont du type 'Lac Supérieur', avec toutefois cette exception que dans quelques cas les dépôts originaux contenaient de la sidérite au lieu d'hématite et que les silicates de fer sont virtuellement absents. Il semble raisonnable d'assigner une origine volcanique sous-marine à une partie du fer. La formation ferrugineuse a été trans- formée en gisements de riches minerais d'hématite compacte à la surface. Elle peut atteindre une profondeur dépassant 100 mètres. Les lentilles de minerai brun foncé poussiéreux (poussière bleue) qu'on trouve à une certaine profondeur dans la plupart des gisements sont attribués à la cristallisation de l'hydroxyde ferrique original en hématite finement cristallisée pendant une période de lixiviation des dépôts par les eaux météoriques. Ce minerai poussiéreux est cons-

titué presque uniquement d'hématite pure avec, dans certains cas, des traînées de kaolin blanc.

Le jaspe-hématite rubané a été transformé en quartzite- magnétite rubanée dans quelques-uns des dépôts de la partie méridionale de Mysore et dans les régions de Madras et d'Andhra où ils ont subi un métamorphisme modéré. Ce sont des couches à cristaux grossiers, mais en tout cas très semblables aux couches de jaspe-hématite.

Dans le Singhbhum oriental (Bihar), dans la partie sud de Mysore, on rencontre des lentilles de magnétite titanifère associées avec des roches ultramafiques telles que des py- roxénites et des gabbros. Ces lentilles sont apparemment des ségrégations magmatiques. La magnétite contient jusqu'à 15 % d'oxyde de titane ; elle contient aussi jusqu'à 2 % d'oxyde de vanadium dans le Singhbhum et une quantité équivalente d'oxyde de chrome dans le Mysore. Ces minerais n'ont pas encore été exploités.

On trouve des minerais de sidérite sédimentaire au Bengale-Occidental, dans un horizon stratigraphique qui se situe entre les formations de Barakar et de Ranigunj avec présence de charbon. L'épaisseur de cette formation est de l'ordre de 300 mètres. Elle consiste en de nombreux rubans minces de minerais sidéritiques entremêlés à l'argile schisteuse, la straticule de sidérite totalisant le dixième de l'épaisseur totale.


ACHARYA, S.; AHMED, S. I. S.; SARANGI, K. 1968, Carbonate facies of iron-formation in Kandadhar Hills, Orissa. Bull. Geocliem. Soc. India, vol. 3, no. 2.

CHATTERJEE, A. 1968. Petromineralogy and stability relations of metamorphosed iron-formations and amphibolites of an area east of Bailadila Range, Bastar district, M.P. Quart. J. Geol. Soc. India, vol. 40, p. 257-74.

CROOKSHANK, H. 1938. The iron ores of Bailadila Range, Bastar State. Trans. Min. Geol. Inst. India, vol. 34, p. 255-82.

DUNN, J. A. 1937. Mineral deposits of Eastern Singhbhum. M e m . Geol. Surv. India, vol. 61, no. 1, p. 214-23.

DUNN, J. A.; DEY, A. K. 1937. Vanadium bearing titaniferous iron ores of Singhbhum and Mayurbhanj. Trans. Min. Geol. Inst. Indfa, vol. 31, no. 3, p. 117-83.

FOOTE, R. B. 1879. On the geological structure of the eastern coast from latitude 15" to Masulipatam. Mem. Geol. Surv. India, vol. 16, no. 1.

GOKHALE, K. V. G. K.; BAGCHI, T. C. 1961. Preliminary investigation of banded iron ore formations of Perumalai Hills Salem district, Madras. Quart. J. Geol. Soc. India, vol. 33, no. 2, p. 49-53.

HOLLAND, T. H. 1892. Preliminary report on the iron ores and iron industries of the Salem district. Rec. Geol. Surv. India,

'

vol. 25, p. 136-59.

__ . 1893. The iron ore resources and iron industries of the Southern districts of Madras Presidency. Imp. Jnst. Handbook of Commercial Products (Calcutta), no. 8, 24 p.

HUGHES, T. W. H. 1874. Notes on the raw inaterials on iron smelting in the Ranigunj coalfield. Rec. geol. Surv. India,

JONES, H. C. 1934. Iron ore deposits of Bihar and Orissa. Mem. geol. Surv. India, vol. 63, no. 2.

KRISHNAN, M. S. 1954. Iron ores, iron and steel, Bull. geol. Surv. India, vol. 9.

PERCIVAL, F. G. 1931. The iron-ores of Noamundi. Trans. Min. geol. Inst. India, vol. 26, p. 169-271.

SASTRY, A. V. R.; VAIDWYANADHAN, R. 1968. Structure and petrography of the quartz-magnetite and associated rocks of Vemparla area, Nellore district. J. geol. Soc. India. vol. 9, no. 1, p. 49-51.

SASTRY, T. H. 1967. Some structural features of Ongole bandde magnetite quartzites. J. Indian Geosci. Ass., vol. 7, p. 67-84.

SPENCER, E.; PERCIVAL, F. G. 1952. The structure and origin of the banded hematite-jaspers of Singhbhum, India. Econ. Geol., vol. 41, no. 4, p. 365-83.

WALKER, H. 1914. Note on the Geological Survey of Raniganj coalfield. Trans. Min. Geol. Inst. India, vol. 7, p. 226-79.

vol. 7, p. 20-3, 122-4.

76

Precambrian iron ores of sedimentary origin in Sweden

R. Frietsch Geological Survey of Sweden, Stockholm

Introduction In Sweden Precambrian iron ores with sedimentary features are encountered in two separate geographical regions, namely central and northern Sweden. Only the ores of central Sweden have been mined and, due to their low content of phosphorus and sulphur, these were the main source of iron in Sweden until the end of the last century, when the basic steel-making processes made it possible to utilize the apatite-rich ores of the country.

Non-apatitic iron ores of central Sweden

The following presentation of the non-apatitic iron ores of central Sweden is mainly based on the papers of Geijer and Magnusson (Geijer and Magnusson 1944, 19520, 1952ó; Magnusson 1953,1960). These ores are known in a very great number of deposits which occur in a broad, semi-circular zone west of Stockholm. The ores belong to a metamorphic volcanic-sedimentary complex of Svecofennian age, which in this part of Sweden is the oldest unit of the Precambrian. The complex begins with a sequence of acid volcanics with intercalations of limestone-dolomite and clastic sediments. The iron ores lie in these rocks. This volcanic sequence is followed by an upper section mainly consisting of detrital sediments.

The volcanics are divided into a lower part with an extremely sodic, quartz-keratophyric composition and an upper part of predominantly potassic, rhyolitic composition. The volcanics are called hälleflintas, leptites or leptite- gneisses according to their grade of inetamorphism. Best preserved are the hüllefliirztas, which are only found in restricted areas,

The supracrustal rocks were folded in connexioii with the intrusion of the oldest group of Svecofennian granites. These form concordant intrusions, consisting of a differentiated series the first members of which are gabbros and

diorites. The volcanic-sedimentary complex was metamorphosed, the hülleflintas being altered to leptites. During this epoch the rocks were subject to widespread metasomatic alterations. As the main element added through this action was magnesium, the process has been called ‘magnesia- metasomatism’. The leptites have, by this process, been changed into rocks characterized by minerals rich in magnesium and ferrous iron (i.e. cordierite and anthophyllite).

After a non-orogenic period marked by the intrusion of greenstone dykes there occurred a migmatization of the supracrustal rocks in connexion with the intrusion of the late Svecofennian granite group. The granites cross-cut the structures and are accompanied by large amounts of pegmatite. The age of the pegmatites, which represent the last phase of the Svecofennian orogeny, is about 1,800 m.y. (Welin and Blomqvist, 1964).

The non-apatitic iron ores of central Sweden all occur in the lower, volcanic section of the volcanic-sedimentary complex and are divided into two main groups, namely quartz-banded ores and skarn- and limestone ores. The ores of the latter group are further divided into a non- manganiferous type and a manganiferous type, the limit put at 1 per cent Mn. This manganese limit has not only a chemical-technical meaning, it also divides the ores stratigraphically: the manganiferous type occurs in the potassic leptites and the non-manganiferous one is in most cases found in the sodic volcanics.

QUARTZ-BANDED ORES

These are characterized by a regular alternation of thin layers of iron oxide and quartz. The iron oxides are hematite or magnetite or both. Magnetite which has usually been formed from hematite indicates a higher degree of metamorphism. The content of iron varies between 30 and 50 per cent; phosphorus (as apatite) is in most cases very low and manganese is usually lower than 0.1 per cent.

Besides the iron oxides and quartz, these ores contain

Unesco, 1973. Genesis of Precambrian iron and manganese deposits. Proc. Kiev Synip., 1970. (Earth sciences, 9.) 77

R. Frietsch

small amounts of gangue forming layers or irregular aggregates. Where the ore is associated with carbonate rocks the later metamorphism has given rise to skarn silicates such as andradite, diopside and tremolite. Where leptite material has been present in the ore, epidote and hornblende may belong to the skarn association.

According to Geijer and Magnusson (19526) the quartz-banded ores of central Sweden were originally deposited as chemical sediments and the source of the iron and silica must be sought in volcanic emanations. The original iron mineral is believed to have been hematite or, in part, limonite, the well-preserved banding and the purity of the bands excluding minerals like siderite or greenalite. The skarn silicates mentioned above are formed through internal reactions between iron, silica and carbonate material and due to higher temperatures caused by the intrusion of the older granites. In connexion with the magnesia-metasomatic alterations there has further been a rather intensive formation of skarn, which is richer in magnesium than that mentioned above.

SKARN A N D LIMESTONE ORES POOR IN M A N G A N E S E

These form a rather inhomogeneous group low both in phosphorus and sulphur. The ore mineral is magnetite which is more or less intimately associated with skarn minerals or carbonates. The distribution of magnetite and skarn minerals is usually very irregular, but in some deposits a fairly regular stratification of ore and skarn occurs. The skarn is of two different types, rich either in Ca or in Mg. To the Ca-rich skarn type belong andradite, diopside- hedenbergite and actinolite, which are always associated with limestones or dolomites. The Mg-rich skarn type is characterized by anthophyllite-gedrite, cummingtonite, talc, forsterite, humite minerals and serpentine. The Mg-rich skarns are younger than the Ca-rich skarns being alteration products of the Ca-rich skarns formed in connexion with the magnesia-metasomatism .

LIMESTONE ORES

The limestone ores differ from the skarn ores by having little or no skarn. They consist of magnetite and limestone or dolomite. This type of ore shows a fairly distinct stratification and is thus considered to be of sedimentary origin.

MANGANIFEROUS SKARN ORES

These are mostly developed as limestones ores. They are low in phosphorus. The content of sulphur is in several deposits higher than 0.2 per cent. The ore mineral is magnetite. The main skarn minerals are spessartite, dannemorite, knebelite or manganiferous fayalite. A pronounced strati-

fication is a rather common feature, and these ores are also considered to have been originally sedimentary. The skarn is a product of internal reactions, as in the quartz-banded ores.

The origin of the skarn iron ores poor in manganese has been the subject of much discussion. According to Geijer and Magnusson (1952a, 19526), the skarn ores might be either sedimentary deposits later affected by a regional metamorphism or true pyrometasomatic deposits. In the first case the skarn originated through internal reactions between the present iron, silica and carbonate material. In the second case the skarn is a result of the addition of iron, magnesia and silica by magnesia-metasomatism to pre-existing sedimentary iron ores or limestones-dolomites. For some of the deposits with a Mg-rich skarn there is strong evidence for an origin by replacement, as they contain E- or F-bearing minerals (ludwigite, fluoborite, humite) typical of contact metasomatic deposits. Regarding the origin of the skarn ores poor in manganese, Geijer and Magnusson (19526) stated that ‘most of the economically important deposits belong to the originally sedimentary ones, but are more or less intensely “worked over” and rearranged in connexion with the intrusion of the first group of Archaean granites, possibly with some additions of iron’. Later Magnusson (1953, 1960) modified this view and indicated that all skarn ores might have been sediments from the beginning. This opinion is mainly based on the fact that there are transitions between the quartz-banded ores and the skarn ores and between the skarn ores and the limestone ores. Geijer (1959), however, defended the existence of the ‘primary skarn’ ores by stressing, among other things, the importance of the borate minerals.

Non-apatitic iron ores of northern Sweden

The non-apatitic iron ores in northern Sweden occur in the Norrbotten county, and most of them lie within a wide zone extending roughly E.-W. on both sides of Kiruna. Scat- tered occurrences are found to the south of this zone. The ores occur in supracrustal rocks of Precambrian age. On the regional map of Norrbotten (Qdman, 1957) the Precam- brian was divided into an older, Svecofennian (Svionian) cycle and a younger, Karelian cycle. In the iron ore zone detailed mapping during the last decade (Offerberg, 1967; Padget, 1970) has required abandonning the subdivision into Svecofennian and Karelian cycles as outlined by Ödman .

Among the supracrustals four different groups can be discerned, but no major unconformity has been observed between them. A tentative stratigraphic scheme for the northern part of Norrbotten is given in Table 1.

The oldest supracrustal rocks belong to a greenstone group which is mainly built up of greenstones and porphyrites. The greenstones are mainly spilitic, often pillow- bearing effusives of basaltic composition. In the greenstones

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TABLE 1. Stratigraphic scheme for the northern part of Norrbotten county

Age (Million years) Iron ores

1,540 Granite group (granites (‘Lina granite’) with pegmatite and aplite)

sandstones with conglomer- Quartzite group (quartzitic

1,605-l,635 ate and phyllite) Apatite iron ore (Kiruna type) Porphyry group (acid and

intermediate volcanics with intercalations of basic volcanics)

1,880 Granodiorite group (granodiorites with diorite and gabbro)

(mica-schists with intercalations of biotite-rich quartzites, and conglomerates)

Greenstone group (basic vol- Skarn iron ores canics with intercalations of and quartz- detrital and chemical sedi- banded iron ments) ores

Schist-conglomerate group

2,800 Granite north of Kiruna .

occur, mainly in the stratigraphically higher parts, intercalations of tuffs, tuffites, phyllites, graphite-bearing schists, limestones-dolomites, mark and cherts. The porphyrites, which are most widespread near and west of Kiruna, have a basaltic or andesitic composition. To the north of Kiruna the greenstones are underlain by a granite which forms the basement of the volcanics. The granite has an age of about 2,800 m.y. (unpublished radiometric determination by Kouvo).

The greenstone group is overlain by mica-schists and conglomerates of moderate thickness. This schist-conglomerate group has a rather restricted extension, but is a distinct marker horizon. It is succeeded by a porphyry group which, in the iron ore-bearing zone, occurs mainly in the central and western parts. This group is built up of predominantly sodic volcanics of rhyolitic or keratophyric composition. Age determinations show that the volcanics at Kiruna and westwards have an age a little over 1,600 m.y. (Welin, 1970).

The porphyry group is overlain by a quartzite group of restricted extent. The quartzites contain intercalations of conglomerate and, occasionally, phyllite.

In the iron ore-bearing zone two groups of intrusives can be discerned besides the ‘basement’ granite north of Kiruna. The older intrusive group, formerly called the Haparanda granite series by Ödman (1957), is a differentiated series with gabbro, diorite and granodiorite, of which the last-mentioned is the most wide-spread. The rocks of the granodiorite group form concordant intrusions in the greenstone group and the schist-conglomerate group and intersect them. The geological relationship to the porphyry group is not known. Age determinations made on gabbros and

granodiorites in the southern part of the Norrbotten county, show an age of 1,880 m.y. (Welin, 1970). This means that the granodiorite group is older than the porphyry group and the quartzite group.

The younger intrusive group which cuts all supracrustal rocks and the rocks of the granodiorite group, is built up of granites, Radiometric age determinations of the Lina granite, which is the most widespread one of the somewhat different granites that belong to this group, has given an age of 1,540 m.y. (Welin, 1970).

Two groups of non-apatitic iron ores can be discerned in Norrbotten, namely the quartz-banded iron ores and the skarn iron ores.

Q U A R T Z - B A N D E D O R E S

These are quartzites in which magnetite and skarn minerals occur in a more or less banded fashion. The quartzites are locally rather high in iron, but the average grade is mostly low and seldom exceeds 20 per cent. The most common skarn minerals are cummingtonite-grünerite, clinoenstatite- hypersthene, hornblende and almandite. The magnetite and the skarn minerals are often accompanied by small amounts of pyrite and pyrrhotite. The content of phosphorus is less than 0.1 per cent. Manganese seems to be low in most deposits.

The quartz-banded ores occur in connexion with detrital or chemical sediments in greenstones of the greenstone group, usually in the stratigraphically higher part of it. Partly the ores are found very near the bottom of the schist-conglomerate group. The ores, which form layer-like bodies concordant with the strike of the host rock, are often associated with limestones or dolomites or occur in the same stratigraphic position as these rocks,

The quartz-banded iron ores of northern Sweden are considered to have been originally (chemical) sediments. Geijer’s (1925) observation that small, rounded aggregates pigmented with magnetite occur in the Käymäjärvi deposit is of genetic interest. These are most likely metamorphosed granules of greenalite or some other iron silicate. Otherwise there is no knowledge of the original iron mineral. These ores were formed by a deposition of iron and silica with varying amounts of limestone or dolomite in the final stage of the basic volcanism that gave rise to the greenstone group. Through later metamorphic processes the iron-bearing sediments recrystallized and the carbonate material reacted with iron and silica, giving rise to the skarn minerals.

Quartz-banded iron ores also occur outside the iron ore-bearing zone, in the southern part of the Norrbotten county, but differ somewhat from those described above. The ores are not associated with basic volcanics but with acid ones. The immediate wall rock may be made up of sediments, such as mica-schists or quartzites, but the ores occur also directly in the volcanics. The most common skarn minerals are diopside, tremolite-actinolite, garnet, epidote and biotite. The content of manganese in some deposits is rather high, in cases rising to 7 per cent. This

79

R. Frietsch

element is concentrated in layers of skarn or carbonate in the ore. In some deposits hematite formed secondarily from magnetite is found in subordinate amounts.

SKARN ORES

These ora in Norrbotten have the same mode of appearance as the main part of the quartz-banded ores in the county. They occur in the greenstone group, mostly in the stratigraphically higher part, where they form lens-shaped bodies concordant with the host rocks, which are either chemical or detrital sediments or the greenstones proper. In many cases the skarn ores are associated with limestones or dolomites or lie in the same stratigraphic position as these. The ore mineral is magnetite or, exceptionally, also hematite. The amount of iron varies in most cases between 30 and 40 per cent. The ores usually contain small amounts of pyrite and pyrrhotite. The content of sulphur usually exceeds 1 per cent. The content of phosphorus, in the form of apatite, is in most cases less than 0.1 per cent, but in some deposits the content is higher and rises locally to 1 or 2 per cent. The content of manganese is usually less than 0.2 per cent.

The skarn minerals that accompany the ore in large amounts are evenly distributed in the ore or form independent masses or layers. A rather common feature is a layering between magnetite and the skarn minerals. Sometimes a layering of calcite with magnetite and skarn minerals is observed. Among the skarn minerals tremolite-actinolite, diopside, phlogopite and serpentine dominate.

The skarn iron ores of Norrbotten were considered by Geijer (1931) and Geijer & Magnusson (19526) as pyrometasomatic replacing limestones and dolomites. This view is supported by the fact that in the Junosuando deposit at Masugnsbyn fluorine occurs in the skarn, mainly bound to chondrodite. The iron-bearing solutions in this deposit should, in the opinion of Geijer, emanate from a pethite granite which forms the foot wall of the ore. However, a sedimentary origin for the skarn iron ores in northern Sweden has postulated most recently by the present author (Frietsch, 1966, 19676, 1970).

The main reasons for a sedimentary mode of formation are that the skarn iron ores and the quartz-banded iron ores, of which the latter are undoubtedly of sedimentary origin, both occur in connexion with sediments in the greenstone group and almost in the same stratigraphic position. In some deposits the two types occur intermingled with each other and a clear division between them is not possible. The skarn-layering and the more rare carbonate- layering in the skarn ores is probably a relict sedimentary texture. As regards the relationship between the older group o€ intrusives and the skarn ores, there now exist indications that the formation of the ore is older than this group. Thus in the small Juolovanjärvet deposit the ore, with a folded skarn-banding, is cut by a granodiorite which most likely belongs to the older intrusive group (Frietsch, 19676).

There is also evidence that the perthite granite at Junosuando cannot have given rise to the ore in the deposit, as the granite cuts the ore and contains inclusions of the skarn. Welin and Blomqvist (1966) investigated the age of an uraninite-bearing sample of the skarn ore from this deposit. The uraninite is enclosed in chondrodite which follows the magnetite. An age of 1,775-1,845 m.y. was obtained. The age of the perthite granite on the other hand is 1,540 m.y. (Welin, 1970).

Further it must be mentioned that the magnetite in the skarn iron ores and the quartz-banded iron ores has a similar trace element distribution (Frietsch, 1970). Of special interest is the relatively high content of magnesium in both types. In the skarn ores it can reach 5.1 per cent (mean value 0.96 per cent) and in the quartz-banded ores 3.3 per cent (mean value 0.53 per cent).

Due to the above-mentioned facts, it is most proble that the skarn iron ores, as are the quartz-banded iron ores, in northern Sweden are original sediments in which iron, silica and carbonates have been deposited in varying proportions simultaneously with the sediments in which the ores occur. The deposition of this iron-bearing formation took place mainly in the final stage of the basic volcanism that gave rise to the greenstone group. In connexion with later regional metamorphic processes the original constituents of the ores have recrystallized and reacted with each other giving rise to the skarn minerals. At the same time the ores were in many cases mobilized and thus here the primary sedimentary features have been almost totally obliterated.

Comparisons between non-apatitic iron ores in central and northern Sweden

In summary the non-apatitic iron ores in central and northern Sweden are most probably of sedimentary origin. Some fundamental differences in the composition of the ore and the gangue of the two regions are here ascribed to differences in the depositional environment. When compared, the following features are relevant.

HOST ROCKS

In both regions the iron ores occur in volcanic rocks, in central Sweden in acid-intermediate ones and in northern Sweden in basic ones. Quartz-banded ores in the southern part of the Norrbotten county also lie in acid volcanics. In both regions the ores are in many cases associated with limestones-dolomites or in the same stratigraphic position as these. The age of the volcanics in central Sweden is more than 1,800 m.y. and in northern Sweden between 1,880 and 2,800 m.y.; a more precise dating is not possible at the moment.

80


IRON MINERALS

In central Sweden the iron is believed to have been deposited in the ferric state (hematite or limonite) but altered to magnetite by later metamorphism. Hematite is often preserved in the quartz-banded ores, especially when such deposits have been subject to only relatively weak metamorphism. Otherwise the present ore mineral is magnetite. As pointed out by the present author (Frietsch, 1967a) the alteration of hematite to magnetite, which is believed to have occurred in connexion with the folding of the volcanics and the intrusion of the older Svecofennian granites, was coupled with reducing conditions. During this epoch there occurred the magnesia-metasomatism by which sulphides were formed also. Oxygen fugacity was then low. The existence of reducing conditions is further supported by the fact that the metasomatism resulted in the formation of silicates rich in magnesium and ferrous iron, silicates with ferric iron not being formed.

In the non-apatitic iron ores in northern Sweden there is no knowledge of the original form in which the iron was precipitated, except that in the quartz-banded ore at Käy- mäjärvi pseudomorphs occur after greenalite or some other iron silicate. Magnetite is, except when hematite is formed secondarily after magnetite, the only iron oxide. That magnetite and not hematite is the present iron mineral is possibly due to the fact that the ores, especially the skarn ores, contain small amounts of iron sulphides. They are probably of syngenetic origin deposited together with the iron (Frietsch, 1966). The presence of the sulphides means a reducing milieu and magnetite is, therefore, probably the primary iron mineral in many cases. Through later oxidation magnetite has been changed to hematite, but only in those deposits where sulphides are missing.

GANGUE

The quartz-banded ores of central Sweden have a gangue that is composed of Ca-Mg-rich or Ca-FeS+-rich silicates, while the gangue in the quartz-banded ores in northern Sweden is composed of FeZ+-rich or Fe2+-Mg-rich silicates. The skarn that follows the quartz-banded ores in the southern part of the Norrbotten county is Ca-Mg-rich or Ca-FeZ+-rich and is thus rather similar, to the skarn following the quartz-banded ores and skarn ores in central Sweden. On the whole there seems to be a similarity both in composition and geologic milieu between the quartz- banded iron ores in the southern part of Norrbotten and in central Sweden.

Also the gangue in the skarn iron ores in central Sweden and in northern Sweden differs in composition. The skarn silicates in central Sweden are in part rich in Ca and in part rich in Mg, the later ones being clearly later and formed through the magnesia-metasomatism. In northern Sweden the skarn silicates are Ca-Mg-rich or Mg-rich. The Ca-Mg-rich skarns are found in almost every ore deposit, the Mg-rich skarns being less common.

The internal relationship between the Ca-Mg-rich silicates (tremolite, diopside) and the Mg-rich silicates (serpentine, phlogopite) is imperfectly known. In some deposits there seems to be a tendency for the Ca-Mg-rich silicates to form independent masses outside the ore and the Mg-rich silicates to be distributed in the ore itself. There are indications from some deposits that the Mg-rich silicates are later than the Ca-Mg-rich silicates, the order of formation being diopsidetremolite-serpentine.

S-P-MN-CONTENT

As previously pointed out, the non-apatitic iron ores in northern Sweden in most cases contain small amounts of iron-bearing sulphides, while the iron ores in central Sweden are more or less free from sulphides. The same is true of the phosphorus content. The relatively high content of phosphorus found in some skarn ore deposits in northern Sweden is certainly a primary feature, as there is no sign of secondary addition. One more important chemical difference between the non-apatitic iron ores in both regions, is that manganese-bearing ores are relatively abundant in central Sweden, but are almost missing in northern Sweden. In the latter region the precipitation of iron was not accompanied by manganese.

AGE OF METAMORPHISM

For the non-apatitic iron ores of central Sweden it seems obvious that the metamorphism with recrystallization and skarn formation occurred in connexion with the folding of the volcanic-sedimentary complex and the intrusion of the older group of the Svecofennian granites. Additional changes took place in connexion with late Svecofennian granites. For the non-apatitic iron ores of northern Sweden it is less clear if the metamorphism affecting the ore was related to the older granodiorite group or the younger Lina granite. The present author previously considered the Lina granite as most important in this connexion, the granodiorite group having only a smaller effect (Frietsch, 1966), but is now inclined to believe that the older intrusives had a significant influence too. The greater age of the skarn formation compared with the Lina granite is shown by the radiometric age determinations from Masugnsbyn, the uraninite in the skarn having an age of 1,775-1,845 m.y. It is possible that the metamorphism in part is older than the granodiorite group. The only proof for this view is the earlier mentioned observation from the Juolovanjärvet deposit which shows that the ore and the skarn had recrystallized before the intrusion of the granodiorite. This does not, however, exclude the possibility that the recrystallization-skarn formation and the granodiorite intrusion belong to the same process, the time gap between these being relatively small.

81

R. Frietsch

Résumé

Minerais de fer précambriens ù caractères sédimentaires, en Suède (R. Frietsch)

A part les minerais de fer oolithique de l'époque jurassique dans le sud de la Suède, les autres minerais de fer suédois qui présentent des caractères sédimentaires remontent au Précambrien. Ils sont confinés dans deux régions géogra- phiques distinctes, l'une au centre de la Suède et l'autre au nord. Dans les deux cas, on trouve des minerais veinés de quartz avec faible teneur en phosphore. Dans la région centrale de la Suède, le minerai consiste en couches alter- nées de quartz et d'hématite, l'hématite étant en général plus ou moins remplacée par de la magnétite. Dans les minerais du nord de la Suède, oii le minerai est la magnétite, avec de moindres proportions d'hématite, la structure rubanée est en quelque sorte moins prononcée. Les minerais veinés de quartz ont dû être à l'origine des sédiments chimiques, formés par la même activité magmatique qui a engendré les appareils volcaniques où on les trouve. Plus tard, des processus métamorphiques en relation avec le plissement des

systèmes volcaniques et l'intrusion de granite ont produit la recristallisation et la réorganisation interne des minerais, comme par exemple la formation de silicates de (( skarn 1) là où les carbonates étaient associés au fer et à la silice. Dans les deux régions on rencontre aussi du (( skarn 1) pauvre en phosphore et constitué de magnétite et de silicates de (( skarn )I. Dans de nombreux gisements, on note une stratification plus ou moins prononcée de magnétite, de silicates et parfois de carbonates. L'origine de ces minerais a donné lieu à bien des discussions, car on y observe à la fois des caractères sédimentaires et pyrométasomatiques. On pos- sède toutefois, maintenant, des indications montrant que presque tout le minerai (( skarn )) de la Suède centrale est sédimentaire, mais les caractères sédimentaires originaux ont été, la plupart du temps, oblitérés par des processus ultérieurs. Le minerai (( skarn 1) du nord de la Suède semble aussi être constitué de sédiments métamorphosés. Cette opinion est basée sur le fait, entre autres, que la magnétite dans les minerais veinés de quartz et les minerais (( skarn 1) ont une composition géochimique voisine.

Bibliography /Bibliographie

FRIETSCH, R. 1966. Berggrund och malmer i Svappavaarafältet, norra Sverige [Geology and ores of the Svappavaara region, northern Sweden]. Sverig. geol. Unders. Afh., C 604. (In Swedish with English summary.)

--. 1967~. The relationship between magnetite and hematite in ihe iron ores of the Kiruna type and some other iron ore types. Sverig. geol. Unders. Afh., C 625.

-. 19676. On the relative age of the skarn iron ores and the Haparanda granite series in the county of Norrbotten, northern Sweden. Geol. Fören. Stocklz. Förh., no. 89, p. 116-18. __ . 1970. Trace elements in magnetite and hematite mainly from northern Sweden. Sverig. geol. Unders. Afh., C 646.

GEIJER, P. 1925. Eulysitic iron ores in northern Sweden. Sverig. geol. Unders. Afh., C 324. __ . 193 1. Berggrunden inom malmtrakten Kiruna-Gällivare- Pajala [The geology within the ore region Kiuna-Gällivare- Pajala]. Sverig. geol. Unders. Afk., C 366. (Tn Swedish with English summary.) __ . 1959. Några aspekter av skarnmalmsprobleinen i Bergs- lagen [Some aspects regarding the problems of the skarn iron ores in the region of Bergslagen]. Geol. Faren. Stockh. Förh., no. 81, p. 514-34. (In Swedish with English sunmary.)

GEIJER, P.; MAGNUSSON, N. H. 1944. D e mellansvenska järnmal- mernas geologi [The geology of the iron ores of middle Swe- den]. Sverig. geol. Unders. Afh., Ca35. _- . 1952a. Geological history of the iron ores of central Sweden. XVZIZ Ini. geol. Congr., Great Britain 1948, Part XIII, p. 84-9. -- . 19526. The iron ores of Sweden. In: Symposium sur les gisements de fer du monde. XZX Congr. Gol. Znt., Alger, p. 477-99.

MAGNUSSON, N. H. 1953. Mulmgeologi[ore geology], Stockholm, Jernkontoret .

-. 1960. Iron and sulphide ores of central Sweden. Guide to excursions nos. A 26 and C 21. XXZ Int. geol. Congr., Norden, 1960. Geological Survey of Sweden.

MAGNUSSON, N. H.; THORSLUND, P.; BROTZEN, F.; ASKLUND, B.; KULLING, O., 1960. Description to accompany the map of the pre-quarternary rocks of Sweden. Sverig. geol. Unders. Afh., Ba16.

ÖDMAN, O. H. 1957. Beskrivning till berggrundskarta över ur- berget i Norrbottens län [Description of the geological map of the primary rocks in Norrbotten county]. Sverig. geol. Unders. Afh., Ca 41. (In Swedish with English summary.)

OFFERBERG, J. 1967. Beskrivning till berggrundskartbladen Kiruna NV, NO, SV, SO [Description of the geological maps of sections Kiruna NW., NE., SW., SE.]. Sverig. geol. Unders. Afh., Af 1-4 (In Swedish with English summary.)

PADGET, P. 1970. Description of the geological maps Tärendö NW., NE., SW., SE. Sverig. geol. Unders. Afh., Af 5-8. (In press.) WELIN, E. 1970. Den svekofenniska orogena zonen i norra Sverige. En preliminar diskussion [The Svecofennian orogenian zone in northern Sweden. A preliminary discussion]. Geol. Füren. Stockh. Förh. (In press.) (In Swedish with English summary.) WELIN E.; BLOMQVIST, G. 1964. Age measurements on radioactive minerals from Sweden. Geol. W r e n . Stoclch. Förh., no. 86, p. 33-50.

-, 1966. Further age measurements on radioactive minerals from Sweden. Geol. Fören. Stoclch. Förh., no. 88, p. 3-18.

82

R. T. BRANDT. What is the relationship, if any, between the sedimentary iron ores of northern Sweden and the massive magnetite deposits of Kiruna and Gällivare?


Discussion

R. FRIETSCH. There is no relationship between them. The sedimentary iron ores and the iron ores of the Kiruna type are quite different in mineralogy, host rock, geological appearance, etc.

A. S. KALUGIN. What is your opinion as to the origin of the Grengesberg deposit?

R. FRIETSCH. The deposit belongs to theKiruna type of ore, and should thus be of magmatic origin,

V. M. CHERNOV. Has the Svecofennian basement been found?

R. FRIETSCH. No, there is no knowledge of such a basement.

V. M. CHERNOV. What is the age of the scapolitization and metasomatism?

R. FRIETSCH. The scapolitization is most probably connec-

ted with the Lina granite. There are, however, indications that the formation of scapolite in some cases is much older.

V. M. KRAVCHENKO. Will you show on your regional map the ‘Stabby’ and ‘Maria’ deposits?

R. FRIETSCH. I regret that I do not know of the existence of such deposits. I suppose they are situated in central Sweden.

V. M. KRAVCHENKO. Have any definitely magmatic basic rocks been found in the stratigraphic section of host rocks in these deposits?

R. FRIETSCH. Such rocks have been found among the basic lavas (greenstones) in the greenstone group. The gabbros belong in most cases to the older group of intrusives.

V. M. KRAVCHENKO. What is the absolute age of the ore- bearing rocks of the deposits cited?

R. FRIETSCH. Unfortunately there are no radiometric age determinations. The rocks of the greenstone group seem to have an age between 1,880 m.y. (older intrusive group) and 2,800 m.y. (granite basement).

83

The ferruginous-siliceous formations of the eastern part of the Baltic shield

V. M. Chernov Karelian Branch of the Academy of Sciences of the U.S.S.R. Geology Institute

The Precambrian of the eastern part of the Baltic shield is divisible into four chronological units or structural stages differing in their tectonic structure, rock composition and peculiarities of magmatism and metamorphism.

The lowest structural stage (the Early Proterozoic (Ar- chaean) foundation) is composed of various gneisses, granitoids, amphibolites and migmatites, which have been subjected to intensive deformation and deep, often repeated, metamorphism and ultra-metamorphism.

The second structural stage is characterized by the development of multiple folded, mottled Early Proterozoic primary sedimentary and sedimentary-volcanic series which have undergone changes from greenschist facies to amphibolite facies of regional metamorphism.

The third structural stage is characterized by the primary development of terrigenous, often rudaceous, deposits, accompanied by carbonate, shungite, pelite rocks and basic volcanites occurring mainly in interstructures of various types (through lines, etc.). The sediments of this structural stage are distinguished by their low grade of metamorphism, are of Middle Proterozoic age and are interpreted by many investigators as formations of the orogenic and subplatform stages in the development of Karelides.

The fourth structural stage-a platform mantle, which is preserved over a small area-is composed of nonmeta- morphosed and weakly dislocated terrigenous Jotnian sediments.

The ferruginous-siliceous formations are confined exclusively to deposits of the second structural stage belonging chronologically to the Lower Proterozoic (2,600- 2,000 m.y.). Some deposits of hematite-martite ore are known also in the territory of Karelia in deposits of the third structural stage.

The results of studies of geology, lithology and geochemical properties of the iron-ore series of the Kola- Karelian region during the past fifteen to twenty years allow us to distinguish between the three genetic types of ferruginous-siliceous and ferriferous formations, differing in paragenetic rock associations, palaeotectonic and palaeofacial conditions of sedimentation as well as in

their stratigraphic position in the Precambrian sequence. The following types can be distinguished: leptite-por-

phyric series of ferruginous-siliceous formations; spilite- diabasic ferruginous-siliceous formations; clastogene ferruginous formations.

The leptite-porphyric series of ferrugirious-siliceous formations. These are widely distributed in the territory of western Karelia in the deposits of the Himola series and are genetically connected with Lower Proterozoic intensive persilicic volcanism. During recent years ferruginous-siliceous rocks, paragenetically connected with persilicic effusions, have been distinguished in the Kola peninsula in the formations of the Kola series (Olenyegora suite).

The ferruginous-siliceous formations of this genetic type have been studied mostly in Karelia where, according to paragenetic rock associations, the schistose-leptite and leptite-porphyric formations are prominent among them. The schistose-leptite ferruginous-siliceous formation comprises the lower part of the Himola series and, according to its volume, is in conformity with the first sedimentary- eruptive cycle of this series. It is composed of various primary sedimentary and eruptive rocks metamorphosed to different degrees among which appear conglomerates, gritstones, arkoses and various aluminiferous gneisses (garnet- biotite and staurolite). Less-common types which occur are tuff breccia, quartz-biotite tuffaceous shales, graphitic, muscovitic, sericitic, talcose and chloritic shales, amphibolitic paraschists and para-amphibolites. Ferruginous- siliceous rocks, mainly biotite, riebeckite and grunerite varieties, are located in the upper part of the formation, where they alternate rhythmically with the above-mentioned shales and gneisses.

It is characteristic of the schistose-leptite ferruginous- siliceous formation that besides eruptive and ferruginous- siliceous rocks, in its constitution there is a widespread development of metamorphosed terrigenous deposits and a facial variation from section to section, caused by the pinching out of lithological units and lateral changes in rock type along the strike.

The leptite-porphyric ferruginous-siliceous formation


V. M. Chernov

conforms in volume to the deposits of the Himola series and in its lower part consists of persilicic volcanic rocks transformed into leptite gneisses, porphyroids, halleginta and various tufogenic crystalline schists and leptites. The widespread development of volcanites deposited in aqueous medium (rhythmic lamination in tuffaceous shales) shows that the development of this formation took place in underwater conditions and was accompanied by intensive volcanism. Along the strike of the western Karelian structural- facial zone, the main features of the paragenesis of rocks of the leptite-porphyric ferruginous-siliceous formation remain unaltered, due perhaps to monotypic palaeofacial conditions of sedimentation.

Spilitic-diubusic ferriiginous-siliceous formations. These are widespread in the Kola peninsula (Kola and Tundra series) and in Karelia in the deposits of the Himola and Parandov series. In contradistinction to the ferruginous-siliceous formations of the leptite-porphyric series, theferrugi- nous-siliceous formations now described are paragenetically closely connected with volcanic series of basic composition. Amphibolitic shales, amphibolites and pyroxene-hornblende gneisses, formed as the result of metamorphic transformations, mainly of basic effusive rocks and their tuffs, are the main members of the rock associations of these formations.

Horizons of paracharnockites and aluminiferous gneisses are often met in the Kola peninsula. Ferruginous- siliceous rocks and graphitic shales enriched by sulphides are observed in the form of thin (0.5-15 m) beds and lenses among the rocks mentioned. A rhythmic structure, consisting of reiterations in the sequence of regularly built rock bands 50-200 m thick composed in the lower parts of metamorphosed volcanites of basic composition and in the upper parts of ferruginous-siliceous rocks and paraschists, is a characteristic lithological peculiarity of the ferruginous- siliceous formations of Karelia.

Clustogene iron ores. The martite-hematite in thin deposits and partings is developed mainly in the southern and south-western part of Karelia (Prionezhye, Suojärvi, Tulo- mozero, Janisjärvij in the Jatulian deposits of the third structural stage. In contradistinction to volcanogenic ferruginous-siliceous formations these ores do not accumulate by volcanic-sedimentary processes. They associate only with terrigenous rocks. The correlation of sections and the paragenetic analysis of the iron-ore strata of Karelia and

the Kola peninsula allow us to distinguish two large epochs of iron accumulation connected with certain stages of the tectonic development of this territory.

During the geosynclinal period, pertaining to the Lower Proterozoic (2,600 f 100-2,000 m.y.), processes of volcanism proceeded in the Karelian and Kola-Norwegian geosyncline zone. Under the influence of these processes ferruginous-siliceous formations of sedimentary-volcanogenic origin were developed. Moreover, there is a certain dependence in the distribution of ferruginous-siliceous formations upon palaeotectonic conditions of development.

Areas of leptite-porphyric formations are distinguished, according to geophysical investigations, by great thicknesses of the earth‘s crust and the ‘granitic layer’. Intensive persilicic volcanism, the thickness of the ‘granite layer’ and the earth’s crust, the wide distribution of terrigenous sediments, breaks in sedimentation and comparatively thin layers of sediments all confirm that a geoanticlinal régime of sedimentation took place during the development of ferruginous-siliceous formations of the leptite-porphyric series.

The spilitic-diabasic ferruginous-siliceous formations were developed in different tectonic conditions. Spatially they are associated with zones of intensive warping and abyssal fractures limiting the Lapland-White Sea and Murmansk blocks. These ‘greenrock‘ warps are cliarac- terized by the thinness of the earth’s crust and ‘granite layer’. It is possible that such a difference in the structure of the earth’s crust in the eastern part of the Baltic shield determined the different types of initial Early Proterozoic volcanism and the difference in palaeotectonic conditions of the development of leptite-porphyric and spilite-diabasic ferruginous-siliceous formations.

The second epoch of iron sedimentation is in connexion with the orogeny stage of tectonic development (Middle Proterozoic 2,000-1,750 m.y.). During this period the Karelian geosyncline zone was transformed into a foldmountain country with accumulations characteristic of this period of orogenic and subplatform deposition of formations of Sariola and Jatulia (molasse; arenaceous, carbonate-terrigenous, shungitej. Ferruginous rocks, developed mainly by clastogenic processes, appeared due to the destruction and weathering of Lower Proterozoic volcanogenic thicknesses and ferruginous-siliceous formations.

Résumé

Les formations de fer siliceux duns lu partie orientale du bouclier baltique (V. M. Chernov)

bien définis du développement tectonique de cette région. 2. L’époque la plus significative de l’accumulation du

fer correspond à la période du Protérozoïque inférieur du 1. L’analyse des formations et la corrélation des développement géologico-tectonique 2 600-2 O00 millions

formations géologiques des structures plissées protéro- d’années) alors que le bouclier baltique était une vaste zoïques des Karélides permettent de choisir dans la partie orientale du bouclier baltique deux époques significati- 3. Dépendant des conditions paléotectoniques de sédi- ves d’accumulation de fer correspondant à des stades mentation, deux séries de formations de fer siliceux se sont

région géosynclinale.

86

The ferruginous-siIiceous formations of the eastern part of the Baltic shield

formées pendant cette période. Elles sont liées génétique- ment à des manifestations de volcanisme acide et basique. Les séries leptite-porphyriques des formations de fer siliceux, qui se sont développées dans la zone karélienne des karélides, se sont formées dans des conditions de régime (( géoanticlinal 1) de sédimentation dans un vaste bloc intra- géosynclinal composé de roches sialiques du début de l'archéen.

4. Les formations de fer siliceux des séries spilite-diabasiques montrent une tendance générale vers les régions de submersion intensive et de haute perméabilité de la croûte (failles profondes). Sur le bouclier baltique, elles sont confi- nées à la dépression de (( greenstone N.

5. La deuxième époque du Protérozoïque où l'on ob-

serve l'accumulation du fer sur une grande échelle est en relation avec un stade orogénique de développement de cette région (Protérozoïque moyen, 2 000-1 750 millions d'an- nées). Pendant cette période, des formations de fer clasto- gène se sont constituées sur le territoire de Karélie, résultant de la destruction des strates du géosynclinal du Protéro- zoïque inférieur.

6. Il n'y a aucune roche ferreuse au stade orogénique de développement dans la péninsule Kolsky.

7. Les informations données ci-dessus permettent de classer les formations de fer siliceux du Précambrien situées dans la partie orientale du bouclier baltique en fonction de la tectonique.

Discussion

A. M. GOODWIN. What is the definition of persilicic volcanics? D o they belong to the calc-alkaline series?

V. M. CHERNOV. These are metamorphosed lavas and tuffs of dacite and keratophyre composition. Yes, they do.

87

Precambrian ferruginous-siliceous formations associated with the Kursk Magnetic Anomaly

N. A. Plaksenko, I. K. Koval, I. N. Shchogolev

The area of the Kursk Magnetic Anomaly (KMA) is huge, about 450 km long and 130-150 km wide (Fig. i). In this area there are two rather well-distinguished lines of magnetic anomalies: northeast and south-west. Their position and configuration are due to steeply dipping beds of ferruginous quartzite in belts of metamorphic rocks. The structure of these beds is not simple. Besides the clearly distinguished, long magnetic anomalies there are many local short anomalies, which are situated both between the main lines and outside them or on their continuation. The crystalline basement of the KMA region is covered by sedimentary rocks of Palaeozoic, Mesozoic, and Cenozoic ages, the thickness of which is 30 m in the north-east and 300 m in the south and south-west . The basement is cut by numerous boreholes and is exposed in iron ore pits.

In the Precambrian rocks associated with the KMA there are two time-structural units: the lower (Archaean) includes different gneisses, migmatites, granites and other rocks of the Oboyanskaya series, and volcanic-sedimentary rocks, products of the metamorphism of spilitic-keratophyric rocks and quartz porphyries of the Mikhailovskaya series. The upper time-structural unit consists of Protero- zoic rocks which are separated from the Archaean rocks by a stratigraphic and structural disconformity. The Protero- zoic rocks are separated into two series: Kursk (lower Proterozoic) and Oskol, which is called the Kurbakin series (lower and middle Proterozoic) in the north-west part of the area. The Oskol (Kurbakin) series overlying the Kursk series with disconformity is clastic ferruginous quartzite, detrital sediment rich in iron, sandstone-shale, shale-carbonaceous rock, and acid effusive rocks.

W e presented samples of rocks of the Kurbakin series containing relicts of algae to the Pollenological Laboratory of Voronezh State University; remnants of the simplest algae were found (Daminorites Eichw . , Oscillatorites Shcp., Leiomarginata Umnova, Rifenites Naum., Turuchanica Tim., Brochosophosphaera facetus Schep., Trachysophos- phaera Naum., Leiosphaeridia tipa Volcova), establishing the middle Proterozoic age of the rocks.

Ferruginous-siliceous rocks (ferruginous quartzites and

gneisses) recur several times in a cross-section of the Pre- cambrian rocks of the K M A . Non-ferruginous rocks enclosing them are closely connected with them in origin and at the same time are separated from one another by discon- formities. It is possible to determine four genetically independent formations of ferruginous-siliceous rocks of different ages, which successively alternated during the history of the region. They are: (a) ferruginous-siliceous gneiss (partings of magnetite gneiss in rocks of the Oboyan series); (b) ferruginous-siliceous metabasite (thick partings of quartz, silicates, and magnetite in rocks of the Mikhailov series). (c) ferruginous-siliceous slate (thick mass of ferruginous quartzite of Kursk series enclosed in schist); (d) ferruginous- siliceous clastic rock (clastic ferruginous quartzite and rich fragmental ore of Oskol-Kurbakin series).

Ferruginous-siliceous gneiss

The rocks of this formation occur in the Kursk-Besedino region of the KMA.

Ferruginous-siliceous rocks which are conformable in pyroxene-amphibole, garnet, and pyroxene-plagioclase gneisses and other rocks are characterized by quartz-magnetite-pyroxene (hypersthene) and garnet-magnetite-hypersthene composition. They are massive and usually vaguely banded or, rarely, distinctly banded.

Interbeds and lenses of ferruginous-siliceous rock have thicknesses of from 0.8 m to 35 m and are of small extent.

Ferruginous-siliceous metabasite

At present this formation is distinguished conditionally. Its ferruginous-siliceous rocks are amphibole-magnetite and chlorite-magnetite quartzites of subore grade, which occur as thin interbeds at the bottom of the upper effusive-sedimentary (keratophyre-shale) suite of the Mikhailov series and alternate with quartz-chlorite, albite-chlorite-biotite, and albite-chlorite-amphibole slates.


N. A. Plaksenko, I. K. Koval and 1. N. Shchogolev

\

Belgorod q

0 Q Nizhnedevitsk

orshechnoe

FIG. 1. Areal map of the Kursk Magnetic Anomaly. Magnetic anomalies: 1, connected with ferruginous quartzites of the Kursk

series; 2, connected with ferruginous quartzites of supposed Obojan and Mikhailov series.

Ferruginous-siliceous slate

This formation consists of ferruginous quartzite of the Kursk series, which is conformable in a sequence of metamorphosed terrigenous sandstone-shale deposits.

Ferruginous quartzites are directly underlain and overlapped by phyllitic, carbonaceous slate and crystalline quartz-sericite and garnet-biotite slate. This formation occurs on various rocks of the eroded Archaean basement with stratigraphic and structural unconformity. Dowii-

wards, rocks underlying ferruginous quartzite give way to metamorphic sandstones, metagravelites, grusses, and conglomerates of Archaean age and also to a metamorphosed crust of weathering.

Ferruginous-siliceous rocks are magnetite quartzite and micaceous hematite-magnetite quartzite with various proportions of magnetite and hematite (jaspilites). A zoning of authogenic minerals is typical. The dimensions of this formation are huge and its productivity (in iron) is the greatest in the region. Below, we give its principal structural characteristics.

90


CLASSIFICATION OF FERRUGINOUS- SILICEOUS ROCKS

Among the varied ferruginous-siliceous rocks that compose the iron ore bands of the formation it is possible to distinguish a limited number of types, the original sediments of which were formed under certain facies conditions and at definite places in the facies profile (cross-section) of the formation. The main genetic types of ferruginous-siliceous rock of the formation are as follows: (a) low-grade ore (with magnetite) and non-ore-bearing silicate-carbonate quartzite (non-ore silicate-carbonate-protoxide facies); (b) silicate (cummingtonite)-magnetite quartzite with ferromagnesian carbonates (ore carbonate-silicate-ferruginous oxide-protoxide facies); (c) magnetite quartzite (ore-bearing magnetite protoxide-oxide facies); (d) micaceous hematite-magnetite quartzite; (e) magnetite-micaceous hematite quartzite; (f) micaceous hematite quartzite.

Types (d) to (f) belong to the ore-bearing hematite- oxide facies. Each of the types may be characterized by a certain mineralogical composition, by geochemical peculiarities, textures and structures, and by a regular position in the cross-section of the formation.

REGULAR ALTERNATION O F ROCKS IN THE FACIES PROFILE OF THE FORMATION

The composition of the iron ore suite is not identical in different parts of the KMA region. This is due to considerable variation in the thickness and to changes in the facies of ferruginous quartzite, and to varying numbers of ferruginous interbeds and bands in the iron ore suite. However, if these genetic types of ferruginous-siliceous rock are arranged according to the order of their abundance as they alternate with underlying and overlying rocks, ignoring interbedding which involves the adjacent types, we obtain a generalized cross-section of the ferruginous silicate slate formation of the following kind (from top to bottom): (a) overlying schist-phyllitic carbonaceous slate-schist; (b) quartzite (non-ore or silicate-bearing and carbonaceous with small amount of ore); (c) quartzite (silicate-bearing and magnetitic with ferromagnesian carbonates); (d) magnetite quartzite; (e) magnetite-micaceous hematite and micaceous hematite quartzites; (f) micaceous hematite-magnetite quartzite; (g) magnetite quartzite; (h) ore-bearing silicate-magnetite quartzite; (i) quartzite (non-ore-bearing or carbonaceous silicate-bearing with small amount of ore). In places, the quartzite has interbeds and lenses of carbonaceous magnetite ore and amphibole schist; (j) carbonaceous magnetite ore with sulphides and pyrite-carbonaceous ore with magnetite. Sulphide-bearing ore occurs close to the contact with schist and in the schist itself. Carbonaceous magnetite ore occurs mostly at the contact with barren quartzite and penetrates into the quartzite; (k) phyllitic carbonaceous slate and schist, considerably pyritized in its upper part. In places, the slate and schist have interbeds of sulphide and carbonaceous magnetite ore; (1) arkosic metasandstone,

barren (non-ferruginous) quartzite containing blastopsammitic structure, metagravelite and metaconglomerate.

Carbonaceous magnetite and pyritic carbonaceous ores occur locally in the formation. Omitting such local occurrences of ore, we have an ideal cross-section of the ferruginous-siliceous shale formation, which can be taken as a manifestation of one simple cycle of sedimentation characterized by similar conditions of sedimentation at the beginning and end of the cycle. The position of some facies in such a cross-section, from the shore seaward to depth, is as follows: Non-ore facies: blastopsammitic barren quartzite-metasandstone-carbonaceous barren silicate-bearing quartzite or with a small amount of ore (silicate-carbonaceous facies).

Ore-bearing facies: silicate-magnetite quartzite (carbonaceous-silicate-magnetitic-hematitic facies). The composition of ore minerals, carbonates and silicates regularly changes towards the deeper deposits, beginning with the silicate-carbonaceous facies part ot the profile. An ideal facies profile of the formation formed under the condition of gradual lowering of the bottom from the shore toward the deepest part of the basin, as represented by the distribution of sediments according to their grain sizes; with increasing depth pelitic material loses its importance and ferruginous-siliceous colloids increase in importance. Organic matter also is disseminated in accordance with grain sizes of sediment; it is enormously concentrated in the pelagic zone of pelite deposition, but gradually decreases in the deeper water deposits. There was a decrease of Eh near the contacts of shale and barren quartzite, and moderate hydrogen sulphide contamination, resulting in the appearance in the shale of a vague sulphide subfacies of the local protoxide facies (represented by carbonate-magnetite and pyritic carbonaceous ores). Far- ther from the shore of the basin, with a decrease of organic matter and a decrease of activity of decayed organic matter, the process of reduction gradually gave way to reduction-oxidation processes, which in turn passed into oxidation processes, resulting in a gradual increase in the importance of iron oxide minerals from the shallowest to the deepest part of the ore-bearing portion of the facies profile. Accordingly, we can precisely outline an authigenic mineral zoning as a feature of the facies profile of the ferruginous-siliceous slate formation (Fig. 2).

The sedimentation cycles represented in the cross-section of the formation demonstrate a regular alteration of conditions of sedimentation from less stable and more shallow in the eastern part of the KMA basin (north- eastern belt) to more stable and deeper water in the western part (south western belt). This accounts for the increased importance of ferruginous-siliceous rocks of the oxide facies and the presence of the ferruginous quartzite series toward the west and south-west.

Ferruginous quartzites having a thickness of less than 200 m and deposited under littoral conditions, mainly in the eastern wing of the north-eastern belt of the K M A , are

91

N. A. Plaksenko, I. K. Koval and I. N. Shchogolev

FIG. 2. Authigenous-mineralogical zoning of facies section of ferruginous-siliceous-slate formation. 1, low ore and barren quartzites; 2, cummingtonite-magnetite quartzites; 3, magnetite quartzites; 4, hematite-magnetite quartzites; 5, pyrite; 6, siderite-pistomesite; 7, iron silicates/cummingtonite; 8, magnetite; 9, hematite.

an ore-bearing facies and consist almost entirely of protoxide and oxide-protoxide minerals. The ferruginous quartzite is interbedded with a relatively large thickness of schist. Towards the south-western belt of the KMA, the proportion of ferruginous quartzite to schist increases sharply and the proportion of oxide facies minerals increases in the ferruginous quartzite. Deposition of sediments of the south-western belt is believed to have taken place, in general, in deeper water than did sediments in the north-eastern belt. W e conclude that the iron-ore building process was sensitive, especially in the shallow water portion of ore facies, to the slightest alteration of conditions of sedimentation, which resulted in a direct relationship between the thickness of different ferruginous quartzites and their genetic type of facies.

A regular change of geochemical and other properties of the ferruginous-siliceous rocks and their main rock-forming minerals takes place in the facies profile of the formation. The essence of the regular geochemical changes is illustrated by Table 1.

TABLE 1. Systematic changes in the geochemical and other properties of ferruginous-siliceous rocks on the facies profile of the ferruginous- siliceous slate formation of KMA

The facies profile of the ferruginous-siliceous slate formation

Facies without ore Facies with ore

Hematitic The main qualitative characteristics of the rocks and minerals Silicate- Silicate carbonate; magnetite Magnetite Specular Magnetite an barren carbonate quartzites iron hematite- specular

magnetite iron (hematite) quartzites quartzites quartzites quartzites

Schists low-ore with

Chemical composition (percentage) Fe general Fe solvent Fe silicate Fez03 Fe0 SiO, Tio,

MnO Ca0 MgO P S C free

A1203

Ratios of themean contents of some SiO, : Fe solution components of the rocks Fe,O,: Fe0

Alzo,: SiO, Tio, : Alzo, Mn : Fe solution P : Fe solution Ca0 : MgO Ti : Y Sr : Ba G e : Fe

6.84 4.95 1.89 3.47 5.62 59.06 0.57 16.71 0.05 0.69 2.34 0.035 O -43 0.45

11.18 0.61 0.28 0.03 0.0084 0.0071 0.29 10.9 1.6 O .o0006

26.25 34.30 17.29 31.10 8.96 3.20 15.70 28.31 19.75 17.77 49.69 44.76 0.24 0.19 2.45 1.94 0.15 0.075 2.18 1.97 3 .O8 2.33 0.075 0.072 0.43 0.225 0.22 0.19

2.68 1.34 0.79 1.59 0.05 O. 043 0.09 0.10 0.0064 0.0021 0.0042 0.0023 0.70 0.84 6.45 3 .o0 3.6 1 .o 0.00003 0.000016

34.59 33.08 1.51 33 -21 14.66 41.83 0.21 0.92 0.07 1.90 1.96 0.072 0.119 0.12

1.18 2.26 0.022 0.23 0.0013 0.0021 0.96 2.40 1.30 0.000024

36.89 35.81 1 .O8 39.09 11.81 40.39 0.15 0.68 0.045 1.74 1.86 0.064 0.053 0.074

1 .O6 3.39 0.017 0.22 0.0011 0.0018 0.94 1.37 1.28 0.000012

39.61 38.67 0.94 48.09 7.62 39.24 0.08 0.41 0.035 1.25

0.048 0.030

-

-

0.94 6.31 0.01 0.19 0.0008 0.0010

1.33 1.20 0.000008

-

92


The facies profile of the ferruginous-siliceous slate formation

Facies without ore Facies with ore

Hematitic The main qualitative characteristics ot the rocks and minerals Silicate- Silicate Magnetite Specular Magnetite carbonate; magnetite quartzites iron hematite- specular an barren carbonate

quartzites quartzites quartzites quartzites

Schists low-ore with magnetite iron (hematite)

The mean contents of move widespread trace elements (percentage) In the rocks M n

Ti V Cu Ni Co Sr Ba G e M n Tn V Cu Ni Ge

In hematite M n Ti V G e Ba

In quartz M n Ti CU Sr Ba

O. 042 0.218 0.020 0.009 0.009 0.0025 0.08 0.05 0.0003 0.31 0.025 0.008 0.10 0.008 0.00065 - - - - -

0.017 0.34 0.031 0.0657 0.034

0.11 0.0284 0.0044 0.0047 0.0058 0.0010 0.040 0.011 0.00045 0.30 0.026 0.0062 0.082 0.041 0.0012 - - - - -

0.049 0.047 0.023 0.013 0.0128

0.067 0.0117 0.0038 0.0030 0.0021 0.0010 0.006 0.006 0.0005 0.020 0.019 0.0040 0.0072 0.0029 0.0010 - - - - -

0.012 0.038 0.011 0.019 0.0069

0.054 0.0075 0.0031 0.0034 0.0023 0.0008 0.0072 0.0053 0.0008 0.019 0.0056 0.0034 0.0053 0.0031 0.00087 0.019 0.007 0.0032 0.00074 0.019 0.0098 0.011 0.007 0.0078 0.0066

0.040 0.0037 0.0027 0.0032 0.0022

0.0064 0.0050 0.00045 0.013 0.0037 0.0031 0.0048 0.0029 0.00065 0.005 0.0063 0.0028 0.00053 0.011 0.0038 0.008 0.0043 0.0064 0.0057

-

0.028 O .O024 0.0018 0.0025 0.0020

0.0040 0.0034 0.00032 0.07 0.0022 0.0021 0.0047 0.0026 0.00047 0.0028 0.0030 O .O023 0.0004 0.004 0.0026 0.005 0.0042 0.0050 0.0055

-

In magnetite

21.12 52.07

20.18 49.42

20.07 47.13

Mean reflectivity for 441-688 n m light waves Pyrite

Magnetite

Microhardness (kg/mm2) Magnetite Pyrite

526.4 1261

531.9 1275

539.2 1278

0.400 0.788

0.410 O. 800

0.410 0.830

Thermo EDS(mv) Magnetite Pyrite

Refractive index of carbonate 1.838- 1.855

1.827- 1.873

1.698- 1.722

1.696- 1.720

1.685- 1.723

1.680- 1.704

Ferruginous-siliceous-clastic formation This formation is not widespread in the KMA region occurring only where beds of ferruginous quartzite in the ferruginous-siliceous formation are overlapped by other rocks with disconformity.

The clastic ferruginous-siliceous rocks of this formation are the products mainly of disintegration, rewashing, and

redeposition of ferruginous quartzite from the ferruginous- siliceous slate formation. In some places, they have pronounced typical clastic texture, but elsewhere clastic texture is obscure.

The differentiation of material into rhythmically alternating iron-rich layers and coarse, sandy, well-laminated ferruginous layers and rudaceous texture, are typical of the formation.

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N. A. Plaksenko, I. K. Koval and I. N. Shchogolev

Summary A comparison of the ferruginous-siliceous rocks in the terrigenous-sedimentary and volcanic-sedimentary sequence enables us to establish considerable distinctions between them. The distinctions relate to the nature of the interlayering of ferruginous and volcanic rocks compared with sedimentary rocks; to the composition and genesis of the rocks that both immediately enclose and separate the ferruginous beds; to the amount of ferruginous material and the composition of the ferruginous zones; to the productivity of the ferruginous suites, i.e. to the amount of ferruginous rock relative to the total thickness of the suites; to the number and thickness of the separate ferruginous layers in suites; to the mineral composition of the ferruginous quartzites and the texture or internal structure of the ferruginous layers; and to the geochemical characteristics and other features.

In general, the Precambrian ferruginous beds of vol-

canic association differ chemically from those of terrigenous association in having more Mn, Mg, Cu, Ca, Co, S, Ba, P and Ni, and less Ge, V and Sr. The ratio of Ti to V generally exceeds 25 in the ferruginous-volcanic association and is less than 25 in the ferruginous-terrigenous association. The ratio of Sr to Ba is greater than 1 in the volcanic association and less than 1 in the non-volcanic association.

Thus, it is quite possible to determine the formational association of the Precambrian ferruginous quartzites.

These facts make groundless the attempts of some investigators to explain the origin of all ferruginous-siliceous rocks of Precambrian age in terms of the volcanic-sedimentary hypotheses.

It is clear that the methods of sedimentary geology, lithofacies and stratigraphic analysis play a great role in the solution of the most complex problems of the origin of the metamorphosed Precambrian ferruginous-siliceous rocks.

Résumé

Les formations de fer siliceux du Précambrien dans la région de l'anomalie magnétique de Koursk (N. A. Plaksenko, I. K. Koval et I. N. Shchogolev)

On peut distinguer quatre types de formations de silex ferrugineux génétiquement indépendants dans une section des formations précambriennes de région de l'anomalie magnétique de Koursk. Elles se remplacent l'une l'autre dans le temps et apparaissent dans des complexes rocheux séparés l'un de l'autre par des ruptures stratigraphiques. Ces formations sont les suivantes (du bas vers le haut) : 1. Ferrugineux-siliceux gneissique ; 2. Ferrugineux-siliceux métabasique ; 3. Ferrugineux-siliceux schisteux ; 4. Ferru- gineux-siliceux-clastogène.

Les deux premières appartiennent à l'Archéen. Elles ont été peu étudiées et leur interrelation n'est pas claire. Les formations de schistes siliceux-ferrugineux se rencontrent avec des interruptions dans la croûte métamorphosée des roches archéennes désagrégées par les facteurs météo- rologiques et qui remontent au Protérozoïque inférieur. C'est la formation la plus productive et par conséquent la mieux étudiée. La formation clastogène-siliceuse-ferrugineuse s'est développée localement. D'après les microfossiles trouvés dans ces roches, elles appartiennent au Protéro- zoïque moyen.

Tous les caractères géologiques des roches siliceuses et ferrugineuses et des roches encaissantes sont caractérisés

dails chaque formation par une individualité spécifique et un caractère qui reflète les conditions de leur formation. Pour le moment, la genèse de deux formations seulement peut être caractérisée avec quelque sûreté : celle des formations schisteuses-siliceuses-ferrugineuses et celle des formations clastogènes-siliceuses-ferrugineuses.

Les roches siliceuses-ferruguieuses (silex ferrugineux) de la première ont une nature colloïde terrigène sédimen- taire ; le fer et la silice, qui participèrent à leur formation, proviennent de la croûte de désagrégation des gneiss archéens, amphibolites et autres roches ferrugineuses. Cela est démontré par tout un ensemble de caractères géochi- miques, lithologiques, pétrographiques, minéralogiques, géologiques et autres.

Les roches siliceuses ferrugineuses de la formation clas- togène se rencontrent sur la surface érodée par l'eau, de la formation schisteuse-ferrugineuse-siliceuse. Elles sont le produit de (( washout )) métamorphosé et de redéposition de quartzites ferrugineuses et des argiles schisteuses qui les composent.

La comparaison de tous les caractères et propriétés dont on dispose pour la recherche sur les roches ferrugineuses des formations volcanogéniques et sédimentaires montre la possibilité d'établir des diagnostics dans lesquels on peut avoir confiance.

Cette communication est surtout consacrée à l'examen de ces caractères et propriétés.

94

Structural-tectonic environments of iron- ore process in the Baltic shield Precambrian

P. M. Goryainov Kola Branch of the Academy of Science of the U.S.S.R.

Within the Baltic crystalline shield, as well as in the other Precambrian regions, iron ores of the cherty-iron series are broadly developed. Acknowledging in the most cases their supracrustal character, the investigators of the Baltic shield Precambrian point out the potentially high correlative properties of cherty-iron rocks. Thus, each regional stratigraphical scheme of the Precambrian involves data concerning iron ore genesis. In this case, in the various schemes, the essential discrepancies arise in the evaluation of events in the lower Precambrian, including the determination of Archaean-Proterozoic boundary, i.e. the period to which the process of ore formation is commonly related. Kratz (1963), in his characteristics of the West Karelian synclinal sub-zone of karelides, points out that within its range the Archaean formations are broadly abundant.

The Archaean formations are presented by gneissose- granites and gneissose-diorites, often migmatized. Supra-

crustal rocks of the Lower Proterozoic attributed to the Gimolian series compose volcanogene and sedimentary- volcanogene formations. According to Chernov (1964) the strata of Kostomuksh deposit comprise tuff breccia, ferruginous, quartzites, leptites, h¿iZZejfi?zta after acid lavas and tuffs, and also amphibolic shales and amphibolites. In the basement of the cross-section granitic conglomerates with plagioclase-amphibolic cement are detected.

The rocks of iron ore strata of the Kostomuksh's deposit form the narrow synclinal zone enclosed between the block ledges of the Archaean foundation.

On the Kola Peninsula the relationships between the rocks of cherty-iron formation and the basement complex are recorded in detail only in the region of KoImozero- Voronya (Fig. 1). Here, in the basement of the formation, the basal granite conglomerates occurring on the oligoclase gneissose granite are distinguishable. The rocks of the cherty-iron formation occur between the blocks of gneissose

FIG. 1. Geological scheme of Kolmozero-Voroiiya river line conglomerates; (b) meta- and ultrabasites of Olenii ridge; (using data of Kharitonov, Garifullin and Maslennikov): I. (Gra- 4. Lenses of magnetitic quartzites; 5. Quartz porphyry, quartz nites): (a) oligoclase gneissose-granites of the basement (in the albitophyre, apokeratophyres ('porphyroides'); 6. Polymictic north-east part granites (gneissose-granites) of Murmansk block conglomerates of Poros suite; 7. Sedimentary rocks of Poros adjoin this line); (b) young microclinic granites; 2. Granitic con- suite (aluminiferous shales, flyschoid bed of quartzites, shales, glomerates; 3. Metabasites: (a) hornblende amphibolites after conglomerates); 8. Faults. diabases, porphyrites, mandelstones with lenses of volcanic

Unesco, 1973. Genesis of Precumbrian iron und manganese deposiis. Proc. Kiev Sump., 1970. (Earth sciences, 9.) 95

P. M. Goryainov

granites in the form of a narrow linear zone spreading over 100 km; from the north the oligoclase biotite porphyroblastic gneisses (‘gneissose granites’) of the Murmansk block adjoin this zone. According to Maslennikov (1969) the complex of supracrustal rocks of the Kolmozero-Voronya zone consists of two suites of different ages.

The lower Polmos suite, with the discontinuity and with basal conglomerates in the base, occurs on the foundation rocks. It is represented by hornblende-amphibolites with relicts of ophitic and porphyritic structures and of amygdaloid textures, In the lower part of amphibolite series some thin lenses of magnetite quartzites occur. In the upper part of the series occur the numerous lenticular bodies of acid volcanites-leptites; quartz-porphyry, quartz albitophyre, with perfectly preserved relict structures of effusive rocks.

In the amphibolite series are found the lenses of ovoid amphibolites. It is supposed that some small concordant lenses of ‘ovoid’ amphibolites, according to their petrological and geological features, are to be considered as metamorphosed volcanic conglomerates. The essentially plagioclasic isolations here are supposed to be the explosive fragments of hypabyssal rocks analogous to the overlying acid volcanites which are cemented by volcanic material.

The overlying Poros suite, with the interruption and with polymictic conglomerates in the base, occurs both on the amphibolites and quartz-porphyry of the Polmos suite and foundation rocks. This section, surprisingly, resembles the Kiruna section taking into account that, firstly, Kiruna volcanites are the deeper magmatic differentiates and, secondly, apatite-magnetitic ore manifestations of the Ki- runa type are absent in the section of the Kolmozero- Voronya zone. A comparison of the two sections is given in Table 1.

The structural relationships between the rocks of cherty-iron formation and basement in the industrial Iman-

TABLE 1

dra lake region on the Kola Peninsula seem to be the most complicated. The main structure in Imandra lake region is defined by an oval arrangement of narrow iron ore bands and corresponding magnetic anomalies. Ferrous quartzites form large lenses wedging out to the sides, stretching up to 4 km and having thicknesses to up 300 m. They are underlain by massive amphibolites with interlayers of biotite gneisses (often with sulphides and graphite), and overlain by aluminiferous gneisses, banded amphibole gneisses, and leptites.

Despite the intense metamorphism for the majority of rock varieties, the relict features of their volcanic origin are determined. Amphibolites preserve the signs of effusive diabases, porphyrites, mandelstones; leptites are the metamorphosed acid and intermediate volcanites, quartz keratophyres, quartz porphyry, dacitic porphyrites.

The rocks of cherty-iron formation in the Imandra lake region are contiguous to coarse-grained, often porphyroblastic gneissoid rocks occurring inside the ovals and also on their outer side. These gneissose rocks are mainly oligoclase-biotitic, with 10-40 per cent quartz content, or hornblende (or pyroxene)-oligoclase-biotitic ones. The presence of microcline is not necessary.

The cherty-iron formation within the main structure of the Imandra lake region and underlying gneissose granites and migmatites were considered to be isoclinally folded. Only in the process of detailed study of samples along the transverse profile did it become evident that the sharp, distinct ‘intense’ gneissosity of migniatized gneissose granites is not the only structural element. It appeared that, in the rocks between the bands specifying the gneissosity, short fragmentary parts appear where biotite is oriented obliquely to the superimposed secondary gneissosity. Orientation of this ‘hatching’ is surprisingly stable. The identified submeridional extension of this weak ‘hatch’ orientation does not coincide with the spatial orientation of the north-western

Kiruna Kolmozero-Voronya

Basement ‘Ancient gneisses’-oligoclase gneissose granite and gneisses.

Granite conglomerates in gneissose and carbonaceous cement. Basic lavas, spilites, diabase porphyrites, metaman- delstones, and amphibolites after them. Lenses of jaspilite, and tufogene rocks. Series of Kirunavaara volcanic conglomerates with fragments of plutonic abyssal analogues of overlying keratophyres. Keratophyres and quartz keratophyres with apatite magnetitic ores, quartz porphyry.

Discontinuity Upper Khauki complex-Vakko series-quartzites, polymictic conglomerates, flyschoid series of phyllites.

Busal level

Oligoclase gneissose granites.

Granite conglomerates in cement of biotite gneisses.

Hornblende amphibolites-massive and banded- with relicts of porphyritic, ophitic structures, and amygdaloid textures, Lenses of magnetite quartzites. ‘Ovoid‘ amphibolites-volcanic conglomerates with acid volcanic fragments(?).

Lenses and bodies of quartz porphyry, and quartz albitophyres in amphibolites.

Poros suite-polymictic conglomerates, aluminiferous shales, quartzites.

96

Structural-tectonic environments of iron-ore process in the Baltic shield Precambrian

secondary distinct gneissosity. Hence, the cherty-iron formation, with its deep and north-western extension of gneissosity and the complex of migmatized biotite gneisses enclosed inside it, refers to structurally autonomous forms of different ages. The contacts of these complexes form the thick zone of blastomylonites where alkaline metasomatites and subalkaline granite rocks are developed.

Thus, it may be admitted that accumulation of volcanogenic cherty-iron formation of the region has taken place under the conditions of differentiation of the crust into foundation blocks already outlined. Between the cracks occurred the eruption of volcanites succeeded by sedimentation.

In the south Pechenga zone of the Kola Peninsula (Allarechensk district) some small bodies of magnetite quartzites occur in amphibolites, aluminous gneisses and other leucocratic supracrustal rocks (apparently leptites). Ferruginous quartzites are here characterized by stratigraphic and lateral zoning exactly the same as in the Iman- dra lake region and possess all the genetic features of this class of supracrustal rocks.

The structure of this region is a combination of isometric or gently oval blocks (more often called ‘domes’) of Archaean granitoid rocks-oligoclase gneissose granites and gneissose granodiorites and, chiefly, linear metabasitic (amphibolitic) series with lenses of magnetitic and magnetite-cummingtonite quartzites or crystalloschists.

Conclusions

Thus, within the Baltic shield to such a level as ‘iron ore’, the Precambrian corresponds to not only qualitatively comparable events, but also to a qualitatively similar state of the Earth’s crust when those events took place. No matter whether ferro-siliceous formations form linear zones of complex structure specified by a combination of isometric foundation blocks and their circumfluent supracrustal iron ore rocks, the cherty-iron formations always occupy another and higher structural position than the rocks composing the blocks of Archaean basement.

These basement rocks are characterized by monot- onous composition: oligoclase biotite gneissose granites, gneissose granodiorites, gneissose diorites, more rarely amphibole or hypersthene gneisses (Karelia and the Kola Peninsula); gneisses of foundation, gneissose granites and gneissose syenites in Finland and Sweden. The cherty-iron formation and the complex of basement rocks often have obvious features of structural autonomy. The rocks of the basement are commonly characterized by the gentle submeridional orientation of structural elements, but younger cherty-iron formations are north-west trending with high- dip elements. The rocks of the foundation may carry traces of two orientations of plane elements: primary and superimposed. If tectonic movements acquire a sharply differentiated character the blocks of the basement lose their shape and elongation; one can see this phenomenon in the Imandra lake region. Here the structural features of the

basement rocks slightly differ from those of the cherty-iron formation. Eventually the blocks of the basement may be converted into narrow elongated bodies which, due to the coming diaphthoresis, cannot be differentiated from the rocks of cherty-iron formation contained between them. Such areas developed for example, to the east from the Monchya-Volchya tundras and in the central Kola area (Chudiavr Lake, Volslipachk, Semb Lake, etc.). They are accompanied by narrow zones, elongated according the strike of the rocks, in which the granulite associations of minerals are distinguished. The appearance of the narrow bands of the granulite facies of the rocks, as well as the different block pattern of the basement, are both, in our opinion, manifestations of the coherent sharply differentiated movements during the process of folding.

According to the present stratigraphical subdivision of the Precambrian, the rocks of the basement and the cherty- iron formation must be referred to the Archaean and Lower Proterozoic respectively, as was accepted for Karelia by Kratz (1963). Subdivision of the Archaean into Upper, Lower, Catarchean is not geologically substantiated. The differentiation of the crust into isometric blocks must be considered as the distinctive feature of Archaean geology. The position of Lower Proterozoic mobile zones was strongly governed by the Archaean blocks’ pattern. Suf- ficient volcanism and sedimentation were confined to the interblock parts, grouping as narrow linear zones.

The volcanic processes in these zones of the Baltic shield began everywhere with extrusions (frequently submarine) of basic lavas-spilites and diabases. Further evolution of volcanism and the iron-ore process connected with it were indicated by the degree of the magma’s differentiation and developed towards either very acid differentiates (e.g. lceratophyric and quartz-porphyritic series of Kiruna with the rich apatite-magnetite ores) or normally acid and intermediate rocks (e.g. the series of quartz porphyries, albitophyres, dacitic porphyries (leptites) of the Imandra lake region and Karelia with iron quartzites). The volcanic process could, al last, end with the same basic volcanites which formed at the beginning of this process, but without the development of large bodies of the ferruginous quartzites (metabasite-amphibolite beds of the Kola Peninsula).

According to Peive (1956) the volcanites of the platforms are characterized by greater differentiation than the volcanites of the orogenic belts. If this is so, the Lower Proterozoic volcaiiites and ores of the Kiruna must reflect the formation of platform conditions. Thus, the Lower Proterozoic mobile zones could differ according to the degree of ‘orogeny’, which was the main criterion for the formational variety of the iron-bearing process which one can see in the Lower Proterozoic of the Baltic shield.

In regarding the isometric block structures as structures of different ages which reflect the primary differentiation of the basement, the formation of interblock mobile zones and intensive deposition in those zones, we must take into account other concepts. A characteristic feature of Early Cambrian structures of all shields was noticed a long time ago; most of these structures are interpreted as domes.

97

P. M . Goryainov

Eskola (1948) considers that the combination of the infra- and supracrustal structures is greatly attributable to the penetration of gneiss domes into sedimentary rocks. Kranck (1959) and Wegman (1930) explained these struc-

tures in terms of granitic diapirism. Kalyaev (1970), ac- centing the principles of granitic diapirism, supposes that sedimentary-volcanogenic rocks of the Ukrainian shield are folded by the increasing granitization.

Résumé

Environnement tectonique et structural des processus de fosmation du minerai de fer duns le Précumbsien du bouclier baltique (P. M . Goryainov)

La différenciation du socle en blocs isométriques a eu lieu avant l’accumulation des formations de silex ferrugineux volcanogène sédimentaire.

Pendant le plissement du Protérozoïque inférieur, ces blocs ont pu perdre leur contour original et au cours du processus de granitisation contemporain les signes d’une indépendance structurale du bouclier et du complexe du minerai de fer sont devenus moins distincts.

La comparaison des complexes de minerai de fer de

Karelija, en Suède centrale et septentrionale, et de la pénin- sule de Kola a montré que, d’abord, elles occupent toutes la même position structurale pénétrant dans le deuxième stade structural et, ensuite, qu’on y trouve de fortes corrélations qui permettent d’identifier des événements géologiques qui pourraient sembler absolument incom- parables.

La présence de silex ferrugineux, calcifère, d‘amphibolites et de gneiss alumineux dans la zone des roches supé- rieures de la croûte dénote un âge K non archéen )) plus jeune de toutes ces roches. Ceci est une des indications qui permettent de différencier les complexes archéens et proté- rozoïques .


CHERNOV, V. M. 1964. Stratigraphy and sedimentary conditions of volcanogene (leptitic) cherty-iron formations in Karelia. Moscow-Leningrad, Nauka.

ESKOLA, P. E. 1948. The problem of mantled gneiss domes. Quart. J. geol. Soc., Lond., vol. 104, no. 4.

__ . 1967. Precambrian of Finland. Docembrii Scandimvii [Precambrian of Scandinavia]. Moscow, MIR.

GEIJER, P. 1931. Berggrunden inom malmtrakten Kiruna- GallivarePajala [Geology of the ore area Kiruna-Gallivare- Pajaia]. Sverig. geol. Unders. Afh., series C, no. 366.

KALYAEV, Y. I. 1970. The problem of connection of granitoid magmatism and folding of the basement. Geotektoniku, no. 1.

KRANCK, E. H. 1959. On the folding movements in the zone of the basement. Geol. Rdsch., no. 46.

KRATZ, K. O. 1963. Geology of the Karelia karelides. Leningrad, Academy of Sciences of the U.S.S.R.

MASLENNIKOV, V. A. 1969. The most ancient Precambrian of the Kola Peninsula (geology and geologic time). Thesis. Moscow, Institute of Geochemistry and Organic Chemistry of the Academy of Sciences of the U.S.S.R.

PEIVE, A. V. 1956. The connection between deposition, folding, magmatism and mineral deposits and tectonics. Glavneishie typy glubinnych razlomov [Main types of deep faults]. Bull. Acad. Sci. URSS, geology series, no. 3. WEGMAN, C. E. 1930. Uber Diapirismus. C.R. Soc. geol. Finl., no. 92.

Discussion

A. M. GOODWIN. In the Canadian shield the oldest rock assemblages of Archaean age contain great thicknesses of clastic sediments, mainly greywacke and shale. Would you care to comment?

P. M. GORYAINOV. I a m familiar with some data on the Canadian shield (Lake Superior and Michipicoten area). I believe that the so-called Archaean supercrystalline rocks, cherty iron rocks included, may be compared with the

Karelian assemblages of the lower parts of the Baltic shield, Strange as it may seem, some difficulties in correlation seem to be due to age-determination figures.

V. M . CHERNOV. What is the stratigraphic relationship between the Olenegorsk suite and the Kola series?

P. M. GORYAINOV. The gneissic iron ore assemblage (Olene- gorsk suite) and the Kola series are not co-ordinate units.

98

Structural-tectonic environments of iron-ore process in the Baltic shield Precambrian

The Kola series is the Archaean undissected basement. The rocks of the gneissic iron ore (Olenegorsk suite) assemblage are Lower Proterozoic.

V. M. CHERNOV. Are there any basal conglomerates between the dome structures which you refer to the basement and iron ore beds? If so, what is the thickness of the conglomerate, its composition and that of the matrix?

P. M. GORYAINOV. The basal conglomerates occur between iron ore beds and the basement blocks (not domes) in the Kolmozero-Voronya region (Kola Peninsula), in the Kos- tamuksh deposits (Karelia) and in Kiruna (Sweden). The conglomerate is composed of rounded granite fragments.

W. M. CHERNOV. Are you certain about the basement of the iron ore beds of Kola Peninsula?

P. M. GORYAINOV. In general, yes. In some cases establishing the basements presents a problem because the rocks have undergone crushing and metasomatism. It is then difficult to distinguish between the basement and the rocks of the second structural stage (the gneissic iron ore beds).

V. M. CHERNOV. What is the attitude of the geologists of the Northwestern Geological Survey to your conception?

P. M. GORYAINOV. M y conception is relatively new and it has not yet been verified. Future investigations will reveal how well-grounded it is.

99

Geology of the Precambrian cherty-iron formations of the Belgorod iron-ore region

Yu. S. Zaitsev Voronezh Geological Prospecting Expedition, Ministry of Geology of the Russian Soviet Federated Socialist Republic, Voronezh (U.S .S .R .>

The Kursk Magnetic Anomaly (KMA) territory is covered with two separate bands of gravimetric-magnetic anomalies, the most intensive ones being produced by steeply dipping iron quartzites. The north-east band consists of the Stary Oskol and Novy Oskol iron-ore regions. The south-west band includes the anomalies of the Belgorod and Igov- Mikhailovsky regions (Fig. 1).

The Belgorod iron-ore region (BIR) is a linear north- west trending zone of crystalline rocks in the south-east area of the KMA (12,000 kin2). The width of the zone ranges from 80 to 100 km and it extends over 150 km.

In the BIR the crystalline rocks lie at a depth of 350- 750 m and are covered with Mezozoic and Palaeozoic sedimentary rocks. The Precambrian basement dips gently (7 m per 1 km) south-west towards the Dnieper-Donetz Basin.

The cross-section of the BIR includes two structural stages divided by the boundary surface with the absolute age of 2,600 & 100 m. y. The lower stage consists of repeatedly dislocated sedimentary-volcanogenic rocks locally granitized and magnetized; the upper stage is mainly composed of sedimentary-volcanogenic rocks forming a distinct structural-facial zone of the Proterozoic geosyncline.

Studies of the crystalline rock sections make it possible to subdivide the BIR iron-ore formation into three types using Plaksenko’s classification (1966): volcanogenic cherty- iron; slate cherty-iron; clastogenic cherty-iron. Each formation characterized a certain period in the evolution of the old basement thus defining the principal stratigraphic unit of the Precambrian.

Volcanogenic cherty-iron type. Rocks of this type differ from others in their chemical composition. In spite of the restricted occurrence in these rocks of the cycle sequence from chlorite and biotite-chlorite slates to stilpnomelane- magnetite slates and quartzites, available data show that silica, ferric iron, manganese and calcium increase. Alu- mina, ferrous iron and magnesium appear to decrease.

Thus, the cherty-iron rocks of the Oboyan-Mikhai- lovsky series of the BIR represent a cherty-iron formation analogous to the Gimo1 unit of the Baltic Shield as well as

to the ferruginous quartzites of the Konsko-Verkhovtsevo series of the Ukrainian Precambrian.

Slate cherty-iron type. Rocks of this iron-ore formation reflect continental deposition prevailing during the early Proterozoic time of geosynclinal development. The formation is a unit of the Kursk metamorphic series. In monocline sections it is characterized by facial changes of ferruginous quartzites deposited under the conditions of shallow sea (slightly ferruginous coarse-banded and silicate rocks) grading into deep-sea facies (finely banded magnetite and mica- iron quartzites). References on its compositionand structural elements are given in the reports of Plaksenko, Chaikin, etc.

Stratigraphic division of the ferruginous quartzites of the BIR is now available only in the Yakovlevo and Gos- tishchevo deposits, where Chaikin and Rusinovich counted up to seven horizons of quartzites. The odd-numbered quartzites correspond to martite (magnetite)-hydro-liema- tite rocks and ferriferous silicates; even-numbered ones to martite (magnetite) rocks with mica-iron facies.

The data show that the quartzites grade into the underlying rocks.

The slate cherty-iron formation of the BIR has evident facial changes in submeridional and north-east directions. The thickness of the mica-iron facies decreases and locally one or more primary oxide facies are absent. Ferrous, ferrous carbonate and silicate (magnetite, silicate quartzites and ferruginous silicate slates) facies prevail.

Facial changes of deep sea sediments to shallower ones extending eastward and north-eastward show that the near- shore line of the depositional basin of the BIR was somewhere to the east of the Prokhorovsko-Korochansk gravimetric and magnetic anomalies. In the same direction, the metamorphism of cherty-iron rocks tends to increase from metamorphic slate (phyllonite) to gneiss.

Clastogenic cherty-iron type. The formation of this type is a part of the Oskol series and generally occurs everywhere. Cherty-iron rocks are metamorphosed re-deposited sediments of the slate cherty-iron formation and are composed of conglomerates, gritstones, debris of ferruginous quartzites, metasandstones and slates.

Unesco, 1973. Genesis of Precambrian iron and manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) O 1

Yu. S. Zaitsev

FIG. 1. Generalized map of the Precambrian of the Voronezh crystalline shield. (Zaitsev and Bogdanov, 1969.) 1. Upper Proterozoic microline and plagioclase-microline granite; 2. Ra- pakivi granite, syenite, granosyenite; 3. Subplatform gabbro- norite, olivine gabbro-norite, olivine gabbro-dolerite, gabbro- dolerite, gabbro-diabase; 4. Middle Proterozoic late orogenic basite-hyperbacite: peridotite, pyroxenite, dunite, olivinite as well as their metamorphosed differentiates, gabbro, gabbro- norite, norite, gabbro-diorite, olivine gabbro and their veined differentiates; 5. Syntectonic plagioclase granitoids; 6. Early and

Along the strike these rocks grade into cherty-iron rocks locally called ‘conglomeratic’ or ‘nodular’ quartzites. Martite mica-iron quartzites average about 50 m in thickness, ‘nodular’ ones being up to 30 m thick.

In contrast to the quartzites of the Kursk series, the ore bands of coarse martite mica-iron quartzites of the Oskol series commonly have sand-like structure. The sandy appearance is due to oval and octahedral martite grains of about 1.5 mm in flaky bands of hematite. Semi-ore bands are mainly composed of oval, lenticular and round jasper debris of brick-red and pink colour averaging 2-3 mm occasionally 8 mm.

Flaky material of iron glance is commonly interstitial and does not form separate bands, as it may be seen in the quartzites of the slate cherty-iron formation. Occasionally jasper debris occupies 25-40 per cent of rock volume. The

synorogenic gabbro, gabbro-diorite, gabbro-amphibolite rare meta-ultrabasite; orthogneiss; 7. Late Svekofeno-Karelian folding (1,700-2,000 m.y.); 8. Early and Late Svekofeno-Karelides (1,700-2,600 my.); 9. Early Karelian (Kursk) folding; 10. Belo- moride analogue influenced by the Late Svekofeno-Karelian folding; 11. Presvekofeno-Karelian granitoids; 12. Belomoride analogue influenced by the Early Karelian (Kursk) folding; 13. Presvekofeno-Karelian median massif; 14. Ancient cores of Presvekofeno-Karelian folding; 15. Deep faults; 16. Bounds of the Belgorod iron-ore region of the KMA.

rocks are composed of iron hydroxides evenly dispersed or in patches. ‘Conglomeratic’ quartzites have lenticular echelon-like debris of hornfels in silicate cherty-iron rocks.

Conglomerates and gritstones grading into metasandstones, metasiltstones and slates are abundant in the cross- section of the clastogenic cherty-iron formation. The bands of conglomerates range from 5 to 30 m. When products of the Kursk quartzites erosion, the conglomerates consist of quartzite debris (mainly semi-ore rocks) cemented by quartz, hydromica and ferruginous material.

Psammitic and psephitic components usually contain jasper hornfels debris, granitoid and quartz pebbles. They are commonly of elongated, lens-flattened, round and irregular form. Typical magnetite (martite) quartzites of the slate cherty-iron formation as well as ore quartzites are absent in the debris material.

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Geology of the Precambrian cherty-iron forniations of the Belgorod iron-ore region

The conglomerates, gritstones and metasandstones grade into martite-mica-iron-sericite slates and metasiltstones. The contact of intergrading rocks with the underlying psephitic ones is not sharp. Martite mica-iron sericite slates and metasiltstones range from 1 to 20 m in thickness, which is proportional to the thickness of conglomerates and metasandstones.

Thus, according to its geological and sedimentary features, the clastogene cherty-iron formation of the BIR corresponds to the carbonaceous terrigenous formation of the Upper Krivoyrog series of the Ukrainian crystalline shield.

On the basis of the geological evidence, the BIR is a complex multistage unit, Comprising extensively dislocated and variously metamorphosed Precambrian rocks of the Belgorod synclinorium. The rocks of the basement include two structural stages. The contact between them is well- defined by folding, magmatic activity, regional metamorphism and old weathering. These stages characterize primary geosyncline and geosyncline cycles of the Precambrian.

Stratigraphic division of the crystalline rocks involving lithological and stratigraphic methods as well as the formation analyses must also be based on the main geological features of the iron-ore formations. These features make it possible to develop reliably correlated stratigraphic sections. The data obtained show that a great period of evolution corresponds to a certain geological formation with specific features of sedimentation and volcanic activity.

The earliest stage-the Archaean-is characterized by a volcanogenic cherty-iron formation associated with volcanogenic spilite keratophyre rocks. The formation was deposited in a changing redox environment depending on the

volcanoes' position in the basin of sedimentation. Ferrous iron is more abundant near the shore line (carbonaceous and silicate forms).

In the open sea environment ferric iron prevails. Fer- rous iron material is much less abundant with decreasing content of alumina, magnesium and titanium. Absence of hematite in this formation within the BIR shows a relatively low oxidizing environment and an abundance of volcanoes. In this respect our results coincide with those of N. Strakhov concerning iron reduction in cherty ores of exhalation type. Silicate cherty-iron rocks of the BIR may be referred to this type.

In the Lower Proterozoic, when the development of the mobile zone began, the slate cherty-iron rocks were formed. They consist of the intergrading metasandstones, slates and ferruginous quartzites. Environmental conditions of this formation are characterized by the changes of the ore and non-ore facies (from silicate and carbonate to magnetite and hematite) which reflect the increase of the depth from the shore line.

The changes in the structure of the geosyncline occurred in the Lower and Middle Proterozoic time and are characterized by the formation of clastogene rocks involving the material of disintegration and redeposition of the metamorphosed rocks of the Kursk series.

The subdivision of the BIR iron-ore formation into different genetic cherty-iron types corresponding to the stages of tectonic and magmatic activity probably reflects the general tendency of the iron ore formations in the early Proterozoic. This tendency may spread over the other regions of the Voronezh crystalline massif and other nearby regions in the East-European platform.

Résumé

Géologie des formations précambriennes de fer. siliceux dans le gisement de Belgorod (Yu. S. Zaitsev)

1. La région du gisement de fer de Belgorod, située dans la structure précambrienne de l'anomalie magnétique de Koursk, est une formation complexe à plusieurs étages constituée par les structures cristallines fortement dislo- quées et différemment métamorphosées qui forment le synclinorium de Belgorod. Les roches basiques y forment deux étages structuraux, séparés l'un de l'autre par un étage de plissement, par des manifestations d'activité magmatique, par une ancienne croûte de désagrégation. Ces étages repré- sentent deux périodes de développement : la période pro- géosynclinale et la période géosynclinale inhérente.

2. La classification stratigraphique des roches cristal-

lines vise à clarifier les particularités et les principes de déve- loppement des formations de fer qui sont les plus typiques dans la section précambrienne.

I1 se trouve qu'à grande échelle chaque étage correspond à un type bien défini deformation de minerai de fer pré- sentant des particularités spécifiques des processus de sédi- mentation et d'activité volcanique séquentiels dans le temps.

3. Trois étages de formation de fer siliceux peuvent être distingués : (a) fer siliceux volcanogénique ; (b) fer siliceux schisteux ; (c) fer siliceux clastogène.

4. La séparation des différents types génétiques de formations de fer de la région de Belgorod en complexes indi- viduels superficiels de la base cristalline semble refléter les principes généraux de la formation du minerai de fer dans le Précambrien ancien.

103

Iron-formation and associated manganese in Brazil’

J. Van N. Dorr II United States Geological Survey Washington, D.C. (United States)

Introduction This paper summarizes data on the major iron-formations2 in Brazil, the related manganiferous deposits, and some unrelated ones, and thus provides a frame of reference for the more detailed discussions of individual areas to be presented to this symposium by my colleagues Drs Barbosa, Grossi, Scarpellí and Tolbert. Unfortunately, because we were separated by great distances, it was not possible to consult during the preparation of our individual papers. I trust that no discrepancies greater than customary between geologists will appear. At the request of D r Xngerson, I shall also discuss a manganese-iron deposit of probable Cambrian and Ordovician age because it seems quite similar to, although richer than, most Precambrian deposits aiid, being essentially unmetamorphosed, may throw some light on those older deposits. In these discussions, I shall try to approach the deposits from the point of view of their sedimentary environments rather than from their detailed mineralogy, economic potential, epigenetic alteration or metamorphic history, although of course these factors cannot be ignored.

That the banded iron-formations enclosed in Precam- brian sedimentary and metamorphic rocks are sedimentary in origin seems so widely accepted today by geologists that thare is no need to labour the point. The acceptance of the concept of sedimentary facies in iron-formation defined by James (1954) is also widely accepted, although epigenetic processes such as weathering, metasomatic and hydrothermal activity and metamorphism may obscure the original nature of these facies. I believe that everyone also accepts the evidence that the iron-formations were dominantly chemical sediments, although in some places contaminated by detrital debris.

So far only two facies of iron-formation have been found in Brazil, the oxide and the carbonate. Of these, the oxide is by far the more widespread both in time and space. It is of course quite possible that much more carbonate- facies iron-formation will be found; the rock oxidizes at the surface to a weathering product almost indistinguishable

from that of oxide-facies iron-formation. Only by exploration below the zone of oxidation can this rock be definitely identified. Until the work of Gair (1962) and Matheson (1956), the presence of sideritic iron-formation in Brazil had not been established.

It is not as widely recognized that the manganiferous sediments also were deposited in sedimentary facies similar to, and to some extent parallel to, the sedimentary facies of the iron-formations, although commonly they are not interbanded with chert. To m y knowledge, the sulphide-facies has never been reported in sedimentary manganese deposits, although carbonate-facies and oxide-facies are very common. Silicate-facies exists if one accepts the premise that braunite may be a primary sedimentary or diagenetic mineral. Other manganese silicate minerals are commonly regarded as hydrothermal, igneous or metamorphic minerals, although the line between hydrothermal and sedimentary deposits becomes quite vague in volcanogene sedimentary manganese deposits. Braunite, bementite, and neotocite may be primary sedimentary minerals in such cases. Fortu- nately, in Brazil we do not have to consider such borderline deposits, for they have not yet been identified; here we have only carbonate- and oxide-facies manganiferous deposits.

Precambrian rocks crop out in perhaps half the area of Brazil, or an area of about 4 million km2. These shield rocks are found from the northernmost to southernmost extremities of the country and from the easternmost areas to those farthest west. Most of these Precambrian rocks are metasedimentary; the task of unravelling their relative and absolute ages is just beginning and it will be many years before all the complexities are resolved and accurate correlations made. Deep weathering in much of Brazil and extreme difficulty of travel in the forested areas have made detailed geologic work difficult and for this very reason

1. Publication authorized by the Director, United States Geological Survey.

2. Iron-formation was defined by James (1954) as a ‘chemical sediment, typically thin-bedded or laminated, containing 15 per cent or more iron of sedimentary origin, commonly but not necessarily containing layers of chert’.

Unesco, 1973. Genesis of Precumùriun iron und r?iungunese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 1 OS

J. Van N. Dorr Il

Brazil is still a country of major undiscovered resources. Brazil's major known iron-formation deposits are those

in Minas Gerais in central Brazil (Fig. i), in Pará in the Amazon area and in Mato Grosso. Smaller deposits are known in Ceará, Bahia and Amapá. Important manganese deposits are found in Minas Gerais, Bahia, Goiás (not shown), Amapá, and Mato Grosso. The deposits in Mato Grosso are Cambrian and Ordovician in age, the others are Precambrian. There is no reason to suppose that all the iron or mangagnese deposits of significant size have been discovered, and care should be exercised in projecting patterns from the deposits now known.

O'

12'

24'

Deposits in Minas Gerais The only carbonate-facies iron-formation yet known in Brazil is in the Nova Lima Group of the Rio das Velhas Series in Minas Gerais, a eugeosynclinal suite of sedimentary rocks not less than 5,000 m thick dated as being older than 2,700 m.y. (Aldrich et al., 1964) by Rb-Sr analysis of muscovite formed in a contact aureole. The most complete description of these ferruginous rocks is that by Gair (1962). They are typical banded metachert-siderite containing varying quantities of magnetite. The lenses of this rock are relatively thin, ranging to perhaps 75 m but generally less, and ranging in length from a few tens of metres to perhaps 10 km or more. Commonly, they are only a few kilometres to a few hundred metres in strike length. The enclosing

FIG. 1. Map showing distribution of major iron (Fe) and manganese (Mn) deposits in Brazil and part of Bolivia.

106

Iron-formation and associated manganese in Brazil

rocks are now phyllite and schist; the original sediments were argillaceous sandstones, claystones, locally carbonaceous, tuffaceous clays and tuffs and probably some extrusive volcanic rocks. The original volcanic rocks are thought to have been mafic or intermediate rocks. The iron- formation beds are interbedded with phyllite aiid may grade into ferruginous phyllite, carbonaceous phyllite or white quartzite (metachert) along strike.

Methane has been detected in the same group of rocks in the nearby Morro Velho gold mine, probably derived from the organic material in the carbonaceous phyllites, which are widespread. The evidence for a reducing environment during deposition of the carbonate-facies iron-forma - tion seems strong.

The carbonate-facies iron-formation in this region has not been altered to significant bodies of iron ore of usable grade either by metasomatism or by supergene enrichment, even though very large bodies of high-grade iron ore of both types of origin are found in younger oxide-facies iron-formation nearby. This may be because carbonate-facies iron- formation generally forms small lenses, has most of its iron in the divalent form and is thus fugitive under weathering conditions compared to trivalent iron in the oxide-facies, or because the carbonate-facies iron-formation is more plastic than the relatively brittle oxide-facies formation, thus reducing permeability, or a combination of factors.

A discontinuous zone of manganese silicate-carbonate rock is found in a belt of rocks correlated with the Rio das Velhas Series; the belt stretches some 200 km north-east from São João del Rey. These metasedimentary manganiferous rocks are enclosed in graphitic phyllite, phyllite, schist and amphibolite. It is not yet known whether the amphibolite is metasedimentary or metavolcanic in origin. The metasediments are similar to the rocks that contain the carbonate-facies iron-formation, but the zone of manganese silicate and carbonate cannot be confidently correlated stratigraphically with the zone containing the iron-formation.

The manganese silicate-carbonate is believed to be the meíamorphic equivalent of original silty manganese carbonate beds deposited in a reducing environment, attested by the uniform but small content of free carbon in the ore (Dorr, Coelho and Horen, 1956) and in the wall rocks. Where the original manganese carbonate content was high and the sediment was relatively uncontaminated by detrital material, metamorphism did not form abundant silicate minerals; where there was a large admixture of silt and clay, spessartite, rhodonite, and many other silicates were formed at the expense of rhodochrosite. Weathering has produced large masses of manganese oxide from the manganese carbonate lenses, but the silicates did not yield significant quantities of oxide ore on weathering.

Thus, these carbonate-facies manganese- and iron-formations were formed in a eugeosynclinal environment, iron and manganese were separated in space during sedimentation, and significant bodies of the two types of deposit were not laid down together, although both are widespread.

There is no good evidence as to the source of the manganese and iron in the Rio das Velhas rocks. Although

volcanism was active in the general region, as attested by the tuffaceous sediments and probable extrusive rocks in the suite, it cannot be proved that these, or thermal waters emanating from volcanic sources, have any genetic connexion with the manganese or iron. The enormous thickness of Rio das Velhas clastic rocks is good evidence that a large land mass was being eroded during deposition of the sediments, and this might have furnished ample supplies of these elements to the basin of deposition.

Spatially within a few hundred metres to a few tens of kilometres from these Rio das Velhas iron- and manganese- formations, but separated from them by vast reaches of time, are the manganese and iron deposits of the Minas Series. The Minas Series overlies the older rocks with profound angular and erosional unconformity; an orogenic event separates the two series. Much of the Minas Series was laid down in a miogeosynclinal or platform environment, and the unit is about 3,500 m thick (Dorr, 1969). The age of the rocks is still uncertain; they were probably deposited between 2,200 and 1,350 m.y. ago.

The iron-formation of the Minas Series will be described in some detail by Professor Barbosa; I shall merely sketch in some of the more important points concerning the relation of manganese to this iron-formation, which crops out widely in the Quadrilátero Ferrífero, a large iron-rich area centring at about 20°15'S., 37O3O'W.

The principal iron-formation, known as the Cauê Itabi- rite, is typically made up of banded quartz and hematite; the pre-metamorphic rock was chert and hematite, and magnetite was probably present locally in significant amounts. The formation was continuous for a minimum distance of 150 km in an east-west direction and 100 km in a north-south direction; the original thickness was probably about 250-300 m. Significant intercalated clastic sediments have not been found, although locally the rock contains some clay. Much more important from the view-point of sedimentary environment is the presence of dolomite interbedded with the iron-formation; in some cases it makes a threefold layering with the quartz and hematite, and in others it substitutes for the quartz bands. Where dolomite is abundant, magnetite is much more common; it is not certain whether this is a diagenetic or a metamorphic feature. Beds of dolomite approximately 1 m to 20 m thick may also be intercalated in the iron-formation. Gradationally overlying the iron-formation is a thick formation largely composed of dolomite marble, dolomitic phyllite, and minor dolomitic iron-formation. Iron carbonate minerals have not been found in the dolomitic iron-formation except in very minor quantity as epigenetic minerals, thus the rock cannot be considered carbonate-facies iron-formation even though it may contain much carbonate. The iron is predominantly in a trivalent state.

Although the manganese content of most of the iron- formation is very low, lenses of manganiferous rock enriched by supergene processes into usable ore deposits range in size to as much as 5 million tons and contain between 30 and 48 per cent Mn. Individual lenses are more than 1 km in strike length in very few cases; normally they are a few

107

J. Van N. Dorr II

hundred metres long, but a single stratigraphic zone containing manganiferous lenses may be more than 10 km in length. The manganiferous lenses are usually less than 3 m thick. They are more common in the upper part of the formation, as are the dolomitic iron-formations and dolomite lenses. The manganiferous lenses are in places, although not invariably, closely associated with the presence of dolomite in the rock. In contrast to the manganiferous beds in the Rio das Velhas Series, these manganiferous sedimentary rocks are within the iron-formation itself.

Unfortunately, because of the deep weathering and the fact that the enclosing rocks are too soft to mine to great depths without excessive timbering, exploration of the deposits lias not gone to a depth at which unaltered rock is found. Thus, the tenor, mineralogy and the general character of the original rock from which these deposits formed by supergene concentration are unknown. Smaller deposits in the dolomite overlying the iron-formation may give a clue as to the origin of these deposits; they are known to have been derived from manganoan dolomite containing from 5 to 40 clarkes of Mn. It is quite probable that some of the deposits in the iron-formation were also derived from manganoan dolomite, either interlayered in the iron-formation in thin bands, as in dolomitic itabirite, or as somewhat thicker beds. It is also probable that some manganese oxide was deposited synchronously with iron oxide in the iron- formation without dolomite, and was concentrated during weathering.

The p H during the deposition of the manganiferous iron-formation must have fluctuated slightly on either side of 7.8, the limestone fence of Krumbein and Garrels (1952), as shown by the intermittent deposition of dolomite. The Eh may have been around O during deposition of the manganoan carbonate, as the oxidation potential needed to convert ferrous iron to ferric iron is much less than that required to oxidize divalent manganese to tri- or quadrivalent manganese, and the iron would oxidize first (Mason, 1949). The manganese would very possibly be deposited as manganoan dolomite or limestone. When the Eh and the p H were higher, the manganese might well have been deposited with the iron in oxide form. Dolomite with included primary manganese oxide is not known in fresh rocks in the region.

Although the rocks have been metamorphosed to the greenschist facies and higher, manganese silicates are rare in the Minas Series, having been found only in manganiferous phyllite, not in dolomite or iron-formation.

Both the oxide-facies iron-formation and the manganiferous rocks of the Minas Series have been enriched to ore grade by supergene enrichment (Dorr, 1964). High-grade hematite deposits of great size and extreme purity have been formed by metasomatic enrichment during the last metamorphism that affected all the Precambrian rocks of the region (Dorr, 1965).

The source of the manganese and iron in the Cauê Ita- birite and the overlying dolomite cannot be proved. It is very probable that they were derived from the weathering of the Rio das Velhas Series, which the Minas Series trans- gresses; an ample source of both elements is present here

and it is known that rocks of the Rio das Velhas were pene- plained before Minas time. Volcanic rocks are rare in the lower and middle Minas Series; in the upper Minas Series they overlie hundreds of metres of nonvolcanic miogeosynclinal sediments that had been deposited on the dolomite overlying the Cauê Itabirite.

Deposits in Bahia

In the Urandí district of southern Bahia, centred about 14'50' S., 42"40' W,, iron-formation and economic manganese deposits are also known. The area is remote and a detailed geologic map of the district as a whole has yet to be completed. Exposures are very poor and weathering is intense. It is understood that the rocks have been metamorphosed to a somewhat higher grade than those in the Quadrilátero Ferrífero, although the argillaceous rocks are still classified as phyllite. Jacobsite is an ore mineral (Ri- beiro, 1966) and the ore zones with jacobsite can be traced by magnetometer (Ribeiro and Ellert, 1969). The enclosing rocks have been correlated with the Minas Series, although this correlation is not absolutely certain.

In any case, the Urandí manganese deposits are lenticular, some are closely associated with iron-formation, and both the iron-formation and most of the original manganiferous sediments seem to have been oxide-facies. Recause some of the ore is extremely pulverulent, similar to some of the manganese ore derived from dolomite in the Quadri- látero Ferrífero, it is possible that some of the ore was derived by weathering of manganoan dolomite or manganiferous phyllite. N o volcanic rocks contemporary with the original rocks have been reported in the region.

Farther north in Bahia, in the regions of Nazaré and Jacobina, manganese oxide deposits derived by supergene enrichment of manganiferous phyllite have produced a small tonnage of commercial ore, but these poorly exposed and superficially explored deposits throw little light on the origin of the manganese, It seems probable that the manganese was an oxide sediment syngenetic with the siltstone or mudstone from which the manganiferous phyllite was formed. The sedimentary suite was probably miogeosynclinal or platform in depositional environment.

The State of Bahia contains many lenses of iron-formation, some of large size, in the highly metamorphosed Precambrian metasedimentary rocks. None of these lenses have been thoroughly studied and the geologic environment is not clear. Metamorphism has transformed much of the iron into rather coarse-grained magnetite and, although it seems probable that the original facies was oxide, this cannot be confidently affirmed in the present state of our knowledge.

'

Deposits in Amapá

The Serra do Navío is a major manganiferous ore-producing district in the Territory of Amapá, north of the

1 O8


Amazon at about 0°59'N., 52"05'W. These deposits will be described for this symposium by Dr Scarpelli, who has studied them in detail in extensive excavations and in many thousands of metres of drill core. The area has few natural outcrops, has been deeply weathered, and is covered by dense rain forest. The structure is highly complex.

The rocks of the Serra do Navío district belong to the Serra do Navío and the underlying Jornal Groups of the Amapá Series. The age of these dominantly metasedimentary rocks is unknown, but they are older than 1,800 m.y. (Almeida et al., 1968). They may well be contemporary with the Imataca Series, the rocks containing great oxide- facies iron-formations and unimportant manganese deposits in Venezuela.

Also included in the Amapá Series are metasedimentary rocks of the Santa Maria Group, containing important iron- formation some 85 km from Serra do Navío. The relative stratigraphic position of the Serra do Navío and Jornal Groups and the Santa Maria Group is not certain; they could be contemporaneous, or the Santa Maria Group could be older than the others, as suggested by Nagell (1962). In the rain forest such matters are not easily clari- fied. The Santa Maria Group iron-formation is oxide-facies.

The Jornal Group is largely amphibolite, considered by Nagell (1962) to be a metasediment, but by Scarpelli (1966) to be an ortho-amphibolite. Work is in progress to determine the origin of this rock, the most consistent in composition and the most widespread of the rocks of the Amapá Series. Scarpelli informs me that he now considers the rock a para-amphibolite (written communication, 1970).

Overlying the Jornal Group is the Serra do Navío Group. Whether or not the contact is conformable is not clear. The Serra do Navío Group consists of dominant quartz-biotite-garnet schist. Scarpelli has subdivided this into three facies: quartzose, biotitic, and graphitic. The manganiferous rocks are in the graphitic facies and, where pure, consist of rhodochrosite marble with very minor rhodonite. Calcite marble lenses are also present. The graphitic facies may contain as much as 20 per cent graphite in the schist.

Rhodochrosite was clearly the original manganese mineral in the rock; as in the Morro da Mina deposit in Minas Gerais and many similar deposits elsewhere in the world, it recrystallized during metamorphism, but new minerals were not iormed in the pure rhodochrosite lenses. Where the rhodochrosjte was mixed with clay and other detrital sediments, a suite of manganese silicate minerals formed which will be described by D r Scarpelli. Silicate minerals are abundant on the walls of the carbonate lenses. The rhodochrosite in the protore ranges from 2 to 99 per cent.

Scarpelli has shown that tlie three facies of the Serra do Navío Group were deposited cyclically; as many as three cycles are present in some localities, each representing tens of metres of sediments. Nagell (1962) suggested that the original sediments were deposited in a euxinic environment, as indicated by the high carbon content of the enclosing rocks and also by the relatively high concentration of arsenic in the oxide ore derived from the protore. Scarpelli

further suggests, and I concur, that the sediments were deposited in an unstable shelf or lagoonal environment.

In the general area of the Serra do Navío manganese deposits, patches of ferruginous laterite cover considerable areas of the high plateaux. I do not know whether these represent lenticular iron-formations in the bedrock or whether they are merely the surface concentration of hydrated iron over iron-rich igneous or metasedimentary rocks expected under these climatic and physiographic conditions, as are found in so many parts of West Africa. The essential point is that the original carbonate-facies manganese sediments are separated spatially, although not necessarily in time, from iron-formation, as was found to be the case in Minas Gerais.

Deposits in Pará

In the State of Pará near 6"S., 51"20'W., an extensive area underlain by thick iron-formation was recently found by D r Tolbert, who will describe the deposit. I had the privi- lege of visiting tlie region in 1968, but much more has been learned about it by drilling and surface geology since that time. The area is one of the most remote and difficult to traverse of any in the rain forest of Brazil and the rocks are deeply weathered, thus we may expect years to elapse before we know most of the details of the geology. Judging from what I could observe, the iron-formation was oxide-facies. According to Dr Tremaine (oral commuiiicatioii, 1970), it is associated with quartzite and underlain by conglomerate. I saw 110 rocks that appeared to be volcanic, but outcrops are rare and scattered and the region vast. From this association the iron-formation seems to have been deposited in a platform environment. Manganese is present in the general region, but we do not yet know whether or not it is in the same stratigraphic unit as tlie iron-formation. W e may confidently expect a notable increment of knowledge concerning these matters in the future.

Deposits in Mato Grosso and adjacent Bolivia

One of the largest and highest-grade known deposits of iron- formation and of unenriched sedimentary manganese oxide in the world is found in a geologically homogeneous area astride the boundary of Brazil and Bolivia near latitude 19"15'S. In Brazil this is in the state of Mato Grosso. Almost all the known manganese is on the Braziliaii side of the border; the iron-formation with which it is ínter- stratified is found in enormous quantity on both sides of the border. The best known area is in Morro do Urucum (Dorr, 1945).

The economically interesting rocks are in the Band' Alta Formation of the Jacadigo Series. The age of these rocks is not certain, as diagnostic fossils have not been described. For many years the rocks were considered to be Silurian in age; some geologists, including Shatskiy (1954)

1 o9

J. Van N. Dorr II

and, for many years, myself, believed that a late Precam- brian age was more probable. Recent regional work by Almeida (written communication, 1969) led him to attribute a Cambrian and Ordovician age to the Jacadigo Series. I have heard at second hand that a United Nations geological team working in Bolivia found brachiopods in the iron- formation, but have not seen this in print and cannot vouch for the accuracy of this statement. Conglomerates in the Jacadigo Series contain cobbles of a nearby granite, which has been radiometrically dated at a minimum age of 888 m.y., and of metamorphosed rocks of the Corumba Series, a metamorphism dated radiometrically as about 550 m.y. (Almeida and Hassui, unpublished data). For these reasons I concur with Almeida's assignment of these strata to the Cambrian and Ordovician.

Unlike all the other ferruginous rocks described above, the Jacadigo Series is only gently and slightly folded and is unmetamorphosed, The degree of weathering and of supergene enrichment is very minor indeed; the rocks are somewhat leached at the surface, but mechanical erosion here dominates over chemical and fresh rock is close to, or at, the surface in most exposures. The iron-formation is very resistant to erosion and stands in high buttes and mesas bounded by steep slopes and nearly vertical cliffs.

The Jacadigo Series consists of a basal formation some 350 m thick composed of clastic rocks, dominantly coarse- grained and dominantly arkosic, together with local channel sands and puddingstone conglomerates and some apparently lacustrine beds, the whole cemented by calcium carbonate. Crossbedding attests continental and near-shore conditions. Gradationally overlying this thick clastic formation is a formation about 100 m thick made up largely of jasper, massively bedded, a cliff-forming unit. Above the transitional zone, the formation contains very little clastic material, although much of the jasper is colitic. This material is not banded and, although the iron content is perhaps 20 per cent, it could not be called a banded iron-formation.

Above the jasper formation, with rather abrupt but completely conformable contact, lies the Band' Alta formation. It is not less than 350 m thick, composed of banded hematite-jasper rock containing lenses and beds of manganese oxide and of detrital rocks, some quite coarse. The iron-formation consists of alternating bands of quite pure, very finely crystalline blue hematite and of red jasper. The hematite bands are generally 1 c m or less in thickness but reach 10 c m locally; the jasper bands are slightly thinner but may range to several tens of centimetres in thickness, though this is rare. Except where contaminated by detrital material both the hematite and jasper bands are essentially monomineralic.

Beds of detrital material and of manganese oxide are intercalated within the banded iron-formation. The detrital material ranges from well-sorted medium-grained sandstone to poorly sorted conglomeratic rock with boulders as much as 30 c m and more in diameter. The boulders appear to be granitic; near the surface all are so altered by through- passing waters that the identification is not secure, for they now consist of iron-stained clay and quartz. Much of the

coarser material is quite angular. The medium-grained sandstone is, in contrast, moderately rounded. The detrital material is in beds ranging from a few centimetres thick to zones 30 m thick in Morro do Urucum. None of the coarser detrital beds have great lateral extent.

The manganese oxide (cryptomelane) beds range from 1 c m to more than 6 m thick. Two main beds are known in Morro do Urucum, both in the lower part of the formation. The lower and most widespread bed averages almost 2 m in thickness, the upper bed perhaps 1 m. An unknown, but probably considerable, part of the manganese beds has been removed by erosion; the part of the main bed that remains is not less than 5 km2 in extent. The upper bed in Morro do Urucum is about 3.3 km2 in extent. Other manganese oxide beds are known in the region in this formation; none are apparently as thick or widespread as the main bed of Morro do Urucum.

The manganese oxide beds are almost all intercalated between clastic beds in the iron-formation. Commonly these clastic beds are only a few centimetres to tens of centimetres thick; the clastic beds below the manganese oxide beds are well-sorted medium-grained sandstone. In many places the overlying beds are similar but locally the overlying clastic bed contains large boulders and cobbles in a poorly sorted medium- to fine-grained sandy and clayey matrix; in such areas the overlying clastic beds may be more than 1 m thick. Detrital grains occur widely scattered in the manganese bed; they are rare in the iron-formation. Average analyses of the iron-formation and of the manganese oxide lenses in Morro do Urucum (Dorr, 1945), are given in Table 1.

TABLE 1

Manganiferous beds Banded hematite average (%) average (%)

Fe 11.1 (range 8-16) 56.9 (range 48.7-62.1) M n 45.6 (range 39.4-50.7) 0.08 (range 0.005-0.60) Si02 1.25 17.3 '%,O, 1.74 0.65 MgO 0.13 0.06 Ca0 0.20 0.06

Iron-silica ratio 8.8: 1 3.29: 1 Iron-manganese ratio 0.24 1 710: 1

Kz0 3.52 0.20

A number of complete analyses of the manganese ores and the iron-formation made by the United States Geologi- cal Survey laboratories are also quoted (Table 2) from Dorr (1945).

All the data so far given apply to Morro do Urucum. Investigations, almost all unpublished, of adjacent areas in Brazil and Bolivia have been carried on in recent years by many geologists and engineers and I have been able to learn of some of the results. Some of the information is particularly important in giving clues to the sedimentary environment in which the Jacadigo Series was deposited.

About fifteen years ago the Bolivian Government spon-

110

J. Van N. Dorr II

200 million tons of manganese ore were present before erosion; I now consider this estimate, made before the discovery of the ore in Rabicho and Mutum, to be extremely conservative, perhaps by an order of magnitude.

Although Shatskiy (1954) suggested that the iron and manganese might have been derived from the weathering of the older iron-formations of the Precambrian shield, this hardly seems probable, as the nearest large known deposits are some, 1,500 to 1,700 km from the Urucum deposits. The nearer ones, and possibly the others also, were probably covered by younger rocks during deposition of the Jacadigo Series. Beneath the Jacadigo Series lies a complex of highly metamorphosed Precambrian crystalline rocks older than the granite mentioned above. It is conceivable that weathering of this complex supplied the iron and manganese. The scale of this, like most major iron-formations, is so vast that the problem of ultimate source is not easily solved.

Summary

To summarize, major deposits of both oxide- an carbonate- facies iron-formation, and manganiferous sediments of Pre- cambrian and early Palaeozoic age are known in Brazil. In several places, the oxide-facies iron-formation is closely associated with manganiferous rocks, the latter being interbedded with the iron-formation. However. in no known case

is banded iron-formation closely associated with carbonate-facies manganese deposits, even though both iron- and manganese-formation may occur in the same sedimentary unit. In all cases, carbonate-facies manganese-formation is closely associated with unusually carbonaceous sediments. That this condition is not peculiar to Brazil is shown by the presence of such ores in Africa; there oxide-facies iron and manganese deposits on a very large scale are interbedded in the Kuruman District of South Africa, whereas the carbonate-facies sedimentary manganese deposits of Ghana, the Ivory Coast, Upper Volta and, probably, the Congo are not associated closely with iron-formation. Gra- phitic or carbonaceous sediments are associated. India shows us that oxide-facies manganese-formations may occur without oxide-facies iron-formation, and Gabon shows us that carbonate-facies iron-formation may be followed iii the sedimentary sequence by carbonate-facies manganese-formation, there also associated with carbonaceous rocks.

All these deposits show that oxide-facies sediments are more likely to be found in a platform or miogeosynclinal environment, but the rule is not invariable; a number of carbonate-facies manganese-formations are found in platform or estuarian environments. By the same token, a eugeosynclinal environment seems to be the most favourable for carbonate-facies manganese-formations and iron-formations too. I know of no major oxide-facies manganese-formation deposited in a clearly eugeosynclinal environment.

Résumé

Formation de fer et de munganèss ell ussociution, air Brésil (J. Van N. Dorr II)

Au Brésil, des concentrations de fer et de manganèse sont connues des points extrêmes au nord et au sud du bouclier. Les minerais commerciaux, à l’exception des minerais de manganèse près de Corumba dans le Mato Grosso, ont été formés épigénétiquement à partir de formations de fer rubanées et de roches métasédimentaires riches en manganèse par de nombreux processus. Le degré de méta- morphisme varie largement. Dans certaines régions, les sédiments mangaiiifères sont étroitement associés aux formations de fer. Dans d’autres, l’association est équivoque. Dans d’autres, on ne connaît aucune association.

L’âge des formations de fer rubanées et des roches sédimentaires manganésifères va du Cambro-Ordovicien dans le Mato Grosso à plus de 2,7 milliards d‘années dans l’État de Minas Gerais. Les principales formations de fer sont celles du Mato Grosso, celle de la partie centrale de Minas Gerais (entre 2200 et 1350 millions d‘années) et celle de Para, récemment découverte et non encore datée, quoique presque sûrement précambrienne, peut-être même du milieu du Précambrien. Les plus grands gisements de manganèse sont les couches non métamorphosées du Paléo- zoïque inférieur interstratifiées avec la formation de fer du Mato Grosso, mais sans doute de bien plus grandes quan-

tités de manganèse se sont déposées au moyen Précambrien et au début de cette période. La plus grande partie du man- ganèse précambrien est maintenant métamorphosée en mi- néraux silicatés réfractaires au processus de décomposition et inutilisable dans l’industrie. L’un des divers dépôts de manganèse supergénés d’Amapa a été formé par la décom- position de picrotéphroïte ; c’est le seul dépôt commercial important que je connaisse qui soit formé en grande partie de silicates. Les sédiments originaux y dépassent l’âge de 1,8 milliard d’années.

Toutes les formations de fer connues ont un faciès d‘oxyde ou de carbonate. La forination de fer à iaciès carbonaté est une série engéosynclinale de plus de 2,7 milliards d’années dans la partie centrale de Minas Gerais. Elle est constituée de lentilles qui en général n’ont pas plus de quelques dizaines de mètres d’épaisseur et quelque 10 ki- lomètres de long. O n trouve beaucoup de lentilles de ce genre dans une même zone, qu’elles caractérisent ainsi sur une vaste surface. La même série contient aussi, sur une zone de quelque 200 kilomètres de long, des minéraux sili- catés de manganèse provenant de vases carbonées et de boues contenant du carbonate de manganèse. Localement, le carbonate de manganèse était assez épais et assez pur pour rester chimiquement inaltéré par le métamorphisme ; le minerai commercial provient d‘une oxydation superficielle. La teneur constante en carbone libre est le signe d’un

112


TABLE 2. Complete analyses of manganese and iron ores1, Morro do Urucum, Brazil (from Dorr, 1945) ~

sairipie ho2 MnO SiOz 17eZO:J Alzo3 MgO Ca0 Na20 K20 1120- I I p TioZ Pz05 V203 Nioz COO 13iiO Liz0 As203 Sb203 Cu0 l'ho Sn02 C P ~ O : ~ s Total NO .I

~

Manganese ores - narrrpies fmn bod N ~ . i .% 71.65 1.0.1 1.26 i5,gï 2.50 NOW 0.22 0.05 4.80 0.19 1.74 0.08 0.01 Nane ~ ~ " e 0.31 0.05 ............... mne ~~~e ............... Y Y . ~

a 61.87 6.24 .40 14.24 1.73 0.07 .33 .1ï 4.07 1.11 2.11 . l u .91 0.01 None .I4 .o7 0.001 N~~~ 0.004 hbne

14 69.82 3.48 .u8 16. 59 . <JZ .42 .$o .i9 3.82 .4ï 2.18 .14 .u ..... .o5 .s3 ................................... 99.91

20 61.18 1.53 3.05 23.07 2.20 .23 .IB .30 :j.uo .4ï 2.51 .21 .72 ..... 'i~ne .o4 .i5 ........................................ 100.4.1

ZG 67.19 3.55 1.77 16.46 2.16 .1<J .21j .20 2.03 1.03 3.42 .I4 . G G Nunc .o2 .95 .................... 100.03

28 ~9.90 ,64 .YO 18.08 2.59 .IO .O.% .53 3.21 .22 2.82 .I0 'Pr. Xone NO"^ .35 ,35 ,002 N~~~ N~~~ NOM NO"^ .......... 99.83

infi 66.30 2.61) 1.71 18.13 2.51 .3G .O7 .SO 2.57 .51 4.39 .IO None Nono m n e .33 .26 ,002 N ~ ~ I ~ N~~~ N~~~ N~~~ .......... 100.43

5 Gs.10 1.50 1.03 17.10 Z.IIR 0.09 0.02 0.63 3.43 0.14 I 2.56 0.05 0.83 N O ~ C ~ o n e 0.20 0.33 0.001 N~~~ xanc xone T ~ . xeric .......... 9y.m

15 ï o . j ~ 3.31 ,36 14.80 .Y? .24 .40 .i7 4.48 .so 1.40 .i5 .za 0.01 xonc .io .i6 ~ r . mne rqone 0.003 N~~~~ .......... ss.ae 23 05.43 4.64 1.31 19.71 1.90 .I1 .50 .13 2.59 .50 2.03 .12 .41 None None .16 .a1 ............... 'xone nune 33 72.10 1.77 1.33 17.43 .i1 Nane .o5 .OG 4.28 .19 1.58 .O4 .28 None None .z4 .lis ............... xone None ............... io0.11

10 72.45 2.58 .56 15.41 1.04 Nuno .a .zu 3.63 .44 2.33 .o5 .34 xone TI-. ,2a .i4 .......... Xons None

..... ...............

Manganese ores - samples ïrarri bed Wo.2

105 69.90 2.80 .96 17.19 1.94 NoDe NOnC .46 3.47 .22 2.34 .O8 None None None .30 None ,002 None Nons None Tr. . None .......... 99.69

Iron ores .............................. ......................... 1 7 ~ IO ..... 0.71 27.8s 70.74 0.36 0.01 0.17 ..... .o.007 0.017 s9.aa

F~ 11 ..... .o9 9.02 119.87 .4a NO"^ 0.54 0.30 NOIE .i9 .................... ,O.iU . 036 100. 60

F~ 25 .o1 11.40 87.60 .a6 0.14 .......... Tr. Tr. 100.29

.......... .......... ..... .......... ..... .................

1. Complete analyses made in the laboratories of the Geological Survey. 2. Location of samples shown on plates 5 and G of Dorr, 1945.

sored investigations in the part of the Jacadigo Series extending into that country in the Serrania de Mutum, and a diamond drill hole about 250 m deep was sunk in the iron- formation. This is said to have revealed an average content of about 50 per cent Fe and 25 per cent SO,; unfortunately I do not have data on the details of the stratigraphy revealed. No manganese was found. Several years later a team of German geologists working in the area found small low- grade lenses of manganese oxide. On the Brazilian side of the border it has been reported that important lenses of manganese oxide have been found in the iron-formation. I know no details of the occurrence. Thirty years ago, I hastily visited the Mutum area and have the impression that the clastic content of the iron-formation is much less than in Morro do Urucum.

In the Serra do Rabicho, about 15 km to the north-east of Morro do Urucum, Haralyi reports (oral communication, 1968) that manganese is present in the iron-formation in a detrital facies, forming the cement in the detrital rocks in some places. It is not known how extensive this lens may be; Haralyi states that the potassium and iron contents of the manganese are lower than that at Morro do Urucum.

Thus, judging from the stratigraphic evidence in hand and subject to correction as more information becomes available from this remote part of South America, it would appear that the iron-formation and the included manganese oxide beds were chemical sediments laid down in an extensive basin filled, in the first instance, by continental clastic sediments and later by chemical sediments. The unusually high content of potassium (3.5 per cent K,O) in the manganese ore suggests that the waters of the basin were not ordinary sea-water, The essential absence of vanadium, nickel, arsenic, antimony, copper, lead, and tin and the relatively low content of barium and cobalt point in the same direction and also militate against any volcanic contribution. The

very high state of oxidation of the iron and manganese ores (MnOJMnO = 23.7) certainly points to a strongly oxidizing environment. The evidence of the association of the manganese oxide beds with clastic beds in the iron-formation clearly indicates an abrupt change in the sedimentary environment that caused precipitation of manganese, for the iron-formation itself contains less than 1 Clarke of Mn. The deposition of jasper was then almost completely inhibi- ted, for much of the SiO, in the manganese beds, only 1.25 per cent in all, is in detrital grains and secondary cryptocrystalline silica. Whether the water in the basin from which the iron-formaiion was deposited was fresh or salt is thus uncertain. I suspect that the basin was estuarian or lacustrine, and that an arid climate may have concentrated the waters to a considerable extent and contributed to the lack of detrital material during most of the time the iron- formation and manganese deposits were being deposited. Rapid and temporary changes in climate may have brought in sudden and local incursions of detrital material and changed the composition of the waters in such a manner that the manganese oxide was deposited. The very local distribution of the coarse clastic material and its angular nature suggest that turbidity currents may have been the agent of transportation. The intermixing of these coarse sediments with manganese oxides and the greater amount of detrital sediments to the north-east suggest that the shoreline may have been to the north-east and not far from the present margin of outcrop of the Jacadigo Series.

The ultimate origin of the manganese and iron is purely speculative. The quantities involved are enormous. Alvo- rad0 (1970) estimated 40,000 million tons of iron-formation in Bolivia and 10,000 million tons in Brazil; this represents small erosional remnants of the original extent. I estimated (Dorr, 1945) that perhaps 500,000 million tons were originally present. I also estimated (Dorr, 1945) that perhaps

111


milieu euxinique au cours de la formation du dépôt. Le faciès sédimentaire original ainsi que le milieu environnant au moment où s’est formé le dépôt de centaines de lentilles minces de formation ferreuse hautement métamorphosées et les roches de silicate de manganèse dans le gneiss et le schiste des régions du bouclier sont inconnus. Le graphite que l’on trouve dans certaines lentilles de silicate de manga- nèse suggère un faciès original carbonaté. A Amapa, un dépôt cyclique de carbonate de manganèse syngénétique presque pur, en couches de 30 mètres d‘épaisseur avec grès, argile schisteuse très carbonée et boue, indique un milieu environnant euxinique. La plus grande partie du minerai d‘oxyde supergène provient de ce carbonate.

Les principales formations de fer au Brésil sont des plates-formes à faciès d‘oxyde ou des dépôts miogéosyncli- naux. Tous les gisements importants de minerai de fer ont été formés épigénétiquement à partir de telles roches. Les sédiments graduellement sous-jacents sont transgressifs. Les principales formations de fer à faciès d‘oxyde ont plus de 200 mètres d’épaisseur. Elles sont continues sur des surfaces qui couvrent des centaines, voire des milliers, de kilomètres carrés ; leur extension originale a dû être beaucoup plus grande. Les sédiments détritiques qui y sont inclus ont un caractère local et sont de peu d’importance. La teneur varie de 30 % à 50 % de fer.

La plupart des formations de fer à faciès d’oxyde

contiennent moins d’un Clarke de manganèse, mais des lentilles ou des zones importantes atteignant 6 mètres d’épaisseur dans la formation de fer peuvent contenir une formation de fer riche en manganèse ou, dans le cas du Mato Grosso, peuvent contenir des couches d‘oxyde de manganèse syngénétique s’étendant sur plusieurs kilomètres carrés, de 1 à G mètres d’épaisseur et contenant jusqu’à 48 % de manganèse. La formation de fer à faciès d‘oxyde manganésifère n’a jamais été reconnue dans un état avant désintégration, on présume que le manganèse apparaît sous forme d‘oxyde syngénétique en basse concentration. D e vastes gisements de manganèse ferrugineux ont été formés à partir de ces roches. Dans la formation de fer précam- brienne, les lentilles manganésifères tendent à se concentrer là où la formation de fer est en transition vers une formation de dolomite superposée, illustrant l’action du p H sur le dépôt. La dolomite peut contenir jusqu’à 30 clarkes de man- ganèse sous forme de calcite ou dolomite manganique.

Ainsi au Brésil, la formation de fer à faciès d’oxyde se trouve en grandes masses continues étroitement associées avec le manganèse, tandis que la formation de fer à faciès de carbonate se trouve en masses discontinues minces sépa- rées des dépôts de carbonate de manganèse dans la même séquence de roches. Les sédiments riches en carbonate de manganèse sont beaucoup plus étendus dans le temps et dans l’espace que les formations de fer à faciès de carbonate.


ALDRICH, L. T.; HART, S. R.; TILTON, G. R.; DAVIS, G. L.; RAMA, S. N. I.; STEIGER, R.; RICHARDS, J. R.; GERKEN, J. S. 1964. Isotope Geology. Annual Report Director, Dept. Terrestrial Magnetism: Carnegie Znst. Washington Yearbook 63, p. 33-340.

ALMEIDA, F. F. M.; MELCHER, G. C.; CORDANI, U. G.; KAWA- SHITA, K.; VANDOROS, P. 1968. Radiometric age determinations from northern Brazil. Bol. Soc. brus. Geol., vol. 17, no. 1, p. 3-14.

ALVORADO, B. 1970. Iron ore deposits of South America. Survey of world iron ore resources, p. 302-380, New York, N.Y., United Nations Department of Economics and Social Affairs, (United Nations Publications Sales No. E. 69. II. C. 4, ST/ECA/ll3 ,)

DORR, J. Van N. II 1945. Manganese and iron deposits of Morro do Urucum, Mato Grosso, Brazil. Bull. US. geol. Surv.,

-. 1964. Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., vol. 59, no. 7, p. 1203-40.

-. 1965. Nature and origin of the high-grade hematite ores of Minas Gerais, Brazil. Econ. Geol., vol. 60, no. 1, p. 1-46.

-. 1969. Physiographic, stratigraphic, and structural development of the Quadrilátero Ferrífero, Minas Gerais, Brazil. Prof. Pap. US. Geol. Surv., 641-A, p. 1-110.

DORR, J. Van N. II; COELHO, I. S.; HOREN, A. 1956. The manganese deposits of Minas Gerais, Brazil. XX Znt. geol. Congr., Mexico City, 1956, Symposium sobre yacimientos de manganeso, vol. 3, p. 279-346.

GAIR, J. E. 1962. Geology and ore deposits of the Nova Lima and Rio Acima quadrangles, Minas Gerais, Brazil. Prof. Pup. U.S. Geol. Surv., 341-A, p. 1-67.

946-A, p. 1-47.

JAMES, H. L. 1954. Sedimentary facies of iron-formations. Econ. Geol., vol. 49, no. 3, p. 235-93.

KRUMBEIN, W. C.; GARRELS, R. M. 1952. Origin and classification of chemical sediments in terms of p H and oxidation- reduction potentials. J. Geol., vol. 60, p. 1-33. MASON, B. 1949. Oxidation and reduction in geochemistry. J. Geol., vol. 57, p. 62-72.

MATHESON, A. F. 1956. The St. John del Rey Mining Co., Limited, Minas Gerais, Brazil; history, geology, and mineral resources. Bull. Canad. Znst. Min., vol. 49, no. 525, p. 37-43.

NAGELL, R. H. 1962. Geology of the Serra do Navío manganese district, Brazil. Econ. Geol., vol. 57, no. 4, p. 481-98.

RIBEIRO FILHO, E. 1966. Jacobsita de Licinio de Almeida, Bahia. Bol. Soc. bras. Geol., vol. 15, no. 2, p. 43-8.

RIBEIRO FILHO, E.; ELLERT, N. 1969. Magnetometria relacionada a jazidas de manganêc do sudoeste da Bahia. Mineraç. e Metall., vol. 49, no. 289, p. 11-13.

SCARPELLI, W. 1966. Aspectos genéticos e metámorficos das rochas do distrito de Serra do Navío, Território Federal do Amapá, Brazil. Anais da VI Conferencia Geológica das Guianas. Avulso Dep. nac. Prod. min., Rio de J., Div. Geol. e Minerai, no. 41, p. 37-57.

SHATSKIY, N. S. 1954. On manganiferous formations and the metallogeny of manganese, Paper 1. Volcanogenic-sedimentary manganiferous formations. Int. Geol. Rev., vol. 6, no. 6, 1964, p. 1030-56. (Translated by V. P. Sokoloff from original article in Akad. Nuuk. SSSR, Zzvesstiya, Seriya Geologi- cheskayu, 1954, no. 4, p. 3-37.)

113

The Precambrian iron and manganese deposits of the Anti-Atlas

G. Choubert and A. Faure-Muret Museum National d'Histoire Naturelle, Paris (France)

A short outline of the Precambrian of the Anti-Atlas

The Precambrian of the Anti-Atlas is subdivided into six systems of which the first five were terminated by important orogenies, resulting in folded and granitized belts, whereas the sixth forms the base of the Palaeozoic cover of earlier folded zones. During the first two systems the Precambrian basement became enlarged aiid resulted in: 1. The Kerdoiis system (Archaean) which granitized at

2,600 m.y. The Zagorides, trending in an E-W direction, are the orogenic belt which corresponds to this system.

2. The Senaga system (Pvecambrian I) is of early Precam- brian age. The orogenic belt that this system constitutes are the Berberides to which three-fold phases correspond and in which granitization has been dated at 1,940,1,850 and 1,750 m.y. The strike varies around the southerly direction.

These two belts cratoiiized the Anti-Atlas in such a way as to cause subsequent belts to mould themselves against the northern front of this craton. They are: 3. The El Gruara system (Precambrian I-II) which is

characterized by important volcanic activity and by the emplacement of ultrabasites. The central Anti-Atlasides have been dated at about 1600 to 1680 m.y.

4. Tfie Limestone and Quartzife system (Precambrian II) which form most of the epicontinental cover of the three previous belts. They are nevertheless strongly folded and locally granitized (1450 to 1500 m.y.). These may be referred to as the Western Anti-Atlasides.

5. The Eastern Anti-Atlasides (900 to 1050 m.y.) constitute the last Precambrian orogenic belt. Like the other two this belt is folded in an E-W direction, moulds the front of the ancient craton and constitutes the Siroua-Sarhro systeni (Precambrian 11-111) which begins with volcanites and tillites dated at 1250 and 1350 m.y. The granitization accompanying this belt is extremely important (Tifnout and Ouzellarh granites).

6. The last Precambrian system begins with the important ensemble of acid volcanites of the Ouarzazate series, which were covered by the Adoudounian dolomites. The sedimentary cycle of the Lower Adoudounian was followed by an important regression separating formations of the sedimentary cycle and embodying the Lower Adoudounian and the Lower Cambrian. The age of this ensemble ranges from 900 to 550 m.y.

There are no important iron or manganese deposits in the first two systems; only some veinlets of oligoclase or veins of quartz with oligoclase and hematite have been located. Among the latter, the more important is the twin orebody of Bou Tazoult on the eastern side of the Archaean massif of Ifni (Fig. 1). Iron-ore deposits are found in the following two systems: (a) at the base of the Precambrian 1-11, and (b) in the schists of the Precambrian Il. The manganese deposits are connected with the acid volcanic series of the platform, partly at the base of the Precambrian 11-111 and partly in the Precambrian II (Ouarzazate series).

The iron ore deposits of the Precambrian 1-11

The extreme base of the Precambrian 1-11 is locally under- lined by a ferruginous formation, the oligoclase schists, belonging undoubtedly to the large family of itabirites (Choubert, 1963). This formation can be observed notably at the southern boundary of the Precambrian window (Bou- tonniere) at El Graara (Central Anti-Atlas south of Ait Ahmane), where it constitutes the ferruginous band of Guelb el Hadid (Fig. 1). These oligoclase schists overlie unconformably the gneiss of Oued Assemiil, which is undoubtedly Precambrian I. It is therefore an ore deposit of the platform cover and consequently closely coiinected with the itabirites. Discovered in 1949 (Jouravsky, 1953) this formation can contain up to 66 per cent Feto, (46- 50 per cent Fe) and a siliceous residue of 24 per cent. In thin section this rock consists uniquely of well-assorted and generally intensely crenulated oligoclase grains and lamellae

Unesco, 1973. Genesis of Precanibriun iron und inringmese deposits. Proc. Kiev Symr,., 1970. (Earth sciences, 9.) 115


and of granoblastic quartz. In places the oligoclase schist may pass laterally into biotite-mica-schists. It is therefore relatively intensely metamorphosed (amphibolite facies). This has probably taken place at origin, and the formation may now be observed as a lenticular body which marks the contact of the Precambrian 1-11 with the Precambrian I. In some parts at Guelb el Hadid it can be followed over nearly 2 km; the maximum thickness is 50 m . Elsewhere it measures no more than a few decimetres or it may be discontinuous.

The oligoclase schists are actually known to occur over nearly 20 km, stretching from Ihrtem in the west along the northern flank of the Takroumt Massif and again to the east of Hassi Atlatat, where they emphasize the axes of anticlinal structures. They are deposits of sedimentary origin connected with the transgression of the Precam- brian 1-11 on to the flattened Berberides, but it is not certain whether they are a marine formation due to that transgression or a continental formation deposited on the gneiss before the transgression-a type of lateritic crust.

The ferruginous formation may be overlain by metamorphic limestones. However, in the presence of over- thrusts, folding and crenulations in the oligoclase schists, it is often difficult to recognize the exact relationship between these two formations. At Guelb el Hadid some small lenticles of limestone measuring 10-20 cm can be observed, closely intermingled with the oligoclase schists and crenulated simultaneously with them. The presence of these limestones could indicate a marine origin.

Indications of the presence of iron in the Precambrian II It is also in the eastern part of the El Graara window (central Anti-Atlas), notably near El Bleida, where there are indications of iron in the schists of the Precambrian II. This system begins here with an important body of quartz diorite which has replaced the original stratigraphic unit for this system, i.e. the limestones of the base. The El Orf quartzites (1529 m) form the second stratigraphic unit; they are present in the shape of a great upright lens 4 km long. The third unit consists of basic volcanites, preceded by some schists and followed by a thick complex of schist and flysch. Weak metamorphism resulted in the formation of greenschists. Iron is also present here as blackish, ferruginous schists with oligoclase. In thin section, they are rocks formed almost uniquely of oligoclase flakes, needles or rounded lamellae and of sericite, well aligned along the beds. There are also some quartz grains, biotite lamellae and chlorite present, as well as rocks formed of quartz and sericite with very fine bands and clouds of oligoclase. These rocks are situated both in the schist band, which separates the quartzites from the volcanites, and in the schists which overlie the latter, for instance in the Gardens of Bleida. These ferruginous occurrences in the schists of the Precambrian II seem of little importance; neither the Fe content nor the tonnages have so far been

evaluated, and they are mentioned here merely as a point of interest. O n the other hand, the same supra-volcanite schists may possibly have Cu mineralization, and investigations are at present being made.

Deposits of manganese in the Precambrian 11-111 Only one deposit of M n dated Precambrian 11-111 merits description. This is the stratiform deposit of Idikel (the village situated below the mine), or alternatively named Aferni (after the nearby mountain pass) (Fig. i), The deposit is situated to the east of Tafraoute (western Anti- Atlas) on top of a cliff. It is interstratiñed with an essentially conglomeratic series about 200 m thick which borders the Archaean massif of Kerdous to the east. This series is the lateral, eastern, reduced equivalent of the thick series of Anzi (1000 m), whose schists and sandstones constitute the country to the west of Tafraoute. The Anzi series surmounts an important volcanic complex, which is essentially rhyolitic and which, up to a short time ago, used to be confused with the massive rhyolitic series of the Precambrian III. Else- where a thin rhyolitic flow can be observed at the base of the detrital series of Idikel. The characteristics of the deposit, as studied and described by Bouladon and Jouravsky (1956), are as follows:

'The conglomeratic series of Idikel is of fluviatile origin and contains intercalations of micaceous pelites (or psam- mites) with floating mica, which originated in the Archaean schists and granites. One of these intercalations (2-3 m thick) comprises a sandstone zone and red pelites with volcanic constituents. It is accompanied by a red dolomite zone of lacustrine origin containing 4-12 per cent MnO. The manganese deposit is interstratified in this red zone and irregularly bedded, the main bed attaining a thickness of 1.5-2 m. Other non-exploitable beds are situated sometimes above, sometimes below, the main bed. The outcrop can be followed over nearly 2 km and dips 20-25" to the east. The rich mineral belt is 800 m long and has been exploited over a width of 150 m. Beyond that the mineralized bed is replaced by dolomite poor in manganese. The marketable grade had 37-51 per cent Mn with 1-13 per cent SO,, 5-10 per cent Bao, 0.02-1.3 per cent Pb, etc. An average of 10,000-20,000 tons per years was mined. Exploi- tation commenced about 1951 and was discontinued about 1959. The total minerals extracted was 100,000 tons.'

The mineral consists of braunite and barium-bearing psilomelane with a little haussmannite and pyrolusite. Rho- donite and rhodochrosite appear locally, the latter close to the dolomites. Hematite, baryte, quartz and albite and different micas are also present. O n the other hand, chalcopyrite and other Cu minerals, baryte, oligoclase and quartz are found in the veinlets which cut the mineralized beds, as well as some acmite and spessartine garnet. The presence of these silicates, for instance rhodonite, indicates a local rise in temperature which was subsequent to the formation of the orebody. The Anzi series could perhaps be locally

117


lightly metamorphosed, i .e. newly formed biotite is present in the rhyolites of Oued Oudrar not far from Anzi.

Several hypotheses can be put forward regarding the origin of the Idikel deposit. 1. According to Bouladon and Jouravsky (1956) it is a

hydrothermal syngenetic deposit. ‘The Idikel deposit was presumably formed during sedimentation which took place at the side of a closed and probably lacustrine basin and emanated from magmatic mineralizing solutions’. Elsewhere Jouravsky (1963) speaks of the Idikel deposit as an oxidation mineral in a superficial zone of carbonatized mineral, doubtlessly in relation with the manganiferous dolomite.

2. The red dolomites, apparently not accompanied by manganese deposits, frequently outcrop on the eastern boundary of the Kerdous Massif (here and there at Izerbi, the Agoujgal window, etc.). These lakes are undoubtedly later than the rhyolitic volcanic activity, which marks the beginning of the sedimentary cycle of the Precambrian II-III. The destruction of these volcanoes, which comprise the zones richest in manganese, and the concentration of manganiferous products in these lakes, could perhaps explain the origin of the Idikel deposit thus making it unnecessary to evoke a hydrothermal origin. It can therefore be assumed that a situation obtains here similar to the Cretaceous manganese deposits along the northern boundary of the Anti-Atlas. The deposit would therefore be syngenetic sedimentary, perhaps enriched by an oxidation process.

3. Finally, according to a third hypothesis, it is an epigenetic hydrothermal deposit: the mineralized zone would have been formed by substitution in the manganiferous dolomite and the embanked pelites by hydrothermal activity.

According to the isopach curves by Mixus (in Bouladon and Jouravsky, 1956) the greater part of the mineralized zone is represented by a rectilinear band orientated N-S and measuring 50-100 m. The thickest parts of the deposit (1.5- 2 m) are aligned to this zone, whereas the manganiferous dolomites border it continuously, except where the deposit is intercepted by the cliff. This rectilinear shape of the mineral zone could explain the feeder of manganiferous solutions, which would have given place to a substitution deposit in the dolomite.

Before passing on to the following section, the presence of a non-exploited manganese lode deposit should be mentioned in the Precambrian 11-111 in the Tidili area, which is situated at the border of the Haut Atlas with the Anti- Atlas (Anfid village). This deposit is of interest because of the presence of rhodonite and spessartine which indicate a rise in temperature (250”) at the time this deposit was formed (Bouladon and Proust, 1959).

Orientation of these lode deposits is ENE-WSW; one of them attains 1 m in thickness. Their paragenesis comprises, moreover, rhodonite (75 per cent Si0,Mn) and spes- Sartine (40.25 per cent MnO), psilomelane, polianite and some hydrated oxides of M n (wad).

It may be concluded therefrom that the Mil deposits

in the Precambrian 11-111 formed at a higher temperature than those in the Precambrian III, to be discussed in the following section. In fact, the latter contain neither rhodonite nor spessartine.

Manganese deposits in the Upper Precambrian (Precambrian íII)

W e owe our present knowledge of the manganese deposits of the Precambrian III above all to Jouravsky. Their study was commended by Neltner (1934), andBondon and Frey (1937) provided us with descriptions and metallogenic details. More detailed studies were made during the 1950s by Jou- ravsky, accompanied by Bouladon and, later, by Pouit.

These authors divide the M n deposits in the Precam- brian III into two great groups: (a) the lode deposits, and (b) the stratiform deposits. The former are very numerous, above all in the region south of Ouarzazate (90 km E-W and 60 km N-S); but they have also been located between the axial plane of Taliouine to the west and as far as east as Sarhro, a distance of about 250 km. Small stratiform lenticular deposits are often associated with lode deposits; but important stratiform deposits (Fig. 1) are not known except in the synclinal basin of Tiouine (40 k m west of Ouarzazate), which measures about 10 x 15 km

The Precambrian III of the Ouarzazate region consists of a succession of slightly unconformable volcanic complexes. Those of the Lower Precambrian III are essentially andesitic, whereas those of the Middle Precambrian III are predominantly rhyolitic (ignimbrites, lavas, tuffs) but generally contain some andesites, latites, even trachytes. Fi- nally, the Upper Precambrian III is characterized by its detrital continental sedimentary basins of the same periods as the lacustrine limestones. Its rhyolites are generally alkaline, and the andesites are usually porphyritic.

The majority of these lode deposits appear in the rhyolites of the Middle Precambrian III. The Tiouine basin, however, consists of conglomerates, sandstones and pelites of the Upper Precambrian III. A description of the characteristics of these two groups follows.

LODE DEPOSITS OF MANGANESE IN THE PRECAMBRIAN III

According to Jouravsky (1963) there are more than 200 lode deposits of manganese in the Ouarzazate area, which extends 90 km E-W and 60 km N-S (Fig. 1, inset). Together they have provided nearly 400,000 tons of mineral containing up to 50 per cent Mn. Lately, their output has been considerably lower and most of these lode deposits are no longer worked, except by local workers.

These lode deposits appear mostly in groups or ‘swarms’. Jouravsky claims that each swarm affects but a single unit or volcanic complex, a single flow or strata of


.loo

o

rhyolite, ignimbrite or andesite (Fig. 2). The lode deposits continue only in exceptional cases from one volcanic unit to another and are bounded top and bottom within the same unit. The adjacent unit may also contain lode deposits, but these are not in any way connected with those in adjacent units. It can therefore be concluded that it is a succession of generations of lode deposits, each generation accompanying a different unit or volcanic complex.

According to the same author, the lode deposits are never deeply situated. None has ever been exploited deeper down than 100 m. The most important ones extended to a depth of 9095 m (Bou Ouzgouar and Tizgui el Illane). Elsewhere exploitation depth did not exceed a dozen or so metres. In most cases, the lode deposit became rapidly wedged in depth in such a way that it became difficult to recognize or to follow the fracture, for instance, the deposit of Taourat which tailed out at - 62 m.

Mn is less frequently replaced by hematite (e.g. the underground workings at Charlot and at Tizgui el Illane). The length of the lode deposits may, however, be considerable, above all in the case of thin deposits poor in gangue mineral, i.e. Tindaf 800 m, Bou Ouzgouar 400 m. The lode deposits, with dolomite as gangue mineral, may attain a thickness of 5 m, but are rarely longer than 50 m.

The mineral composition is constant, the overall composition being braunite, cryptomelane, hollandite, hauss-

w N W

El Borj

El Borj Stratiform deposit

I

mannite, polianite, pyrolusite, psilomelane, todorokite, etc., accompanied by hematite, goethite, barium, quartz, dolomite, calcite, etc. However, the type association is simpler: braunite-cryptomelane-hollandite accompanied by hematite, quartz and barite.

Jouravsky has stated that these minerals crystallized together during the same mineralogical processes, the braunite forming first. This mineral composition hardly changes with depth. Ba0 enrichment occurs near the surface. Jouravsky concluded therefrom that there was no superficial oxidation posterior to mineralization; thus, according to him, mineralization here is hypogene. It must be added here that the manganese lode deposits developed mainly in the Middle Precambrian III, above all in the rhyolites, but also in the andesites, whereas they are rare in the Lower Pre- cambrian III, which is characterized by development of andesites. Finally, Mn lode deposits are practically nonexistent in the detrital formations of the Upper Precam- brian III, though they are volcanic formations, as for example the Tiouine series. The same applies to the Adou- dounian cover (equivalent of the Russian Wendían). It may be concluded that there is a close genetic liaison between the lode deposits and the enclosing flows. The metallogenic processes which led to the formation of lode deposits represent undoubtedly in some way the end of volcanic activity of each mineral complex. Jouravsky qualified these lode

El Bor! Threads E

NSE

300'

200

Ignimbrite Micaceous tuffs A n Grey I:=.=I Dacitic complex O

O 1 Km a Andesites Porphyric andesite

W E

W-f+zt;,;:+2;y + . . . . . ; + + + + * + + + t f + + + t + +\.Iy + + + + + + + + + + + \ + I + - . - - - ~ ? 1 . . . . - + + + + + + + + + + + \I+ltl + + + + + + + f , + .t L . ~ .

Barrage

I

Fault Fault O 1Km '

I I 1

Dacitic micaceous B [....I Rhyolites ignimbrites Recent terrains 1 1 1 Threads of M n

FIG. 2. Sectional diagrams of the thread deposits at El Borj (A) and Tizgui el Illane (B) (Central Anti-Atlas) (Taken from Jouravsky, 1963).

119


deposits as epithermal and closely connected to volcanism. It could also be concluded that the lode deposits were emplaced very near the topographic surface and evidently in a continental environment.

Jouravsky states that saline precipitation of M n took place where the hydrothermal hypogene solutions encountered the superficial waters rich in oxygen. In fact, the salts of trivalent and quadrivalent M n become unstable in diluted solutions and form colloidal Mn(OH),, which precipitates. Furthermore the addition of oxygen necessary for the precipitation of oxidized minerals (here 1.27 per cent) would be clearly deficient. Later it will be seen that the percentage is 2.66 for the minerals of the stratiform deposit at Tiouine. The lode deposits would therefore have formed under conditions of relative lack of oxygen.

STRATIFORM MANGANESE DEPOSITS IN THE PRECAMBRIAN III

As stated previously, stratiform deposits of Mn of some importance exist only in the sedimentary basin of the Upper Precambrian III at Tiouine-Oufrent . There are three such deposits (Fig. 1): Tiouine, Migoudene, and Oufrent, the latter being situated on the eastern side of the synclinal basin.

Elsewhere, in harmony with the main swarms of lode deposits, stratiform lenses of manganese are known to occur, but they are always of minor iniportance. In the area of

WNW

Tiouine, there is a series of lode deposits, the most important being at Tindaf and Taourat. They are situated in the rhyolites and ignimbrites of the Middle Precambrian underlying the Tiouine Series (Fig. 3). This is a sedimentary series of continental origin, red and nearly 1000 m thick. It contains only rarely some isolated flows of trachyte or andesite. However, in other basins of the Upper Precam- brian III, volcanism is both andesitic and rhyolitic and can assume great importance.

At the western side of the basin, the Tiouine series begins with coarse conglomerates (tillites?), which dip 15- 25" to the east. Towards the centre of the basin the conglomerates become medium to fine, thereafter passing into red sandstones and pelites (Fig. 3). The latter have the pronounced red which characterizes the Tiouine basin and the sedimentary rocks of the Upper Precambrian III generally.

This succession is evidently not cyclic, but at all levels there are recurrences of coarse facies in finer facies. In particular, the zones of conglomerates are separated by layers or beds of pelites. O n the other hand, the diminution of the dimensions of the detrital constituents together with the lateral passing of the conglomerates into sandstones and even pelites can be observed in each of the beds going from the side of the basin towards its centre.

Lacustrine limestones with flints containing Collenia are intercalated in the red pelites. These limestones, developed particularly between Oued Ihrir and Migoudene, are also known to occur on the eastern flank of the syncline in the region of Oufrent.

ESE

1400m

. 1300m

. 1100m

FIG. 3. Section of the stratiform deposit at Tiouine showing lateral transition from coarse facies to fine facies (Taken from Colson, in Bouladon and Jouravsky, 1955). 1. Pudding-stones

with very large clasts (tillites?); 2. Normal pudding-stones; 3. Fine pudding-stones going into sandstones and pelites; 4. Mineralized bunch being worked; 5. Subjacent 'principal' ignimbrite.

120


The Tiouine series would appear to be an infilling of this basin by scatter deposits, the detrital material being carried by the floods of transitory inundations which were, however, of short duration. This picture is very different from that evoked by Bouladon and Jouravsky (1955), who envisaged a permanent lake into which various tributaries discharged material. In our opinion only the limestones mentioned above would be lacustrine, whereas the red beds would be comparable with the continental formations of the red sandstone type which generally are assumed to have formed in a semi-arid climate.

The tectonic basin of Tiouine is bounded to the west by important tectonic disturbances. Transverse or oblique faults mark its edge, and these have also been located in the orebody.

The manganese mineralization is present as interstratified beds in the conglomerates; these beds are associated with the pelites and intercalated with them. At Tiouine there are more than twenty beds grouped in three layers: the basal beds; the exploited strata; the upper strata. Their description is part of the present report. However, here are some data by Bouladon and Jouravsky (1955) for each of these successions of strata (Fig. 4).

The lowest beds start at the first pelitic levels some metres above the last ignimbrites. The first level may attain 50 cm and may extend more than 150 m from the border fault. The second level is situated 20 m higher and consists of three beds containing poor mineral.

The exploited layers are situated 50 m higher up. They may consist of up to fifteen beds distributed over 20 m as is

B

/a

O rh yoiitec

FIG. 4. Stratiform deposit at Tiouine: Section (A) and panoramic view (B) of the western edge of the basin. The pudding-stone series is in contact through faults with the ignimbrites (taken from Bouladon and Jouravsky, 1955). 1, 2. Mineralized levels

(pelites, fine pudding-stones); 3, 4. Pudding-stones, some with very large clasts (tillites?); 5. Complexes of ignimbrites, tuffs and breccias.

121


usual in the pelitic layers. Thickness is reduced to 4 m. Extension is 1 km maximum over 300 m according to dip. Updip the beds wedge out rapidly. Downdip they become impoverished and are subdivided into a multitude of layers which disappear slowly into the pelites or sandstones. Certain mineral beds may be eroded by the overlying conglomerates u hich may contain pebbles of mineralized pelites.

The upper sequence of layers is situated 30 m above and occurs only in the northern part of the deposit, and consists merely of 4 thin, poor beds.

The mineral consists of braunite and of cryptorrielane minerals of the hollandite-cryptomelane-coronadite series. Gangue minerals are not abundant and when present consist of quartz, fibrous silica and barytine. Within the mineral the same constituents are found as in the non-mineralized pelites, such as debris of glass, mica, quartz, etc. This could be an indication of the nature of the mineral formed by substitution in the pelites.

The mineralization varies from one sequence of layers to another. Also the lower beds are richer in hollandite than the exploited beds and therefore have braunite and cryptomelane enrichment. On the other hand, the percentage of Pb and of alkalis varies within the same beds from updip to downdip, the Pb being more abundant updip and the alkalis downdip.

Barytine exists only in the lowest sequence. At Migou- dene this mineral is more abundant and is accompanied by quartz and hematite. Locally the latter may even form whole beds.

The average percentages of marketable M n from Tiouine and its principal impurities are: 42-48 per cent Mn; 6-21 per cent SiO,; 2-9.5 per cent Bao; 0.2-2.1 per cent Pb. It is, therefore, poorer than the mineral from the lode deposits and by comparison a little richer in Ba0 and Pb. Its degree of oxidation is higher; 2.66 per cent against 1.27 per cent. The Tiouine deposit was commercially exploited between 1937 and 1962. During this period 50,000 tons of metallurgical rock mineral were mined.

Bouladon and Jouravsky (1955, 1956) consider the Tiouine deposit to be hydrothermal and syngenetic; according to their hypothesis the mineralizing solutions poured into a hypothetical lake. In our opinion the mineral would be hydrothermal pene-epigenetic immediately after the deposition of each bed of pelite. A description of the metallogenic characteristics follows: Its paragenesis is notably the same as that of the lode deposits of the Ouarzazate region, the latter being richer in braunite and poorer in hollandite. They are therefore of common origin.

The rich mineralized beds are situated near the border fault, which bounds the basin to the west. It has been stated earlier in this paper that enrichment occurs along the transverse faults.

Near the same faults, the ignimbrites underlying the deposit show impregnation and substitution by Mn. The manganiferous solutions ‘rose’, therefore, along, these fissures.

The approximate contemporaneity of the Tiouine deposits

and their mineralization have already been shown by Neltner (1934), who noted in particular the break in certain mineral beds by overlying conglomerates as well as the presence of mineralized pebbles in these conglomerates. This has been regarded as proof for the syngenetic nature of the deposits.

Furthermore, these conglomerates may themselves be impregnated by the manganese near the mineralized beds.

Finally, the latter replace the pelite zones or associate themselves with them. The fact that they contain the same detrital constituents as the pelites (vitreous debris, mica, quartz) mentioned earlier on, could show that they are substitution deposits and not deposits formed by precipitation of manganese at the bottom of the lake, as was stated by Bouladon and Jouravsky (1955, 1956).

The formation of the Tiouine stratiform deposits would, therefore, appear to be a complex process which consisted of interaction of three phenomena: infilling of the Tiouine basin by continental deposits; probably abrupt activity of faults during deposition; additions of hydrothermal solutions, doubtlessly related to the re-activation of the faults and, like the latter, taking place during several successive stages. In this way each part of the deposit would be affected by the faults, then impregnated by the solutions. The resulting conglomerate could have eroded the beds which would then have become mineralized, and thereafter the cycle could recommence.

This mineralizing thermalism was undoubtedly coiinected with the end of volcanic activity and thus stratiform and lode deposits have a common origin; both would be hydrothermal epigenetic deposits. The morphology of the lode deposits would be attributable to the Middle and Lower Precambrian III, whereas the stratiform deposits are characteristic of the Upper Precambrian III.

Bouladon and Jouravsky (1955, 1956) quote as comparable deposits those of the Thuringer Wald and the Harz (porphyries, breccia and tuffs of the Permian) and the Ter- tiary deposits in California and New Mexico. In Morocco, the manganese lode deposits may be observed in the volcanoes of the Mio-Pliocene in the Melilla area.

Metallogenic problems of the manganese deposits

It thus follows that in the Anti-Atlas the Precambrian III was an epoch of intense volcanic activity and also a particularly metallogenic one. These two phenomena, that of predominantly rhyolitic volcanism and of mineralization by thermal inanganiferous solutions, appear to be closely connected.

Detailed studies by Jouravsky and Bouladon made it possible to construct a picture of the mode of arrival of these hot solutions, not only in the fissures which traverse the lava or ignimbrite flows, but also in the sedimentary series of continental origin. However, this reconstruction of the mechanism raises a certain number of problems which have so far not been solved.

122


First, why are the M n deposits concentrated above all in the Ouarzazate region over an area measuring 90 x 60 km? The volcanic outcrops of the Precambrian III extend farther than that to east and west; in fact, they characterize a belt of 250 km. However, moving away from the Ouarzazate region, one can note that the ore deposits become more and more rare and disappear altogether. Outside the Ouarzazate area only six manganiferous belts are known, i.e. two to the west (north of Taliouine) and four to the east (in the eastern Sarhro).

The rhyolites themselves are very poor in M n O , and lateral segregation or leaching must therefore be discounted. Elsewhere Jouravsky and Bouladon have shown mineralization to be very superficial and to have taken place in the last flows. Neither can it be shown that certain volcanoes were more conducive to the formation of manganese deposits than others. The selective factor, if it exists, of certain rhyolitic flows would be a very local factor, because in the 90 x 60 km mineralized zone south of Ouarzazate there are several groups of volcanoes with very different volcanic sequences which seem to have had no influence on the distribution of the ore deposits. Finally, lode deposits of manganese are also found in the andesites (Tachgagalt, El Borj, etc.), and one deposit is even situated in the granites of the substratum of the Precambrian 11-111, i.e. the small orebody in the Tamassirt granite (Siroua track) close to the boundary with the Precambrian III. O n the other hand, of the various Upper Precambrian III sedimentary or volcano- sedimentary basins, only that of Tiouine-Oufrent has been mineralized.

The conclusion may therefore be drawn that both lode and stratiform deposits are hydrothermal, but that the origin of the manganese and of the volcanic material is not

necessarily the same. It would appear, however, that volcanism was the factor which set in train the as yet unknown generating processes responsible for the formation of the manganese deposits.

Second, the invariable selective paragenesis of the Precambrian II M n deposits must be emphasized. Chemi- cally, oxides of manganese-always plumbiferous-predominate over all other constituents, i.e. hematite, barytine, quartz and sometimes dolomite (the latter is present in very small, but not negligible, quantities). Therefore, the metallogenic process provoked by the rhyolitic volcanism of the Precambrian 111 evolved in a selective way in providing M n , Fe, Pb, Ba, Si and, eventually, also Ca, M g and some alkalis.

In our present state of knowledge this phenomenon remains inexplicable, but the origin of this mineralization is, however, incontestably connected with rhyolitic volcanism, which is characteristic only of some areas to the exclusion of others and is distributed iii the former over a vast region of 80 x 50 km. This phenomenon occurs only during volcanic epochs and, like eruptions, seems to repeat itself. It must be noted that it is not peculiar to thePrecambrian Ill but that it has also accompanied the rhyolitic volcanism of the Precambrian 11-111.

Despite our knowledge regarding the process which causes the emplacement of the lode and stratiform deposits, the real problem has as yet not been resolved: W h y is this mineralization connected with volcanism, above all acid volcanism? Whence did this mineralization come? Are there deep metallogenic ‘reservoirs’ connected to certain magmatic ‘reservoirs’ of the volcanoes? Or is this metallogenic process autogenic in certain more or less deep zones and under conditions known to be connected with volcanism?

Résumé

Gisements de minerai de fer et de marzgarzèse dans le Précam- brien de l’Anti-Atlas (G. Choubert et A, Faure-Muret)

Les gisements de minerai de fer et de manganèse précam- briens se rencontrent dans le Précambrien moyen et supé- rieur de l’Anti-Atlas.

Dans l’Anti-Atlas, qui est situé dans la partie méridio- nale du Maroc, on ne rencontre pas moins de six systèmes successifs de Précambrien ; chacun d’eux est plissé, granitisé et se superpose au précédent dans la plus complète discordance.

Les deux premiers systèmes, qui sont les plus anciens, à savoir le système de Kerdous, dont l’âge est de 2,6 milliards d‘années (archéen), et le système de Zenaga, dont l’âge est de 1950 à 1 850 millions d’années (Précambrien inférieur), ne contiennent aucun gisement de fer ou de manganèse.

Le système suivant est celui d‘El Graara, dont l’âge

est compris entre I 650 et 1 600 millions d‘années. I1 pré- sente à sa base une couche importante d’oligistoschiste qui pourrait être comparé aux itabirites. Dans le quatrième sys- tème, on rencontre d‘autres oligistoschistes, moins importants et moins continus caractérisés par des calcaires oncho- lites, des quartzites et des schistes-séricites plissés datant de 1 500 à 1 450 millions d’années.

Le manganèse caractérise les deux derniers systèmes : le système Anzi-Siroua-Sarhro plissé et granitisé d’âge compris entre 1 050 et 900 millions d’années ; le système Ouarzazate-oued Adoudou, qui appartient au Précambrien supérieur (900-575 millions d’années).

On observe dans ces deux systèmes un développement important de volcaniques acides du début du cycle ((( volcanisme subséquent ))). Ils sont accompagnés de nombreux gisements de minerai de manganèse veinés ou stratifiés.

Les derniers d’entre eux sont, pense-t-on, d’origine volcano-sédimentaire.

123



AGARD, J.; DESTOMBES, J.; VAN LECKWIJCK, W. 1952. Fer, in Géologie des gîtes minéraux marocains. XIX Congr. géol. int. Alger 1952, Monogr. région. sir. Maroc na 1 et Notes Mém. Serv. géol. Maroc, Rabat, no. 87, p. 103-32.

BLONDEL, F.; BONDON, J. 1936. Le manganèse dans la région du Siroua (Sud marocain). C.R. Acad. Sei. Paris, vol. 202,

BONDON, J.; FREY, R. 1937. Sur la substitution de minéral de manganèse aux rhyolites cambriennes dans la région de Tiouine (Sud marocain). C.R. Soc. géol. Fr., Paris, p. 201-3.

BOULADON, J.; JOURAVSKY, G. 1952. Manganèse, in Géologie des gîtes minéraux marocains. XIX Congr. géol. int. Al- ger 1952, Monogr. région. sir. Maroc no I et Notes Mém. Serv. géol. Maroc, Rabat, no. 87, p. 45-80.

-. ~ . 1955. Les gisements de manganèse de Tiouine (Sud marocain). Notes M é m . Serv. géol. Maroc, Rabat, no. 125, 180 p.

description des gisements du Précambrien III). XX Congr. géol. int. Mexico, 1956, Symposium sobre Yacimentos de Munganeso, vol. 2, p. 217-48. - ; PROUST, F. 1959. Filon manganésifère à rhodonite et spessartite dans le Précambrien II du Haut-Atlas. Mines et Céol., Rabat, no. 8, p. 45-6.

CHOUBERT, G. 1952. Histoire géologique du domaine de l’Anti- Atlas, in Géologie du Maroc. XIX Congr. géol. int. Alger 1952, Monogr. région. Sér. Maroc no 6, et Notes Mém. Serv. géol. Maroc, Rabat, no. 100, p. 75-194. - . 1963. Histoire géologique du Précambrien de l’Anti-Atlas. T. I. Notes Mém. Serv. géol. Maroc, Rabat, no. 162, 352 p. DESPUJOLS, P. 1934. Aperçu sur la géologie et sur les gisements

p. 958-9.

-. , - . 1956. Les gîtes de manganèse du Maroc (suivi d‘une

miniers de la zone française du protectorat marocain. La Science au Maroc, p. 102. Casablanca, Association Française Pour l’Avancement des Sciences (58th session).

JOURAVSKY, G. 1953. Sur la composition minéralogique et chimique des minerais de manganèse des gisements encaissés dans les formations volcaniques du Précambrien III (région d‘Ouarzazate). Notes Mém. Serv. giol. Maroc, Rabat, no. 117, Notes T. 7, p. 289-308. - . 1963. Filons de manganèse dans les f?rmations volcaniques du Précambrien JI1 de 1’Anti-Atlas-Etude métallogénique. Notes Mém. Serv. géol. Maroc, Rabat, no. 170, Notes T. 22,

-; DESTOMBES, J. 1961. Les différents types de minéralisa- tions dans le domaine de 1’Anti-Atlas. Leur cadre géologique et les méthodes de leur prospection. Mines et Géol., Rabat, no. 13, p. 19-57.

LIZAUR Y ROLDAN, J. DE. 1848. Nota sobre unos criaderos de manganeso en el valle del Rio de Oro (Melilla). Notas Inst. geol. España. Madrid, no. 18.

NELTNER, L. 1934. Le manganèse dans les possessions françaises. Les resso~~rces minérales de la France d’outre-mer, vol. 2, Paris, Bureau Géol. et Min. Colon.

POUIT, G.; JOURAVSKY, G. 1960. Gisement de manganèse de Tizi n’Isdid (région de I’Ounein, Haut-Atlas). Mines et Géol., Rabat, no. 11, p. 21-9.

un exemple de minéralisation pénécontemporaine de la série encaissante. Mines et Géol., Rabat, no. 17, p. 41-7.

SITTER, L. U. DE; HUYSE, W. R.; LAGAAIJ, R. 1951. Manganese ores in Eastern Morocco. Geol. en Mij,b., Amsterdam, no. 2,

p. 81-92.

__- ; . 1962. Le gîte de manganèse d‘Ait Isgelt (Anti-Atlas),

p. 52-7.

Discussion

R. P. PETROV. Are the terms ‘itabirite’, ‘taconite’ and ‘jaspilite’ synonymous? W h y do African geologists prefer to use the term ‘itabirite’?

G. CHOUBERT. In France ‘itabirite’ is used for ferruginous quartzite or ferruginous schist. French geologists working in Africa use this term with the same meaning.

S. ROY. Did you determine the thallium and tungsten contents of the volcanogenic manganese ores? Hewett and Fleischer have shown that volcanogenic and hydrothermal deposits have much higher thallium and tungsten contents than those formed as nonvolcanogenic sediments.

G. CHOUBERT. I think that geologists who are at work in Morocco, including Jouravsky, have never analysed the thallium and tungsten contents. I have never heard of such analyses either.

S. ROY. Have you observed any consistent vertical zoning among manganese minerals from higher oxides to lower oxides and silicates along depth?

G. CHOUBERT. As a rule there are no zones in the manganese veins of Anti-Atlas. Jouravsky stated that in veins the Ba0 content is sometimes higher at the surface than at a depth.

124

Tectonic control of sedimentation and trace-element distribution in iron ores of central Minas Gerais (Brazil) A. L. M. Barbosa and J. H. GrossiSad Universidade de Minas Gerais, Belo Horizonte (Brazil)

Introduction Metasedimentary oxide-facies iron-formations are rather frequent in the Precambrian of the state of Minas Gerais, Brazil. The geology of Minas Gerais is not known in sufficient detail to elucidate the mutual relations of all these formations. The area where the most important iron-ore deposits are located, the Quadrilátero Ferrífero, has iron- formations at several stratigraphic levels, but they are neither continuous nor reach high-grade values except at one definite level, namely the Itabira Group of the Minas Series. The conditions of sedimentation of the Minas Series are thus an important part of the history of these ores. The type of ore and the sedimentary environment emerging from the over-all analysis of the rocks are quite similar to those in other Precambrian regions of the world, although the Minas Series stands out among the youngest of the Pre- cambrian iron-bearing strata, all radiochronology methods pointing to an age of about one billion years.

The present investigation was aimed at finding the patterns of distribution of trace-elements in several types of iron ores. The Casa de Pedra mine, near Congonhas, being one of the most important and best explored through drilling in the Quadrilátero Ferrífero, was chosen as a starting point. This paper summarizes the results of determinations made for this mine. A general survey of the geology is given here for the benefit of those who are unacquainted with the literature on Brazilian iron ores.

Sedimentation and tectonic control in the iron-bearing Minas Series

An outstanding amount of iron is found in chemical and detrital deposits, not only in the Minas Series, but also in older and younger metasediments of the Quadrilátero Fer- rífero. Most of the iron has been recycled and, as a whole, mechanical processes, and chemical sedimentary processes at times of marked instability led to the dispersion of iron, while chemical precipitation under quiet conditions formed

the most important iron-formation. This formation, an integral part of the so-called Itabira Group of the Minas Series, shows continuity over many kilometres, and is the only stratigraphic unit that houses metasomatic lenses of pure hematite ore. It thus appears that continuity of the iron-ore beds played an important role in their later enrichment through metamorphic metasomatism.

The sedimentary basin was elongated, with a minimum extent of 500 km N-S to NE-SW and minimum breadth of 160 km in a W-E direction. These figures were reached through geometrical development of major folds and thrust faults. Throughout its depositional cycles this was a mobile zone, with a tectonic control that closely parallels the evolution of Phanerozoic geosyiiclines, and ended, like these, by being compressed into a folded belt.

A strict analogy with typical geosynclines does not hold out. The main differences are: (a) the sedimentary sequences were remarkably thinner here than for any Cale- donian, Hercynian or Alpine geosyncline; (b) they were metamorphosed everywhere, so their non-metamorphic equivalents are unknown, contrary to the rule in younger miogeosynclines; (cj the iron-formation is a unique lithology, not exactly duplicated in younger geosynclinal deposits; (dj all lithologies have a lateral persistence not matched in the Phanerozoic sediments, and no drastic facies changes have been recorded.

The Minas Series has been divided into a lower, clastic group (Caraça Group), a middle, dominantly chemical group (Itabira group) and an upper, dominantly clastic group (Piracicaba group). The latter is a somewhat artificial grouping of five formations, which can be more naturally grouped as a 2-2-1 sequence, as follows: Cercadinho- Fecho do Funil Subgroup; Taboões-Barreiro Subgroup; Sabará formation.

It is generally agreed that the Minas Series, although defined on the basis of lithology, represents a well-defined interval of Precambrian time. The five main divisions adopted above, namely, the first two groups, and the two subgroups plus the isolated formation of the Piracicaba group, are also time-significant .

Unesco, 1973. Genesis of Precambrian iron and manganese deposits. Proc. Kiev Sump., 1970. (Earth sciences, 9.) 125


Each one of the five main divisions is composed of two main lithologies, which can generally be mapped separately and essentially follow one another in an invariable manner, but with gradational and intertonguing contacts resulting from partially simultaneous deposition.

The Minas Series was preceded by deformation, metamorphism and subaerial erosion of earlier formations. Transgression over the mobile zone came from the east and the emerging land to the west supplied most of the detritus, except for the Sabará formation. The transgressive sheet of sand, now quartzite of the Caraça group, was followed by clay deposited in quiet water. The Itabira group resulted from a general lack of detrital supply under continued quiet conditions, together with other factors that led to the most extraordinary deposition of banded iron-formation. When these factors changed, deposition reverted to more normal dolomite. The Cercadinho-Fecho do Funil subgroup shows the return of clastic deposition, with well-balanced rates of basin sinking and uplift in the source of sediments to the west. These shallow water sediments are iron-bearing and show a flysch-type layering. Some layers are iron-rich, detrital iron being derived from then exposed parts of the underlying iron-formation. The Taboões-Barreiro represents a starvation interval, combining rapid sinking with scarce detrital supply. It is followed by the Sabará formation, with the most typical features of flysch deposits. This depositional cycle was closed with the molasse-like Itacolomi Series, followed by deformation leading to structural patterns very similar to the most typical Alpine structures.

Deposition of the iron-bearing Itabira group

With the complete peneplanation of the supply area, only the chemical decomposition products capable of being transported in solution or in a colloidal state could be carried to the sedimentation basin at the Itabira time. Great tectonic stability and uniform climate were the conditions predominant at that time. The sedimentation basin was not too deep through the Quadrilátero area, and oxidizing conditions prevailed in its well-aerated waters. Slight p H variations determined the deposition of the lithologies which distinguish the two group formations: the Cauê formation, whose typical rock is itabirite, and the Gandarela formation, whose characteristic lithology is dolomite.

Typical itabirite is a striped rock, constituted by alternate layers of quartz and iron oxide, this last being generally specularite and subordinately martite or magnetite. W e suppose that this rock looked, before metamorphism, like hematitic jaspilite of the Morro do Urucum, in Mato Grosso, thus being a chemical precipitation product. The local presence of phyllosilicates and amphiboles may indicate clayey detrital contamination. The base of the Caus formation is concordant with the layering of the Batata1 phyllite, and the contact, which is sharp, is probably an isochronic surface. It is possible that the itabirite layering represents an annual sedimentation cycle.

The transportation method and the control of iron deposition are controversial. Admitting the supply by currents fed by meteoric waters which extract iron from the pre- existing continental waters, we should remember that, according to Brusilovsky, thermodynamic calculations favour the transport of Fe(OH), in molecular solution.

Some authors have already speculated on the possible relation of the ferriferous sediments with volcanic activities. This influence is likely when the referred sediments are found in the same stratigraphic levels which typically volcanic rocks occupy, such as the Leptitic Series of Sweden, the Soudan formation of Minnesota, or perhaps the itabirite of the Nova Lima group in Minas Gerais. In the case of the Cauê itabirites, however, this association does not exist and if the iron ore owes its existence to volcanism, it is due in an indirect manner to the volume of carbonic gas which Precambrian volcanoes spewed into the terrestrial atmosphere and, at the time when the biosphere was compounded by the most primitive sea organisms, it was still not extracted on a considerable scale. A larger concentration of CO, in the atmosphere necessarily means a larger quantity of this gas in the sea-water and, therefore, a less than normal pH. Under these circumstances, calcium carbonates would be maintained in solution and iron oxide could be deposited, if it existed in sufficient concentrations. However, the p H was not much below normal as the carbonic acid would havereacted with iron to form siderite, which does not occur in the Itabira group. On the other hand, the association of itabirites with dolomites shows that p H frequently returned to its more normal values above 7.8. Some problems of the genesis of the itabirites are thus still controversial.

The principal component of the Gandarela formation is a dolomitic marble to which dolomitic itabirites associate with phyllites and chlorite schists. The passage between the Cauê and Gandarela formations is transitional and is situated in an extremely variable position in relation to the top and base of the group. The sedimentation of the Itabira time is characterized as predominantly ferriferous at the beginning and predominantly dolomitic at the end, one of these lithologies, however, not excluding the other at any time. The two formations are in fact lithofacies, in a large part synchronic. As a whole, the Itabira Group may exceed the width of 1000 m., but generally it does not extend beyond 200 m.

Metasomatic origin of the high-grade ore

The Precambrian sedimentation in the Quadrilátero Ferrí- fero was followed by intense deformation and metamorphism. In the waning stage of metamorphism a redistribution of the iron took place, with the nearly pure iron oxide bodies as an end product. This process was a metasomatic replacement of quartz and/or dolomite by iron oxide at tectonically favourable sites.

The arguments for a metasomatic origin of high-grade iron ore have been discussed at length by Dorr and Barbosa.

126

Tectonic control of sedimentation and trace-element distribution in iron ores of Central Minas Gerais (Brazil)

The general evidence described in their paper has been confirmed by our more recent investigations in the Casa de Pedra mine. Here the evidence, if anything, is still more pressing, for many isolated itabirite masses, with ill-defined boundaries and only a few decimetres long, are entirely surrounded by hematite ore. It is more diflicult to establish whether the rock that was replaced was a so-called siliceous itabirite (without carbonates), or a dolomitic itabirite. In the specific case of Casa de Pedra our observations seem to confirm Guild's contention that the original rock was dolomitic. The high-grade iron ore body is in contact with dolo- miticitabirite and ferruginous dolomite, both of which show local iron enrichment, with structures very similar to those of the high-grade ore. On the conditions prevailing at the time metasomatism took place, we can add nothing to what has already been established by Dorr and Barbosa.

Geochemical data

OXIDATION RATIO O F THE IRON ORES

The oxidation ratio of oxide-facies iron-formation is primarily a function of the ratio of hematite to magnetite, aIthough maghemite in varying stages of oxidation often takes the place of magnetite.

In the iron-formations of the Quadrilátero. Ferrífero, specular hematite is always the dominant ore mineral. Mag- netite can be seen in the high-grade ores at many places, sometimes in large crystals, but it is known to be present in lower grade itabirite, even if it is generally not megascopically visible in this rock. On most of the outcrops the iron-formation is magnetic enough to disturb the compass, and investigations made in Itabira by Barbour show important amounts of Fe0 both in low- and high-grade ores.

Using Barbour's data, we find that the magnetite/ hematite ratio in hard, unenriched itabirite is about 20 per cent, giving an Fe,O,/FeO ratio of 15.8. This ratio is shown at a depth of nearly 200 m. At shallower depths there is a steady increase in the oxidation ratio.

The high-grade ore is more thoroughly oxidized than the itabirite. In the study mentioned above, only 6 per cent of the hematite ore had oxidation ratios below 35, against 18 per cent of the itabirite samples. The Itabira ores seem to be derived from non-calcareous sediments. On the other hand, in the deposit of Casa de Pedra, which was the object of the present geochemical investigation, the high-grade ore seems to result from replacement Óf dolomitic itabirite, and is distinctly richer in magnetite.

INVESTIGATION O F THE COMPOSITION O F THE IRON ORE TRACE-ELEMENTS AND ASSOCIATED ROCKS

Geochemical investigations of metamorphic and metasomatic rocks of tlie Quadrilátero Ferrífero, as well as some of the weathering products (which sometimes constitute min-

eral deposits), are practically non-existent. In a preliminary investigation made by Herz and Dutra, a small number of samples taken from some of the lithological units of the region discussed had their trace-elements analysed; however, none of these units were geochemically characterized.

During a recent revision and up-dating of the mineral reserves of one of the largest iron ore deposits of the Quadri- látero Ferrífero (Casa de Pedra, District of Congonhas), we had the opportunity to collect about 550 samples of several types of iron ores and associated rocks, as well as soils and laterites derived from thein. These samples had some of their trace-elements determined to establish, for the specific types of ore, distribution standards which would permit the establishment of comparisons and correlations with similar ores in the rest of the Quadrilátero Ferrífero. It was also hoped that ii would be possible to infer tlie nature of the original rocks of iron ore and the behaviour of the suite of the trace-elements during regional metamorphism (and metasomatism) and weathering.

It is important to remember that the averages calculated for the elements investigated are arithmetical (we hope lo broach the problem from the statistical point of view in the future) and the comparisons in this case are largely approximate.

SAMPLING A N D MATERIAL PREPARATION

The samples were taken exclusively from material originating from drill holes (vertical drilling with diamond bits).

Samples from the following units were considered: (1) compact hematite; (2) laminated hematite; (3) schistose hematite; (4) fine granular hematite; (5) average granular hematite; (6) coarse granular hematite; (7) powdered hematite; (8) massive rocks with chlorite and amphibole; (9) chlorite amphibole schists; (1 1) poor itabirites; (12) rich itabirites; (13) dolomites; (14) dolomitic itabirites; (15) decomposed products of dolomilic itabirites and dolomites (clayey-mangano-limonitic material); (I 6) ferruginous aluminous laterites.

In view of the large uniformity of the sampled material, starting from the drill cores, we prepared a small volume sample along one or more metres of column core. In the laboratory, each sample was crushed and ground to 1 O mesh, quartered, reduced in the pestle to 60 mesh and, with successive quartering and reduction, the material was taken to 200 mesh; 5 g were taken for analysis.

ANALYTICAL METHOD

The spectrographical method was employed using a total burn. The standards were synthesized in high purity iron oxide matrix (manufactured by Johnson Mattey Co. and having twenty elements; Ni, Co, Cu, Sc, Zr, V, Pb, Cr, Ba, Sr, Mo, Be, Nb, La, W, Sn, B, Ge, Ti, Mn).

To have a reproducible burn, the samples and standards were diluted in a sodium-quartz-graphite compound.

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The samples and standards were weighed within the craters of the graphite electrodes and burned in the electric arc of the spectrograph until they were totally consumed.

Spectrograph: Jarrel-Ash, 3.4 m focal distance, Ebert assembly with 15.000 traces per inch net. Source: Jarrel- Ash, with a continuous current are of 250 volts and 16 am- peres. Film: Kodak SA-1 (ultraviolet) 5nd Kodak 1-N (visible). Spectral covering: 2500 to 5000 A. Microphotometer: Jarrel-Ash projector-comparator M2 1000.

The spectral lines used were those recommended by Ahrens and Taylor in Spectrochemical Analysis. For each element a work curve was constructed from 1000 ppm to the minimum sensibility limit.

The approximate limits of detection (ppm) for each element are: Ni, 5; Co, 2; Cu, 1; Sc, 5; Zr, 5; V, 10; Pb, 10; Cr, 5; Ba, 1; Sr, 5; Ti, 100; Mn, 100.

The following elements were investigated and were not detected (limits of detection in brackets); Mo(5), Be(2), Nb(20), La(50), W(20), Sn(5), B(10) and Ge(l0).

The spectrochemical analysis were performed at the Geología e Sondagens Ltda. laboratory by C. V. Dutra and Dayse A. O. Lima.

ORIENTATION OF THE INVESTIGATION

In fifty samples taken to determine which elements should be analysed preferentially, we found that, of the list of elements mentioned above, only V, Ni, Cu, Cr, Ba, Ti, Mn and Zr were usually present in the investigated rocks and ores. In the iron ores (hematite and itabirite), the elements Ni, Cu, Cr, Ba, and V systematically occurred and were chosen. In the dolomites, schists, laterites, etc., Ni, Cu, Cr, V and Zr were selected.

RESULTS

Hematites

In the groups of approximately fifty samples of each of the hematite types which had their averages calculated, the results were extremely concordant. The consequences of these unexpected results are considered in another part of this paper.

The combined results for 327 samples are as follows: V, detected in 324 samples; range 12-360 ppm; average 44.5 ppm. Cu, detected in 311 samples; range 2-340 ppm; average 19 ppm. Ni, detected in 219 samples; range 6- 250 ppm; average 22 ppm. Cr, detected in 315 samples; range 8-300 ppm; average 35 ppm. Ba, detected in 54 samples; range 30-1600 ppm; average 91 ppm.

Histograms of the results revealed that a large percentage of the results approached average content of vanadium, copper, nickel and chromium given above, while for barium they did not.

Other elements were estimated in 28 samples chosen at random from the 327. The following results were obtained:

Co, detected in 7 samples; range 6-32 ppm; average 18 ppm. Zr, detected in 11 samples; range 20-60 ppm; average 31 ppm. Ti, detected in 25 samples; range 100-800 ppm; average 657 ppm. Mn, detected in all samples; range 300- 2300 ppm (3 samples with values over 5000 ppm were ignored); average 11 80 ppm .

Sc, P, Sr, Mo, Be, Nb, La, W, Sn, B and G e were not detected.

Itabirites

The average results obtained for rich and poor itabirites were concordant. The combined results for 89 samples are as follows: V, detected in 73 samples; range 6-210 ppm; average 35 ppm. Cu, detected in 71 samples; range 2- 150 ppm; average 22 ppm. Ni, detected in 70 samples; range 7-170 ppm; average 20.5 ppm. Cr, detected in 80 samples; range 8-160 ppm; average 28.5 ppm. Ba, detected in 62 samples; range 34-1000 ppm; average 179 ppm.

Histograms showed that a low percentage of results approached the average.

The following results were obtained from 8 samples (4 of each type of itabirite): Co, detected in 2 samples; 19 and 220 ppm; average 69 ppm. The result does not look significant. Zr, detected in 3 samples; 8, 20 and 24 ppm; average 17.3 ppm. Ti, detected in 6 samples; range 100- 500 ppm; average 216.6 ppm. Mn, detected in all samples; range 500-3500 ppm (1 sample with a value over 5000 ppm was ignored); average 1785 ppm.

Sc, P, Sr, Mo, Be, Nb, La, W, Sn, B and G e were not detected.

Dolomitic itabirites

The following values for dolomitic itabirites can hardly be considered representative; they are only presented as an illustration. Only 6 samples were analysed and in many cases the values were lower than the sensitivity limit. V, detected in all samples; range 26-58 ppm; average 41.6 ppm. Cu, 2 samples with 6 and 8 ppm; average 7 ppm. Ni, 2 samples with 10 and 20 ppm; average 15 ppm. Cr, detected in 5 samples; range 10-110 ppm; average 38.5 ppm (if we do not consider the 110 ppm, the average is 21 ppm). Ba, 3 samples with 10, 36 and 46 ppm; average 27.3 ppm.

Two of these samples showed Ti and Mn, with averages of 150 and 1600 ppm, respectively.

Ferruginous dolomites

Only 12 samples were analysed and in general the results were in agreement. In 8 samples C u was not detected and in 4 samples Ni was not detected. V, range 12-45 ppm; average 30 ppm. Cu, range 13-76 ppm; average 31.5 ppm. Ni, range 8-31 ppm; average 12.5 ppm. Cr, range 13- 44 ppm; average 24.5 ppm. Zr, range 26-72 ppm; average 37.5 ppm.

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Metupelites (massive or schistose)

Phyllites, schists and homophanous rocks (mainly chlorite, talc and amphibole) gave similar results and were grouped together. A total of 57 samples showed Cu, Ni, Cr, V, and Zr, and 4 samples we analysed Co, Sc, Ba, Sr, Ti and Mn. V, range 35-300 ppm; average 107 ppm. Cu, range 2- 340 ppm; average 69 ppm. Ni, range 12-240 ppm; average 67 ppm. Cr, range 13-560 ppm; average 106 ppm. Ba, range 128-240 ppm; average 196 ppm. Co, range 36- 108 ppm; average 58 ppm. Zr, range 38-640 ppm; average 236 ppm. Ti, range 660-22,000 ppm; average 18,850 ppm. Mn, range 1,100-4,500 ppm; average 2525 ppm. Sc, range 36-46 ppm; average 40 ppm. Sr, range 16-42 ppm; average 25.3 ppm.

Mangano- fevruginous material

Physically the mangano-ferruginous material is consistent, with a considerable clayey fraction. Samples were taken in places where it was certain that the material had originated from dolomitic rock decomposition, 27, ferruginous or not, and with a variable content of pelitic material. In 13 samples we did not analyse Ba, and in 16 samples, Zr.

Red und yellow clays

Only 7 samples were analysed. The results were in good agreement. The averages were: V, 123 ppm. Cu, 47 ppm. Ni, 39 ppm. Cr, 163 pprn. Zr, 278.5 ppm.

Laterites

Results for 23 samples were quite similar. The averages were: V, 130 ppm. Cu, 62.5 ppm. Ni, 34.5 ppm. Cr, 123 ppm. Zr, 335 ppm.

The results described above are summarized in Table 1 together with the averages for selected groups of rocks from the literature.

COMPARISONS WITH SELECTED ROCK SEQUENCES

Vanadium

The average results for the hematites and dolomitic itabirites are comparable, as are those for itabirites and ferruginous dolomites. The results for the four units are lie between 30 and 44.5 ppm. In hematite ores (‘quartz-hematite ore’) of central Sweden, twelve samples showed average values of 49 ppm; based on results presented by Lander- green. Our results are also comparable with the 40 ppm of ultramafic rocks. The carbonate content is 20 ppm on average. The metapelites show results (107 ppm) comparable with those for stratified clayey rock (130 ppm).

Decomposition products (residual material over iron ores and associated rocks) set the vanadium in such a manner that there is a larger incorporation of metal in geologically more evolved material, i .e. in mangano-limonitic material the average value is 74 ppm and in the laterites 130 ppm.

All the vanadium is contained in thecrystalline structure of the hematite and magnetite. The constantly lower values for the itabirites than for the high-grade hematite result from the fact that the itabirites contain, besides the iron oxides, quartz with practically no traces of vanadium. Prob- ably the concentration of vanadium in the iron ores was determined at the time of metasomatic metamorphism. The residual enrichment of the itabirite by weathering is essentially a process of quartz leaching, which slightly affected the iron oxides.

Copper

The results for the hematites and itabirites are almost identical and intermediate between the values of the dolomitic itabirites on the one hand and ferruginous dolomites on the other. W e do not have information to compare these results with those for iron ores of other regions. The values for the hematites and itabirites are identical to the values for

TABLE 1. Average results for trace-elements in iron ores and associated rocks (pprn)

1 2 3 4 5 6 7 8 9 10 1 1 12

V Cu Ni Cr Ba Co Zr Ti M n sc Sr

44.5 19.0 22.0 35.0 91 .O 18.0 31 .O 657.0

1,180.0 - -

35.0 22.0 20.5 28.5 179.0 69 .O 17.3 216.6

1,785.0 - -

41.6 7.0 15.0 21 .o 27.3 - - 150.0

1,600.0 - -

30.0 31.5 12.5 24.5 - - 37.5 - - - -

107.0 69.0 67 .O 106.0 196.0 58 .O 236.0

18,850.0 2,525.0

40.0 25.3

74.0 24.0 60.0 51 .O 181.0

101.5 -

- - - -

123 .O 47.0 39.0 163 .O - - 278.5 - - - -

130.0 62.5 34.5 123 .O - - 335.0 - - - -

40 20

2,000 2,000

1 200 30

3 O0 1,500

5 10

200 1 O0 160 200 300 45

1 O0 9,000 2,000

24 440

20 4 20 11 10

19 400

1,100 1

610

O J

130 57 95

1 O0 580* 20 200

4,500 850* 10 450

1, Hematites; 2, Itabirites; 3, Dolomitic itabirites; 4, Ferruginous dolomites; 5, Metapelites; 6, Mangano-ferruginous material; 7, Clays; 8, Laterites; 9, Ultra- mafic rocks (Vinogradov, Geokhimiya, 1962); 10, Mafic rocks (Vinogradov, Geokhimiya, 1962); 11, Carbonates (Turekian and Wedepohl, Geof. Soc. Ant. Bull., Vol. 72, 1961); 12, Stratified clayey rocks (Vinogradov; values with asterisk, Turekian and Wedepoht).

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ultramafic rocks (20 ppm). The carbonate content is 4 ppm on average, Rankama and Sahama gave average values of 20 ppm in limestone and 12.5 pprn in dolomites. The metapelites give average results of 69 ppm, similar to the average (57 ppm) found for stratified clayey rock.

The decomposition products directly connected to the ferruginous dolomites were not enriched during the weathering process with regard to copper; the results show leaching of the element and its fixation in the laterites and clays.

Nickel and cobalt

In the hematites the results for these elements are 22 and 18 ppm respectively, with a Co/Ni ratio of 0.8. In twelve samples of quartz-hematite ore from central Sweden we calculated values of 30 and 20 ppm for Ni and Co respectively (according to data by Landergreen) with a ColNi ratio of 0.7. The itabirites show average contents of 20.5 ppm of nickel and 69 ppm cobalt. (The result for cobalt was obtained from only two very discrepant values.) The average nickel content of the itabirites is comparable with that in the hematites and, when a larger number of cobalt determinations in the itabirites is available, w e can calculate the Co/Ni ratio more accurately.

The nickel contents of the hematites and itabirites are greater than the values for the dolomitic itabirites and ferruginous dolomites. In general terms, the results for these rocks are comparable with the world average for limestone

In the metapelites the Co/Ni ratio is 0.8, while for the stratified clayey rock the ratio is 0.2.

In the mangano-ferruginous material, clays and laterites the nickel shows average values comparable with the original rocks.

(20 PPm).

Chromium

This element is distributed in the hematites, itabirites, dolomitic itabirites and ferruginous dolomites in quantities comparable with values given by Landergreen for ‘quartz-hematite ores’ of central Sweden. The results are noticeably higher than the world average for chromium in limestone. Apparently discrete concentration of the element occurred in the hematites during metasomatism.

The average grade of the metapelites is identical to that indicated for stratified clayey rock.

In the mangano-ferruginous material the average chromium content is greater than in the original rocks, while the clays and laterites show a comparable grade, inferior to that normally reported for similar materials.

Thus the weathering products have an average grade superior to that of the original rocks.

Barium

The distribution of this element in the iron ores and associated rocks is irregular and the contents are not comparable with the average given for selected rock sequences. In view

of its low ionic potential and large size, the barium tends to set by absorption in the clays; this was not confirmed in our case. Nor there was any relationship between the Ba and Mn. A more detailed investigation of the distribution of Ba in the iron ores may explain this.

Zirconium

This element is maintained at a comparable level in the hematites, itabirites and ferruginous dolomites. The averages obtained are similar to the world averages for the ultramafic rocks and limestone.

In the metapelites the distribution of Zr is comparable with the average world values for stratified clayish rocks.

The levels of this element in the decomposed products agree with what we would expect during weathering, i.e. precipitation (by absorption) occurs in the hydrolysed material.

Titanium

A progressive enrichment in this element occurs, which it is difficult to explain. The values are not comparable with the world average for specific types of rock. In the case of metapelites, the average found is exceptionally high.

Manganese

The results for manganese are comparable with those for the Swedish iron ores. Itabirites and dolomitic itabirites show identical grades in Mn, while the hematite shows a markedly lower grade. The average obtained for itabirites and dolomitic itabirites is similar to the world average for ultramafic rocks, while the average for hematite corresponds to that for limestone.

The metapelites are notably Mn-enriched.

Scandium and strontium

Values were obtained only for metapelites and they are not comparable with world averages for specific rock types.

Conclusions

The preliminary study of the distribution of the trace- elements in iron ores and associated rocks of one region of the Quadrilátero Ferrífero suggests: 1. The high-grade hematite bodies present a distribution of

trace-elements essentially comparable with poorer ferriferous rocks. As the investigation stopped at the time of the study of samples of total rock, it is not yet possible to draw conclusions on the migration of these elements in the course of the generating metamorphism metasomatic process of these masses of high-grade iron minerals.

2. The same results are found with representative samples of materials which had suffered weathering alterations,

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i.e. enriched itabirite and masses of high-grade friable hematite.

3. The weathering products show average values comparable with each other and superior to those of the original formations; the results obtained were those expected.

4. In general, the average values for some elements of the hematite-itabirite-ferruginous dolomitic itabirite sequence are concordant with the world averages recog-

nizable for ultramafic rocks, while other elements show averages comparable with limestone rocks. Some elements have averages comparable with either ultramafic rocks or limestones.

5. The metapelites associated with the iron ores are really metamorphosed stratified clayey rocks which are, despite the total absence of quartz, sediments poor in combined silica.

Résumé

Le contrôle tectonique de la sédimentation et la répartition des éléments-traces dans les minerais de fer de la partie centrale de l’État de Minas Gerais au Brésil (A. L. M. Barbosa et J. H. Grossi Sad)

A u cours du Précambrien, des sédiments riches en fer se sont déposés à diverses époques dans l’État de Minas Gerais, au Brésil. Bien que des quantités extraordinaires de fer se trouvent dans des dépôts tant chimiques que détritiques, d‘une façon générale, les processus mécaniques et les processus sédimentaires chimiques dans des conditions calmes ont contribué à la création de la formation de fer la plus importante. Cette formation est continue sur des kilomètres, et elle est la seule source de lentilles métasomatiques de minerai d’hématite pure, la continuité jouant un rôle dans son enrichissement. Le bassin sédimentaire était allongé, son extension étant d’au moins 500 kilomètres entre les directions nord-sud et nord-est - sud-ouest avec une largeur minimale de 160 kilomètres dans la direction est-ouest. Tout au long des cycles de déposition, cette zone a été mobile, sous des actions tectoniques étroitement parallèles à l’évo- lution des géosynclinaux phanérozoïques. Les différences essentielles sont les suivantes : (a) les séquences sédimentaires étaient remarquablement plus minces ; (b) la métamorphose est générale ; (c) la formation de fer n’a pas sa réplique dans les géosynclinaux phanérozoïques ; (d) aucun changement extraordinaire dans le faciès n’a été observé.

Cette communication traite spécialement des dépôts des séries de Minas qui contiennent la plus importante formation de fer. Cette série a été divisée en plusieurs groupes : un groupe clastique inférieur (groupe de Caraça), un groupe moyen essentiellement chimique (groupe d‘Itabira) et un groupe supérieur, surtout clastique (groupe de Piracicaba). Le doyen des auteurs de cette communication considère comme arbitraire le dernier de ces trois groupes. I1 lui semble qu’il serait plus naturel de réunir le sous-groupe Cerca- dinho-Fecho do Funil, le sous-groupe Taboões-Barreiro et

la formation de Sabará. II est généralement admis que la série de Minas représente une époque bien définie de la période précambrienne, Les cinq divisions principales indi- quées ci-dessus ont aussi une signification chronologique. Chacune est composée de deux lithologies dominantes, dé- crites comme des formations, avec un ordre chronologique quasi constant, mais avec des contacts graduels et interdi- gités qui résultent en partie de leur dépôt simultané.

La série de Minas a été précédée par la déformation, le métamorphisme et l’érosion subaérienne des formations antérieures. La transgression de la zone mobile est venue de l’est et les terres émergées à l’ouest fournirent la plupart des détritus, sauf pour la formation de Sabará. La couche de sable quartzite du groupe de Caraça, qui a transgressé, a été suivie par un dépôt d‘argile en eau calme. Le groupe d’Itabira provient d‘une déficience générale de l’apport de détritus pendant des périodes calmes prolongées, qui, avec d‘autres facteurs, a contribué à former un extraordinaire dépôt d‘une formation de fer rubanée. Quand ces facteurs ont changé, le dépôt est revenu normalement à la dolomite. On observe dans le sous-groupe Cercadinho-Fecho do Funil le retour au dépôt clastique avec des périodes bien équili- brées de surhaussement et d‘abaissement du bassin à l’origine occidentale des sédiments. Ces sédiments en eau peu profonde contiennent du fer et présentent une stratification du genre du flysch. Quelques couches sont riches en fer, le fer détritique provenant des parties de la formation de fer sous-jacente qui étaient alors exposées. Le groupe Taboões- Barreiro traduit une période de disette au cours de laquelle un enfoncement rapide s’est trouvé conjugué avec un faible apport de détritus. La formation de Sabará lui fait suite et présente les caractères typiques des dépôts de flysch. Ce cycle de dépôts s’est refermé avec la série du genre molasse d’Itacolomi, suivie d’une déformation qui a donné des sys- tèmes structuraux très voisins des structures alpines les plus typiques.

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Absolute age dating of iron-silicate and ferruginous formations and their position in the Precambrian stratigraphic sequence. précambrienne. Les Analogous formations from formations phanérozoïques the Phanerozoic analogues

La datation absolue des formations de fer et de silicate de fer et leur position dans la série stratigraphique

The iron-chert formations of the Ukrainian shield

N. P. Semenenko Institute of Geochemistry and Physics of Minerals Academy of Sciences of the Ukrainian S.S.R., Kiev (U.S.S.R.)

Iron-chert formations of the Ukrainian Shield are developed in a number of synclinorial zones, representing different structural stages of the earth’s Precambrian crust (Fig. 1).

The oldest is the Konsko-Belozyorsky synclinorial zone, where rocks dated back to 3,500 m.y. have been found. Metamorphic masses composing the Konsko-Belo- zyorsky synclinorium are formations of the oldest defined structural stage of the earth’s crust-presumably the first Precambrian megacycle.

Granites of the age of 3,100-3,400 m.y. are synchronous with rocks of the first Precambrian megacycle.

The Konsko-Belozyorsky folded zone stabilized after the intrusion of the Mokromoskovsky granites-2,700- 2,800 m.y. ago; emplacement of the granites resulted in reworking of older rocks in the western border of the Konsky syncline by processes of granitization and greisen- ization. The above folded zone forms a stable block of the ancient crust. The iron-chert formations remain at lower metamorphic grade (i.e. schist and chert stages) only in their marginal parts, in contact with granites, as they approach a gneissic stage of metamorphism.

Iron-chert formations are present in two cycles-in the Lower Konsky and the Upper Konsky series (Figs 2 and 3).

In the Lower Konsky series-in the Yulyevsky zone of the Konsky syncline as well as in the western zone of the Belozyorsky syncline-iron-chert layers alternate and in- tercalate with thin interlayers of green schist or, at gneissic stages of metamorphism, with amphibolite, forming a rhythmical stratification of chemogenic iron-chert deposits and mixed tuff representing volcanogenic products of basic composition (metabasites). The masses of iron-chert-metabasite formations, in which chemogenic iron-chert products are paragenetically related to tuff-volcanogenic metabasites, attain a thickness of 300-500 m and extend for 10-15 km.

In the upper depositional cycle of iron-chert rocks of the Upper Konsky series iron-chert layers and iron-alumosilicate schists are rhythmically interbanded. Amongst the schists, keratophyre layers are observed as well as acid tuff- schists and tuff-sandstones.

Iron-chert-schist-keratophyre formation constitutes an

Upper Konsky series, replacing an underlying tuff-sandy schist keratophyre suite.

In the Belozyorsky syncline iron-chert layers form a uniform iron-rich zone 100-300 m thick and 30-90 km long. Here, a secondary metasomatic enrichment took place which has resulted in formation of high-grade hematite ore deposits with an iron content to 68 per cent.

In the Konsky syncline the upper iron suite is represented by iron-chert-keratophyre-schist formation which reaches 100-200 m thick and 10-20 km long. However, the iron content of the iron-chert deposits ranges along strike from 35 per cent to 20 and to 10 per cent within an interval of 1-3 km.

Another zone of iron-chert deposits is present in the Bazavluksky synclinorium. It comprises a number of large brachy-synclines-the Verkhovtsevsky, Sursky and Cher- tomlyksko-Solenovsky,-and is composed of metamorphosed iron-chert formations. The oldest minerals found here are 2700-2800 m.y. in age.

The metamorphic formations are found to contain two main cycles of iron-chert deposits corresponding to the Lower and Upper Bazavluksky series (Fig. 4).

In the lower series of the Verkhovsevsky syncline, iron- chert deposits are present mainly in the middle suite. Iron- chert layers alternate here with tuff-schist-metabasite layers. Four iron zones are distinguished each with iron-rich, iron-chert layers. Individual beds of iron-chert layers do not exceed 20-25 m thick; a number of layers approach 20-40 per cent. The iron-chert-metabasite formation is 300-500 m thick; the formation is traced in the area of the Verkhovtsevsky syncline for a distance of 50 km.

Iron-chert intercalations are also found in the upper metabasite suite of the Lower Bazavluksky series of the Verkhovtsevsky syncline. The iron-chert interlayers do not exceed 10-20 cm thick; they amount to not more than 10 per cent of the suite.

In the Sursky syncline iron-chert metabasite formation is also observed in the lower metabasite series. It ranges from phyllite-greenschist to gneiss amphibolite metamorphism, but reaches a gneiss-pyroxene stage in


N. P. Semenenko

Formations:

Iron-chert-schist Iron-chert-keratophyre- schist

@@ Iron-chert-metabasite Iron-chert-ultrabasite

Undivided iron-chert

Enclosing formations:

Rhythmical schist-terrigene Keratophyre (leptite) Porphyrite

a Metabasite Ultrabasite

136

Age provinces:

Konsko-Belozyorsky , m. 2,700-3,500 m.y. Bazavluksky, 2,300-2,700 m.y. Orekhovo-Pavlogradsky , 2,000-2,300 m.y. Korsaksky

Saksagansky metabasite Krivoyroghsko-Kremen- chugsky, 1,700-2,000 m.y.

5 - . .

*I -----

Synclines:

1. Belozyorsky 2. Konsky 3. Chertomlyksky 4. Sursky 5. Verkhovtsevsky 6. Inguletsky 7. Saksagansky 8. Annovsky 9. Zheltorechensky 10. fravoberegliny 11. Kremenchugsky

FIG. 1. M a p showing distribution of iron-chert formations of Ukrainian Shield.

The iron-chert formations of the Ukrainian Shield

FIG. 2. Diagrammatic seismo-geological section of Belozyorsky Lower iron-chert-metabasite series: 7, Porphyrite-schist forma- syncline. 1. Granites and migmatites; 2. Ultrabasite formation. tion; 8. Iron-schist-metabasite formation (iron-chert and green Upper iron-chert-schist-keratophyre series. 3, 4 and 5. Iron- schist zone); 9. Metabasite formation (green schists and chert-keratophyre-schist formation: (3. Schist-iron-chert zone amphibolites); 10. Tectonic dislocations; 12. Refraction and with keratophyre interlayers; 4. Jron-chert zone; 5. Schist- reflection boundaries (seismic evidence); 12. Faults (seismic iron-chert zone); 6. Schist-tuff-sand keratophyre formation. evidence).

400 O (O0 LOO 300 m - II I - -

FIG. 3. Diagrammatic section of the Konsky syncline. 1. Pink Talc-magnetites; 8. Metabasite formation of the upper green granites and migmatites; 2. Grey granites and migmatites. schist suite (KN,); 9. Iron-chert-metabasite formation of the Iron-chert-schist-keratophyre formation of the Veselyanska- Julyevsky suite (KN,): schist-spilite zone and iron-chert zone; Kirpotinsky suite (KB,). 3. Schist zone; 4. Iron-chert zone; 10. Metabasite formation of the lower Julyevsky suite (KN,); 5. Keratophyre-schist formation of the Veselyansko-Kirpo- 11. Tectonic dislocations. Districts: I, Kirpotinsky; II, Medium; tinsky subsuite (KBB). Ultrabasite formation of the Veselyansky III, Julyevsky. suite (a,). 6. Actinolites and chlorite-actinolite schists;

FIG. 4. Diagrammatic section of Verkhovtsevsky syncline. 1. Granites and migmatites. Upper Bazavluksky iron-chert- keratophyre-schist series: 2. Iron-chert-schist-keratophyre formation; 3. Schist-keratophyre formation; 4. Ultrabasite formation. Lower Bazavluksky iron-chert-metabasite series: 5. Upper metabasite formation with iron-chert interlayers; 6,7 and 8. Iron- chert-metabasite formation (6. Amphibolites of the third dividing zone; 7. Actinolites and talcs of the second and first dividing zone; 8. The first, second, third and fourth packet of iron-chert layers intercalated with green schists and apospi- lites); 9. Lower metabasite formation.

sw NO

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the Domotkansky side of the Verkhovtsevsky syncline. The second cycle of iron-chert deposits is found in the

upper suite of the Upper Bazavluksky series; sand-tuff- schist-keratophyre deposits of the middle suite are transitional upwards to iron-chert and tuff-schist-phyllite interlayers.

The iron-chert deposits are associated paragenetically with products of the dacite-rhyolitic volcanism. The thickness of iron-chert layers reaches 50-140 m. They are represented by magnetite-siderite-chlorite-quartz cherts. Iron- chert layers are traced along strike for a distance of 5 km; further their facies pinch out. The iron ratio also changes. The iron-chert formation attains a total thickness of 300- 500 m, iron-chert layers constituting 30-50 per cent of the formation.

In the Sursky brachy-syncline the thickness of the iron- chert-schist-keratophyre formation ranges from 300 to 500 m; On the northern side of the brachy-syncline the formation extends for 30 km.

Individual packets of iron-chert lenses attain a thickness of 90-100 m. Iron content, in most cases, however, is not high, averaging 10-12 per cent and, rarely, 30 per cent.

In the Chertomlyksko-Solyonovsky brachy-syncline iron-chert-schist-keratophyre formation (up to 400 m thick) is found in the Chertomlyksky band. Iron-chert beds are 100-150 m thick and 6 km long. The average iron content is 30-35 per cent. The formation consists of siderite-magnetite-quartz cherts.

The third zone of iron-chert distribution is present in the Orekhovo-Pavlogradsky synclinorium which extends for 100 km. The age of the oldest minerals is 2,300 m.y.

To the east of this zone another smaller iron-chert zone-the zone of the Korsaksky synclinorium-is developed. Some older minerals present in the area of the village of Stulnevo yield an age of 2,800 m.y. Further to the east, in the area of the shield adjacent to Azovsk Sea, small iron-chert deposits occur in the Mangushsky syncline which is evidently synchronous with the Orekhovo-Pavlogradsky zone.

Iron-chert deposits of the Orekhovo-Pavlogradsky syncline are represented by two formations: iron-chert- metabasite and iron-chert-schist-gneiss.

Folded structures are divided here into separate synclinal bands extending for 2-5 km. The rocks are metamorphosed to amphibolite gneiss and pyroxene gneiss; iron- chert rocks consist of coarse hornblende-magnetite quartzite and hypersthene-magnetite quartzite. Iron-chert lenses range in thickness from 10-20 to 100 m and extend for 2-5 km. The facies are less consistent and are often replaced with, and end in, granite. The folded zones represent an inverted synclinorium in which the rocks have been intensely granitized.

The metabasite Saksagansky series of the Kremen- chugsky region, with its iron-chert interlayers, evidently corresponds to the same structural stage.

The main structural zone of development of iron- chert deposits is the Krivoyroghsko-Kremenchugsky synclinorium, traced for 200 km in the meridional direction. The

age of the formation is found to be 2,000 to 1,800 m.y.; the latter age however, is debatable.

Sedimentation of the iron-chert-schist zone began with eruption of ultrabasic products represented by a talcose horizon 200 km long. The origin of the ultrabasic rocks is evidently connected with a deep fault. Keratophyre interlayers occur sometimes in the upper part of the talcose horizon as well, for example, in the Inguletsky syncline along the Timosheva ravine.

The rhythmic banding of iron-chert layers and alumosilicate schists in the Krivoyroghsky series is up to 1,500 m thick. Such a thickness is reached only in two troughs-the Saksagansky and Kremenchugsky, each of 50 km extent. Iron-chert-schist formation of the Krivoyroghsky series is unconformably overlain by the Inguletsky series, which is characterized by a change of sedimentation to a flyschoid type. They occur as compressed isoclinal, echelon-like folds extending along strike for a distance of 10-50 km.

The central structure of the Krivoyroghsky syncline (Fig. 5) passes into the Tarapakovsky anticline in the west; this is closed by the Inguletsky (Likhmanovsky) syncline, truncated from the west by a thrust extending all along the strike.

FIG. 5. Section of Krivoyroghsky series. 1. Inguletsky series; 2. Iron-chert-schist formation of Krivoyroghsky series; 3. Talc horizon; 4. Sand-schist lower suite of Krivoyroghsky series. I, Inguletsky syncline; II, Tarapakovsky syncline; III,, 2, 3, Kri- voyroghsky syncline (III,, Western Krivoyroghsky trough; III,, Sovetsky anticline; III,, Eastern Krivoyroghsky trough); IV, Saksagansky anticline; V, Saksagansky syncline.

The eastern limb of the Krivoyroghsky syncline is represented by the Saksagansky band which, for its own part, is complicated by the isoclinal Saksagansky syncline and anticline.

The curve of the Krivoyroghsky syncline borders the southern margin of the Saksagansky trough, the northern margin being rimmed by the Ternovsky flexure. Within the

138


m - IO0

200

300

400

500

600

100

ao0

4000

ízoo

900

ííoo

4300.

(400. 1500.

i600

limits of the Saksagansky trough some layers are encountered which consist of iron-chert units interbanded with schists.

The total maximum thickness of iron-chert strata is 500-800 m; it amounts to 50-75 per cent of the sequence. The average iron content in the iron-chert strata is 30- 40 per cent. Within the limits of the ñrst and second Saksa- gansky subsuites, a secondary metasomatic enrichment of iron-chert strata occurred and development of iron-rich deposits took place. Ferruginous strata at the surface are oxidized and martitized; they change into magnetite hornfelses at depth.

The cross-sections of the iron-chert-schist suite of the Saksagansky and other troughs are shown in Figure 6.

To the south of the Saksagansky trough, in the area of the Inguletsky syncline, a sharp decrease in thickness of the iron-chert-schist formation is observed for a distance of 20 km. The total thickness of the formation here is 300-400 m; in places it is only 50 m ‘thick. Iron-chert rocks constitute 50 per cent of the formation. The formation is represented by the ñrst, second, fourth and fifth strata each separated by a bench of schists.

To the north of the Saksagansky trough along the strike of the Annovsly syncline, a sudden decrease in thickness is observed along with a gradual pinching out of iron- chert facies.

Further to the north in the parallel ‘Zheltorechensky’ trough, the thickness of iron-chert deposits increases again to 700 m. Still further to the north in the Pravobereghny region a limited thickness (100-300 m) of iron-chert-schist formation is found over a distance of 40 km, with variable

- - . .

.

.

.

.

.

- .

-

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number of iron-chert strata. A facies replacement occurs along the strike of the schist and iron-chert layers; these constitute 50-70 per cent of the sequences as shown at the corresponding cross-sections (Fig. 6).

The Kremenchugsky syncline is developed on the right bank of the Dnieper in the northern part of the synclinorium. It is similar to the structure of the Krivoyroghsky syncline. In the Kremenchugsky trough iron-chert-schist formations once again approach thicknesses of 1,500 m. A total thickness of all the iron-chert strata in the Gales- chinsky syncline attains 500-700 m. The strata constitute 60-70 per cent of the sequence.

The presence of tufogene schists in the sediments of the Kxemenchugsky trough is a prominent feature, as is also a number of clastogene intercalations within the formation. Moroever, a pyrite horizon is encountered in the iron-chert- schist formation; it extends for a distance of 20 km. The lenses of pyrite in the horizon attain a thickness of 90 m.

Deposition of pyrite in the iron-chert formation of the Kremenchugsky region testifies to an alternation of oxidation-reduction regimes during sedimentation.

Tufogene schist intercalations found in the Kremen- chugsky trough suggest a genetic relationship between iron- chert-schists and volcanogene formations.

To conclude the consideration of environments of iron- chert formations it should be noted that they are also present in the Ananyevsky band of the magnetic anomaly in the western part of the Shield on the river Bug.

The present considerations show that cycles of deposition of iron-chert formations occurred repeatedly in structural

FIG. 6. Sections of iron-chert-schist formation along Krivoy- roghsko-Kremenchugsky synclinorium. 1. Iron-chert layers; chensky; V, Pravobereghny; VI, Kremenchugsky. 2. Iron-poor iron-chert-schist layers; 3, Schist layers. Regions:

I, Inguletsky; II, Saksagansky; III, Annovsky; IV, Zheltore-

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stages of the Ukrainian Precambrian ranging in age from 3,500 m.y. to 1,800-1,700 m.y.

Depositional environments and paragenetic associations of interbedded rocks have led to the conclusion that iron-chert formations are chemogenic deposits related genetically to submarine volcanic areas of geosynclines. Evi- dently their deposition may have been related to periods of interruption in volcanic eruption when submarine hydrothermal solfataric and fumarolic activities, developing over great areas, provided a special environment for deposition of siliceous-iron sediments precipitatiiig in a colloidal form. During the periods of interruption, which considerably exceeded in duration the periods of eruption itself, the deliv- ery of tuff-schists detritus did not supress chemogenic iron- chert deposits; these were not smothered in the mass of tuffaceous material but were evidently deposited in a border or peripheral playground of volcanic activity where remote schist-iron-chert deposits were formed.

According to the above paragenetic associations of rocks in interbedded iron-chert formations as illustrated in different structural zones, the author distinguishes three types of iron-formations: (a) iron-chert-schist, (b) iron- chert-schist-keratophyre and jaspilite-leptite, and (c) iron- chert-metabasite and iron-chert-ultrabasite (Fig. 7).

Iron-chert-schist formation is characterized by the highest degree of facies consistency. Continuous iron-chert zones extend for 10-20 km, often with a constant iron content of 30-35 per cent and, rarely, to 45 per cent. Iron- chert layers constitute 50-75 per cent of the sequence. Indi- vidual iron-chert layers with minor schist intercalations attain a thickness of 100-300 m . Schist interlayers are composed of mixed pelitic terrigenous-tufogenic and chemogenic iron-silicate deposits alternating with thin centimetric intercalations of chemogenic siliceous deposits. They are remote formations, deposited in conditions of prolonged and persistent hydrothermal activity of submarine volcanic areas with limited supply of ashy material and a wide distribution of chemogenic colloidal sediments.

The formation of the Krivoyroghsko-Kremenchugsky syncline zone and of the upper suite of the Belozyorsky

I O IOKrn

100 200 300 4ao 500 600 700 a m 900 i000 m

syncline is referred to this type; the formation passes into iron-schist-keratophyre only in the northern part of its strike.

Iron-chert-schist-keratophyre formation is made up of paragenetic associations of tuff-keratophyre-schist and iron- chert deposits formed in the area of submarine volcanism of dacite-rhyolitic lavas. In the area of gneissic stages of metamorphism, acid volcanogenic products have been transformed into leptites; that is why they have acquired the name jaspilite-leptite formations. A lower order of facies consistency and a greater diversity in iron content are their characteristic features.

The proportion of iron-chert deposits in the sequences is not high; it amounts to 20-40 per cent. However, individual lenses of iron-chert deposits are present up to 100- 150 m thick and 5-7 km long. Iron content is 30-35 per cent.

The upper suite of the Verkhne-Konsky series, the Teplovsky band of the Verkhne-Bazavluksky series (the Verkhovtsevsky syncline), etc., are also referred to iron- chert-schist-keratophyre formations.

The above formations are formed in the area of submarine volcanism of rhyolite-dacite andesitic composition. Deposition of iron-chert sediments proper is related to periods of extrusive dormancy and development of long periods of submarine hydrothermal activity.

Areas of acid volcanism are characterized by more abundant subvolcanic activity in comparison with areas of basaltic lavas; that is why iron-chert-schist-keratophyre formations are more intensively enriched with iron-chert deposits. Here they are of local extent only, being restricted to 5-10 km from the immediate proximity of volcanic foci.

Iron-chert-metabasite formations are deposited in areas of development of submarine volcanism of basic lavas. These formations occur in the lower metabasite series of the Konsky, Belozyorsky, Verkhovtsevsky and Sursky synclines. They are characterized by a small number of iron- chert intercalations of relatively limited thickness and extent. However, they compose in some places thick masses up to 500 m thick, in which siliceous intercalations amount to 20-30 per cent.

mI m2 isxi3 [vvv4 mi5 FIG. 7. Types of iron-chert formations. I, Iron-chert-schist 10-30 and 10 per cent of iron-chert deposits. 1. Iron-chert layers; formation with 50-90 per cent of iron-chert deposits; II, Iron- 2. Schist layers; 3. Keratophyre layers; 4. Metabasite layers; chert-schist-keratophyre formation with 30-50 per cent of iron- 5. Ultrabasite layers. chert deposits; III, Iron-chert-metabasite formation with

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The absence of accumulations of iron-chert beds of great extent associated with basic volcanics may evidently be explained by a lesser intensity of hydrothermal volcanogenic processes in such areas compared with areas of rhyolite-dacite volcanism. At the saine time, development of iron-chert intercalations makes it possible to evaluate a regularity of interruptions in extrusive volcanic processes observed during the accumulation of rather thick masses. W e observe a high proportion of iron-chert zones in a mass of tuff-volcanogene materials.

Iron-chert-ultrabasite formations have an insignificant distribution in the Kudashevsky area of the Konsky and Verkhovtsevsky synclines. The alternation of low grade iron-chert interlayers with ultrabasite schist is a prominent feature. The formations testify to manifestations of tuffaceous ultrabasite volcanism with extrusive interruptions also accompanied by submarine hydrothermal activity.

Pure iron-chert rocks consists only of Feo, Fe,O, and SiO, + H,O + CO,. Iron ratio calculated by the formula

(Fe0 + 2 Fe,O,) . 100 Alzo, + M g O + Ca0 + (Fe0 + 2 Fe,O,) ’

derived from chemical analysis, amounts to 95-100. Thus, in chemogene deposits we have a complete differentiation of iron and silica from other rock-forming oxides. With an increase in number of schist interlayers the iron ratio decreases to 60-90 per cent at the expense of admixtures of Alzo, and partly of MgO; in metabasite iron-chert formations Ca0 is added as well.

According to degree of oxidation, iron-chert rocks are divided into oxide, protoxide-oxide, oxide-protoxide and protoxide, characterized by the following coefficients (O), obtained from a chemical analysis by the formula

Résumé

Oxide rocks: O= 10-33 consist only of hematite and quartz;

Protoxide-oxide: O= 1.5-10 contain magnetite, hematite, quartz; magnetite, quartz.

Oxide-protoxide: O= 0.5-1.5 consist of magnetite, iron, silicate, quartz; magnetite, iron silicate, siderite, quartz.

Protoxide: O= 0.05-0.5 consist of iron silicate quartz; siderite, quartz.

The first two groups-oxide and protoxide-oxide-are iron-chert ore rocks in which iron is fully represented by magnetite and hematite.

The third group-oxide-protoxide-is represented by semi-ore iron-chert rocks in which only some iron is in the form of magnetite and the rest is fixed in silicates or carbonate.

The fourth group-protoxide-are nonmetalliferous iron-silicate-siliceous rocks in which all the iron is present in silicates or siderite. By metamorphic stages A slaty stage is characterized by banded iron-chert jaspers. A stage of phyllite and cherts is characterized by iron- quartz cherts or jaspilites (non-silicate iron-chert ores). Gneissic stages of metamorphism are characterized by coarse-crystalline iron quartzites.

According to iron ratio, iron-chert rocks are divided into iron-rich, with an iron content higher than 30 per cent, intermediate with 20-30 per cent, and iron-poor with less than 20 per cent. In primary, unchanged iron-chert deposits an iron content higher than 45 per cent has not been encountered. Increase of iron content to 60-70 per cent and formation of clusters of high-grade ore deposits are related to secondary metasomatic processes of enrichment.

In the process of oxidation iron-chert rocks are martitized to a great depth, exceeding 1 km.

Géologie et genèse des formations de fer siliceux du bouclier cristallin d’Ukraine (N. P. Semenenko)

Des formations de silex ferrugineux se rencontrent à diffé- rents niveaux structuraux du Précambrien du bouclier cristallin ukrainien ; elles remontent à 3 500-1 700 millions d’années.

Les zones de synclinoriums où se sont développées les formations de silex ferrugineux sont les suivantes : Konksko- Belozersky, Bazavluksky, Orekhovo-Pavlogradsky, Priazov- sky oriental et occidental, Krivorozhsko-Kremenchugsky.

La relation entre les formations de silex ferrugineux et les régions volcaniques géosynclinales est établie. O n peut distinguer les types suivants de formation : schistes de fer siliceux ; schistes kératophyre de fer siliceux ; ultrabasites de fer siliceux.

Les dépôts chémogéniques de silex ferrugineux sont représentés par des formations de leptochlorite siliceuse, de sidérite siliceuse, de pyrite siliceuse, et de silice ferrugineuse selon le potentiel redox Eh. Les roches de silex ferrugineux métarnorphisées sont classées en groupes ferriques, ferriques-ferreux et ferreux.

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N. P. Semenenko

Discussion

R. FRIETSCH. A small remark. The Grangesberg deposit cannot be used as an example of iron-chert-schist-keratophyre formation. The ore is mainly a massive magnetite- hematite lying in metamorphosed lavas of intermediate composition. To the ore field belong siliceous hematite impregnations (Lomberg type), but they are probably hydrothermal deposits following the main volcanism.

N. P. SEMENENKO. I first became familiar with the Granges- berg deposit with the help of the Swedish geologist Kautski. There, in subsurface, beds of iron banded quartzites are interbedded with leptites. Some samples of these rocks are in the exposition of the symposium.

R. FRIETSCH. What is the content of hydrogen in the hematite?

N. P. SEMENENKO. The amount of reducing agent in the iron quartzite is sometimes sufficient to reduce hematite to magnetite on heating in a soldered tube.

Z. T. TILEPOV. What is the distribution of the elements Pb, Zn, Mo, G e and Mn along the radius from the centre of igneous rocks?

N. P. SEMENENKO. Under the depositional conditions of iron-chert formations, these elements are not usually deposited.

J. H . GROSSI Sm. H o w much germanium is found associated with magnetite?

N. P. SEMENENKO. The amount of germanium is 20-30 g per ton.

A. F. TRENDALL. Are the ages quoted in the paper ages of deposition or of metamorphism?

N. P. SEMENENKO. The ages of the rocks refer to the time of their metamorphism. The ages of the oldest minerals are mentioned.

A. F. TRENDALL. What methods were used to obtain them?

N. P. SEMENENKO. The ages were determined by the K-Ar method for chert hornblendes, and by the U-Th-Pb method for accessory minerals.

S. J. SIMS. Where the iron chert has been granitized, is the granitic rock iron-rich?

N. P. SEMENENKO. The granites replacing or migmatizing iron quartzites are not iron-rich; the products of assimi- lation are usually removed. The process takes place under conditions of magmatic distillation.

G. A. GROSS. Are the thicknesses of iron-formation mentioned accumulative thicknesses of chert layers, or are these the thicknesses of iron-formation developed by folding?

N. P. SEMENENKO. The thicknesses cited are normal stratigraphic thicknesses.

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Occurrences of manganese in the Guianas (South America) and their relation with fundamental structures B. Choubert Directeur de recherches au Centre National de la Recherche Scientifique (France)

Occurrences of manganese have long been known to exist in the Precambrian formations of the Guiana shield. In Guyana (formerly British Guiana), they were discovered at the end of last century, and Harrison refers to them as early as 1908, but it was only about 1950 that active prospecting began, apparently as a result of the discovery in 1945 of a deposit in the Serra do Navío in the Brazilian territory of Amapá. Prospecting then spread to British Guiana in 1952, to French Guiana in 1955 and to Surinam.

The results were somewhat disappointing from the commercial point of view, but the description of the geological features is of considerable interest and the pages which follow give a summary of what is known.

Basic features of the geological structure of the Guiana shield

The Guiana shield is separated from the rest of South America by the Orinoco and Amazon rivers. In the north it plunges beneath the Atlantic Ocean over a distance of about 1,500 km. Politically this vast area comprises several territories: Venezuelan Guiana, Guyana, Surinam and French Guiana, plus the Brazilian territories of Amapá, Amazonas and Rio Branco.

The geological exploration of this vast territory is still far from complete. To the south, the investigations rarely extend below the second parallel: while Amazonia and the upper Orinoco basin are scarcely known.

The Guianas are a very ancient part of the earth‘s crust, the evolution of which appears to have stopped about 1,700 m.y. ago. Vast stretches of the shield consist of basi- cally granitic terrains. The terrains of sedimentary and volcanic origin are metamorphosed to various degrees. These terrains have been subdivided into a number of series, the names of which vary from country to country. All of them are assignable to the different stages of the geosynclinal evolution, and we get, in ascending order, the following series.

Lustre schists. Yuruari series in Venezuela, Barama series in Guyana, Lower Paramaca series in French Guiana and Surinam, containing thick beds of argillaceous schists, quartz phyllites with quartzite lenses, usually black carbonate rock, horizons of manganese (gondites) and some ferruginous quartzites with iron ore concentrations. Chlor- ite schists, spilites, keratophyres [Guyana) and small massifs of intrusive rock ranging from pyroxenite through diorites and gabbros to granodiorites.

Ophiolitic volcanism. Volcanic series with schists: El Callao and Pastora series in Venezuela, ‘volcanic series’ in Guyana, Upper Paramaca series in French Guiana with intercalary lavas and sediments: basalts andesites, dacites, rhyolites, pillow lava (Venezuela), amygdaloidal lavas, products of submarine volcanism, pyroclastic deposits, jaspers shifting to quartzites, argillaceous schists.

Flysch. Caballape series in Venezuela, Cuyuni series in Guyana, Bonidoro series (north facies) in French Guiana: detrital rocks, a series of very thin layers of variable composition: conglomerates, greywacke, argillaceous schists, etc.

Paramolusses (or lower molasses). Known as Haïma- raka in Guyana, Rosebel and Armina in Surinam, Bonidoro (southern facies) and Orapu in French Guiana. Thick series consisting of arkoses and conglomerates, of argillaceous schists lying in transgression over older rock formations. Schists ranging in colours from grey to violet or sometimes greenish with black intercalations of carbonaceous substances. At the base are polymictic and monomictic conglomerates, mainly becoming important to the east.

According to certain writers, the combined thicknesses of these Precambrian deposits reach 14,000 m in Venezuela and 8,000 m in French Guiana.

Post-paroxysmal molasse and products of terminal basic volcanism, exhibiting marked discordance, overlie all the folded and eroded series described above, to form the Roraïma series which has remained sub-horizontal and consists of coarse pebbles, conglomerates with lava coulees and dolerite and gabbro massifs and veins. This molasse, which has undergone no granitization, is not found in French Guiana. It forms a butte-témoin in Surinam and


B. Choubert

covers vast areas in Guyana and Venezuelan Guiana. Apart from these thick geosynclinal terrains, still older

rocks appear in places which represent the pre-geosynclinal platform with its cover of volcanic and sedimentary rock.

This platform is made up of catametamorphic facies (Venezuela, Guyana and Surinam) or mesometamorphic facies (French Guiana), while the cover is of rhyodacites and pelitic and sandstone series, plus iron ore. The formations were granitized contemporaneously with the series in the geosynclinal domain, and consequently belong to the same chain.

GRANITIZATION

The main sanitization dates back to between 2,050 and 1,750 m.y. and produced the Caribbean granites which consist, on average, of monzonite granite with microcline, plagioclase and biotite. These rocks have engendered a large quantity of pegmatites and vast fields of felspathized rocks. Most of the earlier rocks were rejuvenated by this mechanism which makes their absolute ages identical with those of the granites.

The Guyana granites are of an earlier age which is difficult to determine in view of the general rejuvenatiai. They can be dated after the formation of the volcano-sedimentary series and before the deposition of the paramolasses for which they supplied the material by erosion. They must be deemed the outcome of a remobilization and remodelling of the earlier granitoids which were formed between 2,700 and 2,500 m.y., that is, during the pre-geosynclinal period. They have, nevertheless, retained fairly constant characteristics, which are reflected by their chemical composition. They are relatively poor in potassium, with acid plagioclase and biotite with or without hornblende as their principal constituents: the commonest types are alternatively graiio- diorites and akerite granites.

TECTONICS

For a proper understanding of the tectonics of the chain, it should be remembered that what is today visible on the surface represents the geometry of the deepest zones of this Archaean edifice. It can be conjectured that erosion has removed a thickness of 5-10 km of an undeducible structure of which only the roots are visible. The wide granitized cupolas and the synclinoria in which the parametamorphic rocks have survived represent the folding of the bottom strata among which the dislocations have brought blocks of the pre-geosynclinal basement to the surface here and there, in a more or less remodelled and rejuvenated state.

Occurrences of manganese: their composition

Manganese is closely associated with what may be roughly called the volcano-sedimentary rocks and occurs as lenticular intercalations in the topmost levels of the Barama series (Guyana) and Lower Paramaca (French Guiana and Surinam).

The main occurrences known, with indications of the 'primary' manganiferous rocks, are shown in Table 1.

TABLE 1. Main occurrences of manganese

Longitude/latitude 'Primary' manganiferous rock

Amapá Serra do Navio, on the Gondites, carbonates, Amapari M n 55"05' O"55'

French Guiana Observatory Mountains Kaw Mountains

Upper Sinnamary

Mt Richard Ampouman (Maroni River)

Grand Inini

Surinam Goeje Mountains

Poeketi, on the Tapa- nahoni

Apoema Maripa, Piqué Heuvel

Afoebaka, Brokopondo

Adarnpada

Guyana Saxacalli, on the Esse- quibo

Kutuau (Cuyuni basin) Tasawinni

Pipiani

Matthews Ridge

Arakaka

Manganiferous schists 51"37' 4"08' (( x 52"08' 4"34'

to 52"05' 4"31' Quartz-rich gondites approx. 52"50' 4'30'

Gondites 54"13' 4"43'

54"25' 4"38' 53'45' 3"32'

Gondites 54"Ol' 3"3O' to 5490' 3"2W

(( 54"35' 490' (( 54030' 4"37'

Gondites, braunite (traces) 54"53' 4"42'

Impregnations, small veins in the schists 55"OO' 5'00'

Gondites 56"50 4"23'

Gondites 58"40' 6"35' (( 59"20' 6O.53'

Residual gondite blocks 59"35' 7"28'

Residual gondite blocks 59"43' 7"22'

Mainly gondite, plus braunite 60"lO' 7"29'

Mainly gondite, plus braunite 59"58' 7"35'

As can be seen, most of these are associated with the gondites. These are rocks of rather unusual composition, of which the essential part is spessartite garnet: Mn,AI, (SiO&. Manganiferous rocks of this type are also found

I44

Occurrences of manganese in the Guianas (South America) and their relation with fundamental structures

in Brazil, at Minas Gerais, where they were described by Derby as early as the beginning of the century (1908) under the name of queluzites (Morro da Mina). In the opinion of Hussak (1906), the silicates have their origin in the carbonates of M n under the action of metamorphism, and these carbonates probably represent the protore. In 1909, Fermor, in his study on Mn deposits in India, gave rocks of the same group the name of gondites and it is this name which has gained acceptance in the literature.

Over and above the manganiferous garnet, the gondites may contain quartz, amphiboles (tremolite, actinolite, grunerite), biotite, sericite, carbonaceous matter (LIP to 2 per cent), plus carbonates of M n (dialogite, etc.). The proportions of the mineral elements vary very considerably. The most frequent paragenesis in the Guianas is garnet + quartz, with additions of mica or amphibole. As the garnet content rises we get actual garnet rocks, formed by the action of general metamorphism, and therefore to be distinguished from the skarns which, though rich in garnet, are products of contact metamorphism.

The manganiferous garnet is generally white or grey and shows a wide range of variations in composition. It is thought to represent a mixture of garnets whose individual composition varies very widely-spessarites, pyrope, almandine, grossuralite, andradite-with a little Tio, and carbonaceous matter.

Generally speaking these rocks have undergone a profound alteration, with the formation of black oxides of M n (pyrolusite, polyanite, wad, etc.) which render observation difficult.

To permit a better grasp of the geological characteristics of these deposits, descriptions of three of them are given below.

A M P O U M A N FALLS, M T RICHARD (FRENCH GUIANA)

The geology of the manganese occurrences is fairly well known in this region, which is traversed by the Maroni. In places, the river is as much as 4 km wide, and is dispersed in a multitude of channels which girdle innumerable islands

and rocks. Observation in detail is possible here, whereas in other sectors a thick decomposition layer carried the dense equatorial forest, and masks the hard rock.

The Lower Paramaca (Fig. 1) outcrops extensively with nearby E.-W. alignments, shifting progressively to NE.-SW. as one moves eastward. Overlying it, and in marked unconformity with it, is the Bonidoro series, which begins with a polymictic conglomerate above which comes the usual succession of coarse arkosic sandstones. Between Ampouman Falls to the south and Boëli-Mofou to the north-that is, over a length of 1-2.5 km-it includes chloritoid quartzites of variable 'sandiness' with lenses of carbonate-rich rocks, black carbonaceous schists and quartzites (dark and light stratification), with glances and gondites rich in fine- or coarse-grained garnets and showing an undulating Stratification. In addition to small elongated massifs or veins of gabbros, the series contains diorites and dolerites. Directions of dip are N. or NNW. (30" to 45").

There are two levels rich in lenses of carbonate rock, with surface breadths of 200-300 m and intercalated in quartzose schists. The southern strip is flanked by manganiferous levels each about 100 m thick (see section) at distances of about 600 ni, on either side.

According to Brouwer (1960), the garnet rock lenses form segregations in the quartzose schists where the garnet it less abundant and the chloritoids very common, plus biotite, calcite, epidote and amphibole. The lenses are of all sizes, the smallest being only some 10 m thick.

According to Jaffé, the composition of the Maroni garnets is 57 per cent spessartite to 33 per cent almandine and 10 per cent pyrope. Averaged out, they contain 14 per cent MnO and 38.93 per cent SiO,. The lenses of black carbonaceous rock have a maximum length of 1 km and a maximum thickness of 25 m , and are sometimes of argillaceous sandstone, sometimes entirely quartzites, in all cases very fine-grained. They show only traces of MnO (0.01 per cent), but may have as much as 90 per cent of SiO,, 9.5 per cent of M,O, and 5.6 per cent of Cao.

Southeast of this manganiferous zone, ferruginous rocks are found intercalated in talcose schists. The iron-ore outcrops (85 per cent Fe,O,) are separated from the gondite band by transgressions of the Bonidoro series which make

N . . M A R O N I R I V E R c--- S

m Bkli Mofou Falls Ampouman Falls 200 -, Pe,te SargougFa1l.s . i

!lI 3 4 6 5 Y 12 4

O

FIG. 1. Profile along the Maroni River to Ampouinan. Bonidoro: 1. Coarse sandstone; 2. Conglomerate. Lower Paramaca: calcium carbonate rock lenses; 6. Dioritic rock dykes.

3. Quartzite schists with chloritoid; 4. Gondite levels; 5. Level of

145

B. Choubert

it impossible to see the relationship of this indicator with the manganiferous mineralization. This is a lenticular deposit of low tonnage in association with altered talcose schists (goethite and specularite, with a little quartz: SiOz = 6.5 per cent).

The belt traversed by the Maroni between Ampouman Falls and Boëli-Moufou Falls, extends for about 5 km westwards into Surinam territory and about 20 km NE. into French Guiana, or 29 k m in all. The metamorphism is more marked to the east, where the Paramaca is in contact with a considerable massif o€ Caribbean granite (Massif de l'Espérance). The two zones of fine-grained gondites are intercalated in amphibole rock and compact amphiboles, which have a north-westerly dip and form a fairly high ridge with a lateritic revetment highly manganiferous in places at the base. The texture is often pisolitic.

Ferruginous quartzites have been found in the same region, but their stratigraphic position is not well known. Transformed lavas (green rock) exist on the Maroni near the Surinam bank (Ampouman Falls), as well as inland on the N W . edge of the band.

MARIP A ( s URINAM)

This deposit, the characteristics of which are known through the survey work of Holtrop, is in the basin of Sarah Creek (a right-bank tributary of the River Surinam). Two major indicators are known in this region: Maripa Heuvel and Piqué Heuvel (= hill). Several types of exploratory operations have been effected: detailed surveys, boring, trenching, driving horizontal galleries. Only the essential is mentioned here. Further details can be found in the book by Holtrop (1962).

The Paramaca series here consists mainly of quartzite schists with chlorite, biotite and sericite. In the schists are lenticular intercalations of gondites, carbonaceous schists, ferruginous quartzites, with lenses several hundred metres to 2 km in length and about 100 m thick. The carbonaceous schists occur mainly at the base of the gondite horizons. Elsewhere, the gondites may give place to carbonaceous schists. All this is explicable on the assumption that, in anaerobic conditions and as a result of variations of p H and Eh, Fe and Mn precipitate alternately, M n in the form of carbonates and Fe in the form of oxides. In addition, the appearances of M n and carbonaceous matter are definitely connected, whereas the ferruginous quartzites appear to be independent.

Several parallel alignments directed 10" N. to 20" W. of gondite lenses are known in a zone 700-800 m wide and, rightly or wrongly, it has been held that this distribution came from the folding of a single maiiganiferous bed.

Depthwise, a few trial borings give the stratigraphic succession of the mineralized zones and we find that there are many of thin gondite seams, intercalated in a yellowish-brown decomposition clay. They run to about fifteen (LA-120) for a depth of about 41 m, their thickness varying between 10 c m and a little more than 1 m.

Sample analyses of the manganiferous rock from two borings and a well 21 ni deep show an average composition of: MnO, = 33.28 per cent, Fe20, = 5.74 per cent, A1,0, = 7.89 per cent, and SiO, = 36.38 per cent.

In some places, concentrations of garnet account for 80-90 per cent of the rock and contain an average of 27 per cent of MnO.

Holtrop calculates that there should be a reserve of about 143,000 tons of metallic M n at Maripa.

MATTHEWS RIDGE, A R A K A K A (NW. GUYANA)

Here the manganiferous rock is intercalated in the thick Barama schistose series, equivalent to the Lower Paramaca of French Guiana (see also Holtrop (1962)).

Argillaceous and quartzose schists with sericite occur, with horizons rich in carbonaceous matter and sterile lenticular quartzites, including a roughly 300 m horizon of gondites, ferruginous quartzites and beds of braunite with intercalations of metamorphed chert. Investigations have even pin-pointed two manganiferous zones and thin conglomerate bands with pebbles averaging 2.5 c m in diameter.

The manganiferous zone, which is 150-175 m thick, E.-W. at Matthews Ridge, gradually swings to a NE.-SW. orientation eastwards when, after crossing the Arakaka creek, it runs along the River Barima for a total distance of 22 km. Dips are northerly with 50" to 80" inclination.

Borings show the stratigraphy in depth to be as follows. The ñrst SO m are divided into two parts: metadolerites

topping a manganiferous horizon, which itself incorporates a further intercalated seam of metadolerites. Below, in order of descent over a total depth of 1.5 m, we have 55 cm of brown and yellow clay, then alternating thin beds of gondites and brown clays, or fourteen manganiferous beds in all in a thickness of 95 cm.

In the upper part, lying between two beds of dolerite, we have, at the top, five intercalations of gondites alternating with white clay (to a total thickness of 25 cm), then a series of dark and light clays to a total thickness of 1 m; and once again, two 3-4 c m beds of gondites sandwiched in a white clay, after a seam of black clay and a thin horizon of quartz shingle.

The levels of braunite have an irregular geographical distribution and thicknesses ranging from 1 to 3 cm. O n the other hand, lenses of over 1.5 m thickness are found farther west.

The manganiferous band appears to be, in all, some 30 km long.

Exploitation of the Matthews Ridge deposit started in 1960, with extraction planned at about 1,200 tons daily. Reserves were estimated at some 13 million tons.

EFFECTS OF ALTERATION IN THE SUPERGENE ZONE

In the geological description we only discussed 'primary' M n minerals. Near the surface, the silicates, car-

146


bonates, etc., have been transformed into secondary oxides such as pyrolusite, polianite, psilomelane, manganite, cryptomelane, lithiophorite, or wad. Trial borings provide evidence of impoverishment of the ores as depth increases. In the first 30 m tenor in Mn drops from 40 per cent to 12 per cent, which demonstrates the relative poverty of the primary deposits. Enrichment occurs above the hydrostatic level, which is very high in the Guianas, as these countries belong to the humid sub-equatorial zone. This process of supergene alteration and accumulation is at all points comparable to that of laterization.

Over the whole Guiana peneplain the lateritic cover- ings have been formed at different epochs. At the base of these, accumulations of manganese are to be found here and there (Kaw Mountains, Observatory Mountains, Mt Richard, etc.) in the form of blocks embedded in the clay, pisoliths, nodules, often cemented by the laterite.

The metal content of these secondary ores is relatively high and may exceed 30 per cent (Mt Richard). Large blocks embedded in residual clay-at Tasawinni and Pipiani in Guyana or at Serra do Navío in Amapá-contain up to 40 per cent and 50 per cent of metallic Mn, with some percentages of Fe,O,, Alzo, and SiO, (sometimes > 10 per cent).

Through a diversity of treatments (washings, etc.), satisfactory concentrations can be secured. According to Nagell and Seara (1961) the metallic M n tenor of the commercial ore exported from Serra do Navío in 1957 was 48.5 per cent. The remainder is made up of Fe,O,, A1,0, and Sioz, with some other elements: Cu0 (0.07 per cent), Ba0 (0.12 per cent), As,O, (0.25 per cent), P, etc.

The problems presented in the Guianas by gondite-type deposits are not of a qualitative but of a quantitative order.

GENERAL CHARACTERISTICS O F THE GUIANA DEPOSITS A N D OCCURRENCES

The characteristics we have been describing have points in common, and this is equally true as regards the circumstances of the deposits’ formation.

In every region the primary ore consists mainly either of rocks with a manganiferous garnet content broadly described as gondites, or manganese silicates such as the braunites (Matthews Ridge, Guyana) or carbonates (dialogite) as at Serra do Navío in Amapá and at Upata in Venezuela.

At Serra do Navío it has been observed that here and there the dialogite gives place to garnetiferous rock still containing dialogite. It is thus apparent that the garnet and the calcic amphibole (tremolite) are genetically related to the manganese carbonates. It is legitimate to think that the primary ore everywhere consisted of carbonates and that the subsequent transformations are ascribable to the general metamorphism, which affords some confirmation of the earlier observations made by Derby in Brazil in 1901.

The deposits are aligned in zones several kilometres long-12 kin in the Kaw Mountains, about 30 km in the

Ampouman-Mt Richard sector, 30 km also at Matthews Ridge and Arakaka-with a breadth of 50-175 m . Where the breadth is above 175 m, it is as a result of folding.

Within these zones, the ores form lenticular beds of a thickness varying €rom a few millimetres to several metres. Their length-in the Serra do Navío, for example-is sometimes as much as several kilometres.

The lenses o€ ore alternate with ferruginous clays which are the result of the decomposition of more or less metamorphic argillaceous schists. These rocks contain chlorite, sericite, and often too (Ampouman) biotite and chloritoid and are fairly rich in quartz.

Associated with them are sediments, always the same ones: black carbonaceous rock (argillaceous or quartzose) in lenses, quartzite lenses, ferruginous or non-ferruginous (elsewhere silicified in cherts), all of them, from the spatial point of view, having the same discontinuous character as the ore.

There is no doubt whatever that the amphibolites which occasionally accompany the deposits are in very many cases of sedimentary origin. At Ampouman transitional members have been observed in the series from carbonate rocks to amphibolites. The latter predominate at Mt Ri- chard, where there are no longer any carbonates and where the quartzose and black schistose and more or less graphitic rocks always go with the manganiferous levels.

Again, it is certain that the slightly metamorphic carbonated rocks have generally been replaced by silica. A proportion of very fine-grained quartzites and cherts or phtanites are products of this secondary silification, which is ascribable less to a metamorphic process than to the influence of a hot, damp climate.

W e are dealing with deposits of sedimentary origin. Chemical precipitation alone seems capable of ensuring an alternation of beds whose regularity is revealed by the borings. The relatively low iron content of the manganiferous levels (about 1-2 per cent in fresh rock) contrasts sharply with the abundance of the metal in the shale intercalations (up to more than 10 per cent of Fe,O,) and iii the fresh metamorphic schists, which are often rich in chloritoid (H,Fe Al,SiO,, whose iron tenor js frequently as high as 28 per cent). There is, therefore, nothing surprising in finding ferruginous intercalations (hematite, limonite) in some deposits (Saxacalli, for example), and that in numerous instances beds of quartzites of sedimentary origin are found contiguous with, or close to, the manganiferous beds.

As regards the volcanic and eruptive rocks of the ophiolitic suite which exist in ihe deposits or near the occurrences of M n , their role in the accumulation of the ore continues to be obscure. Following the view of certain authorities (Taliaferro and Hudson in California, Routhier and Arnould in New Caledonia, Zanone in Ivory Coast), it is generally accepted that the M n deposits are genetically related to ophiolites, with jaspers and tuffs. The investigations carried out in the Guianas have not yielded additional proofs for this postulate. The presence in the Guianas of ophiolitic intrusions in the Archaean chain of geosynclinal origin requires no further proof, but no deposit

147

B. Choubert

has so far provided evidence of the mineralizing role of these rocks in this shield.

In some of the deposits mentioned earlier there are small massifs of eruptive rock consisting of quartzite diorites or gabbros (Ampouman) or metadolerites (Matthews Ridge), but these rocks were emplaced in the sediments which had already been deposited and their influence appears to stop at a slight increase in the iron tenors in their immediate neighbourhood.

Webber (1952) has asserted the existence of mineralized ‘tuffs’ in the Barama sector (Pipiani, etc.) in Guyana, but this is disputed by Holtrop (1962). Other authorities see a connexion between Mn and the amphibolites to which they ascribe an igneous origin. However, the discovery in French Guiana of an incontrovertible connexion between amphibolites and carbonate rocks of sedimentary origin renders this attribution suspect in most instances.

It is consequently difficult to contend that the ophiolitic intrusions known to us are the source of the manganese subsequently deposited in the form of chemical precipitates. M y own opinion is rather that the Mii comes from the ultra-basic rocks emplaced along the deep fractures which occurred in the pre-geosynclinal basement and were followed by the formation of troughs. The decomposition and leaching of these rocks prior to sedimentation could have supplied the manganese thus accumulated to the sediments in formation.

I have already published a description of the gilbertite serpentine changing to talcose schists with magnetite crystals in the Kaw Mountains in French Guiana. In the severely altered state, everything is transformed into clay and becomes unrecognizable. Above this come laterite carapaces containing concentrations of M n oxides at the base. These concentrations can reasonably be assumed to be due to the process of Mn migration under the action of the climatic decomposition of the subjacent basic rock; we thus get an indirect proof of the connexion between these rocks and the accumulations of manganese.

Ultrabasic rocks are also known in the south of French Guiana and of Surinam (Goeje Mountains), while Holtrop (1962) reports the existence of pyroxenites giving place to talc-schists in the immediate neighbourhood of the Piqué Heuvel gondite deposits, to the NE. of Maripa. Both seem to be earlier than the Lower Paramaca schists and, ajòu- ti&, than the predominately andesite ophiolitic intrusions of the Upper Paramaca.

This assumption would explain the rectilinear or slightly curved nature of the manganiferous zones, the narrowness of the zones (less than 200 m), as well as their length, which is relatively considerable along the line of the chain and much less in the normal direction. Conversely, the frequentlyregular pattern of the deposits, with alternation of Fe-rich and Mn-rich beds, with the simultaneous formation of carbonate and siliceous strata, accords with the physico- chemical modifications during the process of deposition according to the theories of Pieruccini (1956) on the geochemistry of M n deposits, which fit in perfectly with m y observations.

The concentrations of high tenor iron ore found here and there (Kaw Mountains, Ampouman, Mt Richard) close to the nianganiferous zones do not seem to be directly trace- able to the concentration mechanisms just described, although they share a common origin with the M n (ultrabasic rocks).

As regards the genesis of the ribboned ferruginous quartzites known as itabirites, it can safely be said that nowhere in the Guianas are they to be found in conjunction with the manganese deposits of the type just described. They do, on the other hand, characterize the vestiges of the former semi-platform and their mode of formation appears to be different,

GENERAL DISTRIBUTION IN RELATION TO THE DEEP TECTONIC STRUCTURE

The problem now is to determine the extent to which there is an effective relation between the manganiferous concentrations and the troughs of a geosynclinal system. The present study has shown the existence in various deposits either of carbonate rocks (Ampouman), of argillaceous successions with horizons of shingle (Matthews Ridge), of quartzite successions (Upper Sinnamary) or again of phyllite successions, with fairly considerable quantities of ferruginous quartzites in contact with older serpentines (Kaw). These variations in the formation of deposits hardly agree with the notion of uniformity generally evoked when one talks of lustred schists.

T o attempt to shed light on this question let us turn to the ‘trend‘ map for potassium established on the basis of the analyses of granites (Fig. 2). It satisfactorily reflects (seismic reflection model) the deep structure of the Guiana geosynclinal zone, in particular the distribution of troughs and intermediate masses (Choubert and Vistelius, 1969).

The geological map does not allow an exact idea of this distribution to be secured, as the depth to which erosion has reached leaves no more to be seen than mainly granitic regions and sedimentary zones representing only the roots of the former mountain structures. The isolines plotted from the seismic reflection model reveal the existence of two zones of maximum K20 tenors, separated by a zone of minimum tenors, all with an alignment varying from WNW.-ESE. to NW.-SE., i.e. along the line of the former chain. The maxima correspond to regions where the predominant granites are the Caribbean, considerably richer in potassium than the older Guiana formations, which on the contrary coincide with the zones of minimum tenors. The former indicate the old troughs, the others the intermediate masses.

On plotting the known deposits and occurrences on the map, we find that they are all in the zone of minimum tenors, inside the 3.2 per cent K20 isoline, in other words, along the median ridge (‘Zwischengebirge’) separating the troughs or along the edge of the semi-platform (Marudi, Serra do Navío). This distribution strikes m e as interesting,

148


6+

4 +

o + 6)

+ 61

i 60

+ 58

+ O I O 0 POO 300 km

+

+ '56

+

+

FIG. 2. Situation map of the occurrences of manganese according to the potassium isolines of the granites. 1. Occurrences of manganese (one circle), manganese deposits (two circles); 2. Out-

crops of Lower Paramaca-Barama; 3. Isolines of K,O, 'trends'; 4. Percentages of K,O tenor; 5. Political frontiers.

for it explains the variations in composition of the geological environment, the zone of intermediate masses having been subjected to frequent movements arising from the orogenic evolution of the troughs-two N.-S. eu-mio bi-couples- accompanied by major longitudinal and transverse dislocations still evidenced by the highly variable directions which I observed in the parametamorphic series.

In a work now in press (Choubert, 1969 a), I have tried to establish the geological correlations between the Guianas and West Africa. This comparison brings out the common

features not only in the lithostratigraphy, but also in the tectonic structure and the mineralization of the African and South American continents which are today separated by the Atlantic. The gondite manganese deposits known in the Ivory Coast and Ghana, represent the eastern extension of those discovered in the Guianas, and although their stratigraphy is still unknown, the geological context and mineralization are found to offer close analogies. The same deposition conditions apparently existed between 2,400 and 1,700 m.y. over the entire Guiana-Ivory Coast chain.

1 49

B. Choubert

Résumé

Les indices de manganèse clans les Guyanes (Amérique du Sud) et leurs dutions avec les structures foiidamentales (E. Choubert)

Dans les Guyanes, le Précambrien est subdivisé en un certain nombre de séries d’origine sédimentaire. L’analyse lithostratigraphique permet d’établir qu’il s’agit d‘une succession géosynclinale. Ces séries, plissées et métamorphi- sées, forment les racines d’une chaîne archéenne, profon- dément érodée, dirigée du nord-ouest - sud-est vers l’ouest nord-ouest - est-sud-est.

Le granite qui couvre de grandes surfaces appartient ti deux types principaux, le plus ancien étant moins riche en K,O que l’autre. Les isolignes de K,O, construites en partant de l’analyse chimique des granites, mettent en évidence deux zones de maximums sensiblement parallèles, séparées par une zone de minimums, qui peuvent être considérées comme des dépressions géosynclinales de part et d’autre

d‘une masse médiane. Dans le sud s’amorce le bord de la semi-plate-€orme qui marque la limite de I’airegéosynclinale.

La répartition de ces indices de manganèse dépend de ces structures fondamentales, qui sont presque toujours concentrées le long de la chaîne médiane ou le long de la bordure de la semi-plate-forme. D u point de vue stratigraphique, elles sont liées à la partie sédimentaire volcanique de la chaîne ancienne appelée, en Guyane française et au Surinam, les séries de Paramaca. Génétiquement, elles ont certainement une origine sédimentaire. Parfois, elles sont apparentées aux (( gondites D, parfois elles représentent des concentrations qui ont pu provenir de phillites contenant une certaine quantité de manganèse. En divers points, le manganese est accompagné de concentrations lenticulaires de fer. Les processus qui ont conduit à un enrichissement secondaire ont pris une place proéminente dans la formation des divers dépôts, le plus important étant celui de la Serra do Navío à Amapá (Guyane brésilienne).


BRACEWELL, S. 1947. The geology and mineral resources of british Guiana. Bull. Imp. Inst., Loncl., vol. XLV, no. 1, p. 47-59.

BROUWER, G. C. 1960. Sur la géologie et la métallogénie du massif de l’Espérance. Unpublished technical memorandum, Cayenne. __ . 1964. Feuille de Paul Isnard avec notice explicative, au l/lGO 000. Curte géologique détaillée de la France, Dép. de Za Guyane. Paris, Imprimerie Nationale.

CHOUBERT, B. 1954. Sur ¡es roches éruptives basiques des mon- tagnes de Kaw et de Roura (Guyane française). C.R. Acad. Sci., Paris,, t. 239, p. 185-7. __ . 1965. Etat de nos connaissances sur la gkologie de la Guyane française. B d . Soc. géol., vol. VII, p. 129-35. __ , 1969~. Les Guyano-Eburnëides de l’Amérique du Sud et de l’Afrique occidentale (essai de comparaison géologique). (In press.) - . 19690. Le Précambrien des Guyanes, BRGM Memo. (In press.)

CHOUBERT, B.; BROUWER, G. C. 1960. Stratigraphie de la série de Paramaca en Guyane française. C.R. Acad. Sei., Paris,

CHOUBERT, B.; VISTELIUS, A. B. 1969. Essai d’interprétation de la structure profonde des Guyanes par des moyens mathé- matiques. 4th Venezuelan Geological Congress, Caracas, Minis- terio de Minas e Hydrocarburos. (In press.) -- . 1970. D e la relation entre la composition des granites du bouclier guyanais et la position de ceux-ci dans les structures tectoniques majeures. (In press.)

t. 251, p. 109-11.

DERBY, O. A. 1908. On the original type of the manganese ore deposits of the Quelus District, Minas Gerais, Brazil. An. J. Sci., no. 175, p. 213-16.

DORR, J. Van N. II; PARK, C. F.; PANA, E. G. de. 1950. Deposito de manganeso do Distrito da Serra do Navio, Territorio Federal do Amapá. Bol. Dep. Prod. min., Rio de J., no. 85.

EBERT, H. 1961. Gondites and charnockites as Guide Horizon in the Brazilian Shield. Minutes, 5th Inter-Guiana Geol. Conf., Georgetown, 1959.

FERMOR, L. L. 1909. The manganese deposits of India. Bull. geol. Siirv. Zndia, vol. 37.

HARRISON, J. B. 1908. Geology of the GoId Fields of British Guiana. London, Dulau.

HOLTROP, J. F. 1962. D e mangaanafzettingen van het Guiana Schlid [The manganese deposits of the Guiana Shield]. Leiden (thesis).

HUSSAK, E. 1906. Ueber die manganerzlager Brasiliens. 2. prukt. Geol., no. 14, p. 237-9.

JAPFE, F.; BROUWER, G. C. 1959. Sur la présence de gondites en Guyane française. C.R. Acad. Sei., Paris, t. 249, p. 148-9.

NAGELL, R. H.; SEARA, A. C. 1961. The geology and mining of the Serra do Navío, Manganese deposit Amapá, Brazil. Minutes, 5th Inter-Guiam Geol. Conf., Georgetown, 1959.

PrERuccmr, R. 1956. Analche considerazione sul comportamento geochimico e sul ciclo geochimico del manganese. XX Znt. geol. Congr., Mexico. Symposium sobre yacimientos de manganeso, t. I.

WEBBER, D. N. 1952. Manganese Deposits in the North-West Dist. Brit. Giiiana Geol. Surv., no. 23.

150


Discussion

J. H. GROSSI SAD. The queluzite from Lafayette (ex- Queluz), according to Derby and other geologists, is an Mn-protore with carbonate. Garnet is not essential in the mineralogy of the rock. You have a typical analysis of queluzite on page 6 of the French original.

B. CHOUBERT. Queluzites and gondites are sedimentary rocks which have no stable composition. Besides, these rocks were metamorphosed under quite different conditions: epi-meso and catazone. Transition of dialogite rocks to garnet rocks was observed in Serra do Navío, Amapá. In addition, in the hot and humid climate of Guiana, carbonates are replaced by silicon and are transformed into jasper and granular quartzites. It is, therefore, impossible to set up any subdivisions in gonditoid rocks.

S. ROY. The gondites described in this paper do not exactly simulate those of India first described and named by Fer- mor. The so-called gondites of the Guianas and Brazil were originally made up of MnCO, with associated silica, whereas the Indian gondites are always carbonate-free. Under these circumstances should we accept the term gondite for both carbonate-rich and carbonate-free manganiferous rocks, or should we revive the term queluzite for the South American manganese silicate-carbonate protore?

B. CHOUBERT. I have already answered a similar question by Professor Grossi Sad. Besides, it should be mentioned that the paragenesis of most Guiana gondites is garnet-quartz -a little mica and carbonaceous substances.

E. RIBEIRO FILHO. According to Odman’s paper the rhodonite of the Morro da Mina was formed by a metasomatic process. I would say that the same happened in Chandi District.

B. CHOUBERT. I have never observed rhodonite in the Guiana manganese deposits. U p to now rhodonite is known only in Serra do Navío (Amapá).

G. A. GROSS. Are the volcanic rocks and amphibolites associated with the stratiform manganese deposits rich in potassium or in soda? Has this aspect been investigated?

B. CHOUBERT. This aspect has not been specially investigated, but I have observed the mineralogical composition of amphibolites in sections of quite ordinary hornblende, plagioclase and quartz in different proportions.

W. SCARPELLI. In Serra do Navío the manganese-bearing metasediments have a lateral extension of more than 10 km, and the manganese zones appear in all this extent. Also, all evidence points to a shallow deposition of the manganese rocks. H o w does this fit in with your idea of the carbonate being deposited in deep troughs of geosyncline?

B. CHOUBERT. On our map of K,O trends, calculated from the data of the Guiana garnetoids by means of echo models, we see that the KzO isoline content shows the distribution of deep troughs of geosynclines, which are dissected by a median ridge. All known manganese deposits are concentrated in the area of the median ridge or at the margin of the platform, i.e. where sedimentary conditions are often altered by the fluctuation of this unstable zone depending on syncline evolution. N o manganese deposits have been observed in the syncline zones.

J. VANN. DORR II. You attribute the free carbon in manganese silicate rocks to liberation of carbon from carbonates during formation of manganese silicate, a metamorphic product. On the other hand, where rhodochrosite is quite pure and practically no manganese silicate is formed, as in part of Amapá and Ghana deposits, free carbon is abundant; up to 20 per cent in Amapá and 7 per cent in Ghana.

D o you not believe that this free carbon could not be derived instead from organic debris, producing the reducing environment that caused rhodochrosite instead of oxide to be deposited?

B. CHOUBERT. It is difficult to answer this question, since the presence of organisms in old and 2,000 m.y. sediments has not been proved. Besides, I think that the high temperature and pressure of metamorphosed changes should have turned carbon into a carbonaceous substance, which, however, never reaches the graphite stage.

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Precambrian ferruginous-siliceous formations of Kazakhstan

I. P. Novokhatsky Institute of Geological Sciences, Academy of Sciences, Kazakh S.S.R., Aima-Ata (U.S.S.R.)

Iron-formations are widespread in the Precambrian metamorphosed sedimentary rocks of Kazakhstan. They occur in two main regions: the Ulutau region on the western margin of the Kazakh folded area and the Betpakdala region to the west of Lake Balkhash.

The oldest deposits are of the Archaean Bekturan series, which consists of amphibolite, mica schist, plagioclase gneisses and quartzite. It is up to 4,000 m thick. It consists of arenaceous to argillaceous rocks that have been altered by intense metamorphism, sodium metasomatism and granitization.

A thick sequence (up to 11,000 m) of metamorphic rocks forming the eastern limb of the Ulutau anticlinorium and the Karsakpay synclinorium is Lower Proterozoic in age. There are two important series in this sequence: a lower (Aralbay series) characterized by wide-spread development of liparite-dacitic and andesitic volcanitic rocks; and an upper (Karsakpay series) volcanic green schist series.

Iron-formation is developed mainly in the Karsakpay series composed of quartz-sericite and chlorite-quartz-sericite schist, marble, quartzite and iron-formation, as well as of porphyritoid and green schist formed from basalt and tuff. In the Aralbay series quartzite is less abundant.

In the upper layers of the Karsakpay series porphyritoid and conglomerate occurs sporadically. These series are characterized by an approximate rhythmicity, with porphyritoid and green schist predominating in the lower layers of a given suite, and phyllite and schistose rock with horizons of iron-formation predominating in the upper layers.

The following suites are distinguished in the Karsakpay series: the Burmashin (750 m thick), the Balbraun (800 m thick), the Shaghyrlin (1,500 m thick) and the Biit (1,100 m thick). Iron-formation is most common in the Balbraun suite, where the largest deposits of iron ores (the Balbraun and the Kereghetas) are located. These deposits make up the Karsakpay synclinorium where up to nine horizons of ferruginous quartzite are found.

The Middle Proterozoic in southern Alatau is rep-

resented by sedimentary and acid volcanitic rock. Two series are distinguished there: the Zhiydin (2,000 m thick) and the Maytyubin (9,000 m thick).

The deposits of the Bozdak series, which belong to the Upper Proterozoic, consist of basic and acid extensive rock with sandstone, conglomerate and limestone. The series is up to 3,000 m thick.

Most iron-formation deposits of the Ulutau region lie in the Balbraun suite of the Karsakpay synclinorium. In all, about twenty deposits of ferruginous quartzite are known in the region. They stretch to about 200 km in a narrow strip that extends in a nearly north-south direction. The southern continuation of the strata beneath unconsolidated deposits has been traced by geophysical observations, and its trend changes from NS to NW. In the north iron-formation is traced up to the bend of the Ishim river, where it also is overlapped by younger deposits.

The Balbraun field, the largest field situated to the south of Karsakpay settlement, is briefly described here.

Quartzites, which are sometimes conglomeratic, occur in the base of the ore-bearing strata with angular unconformity on older rock. The older rocks are altered to quartz- sericite, quartz-sericite-chlorite, and graphite-sericite schist, quartzite, and metamorphosed greenstone. Higher up in the section schist is gradually replaced by iron-formation interbanded with quartz-sericite schist. The strata are gathered into a number of isoclinal folds, with a western dip of the layers prevailing. In it a number of synclines stand out with cores bearing ferruginous quartzites (Fig. 1).

Iron-formation is part of the metamorphosed Lower Proterozoic rock that forms complex folds with a north- south strike. Ore outcrops here are about 5 km long and 3.5 to 4 k m wide. Seven ore strips, representing narrow synclinal structures, crop out on the surface. ïheir length along strike is up to 500 m, with an apparent thickness of 200 m or more. U p to 30 sheet-like deposits have been found in the field. Iron-formation and quartz-sericite schist alter- ate throughout the 60-100 m thick ore horizon.

The main ore mineral is hematite. It occurs in flakes, 0.01-0.1 mm in size, forming aggregates with quartz; it

Unesco, 1973. Genesis of Precambrian iron and nianganese deposiis. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 153

I. P. Novokhatsky

FIG. 1. Schematic geological section of the Karskpay synclinorium (Uzbekov, 1960): 1. Ferruginous quartzites with interbeds of quartz-sericite schists; 2. Quartz-sericite and quartz-chlorite schists with interbeds of tuffaceous schists and quartzites;

also occurs as separate strips from fractions of 1 mm to several millimetres thick. Laminar hematite, in the shape of euhedral crystals that sometimes cut the thin flaky aggregates, is less common.

Martite is often represented by porphyroblasts in fine- grained hematite and quartz between 0.01-1.5 mm in size. Magnetite occurs in the ores in small quantities as grains up to 1-1.5 mm in size that are often concentrated along the lamination. Pyrite, which is rare, occurs as single grains. Siderite is an occasional constituent of the ores. Quartz, the main noii-metallic mineral in the ores, constitutes up to about 40 per cent of the rock. Sericite, chlorite, calcite, apatite, epidote and gypsum also occur in the ore horizons.

The iron content of the ores varies from 20 to 63 per cent, with an average of 34-44 per cent. The average abun- dances of other elements are: silica, 29-34 per cent; phosphorus, 0.15 per cent; sulphur, 0.22 per cent. The following elements were determined by spectrographic analysis: Pb, 0.0007 per cent, Zn, 0.015 per cent, Cu, 0.001 per cent, Ba, 0.003 per cent, V, 0.0015 per cent, Ni, 0.0005 per cent and Ge, 0.00017 per cent.

Kereghetas is the southern continuation of the Bal- braun field. The areas have analogous geological structures, except that in Kereghetas greenstone is more abundant and quartzite is less abundant. The strike of rocks is north- south; the dip is steeply westward. Structurally the deposit represents a synclinal fold complicated by finer folding. The anticlinal areas are made up of greenstone, the synclines of schist with iron-formation. Sixteen ore bodies occur in five ore strips within the sequence and are buried by up to 200-250 m.

The ores of Kereghetas are mineralogically similar to the Balbraun ores, but they are leaner; the average iron content here is 37.2 per cent.

3. Green tufogene schists; 4. Quartzites with interbeds of green and quartz-sericite schists occurring with sharp unconformity; 5. Basic effusives and their tuffs.

The other deposits of the region are smaller, and they have not yet been studied in detail. All of them, except the Koldybayshoko deposit, occur in the same regional structure.

Aschitasty

The Aschitasty deposit, which is at the latitude of the town of Arkalyk, lies to the north of the previously described deposits within the same structure. It was discovered recently by geophysical prospecting, and is now being explored. The Precambrian rocks here are covered by younger unconsolidated deposits 40-100 m thick. The deposit lies at the southern part of the Tasoba syncline composed of metamorphic rocks including chlorite-sericite schist, porphyroid, amphibolite, quartzite and iron-formation. The strata are intruded by granitic rock and gabbro. The iron-formation horizon has been traced by drilling over an area of 13 km; it is up to 60 m thick.

In contrast to the other fields, the ores of the Aschi- tasty deposit have been metamorphosed by the intrusions. Contact metamorphism has resulted in recrystallization and coarsening of grain size of quartz, the formation of magnetite instead of hematite, and the formation of amphibole, epidote, apatite and carbonates. In spite of this, the ores have retained their thin banding.

The iron content of the ores varies from 20 to GO per cent, or from 38 to 44 per cent on average. The sulphur content is 0.01-0.11 per cent and phosphorus content is 0.4-1 per cent.

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Betpakdala region Betpakdala, to the west of Lake Balkhash, is another region in Kazakhstan in which iron-formation occurs in the Pre- cambrian sequence. The deposits are located in two strips of metamorphic rocks with a NW orientation. Older rocks, some of which are Archaean, are distinguished by the extent to which they have been metamorphosed; they consist of gneisses and crystalline schists. Overlying them are Proterozoic metamorphic rocks-quartz-mica and quartz- actinole schist containing marble layers, quartzite and iron- formation.

There are two important areas here: the northern region including the Zhuantyube field Gvardeiskoye, and the southern region, known as the Temir zone.

In the northern region, ferruginous quartzite crops out over a distance of about 15 km. Phanerozoic rocks, including Cambrian, Ordovician, Devonian, Carboniferous, Cre- taceous and Palaeogene, occur within this region. Meta- morphic rocks of the Precambrian are represented by quartz-chlorite and quartz-chlorite-sericite schist of green and grey colour, porphyritoid, white and yellowish grey massive quartzite and greyish-black iron-formation. The iron-formation crops out within a belt of metamorphic schist and porphyroid over an area about 10 km long and up to 4 km wide.

The metamorphic rocks here form a syncline elongated in the NW direction and complicated by smaller folds. The rocks dip steeply from about 65 to 82". There are twenty- one are bodies that crop out in zones 800-3,500 m long and 10-100 m wide. They are from 30 to 100 m apart. A magnetic survey has shown the presence of ore bodies up to 6,000 m long which do not crop out. Parallel arrangement of the ore bodies is due to their repetition by folding.

The ores are thin-banded ferruginous quartzite. Band- ing is caused by the alternation of laminae of hematite with laminae of quartz enriched by hematite. The thickness of separate layers varies from 0.1 to 2-3 m m .

Hematite, the main ore mineral, comprises 80-90 per cent of the total amount of the ore minerals. Martite and magnetite are less common.

The composition of the Zhuantyube ores is shown in Table 1.

The following variations in the content of the main elements in the samples were observed: iron, 23.18-59.24 per cent, average, 44.43 per cent; silica, 14.65-66.35 per cent; phosphorus (P,O,), 0.02-0.45 per cent; sulphur, 0.03- 0.28 per cent. The following trace elements were determined spectrographically: vanadium, O .O007 per cent; nickel,

TABLE 1. The composition of the Zhuantyube ores (%)

0.0007 per cent; zinc, 0.007 per cent; copper, 0.001 per cent; barium, 0.0015 per cent; germanium, 0.0003 per cent.

The Zhuantyube (Gvardeiskoye) deposit has not yet been completely prospected so that its total strike length is not known. Geophysical exploration has revealed ore bodies which do not crop out.

The second important area, a southern strip of ferruginous quartzites of the Betpakdala region, is situated in the central part of the Chu uplift, 100 km south of the area described above. It is known as the Temir ore zone. Here the iron-formation is also related to strata of metamorphic rocks built up of quartz-mica, quartz-epidote-actinolite schist alternating with porphyritoid, marble, and quartzite and iron-formation. The strata are gathered into complex folds of NW strike with steep NE pitches (70-88"). The metamorphic rocks are intruded by granite in the direction of the folding.

Within the metamorphic rocks light-grey non-metallic quartzites occur; these contain lenticular bodies of hematite quartzite, sometimes in an echelon-like pattern. Their thickness varies from 5 to 50 m, while their length varies from 50 to 500 m. The total length of the ore outcrops in various areas is 5-7.5 km, and for the whole zone it is up to 40 km.

The thin banding of hematite quartzites is caused by the alternation of laminae enriched by hematite and mag- netite0.1-5 mm thick, with layers of non-metallic quartzite. Martite, hematite and magnetite are the main ore minerals. Martite, which forms isometric grains 0.2-0.5 mm in size, is most abundant.

The iron content of the iron-formation varies considerably; from 10-15 to 32-36 per cent increasing regularly from west to east. The sulphur content is 0.03-0.35 per cent and phosphorus comprises 0.08-0.11 per cent.

The geology of the Temir zone has not been studied in detail. It differs from other regions gf analogous Precam- brian formations in the nature of its ore mineralization. Its economic importance has not yet been dearly evaluated. The most interesting area, 'Magnitny', lies in the south- eastern part of the zone marked by an intensive anomaly up to 3,500 y. It is overlain by unconsolidated deposits.

The features of ferruginous-siliceous formations of the Precambrian in Kazakhstan

The iron-formation of the Precambrian in Kazakhstan have much in common with analogous formations in other

Fe Tio2 A1,03 FenoI Fe0 MnO Ca0 MgO K20 NazO PzOj SO, H,O (ignition loss) Total total SiO, ~ _ _ _ _ _

56.45 16.45 No data 0.85 80.28 0.36 n.f. n.f. 0.27 0.10 0.10 0.36 0.01 0.03 0.57 99.38 49.75 27.45 No data n.f. 69.29 1.65 n.f. n.f. 0.14 0.10 0.10 0.28 n.f. 0.03 0.62 99.66 59.24 14.65 N o data 0.25 81.85 1.22 n.f. n.f. 0.17 0.10 0.10 0.45 0.14 0.01 0.54 99.48

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I. P. Novokhatsky

regions of the world. They are related to a thick series of metamorphic rocks subjected to intensive NS and NW folding. The estimated absolute age of the formations is 2,600-1,900 may.

The ore components are mainly hematite with accessory magnetite. The amount of magnetite increases in contact zones where iron-formation has been intruded by granite. Silicate-magnetite ores are practically absent in iron-formations of Kazakhstan, and this distinguishes them from analogous formations of other regions.

The iron content in hematite quartzite of Kazakhstan s not high, being rarely more than 20-40 per cent.

The ferruginous quartzites are also characterized by a low content of manganese, aluminium, phosphorus, sulphur, calcium, magnesium and such admixtures as lead, zinc, copper, barium, germanium and others.

Greenstones formed from basic extrusives make up an important component of the Precambrian metamorphic rocks. They frequently compose the lower part of a suite, while in the upper part various schists containing iron- formation predominate. Cases of interbedding of quartzite with extrusives and their tuffs are rather rare; this indicates a gap between the deposition of effusives and the main bulk of iron-formation.

Résumé

Les formations de fey siliceux dans le Précambrien du Kazakhstan (I. P. Novokhatsky)

Les formations de fer siliceux du Précambrien sont très abondantes dans le Kazakhstan. On les observe parmi les roches métamorphiques du Protérozoïque, qui forment deux larges ceintures, celles de Karsakpay et de Betpakdala.

La ceinture de Karsakpay est située dans la bordure occidentale de la région plissée de Kazakh et s’étend le long du méridien sur 400 kilomètres depuis le mont Arkalik, au nord, jusqu’à la rivière Beleuta, au sud. A l’intérieur de cette ceinture, on connaît environ une vingtaine de dépôts de minerai de fer siliceux, dont les plus grands sont ceux de Balbraun et de Keregetas. , Le minerai de fer siliceux est concentré dans une strate épaisse de minerais métamorphiques du Protérozoïque supérieur réunis dans des plis complexes qui forment un grand nombre de synclinoria et anticlinoria. La strate est constituée de porphyrites, de schistes verts (avec des veines de porphyrite et parfois de marbre), des quartzites, de minerai de fer siliceux, des schistes de chlorite (avec des veines de minerai de fer siliceux), de micro-quartzites graphitiques, de couches de porphyrite et de schistes d‘actinolite-chlorite et d‘actinolite-épidote-chlorite. Les plus abondants sont les fers siliceux de la série de Karsakpay qui constituent un grand nombre de synclinaux qui s’étendent sur des roches plus anciennes. La formation consiste en quartzites blan- ches et brunes, en quartz séricites gris-verdâtre, en schistes d’actinolite-chlorite et d’épidote-chlorite alternant avec du fer siliceux et des quartzites sans minerai. La puissance de cette formation est d’environ 850 mètres.

Dans le gisement de Balbraun, on trouve du minerai de fer siliceux sur une longueur de 5 kilomètres de long et 300 mè- tres de large. On observe jusqu’à 30 bancs différents de minerai à l’intérieur de cette bande.

L’horizon de minerai qui forme la partie supérieure de la strate est une série alternée de minerai de fer siliceux avec des schistes de séricite-quartz et des quartzites sans

minerai et plus rarement avec des minerais massifs. La puissance est d‘environ 60 mètres.

Le minerai de fer siliceux constitue de façon très dis- tincte des formations zonées d‘hématite, magnétite, quartz, chlorite et séricite et plus rarement d‘apatite, albite et carbonates.

Les réserves de minerai sont évaluées à 158 millions de tonnes. D’autres dépôts dans la ceinture de Karsakpay sont semblables à ceux de Balbraun mais sont de dimensions plus modestes.

Dans les années récentes dans le nord de la ceinture de Karsakpay, dans la région du mont d‘Arkalik on a décou- vert du minerai de fer siliceux par les méthodes géophy- siques. Le dépôt est maintenant connu sous le nom de (( dépôt Ashchitast ». L’horizon de minerai de fer siliceux est connu sur une distance de 13 kilomètres sous une couverture qui peut atteindre 100 mètres d‘épaisseur. La strate qui contient le minerai consiste en schistes de séricite- chlorite, porphyroïdes, minerai de fer siliceux et silex sans minerai. La strate est brisée par des intrusions de granite et de gabbro. L’horizon du minerai a environ 60 mètres d‘épaisseur.

Dans les minerais à part l’hématite, on rencontre aussi de la magnétite. La teneur en fer varie de 20 à 60 %, la moyenne étant de 38,4 %. Les minerais peuvent être concen- trés facilement. On estime les réserves de minerai à environ 500 millions de tonnes.

La ceinture de Betpakdala de minerai de fer siliceux est limitée à la strate de minerais métamorphiques de la mon- tagne de Tchuysk. Les roches métamorphiques sont présu- mées datées de l’époque protérozoïque. Elles s’étendent dans une direction nord-ouest. Elles sont caractérisées par une tectonique complexe et sont divisées en deux districts par une bande de dépôt du Paléozoïque moyen : la région nord (avec le dépôt de Zhuantijube) et la région sud. Dans la région nord, les dépôts paléozoïques sont représentés par des roches métamorphiques rassemblées dans des plis abrupts de direction nord-ouest. Dans le dépôt de Zhuan-

156


tijube, on observe sur une distance d’environ 10 kilomètres un grand nombre d’ameurements de quartzite ; ils représen- tent des couches (qui s’enfoncent rapidement) de minerais de fer siliceux rassemblées en plis complexes. Les minerais se présentent en minces couches de quartzites compo- sées de veines alternées d’hématite et de quartz-hématite. Une veine a de 0,l à 2-3 mm d‘épaisseur. L’hématite est le minerai principal, la magnétite et la martite sont moins abondantes (10 à 20 %).

La région sud (zone de Temirsk) diffère de la région nord à la fois par le caractère et par la composition du minerai. Ici, sur une distance de 40 kilomètres, 011 observe une strate de roches métamorphiques, qui s’étend dans une direction au nord-ouest, avec un plongement raide prédo- minant vers le nord-est. La strate est coupée par une grande

intrusion de granite. Elle est composée de quartzites, schistes, quartz-micacés, schistes porphyroïdes, amphibolites et marbre. Parmi les quartzites, on rencontre des lits ou des formations lenticulaires de minerais de fer siliceux. Dans une bande de 2 kilomètres de large, on trouve trois gisements de minerai de fer siliceux dont l’épaisseur varie de 30 à 50 mètres. Le contenu en fer des minerais de fer siliceux varie largement, depuis quelques centièmes jusqu’à 32 %.

En ce qui concerne la genèse des formations de silex ferrugineux dans le Kazakhstan, les opinions diffèrent : certains auteurs considèrent qu’elles sont des formations sédimen- taires typiques, d‘autres au contraire inclinent à leur assigner un caractère sédimentaire volcanogénique.

Discussion

J. H. GROSSI SAD. Is the iron quartzite (quartz-hematite ore) a chemical or a clastic rock?

I. P. NOVOKHATSKY. Quartz and hematite are mainly chemical sediments. In some deposits of Kazakhstan quartz has certain features of clastogene origin (Temir ore zone).

J. H. GROSSI SAD. What causes the ‘bedded’ aspect of the iron ores of Kazakhstan in your opinion?

I. P. NOVOKHATSKY. The bedding of the iron ores is the result of irregular sedimentation caused by different factors.

J. H. GROSSI SAD. What are the minerals containing germanium? Are they the iron minerals (hematite, magnetite, etc.)?

I. P. NOVOKHATSKY. Germanium is common in hematite and magnetite.

R. T. BRANDT. About what percentage of the Balbraun ferruginous quartzite has an iron content greater than 60 per cent? What are the characteristics of these high-grade zones?

I. P. NOVOKHATSKY. The percentage of ferruginous quartzite containing more than 60 per cent of iron is rather small. High-grade ores are bound in narrow synclines.

R. FRIETSCH. The Aschitasty ore contains 0.4-1.0 per cent P. What mineral is the phosphorus bound to?

I. P. NOVOKHATSKY. Phosphorus is bound to apatite.

R. FRIETSCH. What is the relationship between magnetite- martite-hematite?

I. P. NOVOKHATSKY. Magnetite is of later origin than hematite. Martite is a common mineral of the supergene zone.

S. ROY. At Kumdikul (central Kazakhstan) the iron ores are fairly rich in manganese (4-7 per cent and higher). In view of different Eh and p H limitations for the formation of iron and manganese as oxides, how do you explain this association?

I. P. NOVOKHATSKY. There are different quantitative correlations between iron and manganese, from pure iron ores to manganese ores. The Eh and p H had a great effect during their deposition.

S. ROY. H o w do you explain the facial transition of iron and manganese ores in Karazhal (central Kazakhstan)? D o you attribute it to changes of Eh and p H during deposition as suggested by Krauskopf (1956-57)?

I. P. NOVOKHATSKY. Ore depositional facies are different -from oxide facies to carbonate manganese. In this case I agree with Krauskopf.

S. ROY. What precisely is the grade of metamorphism of the Karazhal ores? M y experience with Indian ores shows that in oxide facies jacobsite and hausmannite appear in manganese ores in a fairly high grade of regional metamorphism.

I. P. NOVOKHATSKY. The grade of metamorphism is not high, it is connected with folding and thermal activity that led to the formation of jacobsite, hausmannite and some manganese silicates (spessartite, rhodonite, friedelite, pyros- malite and others).

157

Geology and genesis of the Devonian banded iron-formation in Altai, western Siberia and eastern Kazakhstan A. S. Kalugin Siberia Research Institute of Geology Ministry of Geology of the U S.S.R.

General geology Deposits of hematite, magnetite and hematite and magnetite ores occur in banded iron-formation in the folded, Devonian volcanogene-sedimentary strata of Altai in western Siberia and eastern Kazakhstan. The study of these ores provides new data concerning the source, environment, mechanism of ore formation, diagenesis, epigenesis and metamorphism of ores of this type in general. Due to their inaccessibility many iron-ore deposits of Altai are not mined; however, the reserves are reported to be 3 million tons.

The ore-bearing district is associated with the Altai Hercynian folds and the depositional environment is considered to have been volcanic island arcs or the margins of a marine basin in the Lower or Middle Devonian. Miner- alization took place in an area 600 km long and 150 km wide, but those iron deposits which have not been eroded are preserved in an area of not more than 20,000 km2. The ores are found mainly in the marine tuffaceous, pelite- aleurite-psammitic lithofacies, but disappear where the rudaceous, carbonaceous and terrigene-carbonaceous marine or terrestrial sediments appear. Outcrops of ore have been studied in detail in an area of approximately 50 km2.

A geologic section through the least-metamorphosed hematite ores follows (Figs. 1 and 2).

ROCKS UNDERLYING THE ORE DEPOSITS

The basement is composed of folded quartz and quartz- feldspathic sandstones, aleurolites, and conglomerates containing Lower Silurian marine fauna. Above the basement lies the only amygdaloidal diabase flow in the region. At Vodopadnyi (Fig. 1) a bed of ancient talus breccia rests on Silurian rocks. Below this breccia the rocks are fractured and have a shell-like cleavage caused by ancient subaerial weathering.

Above the talus breccia and diabase flow, lying on Silurian rocks, are vitrocrystaloclastic ignimbrites, quartz

porphyry and quartz keratophyre tuffs. The ignimbrites contain coarse, detrital grains or quartz, acid plagioclase, albite and K-spar. Biotite and tuffaceous rock fragments with an effusive habit also occur. Devitrified pumice fragments, apatite, zircon, amphibole, titanomagnetite, ilmenite and pyroxene are observed under the microscope. The vitroclastics and crystalline fragments are welded into a pseudofluidal mass with a micro-felsitic, more rarely micro- spherolitic, texture. Depending upon irregularities in deposition, as well as the extent of erosion, the thickness of ignimbrite and tuff varies from several tens to hundreds of metres.

The detrital ferruginous minerals, titano-magnetite, ilmenite, biotite, pyroboles and glass groundmass are irregularly replaced on a large scale by hydromica, quartz, calcite, magnesium chlorite, leucoxene and anatase.

Alteration of ignimbrite and tuff is compared with processes of postvolcanic, hydrothermal rock metamorphism in regions of recent, aerial, explosive volcanism, which is of the argillization type, and is accompanied by the subtraction of large amounts of iron, involving hydromicatization, local opalization and limonitization.

Iron-ore deposits

BASAL STRATA OF CLASTIC ROCKS

In the south-eastern Altai deposits, the subaerial ignimbrites and tuffs are overlapped by psammitic, gravel-bearing, rarely aleurite-pelitic laminated rocks. These rocks consist mainly of poorly rounded quartz grains, acid plagio- clases, potash feldspar, and titanomagnetite as well as zircon derived from underlying tuffs. Because these psam- niitic rocks also contain pyroclastics in the form of pumice forks and fragments, they are called tuffites, tuffogritstone, tuffosandstone, etc. The basal section is characterized by inclined bedding, graded in places, with rare, wave-like cross-bedding .

In the Srednekedrovsdy magnetite-hematite ores of

Unesco, 1973. Genesis of Precumbriun iron und ?nunganese deposits. Proc. Kiev Syinp., 1970. (Earth sciences, 9.) 159

A. S. Kalugin

central Altai the ore-bearing horizon is composed of conglomerates, coarse sandstone and sandstones which consist of fragments of quartzite, silicified tuffs and effusive rocks with hematite concentrated in the fine-grained layers.

D E S C R I P T I O N OF THE O R E - B E A R I N G B E D S

The basal rocks change rapidly upward into horizontal, rhythmically banded hematite ores and finer-grained clastics.

The thickness of the ore beds in the south-eastern Altai deposits attains 65 m, as compared with the normal 10-15 m. Jn many places the ore strata are eroded and overlapped by sandstones or gritstonec containing abundant fragments of ore laminae. Owing to the mountainous relief, the ore beds are exposed vertically for 800-1,000 m and they extend laterally for 15 km. There are five to seven beds with iron contents of 30-35 per cent; in the finer- grained facies the grade increases to 40-45 per cent.

The ore beds are commonly separated by thicker, coarser-grained tuffite layers, some of which are characterized by rhythmic, unidirectional, cross-bedding as well as graded bedding in the upper part of the section. In the psammite-aleurite lithofacies are ores that have been plastically deformed, resulting from landslide deformation, as well as ores that demonstrate a break in continuity. In these fractured ores, breccias with broken fragments as well as detrital pieces of ore layers are found in the tuffite cement.

Non-metallic laminae and beds in the ores consist mainly of reworked volcanoclastics derived from underlying tuffites and tuffs.

The ore laminae consist mainly of thin layers of compact hematite. Under the microscope hematite has micro- to cryptolepidoblastic texture with relicts of pelitomorphic and spherical bodies which range from 0.01 to 0.001 mm in size. Rarely, ferruginous carbonates make up ore laminae and constitute the core of hematite concretions found in carbonate-clay lithofacies.

In some beds of laminated hematite ore, particularly those containing ore-fragment breccias, the ore has been replaced by compact, wavy, pseudolaminated hematite, hydromica and quartzite, the latter minerals having been formed by the decomposition of silicate clastic material in the ore.

The recent formation of authigenic tourmaline, a result of the admixture of clayey matter and thin-bedded clastic material, is common in the ore laminae.

The chemical composition of hematite ores in south- eastern Altai is characterized by a high content of alkalis, particularly potassium, followed by alumina and titania. The presence of the latter two compounds resulted froin the reworking of tuffs, tuffites, pyroclastics of the same age and clayey matter.

The least-metamorphosed hematite ores of Altai have rhythmic horizontal laminae with alternating laminae of tuffite and hematite. These laminae are generally 0.5-1 .O cm, but in places are thicker.

In addition to dominant rhythmic laminae, the Altai ores with psammitic tuffite laminae contain a diversity of structures, including sun cracks, agitation ripples, gas bubble cavities, and halite crystal imprints, which indicate shallow water deposition, including littoral zones.

Agitation ripples, sun cracks, sedimentary breccias with large and small ore fragments, washout or erosional trenches and both cross and wave-like lamination in the aleurite-psammite lithofacies of the ore indicate that the ore was deposited in an environment of multiple sediment deposition that involved turbulent and oscillatory movements of the water. A more detailed study shows that the predominant, rhythmic alternation of ore and tuffite laminae displays all the features of graded bedding.

The rhythmic, graded structure of the layers, persistent or regularly changing thicknesses and the presence of diastems along the borders of the rhythmic layers suggests that the ores may be a peculiar type of ore flysch. This interpretation is strengthened by the occurrence of various hieroglyphs, traces of crawling and burrowing organisms, etc. The ore laminae surface, which is evidently consolidated to some extent, has agitation ripples indicating that the washout (erosion channel) is the result of oscillatory water movement.

Pelitomorphic or microspherulitic structures and indications of syneresis, which are evidences of high absorptive capability, suggest that the ore material was in a colloidal state dispersed in water.

The ore-bearing psammite-aleuritic lithofacies in south- eastern Altai changes to a pelite-aleuritic facies several kilometres away at the Albesin deposit.

The tuffite and hematite laminae coalesce in rhythmic layers. This occurs both in the shallow psammite- aleurites facies and in the pelite-aleurite lithofacies. The intrastratal channels are often not recognized because they are mainly developed in loosely consolidated sediments.

Our investigations show that washout channels in the ore sediments occurred during the formation of the rhythmic laminae and at the time of water fluctuation and turbulence. A single rhythmic layer, which consisted of clastics in the lower part and ore, in the form of mud, in the upper part was deposited. Therefore the rhythmic layers, many centimetres thick in places, did not come from a single source but are the result of a long period of sedimentation.

The thickness of rhythmic layers is in direct proportion to the amount of material supplied and to the length of time between periods of deep turbulence. Therefore, an increase in the thickness of the layers may result from an accelerated rate of supply of material or, at a constant rate of supply, less-frequent turbulence.

It may be concluded that the formation of rhythmic, gradational laminae in the iron sediments of Altai was controlled by the hydrodynamics of sedimentation. It is evident that rhythmic layers of similar types will always be equal to the amount of material that has accumulated between periods of intensive turbulence, although turbulence is capable of penetrating depths much greater than the thickness of a rhythmic layer.

160

Geology and genesis of the Devonian banded iron-formation in Altai, western Siberia and eastern Kazakhstan

O

O

A B

/::.::I .. . 5

lo - 11 Y 12

A 13

a 14

FIG. 1. Geological map of the hematite at Vodopadiiyi in the with large fragments of aerial plant fossil remains; 9. Ignimbrites, south-eastern Altai. 1. Talus deposits, moraines, snow and firn. tuffs, sparse quartz keratophyre and porphyry flows, red- 2-6. Sandstones, mottled aleurolites and conglomerates, tuffites, coloured aleurolites, tephrolites; 10. Aleurolites, quartz sand- algal limestone interbeds, breccias with hematite ore fragments; stones, conglomerates with quartz pebbles; laminae interbeds 7. Volcanomictic sandstones, gritstones and tuffites, in places with brachiopods, bryozoans, trilobites; 11. Fault; 12. Plant with fragments of hematite ores, dolomitic limestones with corals, fossil remains; 13. Fauna fossil remains; 14. Fauna fossil brachiopods and algae; 8. Rhythmically laminated hematite ores remains.

161

A. S. Kalugin

L+ + f + I

I X X X X X X X x X I

FIG. 2. Schematic geological section of the hematite deposits in south-eastern Altai with patterns of main structural types of ores and host rocks. 1. Mottled aleurolites, sandstones, conglom-

erates with fragments and pieces in the basal part of bed; 2. Sil- iceous effusive flows with spheroidal jointing and colloform hematite deposits in fractures; 3. Dolomitized limestones with

162


Rhythmic layered textures in shallow facies of ore beds have resulted from the movement of sediments by waves caused by wind. Evidence of shallow water also tends to support this. It is also possible that the wind effect on the water was a decisive factor concerning the rhythmic laminae in the more abyssal facies. In sedimentary basins in volcanic areas, seismic movements disturbing the bottom and the mass of water may cause turbulence. The following data point out the probability of seismic disturbances of the Altai ore sediments.

Hematite ores occur in the pelite-aleurolitic lithofacies of the Albesin deposit; the individual laminae are closely folded and the axial planes are vertical, thus excluding a landslide effect.

Plication is preserved here inside the diagenetic concretions while the same layers are horizontal. In some cases it is observed that material composed of aleuritic laminae penetrates upward through two or three ore laminae, which are broken into ribbons and folded. This deformation may be due to intensive disturbances.

Finally, breccia occurs with ore debris in sand cement, indicative of instantaneous turbulence and rapid precipitation without lateral movement.

UPPER-ORE DEPOSITS

As shown in Figure 1, overlying the ore strata are volca- nomictosite rocks, tuffites with lenses of dolomitic limestone, corals, brachiopods and algae of the Middle Devo- nian, Eifelian stage (the epoch characterized by extensive advances of the sea in Siberia). Globular quartz keratophyres are seen in the same sections. Deposition of ore ceases with the beginning of these open-sea conditions.

Gritstones, sandstones and breccia, with abundant suspended detritus consisting of underlying hematite ore, are found in many places above the ore strata.

In places, tuffites with intercalations of algal limestones are superimposed on the molasse suite of green and red calciferous sandstones, aleurolites and conglomerates.

EPIGENESIS PHENOMENON A N D ORE DEPOSITION

In the Altai area, €erruginous quartzites are extremely varied as to the type and intensity of metamorphism. In south-eastern Altai, the chlorite-hydromicaceous facies of abyssal epigenesis is observed, while in the central and western Altai zones of green slates, contact hornfels and granitization are observed.

In the process of metamorphism, argillaceous material and vitroclastic ash are altered to thin crystalls- and litho- clastics. Blastic quartz-carbonates, albite, sericite and chlorite aggregrates develop during the first stage of metamorphism. With increasing metamorphism light and dark mica, actinolite, epidote, sphene and other minerals are formed. Idioblasts of magnetite, which are associated with quartz and mica halos, are generated at the expense of hematite. During these stages of metamorphism, many of the original textures are obliterated, but the banding remains in the psephitic lithofacies with relicts of gradational bedding. During high-grade contact metamorphism and in zones of granitization, hematite is completely replaced by magnetite and relicts of sedimentary structures and textures disappear without obscuring the thick bedding and stratification of the ore bodies.

corals, brachiopods; along strike the limestones grade into argillites with hematite and ferrous carbonate concretions; 4. Vol- canomictic sandstones, aleurolites, gritstones, tuffites with fragments of laminated hematite ore; 5. Rhythmically laminated ores of the alenropelitic lithofacies, with submarine slumps, seismic fractures and crushed beds, local mud current deposits with aerial plant fossil remains and ripples; 6. Laminated hematite ore of aleuro-psammitic lithofacies, agitation ripples, sun cracks, gas blister cavities, rhythmic, sometimes cross and wave-like bedding, submarine slumps, breccias with ore fragments, plant fossil remains, possibly partly autochthonous; 7. Basal tuff sandstones, tuff gritstones, tuff aleurolites; diagenetic hematite concretions; 8. Quartz porphyry and keratophyre ignimbrites and tuffs, sometimes effusives; veins and lenses of colloform

dense hematite (black) with jasperoids hydromicatization (dotted) zones, barite veins (slanting hachures); prismatic jointing ignimbrites and rheoignimbrites with bomb zones; 9. Red- coloured aleurolites and sandstones; 10. Tephrolites with siliceous hematite, large Liesegang rings and red-coloured aleurolites and sandstones in the fractures; 11. Quartz porphyry and keratophyre ignimbrite and tuffs, with underlying rock fragments at the base; 12. Amygdaloidal diabase; 13. Conglomerate or fanglomerate; 14. Old talus rock breccia; 15. Quartz sandstones and conglomerates, carbonate interbeds with trilobites and brachiopods. (Structural stretches weremade from specimens and outcrops. Sketch sizes are enlarged compared with the scale of the geological section and reduced compared with their natural size 5-50 times.)

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Résumé

Géologie et genèse de la fosmution dévonienne du tes subatzé dans I’Altai‘, la Sibérie occidentale et le Kuzalc/lstun oriental (A. S. Kalugin)

Des dépôts de quartzites ferrugineuses se rencontrent dans les monts Altaï parmi des dépôts sédimentaires volcanogé- niques, qui datent de façon évidente du Dévonien inférieur et moyen. Les strates à minerai sont essentiellement compo- sées, dans leur partie inférieure, d’ignimbrites, de tufs de porphyres à quartz et d’albitophyre, de tufs et de roches effusives de porphyres trachitiques, de porphyrites tra- chyandésitiques, de roches kératophyres et diabasiques, Des manifestations d‘une ancienne activité solfatarique et de fumerolles avec évacuation de €er jusqu’à la surface d‘alors se rencontrent fréquemment dans ces roches. Des tufites avec participation de roches carbonatées et siliceuses- argileuses, enrichies çà et là de manganèse et de phosphore, prédominent dans la zone qui surmonte le minerai. Les types de plissement des strates à minerai changent depuis le pli ouvert jusqu’au pli linéaire, avec un accroissement

correspondant de la transformation régionale des roches et des minerais depuis un faciès d’épigénèse profonde jusqu’au gneiss et aux schistes verb avec un développement de la région des intrusions granitoïdes, du métamorphisme de contact et de la granitisation.

Les quartzites ferrugineuses des monts Altaï se rencontrent dans les sédiments marins tufogènes libres de carbonates et sont déposées dans un faciès qui va depuis les zones littorales jusqu’aux zones bathypélagiques. Elles contiennent des produits entraînés des formations de solfa- tares-fumerolles provenant de tufs sous-jacents et de cendres volcaniques abondantes fréquemment relavées.

La stratificatioii rythmique des minerais est déterminée par une alternance de vases du minerai original et vol<zo- clastiques résultant de redépositions multiples et du dépôt gradué de vases à minerai et de sédiments clastiques. I1 est suggéré que de tels types de minerai soient qualifiés de minerai de flysch. On peut les distinguer dans une classe par- ticulière de type sédimentaire-volcanogénique avec source superficielle et conditions sous-marines de dépôt de minerai.

Discussion

N. A, PLAKSENKO. In your report you have emphasized the role of clastic material in ferruginous quartzites. D o you find any changes in the granulometric composition of the rocks in the facial profile, i.e. in the direction from the primary source of sedimentation towards the deepest zones of sedimentation?

A. S. KALUGIN. In the Altai region ores like iron quartzites pertaining to shallow facies with sun cracks, agitation ripples, etc., contain abundant psammitic-silt material. On passing to deep-sea facies, characterized by the absence of sun cracks ripples, cross-beddings, etc., we find a predominance of pelitic material.

G. A. GROSS. You mention reproduction of some sedimentary features by laboratory experiments. Would you comment briefly on the technique of these experiments? Did you use silica gels?

A. S. KALUGIN. Using a technique similar to that of Moore and Maynard, we obtained colloidal sediments of ferric and silicon hydrates in artificially prepared sea-water. First, ferric hydrate precipitated out of this colloidal mixture and we observed a two-layer cherty ferruginous sediment. After settling, which lasted from several days to 1-2 weeks, the sediment was again subjected to agitation and settled in a water colum of about 25-30 cm. During the settling we observed sedimentation of undifferentiated cherty-ferruginous inaterial. Our experiments show that rhythmic alter-

nation of ferruginous and siliceous layers in ores of the iron quartzite type cannot be explained by accumulation of colloids alone, because the previously formed texture is broken down when the slightest disturbance is applied to colloids.

A. F. TRENDALL. What was the original nature and origin of the gas in the bubbles?

A. S. KALUGIN. I think that the gas bubbles in shallow facies are probably formed because of air sealed by the near-shore waves in the sand. This phenomenon has been described for recent sedimentations.

I. A. BERGMAN. Have you studied the distribution of accessory elements (P, As, V, M o , etc.) in the ores? What are the main features of their distribution: composition, concentration, etc.? H o w does metamorphism affect the concentration of these elements?

A. S. KALUGIN. The distributions and concentrations of accessory elements in the Altai ores were studied by me and also by S. I. Zubova, A. G. Guzman, V. G. Pono- niarev, E. G. Kassandrov, F. V. Sukhorukov andV. E. Po- pov. The hematite ores contain relatively high amounts of barium and boron. Lhe phosphorus and manganese contents do not exceed tenths and hundredths of 1 per cent. The titanium concentrations are due to clastogene material. Sometimes low concentrations of copper occur. Lead and zinc occur in concentrations of hundredths of 1 per cent;

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molybdenum in concentrations of thousandths of 1 per cent. Slightly metamorphosed ferruginous quartzites are lean in barium. The ores sometimes contain sulphides of non- ferrous metals due to epigenetic mineralization.

Yu. P. MELNIK. What is the percentage of free carbon in the hematite quartzites?

A. S. KALUGIN. Free carbon occurs in the banded hematite ores of Altai, but no data are as yet available on its content.

Yu. P. MELNIK. What are the factors governing the formation of magnetite from hematite? Is it only the rise in temperature, or also the presence of a reducing agent in the rocks?

A. S. KALUGIN. The formation of magnetite requires, besides the rise in temperature, the presence of a reducing

agent in the rocks. You have shown that in your calculations.

Yu. P. MELNIK. What are the metamorphic changes of silica (opal, quartz)?

A. S. KALUGIN. Under metamorphism the fragments of quartz and probably opal, as well as primary auihigeiiic cherty sediments, are transformed into grained, frequently granoblastic rocks.

Yu. P. MELNIK. Have you met with primary iron silicates?

A. S. KALUGIN. Kascandrov’s studies of the iron ores of the Kholzunsk deposit in Altai revealed spheroids of ferruginous chlorite which may be primary unmetamorphosed authigenic silicates.

165

Genesis of high-grade iron ores of Krivoyrog type

Y. N. Belevtsev Institute of Geochemistry and Physics of Metals, Academy of Sciences of the Ukrainian S.S.R. (U.S.S.R.)

Precambrian deposits of high-grade iron ores are confined to thin-banded iron-cherty rocks which bear different names on different continents: banded hematite quartzite, chert, jaspilite, itabirite, taconite, jasper-like beds, ferruginous quartzites, etc.

Iron-rich rocks consist of iron-ore minerals (magnetite, martite, hematite), quartz, carbonates, various silicates (chlorite, sericite, biotite, amphibole) and, to a lesser extent, feldspar, talc and muscovite. The rocks coinposed of the above minerals are classified into two groups: schist and iron chert, and jaspilite (itabirite, taconite, quartzites, etc.).

Two principal varieties of schist are distinguished in terms of their chemical and mineral composition: alumosilicate schists, consisting of sericite, quartz, muscovite, and biotite; and iron alumosilicate schist composed of magnetite, chlorite, biotite, cummingtonite and quartz.

Banded iron-rich cherts consist of two types of bands: metalliferous, made up magnetite, martite or hematite with a subordinate content of quartz, carbonate and iron silicates; and iron-free, composed of quartz and iron silicates, and carbonates. The thickness of the bands ranges from 0.1 to 10-15 m m . Iron-quartz, iron-quartz-silicate, iron- carbonate-quartz, or quartz-silicate cherts are distinguished according to the combination of variously composed intercalations in the rock.

Among ferriferous cherts, jaspilites are distinguished by their fine banding, formed by ore and quartz interbeds under almost absolute absence of other minerals (carbonates and silicates).

Beds of iron cherts are afew metres to 250-260 m thick, alternating with slaty beds form iron-ore suites. The thickness of iron-ore suites may range from 200-300 to 2,000 m. Sometimes effusive rocks represented by amphibolite, talc- carbonate, and talc-serpentine varieties are also found in the suites.

The majority of authors advocate a sedimentary-metamorphic origin of iron ore formations. Different opinions as to the formation of the cited rocks result from a different understanding of the role of sedimentation and volcanism in the formation of the primary composition of

the suite and the effects of its subsequent metamorphism. The Krivoyrog iron-ore basin occupies the central part

of the Ukrainian shield, stretching in the submeridional direction from the southern to northern margin of the shield along the boundary of Upper Archaean rocks in the east and Lower Proterozoic rocks abundant in the west. From the point of view of structural geology the Krivoyrog basin is a deep, relatively narrow synclinal zone which developed from the marginal trough of a Lower Proterozoic geosyncline.

The rocks that make up basin are termed ‘the Krivo- rozhsky series’, which is subdivided into three suites: lower -arkosic sandstone, conglomerate, and phyllite; middle- consisting of seven iron-cherty and jaspilite beds alternating with various slaty beds; upper-quartzite, sandstone, and various shales. All the above rocks underwent regional dynamo-thermal metamorphism to the greenschist and amphibolite stage.

The iron-ore deposits are generally located in jaspilite and iron cherts of the middle suite. They are represented by metalliferous beds, thick flexure deposits, ore lodes, and pockets. Ore deposits are confined to folded or fold-faulted structures, where they are concentrated in groups or patterns that form ore fields.

All the mineral deposits of the basin, in their turn, form three ore fields or ore regions characterized by common geologic-structural conditions, mineral composition and genetic features of ores; the Southern, Saksagan (central), and Northern ore fields or ore areas.

The Southern ore field is located in the southern part of the Krivoyrog basin and is characterized by a predominant development of beds and pods of specularite-magnetite and chlorite-magnetite ore deposits that are confined to the upper part of the middle iron ore suite or to the lower subsuite of the Krivoyrog upper suite.

The Saksagan ore field is situated in the central part of the Krivoyrog basin and is characterized by an abundance of compact and porous martite and friable goethite- hematite-martite and goethite-hematite ores, that form rather complexly shaped deposits with most frequent ore

Unesco, 1973. Genesis of Precumúriun iron and mangunese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 167

Y. N. Belevtsev

columns, stock-like and sheet-like deposits, morphologically connected with the thick articulate deposits (Fig. 1).

The northern ore field occupies the northern area of the basin and is characterized by amphibole-magnetite and hematite-magnetite compact ores confined to complexly- folded aiid faulted block structures.

In the world iron-ore deposits of the Precambrian, as well as in the Krivoyrog basin and Kursk Magnetic Anomaly (KMA), three types of high-grade ores are encountered. Compact magnetite or silicate-magnetite ores are located in areas of fold-faulted structures within rocks, considerably altered by magnesian-iron and iron metasomatism.

Compact magnetite-hematite, mastite-hematite and porous martite ores forming ore columns and sheet-like deposits are abundant in complexly-folded aseas of iron rocks. Magnetite ores replace martite porous and massive ores at shallow or greater depths within one and the same deposit.

Friable or soft hydrated ores are represented by goethite- hematite-martite and goethite-hematite varieties that are largely developed within the superficial crust of weathering or in narrow deep zones of oxidation.

The second type of ore constitutes the bulk of all Precam- brian iron-ore deposits encountered on all continents of our globe. On the other hand, the first and third type of ore are less common aiid of local significance, although there are individual deposits and areas where they are the main, if not the only, ore bodies present (Anshan, northern part of the Krivoyrog basin, etc.).

Magnetite and silicate-magnetite ores

These form the northern and part of the southern ore fields of the Krivoyrog basin, where they occur in amphibole- magnetitechert and shaleof themiddle suiteof the Krivoyrog series. The deposits are confined to areas of metasomatic rock transformation into a magnetite-amphibole variety.

The metasomatic nature of the mineralization has also controlled the morphology of ore bodies, governed by two

factors: (a) degree of tectonic preparedness of rocks (fold or fault deformation) furnishing the ways for downward moving ore-forming solutions; and (b) abundance of amphibole, magnetite-amphibole slate and jaspilite, which are lithologically favourable for rock metasomatism. The ore bodies occur conformably with the host metamorphic rocks.

Four principal varieties of ores are distinguished in terms o€ mineral composition: (1) amphibole-magnetite and amphibole-hematite-magnetite; (2) quartz-amphibole-magnetite (occasionally with hematite); (3) aegirine-amphibole- magnetite; (4) carbonate-magnetite-hematite ores. All these ores are relatively equigranular with a grain-size no more than tenths of 1 m m . The ores are compact, with an average porosity of 3-4 per cent.

Formation of metasomatic iron-ore deposits is associated in time and space with the concluding stages of folding, when the rocks were in the state of the greatest tectonic strain causing high permeability for pore solutions. The role of the ore-enclosing structures was played by zones of interstratal leaf-by-leaf gliding, cleavage, and zones of increased jointing and crumpling, generally originating on the slopes of large folded structures. In consequence, the ores develop in conformity with the rock bedding, enter into the composition of finely-folded structures and exibit no sharp boundaries with the host rocks.

The flow of pore solutions through the weakened zones resulted in the phenomena of dissolution and redeposition of the material, accompanied by accumulation of commercial abundance of iron in favourable structures.

The main ore-forming process resulting iii the emplacement of iron ores is magnesian-iron metasomatism; subsequent stages of metasomatism-alkaline, magnesian- calcium-carbonate, and silicification-caused complication of ore mineral composition and deterioration of its quality due to a dissolution and removal of iron during the stages of metasomatism.

Two stages are distinguished in the ore-forming metasomatic process, namely, magnesian-iron metasomatism proper and iron (iron ore) metasomatism.

The magnesian-iron metasomatism proper is the most active process, involving rocks of different lithologic-petro-

The October Revolution mine

The Dzerzhinsky The Kirov The Karl Liebknecht The Bolshevik The Frunce The XX Congress of CPSU The Red Guards mine mine mine mine mine mine mine

FIG. 1. Longitudinal projection of iron-ore deposits in horizons IV and V of the Saksagan syncline. (The lower portion is shown schematically.)

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graphic composition. It was manifested through the formation of magnesium and iron-bearing amphibole: cummingtonite, grunerite, and hornblende, which replaced biotite, chlorite, and garnet in various iron cherts and shales, as well as quartz and some iron-ore minerals. As a result, amphibole magnetite rocks originated. With a complete replacement of ore-forming minerals by amphibole in shales, amphibolitic cummingtonite (monomineralic) and magnetite-cummingtonite shales appear at the expense of iron-rich cherts. Maximum discharge of magnesium together with bivalent iron from the solutions is confined to tectonically weakened zones-periclinal closures of folds, flexures, structural inversions, etc. The amphibole or magnetite-amphibole shale and chert that are deposited in such areas have the shape of irregularly stretched bands. They are gradually replaced, both along and across strike, first by amphibole-bearing iron-rich chert or shale of quartz- biotite composition, and later, by amphibole-free primary equivalents of the above rocks.

Iron or iron-ore metasomatism immediately followed the magnesian-iron metasomatism proper. In the first stage of amphibole formation, the lack of oxygen prevented the formation of ferric oxides, with the result that iron accumulated in solution. An increase in redox-potential brought about removal of excess iron froin solution, first as magnetite and finally as hematite.

Since iron metasomatism manifested itself immediately after magnesian-iron metasomatism in areas of nearly monomineralic amphibole slate, the essential control of the same structural features, resulting from intense folding, complicated by minor forms of tectonic dislocation, is emphasized.

The deposition of magnetite and hematite of ore generation took place in conformity with the banding in cummingtonite shale, in magnetite-cummingtonite and magnetite-cummingtonite-chert . There is almost no evidence of ore emplacement in amphibole-free iron cherts.

Developing from amphibole or quartz-amphibole bands, the newly formed magnetite and hematite replaced amphibole and quartz. The content of magnetite in ores may vary and, consequently, within the ore deposits, areas of magnetite, hematite-magnetite and magnetite-hematite ores connected by mutual transitions are distinguished. The ores are granular, with textures following those of the original rocks. The most widespread ore texture is a distinctly banded one, which is characterized by a regular alternation of magnetite and amphibole bands or magnetite and amphibole and hematite bands.

Mineralization started with the deposition of magnetite. If, in iron chert and amphibolitic shale, the magnetite content (the product of general dynamo-thermal metamorphism; first generation magnetite) is 20-40 per cent, this may be increased to 60-80 per cent as a result of iron metasomatism, due to neocrystallization (second generation magnetite). In ores with massive texture it may reach 90- 95 per cent. The lamellae of the newly formed amphibole usually coat the grain aggregates of the magnetite or form intergrowth structures. Irregular and elongated grain shape, seldom with crystallographic configuration, O. 1-0.5 mm in diameter, is characteristic of the second generation magnetite.

Next to the second generation magnetite, the second generation hematite crystallized. It also developed in amphibole and quartz-amphibole bands, replacing amphibole and quartz (Fig. 2). The hematite bands in ore consist of

FIG. 2. Example of iron-ore metasomatism. Black, amphiboles; grey, magnetite; white, hematite.

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aggregates of closely intergrown lamellar or elongate grains forming looped or lattice-like networks.

The stage of magnesian-iron metasomatism terminated by deposition of the second generation hematite. Thus, the ultimate products of this stage were amphibole-magnetite, quartz-amphibole-magnetite and ampliibole-hematite-niag- netíte iron ores, with amphibole mainly represented by cummingtonite, to a lesser extent, by grünerite and, exceptionally, by hornblende.

The general scheme of iron-ore formatioil is pictured as follows.

The iron-rich cherty sediments underwent dynamo- thermal metamorphism which caused dehydration of rocks and formation of metamorphic solutions enriched by mobile components (Mg, Fe, Ca, CO,) due to decomposition of authigenous iron silicates. In the final period of folding, the metamorphic solutions migrated into tectonically weakened jointed permeable cavities, which were represented by vertical and steeply dipping structures responsible for the upward character of the solutions movement with a sub-

sequent release of ore-forming components. The lack of oxygen in the first stage of the process promoted development of magnesian-iron amphibole and, later, with a greater supply of oxygen, of successively magnetite and hematite. In this way the magnesian-iron (iron ore) metasomatism originated, resulting in the formation of amphibolic rocks and iron ore deposits in areas of complexly-folded deformations within the iron ore suite of the Krivoyrog series.

Martite and martite-hematite ores

The above-mentioned ores are extensively distributed throughout the Krivoyrog basin. They form numerous deposits of the Saksagan ore field as groups of ore columns located on synclinal limbs, or constitute Aexure deposits. The deposits of martite ores occur in iron-rich chert and jaspilite of horizons IV, V, and VI (Fig. 3) occupying the upper portion of the iron ore suite.

The ore content, or abundance, of deposits in different

D l m 2 4 m S m6 m7

FIG. 3. Distribution of iron-ore deposits: (a) in the Artyom mine, horizon 220 m; (b) in the XX Party Congress mine, horizon 270 m; (c) in the Lenin mine, horizon 267 m. 1. (KiC) Schist

horizon IV; 2. (Ka) Iron horizon IV; 3. (Kac) Schist horizon V; 4. (KS) Iron horizon V, 5. (Kgc) Schist horizon VI; 6. (Kg) Iron horizon VI; 7. Iron ore lode.

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horizons varies widely. The ore coefficient1 of all iron-rich rocks of the iron ore suite amounts to 0.04; however, for different horizons the coefficient varies from O to 0.8-0.9.

The greatest number of deposits is located in the Vth iron horizon composed of jaspilite. The content of iron there is about 37-42 per cent. The total area of ore deposits within this horizon is more than 70 per cent of the total metalliferous area of the Saksagan region. The ore coefficient for iron horizon V ranges, for different mines, from 0.12 to 0.78. The rate of mineralization of the horizon is discontinuous along the strike, thus a chain of ore beds is observed within deposits. However, no ore beds are contained in the jaspilite horizon between the deposits.

The ore coefficient of the VIth iron horizon is 0.024. Ore beds in this horizon constitute 26.4 per cent of the total mineralization area. The rock mineralization in this horizon is also uneven. The intensity of mineralization grows higher in the same places as in horizon V. Ore accumulations in horizon VI also alternate with ore-free or poorly mineralized areas. Such a relationship is observed throughout the Saksagan area and seems to result from the same causes as in horizon V. Consequently, the rate of mineralization of iron-rich chert horizons is clearly discontinuous. Thus, in the southern part of the Saksagaii area, within the synclinal closure (the Dzerzhinsky mine), an intensive mineralization is observed not only in cherty- iron rocks, but also in shale horizons, which resulted in the formation of a single thick flexure bed.

Further northwards in the Artyom mine, two length- wise chains of ore beds are found, confined to iron horizons V and VI and separated by a thick layer of ore-free chert. In all other mines chains of ore beds are also clearly defined and confined to the V, VI and partly the VI1 horizons of iron chert and jaspilite.

Along the strike of the chains alternate bunching and thinning of deposits may be observed, affecting all or nearly all horizons simultaneously. This phenomenon has resulted in lateral mineralization belts, traced in two, three, and sometimes four adjacent iron horizons, with ore-free areas separating the belts.

Along the entire Saksagan band the following six members of lateral mineralization belts, separated by large nonmineralization areas, are clearly distinguished in plan: the belt of the Dzerzhinsky and Artyom mines, the Karl Liebknecht mine, the Bolshevik, October Revolution, and Frunze mines, the XX Party Congress mine, the Krasnaya Gvardiya mine, and, finally, the Lenin mine. The areas separating the above-mentioned ore belts in some cases contain a number of ore lodes. However, the rate of mineralization within them is negligible.

The lateral belts of ore deposits are confined to areas of disturbances in plane-parallel occurrence of one or several chert and jaspilite beds in the form of gentle bends and flexures. The bends are in fact flat cross-folds up to 100 m wide, with a height up to 100 m. Between the mines, where such bends of rocks are not observable at depths of 1,500-2,000 ni reached by boreholes, no ore beds are found. Moreover, in the Dzerzhinsky and Artyom

mines, a dense network of boreholes and mining pits makes it possible to clearly delineate a thick flexure body, dipping northwards in perfect conformity with the Saksagan syncline bend. Consequently, two types of ore-bearing structures are distinguished: the bend of the Saksagan syncline and cross-folds on its limbs (Fig. 1). Thus, the ore deposits in the Saksagan area are closely connected spatially and morphologically with the cross-fold-jointed deformations and are essentially not present in non-folded rocks.

The main cause of the mineralization process involving a bed section or several beds was the ore and silica mobility, which rose under the recrystallization of rock material, and caused the concentration of ore into permeable zones of folding, fine-jointing and residual porosity.

The ore interbeds pass from rocks into ores without any significant changes in thickness. Often a certain increase in thickness of ore bands is observed when additional thin intercalations of newly formed hematite appear, generally at the boundary between ore and semi-ore interbeds. With an intense near-contact folding the ore interbeds become somewhat thinner and may even break, but the thickness of any ore interbed is the same as it is in the non-mineralized rocks. However, when one deals with ore-free and semi-ore interbeds, their thickness sharply decreases in the course of transformation of iron-rich chert into ore. The interbeds wedge out leaving just a thin streak composed of fine relict plates and laminae of hematite that had previously been dispersed in the semi-ore intercalation (Fig. 4).

The decrease in the thickness and the wedging out of semi-ore and ore-free intercalations is a general pheiiom- enon characteristic of various types of ores, including the nonoxidized magnetite ores of the Frunze mine. The thinning out zone is not very wide; in the vast majority of the contacts, it does not exceed 10-20 cm. Often an ore-free layer thins out abruptly disappearing in a distance of 1-2 cm, particularly at sharp tectonic contacts. Cases of gradual, gentle decrease in thickness of the intercalations for several dozens of centimetres and even metres are much less common. Owing to the wedging out of ore-free and semi-ore intercalations, the alternation of iron-rich rocks in ores is always accompanied by a considerable decrease in volume, which may be evaluated by studies on samples or direcily in mining pits.

A substantial number of measurements (more than 100) made it possible to establish that out of a layer of iron-rich rocks 100 c m thick, an ore layer 45-80 c m thick is formed. Consequently, 100 cm3 of rocks yield 45-80 cm3 of ore, the degree of compression or contraction ranging from 20 to 55 per cent. In none of the cases was the original volume of the rocks preserved in the process of mineralization. In some samples the contraction reaches 50 per cent, but the resulting ores still have up to 30 per cent porosity. In other samples the contraction does not exceed 20-25per cent, resulting in very massive ores with a porosity of 4-6 per cent.

1. The ore coefficient of rock is determined by the ratio of ore deposit area on a certain horizon, to the total area of rocks in which these ores are located.

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FIG. 4. Martite chert (1) and ore (2) contact. Dark colour, quartz intercalations; light colour, martite intercalations.

The numerous measurements cited above suggest that the amount of compression does not depend on the value of present porosity, since both massive and friable porous ores are affected. The total decrease in thickness as compared with the original rocks amounts to 15-20 per cent; this decrease, however, is compensated for at the expense of ore-folding .

In individual cases where a complete mineralization of the entire iron zone took place, a certain general thinning out of the zone results. The magnetite ores in the Saksagan area have been found at considerable depths as individual lodes and sections of limited dimensions. These are distributed immediately below the known martite ores or as separate thin beds in non-oxidized iron-rich rocks. Such ores have been encountered in the Kirov mine at a depth of 326 m, in the Lenin mine at a depth of 527 m, and in the Frunze mine at a depth of 280 m .

In most ore deposits of the Saksagan area the ores are fully oxidized and their magnetite content does not exceed 3-5 per cent. The oxidized martite ores are not homogeneous in their composition and physical properties, In numerous deposits compact massive martite ores are encountered, which pass into porous friable ores. In the

southern and central parts of the Saksagan area the massive martite ores are not very common, but in the northern past large deposits composed of homogeneous porous ores are quite sare. Usually a complex structure of ore bodies prevails, with massive ore distributed as inclusions in porous ores or forming a kind of a border at the contact of the ore-bed with non-mineralized iron-rich cherts. The dimensions of the massive ore inclusions range from several centimetres to several metres.

The contacts of massive and porous oses along the strike of the banding are usually rather gradual; sharp contacts are mainly observed along fractures. No changes in thickness of a massive ore layer as it passes into friable ores occur (Fig. 5).

The folding deformation in porous ores, manifested in the form of plication, is perfectly analogous to the deformation in massive ores so far as its character and extent are concerned. The complexly folded bends and fractures of the joints pass from porous friable ores into massive ones without any changes.

Throughout the friable ores remnants of massive ores occur, and niassive ores not infrequently pass into friable ones and then again into massive ores within the same

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FIG. 5. Contact of jaspilite with massive and porous ore.

deposit. The massive ores become friable without any perceptible decrease in thickness. While a distinct thinning is observed at the jaspilite/massive ores boundary, the transition from massive to friable ores is not accompanied by any change in thickness.

Porous ores differ from massive ones only by the value of porosity, the greatest porosity being characteristic of martite ore interlayers that were of a quartz-martite composition in the corresponding massive ores. The ore minerals of massive ores do not change at the transition into porous ores; the microstructure, size, configuration of grains and grain aggregates, as well as relative quantities of martite and laminar hematite, remain the same.

The study of some deposits has shown that friable martite ores are extensively developed near the surface. They are gradually replaced at depth by massive martite and martite-magnetite ores, and at greater depth they are invariably replaced by primary magnetite ores. The magnetite ores have a texture, conditions of occurrence, confine- ment to folded structures and mutual transitions within one and the same deposit exactly the same as those of massive martite ores.

The quantitative and qualitative relatioiiships between massive and porous ores, as well as between porous ores and iron-rich rocks, have shown that it is impossible to get high-grade martite ores through only supergene leaching of quartz out of normal jaspilite. A great number of the determinations of volume, specific weight, porosity and composition of the jaspilite and the resulting high- grade ores suggest that the supergene mineralization in the

conditions of iron immobility should have been inevitably accompanied by a great decrease in the jaspilite volume (up to 50 per cent of the original volume), with the corresponding ‘subsidences’ in the jaspilite series. Numerous measurements (1,200) in 150 sections, through the actual thickness of the jaspilite horizons, have shown that thickness of mineralized and ore-free jaspilite horizons do not differ significantly.

Calculations have shown that the present porous martite ores were formed at the expense of primary massive ores containing 50-57 per cent iron, and could not have been formed directly from jaspilite containing 37-40 per cent iron.

The emplacement of primary massive ores is associated with the process of dynamo-thermal metamorphism of iron-rich cherty rocks accompanied by transport of both iron and silica. The resulting primary ores were of a magnetite or hematite-magnetite composition, had low porosity, and contained about 53 per cent iron. In the process of oxidation the magnetite became oxidized to martite, while the removal of silica caused formation of highly porous (up to 25-26 per cent) martite ores with an iron content of 55-70 per cent.

Friable or soft hydrated ores

Within the Krivoyrog basin such ores are abundant in the Saksagan area, where they form either separate beds or compositionally complex ore deposits together with

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martite ores. In terms of mineral composition, goethite- hematite-niartite and goethite-hematite ores may be distinguished. The main ore minerals are goethite, fine-grained hematite, martite and, to a lesser extent, hydrogoethite. Beds of these ores occur within the zone of oxidation in silicate-iron cherts and iron-rich schist. The formation of hydrated ores is directly related to deep-seated zones of oxidation and the ancient crust of weathering. These ores were formed by disaggregation of iron silicates and carbonates, oxidation of iron, and formation of fine-grained hematite, goethite, and hydrogoethite, as well as by the magnetite to martite transformation.

Mining to SOO m depth and exploration boreholes to depths of 1,500-2,400 m enabled us to get an idea of the depth and character of the oxidation processes. The most conspicuous development of supergene processes is observed in the middle suite rocks-in iron-rich schist and chert. The supergene alteration of the middle suite rocks manifested itself non-uniformly. In some areas it reached depths of only several tens of metres, while in others it reached more than 2,000-2,400 m.

The surficial crust of weathering is widely developed and was located by numerous boreholes throughout the territory of the iron ore basin. The depth of the crust varies widely: in slate horizons from 15-25 to 80-100 m; in iron- rich zones from 20-35 to 150-200 m.

The vast body of evidence accumulated in recent years makes it possible to construct geologic maps and sections of deep-seated oxidation zones in the Saksagan band at depths of 100 and 500 m from the Precambrian surface. In ten most characteristic sections of the Saksagan band both the crust of weathering and deep-seated oxidation zones may be observed.

At present the linear zones of oxidation are traced with the help of deep structural drilling to depths of 1,300- 1,400 and even 2,400 m. However, the lower boundary of the oxidized rocks for most horizons has not been established. The depth of the supergene alteration of rocks seems to exceed 2,000-2,500 m . Manifestations of deep- seated oxidation within the Saksagan band of the Krivoyrog iron ore basin are confined to zones of cross-folding, the contact of the upper and middle suites and fault zones.

Linear zones of oxidized rocks are most abundant in iron horizons of the synclinal structure in the area. The deep-seated zones, including several stratigraphic horizons, strike in the meridional direction, following the principal trend of the structures in the area; at greater depths they plunge in conformity with folded and joint structures. The width of the oxidation zones ranges from several tens to 500-600 ni; in the south of the Saksagan area, in the Dzerzhinsky and Kirov mines, it reaches 1,000 to 1,200 m. The length of individual deep-seated zones ranges from 2- 4 km in the north to 10 kni in the south of the area. The zones are separated by areas of unaltered rocks 200 m wide in the central part of the area and 1,000 to 1,100 m wide in the north.

Within the Saksagan area five deep-seated zones of highly oxidized rocks are distinguished in the following

mines: the first zone, in the area of the Lenin mine; the second zone, in the Rosa Luxembourg mine area; the third one, the XX Party Congress mine; the fourth zone, the Komintern and Bolshevik mines; the fifth, the Karl Liebknecht, Dzerzhinsky and Illych mines.

The factual data suggest that the effect of the supergene processes lias generally manifested itself in a radical alteration of hypogene magnetite and silicate-magnetite ores and in the formation of the goethite-hematite ores and ‘shelestukhas’l due to decomposition of iron chert, jaspilile, and schist.

Percolation of surface waters downwards through the beds in the most jointed zones reached greater depths than in the host rocks and naturally caused a greater alteration in the hypogene ores. As a result of the oxidation of the magnetite ores, massive martite ores originated. In the process of oxidation of the magnetite ores, their texture, form, and dimensions of individual minerals did not change. The only change in the mineral composition due to oxidation of the magnetite ores is inartitization of magnetite.

During the next and the more vigorous stage of development of supergene processes, porous martite ores were formed by the removal of silica.

Comparison of the chemical composition of massive and porous ores has shown that the porous ores resulted from the complete removal of quartz from compact martite ores; this removal of quartz was not compensated for by any introduced material. This brought about a change in the porosity from 4-5 per cent in massive ores to 25- 30 per cent in porous ores. In the process of formation of porous ores, the content of quartz decreased from 15- 26 per cent in massive ores to 0.5-8 per cent in porous ores where oxidation of magnetite was more vigorous (confirmed by a decrease in bivalent iron to 0.6-0.7 per cent). The removal of silica caused a considerable increase in the iron content in ores; massive ores contain 75-80 per cent Fe,O,, while porous ores contain up to 97-98 per cent.

The silicate-magnetite ores in supergene conditions have undergone considerable changes, which were revealed as alterations of iron silicates and carbonates and the oxidation of magnetite. During this process a considerable removal of Ca, Mg, AI and Si took place. Together with decomposition of silicates, carbonates, and the removal of individual elements during supergene alteration of silicate-magnetite ores, fine-grained and dispersed hematite originated, which caused notable enrichment in iron of the goethite-hematite-martite ores. In the second case the processes of weathering resulted in the formation of a new supergene type of ore at the expense of disaggregation of jaspilite and schist (‘shelestukhas’). Consequently, the supergene processes played an essential role in the formation of porous martite and goethite-hematite-martite ores as products of alteration in the magnetite and silicate-

1. In the Krivoyrog basin ‘shelestulchas’ are leach jaspilites, from which a considerable part of SiO? is taken out, and residual quartz is transformed into a mealy substance. According to their content of iron (45- 52 per cent), they are related to iron-rich ores.

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magnetite ores, as well as in the formation of 'shelestukhas' and goethite-hematite ores.

A study of the relationships existing between various types of ores suggested a general scheme of their emplacement, with magnetite ores as primary (original) ores and all martite and silicate-martite ores as products of their subsequent alteration. Magnetite, silicate-magnetite, compact and friable martite, and soft hydrated ores are intimately interrelated, each type representing a certain stage in the development of the inineralization process.

All ore deposits, irrespective of their stratigraphic or structural position, mineral or chemical composition, are epigenetic with respect to iron-silicate and cherty-iron sediments of the Krivoyrog series. Thus, the ore deposits were formed essentially during the secondary enrichment of iron-rich rocks. An exception is furnished by interbeds and nests, non-considerable with regard to their volume and relatively rich in ore, which are found among jaspilites of sedimentary-metamorphic origin.

The following two main processes of ore formation may be distinguished within the Krivoyrog basin: hypogene, which is connected with the process of magnetite and silicate-magnetite ore bodies; and supergene, embracing the formation of goethite-hematite, partially goethite-hematite- martite ores and 'shelestukhas', as well as transformation of massive magnetite ores into porous and friable ones.

The subdivision of ores into two genetic types (hypogene and supergene) is quite arbitrary, since the formation of ores is a very complicated historical process consisting of a consecutive and multistage evolution-transformation of rocks and ores from the period of sedimentation up to present mineralization in the crust of weathering. A historical sequence is as follows: sedimentation, diagenesis, dynamo-thermal metamorphism, supergene alteration (associated with the ancient crust of weathering and, finally, Pre-Tertiary and present supergene alteration). In different deposits the above processes manifested themselves in a number of ways.

The distinctive features of composition, texture and modes of occuri-ence of the Krivoyrog deposits and other similar world deposits of Brazil, Canada and Australia are so original that they cannot be classed with any one of the known genetic types of iron ores (magmatic, hydrothermal).

Among the main features are these: mineral composition of the ores is similar to that of the host rocks; the ores contain the same chemical elements as their enclosing rocks; the ore deposits are distributed in folded-jointed structures; metasomatic processes of mineralization are predominant; no alteration is observed in the rocks adjacent to the ores; a strict combination in time and space of the mineralization process with folding; absence of zoning in the mineral paragenesis; absence of either spatial or time dependence of the deposits on intrusive rocks.

The cited features of ores and ore deposits make it possible to treat them as metamorphogene, originating in the process of dynamo-thermal metamorphism of iron-rich cherty rocks,

It is common knowledge that the main agents of metamorphism and, consequently, of mineralization are hydrostatic and unilateral pressure, temperature, and chemical activity of water solutions (Turner, 1962). These agents cause recrystallization of rocks, considerable removal of the material, and metamorphic differentiation accompanied by rearrangement and concentration of certain metal components.

In order to determine the principal aspects of the metamorphogene process of iron ore development, the following items will be briefly discussed: the nature of ore- forming solutions; causes and routes of the solutions; sources of iron and forms of its transport; causes of the material precipitation during mineralization.

Dehydration of primary sediments takes place during metamorphism. A sedimentary rock that has undergone the stage of diagenesis, still contains a considerable percentage of water present in two forms: free water in pores and coatings on the rock fragments and combined water within the rock material proper. According to Strakhov, the content of moisture in sand is 20-25 per cent; in siltstone, 30-60 per cent; in pelite ooze, 60-80 per cent. Much water is usually present in chemical and colloid sediments, as well as in the hydrates of silica, ferric oxide, iron silicate, etc. The most widespread minerals in sediments contain the following percentages of water: nontronite, beidellite, and montmorillonite, 13.6 per cent; kaolinite, 13.96 per cent; hybbenite, 34.65 per cent, ver- miculite, 19.96 per cent. This water was discharged from the rocks during lower, middle, and upper stages of metamorphism. Water, other than combined, has been isolated at temperatures up to 100" C. Hydroxyl water (combined) is isolated at temperatures higher than 300" C; from kaolinite it is discharged at 400-525" C; from montmorillonite, at 500-800" C; from hydromica, at 300- 600" C; from diaspore, at 5.50" C; from brucite, at 400- 500° C; and from chlorites at different temperatures up to 600-800" C. Most of this water is used for the formation and supplementation of metamorphic hydrothermal solutions.

In the metamorphic rocks of Krivoyrog the percentage of water is commonly 1-2 per cent and seldom reaches 3- 4 per cent. Enormous water masses must have been liberated in the process of dehydration and recrystallization of the rocks. A colossal amount of metamorphic rocks in the Krivoyrog during metamorphism must have supplied more water than could be contained in magmatic rocks of the same volume. If 30-35 per cent by volume of water was contained in the Krivoyrog sediments that had undergone diagenesis, it implies that every 100 m3 of metamorphic rock discharged about 50 ms of water. The total volume of water released during metamorphism is roughly estimated to have been half of the present volume of all the Krivoyrog basin rocks. This water promoted dissolution and recrystallization of the rocks under metamorphism and transported mobile components into areas of tectonic injection due to the inflow of metamorphic solutions. Thus the sources of metamorphic solutions were

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supplied by sedimentary and volcanic rocks that were undergoing metamorphism. They were produced at all stages of progressive metamorphism through the release of free and combined water from various minerals.

The locations of ore deposits are likely ways for niove- ment of the metamorphic solutions. In the conditions of the Krivoyrog basin, ore deposits are confined to folded- jointed structures, through which ore-bearing solutions percolated and where they left iron.

Numerous observations in mines, exposures and quarries, as well as studies on polished samples and sections, point to a close connection between ore mineralization and fine-jointing and porosity, and to ail absence of ore emplacements in large fractures and faults. Joints, clearly developed in the Krivoyrog rocks and ores are always post-ore with no ore mineralization within. The deposition of ores is confined to areas of optimum porosity and fine- jointing that resulted from the formation of folded structures. Consequently, folded-jointed zones favoured movement of ore solutions.

In the Saksagan area the ore solutions became most mobile in the zone of a complex folded-jointed core of the Saksagan syncline in places of shearing strain of rocks. From these main channels the solutions branched off through transverse folded-jointed zones produced by individual flexures or gentle bends. Under such conditions, the movement of metamorphic solutions is confined to folded zones, where the most vigorous movement of the material took place accompanied by lamination and later by fine-jointing. Such a favourable combination is readily matched by an iron-ore suite, characterized by an alternate sequence of iron-quartz cherts and fine-grained shales. The solutions squeezed first out of argillo-silty layers and then also from the shales. These, together with those released through the draining of the iron-rich cherty rocks proper, could not move within the iron-rich chert zones shielded by water-tight slates. Since the waters within each iron- rich chert zone were separated from other zones, they could not mingle, with the result that more or less isolated mineralization zones were produced. Thus the ore-forming solutions moved through folded-jointed zones in iron rocks from the places of tectonic compression to areas of folding where tectonic stresses were relieved and conditions of lower pressures set in.

The amount of iron transported for the formation of high-grade iron ores does not seem too impressive. Iron- rich chert and jaspilite that turned into ores had previously contained ail average of 35-37 per cent iron, while in the ores the figure became 55-60 per cent. Ore deposits occupy as much as 4 per cent of the area of all iron-rich rocks in such a productive areas as the Saksagan region. This implies that the supply of iron for the formation of high- grade ores amounted to no more than 2 per cent. Based on the actual geologic conditions, there are no data that might account for the addition of iron into the Krivoyrog iron-ore suite from magmatic intrusions of from other outside sources. It is postulated that the source of iron for the mineralization was an iron ore suite. This is cor-

roborated both by field and laboratory evidence: (a) transport of the material during mineralization took place within the iron ore suite and even within individual suitable horizons; (b) intimate spatial mineralogical and structural interrelations of iron ores and iron-rich chert and jaspilite.

Insoluble iron compounds are produced at high valency and are represented by oxides and hydroxides. Transformation of bivalent into trivalent iron takes place under the effect of free oxygen or interaction between bivalent iron minerals or oxygen-containing solutions. Essential causes of ore deposition include tectono-physical conditions (pressure regime, temperatures, and gaseous components of the solutions).

Genesis of high-grade iron ores of Krivoyrog type is considered as a natural-historical process of iron deposition, consisting of progressive primary sedimentation, metamorphism and supergene processes.

The earliest process of iron deposition implies sediineii- tation and diagenesis of the iron-rich cherty material, which laid down the foundation of all iron-rich rocks and a certain part of iron ores. The source material for the iron ore suite was derived from crystalline rocks of the Precambrian: gneisses, metabasites, ultrabasites, migmatites and granites. Sedimentation took place under the conditions of geosynclinal regime of the Krivoyrog-Kremenchug sub-geosyncline. Sedimentation is characterized by a single transgressive cycle covering the formation of effusive and coarse-grained rocks of the lower suite, chemogenic products of the middle suite, and clastogene material of the upper suite. Diagenesis resulted in crystallization of sediments as well as in its lithification. This period is characterized by the formation of hydrous minerals, such as chalcedony, opal, hydromica as well as chlorite, siderite, hematite, and pyrite.

The second stage of iron concentration in rocks, culminating in the formation of high-grades ores, is related to dynamo-thermal metamorphism that manifested itself through the formation of the Krivoyrog folded structure. Folding, plastic flow, and lamination of iron-cherty sediments caused heating up and circulation of metamorphic solutions, which resulted in the transport of iron, silica, magnesium, sodium, calcium, and aluminium, as well as in the recrystallization of rocks and the formation of various mineral associations. Such mineral associations as quartz-magnetite, quartz-magnetite-siderite, sericite-biotite-quartz, chlorite-biotite-siderite-quartz, chlorite-magnetite, cummingtonite-magnetite-quartz, aegirine-quartz, hematite-magnetite-quartz, and others are characteristic of rocks of the middle suite.

During the mentioned period the iron-rich cherty sediments were transformed into crystalline iron-rich chert and shale. In areas of development of folding (mainly cross-folding) and fine-jointing, where metamorphic solutions were in vigorous circulation, the transport of rock components, primarily iron and silica, resulted in the formation of high-grade magnetite and hematite-magnetite ores. The tectonic compression active in certain portions of iron complexes brought about instability of quartz, its

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dissolution and removal from the compressed zones. Ulti- mately, as a result of a relative enrichment in iron, high- grade ore bodies were formed. This period includes ore- forming magnesian-iron and iron metasomatism that seem to have developed by the end of the metamorphic transformations of rocks, reflecting a genetically regressive stage.

Two genetic types of ores pertaining to the metamorphic cycle are distinguished metamorphic iron ores making up the middle suite deposits; and metamorphosed high-grade iron ores occurring at the base of the upper suite.

Supergene alteration was the third process of ore formation and modification, and it was responsible for the formation of deep-seated zones of oxidation of iron-rich rocks and high-grade iron ores. Supergene processes brought about considerable transport of iron, silica, and other elements, which caused modification of compact magnetite ores into friable higher-grade and chemically

pure martite ores, formation of martite ores in jaspilite, goethite-hematite-martite ores in silicate-iron chert, goethite-hematite ores in iron-silicate shale, etc.

The data from geological observations are corrob- orated by experimental results and theoretical studies conducted by the Ukrainian Academy of Sciences.

Exploratory boreholes established that high-grade iron ores of the Saksagan area extend as far down as 2,000- 2,400 m. Despite a certain decrease in thickness and worsening in quality of ores with depth due to a gradual dying out of supergene processes, the ore bodies at a depth of about 2,000 m are of substantial thickness and are composed of high-grade ores. This serves as the basis for an optimistic evaluation of the mineralization extent down to an ultimate depth range of the iron-ore suite in the Saksagan syncline lower band. The geophysical data give this depth in the central part as 3.5 km; in the northern part, up to 5 km.

Résumé

Genèse des minemis de fer ii huute teneur de ICrivoyrog (Y. N. Belevtsev)

Les minerais de fer à haute teneur du type de Krivoyrog résultent d‘un processus chronologique naturel d’accumulation du fer, consistant en des processus successifs de sédimentation primaire, métamorphique et hypergène.

La sédimentation et la diagenèse des sédiments de fer siliceux qui sont à la base de toutes les roches ferrugineuses et d’une certaine partie des minerais de fer à haute teneur sont considérées comme les processus les plus anciens d‘accumulation du fer. Les roches cristallines du Précambrien fournirent les matériaux pour la série des minerais de fer : gneiss, métabasites, ultrabasites, migmatites et granites, L’accumulation des sédiments s’est produite à la faveur du régime géosynclinal du subgéosynclinal Krivoyrog-Kre- menchug. La sédimentation est caractérisée par un cycle unique, transgressif, recouvrant la série inférieure des roches effusives, les matériaux chemogéniques de la série moyenne et les roches plastogenes de la série supérieure. Le résultat de la diagenèse a été la formation minérale, le réarrangement et la recristallisation du matériel et sa consolidation, c’est-à-dire sa lithification. La formation de minéraux hydriques tels que la calcédoine, l’opale, l’hydromica aussi bien que la chlorite, la sidérite, l’hématite et la pyrite est caractéristique de cette période.

Le second stade de concentration de fer dans les roches, qui a atteint son plus haut point lors de la formation des minerais à haute teneur, est associé avec le métamorphisme dynamico-thermique, résultat de la formation de la structure plissée de Krivoyrog.

Les plissements, le flux plastique et la formation en bancs des sédiments des silex ferrugineux ont causé le réchauffement et la circulation de solutions métamor-

phiques, avec pour conséquence la migration du fer, de la silice, du magnésium, du sodium, du calcium et de l’aluminium, la recristallisation des roches et diverses parageneses minérales. Les roches de la série moyenne sont caractérisées par des parageneses telles que quartz- magnétite, quartz-magnétite-sidérite, séricite-biotite-quartz, chlorite-biotite-sidérite-quartz, chlorite-magnétite, cummingtonite-magnétite-quartz, aegirine-quartz, hématite- magnétite-quartz, etc.

Durant cette période, les sédiments de silex ferrugineux ont été transformés en cornéennes ferrugineuses cristallines et en schistes. Dans les zones de plissement actif (essentiellement transverses) et de jointements fins, OU les solutions niétamorphiques étaient en circulation intense, la migration des composants rocheux, surtout de fer et de Sioz, ont eu pour résultat la formation de minerais à haute teneur de magnétites et d’hématite-magnétite. La compression tectonique qui s’est exercée sur certaines portions des masses ferrugineuses se traduit par l’instabilité du quartz, causant sa dissolution et son évacuation des zones de compression; plus tard, dans ces zones, des gisements de minerais à haute teneur se sont formés à la suite de la concentration du fer.

On distingue deux types génétiques de minerai associés avec le cycle métamorphique : (a) les minerais de fer méta- morphiques qu’on rencontre dans les dépôts de la série moyenne; (b) les minerais de fer à haute teneur métamor- phosés qu’on rencontre dans la région la plus basse de la série supérieure.

Le troisième processus de formation du fer et d’alté- ration a consisté dans une hypergenèse qui est à l’origine des zones profondes d‘oxydation de roches ferrugineuses et de minerais de fer à haute teneur.

L‘hypergenèse a causé des migrations très significatives

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Y. N. Belevtsev

de fer, de Sioz et d‘autres éléments. Le résultat en a été la transformation de minerais de magnétite compacte en des minerais de martite difîus de plus haute teneur et chimiquement pure, La formation de minerais de martite dans les jaspilites, la formation de minerais de goethite- hématite-martite dans les schistes ferrugineux silicates, etc.

Tous ces processus d‘accumulation du fer dans les sédi- ments et dans les roches sont traités chronologiquement dans leur développement successif depuis la période de sédimentation jusqu’à l’état actuel.

Les résultats des observations et des études géologiques sont confirmés par des études expérimentales et théoriques de 1’Académie des sciences d’Ukraine.

Les puits d’exploration ont permis d’établir que les

minerais à haute teneur de la région de Saksagan s’étendent à une profondeur de 2 O00 à 2 400 mètres avec une épais- seur un peu plus faible et une qualité moindre du minerai au fur et à mesure qu’on s’enfonce, en raison de l’atté- nuation graduelle des processus hypergènes. Les minerais à une profondeur d‘environ 2 O00 mètres sont suffisam- ment épais et d’une haute qualité. Cela pourrait fort bien servir de base pour une évaluation optimiste d‘une extension possible de la minéralisation jusqu’aux plus grandes profondeurs de la série des minerais de fer dans les régions plus profondes de la charnière inférieure de Saksagan. Les données de la géophysique indiquent que la couche a une extension en profondeur de 3 3 kilomètres dans la partie centrale et de 5 kilomètres dans la partie nord.


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KOTLYAR, V. N. 1953. O genezise zheleznykh rud Krivogo Roga [On the genesis of iron ores of Krivoyrog]. Gorny zhurn., no. 12. MANN, V. J. 1953. The relation of oxidation to the origin of soft iron ores of Michigan. Ecori. Geol., vol. 4. MARTINENKO, L. I. 1950. K voprosu ob obrazovanii bogatykh rud Krivogo Roga [On the problem of the formation rich ores in Krivoyrog]. Uclzen. zap. Chernovitskogo in-ta, vol. 8, seriya geol. geogruph. nasrlc, vol. 2. __ . 1955. Rol’ gipergennukh protsessov v obrazovanii rud Saksaganskoi polosy Krivogo Roga [Role of supergene processes in the formation of ores in Saksagan zone of Krivoyrog]. Geologiya igenesis rud Krivorozhskogo zhelezorirdnogo basseina. Moscow, Academy of Sciences of the U.S.S.R. (Trudy soveshchaniya.)

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RAMDOHR, P. 1953. Uber Metamorphose und Secundare Mobi- lisierung. Geol. Rdsch., no. 42.

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also structures of deposits of Krivoyrog basin]. Geologiya i genesis rud Krivorozhskogo busseina. Moscow, Academy of Sciences of the U.S.S.R. (Trudy soveshchaniya.)

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Hiscussion

R. T. BRANDT. Where supergene alteration is found to extend to depths of 2,500 m , is this alteration the result of present-day ground-water activity, or is it related to ancient, Precambrian weathering cycles?

Y. N. BELEVTSEV. Supergene alteration in deep-seated oxidation zones took place after all the tectonic processes and is of Precambrian or, possibly, of Palaeozoic age.

R. T. BRANDT. Is the aerial crust of weathering, which has been located by numerous boreholes throughout the iron ore basin, a recent weathering crust or a Precambrian one?

Y. N. BELEVTSEV. I think it is Precambrian.

G. CHOUBERT. What are the specific weights of the compact and the porous ores?

Y. N. BELEVTSEV. The specific weight of the porous ore is about 4.0 to 4.5; the specific weight of the compact ore is somewhat lower-about 3.7 to 4.0.

S. J. SIMS. Does goethite ever occur in the supergene ores derived from iron-chert sediments?

Y. N. BELEVTSEV. Goethite nearly always occurs in ores formed from iron silicates (chlorite, biotite, etc.).

S. J. SIMS. Is there any relationship between the goethite content and the depth below the surface?

Y. N. BELEVTSEV. In oxidation zones the goethite-containing ores disappear at depths of 400-600 m.

P. M. GORYAINOV. Are rich aluminosilicate ores ever encountered outside iron quartzites? If so, how far outside?

Y. N. BELEVTSEV. High-grade aluminosilicate ores are generally formed from schists. They are most commonly encountered near their contacts with ferruginous horizons.

P. M. GORYAINOV. What is the occurrence of high-grade ores in the magnetite quartzites of Krivoyrog?

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Y. N. BELEVTSEV. Magnetite ores are widespread in the magnetite cherts of northern Krivoyrog.

P. M. GORYAINOV. As high-grade ores are mostly abundant in cross-folding zones (secondary folding), their formation must have taken place in the presence of rocks already metamorphosed (primary folding). Therefore, at the mo-

ment of ore formation the rocks contained little water. D o you not find a contradiction in the mineralization process model that you suggest?

Y. N. BELEVTSEV. The transverse structures are contemporaneous with the longitudinal, and belong to the same stage of folding and metamorphism.

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Effusive iron-silica formations and iron deposits of the Maly Khingan

E. W. Egorov and M. W. Timofeieva Far East Geological Service, Ministry of Geology of the Russian Soviet Federated Socialist Republic Khabarovsk, R.S .F.s.R.

The iron-silica formations of the Maly Khingan stretch for about 150 km, from a zone 40 km north to 10-12 km south of the river Bidjan, southward to the Amur river (Fig. 1).

The sedimentary complex known as the Khingan series has a total thickness of about 7.5 km, and is composed of terrigenous-carbonate rocks, highly metamorphosed in the lower part of the complex and weakly in the upper part. Iron-silica formations and iron ore of Lower Cambrian age are situated in the upper part of the series between dolomite of the Murandavskaya suite below and schist of the Londokovskaya suite above.

Rocks of the Khingan series are in isoclinal folds with steep limbs, complicated by transverse and longitudinal faults which break the whole sequence into isolated blocks. Beds of iron ore are preserved in downthrown blocks and in the cores of big synclinal structures. Iron ore crops out in narrow isolated discontinuous belts that trend southerly, concordant with the general strike of the rocks.

The Maly Khingan ore-bearing sequence is divided into three parts: underore, ore-bearing, and overore. The ore-bearing part consists of ferruginous and ferruginous- manganiferous quartzites. The ferruginous quartzites are banded magnetite, magnetite-hematite, and hematite. The ferruginous-manganiferous quartzites contain assemblages of silica-carbonate, braunite-hausmanite, braunite-hematite, and other iron and manganese minerals.

Rocks bordering the belts of iron ore consist of volcanic and sedimentary-volcanic schists formed by low- temperature regional metamorphism (greenschist facies). The majority of these schists are developed from tuffs and tuffites of basic and intermediate composition and may contain chert-hydromica, chert-chlorite, carbonaceous carbonate, carbonate-mica, and so on. Less important parts of the wall rock consist of expanded breccia of carbonate and carbonate-chlorite (schistose ksenotuffs of the Schalstein type). Rarely there are interlayers of sedimentary carbonate rocks in the volcanic and sedimentary-volcanic schists.

In the north-west part of the Maly Khingan mining district there are concentrated deposits consisting essen-

tially of magnetite. Here, in a single belt of outcrops 40 km long, divided only by Cretaceous effusives, are three iron deposits (Kostenginskoie, Sutarskoie, Kimkanskoie) with total ore reserves of about 1,000 million tons. Seventy per cent of these are of magnetite quartzite and 30 per cent are of magnetite-hematite quartzite. Industrial manganese ores are practically absent in the Maly Khingan district.

The Kostenginskoie deposit is distinguished from the other two in the district by relatively weak tectonic and metamorphic features. The ore field of this deposit is 14 km long and its width, after unfolding, is 3.5 km.

Iron deposits end by the splitting and pinching out of beds of ore, by a decrease in iron content, and by passing into green breccia containing rare interbeds of non-ferruginous quartzite.

In the Kostenginskoie deposit the iron-bearing quartzites are between two sequences of massive unsorted and non-bedded volcanic breccia, which rapidly pass into finer- grained sediments near the iron ore. Interlayers of coarse clastic rock occur at the base and top of the ore body, but are absent in the middle. The ore is distinguished by horizontal banding. Graded bedding is observed in interlayers of fine-grained clastic tuff. All this confirms that the iron ore zone was deposited during an interlude in volcanic activity, in calm water.

The uniform bedding and composition of the upper part of the Khingan series attest to the protracted existence of a vast basin in which carbonate deposition prevailed. At the beginning of deposition of the ore-bearing suite, no terrigenous material was being carried into the basin and a thickness of 500-1,000 m of pure chemical dolomite of the Murandavskaya suite had been deposited. Sudden volcanic activity retarded the normal process of deposition and produced volcanites of the ore-bearing suite during a short episode, after which deposition of carbonate began again in the basin.

Material of the ore-bearing rocks was transported to the basin of deposition by hydrothermal solutions and volcanic exhalations. W e observe channel-ways of solutions

Unesco, 1973. Genesis of Precanrbriaii iron and manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 181


FIG. 1. Map of geological features of the Maly Khingan. 1. Sedi- mentary-metamorphic complex of Proterozoic age; 2. Terri- genous-metamorphic sequence of Lower Cambrian(?) age, containing effusive iron-silica formations; 3. Volcanic rocks of Lower and Middle Cretaceous age; 4. Undivided effusive rocks of Cre-

and exhalations in the form of hematitic microquartzitic silicified zones in places in the underlying dolomite.

After volcanism and ferruginous sedimentation, the periodic deposition of tuffaceous material produced non- ore layers-breccia and schist-containing chlorite and disseminated magnetite.

In rock units consistiiig substaniially of hematite quartzite, interlayers of green brecciated tuff or schist are

taceous age and friable sediments of Palaeogene age; 5. Granitoid rocks of Palaeozoic age; 6. Faults; 7. Main iron and iron-manganese ore deposits: I. Kimkanskoie; II. Sutarskoie; 111. Kosten- ginskoie; IV. Kaylanskoie; V. Bidjanskoie; VI. Yujnokhin- ganskaya.

not uncommon. Here, oreless layers of grey or brownish- grey jasper-like siliceous schist formed at the expense of fine-grained clastic tuff having a silica content of 80-87 per cent. Banding of the hematite ore is broad, with thicknesses of 10-20 cm. With a decrease in the number and thickness of such noli-ferruginous interbeds iii hematitic quartzite, the iron content rises from 30-35 per cent to 35-40 per cent. The average content of iron in the ore is 25-30 per

182


cent. Towards the hanging wall of the ore zone, ore solutions were depleted in iron, but still brought in a significant amount of silica which continued to accumulate after the deposition of the greatest part of the iron, giving rise to the non-ferruginous quartzite. Silica also was added to tuffs, tuffites, or chemical carbonate sediments above the ore-bearing zone. Iron introduced by solution into reducing environments in the basin led to the formation of sulphides.

Siderite, magnetite, and hematite are present together through the entire cross-section of the ore zone without any sign that one mineral formed at the expense of the other. In magnetite quartzite, magnetite forms large crystals of idiomorphic habit which, where abundant in a bed, unite in polyhedral aggregates. Hematite forms dusty grains mainly in non-ore beds and in the interstices between magnetite grains. Siderite may fill the interstices between iron oxide minerals or may form granoblastic growths with other carbonate and quartz in non-ore interlayers. No oolitic, coiloinorphic, or concretionary structures are visible. These particular features indicate the independent for- mzttion of the different iron minerals. Differences in iron- mineral content are considered to reflect gradual changes in the sedimentation environment caused by the mixing of hydrothermal solutions and volcaiiic exhalations with sea- water. In the early stage of iron deposition, the alkaline environment in the basin played a significant role in the precipitation of magnetite and siderite. Continuous in- fluxes of acid solutions from hydrothermal solutions and volcanic exhalations changed the p H of the sea-water and under the new conditions hematite became more stable. Later, with a weakening of hydrothermal activity, the basin environment was dominantly alkaline, which again led to the deposition of magnetite.

Manganese and iron accumulated at somewhat different times. The highest content of manganese, 6-8 per cent, is registered in beds underlying beds of iron ore. Manganese-bearing sediments may be both schist and breccia. Manganese-rich beds may be at different levels, immediately under the ore or at some distance, stratigraphically, from the ore. Manganese-rich layers are composed of one or several lens-shaped silica-carbonate interlayers, 0.1-0.5 m thick, which are macroscopically indistinguishable from the wall-rocks. In the iron ore and in the upper layer of magnetite quartzite, manganese diminishes to 1- 2 per cent.

Among the manganese minerals are small quantities of rhodochrosite and manganocalcite and, very rarely, pyrolusite and psilomelane. Evidently the variations in manganese minerals depend directly upon the availability of COz. It is clear that most of the manganese enters complicated carbonates as parts of isomorphous mixtures. The manganese content increases by the enrichment of beds in pyroclastic material, especially of fine size. So, within the massive breccia beneath the ore zone there are sporadic lens-shaped interlayers, enriched in manganese to 1-2 per cent, where manganese is concentrated mainly in the finest- size fractions of pyroclastic material.

Thus, the abundance of the coarse pyroclastic material in the Kostenginskoie iron deposit indicates deposition of the iron near volcanoes. To the north on the strike of the ore-bearing suite, volcanic rocks become subordinate in the stratigraphic section at the Sutarskoie deposit, and give way to schist at the Kimkanskoie deposit.

Unfortunately, the change of iron ore facies in moving away from volcanic centres is difficult to trace across the Sutarskoie and Kimkanskoie deposits, which are located in contact zones of granite intrusions. The ore of these deposits is intensely metamorphosed and the primary proportions of hematite and magnetite are changed, with an increase in magnetite taking place.

The ‘South band’ district begins 25 Icm south-west of the Kostenginskoie deposit. The ore-bearing layers of the ‘South band‘ extend 40 km and constitute the Yujnokhin- ganskaya group of deposits.

In the west part of the ‘South band’, magnetite ore predominates. Carbonate breccia occurs widely in the wall-rocks and in ore bodies. A bed of silica-rhodochrosite (manganese) ore, 1-3 m thick, is present at the base of the iron ore zone, but hausmanite-braunite ore appears at the base of some iron-ore zones of great thickness, with gradual transition to the protoxide carbonate facies. By and large, the western part of the Yujnokhinganskaya group of deposits is very similar to the Kostenginskoie deposit.

The eastern part of the Yujnokhinganskaya group differs sharply from the western part both by the character of ore and the lithology of the wall-rocks. Here, the ore- bearing zone contains two clearly separated beds, one of iron-manganese at the base of the zone and one of iron ore, higher in the zone. The maximum thickness of the iron-manganese bed is 5-8 m and the minimum thickness of the iron ore bed is 18-20 m . The iron-manganese bed consists of banded hausmanite-rhodochrosite, hausmanite- braunite and braunite-hematite with an iron content of 8.6-11.0 per cent and a manganese content of 19.7-21 per cent. The iron-ore bed consists almost completely of hematite interbedded with cherty jaspilite and fine-grained clastic chloritized carbonate breccia. Wall-rocks are composed of diverse schists with an admixture of carbonaceous matter and give way to carbonaceous and calcareous dolomites. Lens-shaped interlayers of fine-grained clastic breccia are noted only in the underore zone.

Thus, from west to east in the Yujnokhinganskaya deposit a change takes place from coarse clastic littoral sediments to deeper-water deposits. In the same direction, protoxide facies (magnetite and rhodochrosite) change to oxide facies (hematite and hausmanite-braunite), inverse to the distribution of such facies in normal sedimentary deposits, but according to Strahov (1965), a peculiarity of volcanic-sedimentary deposits. Therefore, the regular increase in the thickness and grade of manganese ore in the east part of the Yujnokhinganskaya deposits is attributed to the greater mobility of manganese than iron at a distance from volcanic sources.

Such regularities are also observed in the ‘Eastern ore-bearing belt’ of the north part of the Maly Khingan

183


region, This belt, within which are seven deposits of iron and iron-manganese ore, is located 30 km east of the Kostengiiiskoie deposit and is 50 k m long.

In the north part of the ‘Eastern belt’, ore bodies consist of magnetite quartzite in which are rare members composed of magnetite-hematite quartzite. The thickness of ore bodies is 5-25 m. From north-east to south-west along the belt, the thickness of the ore zone increases and hematite quartzite appears. Still farther to the south, at the base of the ore bed, appear lens-shaped interlayers of siliceous schist of a peculiar lilac colour, in which the content of iron and manganese increases southward. In the southmost part of the ‘Eastern belt’, ore bodies of the Bidjanskoie deposit form persistent beds, 40-50 m thick, but manganese beds, on the contrary, are extremely irregular. In a distance of about 100-200 rn their thickness may change from 3 to 25 m and their manganese content diminishes upward from 27.2 to 12.0 per cent at the roof of the manganese-rich bed. The highest content of manga-

nese-28-31 per cent-is found in the underlying so-called ‘lilac’ schist. The ‘lilac’ schist contains fine-grained semi- transparent granoblastic quartz with dusty ore matter which is described by Mahinin (1952) as a mixture of hematite and hydrohematite. W e believe that these rocks formed from volcanic ash that had been enriched in lightweight particles of acid composition by aerial differentiation. The enrichment of the cinders, in manganese up to 31 per cent and in iron up to 40-50 per cent, supports the possibility of aerial transfer of significant amounts of ore matter.

It should be noted that the wall-rocks, especially those beneath the ore zone, consist almost entirely of uniform carbonaceous sericite-chert shale with an important content of angular fragments of quartz (up to 35 per cent) in cherty cement. Apparently, these rocks are also of volcanic origin. Thus, the Bidjanskoie deposit may be referred to as a volcanic-cherty formation, presumably owing a significant amount of its matter to aerial transfer.

Résumé

Les formations de fer siliceux efliisif et les gisements de fer. chi Maly Khingan (E. V. Egorov et M. W. Timofeieva)

Les formations de silex ferrugineux du Maly Khingan s’étendent sur une surface d‘environ 3 O00 k m 2 ; une bande de leurs affleurements s’étend dans la direction du méridien sur une distance d’environ 150 kilomètres. Au sud et au sud-est du fleuve Amour, les minerais de fer d‘Aii’shan’sky (Chine) et les silex ferrugineux de la région de Lesozavodsk (Primorye) sont considérés comme étant leur continuation.

Les formations de fer du Maly Khingan sont nette- ment stratifiées et peuvent toujours être distinguées dans la série de couches à minerais entre deux bancs épais de roches de carbonate terrigène.

L’âge des séries à minerais semble appartenir au Cambrien inférieur, bien que cela ne puisse être admis sans réserve par suite du manque de déterminations faunistiques.

Associée à la strate encaissante dans la structure complexe de plissements de rupture de la région et dans les stratifications au cœur des plissements synclinaux et des blocs affaissés, la série à minerais se manifeste à la surface sous la forme de bandes séparées qui s’étendent dans une direction méridienne.

La section de cette suite envisagée dans son ensemble est caractérisée par une structure stable à trois horizons : horizon inférieur ((( subore D) : schistes, tuf-schistiques et tufites ; horizon moyen (N ore 1)) : quartzite ferrugineuse et quartzite manganèse-fer avec des bancs de roches effusives ; horizon supérieur (K supra-ore N) : schistes, tufites schisteuses, et, dans la partie supérieure : calcaires.

Les roches qui contiennent ces séries, à l’extérieur des zones soumises à l’influence de l’intrusion, sont caracté- risées par un faible degré de métamorphisme jusqu’au

faciès des schistes verts. Elles sont représentées par des argiles schisteuses siliceuses-séritiques, calcaires et carbo- nifères formées de roches finement fractionnées volcano- terrigènes, volcano-sédimentaires et sédimentaires et des tufites schisteuses grossièrement ou finement fractionnées et des tufs à composition basique.

Les formations ferrugineuses dans l’horizon du minerai sont composées d’un banc unique de roches typique de la région entière : les silex ferrugino-rubanés abondent, alternent et s’imbriquent avec des tufites fractionnées irré- gulièrement et des tufs de composition basique. La puissance de l’horizon atteint de 30 à 60 mètres et elle est marquée par des fractures séparées, par des dislocations et plus rarement par des coincements dans de petites régions.

Les silex ferrugineux sont représentés par des variétés de magnétite clastique, magnétite-hématite et hématites jaspées généralement observées conjointement. Les minerais deviennent essentiellement magnétiques avec de l’amphibole, du grenat et des biotites, qui apparaissent dans leur composition.

L‘accroissement de la teneur en manganèse (en moyenne de 1 à 3 %) et l‘accroissement de la minéralisa- tion du manganèse lorsque l’on s’enfonce sont typiques de l’horizon du minerai tout entier jusqu’à l’apparition d’une strate de ferro-manganèse épaisse de 3 à 4 mètres, qui est composée de braunite, braunite-hausmanite, hématite- braunite, rodochrosite-hausmanite et d‘autres minerais avec un contenu en manganèse allant jusqu’à 30 ”/o. De fortes concentrations en manganèse à la base de l’horizon carac- térisent les parties sud et est de cette région, où les formations de minerai de fer sont étendues.

L’origine des formations de fer dans le Maly Khingan doit être considérée comme sédimentaire-volcanogénique

184


ce que confirment les faits suivants : interstratification de silex ferrugineux avec des roches volcanogéniques ; confi- nement exclusif de la minéralisation aux strates volcano- géniques-sédimentaires ; absence de sources possibles de déplacements des substances du minerai ; consistance dans la composition des formations de minerais de fer sur de vastes régions au cours de leur développement.

Les gisements de minerai de fer et de ferro-manganèse et les manifestations de minerai dans le Maly Khingan diffèrent en dimensions et en réserves. Les plus grandes se rencontrent dans la partie nord-ouest de la région : le

gisement de Kimkanskoye a des réserves d’environ 200 millions de tonnes, celui de Kosten’ginskoye a des réserves atteignant 500 millions de tonnes ; les minerais sont pauvres, siliceux et faiblement phosphoreux. La teneur moyenne en fer est de 30 à 40 %.

Le gisement de Yuzhno-Khingansky est le plus grand gisement de ferro-manganèse. I1 est constitué de diverses sections séparées. Les réserves totales de minerai de manga- nèse de ce groupe atteignent 6,7 millions de tonnes avec un contenu moyen de 20 à 21 % de manganèse et de 9 à 11 % de fer.


MAHININ, W. A. 1952. Geological features of iron-manganese ore deposits of the ‘East band’ in the Maly Khingan (unpublished).

STRAHOV, N. M . 1965. Geochemistry of the sedimentary manganese ore process.

185

Effusive jasper iron-formation and iron ores of the Uda area

E. L. Shkolnik Far East Geological Service, Ministry of Geology of the Russian Soviet Federated Socialist Republic, Khabarovsk (US .S.R.)

Preliminary prospecting of the Uda iron-ore area in the Far East of the U.S.S.R. (Khabarovsk Territory) has shown that the iron-formation and the ore are unusual in several respects. The Uda area covers about 12,000- 14,000 km2 and extends 500 km along the right bank of the Uda River towards the Shantarsk Islands (Fig. 1).

A Okhotsk sea

Iron ore farrn~tio"

I Iran are

Manganese ore

A Phosphorite racks

FIG. 1. Diagrammatic chart of area.

Over a hundred deposits as well as zones of mineralization and many magnetic anomalies have been formed in the area. Potential reserves of ore here exceed tens of billions of tons.

Because fossil remains are absent and exposures are poor, the geology of the area is not yet well-established; stratigraphic relations within the iron-formation and the structural features of the ore-bearing beds are still obscure. Alternating iron ore beds and terrigenous rocks of the formation occur in roughly parallel bands, striking north- eastward for a distance of 100 km.

Some geologists place the volcanically derived silica- rich iron bearing rocks partly in the Uligdan series of the

Lower Cambrian and partly in the Upper Precambrian. They also presume that the formation was repeatedly exposed by folding. Other scientists believe that the sequence consists of several suites of interbedded cherty rock of volcanic origin and terrigenous rock. Another interpretation is that the formation is partially or entirely of Upper (Middle) Palaeozoic age.

Intensively dislocated rocks of Palaeozoic age dip at angles of 60-70" and even 90". Ore bodies dip at the same angles. The general north-eastward trend of the formation is sometimes crossed by intricate folds. Cross folds in the ores are less common than in jasper-rich horizons. The formation is characterized by two sets of faults, one with a north-eastern (longitudinal) trend, the other with a north-western (transverse) trend. Blocks of Upper Palaeo- zoic granite and granodiorite intrude the north-eastern portion of the iron-formation, causing local metamorphism of the host rocks.

Iron and manganese ore and associated phosphorite rocks are located only in the volcanically derived jasper- rich part of the formation, which sometimes contains small amounts of sandstone, shale, limestone and dolomite. There is no ore in the terrigenous rocks.

The formation consists mainly of jasper and argillaceous jasper. Volcanic rocks are less common, but in some places they predominate, laterally grading into jasper and, rarely, into terrigenous rocks. The volcanic rocks are areally related to limestone and dolomite or their epigenetically altered differentiates (silicified rock). Cherty rock members sometimes form tabular or lens-shaped bodies from tens to several hundred metres thick. They commonly alternate with basic lava, pyroclastic rock, iron and manganese ore, carbonate and terrigenous rock. Most jasper is banded; massive homogeneous zones are rare.

Banding is due to alternation of jasper and argillaceous jasper containing 92-98 and 80-90 per cent SiOs, respectively. Occasionally banding is caused by the alteration of coloured jaspers. Iron minerals and subordinate amounts of hydromica and clay produce distinctive colour

Unesco, 1973. Geizesi.~ of Prccunzbriun iron wid marzgunese deposifs. Proc. Kiev Synzp., 1970. (Earth sciences, 9.) 187

E. L. Shkolnik

patterns depending upon whether they are evenly distributed or make up parallel bands in the rock.

Fossil remnants are not usually abundant. Locally chert which contains no iron minerals does contain abundant radiolaria and spicules of siliceous sponges.

Argillaceous jasper is darker and more fissile than common jasper, and has no conchoidal fracture. Thin chips of argillaceous jasper are not transparent. Light- coloured jasper contains negligible amounts of iron minerals; it consists of 97-98 per cent total silica and 80-90 per cent free silica. Subordinate amount of Alzo3 (1-3 per cent) and a high free silica content are characteristic of light-coloured jasper.

An alumina content of 3-4 per cent is found in rocks transitional from common jasper to cherty iron slates. The content of Fe203i-Fe0 varies from 3 to 4 per cent. In specimens grading from jasper to iron ore this figure ranges from 10 to 20 per cent with Alzo, increasing to 6 per cent. It should be noted that the alumina content in ores and in cherty iron slate is nearly the same. MnO,, Cao, M g O , PzO, concentrations vary between tenths and hundredths of one per cent.

Jasper beds commonly alternate with basic lavas and occasionally with pyroclastic rocks. Lava flows may be 10- 15 m thick and about 1-2 km wide. The calc-alkaline rocks of the flows consist of basalt and basalt porphyry. Rarely, spilite with an iron content of 10-20 per cent forms complicated packets 5-20 km long. Pyroclastic rocks are less common than lava flows. The most common are basaltic breccia with some lithocrystalloclastic agglomerate and psephitic tuff.

In many of the ore-bearing rocks, ore horizons are from 10 to 100 km long. In some sections, not exceeding hundreds of metres, the ore bodies form a chain-like pattern. Sometimes the sequence consists of alternating host rocks and echelon-like ore bodies. The number of ore bodies and their thickness vary considerably along the strike. Tabular ore bodies are rarely lenticular and have a complex structure. They are 200-300 m thick and may extend over 6 km, but are commonly 1-2 km long. Zones consisting of almost pure iron ore range from some centimetres to 50-60 m in thickness. The thickness of the ore bodies is directly proportional to their length.

Banded ores are predominant. Banding results from the alternation of layers of different composition and different component ratios. Bands and lenses are irregular in thickness, from microscopic dimensions to tens of centimetres. The average thickness ranges from several millimetres to 5 cm. Random variations from millimetres to tens of centimetres in the thickness of quartz-rich and iron oxide-rich layers results in an irregular banding pattern.

Massive ore texture is not common. In some specimens it occurs as a result of recrystallization of the ore. Brec- ciated ore structure is extremely rare; unmetamorphosed and slightly metamorphosed ore consists of fine and very fine grains. The size of hematite grains is about 8 microns. Magnetite consists mainly of grains of about 50 microns.

Two ore types are predominant: magnetite and hematite

with slight admixtures of magnetite. Other ore minerals present in the area are siderite, leptochlorite, magnetite- goethite-lepidocrocite. Hematite occurs almost everywhere except in zones of contact metamorphism. Brick-red hematite with colloid-like grains is the main oxide mineral. Under hydrothermal metamorphism it is transformed into dark- grey hematite (specularite). The amount of specularite in specimens varies greatly. Magnetite is commonly present in hematite. Its percentage differs in the various stratigraphic sections. There is no evidence of a gradational transition from hematite ore to magnetite ore by relative increase of magnetite.

Sulphides (pyrite, chalcopyrite) in all types of ores rarely exceed 1 per cent. Magnetite commonly contains sulphide blebs. Siderite contains 4-5 per cent manganese. Chlorites (thuringite-chamosite), which are particularly abundant in hematite ores (10-15 per cent), occur as flaky aggregates. Fine-grained quartz is abundant only in relatively poor ores. Silica is abundant in the ores as a constituent of silicates.

Red hematite is commonly not present in magnetite ores. Fine, more or less isometric, crystals of magnetite occur with tiny inclusions of sulphides and non-metallic components. During metamorphism and recrystallization magnetite crystals become ideomorphic and tend to increase in size.

The available data indicate that siderite ores and those with subordinate amounts of siderite are lenticular or elongated and about 10 c m thick. They may be widely distributed all over the area. Some specimens consist of pink-grey manganous siderite globules as large as 2-3 m m . Manganous siderite is rare and consists exclusively of pelitic particles (oligonite). These ores commonly contain sulphide, magnetite and iron carbonate.

Despite their varying mineral composition, the ore types possess certain common features. For example, the iron content in oxide and hydroxide ores is not high. They form two groups of ore with an iron content of 40- 46 per cent and 29-32 per cent. However, some specimens of ore contain from 50 to 60 per cent of iron. Carbonate ores sometimes occur with an iron content, which varies sympathetically with that of manganese, from 10-15 per cent to 25-35 per cent. All types of ores are characterized by a high manganese content ranging from 1-4 per cent to 10-15 per cent. Oxide and to some extent carbonate ores are rich in silica. The SiO, content is roughly inversely proportional to the iron content. The content of alumina, in silicates, is about 4-6 per cent. All ores are slightly carbonaceous. CaOfMgO is about 1-3 per cent. Both high and low iron groups are characterized by a high phosphorus content : from 0.6 to 1.2 per cent in the first group and from 0.25 to 0.3 per cent in the second. Thus the content of P is proportional to that of Fe. The ores are slightly enriched in Ni, V, Cr, Cu, Zn, Pb, As, Ge, Yb.

The ore has a great tendency to localize in certain types of host rocks. Statistical analysis, based on 150 exploration openings, shows that 80 per cent of the ore occurs in cherty rocks, 15 per cent in terrigenous rocks, and only

188

Effusive jasper iron-formation and iron ores of the Uda area

about 4 per cent in lava flows. Iron ores are unevenly distributed in the area; most of the large ore bodies are in the western part. In addition, the following regular features should be noted: (a) contact between ores and cherty rocks is generally gradational; (b) distribution of iron minerals in jasper; (c) jasper and ore are chemogenic sediments; (d) ore and carbonate rocks represents different facies.

A survey of world occurrences shows that the iron- formation and ores of the type described in this report are quite rare. Similar ore occurs in deposits of California,

New Zealand, Japan and in the Lower Palaeozoic of Kazakhstan. All deposits are small.

The number of different ore types, their geologic associations and the great area covered by the iron-formation may be cited as evidence that the Uda iron ore area is rather unusual. The texture of ores in the area is quite different from that of Precambrian jaspilite. Only in the Taikan basin are ores of the Uda area to soine extent similar to the Precambrian ones.

The genesis of iron ore formation in the area inay probably be explained by underwater fumarole activity.

Résumé

La formation de minerai de feu à jaspe effusif’et les minerais de feu de la région d’Oucla (E. I. Shkolnik)

Au cours des dernières années, on a découvert une formation de minerai de fer unique et peu métamorphosée du type eugéosynclinal et dont l’âge se situe entre le Précam- brien récent et le Cambrien. Cette formation, qui a fait l’objet d’une étude préliminaire, est située dans le nord de Khabarovsky Krai le long de la rive droite de la rivière Uda. La prédominance de roches siliceuses du type jaspe et d’effusions de séries de diabase spirite est très significative. Les roches pyroclastiques, terrigènes et carbonatées y jouent un rôle secondaire. Cette formation s’étend à l’intérieur des limites de la région nord-est de la ceinture plissée de Mongolo-Okhotsky dans une bande de quelques centaines de kilomètres de long et de quelques dizaines de kilomètres de large.

Les minerais de fer sont associés dans le gisement et probablement reliés paragénétiquement avec des minerais de manganèse et des phosphorites. D e plus, ils sont exclusi- vement accompagnés de jaspes et souvent ils se transforment en jaspe ; bien que, parfois, on les rencontre dans des effusions basiques et, très rarement, dans des grès aleurolites.

Les minerais de fer se présentent soit sous forme de nappe, soit sous forme de lentilles, qui atteignent quelques kilomètres de long et dont l’épaisseur varie de 50 à 70 mètres. Ils sont généralement rassemblés dans des bancs complexes qui s’étendent en des horizons en chaîne de quelques dizaines de kilomètres de long et de plusieurs centaines de mètres d’épaisseur.

Les minerais sont le plus souvent rubanés dans leur structure, mais ces structures résultent rarement de l’alté- ration des couches de minerai et de jaspe.

Elles sont plus fréquemment constituées par l’intercalation de minerais de richesse et de qualité différentes. Les structures sont surtout à grains fins ou à grains extrê- mement fins.

On rencontre ici des minerais de magnétite ou de magnétite-hématite peu métamorphosés, avec, fréquem- ment, de minces lits de sidérite et de magnétite-sidérite. La magnétite a une structure grainée et subidiomorphe. L‘hématite est représentée par une variété rouge dispersée de forme laminaire, le carbonate de fer est de la sphéro- sidérite avec des proportions variables de manganèse. La faible altération des minerais de magnétite et de magnétite- sidérite, l’absence d’indices et de pseudomorphoses dans le remplacement de la magnétite par l’hématite dénotent le caractère sédimentaire de la magnétite. Les formes miné- rales des combinaisons de fer dans les minerais et les jaspes sont les mêmes.

La teneur en fer des minerais est de 31-46 %. Ils sont en général manganeux, phosphoreux et fortement siliceux (15 à 30 % de silice). La proportion d’alumine est relativement stable et comprise entre 4 et 6 %. Les minerais sont corrélativement pauvres en substances basiques.

L’analyse du matériel recueilli montre que les minerais peuvent être considérés comme des dépôts chemogéniques de champs de fumerolles subaquatiques limités à des zones de paléofractures qui sont restées stables dans le temps et qui sont situées dans des régions de dépression.

189

Differing degrees of Différents degrés de metamorphism, the mineral métamorphisme, faciès des facies and the petrograpbic minéraux et nomenclature nomenclature of ferruginous pétrographique des roches rocks such as ferruginous ferrugineuses telles que quartzites, taconites, jaspilites, quartzites ferrugineuses, it abirit es taconites, jaspilites et itabirites

Mesabi, Gunflint and Cuyuna ranges, Minnesota

G. B. Morey Minnesota Geological Survey (United States of America)

Introduction Iron-formations of Middle Precambrian age in Minnesota are the source of a large part of the iron ore of the world. The iron ores are of two types, iiatural ore and magnetite- bearing taconite. The magnetite-bearing ores occur within the iron-formations as specific magnetite-rich layers which can be mined and concentrated by magnetic methods after fine grinding. The natural ores include iron-rich concentrates that formed locally by oxidation and leaching of the iron-formations and which can be shipped directly, and material that can be shipped after crushing and sizing, or can be beneficated using screening and washing or gravity methods to yield a high-iron product.

Three geographic areas or ranges containing Middle Precambrian iron-formations occur within or partly within the political boundaries o€ Minnesota. Two ranges, the Mesabi and Cuyuna, have been major producers, and the Mesabi range remains as the largest source of iron ore in the United States. The third, the Gunflint range, has never been an ore producer, but its geology has been extensively studied because of good exposures, slight deformation, and minor alteration. As such, it is a prototype of the other more highly altered and deformed ranges in Minnesota.

sequently were eroded extensively prior to deposition of Upper Precambrian strata. Rocks of the Upper Precambrian (the Keweenawan System) bridge the time span between the Penokean orogeny and the beginning of the Palaeozoic era.

There is a gap of almost a billion (lo0) years in the geologic record of northern Minnesota extending from the Late Precambrian to the Late Cretaceous. It seems probable that the area was eroded to its present bedrock topography and the natural iron ores were formed during this interval. Isolated patches of mostly non-marine Cretaceous strata locally overlie the iron-formation on the central and eastern Mesabi range, and marine Cretaceous strata are present as an extensive sheet south of the western Mesabi range and on the Cuyuna range (Sloan, 1964).

Northern Minnesota was glaciated during Pleistocene time and glacial material ranging in thickness from less than 20 ft (6 m) to more than 300 ft (92 m) mantles much of the area. Consequently, the iron-formations crop out only on the Gunflint range and on the eastern part of the Mesabi range. However, thousands of drill holes have precisely defined the distribution of iron-formation on the western Mesabi and Cuyuna ranges.

Stratigraphy GEOLOGIC SETTING

The Precambrian sequence in the Lake Superior region (Fig. 1) is divided into the Lower, Middle, and Upper Precambrian (Goldich et al., 1961) which have radiometric time ranges respectively of greater than 2,500 m.y., 2,500-1,700 m.y., and 1,700-600 m.y. These systems were interrupted by two major orogenies; the Lower Precambrian includes all rocks that were deformed, metamorphosed, and intruded by granitic rocks during the Algoman and older orogenies. These rocks form the ‘basement’ underlying strata of Middle Precambrian age. The Middle Precambrian rocks were deformed, metamorphosed, and intruded by igneous rocks during the Penokean orogeny and sub-

With the exception of an unnamed dolomite unit that underlies the iron-formation on part of the Cuyuna range, all the Middle Precambrian sedimentary rocks in Minnesota are assigned to the Animikie group. The stratigraphic position of the dolomite is uncertain, but it appears to be separated from underlying and overlying rocks by uncon- formities (Grout and Wolff, 1955). Possibly it is equivalent to other dolomitic formations that occupy a similar stratigraphic position in Wisconsin and Michigan (Fig. 1).

The correlation of the Animikie group with other Middle Precambrian rocks in the Lake Superior region has long been a problem. The term ‘Animikie’ was first used by Hunt (1873) for exposures around Thunder Bay

Unesco, 1973. Genesis of Precambrian iron arid manganese deposits. Proc. Kiev Synzp., 1970. (Earth sciences, 9.) 193

G. B. Morey

EXPLANATION OF SYMBOLS USED ON MAP

PHANEROZOIC COVER

UNCONFORMITY AL-

KEWEENAWAN ROCUS

UHCONFORMITY 2-

UNCONFORMITY , LLud ( B L A C ~ UON FORMATION

REPLACES SYMBOLIC DOTS WHEN AT OR NEAR UNCONFORMITYI

AFFECTED BY PENOKEAH LOWER PRECAMBRIAN ROCKS

EVENT

LITHOLOGIC SYMBOLS Volcanic Sedimentary Crystalline Rocks Rocks Rocks

E-/ 1-1

Lm Fq m j

Sandstone and Gabbro and granite of post-Middle

1-j Kewecnawan age

ond breccia Conpiornerote, and Granitic rocks of k-] port-Anmilie aqa

Basalt quartzite

G~aenstone tuff

(Irkm.

Greenston., in port with presarved pillow

strucbre~ Graywacle-shah Gneissic gronils af Iate prc-Animikie age pj

Amphibolite = Gronitic gnslrs of Iran-formation known ar probable

eorly pre-Animikie oqe EsZl Doiamiie

Schist praboily includes some rwk5 of volcanic origin

FIG. 1. Geologic m a p of the western Lake Superior area (modified from Trendall, 1968) and a stratigraphic summary of Precambrian rocks in the region (after James, 1960).

194

Mesabi, Gunflint and Cuyuna ranges, Minnesota (United States of America)

on Lake Superior. Irving (1 883) equated the ‘Animikie’ with Murray’s (1857) ‘Huronian’ and since that time many workers have accepted this correlation (Van Hise and Clements, 1901; Van Hise and Leith, 1911; Clements, 1903).

Because of lithologic and stratigraphic similarities, the Animikie Group in Minnesota can be correlated with part of the Marquette Range Supergroup in Michigan (Fig. 1). However Cannon and Gair (1970) have suggested that I . . . continued investigations have failed to show unequivo- cal correlation between Middle Precambrian rocks of Michigan and the Huronian Supergroup of Ontario . . .’. Therefore the term ‘Huronian’, especially in a time-stratigraphic sense, is no longer used in Minnesota.

The Animikie rocks in Minnesota are inferred to be younger than 2,000 m.y. (Hanson and Malotra, 1971); their actual age is still uncertain. Various attempts to determine the actual age have resulted in values ranging from, 1,685 & 24 m.y. (Faure and Kovach, 1969) to 1,900 & 200 m.y. (Hurley et al., 1962). As yet this problem is still unresolved.

MINERALOGY AND TEXTURES OF THE IRON-FORMATIONS

The Middle Precambrian iron-formations are ferruginous chert? containing from 25 to 30 per cent iron. Winchell (18933 originally referred to the lower part of the Biwabik iron-formation on the Mesabi range as ‘taconyte’ because of a supposed correlation of these rocks and Taconic rocks of the eastern United States. Later Spurr (1894) used this term ’. . . as a designation of the iron-bearing rock in general . . .’, and by the turn of the century, the spelling had become firmly established as ‘taconite’. Through common usage, the word taconite has since been extended to include all rocks of the Biwabik iron-formation except the oxidized ores (Gruner, 1946), and has been used as an informal rock name to describe the iron-rich rocks in other iron-formations of the Lake Superior region.

Most of the iron-formations in Minnesota, where unaffected by younger metamorphic events, are mineralogically complex rocks that are intermediate between James’ (1954) carbonate and silicate facies of iron-formation. The chief minerals are quartz, magnetite, hematite, iron carbonates, and iron silicates (Gruner, 1946; White, 1954). Most of the quartz is microcrystalline, although a few detrital particles as large as 0.2 mm in diameter are present. The iron silicates (greenalite, minnesotaite, stilpnomelane) may occur singly, but more commonly in combination with each other as matted or radiating aggregates. Generally they are difficult to differentiate because of intimate intermixing and fine-grain size and thus they are generally treated as a group. Magnetite forms tiny octahedra; commonly these may also occur in irregularly banded, regularly banded, laminated, patchy, or mottled concentrations. The carbonates form small to large rhombs or irregularly rounded grains.

In general, taconite composed dominantly of chert

with iron silicates or magnetite is apt to have a coarse- grained or granular texture, whereas rock that is mostly carbonates and/or iron silicates is apt to have a fine- grained or slaty texture. Thus, two fundamentally different kinds of iron-formation can be distinguished (Wolff, 1917; Gruner, 1946; White, 1954) : (a) cherty taconite, which is characteristically massive, quartz-rich, and has a granular texture, and (bj slaty taconite, which is generally fine grained, finely laminated, and composed mostly of iron silicates and carbonates.

DESCRIPTION O F INDIVIDUAL RANGES

Mesabi range

The name Mesabi range refers to the subcrop belt of the Biwabik iron-formation, most of which is buried beneath glacial deposits (Fig. 2). This subcrop belt, one-quarter to three miles (0.4-4.8 km) in width, extends for about 120 miles (192 km) in an east-northeast direction. The eastern end of the iron-formation is truncated by the Duluth Complex of Middle Keweenawan (Late Precambrian) age. The western end is covered by Cretaceous strata and by thick glacial deposits; the extent and trend of the formation west of Grand Rapids are not as well known as on the remainder of the range.

The Biwabik iron-formation, ranging in thickness from 100 to 750 ft (30-225 m), is underlain by quartzite and impure argillite of the Pokegama Formation and overlain by the Virginia formation, a thick succession of dark-grey argillite and intercalated greywacke. The base of the iron- formation is well defined by an abrupt change from iron- poor quartzite to iron-bearing and granular rocks. Through- out much of the range, the top of the formation is the top of a limestone-bearing unit containing little iron, but some iron silicates, and a few interbeds of granular chert (Fig. 2). The limestone-bearing unit pinches out in the western Mesabi and to the west, the top of the iron-formation is the top of a cherty siderite unit. The cherty siderite unit fingers-out to the west, and on the western-most Mesabi the top oí‘ the iron-formation is placed at the top of a graphitic argillite unit that is commonly an iron-bearing rock. On the western-most Mesabi, several iron-bearing members, one as much as 200 ft (60 m) thick, are interlayered with the argillites of the Virginia Formation. Thus, there is ample evidence of interfingering between the Biwabik iron-formation and the Virginia formation in this area.

Because recognizable lithologic units consisting of various proportions of rock strata having ‘cherty’ or ‘slaty’ textures occur over long distances, WOW (1917) was able to sub-divide the iron-formation into four units. From bottom to top they are: Lower Cherty, Lower Slaty, Upper Cherty, and Upper Slaty. These units, which have been retained subsequently as members (Gruner, 1946; White, 1954), can be traced along most of the main

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range and are recognizable even in the highly metamorphosed rocks near the Duluth complex. On the western- most Mesabi, however, the four-fold subdivision is not present. The Lower Slaty member, including an ash-fall unit called the Intermediate slate, pinches out near Grand Rapids; to the west, the Lower and Upper Cherty members are joined as a single member. O n the basis of texture and mineralogy, the lower part of the formation is referred to as a Cherty member and the upper part as a Slaty member.

Both the thickness and the iron content of the cherty rocks diminish westward. At the far western end of the range these rocks are only 20 ft (6 m) thick and contain almost no iron-bearing minerals. However, as elsewhere on the range, the basal conglomerate still contains algal structures. O n the other hand, the slaty rocks consist of a chert and abundant siderite and are intimately associated with argillites that generally have a higher iron content than the overlying Virginia formation.

Because of vertical and lateral lithologic changes within members, the subdivision of the iron-formation is at places arbitrary, but nevertheless useful from both a genetic and an economic standpoint. In general, lithic units included in the Slaty members contain 40 per cent or more slaty taconite (White, 1954), although locally within a given member the proportion of thin-bedded rocks may be less. The slaty strata characteristically contain sparse iron oxides, and the associated granular or cherty rock most commonly are silicate-rich. The cherty members on the other hand contain 10-30 per cent slaty material and are rich in magnetite, although they also contain abundant cherty or cherty-silicate taconite.

Gunflint range

The Gunflint range is more or less continuously exposed from west of Gunflint Lake on the international boundary to Thunder Bay 011 Lake Superior, a distance of approximately 110 miles (176 km) (Fig. 3). Isolated exposures on the north shore of Lake Superior indicate that these rocks once extended at least an additional 70 miles (112 km) to the east.

Rocks exposed on the Gunflint range are the north- eastern extension of those on the Mesabi range; the two ranges are separated over a distance of approximately 40 miles (64 km) by the Duluth Complex. As on the Mesabi range, three formations comprise the Animikie group. Locally the Gunflint iron-formation is underlain by a quartzite and overlain by the Rove formation, an interbedded argillite and greywacke sequence. The basal quartzite is thin and locally absent; where present, it is commonly included as a basal member in the iron-formation. The Gunflint iron-formation and the Rove formation have a gradational contact and, as on the Mesabi range, the top of a limestone-bearing member is considered to be the top of the iron-formation.

Goodwin (1956) divided the iron-formation into six sedimentary facies, each of which ' . . . is an areally restricted unit with unique lithic characteristics . . .' (Fig. 3). The

boundaries of these members do not coincide with the boundaries of the older four-fold classification scheme used on the Mesabi range, but fortunately the two schemes can be correlated with only slight difficulty.

The lower-most facies of the Lower Gunflint member consists of algal chert and lies on basement or on the conglomerate facies. It is overlain by tuffaceous shale (equivalent to the Intermediate slate) which in turn is succeeded by a thick granular taconite facies which grades north-eastward into banded chert and carbonate. This unit in turn grades northeastward into a granular taconite facies. The basal facies of the Upper Gunflint member is confined largely to the south-western part of the district, where it consists of algal chert. The algal chert is overlain by a second tuffaceous shale that forms the most persistent unit in the iron-formation, making it a marker horizon of time-stratigraphic significance. To the south-west, the shale is overlain by a thick taconite unit which grades north-eastward into banded chert and carbonate. The Limestone member forms the top of the Gunflint iron- formation, and is a thin but persistent unit separating iron- and silica-bearing rocks from the overlying Rove formation.

The lower and upper tuffaceous shale units are products of explosive volcanism, as are the several lava flows of basaltic composition that occur in the Gunflint. The flows and the tuffaceous rocks are considered indicative of the contemporaneity of volcanism and iron-formation deposition.

Cuyuna sange

The Cuyuna range is near the geographic centre of Minnesota (Fig. 4) in an area that is generally flat and mantled by a thick layer of Pleistocene glacial drift. Natu- ral exposures of iron-formation are lacking; consequently the geologic interpretation is based on artificial exposures in open-pit mines, on diamond drilling, and on various geophysical techniques.

The relationship of the Middle Precambrian rocks to the older rocks of the area is not adequately known. In a simple interpretation, the Middle Precambrian rocks occupy a complex north-east-trending synclinorium, bounded by older rocks, except at the north-east end where the structure widens. The Middle Precambrian strata probably wedge out against older rocks to the west and south, connect northward with the Mesabi range, and disappear eastward under a thick Cretaceous and Pleistocene cover.

The Troinmald, the main iron-formation of the Cuyuna range, is a distinct stratigraphic marker unit throughout the north range. It is underlain by the Mahnomen forrna- tion, an intercalated sequence of light-coloured argillite or slate and quartz-rich siltstone that is at least 2,000 ft (600 m) thick, and is overlain by at least 2,000 ft (600 m) of the Rabbit Lake formation, an argillite and slate unit that contains some carbonaceous and ferruginous material. The iron-formation ranges in thickness from 45 ft (13 m) to more than 500 ft (150 m) and is divided into two map- pable units; (a) an iron oxide-rich, thick-bedded facies,

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and (b) an iron silicate and iron carbonate-rich, thin- bedded facies. In about one-third of the north range, the thick-bedded facies overlies the thin-bedded facies and grades downward into it; elsewhere each facies comprises the entire iron-formation.

Intermixed chert and iron oxides in relatively massive beds characterize the thick-bedded facies. The cherty layers in this facies have a granular texture. Most granules and oolites are composed of various combinations of quartz, carbonates, fine-grained iron silicates, magnetite, and hematite. Locally abundant amounts of detrital quartz occur near the top of the thick-bedded facies, especially in chert beds having an oolitic texture.

The thin-bedded facies is characterized by individual bedding laminae that are less than half an inch (1.25 cm) thick; many laminae are less than one-eighthinch (0.3 cm) thick, and some laminae can be distinguished only with a microscope. The principal minerals in this facies are siderite, magnetite, s tilpnomelane , minneso taite , chlorite, and quartz. The silicates, carbonates, and magnetite comprise various proportions of the facies in all parts of the range and from bottom to top of the unit; thus there is no apparent mineralogic zoning. Argillaceous layers are intercalated with the thin-bedded facies at several places; some of this material may be reworked and altered ash falls,

INTERRANGE CORRELATIONS

There is little doubt that the rocks on the Mesabi and Gunflint ranges are correlative, as there are broad similarities in general stratigraphy, texture, and mineralogy (Broderick, 1920). The correlation of rocks on the Mesabi and Cuyana ranges is less certain in that they appear to have little stratigraphic similarity. However, exploration in an area approximately midway between the Mesabi and Cuyuna ranges (the Emily district) has shown a sequence of rocks intermediate between that found in either range.

Commonly the Emily district has been considered part of the Cuyuna range because of structural similarities (cf. Grout and Wolff, 1955); however, it has a stratigraphic succession much like that found at the western-most end of the Mesabi range. The lower-most formation contains quartzite like that of the Pokegama formation intercalated with light-coloured sericitic argillite similar to that in the Mahnomen formation. The iron-formation in the Emily district consists of a lower cherty member 255-300 ft (78.5-92 m) thick, a thin-bedded member 25-90 ft (7.5- 27 m) thick and a second cherty member of unknown thickness. In detail, the Lower Cherty member resembles stratigraphically the Lower Cherty member on the Mesabi range. A thick section of carbonaceous argillite, lithologically similar to both the Rabbit Lake and Virginia formation, overlies the main iron-formation. Thus, in general, the three-fold subdivision of the Animikie group extends from the Mesabi range to the Cuyuna range. Although some of the necessary detail is still lacking at this time, there is no evidence refuting a correlation between

the Mesabi and Cuyuna ranges and the iron-formation is projected along a sinuous path from the Mesabi range through the Emily district to the iron-formation of the Cuyuna range. The correlated segments of Middle Precambrian iron-formation thus extend from 70 miles (1 i2 km) east of Thunder Bay, Ontario to the central part of Minnesota, a distance of about 400 miles (640 km).

D EP O S I T I O N A L HISTORY

The Middle Precambrian rocks in Minnesota were deposited in part of an elongate basin, the configuration of which was probably controlled by a pre-existing grain in the older rocks (Van Hise and Leith, 1911). The western limit of the basin is unknown, but White (1954) suggested that the western shoreline was located somewhat to the west of the present Animikie exposures; the rocks extended to the east beyond Thunder Bay, Ontario (Goodwin, 1956) and to the south into Wisconsin and northern Michigan, where typical eugeosynclinal accumulations are now exposed (James, 1958).

In Animikie time a sea spread slowly from south to north across a broad, relatively flat plain. During the sea’s advance, shallow water clastic sediments were deposited (Fig. 5(a)). Relative subsidence was greater in the southern part o€ the basin, but sediment infilling was more or less able to keep pace with subsidence. Thus a thick wedge of fine-grained clastic detritus, fringed by a thin strand-line deposit of sandstone and conglomerate, was deposited prior to deposition of the iron-formation. The change from clastic to chemical sedimentation was abrupt (Fig. 5(b)) and there is no evidence that this change is not time- transgressive; it is possible that the iron-formation-clastic rock contact represents an onlap of considerable importance. This contact also indicates that stable equilibrium conditions were achieved between the source area and the depositional basin, and that during iron-formation deposition, little or no clastic detritus was supplied to the basin beyond that which is found near the inferred position of the strand-line.

The various textural and compositional aspects of the iron-formation result from deposition under differing environmental conditions, and a close relationship exists between the inferred physical and chemical environment, the composition of the precipitate, and its textural character. LaBerge (1967) suggested that the non-granular or slaty taconites are similar in many respects to siltstones or argillites, and that many of the granules in the cherty taconite were derived from texturally similar fine-grained material. In cherty taconites, granules commonly occur in strata having graded bedding or cross-bedding, and mixtures of granules with chert and carbonate pebbles, fragments of algal structures, oolites, and detrital quartz indicate that the granules behaved as particulate detritus (Mengel, 1965). Microscopic features indicate that many granules were reworked from previously deposited material; the cherty taconites are somewhat akin to clastic limestones. Thus,

200


Mesabi Emily Cuyuna

Greenschist Facies Amphibolite Facies

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Iron -Formation, .,Volcanic5

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(4 FIG. 5. Schematic diagrams illustrating the evolution of the Penokean geosyncline and associated orogeny in east-central Minnesota, (a), early Middle Precambrian time; (d), late Middle

Precambrian time; and (e), present configuration (modified from Goldich et al., 1961).

during iron-formation deposition, granule-bearing sediments were deposited in a shallow-water, agitated environment, whereas slaty or thin-bedded sediments were deposited in deeper, less active water. Fluctuations in the relative position of the strand-line or alternate basin deepening and infilling would result in a vertical sequence containing intercalated beds of granular and slaty rock types.

The iron-formations also reflect certain aspects of the chemical environment of deposition (James, 1954, 1966). In general, oxide-rich facies are to be expected in shallow wave- and current-swept areas, whereas iron sulphides are

deposited under quiet, deep, and stagnant conditions; the other facies are deposited under intermediate conditions, and more or less overlap these end members. Thus, the Trommald iron-formation was deposited in two distinct but gradational environments. Relatively deep and reducing water yielded a thin-bedded facies consisting of chert, carbonates, and silicates and as the basin filled, shallower, probably agitated, oxidizing water yielded a thick-bedded granular facies consisting of chert and iron oxides.

White (1954) outlined a similar lithofacies pattern in the Biwabik iron-formation and suggested that the intercalated cherty (oxide-silicate) and slaty (silicate-carbonate)

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G. B. Morey

facies resulted from deposition under transgressing and regressing conditions. Goodwin (1956) suggested that a similar facies pattern in the Gunflint iron-formation resulted from deposition at various water depths during periods of crustal instability, and that subsidence period- ically modified the basin configuration and, in turn, the facies distribution. Extensive and continuous instability ultimately modified the equilibrium state between basin and source area, resulting in a great influx of shaly material far into the basin (Fig. 5(c)). Volcanism has been suggested as the cause of the abrupt change from chemical to clastic sedimentation that characterized the entire basin in Minnesota. Although the basin in Minnesota was situated near a volcanic region, it was somewhat distant from the centre of igneous activity inasmuch as the proportion of recognizable volcanic material present is relatively small. Kowever, there is ample evidence of extensive volcanic activity elsewhere in the basin. James (1958) noted that approximately equivalent rocks in northern Michigan contain much volcanic material and that individual formations are lenticular, indicating deposition in a series of small and short-lived basins iii a volcanically active and tectonically unstable area. Thus is seems reasonable to postulate that tectonic instability in the eugeosynclinal part of the basin was reflected in Minnesota by fluctuations in the relative position of the strand-line and in systematic changes in the kinds of sediments that were deposited (Morey, 1969; Morey and Ojakangas, 1970).

During later stages of deposition, the basin became unstable (Fig. 5(d)), and the near axial part was folded and intruded by igneous rocks (Goldich et al., 1961). This deformation, called the Penokean orogeny, marked the end of Middle Precambrian deposition in Minnesota.

STRUCTURE

ïhe present structural configuration of Middle Precambrian rocks in Minnesota (Fig. 5) has resulted from the partial superposition of at least two tectonic events. The rocks first were folded and metamorphosed during the Penokean orogeny. As most of this deformation was restricted to the axial region of the Penokean geosyncline, the structural complexity increases in a south-westerly direction. The rocks of the Cuyuna range were markedly affected, the rocks of the Mesabi range were moderately affected, and the rocks of the Gunflint range were only slightly affected. Later, in Middle Keweenawan time, the structural configuration of these rocks was modified through the formation of the Lake Superior syncline, the axis of which approximates the present position of Lake Superior. As a consequence, the strata on the Mesabi and Gunflint ranges, being on the north limb of this structure, now dip gently south. The Cuyuna range, however, was little affected by this younger structural event.

The Gunflint range is a homocline that strikes east-north-east and dips 5-15"SE, except near Middle Keweenawan intrusive rocks, where dips may be as high

as 60". Gravity faults are widespread and dominate the structure. Most faults are steeply dipping and have vertical movements of as much as about 300 ft (90 m) (Tanton, 1931); however, most displacements are small, ranging from 20 to 100 ft (6-30 m). The majority of the faults have an easterly or east-north-easterly trend, but several have northerly or north-westerly trends.

The structure of the Mesabi range superficially resembles, but is more complex than, that of the Gunflint range. It is a gently dipping homocline that strikes east- north-east and dips 5"-15" SE. This general trand, however, is interrupted by several prominent structural features that have caused noticeable bends in the strike of the Animikie rocks or have produced pronounced changes in the outcrop width of the iron-formation. Among the more prominent structural features are (Fig. 2): (a) the 'Virginia Horn', a broad cross-fold composed of a gentle south-westerly plunging syncline and a parallel anticline; (b) the Siphon structure, which may be either a fault or a northward- trending monocline; (c) the Eiwabik fault, a fault about N 30" W , dips steeply south, and has a vertical displacement of about 200 ft (60 m); (d) several northerly trending cross- faults east of the 'Virginia Horn'; and (e) the Sugar Lake anticline, a broad southward-plunging structure west of Grand Rapids. In addition to these rather conspicuous structures, other less apparent but nevertheless important structural features have been delineated only recently (Marsden et al., 1969). For example, recent work in the Nashwauk-Keewatin area (Fig. 6) has indicated the presence of a more extensive structural pattern than was previously recognized. This structurally complex area consists of several cross-faults that, together with faults and monoclines in the intervening areas, fan out and dip toward each other, forming broad steps into a centrai graben.

Other minor structural features also are common, and include: (a) small folds and monoclines that locally produce steep dips in the iron-formation; (b) faults that commonly strike N75" W or N 20" W, and have displacements of less than 50 ft (15 m); and (c) very prominent joints having N 10"E, N40" W, or east-west directions (Gruner, 1924, 1946). These minor structures were important factors in localizing the extent and distribution of the natural ore bodies.

Not all the structures on the Mesabi range can be related to a particular tectonic event. Sims et al. (196%) have inferred that many faults that cut the iron-formation represent rejuvenated movements on older structures that developed initially in Early Precambrian time. Gruner (1946) and White (1954) concluded that both the present southward dip and the joint system are consistent with, but not necessarily related to, the development of the Lake Superior syncline. Many of the other structures have trends consistent with their having been developed during the Penokean orogeny. Thus, it may be concluded that the structure of the Mesabi range is complex and is related in time to both tectonic events that affected the area.

In contrast to the Mesabi range, the structure of the

202


E X P L A N A T I ON y Fault --

S C A L E

FIG. 6. Generalized geologic map of the Nashwauk-Keewatin area, Mesabi range, Minnesota showing the spatial relationship

Cuyuna range is the result of a single deformational event. The bedrock pattern is one of a large conspicuous syncline and several relatively inconspicuous but major anticlines. The axial planes of almost all the folds strike north-east and dip steeply south-east; in most places, the south-east limbs of the synclines are overturned. Drag-folds of all sizes are abundant, and have a normal and systematic relationship to the principal folds. Cross-folds or undulations of fold axes, or reversals of plunge directions, only are exposed locally although they may be widespread. In general, the axial trace of the cross-folds is about N10- 20" E.

Faults that have been definitely recognized and mapped (Schmidt, 1963) are limited in number and size. Although they may be important within a particular mine, they are insignificant to the district as a whok. Joints are abundant in most of the iron-formation, and served as major loci for the development of some of the Cuyuna

between various structural features and natural ore bodies (Owens et al., 1968, and modified from Marsden et al., 1969).

natural ores. The dominant joints strike about N 35" W and are vertical; hence they are approximately perpendicular to the fold axes.

METAMORPHISM

The Middle Precambrian sedimentary rocks have been affected by at least two metamorphic events, which can be readily distinguished. The first, a dynamo-thermal event, occurred over a wide area in late Middle Precambrian time, and is considered to be a consequence of the Penokean orogeny. The second, a thermal event related to the emplacement of gabbroic rocks in Late Precambrian time, is superposed on the older event within a narrow aureole adjacent to the gabbroic rocks.

The Gunflint iron-formation is inferred to be the least metamorphosed of the three iroii-formations. Greenalite,

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G. B. Morey

commonly believed to be a primary mineral, is the major Inasmuch as there is no stratigraphic control of the distri- silicate mineral; stilpnomelane occurs only where the iron- formation has been metamorphosed by Keweenawan intrusions, and minnesotaite is rare.

In contrast to the Gunflint iron-formation, minnesotaite and stilpnomelane, minerals considered by some to be indicative of the greenschist facies (James, 1955), are abundant constituents in parts of the Biwabik iron- formation. White (1954) has shown that the distribution of these silicate minerals is stratigraphically controlled on a dominantly regional scale and he concluded, therefore, that they were primary in origin. However, textural data (French, 1968) indicate that much of the stilpnomelane and minnesotaite is secondary. Most of the minnesotaite is restricted to cherty layers, where it appears to have replaced greenalite, according to the reaction; greenalite + quartz = minnesotaite + water, On the other hand, stilpnomelane is restricted to slaty beds which, on the average, contain a higher percentage of alumina than do the cherty beds (Gruner, 1946). Because many workers agree that stilpnomelane is one of the first minerals to form in an iron chlorite-rich sediment (Yoder, 1957) or in a sediment containing volcanically-derived material (LaBerge, 1966), stilpnomelane may have formed in beds originally containing this material. Thus, it can be postulated that selective transport concentrated clay-size material such as chamosite and/or altered volcanic detritus into particular stratigraphic horizons, and that the original bulk coinposition influenced the mineral phase that formed. This explanation can account for the apparent stratigraphic control of the silicates.

Unfortunately there is no textural evidence from the Biwabik iron-formation that can be used to indicate with certainty the physical conditions under which the mineral assemblages formed. Perry (personal communication, 1970), however, has demonstrated, using oxygen-isotope ratios in quartz-magnetite pairs, that when the magnetite equi- librated, temperatures on most of the main Mesabi range never exceeded around 125"-150° C. If these temperatures are indicative of the conditions under which the silicates formed, any distinction between diagenesis and low grade metamorphism is meaningless.

On the Cuyuna range, evidence for a Penokean metamorphic event is much more definitive. Schmidt (1963) has shown that minnesotaite and stilpnomelane are abundant constituents of the Trommald iron-formation; the iron- formation is intercalated with schistose rocks that contain chlorite and biotite. Moreover, grunerite, considered in-

bution of any of the silicates, these minerals must have formed via reactions involving primary iron-formation minerals such as quartz, carbonates, and greenalite.

A second metamorphic event, related to emplacement of Keweenawan gabbroic rocks, is superposed on the broad mineralogic pattern described above. The effects of this event are restricted to a narrow aureole in the east Mesabi district of the Mesabi range and to that part of the Gunflint range exposed in Minnesota. French (1968) demonstrated that metamorphic affects in the Biwabik iron-formation decrease from east to west away from the contact and was able to define four metamorphic zones. Zone 1 is unaltered taconite characterized by being fine-grained and composed of quartz, iron oxides, iron carbonates, and the iron silicates chamosite, greenalite, minnesotaite, and stilpiiomelane. Zone 2 is a transitional zone exhibiting no mineralogic changes, but having considerable secondary replacement of the original minerals by quartz and ankerite. Zone 3 comprises moderately metamorphosed taconite characterized by the development of grunerite and by the disappearance of layered silicates and carbonates (Griffin and Morey, 1969). Zone 4 is highly metamorphosed taconite characterized by increased hardness and grain size and by the presence of iron-bearing pyroxenes. A similar zonation occurs on the Gunflint range, and Morey et ul. (in preparation) have recognized three metamorphic zones that are roughly equivalent to French's zones 2, 3, and 4. The mineral assemblages near the Duluth Complex on the Gunflint range are similar to those from the easternmost Mesabi range (French's zone 4) described by Bonnichsen (1969). These studies have shown that the metamorphism was largely isochemical and characterized chiefly by progressive loss of H20 and CO,. There is no indication in either area that large quantities of components were introduced into the iron-formations from the Duluth Complex.

The moderate grade of metamorphism, characterized by grunerite and cummingtonite, is approximately equivalent to the garnet grade of regional metamorphism; pyroxene-bearing rocks are indicative of the sillimanite grade of regional metamorphism (James, 1955). Perry and Bonnichsen (1966), using oxygen isotope fractionation in quartz-magnetite pairs, estimated that the Biwabik iron- formation attained a temperature of between 700" and 750" C near the Duluth Complex contact. French (1968), using experimental data, concluded that the moderately metamorphosed taconite attained temperatures of 300- 400" C 2-3 miles (3-4.5 km) from the Duluth Complex

dicative of the garnet grade of regional metamorphism (James, 1955), is present in the iron-formation in the southeastern part of the north range where garnet occurs in intercalated schistose rocks. The close association of metamorphic minerals in the iron-formation and in the intercalated pelitic rocks led Schmidt (1963) to conclude that the Cuyuna range was metamorphosed during the Penokean orogeny (later dated by Peterman (1966) as Magnetite occurs throughout the unoxidized Middle Pre- having occurred 1,750 m.y. ago) and that the silicate cambrian iron-formations in amounts ranging from very minerals in the iron-formation are a product of that event. minor to abundant. It may occur as (a) disseminations of

contact.

Origin and distribution of the ores

MAGNETITE TACONITE ORES

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individual octahedra, (b) aggregates of individual octahedra, or (c) layered clusters formed by interconnecting aggregates of grains. Very fine-grained magnetite, which occurs in both granules and matrix as disseminated and diffuse crystals 5 microns or less in size, is probably primary in origin. However, definite secondary magnetite euhedra, 0.05 to 0.1 mm in size, commonly replace earlier iron silicates in granules and also surround and vein fine-grained siderite in the thin slaty bands associated with cherty taconite (LaBerge, 1964; French, 1968). Replacement of granules by magnetite is most common at their margins, yielding an inner core of greenalite or minnesotaite surrounded by a rim of coarser magnetite that in most situ- ations preserves the outline of the granule; more rarely an entire granule may be pseudomorphosed by magnetite. Thus, in many cases, there is little relationship between the primary sedimentary texture and the distribution of much of the magnetite.

Thick layers from which magnetite can be mined and concentrated using modern technology are found only on the Mesabi range. The ore bodies commonly occur within the cherty members; the slaty units are either too thin or too lean to be mined. The ore bodies are tabular stratigraphic units that have arbitrary boundaries defined by vertical and lateral changes in magnetite content and grain size. The lower cherty ore zone, or the middle part of the member, is of the greatest extent and is fairly uniform in magnetite content. In contrast, the Upper Cherty member is much less uniform in iron content, and the magnetite has a very erratic distribution; it is a less suitable, although workable, ore horizon.

The magnetite ores occur in two principal areas, the Main Mesabi district between Nashwauk and Mesaba, and the East Mesabi district east of Mesaba. The distinction of these two areas is more than geographical inasmuch as the East Mesabi ores were modified by metamorphic changes associated with the contact aureole of the Duluth Complex, West of Nashwauk there is a profound lateral change from magnetite- to carbonate-bearing strata; consequently, only relatively small ‘islands’ or tongues of unoxidized magnetite- bearing strata still exist.

Minable magnetite deposits in the East Mesabi district are confined to the Upper Cherty and lower part of the Upper Slaty members. Although the iron-formation has been modified to a completely recrystallized, sandy textured, granular rock near the Duluth Complex, the recoverable magnetite-concentrate from any particular horizon is about the same as that fromits non-metamorphic equivalent, even though the magnetite is somewhat coarser-grained. This may be due in part to the extensive development of an intimate magnetite-silicate fabric in themore metamorphosed rocks, making grinding and liberation more difficult.

NATURAL ORES

Natural ore bodies occur only on the Mesabi and Cuyuna ranges. They have been described by Leith (19031, Wolff

(1915), Gruner (1946), White (19541, Grout and Wolff (1955), and Schmidt (1963), and will be described only briefly here. The ore bodies occur in a wide variety of shapes and sizes within the Trommald and Biwabik iron- formations, where there is a marked correlation between their location and structural features that caused fracturing in the original taconite (Fig. 6). The deposits range in shape from fillings along narrow fissures, through channel- type deposits that occupy a system of fractures, to blanket- type deposits where extensive areas of ore formed in favourable stratigraphic horizons. In some ore bodies, ore was developed essentially in the entire iron-formation from the foot-wall to the hanging-wall.

There is little doubt that the natural ores are the products of secondary oxidation and leaching of original iron-formations. The original minerals were oxidized to ferric oxides, mainly hematite and goethite, while at the same time, calcium, magnesium, and much of the silica were removed by leaching. There have been extensive discussions as to the source and nature of the oxidizing and leaching solutions, but only two principal hypotheses pertaining to the origin of the natural ores have been proposed. In the first, the oxidation and leaching are postulated to have been accomplished during weathering by downward meteoric waters; in the second, by upward or hot hydrothermal solutions. Although much data pertaining to the ore bodies have been obtained, no field, experimental, or theoretical evidence can be considered as absolutely indicative of either mechanism. However, there is a general concensus today that most of the Mesabi natural ores probably were developed through normal weathering processes (cf., Marsden et aZ., 1969). All the ore bodies are related to an erosion surface which, in general, is the present bedrock surface. They underlie either glacial drift, or in some areas a veneer of Cretaceous strata. The ore extends to various depths, but most is concentrated fairly near the surface. Thus it is inferred that surface waters following permeable zones, such as faults, joints, or fractures, acting over a long period of time, could have produced the observed configuration, distribution, texture, and composition of the Mesabi natural ores.

The origin of the Cuyuna ores is somewhat more complex. Schmidt (1963) believes that the ores were formed by two different processes. The first, probably hydrothermal, process is characterized by deep oxidation and the formation of ore in a rough spatial relation to fractured zones near the southeast edge of the north range where deformation was most intense. The resultant ore bodies are large, deep, tabular, and hematite-rich. A second process apparently took place at a later time as the result of ordinary weathering. An irregular blanket of goethitic ore formed on all exposed surfaces of the unoxidized iron-formation, along joint planes previously opened during the hydrothermal event, and at the peripheries of all the pre-existing hydrothermal ore bodies. Except where developed along the edges of hydrothermal ores bodies, no second-stage ore occurs more than 100 ft (30 m> below the present bedrock surface; the ore distribution, although

205

G. B. Morey

modified by erosion, indicates that the bedrock surface nearly parallels the present surface and was, therefore, one of low relief.

The actual time of ore iormation is not completely documented. Peterman (1966) showed that the Cuyuna rocks were metamorphosed about 1,750 m.y. ago and later affected by ‘. . . “hydrothermal” leaching . . .’ about 1,460 m.y. ago, an interpretation consistent with the geologic evidence outlined by Schmidt (1963). Unfortunately, the time of second-stage ore formation cannot be as precisely established. If it is inferred that both the second- stage Cuyuna ores and the Mesabi ores developed approximately during the same time-interval, some limits can be

placed on that interval, The ore deposits must have been formed prior to Late Cretaceous time on a previously developed bedrock surface of low relief. Parham (1970) has shown that a thick regolith was developed in Mesozoic time prior to the early Late Cretaceous on a peneplain that extended from Manitoba, Canada to at least southern Minnesota. Previously Symons (1966) suggested, on the basis of palaeomagnetic data, ‘. . . that meteoric solutions weathered the primary Animikie iron-formations during the Mesozoic-Cenozoic to form. . . ore deposits.’ Thus, it appears likely that the weathered natural ores were the consequence of a prolonged period of weathering during late Mesozoic time.

Résumé

Les chaines de Mesabi, Gunflint et Cuyuna dans le Minnesota, aux États- Unis d’Amérique (G . B. Morey)

Le Minnesota est l’un des plus grands producteurs de minerai de fer du monde. La plus grande partie du minerai provient des formations de fer du Précambrien moyen dans les chaînes de Mesabi, Cuyuna et Gunflint. La chaîne de Mesabi est une bande étroite de formation de fer qui s’étend sur près de 200 kilomètres à travers la partie septentrionale du Minnesota. C‘est le plus grand producteur du monde avec une production de 2,7 milliards de tonnes brutes de minerai depuis 1892. Sur ce total, environ 809 millions de tonnes brutes (soit 30 %) ont été concen- trées soit à partir de minerais naturels à faible teneur, soit à partir de taconite contenant de la magnétite. Durant les quinze dernières années, la production de minerai naturel a diminué et, en 1968, 59 % de la production totale de la chaîne était du concentré de taconite.

L’extrémité ouest de la chaîne de Mesabi disparaît sous des strates du Crétacé et du Pléistocène ; cependant, la formation de fer suit un tracé sinueux pour se rattacher aux formations de fer de la chaîne de Guyuna dans la partie centre-est du Minnesota. Depuis sa découverte en 1904, la chaîne de Cuyuna a produit et expédié environ 103 millions de tonnes brutes. Dans les années récentes, la production de minerai naturel a diminué de 81 %, passant de 3,6 millions de tonnes en 1955 à 698 O00 tonnes en 1968. Cependant, à l’inverse de ce qui se passe dans la chaîne de Mesabi, aucun concentré de taconite n’est en production courante.

La partie est de la chaîne de Mesabi est tronquée par le complexe de Duluth du Précambrien supérieur mais un prolongement de la formation de fer émerge de nouveau du complexe de Duluth ?i environ 60 kilomètres au nord-est sur la chaîne de Gunflint dans le district de Thunder Bay dans l’Ontario et dans la partie adjacente du Minnesota. On ne trouve aucun minerai naturel sur la chaîne de Gunñint et l’on ne peut guère espérer trouver du minerai de taconite que dans la petite partie de la chaîne qui se trouve dans le Minnesota.

La minéralogie des formations de fer inoxydées comprend du quartz (silex), de la magnétite, de la sidérite, de la stilpnomelane et de la minnesotaïte avec de moindres quantités d’hématite, de calcite, de dolomite, de chamosite, de greenalite et de chlorite. A l’intérieur d‘une zone de contact métamorphique autour du complexe de Duluth, la formation de fer contient du quartz, de la magnétite, des amphiboles, des pyroxènes, du grenat et de la fayalite. La teneur moyenne de la formation de fer inoxydée est d’environ 29 % de fer, 46 % de silice et 0,9 % d‘alumine.

Quant à leur structure, les chaînes de Mesabi et du Gunflint présentent un homocline peu accusé de direction est-nord-est et plongent de 5 à 15” vers le sud-est. Cette direction générale est modifiée par plusieurs plissements en travers dirigés vers le nord et par de nombreuses failles d‘orientation nord-est, est et nord-ouest. Par contre, la structure de la chaîne de Cuyuna est complexe. Les roches sont étroitement plissées dans une série de plis de direction générale nord-est, localement isoclinales et généralement renversées vers le nord-ouest ; on trouve des plissements en travers dirigés vers le nord et quelques petites failles. Des structures moins importantes (petits plissements, failles et monoclines) ont joué un rôle important pour la locaíisa- tion des gisements de minerai naturel sur les chaînes de Mesabi et de Cuyuna.

Les concentrations de minerai naturel offrent une grande variété de formes et de dimensions. Les données géo- logiques et chimiques dont nous disposons actuellement indiquent que les minerais naturels sont le résultat de solutions qui se sont déplacées le long des zones perméables OU l’oxydation et la lixiviation se sont produites. La source et la nature de ces solutions sont inconnues. Les observations faites sur la chaîne du Cuyuna laissent peiiser qu’il y a eu deux périodes d’altération, une période ancienne hydrothermale suivie beaucoup plus tard par une période de désagrégation et de désintégration. Les observations faites sur la chaîne de Mesabi corroborent la thèse que la lixiviation et l’oxydation se sont produites à l’époque cénozoïque par l’action des eaux de surface et des procédés normaux

206


de désagrégation. Les gisements de minerai naturel sont généralement composés de quartz, de martite, d’hématite et degoethiteet, en quantitémoindre, demagnétite, d‘oxyde de manganèse et de kaolinite. Bien que la chimie des gisements

de minerai naturel soit étroitement reliée à la composition des strates dont ils sont dérivés, la teneur moyenne en fer est d’environ 59 % ; la quantité de silice varie entre 2 et 10 % et celle d’aluminium varie de moins de 1 à 6 %.


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208

Physico-chemical conditions of the metamorphism of cherty-iron rocks

Y. P. Melnik and R. I. Siroshtan Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrainian S.S.R.

Cherty-iron rocks make up a considerable part of any iron-formation. These rocks are characterized by a variety of mineral phases (from hematite-bearing nonsilicate jaspilites to silicate-bearing cherts and slates), and by a similarity in bulk chemical composition. For example, jaspilites and slates contain different mineral phases, but are very similar to each other in their silica and total iron content, whereas other rock-forming components are in subordinate quantities and do not play any important role in mineral formation. Thus, the chemical composition of such rocks is characterized by the predominance of iron and silica over other components. This characteristic enables us to separate cherty-iron rocks into a separate iron-siliceous isochemical group. This peculiarity of mineral formation in the system Fe,O,-Fe-SiO, has been described by Korzhins- ky (1940) and Semenenko (1966).

Korzhinsky maintains that, at the early stages of metamorphism, the formation of paragenetic hematite to magnetite was accompanied by the inert behaviour of oxygen. Increased metamorphic alteration is accompanied by a certain activity of oxygen, which results in the replacement of hematite by magnetite ; the former becomes unstable in high-temperature mineral associations. Seme- nenko considers that the activity of iron depends on the presence of ferrous oxide, which reacts with SiO, to form silicates; ferric oxide enters into reactions with silica only when chemical potential of Na,O is high. When Fe0 and Fe,O, are both present in the sediment, the ore mineral magnetite forms first; the remaining Fe0 reacts with SiO, to form ferrous silicates.

The equilibrium of the iron ore minerals (hematite, magnetite and siderite) in metamorphic rocks was investigated both on the basis of thermodynamic calculation (Hawley and Robinson, 1948; Holland, 1959; Kornilov, 1969; Melnik, 1964a, b, 1966a, 196901, and on the basis of experimental data (French and Rosenberg, 1965; Melnik, 19666; Seguin, 1968; Shunzo, 1966). Peculiarities of mineral equilibrium involving the participation of fayalite also have been studied (Melnik and Jarotschuk, 1966). The data obtained helped the more complete understanding

of the role of other components (graphite) and mineral associations (hematite + magnetite, magnetite + fayalitej in elucidating the conditions of metamorphism in cherty- iron rocks, and in the creation of the controlled oxidation-reduction system-via buffers-that fix the fugacity of oxygen-foz (Eugster, 1961; French, 1966; dames and Howland, 1955).

The physico-chemical investigations cited above were, to a certain extent, approximate and not exhaustive because of the absence of thermodynamic constants for a number of rock-forming minerals (amphiboles, micas, chlorites), inaccuracy in calculations of tabular constants (siderite, ferrosilite, fayalite, etc. j, considerable discrepancies between calculated and experimental data, and the absence of the data pertaining to the characteristics of fluid phases under high pressure. Apart from this, many thermodynamic calculations were treated only approximately, without due regard to the effect of pressure.

This paper reports thermodynamic analysis of mineral equilibria carried out using a new system of thermodynamic constants, which included such hydrosilicates as grunerite and minnesotaite. All calculations have been made in accordance with a very precise technique with due regard to the effect of pressure on solid phases (correction for AV> and for fluid phases (correction for the fugacity coefficient - y>.

Below we consider the metamorphic peculiarities of cherty-iron rocks of different lithological composition.

Metamorphism of silicate iron-formation

It is believed that, after the processes of sedimentation and diagenesis, in rocks of this type the stable silicate containing ferrous iron is a mineral of the minnesotaite type (ferrous taicj, silica is in excess, siderite is absent, the presence of ferriferous oxide is possible, and the fluid phase in the intergranular space consists dominantly of water.

Unesco, 1973. Genesis of Precanzbriari iron and manganese deposits. froc. Kiev Syrnp., 1970. (Earth sciences, 9.) 209

Y, P. Melnik and R. I. Siroshtan

Under progressive metamorphism the low-temperature transformation of minnesotaite to grunerite takes place according to the dehydration reaction:

7 Fe,Si,O,,(OH), tt 3 Fe,Si,O,,(OH), + 4 SiO, -I- 4 H20 (i)

The P-T curve for this reaction (Fig. l), at low and moderate pressures, lies in the interval of 250°-280" C and is characterized by a reverse slope. It is worth mentioning that the given position of the curve cannot be thought of as reliably fixed because thermodynamic constants of minnesotaite are based on scarce experimental data. The absence of minnesotaite may testify to the beginning of metamorphism under green schist facies conditions. Gru- nerite is stable from the beginning stages of regional metamorphism and is a typical mineral in iron cherty rocks that have been metamorphosed under both green schist and amphibolite facies conditions. But at the top of the amphibolite facies at temperatures of 640°-690" C grunerite is decomposed (Fig. 1) according to the reaction:

2 Fe,Si,OzZ(OH), +. 7 Fe,SiO, + 9 SO, + 2 H,O (2)

1°C

800

700

600

500

400

300

200

100

FIG. 1. P-T curves of metamorphic reactions in iron cherty rocks with excess of silica. To the right of the curve the predominant fluid component is shown. C, graphite; Hem, hematite; Mgt, magnetite; Fay, fayalite; Cru, grunerite; Min, minnesotaite; Sid, siderite.

into minerals (fayalite and quartz) stable under granulitic facies conditions of association. Ferrous pyroxene-ferrosilite-bearing assemblages are not stable at any temperatures whenever the pressure is below 15,000 bar, as indicated by the positive value of AZ, for the solid phase reaction:

2 FeSiO,+Fe,SiO, + SiO, (3) but where more than 10-15 per cent molecular magnesium is present, the direction of the reaction is reversed,

A summary mineral equilibria diagram with Ig fol-T co-ordinates at a pressure of 5,000 bar is given in Figure 2.

O

-10

-2c

I -30

1 -4c

I P foz

-50

arnphiboliie granulite ' green schist

facies facies facies

-60 IO00 I100 1200

T'K - FIG. 2. Metamorphism of silicate iron-formation (diagram Ig fon-T). Diagram for P, = P/ = C (Pa,o, P,?, Po,> = 5 kbar. Isolines for lgf;I,/fH,o are shown as dotted lines. Fe, iron; Hem, hematite; Mgt, magnetite; Fay, fayalite; Cru, grunerite; Min, minnesotaite.

As can be seen, silicate equilibria with magnetite according to reactions

6 Fe,Si,0Z2(OH), + 7 O, 3c 14 Fe,O, + 48 SiO, + 6 H,O (4)

(5)

are buffered and control, with stable T and P, = the fugacity of oxygen. The facies boundaries also are shown on this diagram.

W e consider as very important the confirmation by thermodynamic data of the instability of hematite with ferrous silicate under P-T conditions characteristic of any metamorphic facies.

Because in essential fluids water can consist of decomposition products only (not counting neutral gases), we

3 Fe,SiO, + O, e 2 Fe,O, + 3 SiO,

21 o


' Total diagrams of equilibria

i, Gru \

t -6 -4 - 2 O c2 +4 16

Detailed diagrams of mineral equilibria

-6 -4 -2 O 1.2 1 4 16 -6 -4 -2 O 1.2 t 4 t6

t 4

+2 -3 -2 -I o tl

green schist facies

(ai

I I -3 -2 *I o +i amphibolite facies

-3 -2 -I o tl

granulite facies

(cl

FIG. 3, Metamorphism of silicate iron-formation (diagrams lgfE20-lgfHz). In (a), (b) and (c), T=6OO0 K (327O C), 800" K

(527" C), and 1,000" K (727" C), respectively. Fe, iron; Hem, hematite; Mgt, magnetite; Fay, fayalite; Cru, grunerite.

can also construct diagrams with PHz0-PH2-T co-ordinates. However, diagrams with lg - lg fH, - T co-ordinates and, especially their isothermal section, are more useful. Such sections are shown in Figure 3 for temperatures corresponding to the changes in metamorphic facies conditions .l From diagrams (b) and (c) the composition of fluid and, in particular, the hydrogen content, which is in equilibrium with the mineral association in question, can easily be defined.

Pure water is an oxidizer in relation to ferrous silicate. When a considerable quantity of water enters the rocks, this can lead to replacement (partial or complete) of barren minerals by magnetite according the reactions:

3 Fe,Si,O,,(OH), + 4 H,O

3 Fe,Si04 + 2 H,O = 7 Fe304 + 24 SiO, + 7 Hz

= 2 Fe304 + 3 SiOz + 2 H,.

(6)

(7) Thus, under the above-mentioned amphibolite facies conditions, approximately 100-150 g pure water-a quantity

quite possible in hydrothermal activity-is required for the oxidation of 1 g grunerite into magnetite. But for the oxidation of magnetite into hematite via a similar process (a variant of hypogene martitization) an enormous quantity of water is needed and, as such, the phenomenon can be only of local importance.

Metamorphism of carbonate iron-formation

The diagnostic features of carbonate iron-formation are the occurrence of siderite in paragenesis with quartz and the occurrence of Fe oxides, both magnetite and hematite; hydrosilicates with ferrous iron do not occur.

The nature of carbonate iron-formation metamorphism depends, to a certain extent, on the presence of hematite

1. Isobar numbers correspond to P,=P~=~(FH,IJ +PE,+ Po?), kbar. Thin incline lines-isobars Ig fo,.

21 1


because, in such cases, under comparatively low temperatures (300" C) and PeO1 = 2,000 bar (Fig. i) the following reaction is possible:

FeCO, + Fe,O, = Fe,O, + CO,. (8) But the equilibrium assemblage siderite + heinatite + magnetite depends greatly on pressure in as much as the slope of the P-T curve defining a decarbonatization reaction is much steeper than that for dehydration reactions.

It is believed that Auid consists dominantly of carbon dioxide, which is why the above-mentioned reaction does not define precisely the lower temperature limit of the green schist facies and why siderite and hematite sometimes occur with grunerite at temperatures up to 390"-420" C and Pcoz = 5,000-7,000 bar.

After disappearance of hematite at the completion of this reaction, pure FeCO, remains unchanged because, at temperatures lower than 400'-500" C, the formation of magnetite from carbonate requires the presence of oxidizers that must be derived from outside the system. In Figure 4, with Pco2 = 5,000 bar, the phase limit of bivariante qui- librium of the siderite + magnetite assemblage corresponds to the temperature interval 38O0-5OO0 C.

Above 400°-500" C dissociation of siderite is theoretically possible according to the reaction:

.3 FeCO, = Fe,O, + 2 CO2 + CO (9) but the proportion CO:CO,=l :2, as required by the

O

-10

-20

1 -3c

1 - 4 1

kfo,

-51

c 51 500 600 700 800 900 1000 II00 1200

T"K - FIG. 4. Metamorphism of carbonate iron-formation (diagram lgfo,-T). Diagram for Ps = Pf = C(Pco,, Pco, Po- = 5 kbar. In a dotted line isolines Ig fe0/fco~ are shown. Area of metastability under the line of graphite is shaded. C, graphite; Fe, iron; Hem, hematite; Mgt, magnetite; Fay, fayalite; Sid, siderite.

equation, is metastable because of the dissociation of carbon monoxide to form graphite:

2CO=C+COz. (10) It has been shown by many investigators that graphite, whether newly formed or already present in the rock, is an oxygen buffer which can regulate fo, and fc0 in the carbonate ñuid. The line of graphite stability divides the diagram (Fig. 4) into two parts. Only the minerals whose fields of stability are crossed by this line-siderite, magnetite and fayalite-can be found in equilibrium with graphite. Mineral associations in the shaded area of Figure 4 cannot exist, as it is physically impossible to create such low values of fog in carbonate rocks. Mineral assemblages found in an unshaded field are stable only in the absence of graphite.

Analogous observations should be taken into consideration when studying the isothermic sections of diagrams with lg fco2-lg feo co-ordinates for the temperatures of various metamorphic facies (Fig. 5). Five petrological conclusions drawn from the analysis of Figures 4 and 5 are as follows.

First, heinatite cannot exist in equilibrium with graphite at any temperature. Under green schist facies conditions, hematite must react to form siderite:

2 Fe,O, + C 4- 3 CO, + 4 FeCO,

6 Fe,O, + C+ 4 Fe,O, 4- CO,

(1 1)

(12)

or magnetite:

depending mainly on feo,. Under amphibolite and granulite facies conditions, reduction is only possible according to reaction (12).

Second, siderite is a stable mineral up to temperatures of the beginning of amphibolite facies.

Third, the equilibrium transformation of siderite into fayalite is thermodynamically impossible as the stability fields of these minerals are separated at any temperature by the magnetite field along the graphite join. Reaction

2 FeCO, + SiO, = Fe2Si04 + 2 CO, (13) is not an equilibrium reaction. Only the phase transformation of siderite into magnetite by reaction (9) is possible, reduction of magnetite to fayalite then follows.

Fourth, at temperatures higher than 500"-600" C under amphibolite facies conditions, the association of magnetite with graphite (Fig. 5(b) and (c)) becomes unstable as a result of the reaction:

2 Fe,O, + C + 3 SiO, + 3 Fe,Si04 + CO,. (14) Finally, under granulite facies conditions, graphite is stable only with fayalite.

By using the diagrams (Figs. 4 and 5), one can find equilibrium fluid compositions. Because increased temperature causes the graphite field to become smaller, the carbon monoxide content in any fluid in equilibrium with graphite must increase.

Carbon dioxide, as well as water, can be an oxidizer for silicates containing ferrous iron, but in this case still

21 2


greater quantities are required. This is why formation of could appear in the course of conjugated oxidation-re- magnetite in such a way can hardly be of ore-making duction reactions involving the participation of COz, importance. evolved in the disassociation of carbonates, for example

reaction (9) or the reaction

Metamorphism of silicate-carbonate iron-formation

A great number of cherty-iron rocks are not represented by purely silicate or carbonate types, but by combined silicate-carbonate ones. Because the rock-forming minerals are ferrous silicates and siderite, the presence of a certain quantity of graphite, in the role of oxygen buffer, is required. Graphite could be formed by the metamorphism of organic carbon originally present in the sediment, or it

3 Fe,SiO, + 2 CO, $ 2 Fe,O, -1- 3 SiO, + 2 CO (1 5)

with further dissociation of CO as per reaction (10). For this reason it is necessary in the analysis of mineral

equilibria to build sectional diagrams along the line of graphite stability (Fig. 5). Using fGo, and fHs0 as independent variables, such a section at a constant temperature will represent the surface on which oxygen fugacity is controlled everywhere by the presence of graphite (Fig. 6), as per the reaction:

c -1- O, CO,. (16)

I I I t6

+5

+4

+3

+2

+ I - 4 -2 0 +2 +4 +6 -4 -2 O +2 +4 + 6

green schist facies amphibolite facies

ia) . íb)

-4 -2 O t2 t4 +6 granulite facies

(C)

FIG. 5. Metamorphism of carbonate iron-formation (diagrams (a), (b) and (c), T=600° K (327O C), 800" K (527' C), and Ig fco,-lg feo). Numbers of isobars correspond to Ps=P~ 1,000" K (727" C), respectively. C, graphite; Fe, iron; Hem, =C(Pco2, Pco, Po,), kbar. Area of metastability is shaded. In hematite; Mgt, magnetite; Fay, fayalite; Sid, siderite.

+?

C6

Ig fco2 +3 +2 I ti

green schist facies

(a)

-1 o 41 +2 +3 +4 +5 amphibolite facies

íb)

.I o +1 c2 +3 +4 +5 granulite facies

(C)

FIG. 6. Metamorphism of silicate-carbonate iron-formation (dia- denote isobars lg foz; thin hachures denote isobars Ig fco. In (a), grams Ig feo,-lg fH,o in the plane of graphite stability). Isobar (b) and (c), T= 600" K (327" C), 800" K (527" C), and 1,000" K numbers correspond to P,=PI=C(PC,?, PH~o, PH?, PGO, PO,), (727" C), respectively. Mgt, magnetite; Fay, fayalite; Gru, kbar. Thin incline lines denote isobars lg fH2; thin dotted lines grunerite; Sid, siderite.

213


This enables us to include in the diagram isobars of oxygen fugacity and isobars of carbon monoxide fugacity as per reaction (10). If foz and ffis0 are known, one can easily determine for every point in the diagram the fugacity of hydrogen, according to the reaction of water decomposition:

2 H2O 2 H2 -1- 0 2 (17)

and draw corresponding isobars. The resulting diagrams make it possible to carry out

a detailed analysis of mineral equilibria and to determine at a given T and P the content of any of five volatile components of the fluid phase. Let us consider, first of all, the relation of siderite and grunerite, as determined by the reaction:

7 FeCO, + 8 Si02 + H,O = Fe,Si,O,,(OH), + 7 COz (18)

Siderite + grunerite (-I- quartz) in equilibrium occupy a wide field under P-T conditions of low-temperature metamorphism, and this is the main mineral association in green schist facies rocks (Fig. 6(a)). However, under amphibolite facies conditions (Fig. 6(b)) this mineral association theoretically can remain only under a very high fluid pressure of more than 7,000-8,000 bar.

Within this stability field the formation of grunerite depends not on temperature but rather on the relation of CO, and HzO in fluid. Thus, the main cause for the development of grunerite in metamorphic rocks is the presence of a sufficient quantity of water, which provides a stable high value for PHto and the relation Pzz0 : Pco, with the expen- diture of water according reaction (18). One can assume that the development of grunerite through metamorphism is connected with processes involving the dehydration of chlorites, hydromicas and other hydro-minerals, to form the slate beds which are interlayered with iron cherty formation.

The equilibrium siderite + grunerite +magnetite is mo- novariant, and at pressures of 4,000-8,000 bar corresponds to a temperature of 420°-530" C. At higher temperatures, the bivariant equilibrium grunerite + magnetite exists, at 550°-650° C, it is replaced by the equilibrium grunerite + fayalite and above 670"-700" C by fayalite only (Fig. 6(c)).

The above data prove that, in fluids of complex composition, temperatures of equilibrium reactions involving the participation of only one volatile component drop markedly, and the appearance of fayalite becomes possible under amphibolite facies conditions.

Metamorphism of oxide iron-formation

The rocks of this type of iron-formation are represented by sediments originally having iron hydroxides that were transformed during diagenesis into goethite and, possibly, into hydromagnetite or magnetite. The transformation .of

goethite into hematite takes place prior to metamorphism at temperatures up to 120°-180" C. The metamorphic transformation of hematite into magnetite according to the reaction

(19) 6 FezO, & 4 Fe,O, + O2 is possible only in the presence of reducing agents (free carbon, gaseous CO or HJ. Where no reducing agents are present, the association of hematite + magnetite is quite stable in all facies of metamorphism, including the granulite facies.

The equilibrium of hematite with other minerals has been considered in previous sections of this report.

Certain peculiarities of low-temperature metamorphism of iron cherty formation

A great number of iron cherty rocks which have undergone metamorphism to the green schist facies are characterized by the close and frequent interlayering of various beds that have different compositions and modes of formation. In some places, within a distance of some centimetres, a bed of siderite-hematite is replaced by one of grunerite- magnetite, and silicate-rich beds having graphite are interlayered with hematite-magnetite beds that contain no free carbon. A detailed investigation of layered rocks has provided enough evidence to suppose that specific equilibrium conditions occur in a limited volume (mosaic or local equilibrium); this equilibrium volume is characterized by a fluid phase of quite different composition.

Probably diffusion, not only of the solid phases, but also of the fluid phase as well, was limited at low temperatures.

These separate volumes of the metamorphosed rocks can be considered as closed systems, and iron cherty formation, taken as a whole, can be treated as a number of closed systems.

Under amphibolite facies conditions, there is a tendency towards the equalization between beds of the fluid composition and thus towards a reduction in the diversity of mineral associations.

The analysis of separate groups of rocks that belong to different metamorphic facies reveals the presence of a certain metamorphic zoning in the iron cherty formations of the Ukrainian Shield.

Metamorphism of iron cherty formations and ore deposition

Banded iron cherty rocks make poor ores, their value being determined by their magnetite contents. Magnetite crystallization probably took place in early stages of metamorphism via the reduction of hematite by carbon according to reaction (12), or by the reaction of hematite and

214


siderite according to reaction (8), and under metamorphic conditions near the amphibolite facies grade by the thermal dissociation of siderite according to reaction (9). Thus, metamorphism at low and moderate temperatures contrib- utes greatly to an increase in ore quality. However, the introduction of water at these metamorphic stages has a negative effect, as it results in the formation of grunerite according to reaction (18) instead of magnetite.

Under amphibolite and granulite facies conditions, a surplus of graphite leads to the development of silicate -grunerite or fayalite-at the expense of magnetite, thus reducing the ore quality to a certain extent. However, introduction of a considrable quantity of pure water sometimes may lead to a reverse process; that is oxidation of silicates to magnetite.

Résumé

Les conditions physico-chimiques du métamorphisme des jòrmatioiis de jeu siliceux (Y. P. Melnik, R. I. Siroshtan)

1. Les roches de fer siliceux, qui ont une grande extension dans les limites du bouclier ukrainien, se rencontrent non seulement en strates qui contiennent de riches dépôts de minerai de fer, mais aussi en minerais à faible teneur, dont la concentration est facile, Les formations de fer siliceux sont des sédiments chimiques métamorphosés ; le fer et la silice dans ces roches proviennent de fumerolles hydro- thermales actives au cours d'un volcanisme subaquatique dans des zones géosynclinales.

2. Les roches de fer siliceux sont de différents types pétrographiques . En raison des proportions différentes des minerais, des silicates et des carbonates, ces roches peuvent être classées en trois types : les roches de minerai, dans lesquelles le composant principal est la magnétite ou l'hématite ; des roches à faible teneur, dans lesquelles le quartz joue un rôle important avec d'autres minerais ; les roches sans minerai, dans lesquelles seuls les silicates et le quartz sont les minéraux constituant la roche.

3. Les principaux minéraux des roches de fer siliceux (qui consistent en Sioz, Feo, Fe,O,, H,O et CO,) sont le quartz, la grunerite, la fayalite, la magnétite, l'hématite et la sidérite. La silice est un composant en excès toujours présent dans les associations minérales sous forme de quartz. Il en est de même de Fe,O,, qu'on rencontre dans les conditions de basse température et qui détermine l'apparition de l'hématite. La présence de grunerite ou de sidérite dépend du rapport H,O/CO,.

4. Suivant les conditions dans lesquelles les roches de fer siliceux se sont métamorphosées, on peut distinguer trois faciès de métamorphisme : le schiste vert, l'amphibolite et la granulite. Les roches de fer siliceux des trois faciès se rencontrent dans la région structurale Krivoyrog- Kremenchug. Cela reflète l'irrégularité des conditions thermodynamiques au cours du métamorphisme. Des diffé- rences dans la composition minérale des roches de fer siliceux sont le résultat de changements dans les conditions du métamorphisme tout le long du filon.

5. Les paramètres physicochimiques de la stabilité de i'hématite, de la magnétite, de la sidérite, de la grunerite, de la fayalite et du graphite dans les minerais de fer siliceux métamorphosés sont à la base d'un nouveau système de

constantes thermodynamiques des minéraux. Les diagrammes d'équilibre minéral ont été établis pour des condi- '. tions P-T de faciès de schiste vert, d'amphibolite et de granulite pour des compositions H,O/CO, d'un fluide (en coordonnées log foi et T ; log fco, et log fco ; log fHz0 et

6. Les limites supérieures de température des associations minérales typiques sont calculées à partir des données thermodynamiques ; elles s'accordent avec les observations pétrographiques : Sidérite + hématite +magnétite Sidérite + grunerite +magnétite + graphite Sidérite +magnétite +graphite Fayalite +magnétite +graphite Grunerite + fayalite + quartz

7. On trouve que l'hématite ne peut pas coexister en équilibre avec les silicates de Fe2+ (grunerite, fayalite) ou avec le graphite dans les roches de fer siliceux quels que soient P, T et la composition du fluide. La transformation métamorphique de la sidérite (+ quartz) en fayalite est peu probable, puisque dans les conditions P-T du faciès de schiste vert la sidérite doit se transformer en grunerite ou en magnétite, suivant la valeur du rapport CO,/H,O dans le fluide. L'association magnétite -1- quartz + graphite devient instable dans les conditions du faciès amphibolitique et se transforme en grunerite ou fayalite stable, suivant la valeur de Ces processus expliquent l'abaissement de la qualité du minerai du fait du métamorphisme, résultat des transitions des minerais de fer (magnétite, sidérite) en silicates.

8. La comparaison des diagrammes et des données pétrologiques permet de montrer que l'équilibre a le carac- tère d'une mosaïque aux premiers stades du niétanior- phisme. Les roches situées dans des régions séparées diffèrent beaucoup, après métamorphisme, quant à la composition du hide (variétés fo,, fEZo, fC,> et quant à la teneur en composés volatils, essentiellement l'eau.

9. Si l'on suit la composition minérale des roches de fer siliceux tout le long du filon de Krivoyrog-Kremenchug, on observe un zonage dans le faciès métamorphique. La partie centrale de cette région consiste en roches de faciès granulitique qui, en direction du nord et du sud, remplace les faciès amphibolitique et de schiste vert. Vers le sud de Krivoyrog le stade du métamorphisme est encore plus

log fH2 ; log fco, et 1% fH@ etc.).

200-400" C ;

400-550" C ; 500-600" C ;

370-500" C ;

640-690" C.

21 5


élevé, avec des roches représentées par un faciès aniphi- bolitique. Les variations qu'on observe dans les types de rocheç de fer siliceux dénotent des variations de P-T au

cours du métamorphisme. L'âge du métamorphisme est de 1 900 à 2 100 millions d'années d'après des datations par la méthode K-Ar .


EUGSTER, H. 1961. Physico-chemical problems of rocks and ores formation. Vol. I, Moscow, Academy of Sciences of the U.S.S.R. (In Russian.)

FRENCH, B. M. 1966. Rev. Geophys., vol. 4, no. 2, p. 223. FRENCH, B. M.; ROSENBERG, P. E. 1965. Science, vol. 147,

HAWLEY, J. E.; ROBINSON, S. C. 1948. Econ. Geol., vol. 43, p. 603. HOLLAND, H. G. 1959. Econ. Geol., vol. 54, no. 2. JAMES, H. L.; HOWLAND, A. L. 1955. Bull. geol. Soc. Amer., vol. 66, no. 12, p. 1580.

KORNILOV, N. A. 1969. C.R. Acad. Sei. URSS, vol. 184, no. 4, p. 939. (In Russian.)

KORZHINSKY, D. S. 1940. Factors of mineral equilibria and mineralogical facies of depth, Bull. Inst. Geol. Acad. Sei. USSR, vol. 12, no. 5. (In Russian.)

MELNIK, Y. P. 1964a. Acad. Sci. USSR, Geol, Rudn. Mesto- rozdenij, no. 5, p. 3. (In Russian.) __ . 19646. Acad. Sci. UIcSSR, Geol. Zum., vol. 5, no. 5, p. 16. (In Ukrainian.)

no. 3663, p. 1283.

-. 1966a. In: Problems theory and experiment in ore formation, p. 58. Kiev, Naukova Dumka. (In Russian.)

--. 19666. In: Research on nature and artificial formation of mineuals, p. 120. Moscow, Nauka. (In Russian.)

-. 1969a. In: Problems of genesis of precambriun iron rocks, p. 259. Kiev, Naukova Dumka. (In Russian.)

-. 1969b. Acad. Sei. UIcSSR, Geol. Zurn., vol. 29, no. 4, p. 13. (In Russian.)

MELNIK, Y. P.; JAROTSCHLK, M. A. 1966. In: Problems theory and experiment in ore formation, p. 98. Kiev, Naukova Dumka. (In Russian.)

-, 1970. Acad. Sci. USSR, Zapiski Miner. Ob., vol. 99, no. 1, p. 3. (In Russian.)

SEGUIN, M. 1968. Nat. canad., vol. 95, no. 6, p. 1195, p. 1217. SEMENENKO, N. P. 1966. Metamorphism of Mobile Zones, Kiev,

SHUNZO, Y. 1966. Econ. Geol., vol. 61, no. 4, p. 768. Naukova Dumka. (In Russian.)

21 6

The Serra do Navio manganese deposit (Brazil)'

W. Scarpelli Industria e Comercio de Minerios S.A. (Brazil)

Introduction The Serra do Navio manganese deposit is in the Federal Territory of Amapá, in northern Brazil (Fig. 1). Production started in 1957 and up to the end of 1969, 10 million metric tons of washed ore had been produced, most exported to North America and Europe. The average grade of the commercial, beneficiated ore varies from 48 to 50 per cent of manganese. Beneficiation consists of crushing, washing and classification by size and density.

The deposit is part of the Precambrian Guyana Shield, at the left bank of the River Amazon. In Serra do Navio and vicinity this shield is composed essentially of gneisses, amphibolites, schists and quartzites (Table 1) plus pegmatites and quartz veins.

TABLE 1. Stratigraphic column of the Serra do Navio district

Series Group General description Lithologic units

Amapá

Serra Manganese pro tores

Navio Biotitic facies do Metasediments Graphitic facies

Quartzose facies

Amphibolites

Quartzites Jornal Amphibolites Schists

? ? ? - Gneisses Gneisses

The oxide ore bodies are the product of secondary enrichment of protores which occur in the Serra do Navio group, outcrop in topographic ridges, and are mined from open pits.

The Gneisses Gneisses are the most common rocks in the neighbourhood of Serra do Navio. The other metamorphic rocks seem to occur as inclusions in them. The predominant type of gneiss is leucocratic and composed essentially of quartz, microline and/or oligoclase, and biotite, occasionally with dark hornblende-rich zones parallel to the foliation. In some places a gneiss with very high quartz content forms prominent ridges. This type of gneiss probably is the product of metamorphism of a silica-rich rock, possibly quartzite.

The Jornal group

The amphibolites of the Jornal group are second to the gneisses in areal extent. They are not uniform in texture or mineralogical composition, varying considerably even in short distances. The predominant mineral is green hornblende, followed by andesine-oligoclase and, in variable percentages, magnetite, titanite, diopside, tremolite, carbonate, and sulphides. Quartz occurs in small veins.

Differences in adjacent bands of amphibolite are conspicuous. The bands vary in grain size (fine to medium or coarse-grained), in structure (well or poorly developed foliation), in texture (presence or absence of oriented minerals) and in mineralogical composition. It is very possible that the amphibolites are derived from a heterogeneous rock sequence. This possibility is reinforced by the occurrence within the amphibolites of belts of quartzites, gneisses and biotite schists concordant with the foliation. These belts show that the rock column from which the amphibolite originated by metamorphism was not homogeneous.

From these observations the origin of the amphibolites cannot be inferred. It is relatively certain that at least part of the sequence is of sedimentary origin, as testified by the

1. With permission of Industria e comercio de Minerios S.A., ICOMI, Rio de Janeiro, Brazil.

Unesco, 1973. Genesis of Precambrian iron and mungunese deposits. Proc. Kiev Synzp., 1970. (Earth sciences, 9.) 217

W. Scarpelli

* Y)

S C A L E - 1:6.000.000

C O N V E N T I O N S

0 city Village

__t_ Roilroa,d

FIG. 1. Geographic location of the Serra do Navio manganese district.

218

The Serra do Navío manganese deposit (Brazil)

belts of quartzites and the belts of diopside-calcite-tremolite-titanite. O n the other hand, some intercalated bands of gneiss have remnants of a typical porphyritic texture, indicative of an igneous origin. N o field or microscopic evidence was found indicative of the origin of the amphibolite layers.

At one point in the River Amapari there is a good exposure of the contact zone between the gneisses and the amphibolites, No gradational change was observed in either of these two rocks toward the contact, which is parallel to their foliation. At the actual contact zone the gneiss alter- nates with the amphibolite in a series of continuousunfolded bands, each one about 0.3-1 .O m thick. It is possible that there is no great difference in age between the rocks from which the gneisses and the amphibolites originated, as both had the same metamorphic history.

The Serra do Navio group

Above the amphibolites there is a sequence of metasediments composed of quartzites, schists and carbonate-rich layers (Fig. 2). These units alternate in a relatively cyclic pattern and are subdivided into three distinct facies, a quartzose, a biotitic, and a graphitic. The minerals which occur in these facies are almost the same, but occur in variable percentages.

Table 2 shows the mineral compositioii of these facies

as determined from thin sections. It must be emphasized that the mineral percentages of these rocks change markedly from layer to layer and place to place, thus the tabulated data are given only to illustrate the variety of observed minerals and mineralogical composition. Quartz and biotite are the most common minerals, followed by graphite, muscovite, sillimanite, garnet (usually almandine), plagioclase (oligoclase, rarely andesine), andalusite, sulphides, and other less frequent minerals. The predominant sulphide is pyrite, followed by chalcopyrite and arsenopyrite.

The manganese protores occur as lenses in the upper part of the graphitic facies. There are two types of protore. The thicker and richer in manganese is composed essentially of rhodochrosite, followed by manganese-bearing silicates such as spessartite, tephroite and rhodonite. The thinner lenses, poorer in manganese, are composed of spessartite, amphiboles, quartz, and graphite.

In the quartzose facies there are layers very rich in calcium carbonate and silicates. They occur in two types. The thicker, which can be described as marbles, are composed essentially of calcite pIus some calcium silicates, The thinner are composed of coarse-grained diopside, tremolite, calcite, and pyrrhotite and can be called a calc-schist.

The bedding of the metasediments, which survived at least three metamorphic phases and is recognizable in the field, was preserved by the development of the metamorphic foliation parallel to it (with some exceptions) and by the great differences in composition of individual layers.

TABLE 2. Composition of the metasedimeiitary facies of the Serra do Navio groupl

Quartzose facies Biotitic facies Graphitic facies (16 samples) (13 samples) (14 samples)

Mineral Maximum Number M a x i m u m Number Maximum Number

minimum positive minimum positive minimum positive ( %) ( %) samples ( %) ( %) samples ( %) ( P4) samples

Average and of Average and of Average and of

Quartz Biotite Graphite Muscovite Plagioclase Silimanite Andalusite Staurolite Garnet Cordierite Tremolite Diopside Titanite Carbonate Hornblende Tourmaline Sulphide Epidote Chlorite

30 9 4 8 12 3 4

traces 5 O 10 I 1 1 3 3 2 3 1

60-3 33-0 9-0 25-0 340 35-0

traces-O 35-0

50-0 10-0 6-0 7-0 8-0 20-0 6-0 15-0 5-0

17-0

-

16 10 12 12 9 5 5 1 15 O 8 2 6 2 3 12 10 5 6

29 34 4 3 4 8 4

traces 8 2 O O O O O 1 1 1 1

50-20 53-20 8-traces 8-0 25-0 25-0 16-0 1-0 25-0 18-0 - - - - - 3-0 3-0 5-0 5-0

13 13 13 10 5 10 6 5 11 2 O O O O O 12 9 1 4

25 14 26 4 3 2 7 6 6 O

- 0 O

traces traces

O 2 1 5 5

38-8 340 45-15 13-0 25-0 5-0 20-0 10-0 10-0 - - - 1-0 1-0

6-0 2-0 30-0 45-0

-

14 13 14 8 4 8 8 10 10 O O O 1 1 O 10 10 4 5

1. Excluding the manganese protorcs and the quartz-free calc-siliceous layers of the quartzose facies.

219

FIG. 2. Geological map of the Serra do Navio District. Contour lines at 50 metre interval. Contacts are inferred from the known data. At the centre the ‘y’ shaped metasediments of the Serra do

Navio group (SNV) overlying the Jornal group (J), and, at West, the gneisses (G). The dotted areas represent the outcrops and float of the manganese ore bodies.

The Serra do Navio manganese deposit (Brazil)

M A N G A N E S E P R O T O R E S

There are two types of manganese protore, one carbonatic and one siliceous, or garnetiferous (Table 3). The carbonatic protore has an average 31 per cent of manganese and is composed principally of rhodochrosite, followed by spessartite, occasionally with veins and bands of tephroite and rhodonite. As accessories there are sulphides (sphalerite, niccolite, gersdorfite), graphite and orthoclase. The texture is mosaic, sometimes disturbed by shearing. In the groundmass of rhodochrosite spessartite crystals grew to a maxi- m u m diameter of 0.5-1.0 mm.

TABLE 3. Chemical analysis of protore samples

A B C D E F G

M n Fe SiO2 A1203

CO2 C Ca0 MgO NazO K2O s As P Ignition loss

36.6 1.3 6.6 2.9 33 $4 na. 4.9 2.9 na. na. n.a. na. n.a.

na.

31.8 0.4 3.5 2.6 32.4 na. 8.1 3.7 0.3 0.1 1 .o n.a. n.a.

n.a.

24.1 3.6 34.7 8.9 8.6 4.3 1.7 1.7

t0.05 <0.05 0.6 0.1 0.03

10.2

33.7 3.7 17.1 2.4 26.3 9.3 0.7 0.3

<0.05 <0.05 0.3 0.1 0.04

25.5

35.7 3.7 15.3 1.9 27.5 7.9 0.6 0.4

<0.05 <0.05 0.01 0.2 0.04

26.7

27.7 2.8 32.6 3.6 19.3 8.2 0.8 0.2

<0.05 <0.05 0.1 0.2 0.03

19.7

3.6 4.8 49.7 12.7 1.2 3.4 3.5 3.1

<0.05 <0.05

1 .o <0.1 0.08

9 .O

A, Carbonatic protore (TG-59, 106 m); B, Carbonatic protore (C2); C, Transition between garnetiferous and carbonatic protores (T6-72, 121 m); D, Carbonatic protore (TG-72, 124 m); E, Carbonatic protore (TG-72, 125 m); F, Transition between garnetiferous and carbonatic protores (T6-72, 128 m); G, Garnetiferous protore (TG-72, 130 m); n.a.. Not available.

The siliceous or garnetiferous protore has 5-25 per cent of manganese and is formed essentially of manganese- bearing garnets up to 3.0 mm in diameter. Graphite is a minor component, occurring as inclusions in the garnets or in interstices between the garnets. Quartz appears in small veins or between the garnets. A very fine-grained amphibole, probably manganesiferous, locally replaces the garnets.

The carbonatic protore, the more important of the two, occurs as lenses of variable extent and thickness. The largest known lens in the Serra do Navio district is almost 1 km long, with a thickness of 20-30 m , the thicker parts in fold axial zones. On average, the lenses of this type of protore have thicknesses of 10-20 m , and lengths of 200-400 m. The garnetiferous protore appears as layers, only occasionally thicker than 2.0 ni, generally at the contact between the carbonatic protore and the enclosing schists. It also inarks the lateral continuity of the manganiferous horizons away from the carbonatic lenses.

The carbonatic protore is the product of metamorphism

of the original manganese-rich sediment, which was a carbonate almost free of silica-bearing clastic fragments. O n the other hand, the garnetiferous protore represents either the metamorphism of an impure manganiferous sediment or the product of metamorphic reactions between the carbonatic protore and the adjoining sediments.

Q U A R T Z O S E E A C I E S

The quartzose facies is a fine-grained metamorphosed chert, with intercalations of calc-silicate layers. Its foliation is parallel to its bedding planes, which are recognizable by variations in mineral composition between adjacent layers. It frequently has calc-silicate layers (not represented in Table 3) which are coarse-grained and composed of variable amounts of calcite, diopside, pyrrhotite, grossularite, and tremolite, plus some titanite.

Probably the original sediment was a mixture of chert, calcite and some clay, and during diagenesis the carbonatic material was segregated from the siliceous in the form of intercalated and discontinuous lensoid layers. Later metamorphic reactions between the carbonatic and the siliceous minerals resulted in the calc-silicates.

B I O T I T I C F A C I E S

The biotitic facies is a biotite-schist. Its foliation is very well developed and commonly discordant with the bedding planes. It has a grain size coarser than the other two facies due to a more intense recrystallization, which also obscured the bedding planes. Sillimanite, andalusite, and a pink almandine occur as porphyroblasts of several millimetres in a mass of quartz and biotite.

The biotitic facies is the product of metamorphism of a pelitic sediment. In the sedimentary column it grades both into the quartzose facies and the graphitic facies.

G R A P H I T I C F A C I E S

Darker than the other two facies, the graphitic facies has its high concentration of graphite as the only real distinction froin the other two. Sometimes it resembles the quartzose facies, with its fine-grain size, high quartz content, and very well-preserved bedding, and sometimes it is more like the biotitic facies, with its coarser grain size, good foliation, poorly recognizable bedding planes, and presence of larger porphyroblasts of andalusite.

The graphitic facies is possibly the metamorphic product of a clayey sediment rich in carbonaceous material which probably originated from organic matter accumulated with the clays and preserved through diagenesis and metamorphism. The free carbon did not originate through decomposition of carbonates, because carbonatic layers almost free of graphite are frequent in the quartzose and gra-phitic facies, none of them revealing signals of decomposition of

221

W. Scarpelli

the carbonates into graphite. Furthermore, the more graphitic zones have no sign of carbonates, although they sometimes are close to lenses of carbonates.

S T R A T I G R A P H I C S E Q U E N C E

In Serra do Navio there are several repetitions of each metasedimentary facies in the stratigraphic column. A statistical verification of the vertical distribution of the facies and their mutual relations revealed a clear transitional pattern between the facies and a tendency for cyclic recurrence of that pattern. This analysis was made possible by the availability of fresh rock core samples collected by diamond drill holes throughout the district.

In a preliminary verification of the preferential associations between the rock units, 421 contacts between the metasedimentary facies and the manganese protores were investigated. At each contact it was noted which rock unit

I occurred at each side of the contact and whether the contact was gradational or not. Although they were known, the thickness of the rock units and the local structure were disregarded. Table 4 presents the numerical results of that analysis.

The most common rock unit encountered was the gra-

phitic facies (with 31 per cent of the analysed contacts), and followed by the biotitic facies (with 25 per cent), the quartzose facies (with 22 per cent), the garnetiferous protore (with 16 per cent), and the carbonatic protore (with 6 per cent). From the frequency of observed contacts of each lithologic unit the probability that each rock unit would occur in contact with each one of the others if they occurred in a random order, and also the actual distribution of the contacts was calcdated by dividing the actual data by the values calculated for a random distribution of the contacts, a correlation coefficient was obtained that indicates if each pair of the rock units occur in contact more or less frequently than expected in a random distribution. Correlation coefficients greater than 1 .O indicate those contacts preferred by nature and those of values lower than 1.0 show the contacts which were not preferred.

The clear predominance of the gradational over the nongradational contacts suggests that the changes in rock lithology during sedimentation proceeded preferentially in a gradative way. The transition zone of the gradational contacts is normally only a couple of metres thick, but it can be expected that it was thicker during sedimentation. Although some of the gradational contacts can be attributed to metamorphic reactions, the metamorphic conditions did not strongly favour these reactions, as is testified by the

TABLE 4. Contacts between the metasedimentary rock units

Probability of contacts Actual distribution by random distribution of contacts Correlation coefficient

Rock unit ( %) ( %) (actual/probability)

Fre- N a m e quency M c M g G B Q M c M g G B Q M c M g G B Q

( %)

All contacts (421) M c 6 - 16 33 27 24 - M g 16 7 - 37 30 26 17 G 31 9 23 - 36 32 6 B 25 8 21 41 - 30 3 Q 22 8 19 40 33 - 4 Standard deviation of the correlation coefficients = 0.45

Gradational contacts (306) M c 5 - 17 34 27 22 - M g 16 6 - 38 31 25 20 G 32 7 24 - 38 31 4 B 26 7 22 43 - 28 O Q 21 6 20 41 33 - 2 Standard deviation of the correlation coefficients = 1.13

Nongradational contacts (11 5) M c 9 - 15 30 25 30 - M g 14 10 - 32 27 31 9 G 27 12 19 - 32 37 11 B 23 12 18 35 - 35 13 Q 27 12 19 37 32 - 7 Standard deviation of the correlation coefficients = 0.04

44

31 5 11

-

65

36 2 6

-

14

16 11 21

-

29 60

46 34

-

25 69

49 31

-

33 31

39 39

-

13 8 38

51 -

O 4 40

61 -

33 19 34

33 -

- 14 15 2.4 25 0.6 46 0.4 - 0.5

- 10 7 3.3 20 0.6 49 0.0 - 0.3

- 20 41 0.9 39 0.9 37 1.1 - 0.6

2.7

1.3 o .2 0.6

-

3.8

1.5 0.1 0.3

-

0.9

0.8 0.6 1.1

-

Note: Mc, Carbonatic protore; M g , Garnetiferous protore; G , Graphitic facies; B, Biotitic facies; Q, Quartzose facies.

O .9 1.6

1.1 0.8

-

0.7 i .8

1.1 0.8

-

1.1 1 .o 1.1 1.1

-

0.5 0.3 1.1

1.5 -

0.0 0.1 1.1

1.8 -

1.3 0.7 1.1

1 .o -

0.6 0.6 0.8 1.5 -

0.5 0.3 0.6 1.7 -

0.7 1.3 1.1 1.1 -

222


relatively large number of nongradational contacts. Prob- ably metamorphic reactions played a relevant role in the formation of garnetiferous protore in the flanks of carbonatic protore, but it seems that this was the only case whereit was relatively effective in originating transitioiizones.

The correlation coefficients and their standard devi- ations show great differences in distribution between the gradational and the nongradational contacts, While for the gradational contacts the standard deviation is 1.13, for the nongradational contacts it is only 0.04, indicating a tendency for gradational contacts to follow preferred orders, while contacts of the nongradational type follow a random distribution more closely.

The data OF the gradational contacts permit the construction of a sequence of the preferred gradational contacts in the following order: carbonatic protore/garnetiferous

protore; garnetiferous protore/graphitic facies; graphitic facies/biotitic facies; biotitic facies/quartzose facies.

This sequence contains all the positive correlation coefficients (higher than 1 .O), for the gradational contacts. It indicates the other most probable rock unit into which a given rock unit is most likely to change.

The vertical disposition of the sequence of contacts was analysed in detail for 160 contacts from the C-2 ore deposit (Fig. 2). This deposit is specifically good for this test since it has a thick and gently folded column of metasediments (Fig. 3), with no possibility of interference of overturned contacts.

The method of calculation was the same as already described except that the vertical disposition of each two rock units in contact was annotated. These results are given in Table 5.

TABLE 5. Vertical sequence of the rnetasedimentary rock units (160 contacts from the C-2 ore deposit)

Top-to-bottom correlation coefficients Rock unit at bottom All contacts Gradational contacts Nongradational contacts of contact (160 contacts) (124 contacts) (3U contacts)

Fre-

PA) N a m e quency M c M g G B Q M c M g G B Q M c M g G B Q

Mc 11 - 2.2 0.7 0.6 0.8 - 3.4 0.8 0.0 0.4 - 0.0 0.5 2.0 1.5 Mg 18 2.4 - 1.1 0.1 1.1 2.7 - 1.4 0.0 0.7 1.3 - 0.0 0.6 2.6

0.3 2.1 G 29 0.4 1.9 - 0.2 1.4 0.2 2.4 - 0.2 1.1 B 22 0.8 0.3 1.9 - 0.5 0.0 0.2 2.4 - 0.3 3.6 0.5 0.0 - 1.5 Q 20 0.4 0.6 0.2 2.8 - 0.5 0.3 0.2 2.8 - 0.0 1.1 0.0 2.7 -

1.1 0.7 -

Note; Mc, Carbonatic protore; M g , Garnetiferous protore; G , Graphitic facies; B, Biotitic facies; Q, Quartzose facies.

The data on the transitional contacts suggests the following sucession of rock units to be the most common, from top to bottom:

graphitic facies

garnetiferous protore carbonatic Drotore

(coeff. = 1.4) ............................................................................

(coeff. = 3.4) (coeff.

(coeff.

(coeff.

(coeff.

(coeff.

= 2.7)

= 2.4) = 2.4) = 2.8) = 1.1)

garnetiferous protore graphitic facies

biotitic facies

quartzose facies ...........................................................................

graphitic facies All the other thirteen possible contacts were less favoured than the listed seven. The truly dominant sequence is composed of only five contacts, starting with the garnetiferous protore at the top and ending with the quartzose facies at the bottom.

The majority of the nongradational contacts have the quartzose facies at the top. This Facies shows a marked tendency to have nongradational contact at the bottom,

which does not occur at its upper contact. The quartzose facies seems to constitute the lower unit of the sequence of metasedimentary units that grade upwards, one into another.

From these data a model of the ideal sedimentary sequence can be inferred. It starts at the bottom with the quartzose facies and grades upwards first into the biotitic facies and then into the graphitic facies: in which the manganese protores occur. Actually this model is disturbed by occurrences of rock units not in that order, but these occurrences do not usually disturb the general pattern and are commonly represented by layers thinner than the usual. Although the manganese protores occur essentially in the graphitic facies, occasionally one of them is associated with the other two facies. These occurrences are actually minor and represented by lenses of protore only some decimetres thick.

CYCLICITY O F THE SEQUENCES

There are several sequences of metasediments in Serra do Navio, each one following roughly the described pattern. In the ideal sequence, the graphitic facies occur at

223

E O

O

o

O

m

:: O


the top, the quartzose facies at the bottom and the biotitic facies is intermediate between the two. As was already pointed out, exceptions to this ideal sequence are not rare. The ideal sequence has more chance to occur where the contacts are gradational.

The lower contact'of each sequence is considered to be at the bottom of the quartzose facies, because it is here that there is the greatest number of nongradational contacts and the greatest differences in the rocks of the top and the bottom.

In Serra do Navio the maximum number of sequences drilled in a single section was three, and in that section the complete column of metasediment was not out. In the mine benches and road cuts the weathering of the rocks and the poverty of outcrops makes the recognition of each sequence very difficult.

ORIGIN OF THE METASEDIMENTS

As mentioned previously, the quartzose facies is a metachert with intercalated lime-rich layers, the biotitic facies is a meta-pelite and the graphitic facies a nietasediment rich in carbonaceous matter. They occur preferentially in that order, from bottom to top, and form repeated sequences, Coarse- and medium-grained clastics are notably absent from these rocks, which were essentially clayey and chemical sediments. This €act may indicate that their source was mature.

The present thickness of each metamorphosed sequence, from the quartzose to the graphitic facies, varies generally from 40 to 60 m . Of the three facies, the quartzose is the thickest and has better lateral continuity and hom- ogeneity. The other two facies are rather variable in thickness and not very homogeneous laterally.

The meta-chert with lime is a chemical marine sediment, deposited at shallow water depth and moderate temperatures. The p H was above 7.5, a necessary condition for the deposition of calcite. Such an environment was oxidizing and relatively free of clastics. Clastics are represented by a few clayey beds, now metamorphosed to biotite-schists.

The gradual transition of the meta-chert into the meta- pelite indicates a change in the conditions which favoured the deposition of chert, an increase in the rate of deposition of clays, or both.

As the meta-pelite became gradually richer in organic materials, it changed into what is now the graphitic facies. This sediment was deposited in a reducing and probably restricted environment, of low p H and negative Eh, conditions also favourable for the deposition of the lenses of rhodochrosite.

A possible model for the deposition of these sediments in their rhythmic sequence will now be described and is illustrated in Figure 4.

The sedimentation probably initiated (Fig. 4(a)) on a shallow marine platform which received small amounts of clastics from an aajoining land. Cherty and limy beds formed farther from the shore line where the deposition of

OXIDIZING P H A S E

ia1

T R A N S I T I O N A L P H A S E

íbl

REDUCING P H A S E \ . (Cl

( J I

FIG. 4. Schematic model of the sedimentary phases of the Serra do Navio group, as inferred from their tendency to form rhythmic sequences. (a) On a flat platform chert and lime form seaward while clay is deposited close to the shore line. (b) The clayey beds eventually cover the chert beds. Water depth diminishes and clayey beds rich in carbonaceous beds start to appear. The areas of deposition start changing to a reducing environment. (c) Organic-rich clayey beds continue to form. The shore line migrates seaward, leaving behind a swampy area and lagoons. Manganese carbonate precipitates as lenses in the strongly reducing lagoons. (d) After subsidence a new cycle of sedimentation starts, with chert seaward and clay landward, covering the previous cycle. In the early stages of subsidence clayey beds covered the manganese-rich lenses.

clastics was minimal. At the same time clays were deposited closer to the land.

As the sedimentation proceeded, the water depth decreased and the zone of deposition of clay moved seaward. The deposition of layers of chert was interrupted and replaced by layers of clay (Fig. 4(b)).

Eventually, as the pile of sediments increased and the water depth decreased, the environment became reducing, with acid pH, negative Eh and diminished water aeration. Organisms deposited with the clay were not decomposed. With the continuous accumulation of material, the environment slowly changed and probably a large swamp was formed, and several lagoons. At this phase the sedimentation became essentially laguna B, and in the lagoons, where the addition of clastics was minimal or absent, the manganese carbonate was precipitated (Fig. 4(c)). The manganese probably stayed in solution as a carbonate during the whole process, precipitating when the environment was reducing and favourable for its deposition. Although the important manganese protores occur at the upper part of the graphitic facies, the manganese was also deposited during all the phases of deposition of the graphitic facies, as is demonstrated by the presence of thin protore layers throughout the graphitic facies.

The next step of the model is a subsidence of the pla.tform, renewing the original conditions of the cycle (Fig. 4(d)). A thin layer of graphitic facies covering the

225

W. Scarpelli

protore is the first indication of that recurrence, but the greatest evidence is the presence of the quartzose facies covering the graphitic facies in nongradational contacts. These two facies represent two strongly different environments of deposition and by themselves suggest such sudden change.

Actually there are exceptions to the stratigraphic cycle as it is described in the model. In some sequences the biotitic facies is locally absent, in others it occurs also as thin lenses inside the other two facies (this may represent some sort of sudden clastic sedimentation). Occasionally the graphitic facies or the quartzose facies are locally absent. These exceptions do not invalidate the model, which was constructed to put together the known tendencies of the sedimentation processes. Of course, the model might be modified in the light of additional data.

METAMORPHIC PHASES

The metasediments were metamorphosed three times. The older metamorphism was dynamic, of the amphibolite facies of metamorphism (according to the classification of Turner, 1968), accompanied by minor folding. It had a strong component of load pressure, which favoured the development of foliation parallel to the bedding planes. Quartz, biotite, and plagioclase are the most common minerals of this phase and are preserved as oriented inclusions in younger porphyroblasts. As inclusions, they occur with mosaic texture, have a very fine grain size, and do not present evidence of tectonic deformation.

Next a thermal metamorphism occurred, of the hornblende-hornfels facies of Turner (1968), with the growth of porphyroblasts of andalusite and almandine in the three metasedimentary facies, of grossularite, tremolite and diopside in the calcic zones of quartzose facies, of spessartite, rhodonite and tefroite in the protores, and of cordierite iri the biotitic facies. Andalusite is more frequent and coarser in the graphitic facies, where it is very rich of inclusions of graphite.

The youngest metamorphism was a dynamic nieta- morphism, of the amphibolite facies, characterized by rock deformation, folding and mineral recrystallization. The textures of the quartz-rich zones became sutured and broken minerals are common. In general only the porphyroblasts and their inclusions did not recrystallize or become deformed. In the three facies sillimanite replaced biotite and andalusite and was accompanied by muscovite. In the protores and the calcic zones amphiboles developed across older minerals. During this metamorphic phase a pneuma- tolitic phase took place, characterized by the formation of pegmatites and local tourmalinization.

Samples of the metasediments were analysed by the Laboratory of Geochronology of the University of São Paulo, Brazil. The K-Ar method gave an age of 1,710- 1,770 m .y. It is assumed that this age represents the age of the Iast dynamic metamorphism and folding.

STRUCTURE

The overall structure of the district is poorly known due to the scarcity of outcrops and the presence of identical rocks in several levels of the stratigraphic column, making lateral correlations difficult. Some informations is known from drilling and mine exposures.

The folds vary from open (Fig. 3) to closed, plunging northwest and southeast at small to moderate angles. Their axial planes dip northeast. In some areas they turn into isoclinal folds, with the parallellimbs dipping steeply northeast.

The carbonatic protore has its maximum thickness in the nose of closed folds due to plastic recrystallization and concentration of rhodochrosite in the axial zones.

The biotitic facies occasionally appears deformed, like an incompetent layer. The quartzose facies was more resistant to deformation, forming concentric folds, which helped to avoid strong tectonic deformation of the column of metasediments.

SECONDARY ENRICHMENT

The predominant ore mineral is cryptomelane, followed by pyrolusite and occasionally other minor manganiferous minerals. In hand specimens one can distinguish between an amorphous mass (essentially cryptomelane) and crystalline zones (pyrolusite). These minerals replace the manganese carbonates and silicates of the protores down to the water table, leaving behind decomposed remnants of the manganese silicates in some cases.

The ore is in a continuous process of secondary enrichment. Vugs are formed during the weathering and subsequently filled by successive bands of cryptomelane and/ or pyrolusite, in mammillary textures, probably formed by deposition from colloidal solutions.

The high-grade ore replace the carbonatic protore in situ (Fig. 3), with the original bedding still occasionally recognizable. The garnetiferous protore produces an ore of lower grade, due to both its original lower content in manganese and the formation during weathering of stable minerals rich in silica and iron.

The manganese oxides generally occur in the enclosing schists and quartzites only as small, isolated, and unec- onomic replacement pockets and veins, but when the structure is favourable for their accumulation, as in the axial zones of synclines, the manganese oxides replace the weathered schists intensively, forming schistose ores, which arecharacterized by theabundance of remnants of the schists.

Float ore is common in the hill sides below the outcrops of manganese ore. It is composed of detrital ore fragments plus spheres of manganese oxides chemically precipitated around a nucleus of quartz or other minerals (granzon).

The bulk of the gangue is composed of kaolinite and goethite, produced by the weathering of schistose rocks and manganese silicates. These minerals are followed by gibbsite, also a product of weathering, and residual quartz and graphite.

226


Résumé

Le gisement de manganèse c h Serra do Navio, an Brésil (W. Scarpelli)

Les dépôts de minerai de Serra do Navio se rencontrent dans une séquence métasédimentaire qui se trouve, au point de vue stratigraphique, au-dessus d’une ceinture d‘amphibolites. Ces roches ont été métamorphosées au moins trois fois, et le métamorphisme le plus récent remonte à 1,7 milliard d’années.

Les sédiments en cours de transformation se succèdent et peuvent être classés en trois faciès : siliceux, à biotite et à graphite. La distribution verticale de ces trois faciès révèle une tendance à une périodicité rythmée. Ils forment des séquences sédimentaires successives, chacune d‘elles partant à la base du faciès siliceux pour se développer plus haut dans le faciès à biotite, et enfin dans le faciès à graphite. Des lentilles de rhodochlosite, qui ont eu leur origine dans le gisement de minerais, se rencontrent à la partie supé- rieure du faciès à graphite, du sommet duquel part la séquence suivante, avec encore le faciès siliceux à la base. On observe au moins trois séquences de ce genre dans la région.

On suppose que les sédiments originaux se sont déposés dans un milieu néritique peu profond en cours de subsidence. On constate en général qu’à la base de chaque cycle sédimentaire se trouvent du chert et de la chaux, tendant

graduellement, en s’élevant, vers une zone pélitique qui plus tard s’est enrichie en matière organique. Il semble que chaque cycle a commencé après une transgression du niveau marin, les sédiments pélitiques se déposant du côté de la terre, tandis que les sédiments chimiques riches en silice se déposaient du côté de la mer. Au fur et à mesure que l’empilement des sédiments croissait et que la zone pélitique recouvrait la zone riche en silice, la profondeur de l’eau diminuait. C‘est alors que les sédiments pélitiques s’enri- chirent de débris organiques, Lorsque la profondeur de l’eau a atteint son minimum, la circulation de l’eau a été contrariée et quelques lagons se sont formés, où se sont trouvées des conditions favorables au dépôt de manganèse sous forme de carbonate.

On peut encore distinguer trois phases de métamor- phisme dans les sédiments au cours de leur transformation : deux phases dynamiques et une phase thermale. La consé- quence des deux métamorphismes dynamiques est la présence de structures plissées superposées.

Les gisements de minerais actuels ont été formés par l’altération atmosphérique et par l’enrichissement subsé- quent de lentilles de protore riche en rhodochrosite, qui ont été remplacés par des cryptomélanes et des pyrolusites. Le protore inaltéré contient environ 31 % de manganèse, tandis que les gisements supergenes en contiennent environ 50 %.


CRESSMAN, E. R. 1962. Non detrital siliceous sediments. Prof. Pap. US. Geol. Surv., 440 T.

DORR, J. VAN N., II; PARK, C. F., JR.; DE PAFIA, G. 1949. Manganese deposits of the Serra do Navio district, Territory of Amapá, Brazil. Bull. US. Geol. Surv., 964 A. __ ; SOARES COELHO, I.; HOREN, A. 1956. The manganese deposits of Minas Gerais, Brazil. XX I72t. Geol. Congr., Mexico, vol. III, p. 279-346.

KRAUSKOPF, K. B. 1967. Introduction to geochemistry, New York, N.Y., McGraw-Hill.

KRUMBEIN, W. C.; GARRELS, R. M . 1952. Origin and classification of chemical sediments in terms of p H and oxidation- reduction potentials. J. Geol., vol. 60, p. 1-33.

LEINZ, V. 1948. Estudo genético do minério de manganês da Serra do Navio, Territorio do Amapá. Vol. 20, no. 2, p. 211- 21. Rio de Janeiro, Academia brasileira de ciências.

NAGELL, R. H. 1962. Geology of the Serra do Navio manganese district, Brazil. Econ. Geol., vol. 57, p. 481-98.

PARK, C. F., JR. 1956. Manganese ore deposits of the Serra do Navio district, Federal Territory of Amapá, Brazil. XXInt. Geol. Congr., Mexico, vol. III, p. 347-76.

SCARPELLI, W. 1966. Aspectos genéticos e metamórficos das rochas do distrito de Serra do Navio. Avulso Dep. nac. Prod. min., Rio de J., no. 41, p. 37-56.

-. 1968. Precambrian metamorphic rocks of Serra do Navio, Brazil, unpublished work, Stanford University.

SUJKOWSKI, ZB. L. 1958. Diagenesis. Bull. Amer. Ass. Petrol. Geol., vol. 2, p. 2692-717.

TURNER, F. J. 1968. Metamorphic petrology, mineralogical und field aspects. New York, N.Y., McGraw-Hill.

VALARELLI, J. V. 1966. Contribuição á mineralogia do minério de manganês da Serra do Navio, Amapá, Avulso Dep. nue. Prod. min., Rio de J., no. 42, p. 83-98.

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Discussion

I. P. NOVOKHATSKY. What are the manganese minerals in carbonate rocks?

W . SCARPELLI. In fresh carbonatic protore the manganese mineral sare rhodochrosite, spessartite plus rhodonite and tephroite. In the oxide ore the main manganese minerals are cryptomelme and pyrolusite.

I. P. NOVOKHATSKY. What is the depth of the oxidation zone?

W. SCARPELLI. Usually the oxidation goes down to 70- 100 m in the hills where the ore bodies occur.

R. T. BRANDT. What do you consider was the reason for the deposition of large amounts of manganese in these sediments, and where did it come from originally?

W . SCARPELLI. The reason for such a deposition was the conjunction of favourable conditions for the accumulation of manganese in solution and the existence of environments favourable for the deposition of the manganese as carbonates with minimum admixture of clay.

W e do not know where the manganese came from.

J. H. GROSSI SAD. The second metamorphism is called thermal. Why? Did you find intrusives related with this metamorphic phase?

W . SCARPELLI. W e called that metamorphism a thermal one because temperature seems to be the major or the only factor which affected the rocks. Porphyroblasts grew in all the rock units described, and they grew without deformation, indicating that pressure did not affect the rock units during that metamorphic phase. W e did not find intrusives related with this metamorphic phase.

J. VAN N. DORR. Most of the protore is MnCO,, but in one deposit it is picrotephroite, I believe. D o you have

any idea why picrotephroite formed at this particular deposit (Chumko)?

W. SCARPELLI. Tephroite, like rhodonite, was formed through chemical reactions between silica and rhodochrosite. The quantity of silica and rhodochrosite involved in the reactions dictated whether rhodonite or tephroite was formed. All the rocks of the district, except the protore, show small quartz veins. In the protore we do not get veins of quartz, but we get veins of tephroite and/or rhodonite in areas where the enclosing rocks have quartz veinlets, W e accept the possibility that silica entered the protore along fractures and reacted with the carbonate forming the silicates, under high temperature, during the thermal metamorphism. Tephroite and rhodonite occur mainly as large crystals.

J. VAN N. DORR. Can you tell us anything about the pelletizing plant for manganese oxide pellets under construction?

W. SCARPELLI. The pellet plant is under construction, and we expect that it will start operating in the second semester of 1971, producing the first manganese ore pellets in the world.

B. CHOUBERT. Can you explain the technique you use to obtain the numbers to calculate the correlations between elements of various formations.

W. SCARPELLI. I compared the real data obtained through detailed examination of all the diamond drill cores available with a set of theoretical numbers calculated to express the probability of occurrences of contacts at random. The coefficients represent the ratio between the real data and the theoretical set for a random distribution of contacts. They show how nature preferred some contacts over others.

228

Genetic studies on the Precambrian manganese formations of India with particular reference to the effects of metamorphism S. Roy Department of Geological Sciences, Jadavpur University (India)

Extensive deposits of manganese ore occur in the Precam- brian shield of the Indian subcontinent in the states of Madhya Pradesh, Maharashtra, Orissa, Bihar, Andhra Pradesh, Gujarat, Mysore and Goa (Fig. 1). Both syngen-

I SCALE I W O IO0 ZPO MILES I ODELHI

FIG. 1. Map showing manganese ore deposits in India. Madhya Pradesh-Maharashtra ore belt; 2. Gangpur-Bamra deposits; 3. Panch Mahals Dt deposits; 4. Jhabua Dt deposits, Madhya Pradesh; 5. Srikakulam Dt deposits, Andhra Pradesh; 6. Bonai- Keonjhar deposits, Orissa; 7. Kalahandi-Koraput-Patna deposits, Orissa; 8. Sandur-Bellary deposits, Mysore; 9. Shimoga Dt deposits, Mysore; 10. North Kanara deposits, Mysore; 11. Banswara deposits, Rajasthan; 12. Goa deposits.

etic and supergene epigenetic deposits are present, of which the first type is by far the majority (Roy, 1966). This paper will deal with the syngenetic manganese formations consisting of interbanded manganese ores and manganese silicate rocks that show evidences of regional (and locally thermal) metamorphism to different grades.

Distribution and geologic setting

The syngenetic manganese formations of India occur in the Sausar (Madhya Pradesh-Maharashtra), Aravalli (Rajas- than-Madhya Pradesh), Champaner (Gujarat), Gangpur (Orissa) and Khondalite (Andhra Pradesh) groups, all meta-sedimentary formations of Precambrian age. The geochronological relationships of the above groups are given in Table 1.

MANGANESE FORMATIONS O F THE SAUSAR GROUP

In India, syngenetic manganese formations are best developed in the Sausar Group that covers the districts of Nagpur and Bhandara (Maliarashtra) and Chhindwara and Balaghat (Madhya Pradesh). The manganese formations run NE.-SW., E.-W. and NW.-SE. as an arcuate belt, the obtuse bulge facing south, for more than 130 niiles (208 km) with an average width of 20 miles (32 km). The Sausar Group is represented by a miogeosynclinal orthoquartzite- carbonate sequence and the geological succession has been established, based on the study of several workers (Naraya- naswanii et al., 1963). The different formations of the Sausar Group are represented by pelitic, psammopelitic, psammitic and calcareous rocks (Fig. 2). Igneous rocks are almost absent in the Sausar sequence, except for rare occurrences of late and post-tectonic granite plutons in northern Nagpur, northern Bhaiidara and Balagliat districts.

The manganese formations of the Sausar Group are entirely stratigraphically controlled. They are particularly associated with pelitic rocks in the bottom, middle and

Unesco, 1973. Genesis of Precambrian iron and nianganesc deposits. Proc. Kiev Sump., 1970. (Earth sciences, 9.) 229

S. Roy

BHANDARA DT. CHHINDWARA OT. N A G P V R DT.

GOYIARI WbDHONb- GuMGAON- MINSIR-KbNORL CHIIL4- SITA-

MAHbRKUND AllEA UOHBAON AREI JUNAWAWI AREA sho*G'-Do*cR' BUZURG A R E I

RAMDONGRI-

TABLE 1 (after Sarkar, 1968)l

B A L A G H A T DT.

TIIOOI-SlTAPbTORE RAMIAMh-IPETIA I8ARWELIJlKWA AREA A R E A bREA

Madhya Pradesh and Maharashtra Gujarat Orissa Andhra Pradesh

Satpnua Cycle Sausar Group 846-986 my.

Asavalli Cycle Aravalli Groupz Champaner 950-1,500 m.y. Group Sed., c. 2,000 my.

Eastern Ghat Cycle

Gangpur Group 846-946 m.y. Sed., 1,700-2,000 m.y.

Khondalite Group Phase II, c. 1,600 my. Phase I, c. 2,650 my.

1. The ages indicate closing of events. Sedimentation (Sed.) ages are also given where determined. 2. The Aravalli Group mainly covers the state of Rajasthan and partly Madhya Pradesh. Syngenetic manganese formations in this group are only found in

Madhya Pradesh.

upper parts of the Mansar formation. Besides the occurrences in the Mansar formation, manganese deposits are also enclosed in Lohangi (marble and calc-silicate rocks) and Tirodi biotite gneiss formations on a very minor scale

The manganese formations are independent beds of

ditel) that are intimately interbanded among themselves and with the enclosing pelitic meta-sediments of the Mansar formation. Primary sedimentary lamination is discernible

(Fig. 2). 1. Regionally metamorphosed manganese silicate rocks, essentially composed of spessartite and quartz, with or without other manganese silicates. First named gondite by Fermor (1909), later elaborated by

oxidic manganese ores and manganese silicate rocks (gon- Roy and Mitra (1964) and Roy and Purkait (1968).

LEGEND

BICHUA AND JUNNANI FDRMATIOU.

CHORBAOLI FDRMATIO~.

M A N S A R FORMATICM

@ GREENSCHIST FACIES

@ ~ u ~ F ~ ~ : E -ALMAMOIME

KYAUITE -ALMANDIME @ SUBFACIES. SILLIMANITE-ALMANDINE- @ MUSCOVITE SUBFACIES. LOHAHGI FCRMATIOM.

SITASAOHGI FORMATION.

TIRODI BIOTITE GNEISS FORMATION.

BEOS ABSENT.

9 MANGANESE ORE. FIG. 2. Synoptic diagram showing stratigraphic positions of the manganese ore horizons in the manganese ore belt of Madhya

Pradesh and Maharashtra. (Thickness and lateral extent not to scale.)

230

Genetic studies on the Precambrian manganese formations of India with particular reference to the effects of metamorphism

in the ores themselves. The orebodies, gondite, and the pelitic nieta-sediments show concordant relationships and are Co-folded in different scales. These manganese formations were originally laid down as sediments from a nonvolcanogenic source.

The pelitic rock of Mansar formation show effects of regional metamorphism to different grades. The rocks in the Bharweli-Ukwa area, Balaghat Dt, Madhya Pradesh, have been metamorphosed to the quartz-albite-epidote-almandine subfacies of greenschist facies. At the Dongri Buzurg- Kurmura area, Bhandara Dt, Maharashtra, the same grade of metamorphism has been attained by the rocks, increasing progressively eastward to staurolite-almandine and kyanite- almandine subfacies (almandine-amphibolite facies) in the Chikla-Sitasaongi area. The major part of the Sausar tract, including the areas around Tirodi-Sitapatore, Netra-Ram- rama (Balaghat Dt), Mansar-Kandri, Gurngaon-Ramdon- gri (Nagpur Dt) and Gowari Wadhona, Sitapar-Kachidhana (Chhindwara Dt) have been metamorphosed to sillimanite-almandine-muscovite subfacies (almandine-amphibolite facies). Since the pelitic rocks and the manganese formation are associated intimately in a syngenetic sequence and were later metamorphosed together, it is assumed that both of them have been subjected to the same range of pressure- temperature changes.

MANGANESE FORMATION O F THE G A N G P UR GROUP

The manganese formation occurs as an important member in the meta-sedimentary non-volcanogenic sequence of the Gangpur Group. Syngenetic manganese orebodies and gondite occur interbanded and Co-folded in pelitic rocks in Ghoriajor-Monomunda area, Sundargarh Dt, Orissa. The manganese formation occupies the core of the Gangpur anticlinorium (Krishnan, 1937) and constitutes, along with the pelitic rock, the Ghoriajor formation at the base of the Gangpur Group. The pelitic rocks of the area have been metamorphosed to staurolite-almandine subfacies.

MANGANESE FORMATIONS OF THE ARAVALLI AND CHAMPANER GROUPS

The Aravalli Group is well represented in the states of Rajasthan and part of Madhya Pradesh (Jhabua Dt) and its southward continuation in Gujarat has been designated as the Champaner Group. The Aravalli (and equivalent Champaner) Group is constituted of quartzite, limestone and calc-silicate rocks, and phyllites and mica schists with syngenetic beds of manganese oxide ores and manganese silicate rocks. The syngenetic manganese orebodies at Shi- varajpur (Panch Mahals Dt, Gujarat) are interbanded and Co-folded with phyllite and quartzite in a non-volcanogenic sedimentary sequence that has been regionally metamorphosed to the quartz-albite-muscovite-chlorite subfacies. Manganese orebodies and gondite, interbanded

with quartzite and phyllite at Kajlidongri, Jhabua Dt, Madhya Pradesh, have been metamorphosed to the quartz- albite-epidote-biotite subfacies. A n isolated deposit of manganese oxide ore and manganese silicate rock (kodu- ritel), interbanded and Co-folded with wollastonite-diopside hornfels, has been thermally metamorphosed to the pyroxene-hornfels facies by porphyritic biotite-granite at Jothvad, Panch Mahals Dt, Gujarat.

MANGANESE FORMATIONS O F THE KHONDALITE G R O U P

The Khondalite Group meta-sediments cover the eastern Ghats region of south India and are made up of rocks metamorphosed to the granulite facies (garnet-sillimanite- graphite granulite, calc-granulite, garnetiferous quartzite, charnockite etc.). At Kodur-Garividi-Devada and Gar- bham, Srikakulam Dt, Andhra Pradesh, syngenetic manganese orebodies (and rarely manganese silicate rocks) occur interbanded and Co-folded with the meta-sedimentary members of the Khondalite Group. Igneous rocks, other than much later granite plutons and pegmatite, are absent in this area. In the Kodur group of mines, the manganese orebody, occurring as conformable beds, is enclosed in calc-granulite, whereas in the Garbham area, the orebody occurs interbanded with garnetiferous quartzite. In both the areas, garnet-sillimanite-graphite granulite are present in close association, though never in direct contact with the orebodies.

Mineralogy and texture

The mineralogy of the manganese oxide orebodies and the manganese silicate rocks of the metamorphosed manganese formations of India has been summarized by Roy (1966). The mineral assemblages representing different grades of metamorphism have been tabulated in Tables 2 and 3. In these tables, the mineral assemblages of metamorphosed manganese formations from other countries are also listed from various sources (Dorr et al., 1956; Hewett et al., 1961; Horen, 1953; Huebner, 1967; Hutton, 1957; McAndrew, 1952; Mohr, 1964; Roper, 1956; Segnit, 1962; Servant, 1956; Watanabe, 1959; Westerveld, 1961; Woodland, 1939).

Besides the metamorphosed manganese formations listed in Tables 2 and 3, several other important deposits deserve particular mention. These could not be included in the tables due to uncertainties about their precise grade of metamorphism. The metamorphosed manganese formations of Ghana (spessartite-rhodonite-rhodochrosite (Service, 1943)); Madagascar (spessartite-rhodonite-tephroite-apatite-quartz (Boulanger, 1956)); southern Ural (spessartite-rhodonite-bustamite-piedmontite (Betekhtine, in

1. The n a m e kodurite was proposed by Fermor (1909) to designate thermally metamorphosed manganese silicate rocks of spessartite- andradite garnet, potash felspar and apatite.

23 1

S. Roy

TABLE 2. Sedimentary deposits of manganese metamorphosed to greenschist facies

Mineralogy Area Type of deposit Silicate-carbonate

Oxide Silicate

India Shivarajpur, Gujarat Quartz-albite-muscovite- Braunite - -

Kajlidongri, Madhya Quartz-al bite-epidote- Braunite, bixbyite, Spessartite, - chlorite subfacies

Pradesh biotite subfacies hollandite, rhodonite, blanfordite, jacobsite juddite, winchite,

manganophylite, alurgite, quartz

Bharweli-Ukwa, Quartz-albite-epidote- Braunite, bixbyite, - Madhya almandine subfacies hollandite Pradesh Dongri Buzurg- Quartz-albite-epidote- Braunite, Spessartite, Kurmura, almandine subfacies hollandite, rhodonite, Maharashtra jacobsite, quartz

manganite

United States of Anierica Big Indian Deposit, Quartz-albite-muscovite- - - Sierra Nevada chlorite subfacies

Smith Prospect, Sierra Zeolite or lower part Hausmannite - Nevada of greenschist facies

Calaverous Formation, Greenschist facies - Sierra Nevada

Buckeye Deposit, Blueschist facies Hausmannite, - California braunite

Brazil Merid Mine, Minas Quartz-albite-epidote- Hausmannite - Gerais almandine subfacies

United Kingdom Merionethshire, Wales

Australia Tamworth District, New South Wales

New Zealand Western Otago

Africa Tiéré

Lower part of greenschist facies

Greenschist facies Hausmannite, jacobsite

Quartz-albite-muscovite- - chlorite subfacies

Quartz-albite-muscovite- - chlorite subfacies

Rhodochrosite, rhodonite, tephroite

Rhodochrosite, rhodonite, tephroite, spessartite yellow IA silicate Rhodochrosite, rhodonite, tephroite, spessartite, neotocite, bementite

Rhodochrosite, bementite IA and i2Â silicate

Rhodochrosite, manganoan calcite, rhodonite, pyroxmangite, tephroite, spessartite, kupfferite, bementine, neotocite, graphite, quartz

Rhodochrosite, spessartite, quartz

Mangandolomite, rhodonite, knebelite, tephroite, quartz

Rhodochrosite, rhodonite, spessartite

Spessartite, quartz -

232


TABLE 3. Sedimentary manganese deposits metamorphosed to almandine-amphibolite and pyroxene-hornfels facies

Mineralogy Area Type of deposit

Oxide Silicate Silicate-carbonate

India Chikla-Sitasaongi area, Maharashtra

Staurolite-almandine and kyanite-almandine s ubfacies

Braunite, bixbyite, hollandite, jacobsite, vredenburgite

Spessartite, rhodonite, - tirodite, alurgite, manganophyllite, plagioclase, quartz Spessartite, rhodonite, - blanfordite, brown manganiferous pyroxene, winchite, juddite, tirodite, alurgite, manganophyllite, . apatite, quartz Spessartite rhodonite, - manganoan diopside, brown manganiferous pyroxene, blanfordite, winchite, tirodite, piedmontite, manganophyllite, apatite, plagioclase, quartz Spessartite, rhodonite, - manganoan diopside, brown manganiferous pyroxene, blanfordite, winchite, tirodite, juddite, piedmontite, manganophyllite, apatite, calcite, quartz Spessartite, rhodonite, - apatite, quartz

Tirodi-Stapatore area, Madhya Pradesh

Sillimanite-almandine- muscovite subfacies

Braunite, bixbyite, hollandite, hausmannite, jacobsite, vredenburgite

Netra-Ramrama area, Madhya Pradesh


Braunite, bixbyite, hollandite, hausmannite, jacobsite, vredenburgite

Gowari-Wadhona, Madhya Pradesh


Braunite, bixbyite, hollandite, jacobsite, hausmannite, vredenburgite

Kodur-Garbham area, Andhra Pradesh

Granulite facies Braunite, hollandite, jacobsite, hausmannite, vredenburgite

Norway Mount Brandnuten Almandine-amphibolite

facies Braunite, jacobsite, hausmannite

Spessartite, rhodonite, - quartz

Africa Otjosondu Almandine-amphi bolite

facies Braunite, bixbyite, hollandite, jacobsite, hausmannite, vredenburgite

Spessartite, rhodonite, - diopside, acmite, quartz

India Jothvad, Gujarat (thermally metamorphosed deposit)

Pyroxene, hornblende hornfels facies

Braunite, bixbyite, hollandite, hausinannite

Spessartite, andradite, - rhodonite, blanfordite, brown manganiferous pyroxene, winchite, alurgite, manganophyllite, apatite, quartz

Japan Noda-Tamagawa, Iwate prefecture (thermally metamorphosed deposit)

Pyroxene/hornblende hornfels facies

Hausmannite, braunite, manganosite, pyrochroite

- Rhodochrosite, rhodonite, pyroxmangite, tephroite, spessartite, galaxite, bementite

233

S. Roy

Zvéreff, 1953)); sou thernKhingan (rhodochrosite-bustamite- rhodonite-tephroite and braunite-hausmannik-hematite- magnetite (Chebotarev, 1960)), and the Postmasburg-Tha- bazimbi area (braunite-bixbyite-hausmannite-jacobsite-vre- denburgile-chalcophanite (De Villiers, 1956)) belong to this group.

The metamorphosed inanganese orebodies of India characteristically exhibit banding both on a macro- and micro-scale. In the quartz-albile-epidote-biotite subfacies at Kajlidongri, relict colloform texture and relict crustified

fracture-filling veins are exhibited by braunite and bixbyite (Figs. 3 and 4). Such textures indicate that colloform higher oxides (pyrolusite, cryptomelane, etc.), originally deposited in sediments, were reduced by rising temperature to form braunite and bixbyite, though the broad textures remained virtually undisturbed. On recrystallization, bixbyite usually shows idioblastic habit (Fig. 5) owing to its strong force of crystallization. Equant and strain-free recrystallized grains of hollandite, of varying size, often constitute entire bands (Fig. 6). In almandine-amphibolite facies, vredenburgite

FIG. 3. Relict colloform texture with bixbyite (white) at the core and braunite (light grey) forming the rim. A relict crustified vein showing these two minerals is also seen. Kajlidongri. ( x 90.)

FIG. 4. Same feature as in Figure 3, further enlarged and showing details. Braunite (grey) at the core is enclosed by bixbyite which is further rimmed by braunite. The subhedral to euhedral habit of bixbyite is characteristic. Kajlidongri. ( x 450.)

FIG. 5. Idioblastic crystals of bixbyite in quartz gangue. Hematite (white) is released when part of bixbyite converted to braunite (grey). Bharweli. ( x SOO.)

FIG. 6. Strain-free equant grains of hollandite formed by recrystallization during metamorphism Gowari Wadhona. ( x 200.)

234


showing widmanstatten intergrowth (due to exsolution) of hausmannite in jacobsite (Fig. 8), exhibits xenoblastic to granoblastic texture with braunite (Fig. 7). Deformation texures are commonly exhibited by the metamorphosed orebodies. Braunite and hollandite are stretched and elongated and show preferred dimensional orientation parallel to the banding, imparting a schistosity to the ore (Fig. 9). Translation twinning is developed in hollandite (Fig. 10) and hollandite-bixbyite-braunite ores show evidences of microfolding (Fig. l l). Micrographic intergrowth between rhodonite and jacobsite (Fig. 12), observed in gondite may be interpreted as due to oxidation of FeSi0,- rich rhodonite or pyroxmangite (see Table 4). Such

intergrowth has also been described from Dongri Buzurg, India, by Chaudhuri (1967). The transformation of bixbyite to braunite along crystallographic planes is characteristic in all grades of metamorphism (Fig. 13), and very often hematite is released from bixbyite during this transformation. The resultant braunite shows slightly different optical characters (cf. Roy, 1966) and is possibly lower in silica content1 than that formed directly by the

1. This variety may refer to the silica-poor braunite described by De Villiers anci Herbstein (1967). This braunite was separated from a sample containing only braunite and bixbyite and analysed. The Si03 content was determined as 3.89 per cent. The possibility of contamination of bixbyite in the sample is, however, not entirely ruled out.

FIG. 7, Braunite and vredenburgite (hausmannite lamellae etched black) exhibiting granoblastic texture in recrystallized ores. East of Dongri Buzurg. (x 125.)

FIG. 8. Hausmannite (lamellae) and jacobsite (host) intergrown to form vredenburgite. Hausmannite (grey) has been largely oxidized to cryptomelane (white). Netra. ( x 600.)

FIG. 9. Deformed braunite (dark grey) and hollandite (white, pitted) showing preferred dimensional orientation parallel to banding. Bharweli. ( x 125).

FIG. 10. Translation twins formed in hollandite due to deformation. Gowari Wadhona. Crossed nicols. ( x 250.)

235

S. Roy

FIG. 11. Microfolded bands of braunite and hollandite (coarse- grained). Bharweli. ( x 110.)

FIG. 12. Micrographic intergrowth between jacobsite (white) and rhodonite (black). Tirodi. ( x 300.)

FIG. 13. Braunite (darker grey) formed as a conversion product of bixbyite along the crystallographic directions. Black portions represent pits. Bharweli. ( x 400.)

reaction7 MnO,+3 Si0,=3 Mn203. MnSiO,+O, (Table4). Hausmannite sometimes envelops hollandite grains and also occupies the cleavage planes of the latter (Fig. 14).

The gondite generally shows xenoblastic to granoblastic texture. Banding with oxide ore or within its own components is common. Textural relationships indicate that the formation of spessartite and rhodonite overlapped and a few other manganiferous pyroxenes and amphiboles such as brown manganiferous pyroxene (Mn. aegirine- augite), tirodite and wiiicliite formed by metamorphic process from the original bulk composition, Prolific development of manganiferous silicates other than spessartite and rhodonite in gondite is, however, always related to alkali aiid Fe3+ metasomatism from later intrusivepegmatites (Roy, 1966; Roy and Mitra, 1964; Roy andPurkait, 1968).

FIG. 14. Hausmannite (dark grey) in the periphery and cleavage planes of hollandite (white). Subrounded grains of braunite (medium grey) are also present in hollandite. Gowari Wadhona. (% 300.)

Discussion The trends of transformation of different phases of manganese with rising temperature, as determined in the laboratory and observed in natural assemblages, are shown in Table 4. Higher oxides of manganese (E, ß, y, and 8-MnO,) when present exclusively, convert to E-Mn,O, (bixbyite) at temperatures between 500"-600" C, and further to hausmannite at 877" C in air (Faulring et ul., 1960; Hahn and Muan, 1960; Klingsberg and Roy, 1959; Okada, 1959~; Ukai et al., 1956). Okada (19596, 1960) has shown that todorokite [(Mn, Ca)Mn,O, . 2Hz0] and birnessite [S-MnO, : (Na, Ca)Mn,O1,2.8H,O] which, incidentally, are the main constituents of recent deep-sea manganese nodules, transform to hausmannite at 600"-700° C.

23 6


TABLE 4. Phase transformation of manganese minerals with rising temperature

1. h- MnOz(GRYPT0MELANE) íA) ß- MnO?(PYROLUSiTEi

-31c- MANGANOSITE - P'YROCHROITE (BI 1- (MnO) i MniOH)21

2.

3.

4

5

6

7

8

9

550°-60dC- BIXBY ITE- 8770- HAUSMANNITE - 900.c - HOLLANDITE -500'~- PSILOMELANE (c) I(Bo.H2012 Mn5 0101

iMn O21 W M n OOH)

(MnDl IMn CO3)

(Ba R806) - 9 5 0 ~ ~ -BIXBY IT€ - 565% - PYROLUSITE-375.C - MANGANITE (D)

MANGANOSITE -RHODOCHROSITE (E)

6- MnO;(NSUTITE) ô- MnOz(B1RNESSITE) (L-Mn;! O3 1

(A)7Mn02+5¡02~3Mn~O3~nSiO3tO2 ß- MnOz(PYR0LUSITE) + sioz (BRAUNITE)

6- MnOz(BIRNESSITE) 6- MnO2 INCUTITE) 2MnO2+2S102=2MnS103tC2

(RHODONITE) (C) 3Mn2 O3 Mn C103t6S102= 7Mn SI O3 i312 Oz

(BRAUNITE) (RHODONITE) BEMENTITE - BRAUNITE -RHODONITE

i(Mn.Mg, Fe16 Si4QOH) 181 (3Mn2 OaMnSiO3) (MnSiO3)

PYROXMANGITE 2[IMn,FelS10~3/2 Oz- MnFegO4+MnSiO3+ SIOZ (PYROX M A N G I TE (JACOBSITE) (RHODONITE)

d Mn02(CRYPTOMELANE) \ BIXBYITE - JACOBSITE A N D / O R VREDENBURGITE ß- Mn02iPYROLUSITE) IRON OXIDE

6-MnOz(NSUTITE) + OR 4- MnOzIBIRNESSITE) 60% Fe2031 Mn304 IN SOLID SOLUTION)

(Mn Fez 04) íMnFe204 WITH EXCESS

(RHODOCHROSIT E) (RHODONITE)

Mn CO3 + MnSiOJ = MnZCi04 + CO2 (RHODOCHROSITE) (RHODONITE) (TEPHROITE)

(A) MnzS104 + Ciop = 2 M n Si03 (8) 6Mn2 si04 + O2 (TEPHROITE) (RHODONITE) (TEPHROITE)

(LU v n Co3 + Ci O 2 = Mn C1O3 + CO2 (8) 7 M n C O 3 + S102= 3Mnz03.MnSIO~ ~ C ~ 2 M n C O ~ +Si0z=MnpCi0~~2CO~ (BRAUNITE) ( T E PH ROIT E )

= 6Mn '51.03 t 2 M n 3 0 4 (RHODONITE) HAUSMANNITE

Manganite (y-MnOOH) transforms to pyrolusite at 375" C which is again reduced to bixbyite at 565" C and further to hausmannite at 950" C (Table 4, lD, Das Gupta, 1965). Groutite (cr-MnOOH) transforms to pyrolusite at 300" C (Lima-de-Faria and Lopes-Vieira, 1964) and at higher temperatures, the usual transformation of pyrolusite to bixbyite and hausmannite follows. Pyrochroite [Mn(OH),] transforms to manganosite (MnO) at 32" C (Table 4, 1B; Klingsberg and Roy, 1959) which easily oxidizes to hausmannite. Psilomelane [(Ba,H,O),Mn,O,,] transforms to hollandite at about 550" C (Fleischer, 1960; Fleischer and Richmond, 1943; Wadsley, 1950) and hollandite transforms to hausmannite at 900" C (Table 4, lC, Fleischer, 1964).

Braunite and rhodonite are formed by reaction of higher oxides of manganese with silica in rising temperature depending on the oxygen fugacity (Table 4, 2A, and 2B; Huebner, 1967; Roy, 1968). The low temperature manganese silicate, bementite [(Mn,Mg,Fe),Si,(O,OH),,] con- verts to braunite first and further to rhodonite at higher temperatures (Ito, 1961). Higher oxides of manganese, admixed with iron, convert first to bixbyite with varying amounts of Fe,O, (according to temperature (Mason, 1944)) and then to jacobsite and/or vredenburgite with rising temperature.

In a metamorphosed oxidic orebody, braunite and bixbyite may form together or one in exclusion of the other,

depending upon bulk composition (availability of silica, iron, etc.), temperature and oxygen fugacity (Muan, 1959~). Braunite, once formed, remains stable with increasing temperature as the presence of silica in the structure has a strong stabilizing effect on Mn2+ (Muan, 1959~7, 19593). Huebner (1967), discussed the possibility of reaction between braunite and silica to give rhodonite (Table 4, 2C) and Pavlovitch (1931) showed that braunite (8.92 per cent Sioz), on heating to 1,400" C for five hours, converted to tephroite and an eutectic intergrowth of tephroite and hausmannite. From the study of natural deposits in India (quartz-albite-muscovite-chlorite to sillimanite-almandine- muscovite subfacies) and elsewhere, however, no such evidence of transformation of braunite or its reaction with other phases has been obtained. Evidences of formation of braunite as a conversion product of bixbyite (Fig. 13) are, on the other hand, ubiquitous in Indian ore deposits.

The formation of hausmannite in metamorphosed oxidic (carbonate-free) ores, is a function of high temperature and concomitant reduction. Its presence, in the absence of jacobsite, indicates a bulk composition low in iron (cf. Jothvad deposit, Table 3). In highly metamorphosed ores with iron-rich bulk composition, hausmannite is associated with jacobsite and/or vredenburgite. From a bulk composition rich in manganese carbonate, however, hausmannite may form either by dissociation of rhodochrosite at high temperature (Table 4, 1E; Smith Prospect,

23 7

S. Roy

Buckeye deposit, United States; Merid mine, Brazil; Tam- worth District, Australia, etc.) or by controlled supergene oxidation of the latter (Bricker, 1965).

The formation of jacobsite is dependent on the iron content in the original bulk composition, temperature and oxygen fugacity. It has been reported from low grade metamorphic deposits (cf. Kajlidongri, Table 2) and even from unmetamorphosed colloidal ores (Dongri Buzurg; Roy, 1959; Morocco; Vincienne, 1956; Saint Béat, Pyrénées; Perseil, 1966). The manganese content of these jacobsites is invariablylow. The mineral is also very widely distributed in the high temperature assemblage of almandine-amphibolite facies in India. The composition of jacobsite varies as a, function of temperature and, with rising temperature, more and more Mn,O, goes into solid solution in Fe,O, (Mason, 1943; Muan and Somiya, 1962; Van Hook and Keith, 1958). The micrographic intergrowth of jacobsite and rhodonite recorded by Chaudhuri (1967) from Dongri Buzurg and by the author from Tirodi (Fig. 12), formed by oxidation of FeSi0,-rich rhodonite or pyroxmangite. In carbonate deposits where rhodochrosite has a high content of FeCO, in solid solution, the carbonate stability field is reduced and Fe-Mn spinel or jacobsite may form as a decarbonation product in high temperature (Huebner, 1967; cf. Tamworth District, Australia).

The presence of vredenburgite ensures a high temperature and testifies to an original bulk composition rich in manganese and iron and a sufficiently slow cooling for the exsolution intergrowth to form, postdating deformation. The stability field of vredenburgite has been studied by Mason(1943),VanHookandKeith(1958), andYun(1958).

Rhodochrosite, under laboratory conditions in air, dissociates to M n O + COz at 610-635" C and the resultant manganosite rapidly oxidizes to hausmannite at 670- 735" C (Cuthbert and Rowland, 1947; Kulp et al., 1949, 1950, 1951). When rhodochrosite is associated with silica, the reaction at high temperature produces rhodonite (Table 4, 6A). If the rhodochrosite is present in larger proportion with respect to silica (Mn: Si ratio very high), tephroite may be produced (Table 4, 6C, 7). Rliodo- chrosite may also react with silica with rising temperature and high oxygen fugacity to produce braunite (Table 4, 6B; Huebner, 1967).

The manganese silicates formed at elevated temperatures, either by reaction between higher oxides of manganese and silica or between manganese carbonate and silica, may take part in further reactions with a rise of temperature. Thus rhodonite may be oxidized to give rise to hausmannite and silica (Table 4, 9; Huebner, 1967) and tephroite may react with silica to produce rhodonite (Table 4, 8A), or it may be oxidized to form rhodonite and hausmannite (Table 4, 8B).

Studies of metamorphosed manganese formations of India and the data accumulated on other deposits suggest that thenature of the bulk composition of syngenetic manganiferous sediments largely determines the later metamorphic reactions and the products. The unmetamorphosed syngenetic oxide sediments contain manganese largely in the

tetravalent state and only a minor part in divalent or trivalent state (Hokkaido, Japan: Hariya, 1961; Cuba: Hewett, 1966; Irnini Tasdremt, Morocco: Bouladon and Jouravsky, 1956; Timna Dome, Israel: Bentor, 1956; Deep Sea nodules: Mero, 1965; Roy, 1969). The manganese carbonate sediments contain divalent manganese and CO, (Usinsk and Labinsk deposits, U.S.S.R.: Varentsov, 1964; Urukut deposit, Hungary: Nemeth and Grasselly, 1966, etc.) and sometimes the oxidic sediments transgress to carbonates (Nikopol, Bol'she Tokmaksk and Chiatura deposits, U.S.S.R.: Varentsov, 1964). The syngenetic sediments sometimes contain low-temperature manganiferous silicates, represented by layer silicates containing mainly divalent manganese and these phases are generally hydrated.

As shown earlier, the metamorphosed manganese formations of India are characterized by oxidic ores sharply interbanded with manganese silicate rocks (gondite) in most places. The mineralogy of the manganese ores and gondite indicates that during metamorphism the individual bands act as separate entities, and the transformation of phases and reactions with rising temperature proceeded according to the physico-chemical conditions restricted to individual bands. The total absence of rhodochrosite and tephroite is characteristic, indicating that the original sediments were devoid of manganese carbonates (Roy, 1966, 1968; Roy and Purkait, 1968), and were, by and large, made up of higher oxides of manganese with variable admixtures of silica, alumina, iron, etc. With the onset of metamorphism, braunite formed as the only metamorphic mineral in quartz-albite-muscovite-chlorite subfacies (Shivarajpur) and its formation, by reaction of higher oxides and silica, in preference to rhodonite, was possibly facilitated by high oxygen fugacity. Rhodonite and spessartite, together with other manganese silicates, appeared in quartz-albite-epidote-biotite subfacies at Kajlidongri, and were restricted to the gondite bands, separated from braunite-rich oxidic ore bands. In the quartz-albite-epidote-almandine subfacies at Dongri Buzurg-Kurmura area they are similarly developed, whereas at the same grade of metamorphism, they are absent in the Bharweli-Ultwa deposits. Thus, the formation of rhodonite by the reaction of higher oxides and silica is not controlled entirely by temperature, and its appearance in gondite bands, in preference to braunite, is a function of oxygen fugacity. Spessartite-rich garnet usually appears in low grade metamorphic rocks (quartz-albite- muscovite-chlorite subfacies: Chimer, 1960; Hutton, 1957; Woodland, 1939). Fermor (1909) suggested its formation by reaction between kaolinite, silica and manganese oxide, and Huebner (1967) considered that the decomposition of manganiferous sheet silicates with rising temperature gives rise to the formation of spessartite. However, braunite forms earlier than spessartite during metamorphism (Hueb- ner, 1967). Bixbyite appeared first in the quartz-albite- epidote-biotite subfacies (Kajlidongri) and the relict colloform texture shown by the mineral together with braunite strongly indicates their derivation from originally colloforin higher oxides by transformation with rising temperature. Hollandite, accompanying braunite and bixbyite at Kajli-

23 8

Genetic studies on the Precambrian manganese formations of India with particular reference to thr effects of metamorphism

dongri, apparently formed by transformation of barium- bearing psilomelane at higher temperature. Low-manganese jacobsite, approaching the magnetite composition field, appeared first at Kajlidongri and its manganese content gradually increased at higher grades of metamorphism.

At Chikla-Sitasaongi (staurolite-almandine and kyanite-almandine subfacies), Tirodi-Sitapatore, Netra-Ram- rama, Gowari Wadhona (sillimanite-almandine-muscovite subfacies) and Kodur-Carbham (granulite facies), India, jacobsite (Mn,O,-rich), hausmannite and vredenburgite appeared and continued in stable assemblage with braunite, bixbyite and hollandite. The appearance and stability of these phases at higher temperature has been explained by Muan and Somiya (1962) and by the previous discussions on the mineralogenetic trend of pure and admixed (with iron) higher oxides of manganese with rising temperature, summarized in Table 4. The thermally metamorphosed (pyroxene-hornfels facies) deposits of Jothvad exhibit a braunite-bixbyite-hollandite-hausmannite assemblage in the ores, the absence of jacobsite and vredenburgite being explained by low content of iron in the sediments. The gondite contains spessartite and rhodonite in staurolite- almandine, kyanite-almandine and sillimanite-alniandine- muscovite subfacies, which shows that the phases are stable at higher temperature. Spessartite and rhodonite are often accompanied by other manganiferous silicates. Iron-rich rhodonite os pyroxmangite, however, oxidized to give rise to iron-poor rhodonite and jacobsite in micrographic intergrowth (Fig. 12) at Tirodi in the sillimanite- almandine-muscovite subfacies .

Manganese silicate-carbonate protore, developed in many parts of the world, contains principally rhodochrosite, rhodonite, spessartite, tephroite and quartz (Tables 2 and 3). The mineral assemblages indicate that in the original bulk composition, manganese was present as carbonate (rhodochrosite) with admixtures of silica, alumina, etc. Rhodonite and tephroite were formed by decarbonation of rhodochrosite and reaction with silica in rising temperature (Table 4, 6A, 6C, 7, SA, 8B), though no direct relationship of the formation of rhodonite and tephroite to increasing grade of metamorphism has been established. Tephroite has been described from the SmithPros- pect, United States (Huebiier, 1967) which has been metamorphosed only lo zeolite or lower greenschist facies, while

the mineral is absent (in similar bulk composition) in the same metamorphic grade at Merionethshire, United King- dom, and Western Otago, New Zealand. It is also characteristically absent in the carbonate-free manganese silicate deposits including the rocks of sillimanite-almandine-muscovite subfacies of India. Thus, temperature and total pressure cannot explain all the phase assemblages and intensive parameters of @Oz and p.0, are most important (cf. Huebner, 1967). Most of the silicate-carbonate deposits are devoid of manganese oxide phases (excepting the supergene oxidation products). In rare cases, hausmannite, h ausmannite-brauni te, liausmanni te-jacobsite and hausmannite-braunite-manganosite-pyrochroite are developed (Tables 2 and 3), apparently formed by decarbonation of rhodochrosite and ferroan rhodochrosite and by the reaction of rhodochrosite and silica at high oxygen fugacity.

Thus, the syngenetic manganese ore deposits of India were originally laid down as higher oxides and show, on metamorphism, a reaction sequence characterized by the progressive reduction of manganese. The syngenetic oxides contained only minor amounts of silica and iron in the orebands, and these, with rising temperature, reacted to form braunite-bixbyite-hollandite-jacobsite-hausmannite- vredenburgite assemblages. The interbanded gondites were originally represented by manganese oxide, considerable amounts of silica, clay, etc., in an admixed sediment. The oxidic ores and gondite bands acted as separate entities during metamorphism. The latter, by reaction at rising temperature, gave rise to spessartite-rhodonite-quartz assemblages, often containing manganiferous pyroxenes, amphiboles and micas. Metamorphism of theIndian manganese formations took place at oxygen fugacities higher than those in the carbonatic rhodochrosite-rhodonite- tephroite deposits (cf. Smith Prospect, Western Otago, Merid mine, etc.). Huebner (1967) pointed out the great disparity between the oxygen fugacity of certain manganese deposits compared to the foz prevailing during metamorphism of non-manganiferous country rocks. H e suggested that oxygen is not free to pass between the manganese deposit and the surrounding system and the country rock is not responsible for imposing a high fo, on the manganese deposits, rather the manganese minerals themselves buffer or internally define the high oxygen fugacity.

Résumé

Étude génétique des furmatiuits de manganèse précambrien stratigraphiquement dans les mêmes plissements que les en Inde avec références particulières aux effets du métamor- quartzites de schiste pblitique, les calco-silicates et les plzisme (S. Roy) marbres du groupe précambrien de Sausas (province de

Madhya-Pradesh et Maharashtra), du groupe Aravalli (ré- Les masses minéralisées d'oxyde syngénétique de manga- gion de Kajlidongri, Madhya-Pradesh), du groupe Cham- nèse de l'Inde, intimement imbriquées avec des roches paner (Shivarapour, Gujarat), du groupe de Gangour (Sun- silicatées de manganèse (groupées ensemble en formatioii dargarh District, Orissa) et du groupe de Khondalite de manganèse), se rencontrent en lits continus ordonnés (district de Srikakulam, Andhra-Pradesh). Les formations

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syngénétiques de manganèse ont éré métamorphosées régio- nalement en schistes verts et en faciès d'almandine-amphibolite. L'intensité du métamorphisme subi par la formation de manganèse a été déterminée par le degré précis atteint par le méta-sédiment pélitique englobé. A Jothvad, dans l'État de Gujarat (groupe de Champaner) la formation de manganèse a subi un métamorphisme de contact dans les faciès hornblende-hornfels et pyroxène-hornfels .

Il a été possible de suivre les changements minéralo- giques et structuraux dans le minerai d'oxyde de manganèse aux différents degrés du métamorphisme régional ou de contact. La braunite est le premier minéral métamorphique qui apparaisse et cela est la seule phase stable dans les dépôts des sous-faciès quartz-albite-muscovite-chlorite. La braunite a un domaine de stabilité très étendu et continue jusqu'au sous-faciès sillimanite-almandine-muscovite. La bixbyite (avec une faible teneur de Fe,O,) apparaît d'abord dans le sous-faciès quartz-albite-épidote-biotite en association avec le sous-faciès épidote-almandine. L'hausmannite, la jacobsite et la vredenburgite (présentant un développe- ment enchevêtré & deux phases) apparaissent dans le sous- faciès staurolite-almandine continuant le sous-faciès sillimanite-almandine-muscovite, en association avec la braunite et la bixbyite (avec une haute concentration de Fe,O,).

Les formations minérales d'oxyde de manganèse pré- sentent des changements dans leur contexture en fonction du degré de métamorphisme. Les minerais syngénétiques non métamorphosés sont constitués d'oxydss supérieurs à grains fins et présentent parfois une structure rubanée et des marques évidentes de dépôt à Yair libre. La survivance de la contexture et la structure rubanée encroûtée sont mises en évidence par la présence de bixbyite et de braunite dans les gisements de minerai du sous-faciès quartz-albite- épidote-biotite. Apparemment, là où le métamorphisme est peu développé, les oxydes sédimentaires supérieurs ont été recristallisés pour former de la braunite et de la bixbyite bien que la contexture originale n'ait pas disparu. La structure rubanée sédimentaire originale est conservée dans le minerai. Des contextures particulières dues à la déformation abondent en particulier dans les minerais métamorphosés à faible teneur. D e telles contextures se traduisent par un allongement et une orientation dimensionnelle préféren- tielle dela braunite et de la hollandite, par des déformations et des déplacements des macles de la hollandite, par des

micro-plissements des filons de minerai, etc. Dans les minerais hautement métamorphosés, la texture granoblas- tique à grains grossiers est caractéristique. On trouve dans les minerais hautement métamorphosés une texture d'exsolution mise en évidence par la jacobsite et l'hausmannite (vredenburgite). On a observé la transformation de la bixbyite en braunite suivant les plans cristallographiques à tous les stades du métamorphisme. Une formation inters- titielle, filons ou couches minces d'un minerai dans l'autre, se rencontre à l'occasion et peut résulter des phéno- mènes de croissance et de tension superficielle entre des particules qui sont devenues simultanément plus grossières.

Les roches de silicate de manganèse métamorphosées régionalement (gondite) sont caractérisées par l'association spessartite-rhodoiiite-quartz-apatite-tirodite-pyroxène man- ganifère brun (aerigine-augite)-braunite (et autres oxydes inférieurs). La blanfordite, la winchite, la juddíte, etc., se sont développées au contact de la gondite et de la pegmatite. Les kodurites (roches silicatées de manganèse résultant d'un contact métamorphiquej diffèrent des gondites par la composition des grenats (spessartite-andratitej et par I'abondance des feldspaths potassiques. La paragenèse des phases minérales dans les roches de silicate de manganèse, en relation avec le degré de métamorphisme, a été établie.

La tendance minéralogénétique qu'on observe aux différents degrés de métamorphisme régional dans les gisements de minerai oxydiques et dans les roches silicatées de manganèse a été rattachée aux résultats déduits de l'étude de l'équilibre des phases. Xi a été établi que les associations minérales des minerais d'oxyde de manganèse aux différents stades du mktamorphisme proviennent de la réduction progressive des oxydes plus riches à l'origine, oxydes qu'on rencontre dans les systèmes ferrugineux analogues et au cours du métamorphisme d'associations de minéraux comportant de la rhodochrosite [cf. Brésil). Les associations minérales dans les roches silicatées de manga- nèse métamorphosées avec absence de tephroite et de rhodochrosite, indiquent aussi que les carbonates étaient origi- nairement absents dans la composition d'ensemble et que les oxydes élevés, silicates, silice, alumine, etc., réagirent par réduction progressive pour donner naissance aux associations de spessartite-rhodonite-quartz-braunite avec plus ou moins de pyroxènes manganifères, d'amphiboles et de micas,

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STRACZEK, J. A. et al. 1956. Manganese ore deposits of Madhya Pradesh, India. XX Znt. geol. Congr., Mexico. Symposium on Manganese, t. 4, p. 63-96.

UKAI, Y.; MAYEDA, T. 1956. Mineralogical study on manganese dioxide minerals. Mem. Coll. Sci., Kyoto, Series B, vol. 23, p, 203-22.

VAN HOOK, J. J.; KEITH, M. L. 1958. The system Fe,O,- Mn,O,. Amer. Min., vol. 43, p. 69-83.

VARENTSOV, 1. M, 1964. Sedimentary manganese ores. Amster- dam, London and New York, Elsevier.

VINCIENNE, H. 1956. Observations géologiques sur quelques

gîtes marocains de manganèse syngénétique. XX Znt. geol. Congr., Mexico. Symposium on Manganese, t. 2, p. 249-68.

WADSLEY, A. D. 1950. Synthesis of some hydrated manganese minerals. Amer. Min., vol. 35, p. 485-99. WATANABE, T. 1959. The minerals of the Noda-Tamagawa mine Iwate Prefecture, Japan. I. Notes on geology and paragenesis of minerals. Miner. J., vol. 2, p. 408-21. WESTERVELD, J. 1961. The manganese vein on Mount Brand- nuten, Botnedal, S. Norway. Norg. geol. Unders., vol. 213,

WOODLAND, A. W. 1939. The petrography and petrology of the lower Cambrian manganese ore of West Merionethshire. Quart. J. geol. Soc., Lond., t. XCV, no. 1.

YUN, I. 1958. Experimental studies of magnetic and crystallographic characters of Fe-bearing manganese oxides. Mem. Coll. Sci. Kyoto, Series B, vol. 25, p. 125-37.

ZVÉREFF, R. 1953. Les gisements manganésifkres de l'Oural méri- dional (d'après A. Betekhtine), Ann. Min., Paris, 142" année, vol. 8, p. 44-5.

p. 203-23.

Discussion

J. VAN N. DORR. Are any of the many iron-formations in India contemporary with the manganese deposits of the Nagpur region?

S. ROY. Syngenetic manganese deposits of the Sausar, Aravalli and Khondalite groups are not associated with contemporary iron-formations. In the Gangpur Group the syngenetic manganese formation is overlain by iron ore formation. In Sandur, Mysore State, iron-formation is associated with manganese ore bodies, but it is not absolutely certain whether some of these manganese ore bodies are syngenetic, rather many of them are distinctly epigenetic.

Y. P. MELNIK. D o the temperatures of the phase transformation of manganese minerals (Table 4 and text) correspond to 1 bar pressure or to high pressures under metamorphism?

S. ROY. The temperature values cited in Table 4 are taken from results of laboratory phase-equilibrium studies carried out by different workers in air. So these absolute values will be different during metamorphism according to prevalent foz, feo,, etc.

Y. P. MELNIK. What agents caused the reduction of the manganese minerals: natural reducing compounds in rocks -carbon, etc.-or the temperature rise only?

S. ROY. Except for the Khondalite group, the mkganese ore bodies in India are not associated with graphitic sediments. It may, thus, be assumed that the reduction of sedimentary manganese oxide minerals during metamorphism was effected by rise in temperature only.

I. P. NOVOKHATSKY. What is the temperature range for formation of hausmannite-braunite and tephroite-spessartite?

S. ROY. The temperature of formation of hausmannite varies according to the nature of the original sediments, i.e. carbonate or oxide, as shown in Table 4. In oxidic sediments of India it is formed at fairly high temperatures in almandine-amphibolite facies. Once formed, braunite continues to be stable in higher grades. Spessartite forms in low grade (at the beginning of greenschist facies), and continues to persist in higher grades in gondite. The temperature of formation of tephroite depends very much on foz and feo., as pointed out in the paper.

I. P. NOVOKHATSKY. H o w do you estimate pressure values in the process of metamorphism?

S. ROY. The pressure-temperature values of the ore deposits were determined by referring to the enclosing pelitic meta- morphites which were metamorphosed to the same grade.

242

Precambrian ferruginous formations of the Aldan shield

I. D. Vorona, V. M. Kravchenko, V. A. Pervago and I. M. Frurnkin Yakutia Geological Service (U.S.S.R.)

The Aldan shield is an outcropping crystalline basement ridge on the southeastern margin of the Siberian platform and is a unique area in which to study laws governing iron-ore accumulation during Precambrian time. Old ferruginous formations within the shield consist of three genetic rock units, i .e. metamorphosed sedimentary fer- rosiliceous (jaspilitic), metamorphogenetic ferromagnesian and sedimentary oolite-hematitic (Table 1). Both the first and the second units are interspersed with the Precambrian sequence, but they attain great thickness and comprise deposits of commercial value within the only two series, i .e, Deoss-Leglierian (Middle Archaean) and Subganian (Upper Archaean). The third unit does not contain any commercial deposits and it is therefore not discussed below.

Description of ore-bearing formations

The Precambrian within the Archaean shield (Fig. 1) consists of Archaean formations forming a crystalline basement as well as of Proterozoic formations building up the old mantle of the platform. The following are geotectonic features occurring within the Archaean formations (from older to younger): Yengrian, Timptono- Jeltulian, Olecmian and Subganian. Physical measurements of geologic time also proved the Archaean (from 2,500 to 4,500 m.y.) age of the formations (Geochronology, 1968). Muscovite-pegmatites crosscutting the Subganian strata were dated at 2,500 m.y., while pyroxenes belonging to Yengrian metamorphic rocks were dated at 4,500 m.y. old.

Proterozoic strata consist of Udocan, Maimacan and Engilian deposits of Lower, Middle and Upper Proterozoic age respectively. As Maimacan and Engilian strata are not ferruginous they are not discussed below.

Deoss-Leglierian metamorphogenetic ferromagnesian formation bearing magnetic iron ore

Magnetic iron-ore deposits of this formation are not known in Precambrian strata found in other regions of the U.S.S.R. Similar Archaean skarn magnetite deposits have been found in the Central Sweden (Gejer, 1939; Gejer and Magnusson, 1955). Practically all commercial deposits of the types to be described herein are found in the central part of the Ungra-Timptonian graben-synclinorium (Fig. 1).

The iron-ore deposits are found on the flanks or occasionally in the core of secondary and younger synclines. They are associated with two producing horizons in the Deoss-Leglierian strata which are each 400-500 m thick and are interbedded with barren bands having thicknesses of 500-800 m (Pucharev, 1959). The producing horizons are interbedded with beds and lenses of dolomite marbles and calciphyres up to 200 m thick, diopside rocks, amphibolites and various gneisses (amphibole, pyroxene, biotite, garnet, graphite, etc.), as well as with magnesia skarn metasomatic replacements which contain magnetite occurrences.

Bed-like or flattened-lens ore bodies lie conformably in the country rocks. The ore bodies strike over 270- 4,200 m and are 11-140 m thick.

Serpentinized forsterite and diopside-magnetite varieties are predominant (about 75 per cent); clinohumite and sahlite-scapolite-magnetite ores are less commun (about 20 per cent), whereas hypersthene-, hornblende- and salite-feldspar-magnetites are rare. Sulphides, i.e. pyrrhotite and, more rarely, pyrite and chalcopyrite, are common in all of the rock varieties. The ores have massive banded-lenticular and banded structure. The ores have medium-grained structure. The mean iron content in rich varieties of ore amounts to 45-50 per cent, in disseminated ores to 35 per cent, and that of sulphur amounts to 1.5- 2.5 per cent.


I. D. Vorona, V. M. Kravchenko, V. A. fervago and I. M. Frumkin

TABLE 1. Precambrian ferruginous formations of the Aldan shield

Geologic Absolute Rock Geotectonic Regional Conjugate Genetic type Type of Minerai Iron age age complex position of the metamorphism magmatic and ferruginous composition content and

(m.y.) or series ore-bearing strata facies formation composition formation and and texture size of of the ore- deposit type of the deposits

Lower 2,000- Udocan Protoplatform- Nil Proterozoic 2,400 troughs

Lower over Subrranian Palaeoaulacogenes' From green- - Proterozoic- 2,500 Upper Archaean

Upper 2,750 Olecmian Archaean 3,100

Middle 2,750 Deoss- Archaean Leglierian

series

4,500 Lower (3,800- Yengriaii Archaean 5,400?)

- schistose to amphibolitic

Geosyncline Amphibolitic

Superposed From Graben- amphibolitic synclinorium to granulitic (pretoaulacogene?)

Protogeosyncline Granulitic

1. Aulacogene: narrow sub-longitudinal trench-like synclinorium structure.

Nil

bearing formation

Terrigenous: quartzose sandstones, aleurolites, argillites

Gabbrodiorite Terripenous- igneous; quartzites, aluminiferous schists biotite and amphibole gneisses; ortho- amphibolites, meta- porphyrites

Ophiolite Igneous- sedimentary; biotite, amphibole gneisses, schists, amphibolites

Gabbrodiorite Igneous- sedimentary; dolomitic marbles, calciphyres, basic crystalline schists, metadiabases; amphibole, pyroxene and other gneisses

Igneous- chemogenic; quartzites, gneisses of high alumina content and basic ortho-schists

Ophiolite Sedimentary; dolomitic marbles; calciphyres, pyroxene and amphibole gneisses

iron ores

Sedimentary Laminated Possible colite-hematite; colite- minor the Atugeian hematite deposits and the Hugdin ores deposits

Ferro-siliceous Banded Major (jaspilitic); the ferruginous deposits of Ymalic and the quartzites poor but Neliukin (magnetite, easily deposits grünerite- beneficiated

magnetite, ores biotite- magnetite, etc.)

Ferro-siliceous; Lenticular- Minor the Olecmian banded occurrences occurrence ferruginous of poor but

quartzites easily (magnetite, beneficiated amphibole- ores magnetite)

Ferromagnesian Massive, Major (metasomatically lenticular deposits of altered); the banded, rich ores Tayojiioye and magnetite Desovian ores deposits containing

forsterite, diopside, scapolite, serpentine, phlogopite

Ferro-siliceous; the Yagindya, Holednikan deposits

Ferromagnesian (metasomatically altered); the Emeljac deposit

Lenticular- banded ferruginous quartzites (amphibole- magnetite, pyroxene magnetite) Massive, mottled lenticular, banded magnetite ores containing forsterite and hypersthene

Iron ore occurrences, minor deposits of poor but easily beneficiated ores Minor deposits of rich ores

Magnetite ore deposits associated with magnesium-ferruginous formation are explored on the surface and down to the depths of 400-1,000 m. The deposits vary from minor ones estimated at 25-45 million tons to major ones of 300- 1,500 million tons. The Tayoshchnoe, Pionerskoye and Desobskoye deposits, are among the largest ones located in

Leglierskaya and Dios-Khatiminskaya synclines. The potential reserves of this type of deposit (Pervago, 1966) are estimated at about 3 billion tons, including 983.8 million tons of commercial reserves.

244

FIG. 1. Geologic-tectonic scheme of the Charo-Aldan iron province. 1. Cenozoic sediments in superimposed troughs; 2. Mesozoic-Middle Proterozoic sediments of the platform cover; 3. Proterozoic sediments of the Bajkal folded zone; 4. Lower Proterozoic, the Udocan complex in platform depressions; S. Upper Archaean, the Subganian complex in palaeoaulacogenes; 6. Upper Archaean, the Stanovoj complex in the outer structural-facies zone; 7. Upper Archaean, the Olecmian complex in structural-facies zones, (a) outer, (b) internal; 8. Middle Archaean, Deoss-Leglierian series in a protoaulacogene; 9. Middle

Metamorphic sedimentary ferro-siliceous formation in the Subganian series

Deposits of poor iron ores (ferruginous quartzites) of the Subganian series are isolated from the deposits described above. They are concentrated on the western margin of the Aldan shield, in some narrow sub-longitudinal trench- like synclinorium structures.

The ore-bearing Subganian complex consists of ferruginous quartzite beds 10-40 m thick, in places 100-200 m, at the top of the sequence. These are lying among biotite gneisses occasionally interbedded with pyroxene-amphibole schists, amphiboles and quartzites. Ferruginous quartzites are commonly associated with a certain division of the sequence; less commonly these are interbedded with barren rocks up to 200 m thick. Ferruginous quartzites form simple and complex beds and extend for several kilometres, but usually do not exceed 10 km. These are clearly banded and, less commonly, plicated.

Archaean, the Timptono-Jeltulian complex in structural-facies zones, (a) outer, (b) internal; 10. Lower Archaean, the Yengrian complex in structural-facies zones, (a) outer (the Yengrian series), (b) internal (the Kurultian series).

Deposits of iron ores : 11. Magnesia-iron-formation, (a) large deposits, (b) minor deposits; 12. Iron-chert formation, (a) potentially large deposits, (b) minor deposits; 13. Sedimentary oolite-hematite formation; 14. Deep faults-structural sutures; 15. Faults.

The main minerals of the ferruginous quartzites are magnetite, quartz and amphiboles; plagioclase, garnet and biotite are less common. Clinopyroxene and fibrolite are rare. Zircon, orthite and titanite are accessories. Magnetite occurs in practically monomineral bands and is interspersed, to some extent, in quartz grains as small inclusions. Mag- netite lying in the subsurface turns to martite. Amphiboles are represented by two types, actinolite-ferroactinolite (40-59 per cent iron-bearing) and cunimingtonite-grunerite (48-74 per cent grunerite molecule). The two varieties occur both together and separately.

Table 2 shows the chemical composition of main mineralogical types of ferruginous quartzites.

These characteristics classify the ferruginous quartzites as poor but easily beneficiated ores. Rare occurrences of rich metamorphic magnetite ores which contain 64 per cent iron are associated with ferruginous quartzites (col. 5, Table 2).

Iron-ore deposits ascribed to the Subganian ferro- siliceous formation are poorly studied and almost unexplored. From the aeromagnetic surveys, potential iron ore

245

I. D. Vorona, V. M. Kravchenko, V. A. Pervago and I. M . Frumkin

TABLE 2. Chemical composition of the Subganian ferruginous quartzites (%)

Compounds 1 2 3 4 5

Loss on ignition Total H20+ Fe total Fe soluble

47.96 0.2 2 .O3 35.72 9.71 0.08 2.68 1.33 0.1 0.08 0.13 0.04 nil

100.04

32.76 26.3

-

44.62 0.27 2.06 24.63 21.78 O .40 2.72 1 .o 0.05 0.13 - - 2.01 99.67 nil 34.66 29.5

43.28 0.22 0.08 36.14 16.54 0.04 1.33 0.62 0.13

traces 0.14 0.02 1.84

100.38 0.09 38.15 37.0

48.4 0.13 0.21 49.14 0.79 0.03 0.015 0.63 0.29 traces 0.045

0.82 99.8 0.13 40.62 34.6

-

5.21 0.14 1.94 60.55 29.66 0.07 3.34 0.35

0.29 0.103 0.005 0.85

100.508

64.14

-

- -

1. Hornblende-grünerite-magnetite ferruginous quartzite (Charsk deposit); 2. Grünerite-magnetite ferruginous quartzite (Charsk deposit); 3. Magnetite ferruginous quartzite (Tarikhan deposit); 4. Martite ferruginous quartzite (Tarikhan deposit); 5. Rich quartz-hornblende- magnetite ore, thin-banded (Olecmian occurrence).

reserves within the Subganian strata (down to 500 m) amount to about 5-7 billion tons.

Thus, the Precambrian Charsk-Aldan iron ore province is found within the central and western parts of the Aldan shield. The potential iron ore reserves in the province amount to 8-10 billion tons. Commercial deposits are associated with two Archaean formations, a metamorphogenetic magnesio-ferruginous formation and a metamorphosed sedimentary ferro-siliceous formation. Areas of their distribution are separated, thus forming two iron ore subprovinces, namely Aldan and Charsk (Kravchenko and Vorona, 1968). The two iron-ore formations occur repeatedly in the Precambrian sequence. Each formation becomes commercial only at one stratigraphic time, i.e. Middle and Upper Archaean.

Linear-oriented synclinorium structures of the aulacogene types developed on the Archaean protoplatforms played a decisive role in forming and distributing com- niercial iron-ore formations. Though iron-ore formations are sometimes found in geosynclinal structures, these never acquire commercial value.

A close positional connection of the iron-ore formations with the Archaean pre-orogenic formations of the basic igneous rocks was revealed which allows their relationships to be traced when studying the problem of the primary concentration of iron in iron ores.

Résumé

Formations feuvifères du Précambrien inférieur du bouclier d’dldan (I. D. Vorona, V. M. Kravchenko, V. A. Pervago, I. M. Frumkin)

Les formations ferrifères du bouclier d’Aldan se caracté- risent par leur vaste étendue, par une grande diversité géné- tique et par l’énorme période (au moins deux milliards d’années) qui a présidé à leur constitution.

Par rapport aux autres régions du monde, c’est sur le

bouclier d‘Aldan qu’apparaît au mieux l’histoire géologique précambrienne du globe, ce qui offre d‘excellentes conditions concrètes pour l’étude des lois qui ont régi le processus de formation des gisements de fer précambriens.

Cette étude fournit des renseignements de base sur les particularités géologiques, génétiques et minéralogiques des formations ferrifères et sur les degrés de concentration du minerai de fer dans le bouclier d‘Aldan.


FROLOVA, N. V. 1951. Ob usloviyach osadkonakopleniya v arheyskoy ere [On the conditions of accumulation of deposits in the Archaean]. Tr. Irlciitskogo gas. univ., vol. 5, ser. geol., vip. 2. Moscow, Gosgeolizdat.

FRUMKIN, I. M. 1967. Structurno-litologichesckiy method carti- rovaniya dokembriyskich obrazovaniy i rezultaty ego prime- neniya na Aldanskom schite [Structural-lithological method of mapping of Precambrian formations and the results of its utilization on the Aldan Shield]. Problemy izucheniya geologii dolcembriya, Moscow, Nauka. -_ . 1970. Napravlennost geologicheskogo razvitiya zemnoy kory Aldanskogo schita v arheiskoe vremya [Trend of geo-

logical development of the earth crust of the Aldan Shield in the Archaean time]. Tectonica Sibiri., vol. III, Moscow, Nauka.

GEJER, P. 1939. The paragenesis of ludwigite in Swedish iron ores. Geol. Fören. Stoclth. Förh., vol. 416, p. 61.

-; MAGNUSSON, N. H. 1955. Zheleznie rudy Swecii [Iron ores of Sweden]. Zhelezorudnie mestorozlzdeniya mira, vol. 1, IL.

Geochronologia dolcembria Sibirdtoy platformy i ee sclai~tchatogo obramlenia [Geochronology of Precambrian of the Siberia platform and its plicative framing]. Leningrad, Nauka, 1968.

KITSUL, V. I.; LASEBNICK, K. A. 1966. Geologia i petrographia docembriysckich kristallitchesckich obrazovaniy rayona sliania

24U

Precambrian ferruginous formations of the Aldan shield

Aldana i Ungry (k probleme ‘Ungrinsckogo clina’) [Geology and petrography of Precambrian crystalline formations of the region of the confluence of Aldan and Ungra (on the problem of ‘Ungra Wedge’)]. Geologia i petrologia docembria Aldans- ckogo schita, Moscow, Nauka.

KRAVCHENKO, B. M.; VORONA,I.D. 1968.Tcharsko-Aldanskaya. Zhelezorudnaya provincia [Tcharsk-Aldan iron ore province]. Materialy po geologii i poleznym iscopaemytn Yalrutsclcoy ASSR. Yak. knign. izd., vip. 18. Yak.

LEYTES, A. M.; MURATOV, M. V.; PHEDOROVSKY, V. S. 1970. Paleoavlacogeny i ich mesto v razvitii drevnih platform [Palaeoaulacogenes and their place in the development of ancient platforms]. DAN SSSR, vol. 191, no. 6.

MARACKUSCHEV, A. A. 1958. Petrologia Taezhnogo Zhelezo- rudnogo mestorozhdeniya v archee Aldansckogo schita [Petro- logy of Taezhniy iron ore deposit in the Archaean of the Aldan Shield]. Tr. DVF AN SSSR, ser. geol., vol. 5, Magadan.

NUZNOV, S. V.; YARMOLUCK, V. A. 1968. Novie dannye po stratigraphii dokembria na primere Aldanskogo schita [New data on Precambrian stratigraphy on the example of the Aldan Shield]. Sov. geologia, no. 5.

PERVAGO, V. A. 1966. Aldanskaya Zhelezorudnaya provincia [Aldan iron ore province]. Moscow, Nedra.

PUCHAREV, A. I. 1959. O geologii i osobennostyah lockalizatsii orudeneniya Yuzno-Yakutsckich Zhelezorudnich mestorogdeniy [On the geology and particularities of the localization of ores in the South-Yakutian iron ore deposits]. Geologia rudnich mestorogdeniy, no. 1.

SCI-IADYNIN, L. I. 1958. O genezise Yuzno-Yakutsckych Zhelezo- rudnyh mestorozhdeniy [On the genesis of the South-Yakutian iron ore deposits]. Izv. AN SSSR, ser. geol., no. 1.

SERDYUTCHENKO, D. P. 1960. i dr., Zheleznie rudy Yuznoy Yakutii [Iron ores of South Yakutia]. Moscow, Academy of Sciences.

Discussion

A. M. GOODWIN. What radiometric techniques were used to determine the absolute ages of the Archaean rocks?

I. D. VORONA. Mainly the K-Ar method for amphibolites, micas and pyroxenes; to a lesser extent, the U-Pb method.

247

On the issue of genesis and metamorphism of ferromanganese formations in Kazakhstan

V. M. Shtsherbak, A. S. Kryukov and Z. T. Tilepov The Institute of Geological Sciences of the Academy of Sciences of the Kazakh S.S.R.

In Kazakhstan the ferruginous and ferromanganese formations are found in most of the geological provinces and epochs (Fig. 1). Recently, apart from the well-known Early Proterozoic ferruginous quartzites in Karsakpay Kara Tau and Famensk, and ferromanganesian and siliceous-carbonaceous formations in the Atasuy region, new ferruginous and ferromanganesian formations and ore deposits of different epochs have been distinguished in the Bet-Pak- Dala steppe, Uspensk Synclinorium, Chingiz Anticlin- orium, the Altai and in Turgai.

In the Bet-Pak-Dala steppe of Kazakhstan within the Chu anticlinorium the Precambrian formations of volcanic ferruginous chert and jaspilite have been distinguished. These have subjected to various stages of metamorphism. The anticlinorium includes Proterozoic, Cambrian, Ordo- vician, Devonian, Carboniferous, Cretaceous and Palaeo- gene formations. Proterozoic deposits with ferruginous formations are porphyroids, slate, quartzite and ferruginous quartzite, stretching north-west for 50 km in a belt 15 km wide. Cambrian deposits are composed of sandstone, metamorphic schists, limestone, jasper quartzite and porphyrites. The Ordovician is represented by arkose and micaceous sandstone and platy flints; the Devonian is expressed by effusions of acid composition, while Carboniferous deposits are mainly sandstone, limestone and conglomerates. There are basic and ultrabasic intrusions of Upper Ordovician or Cambrian age in the north-eastern part of the anticlinorium. In most of the region all these rocks are over-

lain by friable sediments of the Mesozoic and the Cenozoic. The volcanic ferruginous chert occurs in the Protero-

zoic deposits in the north-western part of the anticlinorium. It is associated with the recently discovered large ore deposit at Zhuantobe (Gvardeiskoe) (Fig. 2), where the section of the formation has been most thoroughly investigated. The bedrock of the studied part of the section is composed of grey and light grey porphyroids (on quartz- porphyry) with intercalations of green quartz schists. It is overlain by 50 m of ferruginous quartzite with lenses of porphyroid and green and grey quartz-sericite slates. The next layer is also composed of ferruginous quartzites, but here they are black and brown, 55 m thick, with lenses of porphyroids. The total thickness of the Proterozoic ore- bearing deposits is approximately 700-750 m.

At Zhuantobe more than twenty lenticular ore bodies between 5-50 m thick and 50-3,500 m long have been discovered. The ores in the deposit are laminated ferruginous quartzites, with hematite laminae 0.1-3 mm thick intercalated with quartzite-hematite. Associated minerals are magnetite (10-20 per cent) and martite. Hematite occurs as lathes (0.01 x 0.05 mm) and plates (0.003 x 0.008 mm), forming a thick net. Crystals of magnetite are dispersed in the hematite mass as porphyroblastic grains often with an octahedral shape and dimensions 0.01-0.5 mm. Quartz forms isometric grains with dimensions 0.007-0.01 m m . The chemical composition of the ferruginous quartzites is given in Table 1.

TABLE 1. Chemical composition of ferruginous quartzites at Zhuantobe

Sample Fe No. (tota,) SiOs Alzo3 Fe,O, Fe0 M n O MgO C a 0 KzO Na,O P20, SO, H,O- irznition 'Ota1

1. 56.45 16.45 0.85 80.28 0.36 0.27 0.10 0.10 0.36 0.01 0.03 0.57 99.38 2. 49.75 27.45 - 69.29 1.65 0.14 0.10 0.10 0.28 0.03 0.62 99.68 3. 59.24 14.65 0.25 81.85 1.22 0.17 0.10 0.10 0.45 0.14 0.01 0.54 99.48 4. 39.75 39.07 1.63 56.52 0.83 0.02 0.04 0.75 0.20 0.20 0.39 100.24 5. 23.32 65.70 0.95 32.34 0.88 0.04 0.07 0.30 0.10 0.20 0.01 100.71 6. 46.22 29.73 0.97 67.14 0.83 0.08 0.04 0.25 0.10 0.20 0.01 99.34 7. 37.95 43.73 1.47 54.21 1.39 0.04 0.11 0.20 0.10 0.20 0.06 101.33

Unesco, 1973. Genesis of Precambrian iron and »fangonese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 249

V. M. Shtsherbak, A. S. Kryukov and Z. T. Tilepov

FIG. 1, Regions of development of ferruginous and ferruginous- siliceous formations in Kazakhstan. 1. ferruginous-siliceous formations of the Precambrian; 2. ferruginous-siliceous forma-

The ore-bearing rock mass of Zhuantobe is strongly folded with dips of 65-85' (Fig. 2). Crumpled layers often occur.

The appearance of boudinage structures in ore-free and low-ferrous quartzites, with sections of 0.5 m , in ferruginous quartzites is evidence of intensive dynamic reworking of rock mass. Regional metamorphism is only of the greenschist facies, as is judged by intensive chloritization, seri- citization, presence of stilpnomelane in ores, little recrystallization of fine-grained hematite and its transition to magnetite, as well as preservation of rnicrolamination of quartz-hematite ores. No metamorphic differentiation has been observed.

The ferruginous chert formation occurs in the Protero-

[ Y J 4

tions of the early Palaeozoic; 3. siliceous-carbonate ferruginous- cherty formations of the middle Palaeozoic (Fammenian); 4. the iron-ore deposits.

zoic deposits in the central and south-eastern parts of the Chu aiiticlinorium. The formation includes the epidote- biotite-chlorite-quartz and amphibole slates, ferruginous quartzites, dolomitized limestone and marble. The Temir group of quartz-hematite-magnetite ores also belongs to it. This formation differs from the formation of the first type by the absence of volcanics and a minor iron content. Wide occurrence of quartz-chlorite-epidote, quartz-biotite-epidote and amphibole slates in these rocks, together with the transition of fine-grained quartz-hematite ores into medium- and coarse-grained quartz-hematite-magnetite ores places this formation in the amphibole facies of metamorphism.

In the southern part of the anticlinoriuin, microcline and biotite gneiss, gneissose granite, quartz-andalusite

250


Section A - Ei

FIG. 2. Schematic geological map of Zhuantobe deposit. 1. friable depositions of the Mesozoic and the Cenozoic; 2, quartz-chlorite and quartz-sericite schists; 3. ore-free quartzites;

slates and low-ferrous magnetite quartzites belonging to granulite facies are included among the Proterozoic formations. The low ferrous content in the rocks of the formation is evidently conditioned by the high metamorphous differentiation and probably by the remoteness from the volcanic centre.

In the north-eastern part of the Bet-Pak-Dala steppe, along the eastern part of Zhailmy synclinal limb for 50 km, the Ordovician deposits appear as quartz conglomerates, polymict, sandstones, aleurolites, chlorite sericitic schists, quartzites and porphyrite. In the eastern part they are interrupted by the large Upper Palaeozoic granitoid intrusions. The overlying part of the rock mass, datable

and known as the Kosagaly formation, interbeds the persistent 300 m thick horizon of quartzites and jaspers with strata and lenses of ferrous and ferrous-manganese ores. These form the Tuyak-Kosagaly group of deposits. The ores are intercalated with quartzites, jaspers and amphibole-epidote-chlorite slates. In the south andesite and basalt porphyrites are also intercalated in the sequence. The ore-bearing rock is of rhythmic fine-grained composition and includes hematite and magnetite. Near Kosagaly Hill it contains up to 13 per cent manganese. The presence of volcanics and jaspers in the ore-bearing siliceous horizon is evidence for the volcanic sedimentary origin of the formation.

At the contact zones of late Palaeozoic granitoid intrusives, the ore-bearing rock mass appears to be of hornfels

4. porphyroids; 5. ferruginous quartzites; 6. dislocations with a break in continuity.

texture. Some distance from the intrusives, along the crush zones and zones of ore jointing, the enclosing rocks are also metasomatically altered. In addition to contact metamorphism, another ore recrystallization took place as the result of which jasper quartzites, volcanites and cherts have been locally changed to new metasomatites of epidote- actinolite-chlorite, sometimes with small amounts of pyroxene and garnet, as well as massive magnetite ores with relicts of jasper quartzites, heavily modified volcanics and thin laminated magnetite-hematite ores. Metasomatic processes contributed to the formation of magnetite deposits in beds five to ten times thicker than those of the primary sedimentary magnetite-hematite ores.

Ferruginous-cherty-carbonate-volcanic formations of the Fammenian stage are widely spread in superimposed troughs within the limits of the Uspenk synclinorium and Chinghiz anticlinorium. By analogy with the Atasuy ferruginous-manganese deposits, these formations are regarded as having volcanic sedimentary origin. In the same synclinal fold, apart from the Atasuy deposits, there are volcanic sedimentary iron-ore deposits with a high intensity of metamorphism, mainly that of contact and contact-metasomatic character, and with a considerable differentiation of ore.

Some researchers consider the magnetite ores and associated skarns of the Kentyube deposit to be the result of the metasomatic reworking of volcanic sedimentary ferruginous rocks of the Fammenian stage.

251

V. M. Shtsherbak, A. S. Kryukov and 2. T. Tilepov

The iron-ore deposits at Tortkul and Kirghizia in the East Karagaly ore region also serve as examples of formations of intensive metamorphic differentiation following local metasomatic phenomena.

In the formation of the Tortkul deposit (Fig. 3) the sandstone, aleurolite, aleuropelite, limestone, andesite porphyrite and tuffs of the Fammenian participate. In the

middle part of the rock mass jasper and carbonate-siliceous sinters with magnetite-hematite mineralization occur. The Fammenian formations are conformably overlain by conglomerates and arkose sandstone.

All these deposits form the brachy-syncline with submeridional shearing induced by the intrusion of Upper Palaeozoic gabbro-diorites and granodiorites in the north-

\

FIG. 3. Geological structural map of the Tortkul deposit. 1. recent deposits; 2. tufogene conglomerates, arkose sandstone, limestone; 3. tufite, cherty aleurolite, tufogene sandstone, ore; 12. hematite-magnetite ore. limestone; 4. sandstone, aleurolite, aleuropelite, limestone;

5. porphyrite and tuffs; 6. diorite porphyrite; 7. granodiorite; 8. diorite; 9. gabbro-diorite; 10. skarn; 11. magnetite-martite

252


west. In the south-eastern part of the brachy-syncline, the carbonate-siliceous sinters and jaspers over a distance of 1.4 km include beds and lenses of thin-banded magnetite- hematite ores, 50-400 m long and 5-30 m thick (Tortkul region I, Fig. 3). The ores have a rhythmic lamination structure. Fine laminae of hematite and magnetite-hematite (0.5-3 .O mm) with carbonate-cherts aleurite alternate with finer laminae of carbonate-siliceous sinters and aleurolites . In the richest ores of this type the average content of iron is 50per cent, sulphur 1.2per cent and phosphorus 0.02 percent.

As can be judged from the morphology of the deposit, the composition and structure of ores as well as their association with jaspers and volcanites, the finely laminated formations are of volcanic sedimentary origin.

In the northern and southern parts of the brachy- syncline there are also stratiform ore deposits associated with volcanic carbonate-chert formations (ore regions Tortkul II and III, Fig. 3), but there are also ore bodies of different types such as intersecting veins, stockworks and others. Magnetite ores are mainly associated with pyroxene-garnet skarns and albitized rocks, which appeared as an aftermath of the metasomatic phenomena in the zone of contact with the Upper Palaeozoic intrusion.

In the skarn-magnetite laminated deposits relicts of

thin-banded hematite-magnetite ores often occur. Some- times breccias are observed, with fragments of thin-banded hematite-magnetite ores and cementskarns or metasomatic magnetite.

All the above proves that magnetite ores of Tortkul II and III regions, related to the horizon of the carbonate- ferruginous chert formations, are the regenerated analogues of the primary sedimentary (volcanic sedimentary) magnetite-hematite ores of the Tortkul I region. In the ore- bearing rock mass, near the contact of the intrusion, hornfels has been formed and recrystallization of the ores has taken place. Then the metasomatic metamorphism produced albitization and skarning of the intrusions and the ore- bearing rocks, as well as further recrystallization and redeposition of ores.

The above examples, together with data from other provinces and deposits, show that the formation of volcanic- sedimentary, ferruginous and ferruginous-manganese rocks occurred in Kazakhstan during the Precambrian and early to middle Palaeozoic. The compact deposits of rich and easily dressed magnetite ores (Kosagaly, Tuyak, Tortkul, Kentyube deposits and others), were formed at the expense of poor sedimentary, mainly hematite, concentrations by metamorphic differentiation and local metasomatism.

Résumé

Formation et métamorphisme des roches ferrugineuses de rlivesses époques dans les psovinces du Kazakhstan (V. M. Shtsherbak, A. S. Kryukov, Z. T. Tilepov)

L'anticlinal de Chu, la région de minerais de Tuyak- Kosagaly et le synclinorium d'Uspensk permettent de mettre en lumière certaines particularités de la formation des roches ferrugineuses et des métamorphismes de diffé- rentes époques dans le Kazakhstan.

1. Dans l'anticlinal de Chu, les formations précam- briennes de silex ferrugineux volcanique et de jaspilite ont distinctement subi divers stades de métamorphisme. La formation du premier type se rencontre dans la partie nord- ouest de l'anticlinal. Elle est représentée par des porphy- roïdes, des quartzites ferrugineux et des ardoises. A cette formation est rattaché le vaste gisement de minerai de magnétite-hématite de Zhuantobe. Ici les volcanites s'inter- calent avec des quartzites et des minerais de fer qui mettent en évidence une origine sédimentaire volcanique pour ce dernier. Si l'on en juge par une chloritisatioii assez intense, par la présence de minerais de stilpnomélane, par la recristallisation partielle (jusqu'à 20 %) d'hématite à grains fins et sa transition en magnétite, ainsi que par la préservation de la microlamination des minerais de quartz-hématite, le métamorphisme régional des roches ferrugineuses de cette section de l'anticlinorium atteint le faciès des schistes verts. La formation des silex ferrugineux domine dans les parties centrales et sud-est de l'anticlinorium. Les ardoises mi-

grantes, les quartzites ferrugineux et les marbres dolimi- tisés qui encaissent toutes les manifestations de minerai de Temir peuvent être considérés comme appartenant à des faciès d'amphibole en raison de la présence de quartz- chlorite-épidote, de quartz-biotite-épidote, de biotite-chlorite-quartz et d'ardoises amphiboles et de minerais à grains moyens et ñns de quartz hématite-magnétite. Cette formation est différente de la formation du premier type par l'in- tensité du métamorphisme, l'absence de volcanites et sa moindre teneur en fer. Ces quartzites ferrugineux semblent avoir leur origine dans un volcanisme ancien.

2. Le groupe de gisements de quartz-hématite-magné- tite de Tuyak-Kosagaly a ses limites dans la partie centrale des séries de Kosagaly (O&, constituée de grès polimicte, de quartzites, de quartzites à jaspe, de silex et de volcanites modifiées. La présence de volcanites et de jaspes dans les roches à minerai, conjointement avec des minerais de quartz-hématite-magnétite à grains fins, est un autre argument en faveur de l'origine volcanique sédimentaire des dépôts de Tuyak-Kosagaly (formation de silex ferrugineux volcaniques). La majorité des roches à minerai est pure- ment schisteuse. Au niveau des quartzites, les structures de boudinage se distinguent tandis qu'on trouve des grès, volcanites et silex qui ont donné naissance à un dévelop- pement intense de chlorite, actinolite, épidote, séricite, calcite, etc. La composition pélitique des silex et des grès indique une recristallisation d'ensemble. Dans les minerais de quartz-hématite-magnétite, on a mis en évidence une

253

V. M. Shtsherbak, A. S. Kryukov and 2. T. Tilepov

recristallisation faible et une augmentation de la dimension des grains de magnétite d‘un ordre de grandeur. Tout cela indique un degré peu avancé de métamorphisme régional dans la masse rocheuse à minerai jusqu’au faciès de schiste vert et peut-être au premier stade du faciès d‘amphibole. Dans ce processus, aucune différenciation importante des éléments qui constituent le minerai ne s’est produite. Aux zones de contact des intrusions granitoïdes du Paléozoïque récent, la roche à minerai présente une texture de hornfels tandis qu’à quelque distance des intrusions, le long des zones d’écrasement et de celles de jointement des minerais, la roche encaissante change aussi de façon métasomatique. Parallèlement avec le métamorphisme de contact, une autre recristallisation du minerai s’est produite, et, comme résul- tat de phénomènes métasomatiques qui se sont produits localement aux endroits où se trouvaient des quartzites à jaspe, des volcanites et des silex, de nouvelles métasomatites d’épidote-actinolite-chlorite sont apparues quelquefois avec de petites quantités de grenat, ainsi que des minerais de magnétite massifs avec des restes de quartzite à jaspe, des volcanites profondément modifiées et des minerais de ma- gnétite-hématite en fines lamelles. Le fer, entre-temps, s’est redéposé, provenant des parties basses des couches plon- geantes. Les processus métasomatiques ont contribué à la formation de dépôts de magnétite d’épaisseur 5 à 10 fois supérieure à celle des lits des minerais de magnétite-hématite sédimentaires primaires.

3. Les formations ferrugineuses (silex, carbonate) volcaniques de l’étage fammenien sont largement répandues dans les dépressions superposées dans les limites du syncli-

norium d‘Uspensk et de l’anticlinorium de Chinghiz. L‘origine sédimentaire volcanique de ces formations a été démon- trée par de nombreux chercheurs. Sauf pour des dépôts d’Atasuy de la dépression de Zhailmyn dont le métamor- phisme des roches a à peine atteint le stade du schiste vert, on considère que ces structures contiennent des formations de carbonate, du fer siliceux, sédimentaire et volcanique, avec un haut degré de métamorphisme, essentiellement de contact, avec une certaine différenciation métamorphique. Les minerais de magnétite et les skarns du dépôt de Keny- tuybe qui les accompagnent résultent d‘une reprise méta- somatique des roches ferrugineuses sédimentaires volcaniques de l’étage fammenien. Les gisements Tortkul et Kirghibia peuvent aussi servir d’exemples de formations après une différenciation métamorphique intensive faisant suite à des phénomènes métasomatiques locaux. Au détri- ment des dépôts d’hématite-magnétite en couches minces, les minerais de magnétite apparaissent ici en association avec les skarns.

4. Les exemples mentionnés ci-dessus indiquent que dans les limites du Précambrien du Kazakhstan, le Paléo- zoïque ancien et moyen a été caractérisé à certaines époques par des formations de roches ferrugineuses sédimentaires surtout d‘origine volcanique. Ces dernières ont eu à subir différents stades de métamorphisme. La différenciation métasomatique des matières qui ont constitué le minerai a été surtout intense au cours de phénomènes métasoma- tiques locaux d‘où ont résulté des dépôts compacts de minerais de magnétite ayant une valeur commerciale (dépôts de Kosagaly, Tuyak, Kenytuybe, Tortkul et autres).

254

Genesis of high-grade secondary iron and manganese ores from iron- silicate and ferruginous formations and ores, metasomatic processes and processes of oxidation in them

Genese des minerais de fer et de manganese secondaires à haute teneur, à partir des formations de minerais de fer et de silicate de fer; processus métasomatiques et processus d'oxydation qui s'y rattachent

Iron-formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites A. F. Trendall Geological Survey of Western Australia, Perth

Introduction The iron-formations of the Precambrian Hamersley Group of Western Australia crop out over an area of about 85,000 km2 bounded by latitudes 21" and 23" S. and longi- tudes 116" and 122"E. Comparatively little was known of them until 1961, when systematic mapping of the Hamersley Range area by the Geological Survey of Western Australia began, under the supervision of W. N. MacLeod, concurrently with intensive private company exploration for iron ore. The main earlier geological records are those of Maitland (1909), Miles (1942), and Talbot (1920), who described the rocks collected by Talbot and others. These authors referred to the iron-formation either as part of the 'Nullagine Series', or as of 'Nullagine age'; the term 'Nullagine' has now been replaced by the stratigraphic nomenclature given in this paper.

MacLeod et ul. (1963) reported the results of the systematic mapping, and MacLeod (1966) later provided a comprehensive account of the geology of the area, with special emphasis on the economic iron deposits. Between 1964 and 1967 the Geological Survey carried out a study of the crocidolite deposits associated with the Hamersley Group iron-formations, involving attention to the general depositional environment of the iron-formations and their subsequent diagenetic and deformational history. Various aspects of the results of this study were reported concurrently by Ryan and Blockley (1965), Trendall (1965a, 19656, 1966u, 19666, 1966c, 1968, 19691, Blockley (1967, 1969) and Trendall and Blockley (1968); the complete results were finally compiled hyTrendal1 and Blockley (i 969).

Trendall and Blockley (1969) suggested that the Hamersley Group iron-formations were laid down as seasonally varved evaporitic chemical precipitates in a barred basin with a warm desert climate. M y purposes here are to summarize the evidence and arguments for this suggestion in a single short paper, to draw comparison between the Hamersley Group iron-formations, other varved evaporites and other iron-formations, and to exam- ine the geological consequences of these comparisons.

.

Hamersley Group iron-formations

REGIONAL SETTING

The Hamersley Group is one of three constituent groups of the Mount Bruce Supergroup; conformably below it lies the Fortescue Group, and above it, with some local discontinuity, the Wyloo Group. The present outcrop area of each of these three groups appears in Figure 1. They were laid down sequentially in an ovoid depositional basin (the Hamersley Basin) about 500 km long and 250 km wide, with a west-north-westerly elongation, which formed by the steady depression of an evenly eroded surface of Archaean granites, metasediments and metavolcanic rocks. These older rocks of the basin floor have now been re- exposed over a wide area around the present main outcrop area of the Mount Bruce Supergroup, and also within it in local inliers in anticlinal cores (Fig. 1).

The Fortescue Group has a maximum thickness of 4,350 m, and consists largely of basic lava, pyroclastic rocks, sandstone and shale. The Hamersley Group is about 2,500 m thick and is characterized by an abundance of iron-formation; details of its stratigraphy appear below. The Wyloo Group consists of mixed clastic sediments with thick local developments of dolomite and basalt; it reaches a thickness of 9,500 m.

STRATIGRAPHY

The lithostratigraphic subdivisions of the Hamersley Group, which have been formally named in accordance with the Australian Code of Stratigraphic Nomenclature (Geologi- cal Society of Australia, 1964), are shown in Figure 2. As emphasized by Trendall and Blockley (1969), these units are named principally as a convenience for field mapping. In practice the formal selection and naming of formations will depend upon accidents of physiographic development, the judgement of the mapping geologist, and the purpose and scale of the mapping; in consequence the named units

Unesco, 1973. Genesis of Precanibrian iron and manganese deposils. Proc. Kiev Syi?7p., 1970. (Earth sciences, 9.) 257

A. F. Trendall

LOCALITY MAP

AUSTRALIA

u

, :. ... :., ;. .. .i :. .. 2. '. :. :. 24'0C

REFERENCE

PHANEROZOIC ROCKS

u UNCONFORMITY

FORTESCUE GROLIP

UNCONFORMITY m' ARCHAEAN ROCKS

MAJOR FOLD AXES: -r -SYNCLINAL -+- -AN:ICLINAL -- STRIKE AND 'DIP OF

BEDDING

FIG. 1. Map showing the outcrop areas of the Fortescue Group, Hamersley Group and Wyloo Group, and their regional geological setting.

25 8

Iron-formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites

-

VERTICAL SCALE

300

i j; HAN GI

-

BFLGEEDA IRON FORMATION

v v

ILEY IP

W O O N G A R R A VOLCANICS

IROCKMAH IRffl FORMATION

V v v V

v v V

v v V

v v V

v v V

v v V

v v V

WITTENOOM DOLOMITE 50

IRON FORMATION

I I 516 4.4

VERTICAL SCALE 20-

S I I 5.2 I l

I I S 9 4.1

56 S S I I

I I s4 3.9

10.

O.

FIG. 2. Stratigraphic column of the Hamersley Group, and internal details of the Dales Gorge Member of the Brockman Iron Formation.

259

A. F. Trendall

may not reflect the real sedimentological significance of the rocks to which they are applied.

This situation has arisen in the Hamersley Group, and from the upper part of the Fortescue Group to the base of the Woongarra Volcanics the most significant large- scale regularity in the sequence of major sediment types seems to be an alternation between shale, or shale and dolomite, and iron-formation, with or without subordinate shale. The rearrangement of the formal units given in Table 1 illustrates this alternation:

TABLE 1.

No. Approximate thickness

in metres

8. Weeli Wolli Formation 7. Yandicoogina Shale Member of

Brockman Iron Formation 6. Joffre Member of Brockman Iron

Formation 5. Whaleback Shale Member of

Brockman Iron Formation 4. Dales Gorge Member of Brockman

Iron Formation + top of Mt McRae Shale

Formation + lower part of Mt McRae Shale

3. Wittenoom Dolomite + Mt Sylvia

2. Marra Mamba iron Formation 1. Jeerinah Formation (shale) of

Fortescne Group 1. Figure 2 shows c. 450 m. 2. Figure 2: 150 m.

185l

60

315

60

185

185

150-300

Each of the numbered iron-formation units in this sequence (2, 4, 6, 8) is lithologically distinct from the other three, as well as from the iron-formation of the Boolgeeda Iron Formation. These differences are expressed in the field by such secondary features as topographic response to erosion and colour of the weathered rock-face, which depend on subtle primary differences in composition, sequence and regularity of the mesobands (see 'Scales of Banding', below) and in minor but consistent differences in mineralogy.

AGE

Acid lavas of the Woongarra Volcanics have been dated by the Rb-Sr method to give an isochron of 2,100 m.y. (Leggo et al., 1965), later modified to 2,000m.y. (Compston and Arriens, 1968). An age of 2,200 m.y. is given by the Fortescue Group, from unsatisfactory material, and a minimum age for the Wyloo Group of 1,700 m.y. is given from intrusive granite.

No chronostratigraphic allocation of the Hamersley Group is here suggested.

STRUCTURE

In the northern part of the Hamersley Range area the Fortescue Group rests directly on Archaean rocks and dips off them southwards at only a few degrees. The rocks are perfectly undisturbed. Along the Hamersley Range itself the overlying Hamersley Group also has low dips, normally less than lo", and the range is formed by a broad open synclinorium. Southwards from the Hamersley Range the intensity of folding steadily increases. Thus, in the central part of the area, dips of 30-40" are common, and the amplitude and wave-length of folding are sufficiently great for the complete succession from Archaean to Wyloo Group to be re-exposed. In the southernmost part of the Hamersley Group outcrop, the folds are smaller and tighter, and locally overturned beds occur.

METAMORPHISM

The Hamersley Group has nowhere undergone metamorphism. Trendall (1966~) argued from the evidence of Hoering that no part of the Brockman Iron Forma- tion had reached a temperature above 160°C. Since then Grubb (1967) has successfully synthesized fibrous riebeckite, which had previously been regarded as a metamorphic mineral, at 35" C. Oxygen isotope analysis (Becker and Clayton, personal communication) also suggests that the Hamersley Group iron-formations have never reached higher temperatures than those of a normal geothermal gradient.

SCALES OF BANDING

All the iron-formations of the Hamersley Group, except the Boolgeeda Iron Formation, are banded. For various reasons most work in the Hamersley Range area has been carried out on the Dales Gorge Member of the Brockman Iron Formation, and the follawing description of the different scales of banding in that member serves as a basis for later comparison with the banding in other iron- formations of the Hamersley Group.

In the Dales Gorge Member there are three distinct scales of banding (Trendall, 1965b): macrobanding, mesobanding and microbanding.

Macrobanding (Fig. 2) is the name given to the major alternation between the two contrasted lithologies of the member-banded iron-formation, and mixed shale and chert-siderite iron-formation. Seventeen macrobands of banded iron-formation, numbered upwards from BIFO to BIF16 and ranging in thickness in the type section (Trendall and Blockley, 1968) from. 2.29 m (BIF8) to 15.05 m (BIFló), alternate with sixteen macrobands of the mixed lithology, numbered upwards from S1 to S16 and ranging in type-section thickness from 0.62 m (S2) to 5.47 m (S16), to give a total of thirty-three iiumbered macrobands,

260


Mesobanding (Fig. 3) is the name applied to the banding of the scale which is usually referred to in general use of the term ‘banded’ iron-formation. Within the BIF macrobands of the Dales Gorge Member, there is a conspicuously striped succession of internally consistent bands of different composition of average thickness 7.9 mm (860 measurements), and a range of 1 mm (by definition) to 66 mm. The major mesoband types in the Dales Gorge Member are chert (about 56 per cent of total thickness), chert-matrix (about 21 per cent) and magnetite (about 13 per cent). The mesobands forming the remaining 10 per cent of the thickness of the member consist mainly of stilpnomelane, carbonates, riebeckite, and minor miscel- laneous types. Chert-matrix is the name given to the fine- grained, iron-rich, homogeneous, structureless or finely laminated material, composed mainly of a variable mixture of quartz, carbonates (ankerite or siderite), stilpnomelane, hematite and magnetite, with which the chert mesobands alternate.

Most chert mesobands are internally microbanded. Microbanding (Fig. 3) is an alternation of regularly repetitive laminae of even thickness. The microbands are defined by a varying content of some iron mineral, most commonly by either hematite, ankerite, siderite or stilpnomelane, or by some combination of these. The usual thickness of

microbands (the combined thickness of one iron-poor and one iron-rich lamina) is within the range of 0.2-2.0 m m . They are clearly visible to the naked eye as a light and dark colour alternation in fresh samples of the iron-formation.

Of the five main stratigraphic units of iron-formation already referred to (2, 4, 6, 8 and the Boolgeeda Iron Formation) only the Dales Gorge Member (4) has clear macrobanding, although both the Joffre Member (6) and the Marra Mamba Iron Formation (2) have thin iiiterca- lated stilpnomelane-rich shales. The Boolgeeda Iron For- mation is exceptional in having only a faint lamination in otherwise massive black iron-formation, while the Weeli Wolli Formation (8) differs from the others in the presence of great thicknesses of thinly microbanded iron- formation with no mesoband development. The Joffre Member (6), the Dales Gorge Member (4), and the Marra Mamba Iron Formation (2) all exhibit mesobanding and microbanding, but there are distinctive differences in the proportions of the mesoband types.

LATERAL STRATIGRAPHIC CONTINUITY

With the conspicuous exception of one small area in the south-west, and to a lesser extent in the extreme east,

A FIG. 3. Mesobands and microbands in drillcore of banded iron- formation of the Dales Gorge Member. The darker mesobands are of chert-matrix or magnetite, and the paler mesobands are of chert, within which the fine microbands are clearly

B visible. Note that the left-hand and right-hand parts of B are photographs of the same stratigraphic level from two drillholes separated by a distance of kilometres.

261

A. F. Trendall

the constituent formations of the Hamersley Group persist throughout the outcrop area. Local erosion of the upper formations of the group in the northern and eastern parts of the outcrop means that the lateral continuity of the basal Marra Mamba Iron Formation is capable of more convincing demonstration than that of the uppermost Boolgeeda Iron Formation, whose outcrop area is now relatively small.

The most accurate information available on regional thickness variation of any stratigraphic unit is that for the Dales Gorge Member. With the one exception already mentioned, the thirty-three numbered macrobands of this member can be identified and measured throughout its outcrop area, providing not only thickness data for the member but also internal confirmation that a true depositional thickness unaffected by penecontemporaneous erosion is obtained. Trendall and Blockley (1969), report an extreme range in the thirty-one sections measured from 87 to 186 m, with a rate of change over the outcrop area of about 0.5 per cent per km. Although in detail the pattern is complex, there is a general thinning of the member away from the central part of the outcrop area, where the thickest sections were recorded.

In the Wittenoom Gorge area (22"52'S.; 117"08'E.), excellent core is available from the Dales Gorge Member to study correlation between a series of drillholes for crocidolite (blue asbestos) exploration. Mesoband correlation was quickly established between the most widely spaced drillholes (about 10 km) of this group, and was later extended, using both further drillholes (about 80 km) and surface exposures, over most of the available outcrop area, in some of the lower macrobands of the member selected for detailed study (Fig. 3; see also Trendall and Blockley, 1969).

Within some chert mesobands of the Dales Gorge Member selected for detailed study lateral microband correlation equal to that of the mesobands has also been established (Trendall and Blockley, 1968). A difficulty facing any microband correlation is the usual regularity of the microband sequence: some degree of irregularity is needed before correlation can be convincingly argued. The greatest lateral distance over which microband correlation has so far been achieved is 296 km, but since continuity has always been found with sufficiently detailed search, it is provisionally accepted that microbanding, like mesobanding, is continuous over the whole Hamersley Group outcrop, not only within the Dales Gorge Member, but equally in all other iron-formation units of the Hamersley Group in which microbanding occurs.

In summary, lateral stratigraphic continuity in the Hamersley Group is almost outcrop-wide, not only at the formation and member level, but also at the scale of macrobands, mesobands and microbands.

INTERPRETATION OF MICROBANDING

Some chert mesobands contain several hundred microbands which are believed to be continuous over the outcrop area

of the Hamersky Group. The three main features of microbanding so displayed (even spacing, constant repetition, wide lateral extent) are the points which both demand explanation and provide the main arguments for origin.

The first question to be answered is whether microbands are primary sedimentary features or secondary-possibly some form of Liesegang banding which developed during diagenesis. This latter possibility is difficult to disprove, but Trendall and Blockley (1969) have summarized the relevant arguments and, for the purposes of present discussion, their conclusion that microbands are probably primary is here accepted.

If they are primary, then the second question arises, of how such thin layers, successively siliceous and ferruginous, could have been laid down successively throughout such a vast area. It is virtually impossible to imagine any form of mechanical transport of material that could have achieved this, and the only reasonable conclusion is that the material is chemically precipitated; this is, of course, entirely consistent with both the chemical composition and the evident absence of normal clastic textures.

If the microbanded chert results from chemical precipitation, a third main question follows: what events controlled the continuous regularity of the microbanding? The alternate precipitation of thin and even laminae of such disparate composition in regular pulses suggests some alternation of the basin chemistry with a regularity only matched in the natural surface environment by the two astronomical rhythms controlled respectively by the earth's rotation and revolution-the day and the year.

In deciding which of these is more likely to have controlled microbanding, two lines of argument may be used. The mean microband thickness in 300 chert mesobands of the Dales Gorge Member was 0,65 mm (Trendall and Blockley, 1969). If it is assumed that the number of days in the year during deposition was 365 then an annual accumulation of 237 mm of material is indicated. In fact, it is likely that the rotation rate at that time was much higher (Wells, 1963), so that this figure is a minimum. Nevertheless, it is unacceptably fast as a likely sedimentation rate for any depositional basin, and in terms of water circulation would raise problems of even distribution and supply of material. The second argument is a consideration whether, by modern comparison, the alternation of day and night represent changes of sufficient physical intensity to radically affect the depositional chemistry of a body of water at least 85,000 kmz in extent. It seems unlikely, and both arguments lead to the conclusion that the year is a more likely primary control for the chemical precipitation of microbands: they are essentially non- glacial seasonal varves.

If microbands are varves, it may be calculated from chemical analysis of microbanded cherts that about 22.5 m g of iron were precipitated annually per square centimetre of the basin area. The fact that such estimates are independent of the microband thickness of the analysed cherts is consistent with the interpretation here proposed,

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INTERPRETATION OF MESOBANDING

If the microbands within a microbanded chert mesoband form a regular record of the passing years from the lower to the upper surface, then what events do these surfaces themselves represent, at the transitions from and into chert- matrix or magnetite? This question is answered here by reference to a type of laterally discontinuous or podded chert, which occurs in most of the Hamersley Group iron- formations.

In the description of mesobanding above, it was stated that mesobands have, relative to their thickness, extreme lateral continuity. This is not always so. At some levels strings of flat chert lenticles (often connected when their full areal extent is observable) appear to represent mesobands with local small-scale discontinuity. Such pods are as commonly microbanded internally as are normal chert mesobands, and the behaviour of microbands at the lateral termination of pods is illustrated diagrammatically in Figure 4. The microbands are not sharply truncated at the chert pod margins, but pass smoothly into the adjacent chert-matrix, where it is normally represented by a pervasive lamination lacking the distinctive twofold compositional disparity of the chert microbands. -- - ---- - -- -- -- A -

- - - zz- ------ - - ------ -- --

FIG. 4. Diagram showing the structural behaviour of microbands at the lateral termination of a chert pod.

It is deduced from the relationships in Figure 4 that the thickness of chert-matrix, t2, at Y, represents the compacted remains of chert similar to that of thickness t, which remains uncompacted at X. A closely similar situation has been similarly interpreted by Bryant and Koch (1969). The ratio t, / tz in Dales Gorge Member chert mesobands is commonly close to 7 : 1. It is found that the total amount of iron present in the lesser thickness of chert-matrix approximates to that in the equivalent greater thickness of chert. Trendall and Blockley (1 969) have argued in some detail the case for regarding chert-matrix as compacted chert or, alternatively, that chert and chert- matrix are the less strongly and more strongly compacted products of some unknown common parent.

The next step of the argument follows naturally. The chert-matrix at Y is petrographically identical with that immediately above and below, in which lamination is not demonstrably continuous with microbanding in laterally adjacent chert. Why, therefore; should a different

origin be assumed for it, or indeed for any chert-matrix? If it is assumed that the chert-matrix at Z is compacted chert, the history to be read from a regularly interbanded , sequence of chert-matrix and chert mesobands becomes apparent.

The preserved microbands of a typical laterally continuous chert mesoband are read as a sequence of recorded annual depositional events. The transition to chert-matrix at its upper surface is now interpreted as the start of a period when similarly microbanded sediments were laid down which differed in their response to compaction. This hypothesis, originally argued for the Dales Gorge Member by Trendall later found strong supporting evidence in a type of continuously microbanded iron-formation restricted to the Weeli Will0 Formation.

The nature of the control of compaction in the parent material has not so far been determined. It is assumed, however, that some very minor chemical or physical property critically controlled reaction to the pressure of continuing sedimentation, and that this property was controlled by rather regular changes in the environment of the depositional basin, to give the regular alternation of chert and chert-matrix mesobands (Fig. 3).

CYCLICITY OF MESOBAND SEQUENCE

In the Dales Gorge Member, the Joffre Member, and in the Weeli Wolli Formation, there is a small-scale (about 10- 20 cm) cyclicity of mesoband type. In general, groups of coarsely microbanded cherts alternate with groups of finely microbanded cherts, alternating with chert-matrix and magnetite in the normal way, but there are other characteristics of the cherts which help to differentiate the two cherts and to define the cycles. These are seen by Trendall and Blockley (1969) as the reflection of long-term cyclic environmental changes, and their expression in the Dales Gorge Member and Joffre Member have been called by them the Calamina cyclothem and Know cyclothem respectively.

The Hamersley Basin and its environment

INTRODUCTION

If the foregoing arguments are accepted, the deposition of the Hamersley Group iron-formations must be envisaged as involving the annual deposition of a thin basin-wide skin of precipitate with a clock-like regularity for up to a million virtually undisturbed years continuously. The impression is one of exceptional stability. With such a close relationship of precipitation to the year, it is difficult to escape the conclusion that the triggering mechanism of the- precipitation lies in direct seasonal effects on the basinal environment. The simplest such mechanism to envisage

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appears to be an annual increase of the silica and iron concentrations above their permissible solubilities by evaporation. Trendall and Blockley have built up a detailed reconstruction of the depositional basin of the iron-formations based on this central suggestion, and have shown that it is consistent with many other features of these rocks. Their arguments are not repeated in detail in this summary of their reconstruction, but the main types of evidence used are indicated for each aspect of the reconstruction.

INITIATION OF THE BASIN

The name Hamersley Basin is applied to the basin in which the Fortescue, Hamersley and Wyloo groups were successively deposited. The initial event in the life of the basin was the accumulation of the thick and largely volcanic Fortescue Group. The sedimentary rocks interstratified with the lavas of this group bear abundant evidence of shallow-water deposition, while the lavas themselves locally have pillow structure. During this period the whole surface of the accumulating material was probably almost flat, with local shallow lakes. N e restricted centres of this vulcanicity are apparent, and it may be assumed that the abundant dolerite dykes that transect the presently exposed areas of Archaean rocks represent the fissures along which the basalt sheets were comparatively quietly extruded.

SIZE, SHAPE AND STRUCTURE OF THE BASIN

The basin, during deposition of the Hamersley Group, is believed to have been ovoid in shape, with a long axis of about 500 km trending west-north-west and a shorter axis of about 250 km, to give an area of about 100,000 k m 2 . It is believed to have been enclosed on all but the north- western side, where there was at least a partial connexion with the open ocean.

The main argument for the shape of the basin is the isopach pattern of the Dales Gorge Member, which appears to be roughly followed by most other stratigraphic units of the basin. The limits of size are provided by an outward extrapolation of the isopachs, and by the stratigraphic impression of proximity to a basin margin along the eastern edge of the outcrop. It appears that the basin formed a rough ovoid enclosing on, and never extending far outside, the present outcrop. Closure of the basin is consistent with the isopach pattern, and is required by the suggestion that precipitation was triggered by evaporative concentration. A north-western oceanic connexion is argued partly by a local thickening of the Dales Gorge Member at that extremity of outcrop and partly by doubt concerning the possibility of maintaining a sufficiently delicate balance between evaporation and water intake (from where?-see below) to maintain a stable water level in the basin without benefit of control by an oceanic connexion.

DEPTH, CIRCULATION A N D IRON CONTENT O F W A T E R

From the perfect preservation of annual microbanding in what was probably a delicate gelatinous precipitate, it is argued that there was negligible bottom current, and no disturbance of the bottom by waves or storms. By modern comparison a minimum depth of 50-200 m is indicated (Kuenen, 1950). The validity of such a direct comparison is open to much discussion, however. Nothing definite is known of the likely turbulence of the atmosphere at that time, for example, and the possible presence of an algal raft (see below) would affect the validity of the argument.

A second argument for water depth runs as follows: There is negligible lateral stratigraphic or lithological variation throughout the present outcrop.

Therefore all parts probably had a similar environment and were laid down in water of the same order of depth.

But certain slump structures (omitted from descriptions above) indicate at least some bottom slope towards the centre of the basin, as would be expected if this were the area of most rapid depression.

Even if this were as little as 1 : 1,000 there would be a depth differential of 100 m between centre and edges.

Therefore, some figure above this is a minimum depth estimate for iron-formation deposition in the Hamersley Basin.

A depth estimate of about 200 m is accepted for the sake of continuing discussions. A concentration of 10-20 ppm of iron seems reasonable for the water of the basin; this is vitally dependent on the atmosphere of the time, a complex issue which is not discussed here in detail. With this concentration range, each square centimetre column of basin water would contain 200-400 m g of iron. The annual deposition of 22.5 g of this represents only 5-10 per cent of the total available iron. If circulation were confined to the upper 100 m of the basin, iron concentrations of less than 20 ppm in the incoming water would be adequate to replace the precipitated iron, and would require circulation speeds of less than 1 m/sec, which is consistent with the evident lack of disturbance of the precipitated microbands after deposition.

CONDITIONS ON THE SURROUNDING LAND

During the long stable periods of Hamersley Group iron- formation deposition virtually no clastic debris was transported into the basin. Two contrasted explanations are conventionally available to explain this; either there were no rivers because the surrounding area was desert, or the existing rivers were so sluggish as to be incapable of carrying a significant undissolved load.

The first of these hypotheses is preferred for three main reasons. First, the second hypothesis, although widely suggested in palaeogeographic interpretation, is inconsist- ent with the fact that all existing rivers sufliciently large and

264


mature to have low velocities have deltas built from transported debris. And secondly and thirdly the desert hypothesis seems to fit both with the high evaporation rate, independently suggested as a trigger for precipitation, and with the apparent regularity of annual climatic conditions suggested by the regularity of microbanding.

The maintenance of a desert climate through a least a million years suggests that the surrounding topography lacked mountainous areas tending to stimulate at least occasional storms, flash floods and consequent clastic intercalations in the iron-formations.

PALAEOLATITUDE

No good palaeomagnetic estimate of latitude is yet available for the Hamersley Group. Work so far (Irving and Green, 1958) is, frankly, preliminary, and based on too few data. If the arguments for desert climate, intensive evaporation and marked seasonal variation (to define the microbands) are accepted then alow latitude off the equator, perhaps at one of the tropics, is indicated.

VULCANICITY IN AND A R O U N D THE BASIN

It has already been noted that the Hamersley Basin started with intense volcanic activity, resulting in the present Fortescue Group. It can be seen also from Figure 2 that the Woongarra Volcanics, a stratigraphically concordant sequence of acid tuffs and lavas about 450 m thick, separate the Boolgeeda Iron Formation from the remaining iron- formations of the Hamersley Group. It is evident from these facts alone that there is a close general association between volcanic activity and the development of the basin.

There is further evidence of a closer association in time of volcanic activity and the iron-formations. Within the Mount McRae Shale, the Dales Gorge Member (in the 5 macrobands) and the Joffre Member, stilpnomelane bands 1-5 c m thick have clearly defined relicts of volcanic shards, and are ash-fall tuffs (Trendall and Blockley, 1969). Their presence was first noted by LaBerge (1966). Through- out Brockman Iron Formation time, at least, there was periodic explosive vulcanicity in the general vicinity of the basin.

The significance of this association between Hamersley Group iron-formations and vulcanicity is discussed further in the final section of this paper.

LIFE IN THE BASIN

Although Edge11 (1964), regarded some microbanded chert pods as silicified algal (stromatolitic) growths, this no longer seems likely. However, many of the Hamersley Group shales are highly carbonaceous, and LaBerge (1967) dis-

covered structured skeletal carbonaceous bodies from chert of the Dales Gorge Member. Trendall and Blockley (1969) later found the same bodies in Marra Mamba Iron For- mation chert. Algal stromatolites occur in the Carawine Dolomite (correlative with the Wittenoom Dolmite, Fig. 2) in the extreme eastern part of the outcrop, and stromatolites occur in both the Fortescue and Wyloo Groups.

There is thus abundant indication of life in the basin, though the nature and extent of this is uncertain. Trendall and Blockley (1969) have suggested the possible presence during iron-formation deposition of a floating algal raft. The hypothesis has a number of advantages, for example: (a) by acting as a quantitative biochemical buffer between some simple climatic factor, possibly annual insolation, and the basin chemistry an algal layer tolerant of the suggested range of iron concentrations may provide an equal annual precipitation of iron regardless of a relatively spasmodic supply; (b) the interposition of a biochemical process may overcome the difficulty of finding any climatic factor which is likely by itself to affect the basinal chemistry sufficiently to provide the chemical contrast between the parts of the microbands; (c) an algal raft may be a contributory cause of the apparent absence of any influence by waves or weather on the light and incoherent precipitate.

SINKING RATE A N D E N D OF THE BASIN

With a number of assumptions, discussed at length by Trendall and Blockley (1969), the average rate of sinking of the Hamersley Basin during Fortescue Group and Hamersley Group tirnewas between 16,000and 8,00Oyears/m. At the end of Hamersley Group time the basin was full in the sense (Dallmus, 1958) that the floor of the basin was planar, not geoidal. The Wyloo Group was subsequently laid down in a marginal trough which formed only along the southern edge of the basin.

Are Hamersley Group Iron Formations exceptional?

Judged by published descriptions of Precambrian iron- formations, those of the Hamersley Group appear to have some features which are uncommon; notable among these are the striking regularity, abundance, and strength of expression of microbanding in many chert mesobands (Fig. 3). Of the 20 per cent of chert mesobands in the Dales Gorge Member which are not clearly microhanded, many have a faint trace of microbanding, and it would be possible to select a sequence of chert mesobands in which the microbanding varied by insensible degrees from its normal development, as in Figure 3, to non-existence.

It may be postulated, from this, that all chert mesobands in Hamersley Group iron-formations were initially microbanded, but that in some it has been obliterated by

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diagenetic processes, If no chert mesobands in Hamersley Group iron-formations bore microbanding, there would certainly be much less reason to regard them as at all unusual or distinctive. It follows, too, that if microbanding proved to be similarly present in many other iron-formations, not only would their differences from those of the Hamersley Group again lessen, but there would arise at least the possibility that the genetic hypotheses argued here, and in prior publications, are rather widely valid.

In my own observation, microbanding is not as uncommon in other iron-formations as may appear from the literature. Attention has already been drawn (Trendall, 1968) to the equivalence of Hainersley Group microbands and both the finest laminae of the Transvaal System of South Africa (Cullen, 1963), and the 'second order laminae' in iron-formation of the Marquette Range noted by Tyler and Twenhofel (1952). In Figure 5 are illustrated further examples from iron-formations of different ages in South Africa, North America, India and Australia. The frequency with which examples of lamination at least geometrically comparable with microbanding may be discovered in many iron-formations when sought suggests that, for the purposes of further discussion, the implications may be explored of a supposition that many iron-formations have an evaporitic origin.

Other varved evaporities in the stratigraphic record

The two best known occurrences of varved evaporites in basins whose palaeogeography can be reconstructed as confidently as that of the Hamersley Basin are in the Permian rocks of Germany and of Texas-New Mexico. Varved evaporites have also been reported from, at least, the Devonian of Canada, the Permian of England (Stewart, 1963a), the Miocene of Sicily and the Pleistocene of Israel (Bentor, 1968), but none of these has so far been studied closely in relation to basin development,

The varves of the German Zechstein basin have been extensively recorded by Richter-Bernburg, most recently in a summary which gives references to earlier work (Richter- Bernburg, 1963). In a particular anhydrite band, about 2 m thick, there are about 1,200 varves 0.5-3 mm thick which are defined by thin bituminous layers in the anhydrite. Mostly the varves are rather evenly spaced (Richter- Bernburg, 1963), but distinctively spaced sequences permit lateral varve correlation over an east-west distance of 320 km, and of 350 km from north to south, to give a total area of about 100,000 km2 over which the varves can be correlated.

A general account of the Permian basins of West Texas

FIG. 5. Examples of microbanding in chert mesobands of for- B. Archaean iron-formation, Canada (Timagaini; about 43" N., mations other than those of the Hamersley Group, from different 77" W.). C. Transvaal System, South Africa (Derbi, about 28"S., continents and of various ages. The vertical bar with each letter 23' E.). D. Archaean iron-formation, Jndia (Noamundi; represents 1 mm. A. Archaean iron-forniation, Western Aus- about 22" N., 85" E.). E. Archaean iron-formation, Canada tralia (Mt Crawford; 28"35' S., 122"23' E.; G S W A No. 2/3598). (Kaministikwa; about 48" N., 89" W.).

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and southern New Mexico has been given by Hills (1968), who shows the depositional limits of the Castile evaporite deposits to be rather crudely ovoid, with an area of 17,500 km2. Within this area, Anderson and his associates (Anderson and Kirkland, 1966; Anderson, personal communication) have established that the millimetre-scale calcite-anhydrite couplets in the Castile anhydrite, which were first described as varves by Udden (1924), may be correlated with confidence throughout the depositional area.

, The varves of both these basins thus have their thickness, regularity, and lateral extent all comparable with the microbands of the Hamersley Basin. In all three respects these evaporitic varves are geometrically distinct from all other forms of sedimentary stratification.

Are iron-formations the Precambrian evaporites?

Although it has been argued that the Hamersley Group iron- formations originated by the annual accumulation of an iron-rich precipitate whose deposition was triggered by evaporation from a partially enclosed basin, the ultimate origin of the precipitated iron has not so far been discussed. Trendall and Blocltley (1969) gave revised statistics for an argument concerning this which was first put forward by Trendall (19654.

The argument was designed initially to negate the classical weathering hypothesis for the derivation of the iron in iron-formations, as proposed by Gruner (1922) and later supported by, for example, Sakamoto (1950), James (1951, 1954, 1966), Lepp and Goldich (1964) and Govett (1966). It is based on the problem of disposal of surplus material after chemical ‘processing’ (weathering) of enough Precambrian crust of average chemical composition to supply all the iron in the sediments of the Hamersley Group.

The relevant data are displayed in Table 2. In column 1 appear the total weights of metals and silicon in the Hamers- ley Group sediments, estimated from the assumed area of the basin, and the known proportions, average chemical compositions, and densities of the sediment types. In column 2 are shown the total weights of these constituents in the weight of average Precambrian crust needed to supply the controlling 9.4 x 1013 tons of iron in column 1. Column 3 shows the difference between columns 1 and 2. The figures show that a weathering hypothesis for the derivation of the iron creates more problems than it solves. If column 2 is recalculated in terms of total rock volume then 1,000,000 km3 of crust would have been required to produce the 120,000 k1n3 of the Hamersley Group. Where did all the iron-free surplus go?

If the weathering hypothesis solves no problems, Trendall and Blockley (1969) argued that some alternative

TABLE 2. Total original metal of Hamersley Group sediments, and comparison with metal content of average Precambrian crust (all figures in units of 1013 tons).

Precambrian Excess of crust for column 2 Hamersley over Group iron column 1

Hamersley Group

sediments

1 2 3 Si 7.7 86.2 78.5 Al 0.6 21.1 20.5 Fe 9.4 9.4 - M g 1.1 2.8 1.7 Ca 1.1 6.6 5.5 Na o. 1 6.2 6.1 K 0.5 8.1 7.6 Total 20.5 140.4 119.9

derivation of the Hamersley Group iron must be sought, and the constant vulcanicity associated in particular with the iron-formations and in general with the basin makes a volcanic source the most immediately attractive, the more so as the efflux of iron in association with vulcanicity is a demonstrable modern occurrence. However, there is another alternative if the similarity of iron-formations and later evaporites is pursued further. If the classical barred basin hypothesis for saline evaporite formation is followed, it may be that the comparatively low absolute concentrations of iron in the Hamersley Group depositional water suggested in prior discussion could easily have been produced by gentle continuous circulation of ocean water across a barrier separating the evaporative concentrate from the less ferruginous external body.

It is widely accepted as a statistical truth that banded iron-formations are, if not completely confined to, at least very much more abundant, in the Precambrian rocks (James, 1966 and further references therin) in spite of converse arguments (O’Rourke, 1961).

It is equally clear that saline evaporites of the type implied by normal usage of the term have a comparable restriction to the Phanerozoic (Kozary, Dunlap and Hum- phrey, 1968; Stewart, 1963b). Could it be that these two, hitherto unrelated, observations could be linked to each other via the varved evaporite hypothesis for Precambrian iron-formations outlined in this paper, and that the transition from ferruginous to saline evaporites near the end of Precambrian times marks an abruptly transitional period of oceanic composition of fundamental significance in the surface geochemical evolution of the earth?

Acknowledgements

This paper is published with the permission of the Director, Geological Survey of Western Australia.

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Résumé

Formations de feu du groupe de Hameusley, en Australie occidentale : exemples typiques d'évuporiíes précumbriennes en vurve (A. F. Trendall)

Le groupe précambrien de Hamersley affleure sur une surface de 85 O00 k m 2 environ dans la partie nord-ouest de l'Australie, entre 116 et 112"E et 21" et 23"s. Avec le groupe de Fortescue au-dessous et le groupe de Wyloo au-dessus, il fait partie du supergroupe du mont Bruce - succession volcanique-sédimentaire dont l'épaisseur locale maximale est d'environ 10 O00 mètres et qui s'est déposée dans un bassin ovoïde [le bassin de Hamersley) d'environ 500 km de long et 250 km de large, avec un prolongement dans la direction ouest-nord-ouest . Le bassin existait il y a 2 200 à 1 800 millions d'années. Le groupe de Hamersley a une épaisseur de quelque 2 500 mètres, dont 1 O00 mètres environ sont une formation de fer qui se présente en cinq unités stratigraphiques principales séparées par des filons-couches de schiste, dolomite, lave, tuf et dolérite. Une de ces cinq unités a été étudiée en détail : le Dales Gorge Member (180 mètres d'épaisseui;) dans la formation de fer de Brockman. On distingue à son intérieur trois échelles de formation de bandes : a) à grande échelle, des macrobandes, définies par 16 alternances de fines couches de schiste avec des formations de fer plus épaisses ; 6) à échelle moyenne, des mésobandes, définies par des couches alternées de 1 à 30 mm d'épaisseur de silex noir et d'une (( matrice de silex n riche en fer ; et c) à petite échelle, des microbandes, de 0,2 à 2 mm d'épaisseur, définies par les lignes d'un minéral contenant du fer, espacées réguliè- rement dans les mésobandes de silex noir. Dans l'expression formation de fer (( en bande D, le type de bande normalement considéré est celui des mésobandes. Les macrobandes, les mésobandes et les microbandes du Dales Gorge Member s'étendent généralement sur toute la surface d'affleurement. Nous soutenons, principalement pour des raisons de texture,

que : CI) les microbandes constituent la seule structure primaire préservée actuellement; h) les microbandes reflè- tent des alternances saisonnières de précipités colloïdaux originels riches et pauvres en fer (ce ne sont donc pas des varves non glaciales) ; c) les mésobandes exemptes de microbandes de matrice de silex noir sont le résultat du compac- tage de précipités également en varve, si bien que les méso- bandes sont une structure secondaire et diagénétique. Les conditions dans le bassin de Hamersley et aux alentours ont été exceptionnellement stables. La profondeur probable de l'eau pendant le dépôt de la formation de fer était de 50 à 250 mètres ; il y avait une circulation réduite vers un océan ouvert au nord-ouest, avec des courants internes de vitesse inférieure à 1 m/s et une teneur en fer de 10 à 20 ppm. La région environnante était plate, avec un climat déser- tique, un drainage négligeable, une évaporation annuelle de 3 mètres environ et de fortes variations saisonnières, impliquant une latitude quasi tropicale. Des phénomènes volcaniques intermittents étaient fréquents et ils ont apporté probablement une grande quantité de matières ; le bassin baissait à l'allure moyenne d'environ 1 mètre en 15 O00 ans. L'interprétation donnée s'applique spécifiquement aux formations de fer du groupe de Hamersley. La comparaison détaillée de diverses formations de fer révèle des différences importantes, qui peuvent être attribuées à une variété de facteurs. L'environnement le plus semblable à celui qui est supposé pour le bassin de Hamersley, d'après les don- nées géologiques d'autres régions, semble être celui des évaporites permiennes des bassins de Delaware et de Zechstein, et ce n'est pas par hasard que la géométrie stratigraphique de ces derniers est aussi très similaire ; cependant, la composition chimique du contenu est très différente, et la question reste posée de savoir si cette composition reflète quelque circonstance spéciale de l'environnement local de Hamersley ou résulte réellement et nécessairement de l'évolution géochimique de la terre.


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KOZARY, M . T.; DUNLAP, J. C.; HUMPHREY, W . E. 1968. Incidence of saline deposits in geologic time. Spec. Pap. geol. Soc. Amer., no. 88, p. 43-57.

p. 251-66.

vol. 49, p. 235-91.

KUENEN, P. H. 1950. Marine geology. New York, Wiley. LABERGE, G. L. 1966. Altered pyroclastic rocks in iron-formation in the Hamersley Range, Western Australia. Econ. Geol., vol. 61, p. 147-61.

-. 1967. Microfossils and Precambrian iron-formations. Bull. geol. Soc. Amer., vol. 78, p. 331-42.

LEGGO, P. J.; COMPSTON, W.; TRENDALL, A. F. 1965. Radio- metric ages of some Precambrian rocks from the Northcrest Division of Western Australia. J. geol. Soc. Arcst., vol. 12,

LEPP, H.; GOLDICH, S. S. 1964. Origin of the Precambrian iron- formations. Econ. Geol., vol. 59, p. 1025-60.

MCLEOD, W . N. 1966. The geology and iron deposits of the Hamersley Range area, Western Australia. Bull. W. Aust. geol. Surv., no. 117, 170 p.

MCLEOD, W. N.; DE LA HUNTY, L. E.; JONES, W . R.; HAL- LIGAN, R. 1963. A preliminary report on the Hamersley Iron Province, North-West Division. W. Aust. geol. Snrv. Annu. Rep. for 1962, p. 44-54.

MAITLAND, A. G. 1909. Geological investigations in the country lying between 21'30' and 25"30'S latitude and 113"30' and 118"30' E longitude, embracing parts of the Gascoyne, Ashburton and West Pilbara Goldfields. Bull. W. Aust. geol. Surv., no. 15.

MILES, K. R. 1942. The blue asbestos bearing banded iron- formation of the Hamersley Range, Western Australia. Bull. W. Aust. geol. Surv., no. 100, 37 p.

pt. 1, p. 53-65.

O'ROURKE, J. E. 1961. Palaeozoic banded iron-formations. Econ. Geol., vol. 56, p. 331-61.

RICHTER-BERNBURG, G. 1963. Solar cycle and other climatic periods in varvitic evaporites. p. 510-9. In: A. E. M . Nairn (ed.), Problems in palaeoclimatology, London, New-York, Sydney, Interscience Publisher.

RYAN, G. R.; BLOCKLEY, J. G. 1965. Progress report on the Hamersley blue asbestos survey. Rec. W. Aust. geol. Surv., no. 1965/32. (Unpublished open file report.)

SAKAMOTO, T. 1950. The origin of the Precambrian banded iron ores. Amer. J. Sci., vol. 248, p. 449-74.

STEWART, F. H. 1963a. The Permian Lower Evaporites of Fordon in Yorkshire. Proc. Yorlcs. geol. Soc., vol. 34, pt. 1, no. 1, p. 1-44.

-. 19636. Marineevaporites. Prof: Pup. U.S. geol. Surv. 440-Y. TALBOT, H. W. B. 1920. The geology and mineral resources of the North-West, Central and Eastern Divisions. Bull. W. Aust. geol. Surv., no. 83.

TRENDALL, A. F. 1965~. Origin of Precambrian iron-formations (Discussion). Econ. Geol., vol. 60, p. 1065-70. - . 1965b. Progress report on the Brockman Iron Formation in the Wittenoom-Yampire area. W. Aust. geol. Surv. annu. Rep. for 1964, p. 55-65.

-. 1966a. Second progress report on the Brockman Iron Formation in the Wittenoom-Yampire area. Rec. W. Aust. geol. Surv., no. 1966/1. (Unpublished open file report.)

-. 19666. Second progress report on the Brockman Iron Formation in the Wittenoom-Yampire area. W. Aust. geol. Surv. annu. Rep. for 1965, p. 75-87.

__ . 1966c. Altered pyroclastic rocks in iron-formation in the Hamersley Range, Western Australia (Discussion). Econ. Geol., vol. 61, p. 1451-8.

-. 1968. Three Great Basins of Precambrian Banded Iron Deposition. Bull. geol. Soc. Amer., vol. 79, p. 1527-44. - . 1969. The Joffre Member in the gorges south of Wittenoom. W. Aust. geol. Surv. annu. Rep. for 1968, p. 53-7.

TRENDALL, A. F.; BLOCKLEY, J. G. 1968. Stratigraphy of the Dales Gorge Member of the Brockmann Iron Formation, in the Precambrian Hamersley Group of Western Australia. W. Aust. geol. Surv. annu. Rep. for 1967, p. 48-53.

-. 1969. The Iron Formations of the Precambrian Hamersley Group, Western Australia. Bull. W. Aust. geol. Surv., no. 119, 350 p.

TYLER, S. A.; TWENHOFEL, W. H. 1952. Sedimentation and stratigraphy of the Hurnian of Upper Michigan. Amer. J. Sci., vol. 250, p. 1-27, 118-51.

UDDEN, J. A. 1924. Laminated anhydrite in Texas. Bull. geol. Soc. Amer., vol. 35, p. 347-54. WELLS, J. W. 1963. Coral growth and geochromometry. Nature, Lund., vol. 197, no. 4871, p. 948-50.

Discussion

A. S. KALUGIN. There are indications by LaBerge of the Geology I criticized both his stratigraphic vagueness and presence in the iron quartzites of Australia of clastic ma- the conclusions which he drew from the presence of volcanic terials in the form of volcanic ash. Will you comment on material. Volcanic ash, replaced by stilpnomelane occurs in this point? at least the Marra Mamba iron-formation and the Dales

Gorge member and Joffre member of the Brockman Iron A. F. TRENDALL. LaBerge's descriptions are accurate, Formation. The presence of this volcanoclastic material although in a published discussion of this paper in Economic falling into the basin does not affect the origin for the

269

A. F. Trendall

iron-formation suggested in the paper. The volcanoclastic material is restricted to distinct bands, from a few centimetres to a few metres thick, which consist almost entirely of stilpnomelane, and does not occur in the banded iron- formation itself.

S. ROY. In case of folding, because of the difference in consistency of chert, hematite or magnetite bands, there should be pinching and ñowage in different bands. Does it not affect the stratigraphic correlation of the mesobands?

A. F. TRENDALL. Only in the extreme south of the basin, where the folding is intense. Certain minor folds in the northern part of the basin, which are described in Bull- etin 119 of the Geological Survey of Western Australia, do not affect the correlation of mesobands.

I. P. NOVOKHATSKY. Are there any other indications of evaporites?

A. F. TRENDALL. I a m not sure what other indications of evaporites are expected in the question. Apart from the complete absence of terrigenous clastic material, the main indications of similarity between evaporites in the usual sense and iron-formation of the Hamersley Group are the three features of microbands to which attention is drawn in the paper: their regular repetition, their small thickness, and their very large areal extent.

I. P. NOVOKHATSKY. What is the approximate concentration of readily soluble salts in the water of the evaporite basins?

A. F. TRENDALL. The concentrations of the readily soluble salts in the waters of the Hamersley basin are not known quantitatively, but the presence of diagenetic riebeckite and stilpnomelane in the iron-formations leaves no doubt that sodium and potassium were both present.

E. C. PERRY. Comparing microbanding mineralogy and age, what relation in space and time do you envisage between the Hamersley basin, the South African basin and perhaps the Brazilian basin? Does the geometric shape of

the Hamersley basin that you suggest preclude direct correlation between Hamersley and Transvaal?

A. F. TRENDALL. The age of the Hamersley basin is rather accurately known at 2,000 m.y.; the age of the Transvaal and Cape Province basins could be about the same, but the interbedded volcanics from which an exact age could be obtained have not yet been dated. W e have heard at this Symposium of the difficulty of dating the Brazilian itabirites, Although there are obvious similarities among all three basins, it is dangerous to suppose, before evidence is available from all three, that they necessarily have exactly the same age. There may be no overlap in time in the deposition of the iron-formation in each. The answer to the supplementary question is ‘Yes’-direct correlation is impossible.

E. C. PERRY. Is it likely that differential weathering would affect all rocks equally (crust of average chemical coinposition) or might basaltic rocks be selectively weathered?

A. F. TRENDALL. Differential weathering is of course possible, but it is so difficult to set up any reasonable model for the provision of sufficient basalt to be weathered. The iron certainly was not derived from the basalts of the underlying Fortescue Group, which have normal basaltic composition.

N. A. PLAKSENKO. Could you tell us how the mineralogical composition of the iron-formation changes vertically and horizontally? It would be interesting to know the variation of iron ore minerals.

A. F. TRENDALL. It was mentioned in the paper that the five main stratigraphic units of iron-formation are each chemically, mineralogically and lithologically distinct. It is not possible in a brief answer to give the details of these vertical differences and I can only refer Prof. Plaksenko to Bulletin 119 of the Geological Survey of Western Aus- tralia for a full account. As far as horizontal variation is concerned there appears to be no significant variation throughout theareaof the basin; this fact is emphasized, since it is contrary to published accounts of other iron-formations.

270

Geology and iron ore deposits of Serra dos Carajás, Pará, Brazil

G. E. Tolbert, J. W. Tremaine, G, C. Melcher and C. B. Gomes Cia. Vale do Rio Doce, Div. de Desenvolvimento, Brazil

Introduction The Serra dos Carajás iron deposits were discovered in August 1967, by geologists of the Companhia Meridional de Mineração while undertaking a systematic exploration programme for economic mineral deposits (Tolbert et al., 1968). Supported by boats, small aircraft and helicopters, exploration began on the Xingú River and progressed eastwards to the Itacaiunas River, a tributary of the Tocantins River, from where the initial penetration to the Carajás ranges was made. Additional iron deposits of the same type, but considerably smaller in size, were also discovered 175 km west of the Serra dos Carajás and 30 km north-east of the town of São Felix do Xingú (6"19'-6"30'S. and 51 "41'-52"00' W .). All of the deposits are now controlled by a new company formed by Companhia Meridional de Mineração and Companhia Vale do Rio Doce.

The Carajás region is in a remote, unexplored part of the Amazon rain forest, between two major southern tributaries of the Amazon, the Xingú and the Tocantins rivers (5"54'-6"33'S. and 49"35'-50°34'W.; see Fig. 1). Serra dos Carajás, or the Carajás Range, does not constitute a single mountain range, but, as defined in this report, refers to the two principal iron-bearing ranges, Serra Norte and Serra Sul, and the intervening area.

The field work to date has consisted of reconnaissance geologic mapping of the iron deposits at a scale of 1 : 5,000, airborne and ground magnetometer surveys and topographic surveys. Detailed exploration, including geologic mapping of the ore bodies, diamond drilling and the excavation of adits, has been confined to one area, called N-1, where the main camp and an air-strip are located.

Physiography

Three physiographic features are prominent in the Carajás region: (a) jungle-covered lowlands; (b) long, nearly straight ridges or chains of hills; (c) discontinuous iron-bearing plateaux.

In the state of Pará, the Itacaiunas-Carajás region, between the Tocantins and Xingú rivers, consists of fairly flat lowlands, a few widely scattered ridges and hills, and iron-bearing plateaux. Except for the plateaux, the region is covered by dense, tropical rain forest. North and west of the iron deposits are several long, narrow, westward- trending ridges and chains of low, rounded hills. In contrast, the Carajás ranges are more extensive and sinuous. These uplands are 600-700 m in elevation and have a relief of 200-300 m. Their uniform elevation suggests they are erosional remnants of a former widespread peneplain.

Where the Carajás ranges are underlain by iron-formation, a hard hematite crust or canga capping has prevented the growth of tall rain forest resulting in a series of clearings or savannah areas covered by low brush which stand out in contrast to the surrounding dense jungle. Some clearings have small lakes or swamps in low areas. The topography of these clearings, which is strongly influenced by the underlying rocks, is formed by undulating plateaux and hills, or monadnocks, which rise as high as 100 m above the general plateau surface. These savannah, like plateaux, which coincide with the ore deposits, range from 1 to 30 km in length and from 500 to 2,000 m in width. Nearly vertical scarps, 10-20 m high, and 30-40" talus slopes, are common along the borders where undercutting of the canga rim has caused progressive erosion of the plateaux.

The climate is tropical and temperatures during the day are often high; however, the plateaux enjoy a far less humid and more healthy climate than the lowlands. At the project site the average annual rainfall for 1968-69 was 180 cm, the wet season extending from November to May.

Regional geology

The area underlain by the Carajás iron-formations constitutes a small part of an extensive Precambrian terrain which extends from the Tocantins to the Madeira River and is bordered on the north by sediments of the Amazon basin.

Unesco, 1913. Genesis of Precunibriun iron and manganese deposifs. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 271

G. E. Tolbert, J. W. Tremaine, G. C. Melcher and C. B. Gomec

1 O' i0 O 50 100 150 200km

1 55'

FIG. 1. Index map showing location of Serra dos Carajás.

Before 1950 very little was known about the geology of southern Pará, except for occasional observations made by a few river expeditions. Recently, aerial photographs have provided some geologic information enabling geologists, under the auspices of the Departamento Nacional da Produção Mineral, to establish a generalized sequence of rock types in the Araguaia region (Barbosa et al., 1966). This sequence, which extends as far north as the Itacaiunas River, is composed of five separate units: Undifferentiated Precambrian rocks, consisting of granite, migmatite and paragneiss.

Precambrian metasediments (greenschist facies) correlated with the Araxá Series (described in south-central Brazil) which include mica schist, quartzite and paragneiss.

Slightly metamorphosed phyllite with intercalations of quartzite, itabirite, conglomerate, greywacke and limestone of the Tocantins Series.

Unmetamorphosed, Eopalaeozoic Gorotire Formation consisting of sandstone and conglomerate.

Carboniferoiis sediments of the Piauí Formation, which include carbonate rocks, sandstone, chert and shale.

Parada et al. (1966) mapped the Rio Naja area in the Rio

5'

Fresco region and found a north-east-trending sequence of quartzites and itabirites which they named the Tocandera formation. This formation may correspond, in part, to the Tocantins Series.

Almaraz (1967) dated various rocks (granite, amphibolite, migmatite) in the Marabá-Itacaiunas area using the potassium-argon 'method and obtained an average age of approximately 2,000 my. Almeida and co-workers (1968) have included this area in the 'Tocantins-Tapajós craton', the last metamorphism of which is also dated at about 2,000 m.y.

Amara1 (in preparation) has recently obtained additional age information on rocks collected during the current project. His results also indicate that the last principal metamorphic event occurred at about 2,000 m.y., as confirmed by determinations on seven samples of gneiss, amphibolite and muscovite schist from the Itacaiunas River. The age of one of the samples, an amphibolite from the western part of the Serra de Itapirapé, situated about 65 km north-west of area N-1, was determined to be 3,2805 113 m.y. Interestingly, this is the oldest date yet determined for a Brazilian rock.

272

Geology and iron ore deposits of Serra dos Carajás, Fará (Brazil)

A considerable amount of geologic information has resulted from current field work, including the completion of surveys in previously inaccessible areas made possible by the utilization of helicopters. Owing to the extensive size of the region, the ubiquitous forest and laterite cover, and the complete lack of population or supply facilities, it has been extremely difficult to establish a stratigraphic sequence or to correlate the principal lithologic units mapped during the current exploration programme with those enumerated by Barbosa et al. (1966). The writers recognize that some of the lithologies found at Serra dos Carajás are similar, in part, to those described in the Araguaia region; for example, quartzite and itabirite of the Tocantins Series and mica schist of the Araxá Series (?) have counterparts in the Carajás rocks. Nevertheless, only after considerably more work has been done on the structure and stratigraphy in both regions will reliable correlations be possible. Several rock types distinguished by geologic mapping are described below and shown in Figure 2.

GNEISS, GRANITE AND AMPHIBOLITE

Much of the Itacaiunas-Carajás region is underlain by rocks which are dominantly gneiss with subordinate granite, amphibolite and granulite. Many of the rocks called gneiss in the field are found to be mylonitized granite when examined under the microscope. The varieties of gneiss include orthogneiss, injection gneiss and migmatites. In the area west of the Itacaiunas River, permatite dikes cut granite, gneiss and metasediments.

Bands of amphibolite several metres thick are commonly associated with gneiss and appear to be derived from mafic rocks. Blastophitic texture is preserved, and uralitized pyroxene or relicts of pyroxene are altered to hornblende along border zones. Some varieties with garnet crystals aligned parallel to the lineation may be of sedimentary origin.

Granulite, with typical granoblastic texture, crops out in a few places. It is composed of quartz, andesine-labradorite, augite, hypersthene, garnet and biotite.

of phyllite and schists are very rare. Those few phyllite samples collected are composed of quartz and sericite with subordinate chlorite, biotite, actinolite, epidote and opaque minerals.

Muscovite schist is the dominant mica schist, but chlorite, actinolite, biotite and graphite schists are also found. Most of these rocks, which are fine-grained with lepidoblastic texture, contain epidote, garnet, albite, microcline and carbonate.

Deposits of manganese oxides believed to be derived from the oxidation of quartz-spessartite schist are found in a few places along ridges north-east and north-west of Serra Norte.

IRON - FORMATION (ITABIRITE) Owing to the fact that leaching and enrichment have been so thorough in Serra dos Carajás, fresh, unaltered itabirite is rarely encountered. Ferruginous quartzite and itabirite are found on ridges north of the iron deposits which suggests that iron-rich sedimentation was not confined exclusively to the Carajás ranges.

Itabirite is composed of alternating laminae of quartz and iron oxides that range in thickness from 0.05 mm to 10 m m . The main minerals in this rock are quartz (recrystallized chert), magnetite and hematite. Secondary minerals are goethite, martite, gibbsite and, rarely, sericite. Xeno- morphic quartz, both granular and elongate, occurs in two grain sizes. In places, very fine-grained quartz (0.005- 0.01 mm) shows a typical mortar texture and appears to have a cataclastic origin. A possible third generation of quartz is represented by tabular crystals, with grain sizes from 0.005 mm to 0.03 mm, that grow perpendicular to crystal faces and ñll the interstices of opaque minerals.

Hematite is nearly always peripheral to magnetite grains evidencing the growth of hematite at the expense of magnetite. Idiomorphic magnetite crystals, which are equi- dimei-isional and submillimetric, are aligned parallel to the banded structure but, as observed in a few specimens, they truncate this structure, suggesting a later stage of magnetite crystallization.

QUARTZITE, PHYLLITE AND MICA SCHIST SANDSTONE A N D CONGLOMERATE

Pure, granoblastic quartzite is the dominant type with subordinate ferruginous schistose varieties. Sericite, epidote, biotite, limonite and pyrite are common mineral assemblages in these rocks. The grain size ranges from 0.1 mm to 1.2 mm; cataclastic and mylonitic textures are common. Most of the quartzitic rocks are metamorphosed to low grades, although the presence of garnet, amphibole and pyroxene indicate higher grades in the area north of Serra Norte toward the Itacaiunas River.

Argillaceous metasediments and schists are interbedded with iron-formation and undoubtedly constitute a large part of the section. However, intense tropical weathering has converted these rocks to reddish laterite and fresh exposures

The area between the western part of Serra Norte and Serra Sul is underlain by friable sandstone, quartzitic sandstone, quartzite and conglomerate. These rocks are considered to be younger than the metasediments and may be equivalent to the Gorotire Formation. The sandstone is composed of fine-grained, xenomorphic quartz with subordinate sericite, biotite, chlorite, zircon and opaques. Some varieties are conglomeratic with fragments (0.4 mm to 1.5 cm in diameter) composed of quartzite and banded itabirite. The matrix is fine-grained quartz, iron oxide, sericite and biotite. Angular pebbles found in some places suggest that these sediments were derived from a local source.

273

G. E. Tolbert, J. W. Tremaine, G. C. Melcher and C. B. Gomes

. . , , ... ..' .. ..::,

Gorotire f o r m a t i o n I?). quartzitic s a n d s t o n e , partly conglomeratic

Quartzite. phyllite. a n d m i c a with intercalated.

b a n d s of iron formatinn.(if). c o m m o n l y c a p p e d b y tanga

Gneiss a n d granite

LlthologiC contact

Fa U I t ,d a s i e d w h e r e in fe rred. questioned w h e r e doubtful. mostly photogeologic interpretation

__------ Structural lineament: p h o t o - geological interpretation

-i--+ A X ~ S o f anticline or syncline with plunge: photogeological interpretation

SCALE O 2 4 6 8 1Okm

FIG. 2. Generalized geologic map of Serra dos Carajás region.

MAFIC ROCKS

Several varieties of mafic rocks, probably of more than one age, are found in the region. A petrographic study of mafics from the Itacaiunas area resulted in the identification of diabase, olivine diabase, metadiabase, basalt,

aiid andesite porphyry. Diabase dikes with ophitic texture are the most common aiid are comprised of dominant pyruxene aiid plagioclase (generally labradorite) with minor quartz, olivine and biotite. A n age determination on andesite porphyry from the Itacaiunas River gave 507&99 m.y., which corresponds to the age of otlier

274

Geology and iron ore deposits of Serra dos Carajás, Pará (Brazil)

rocks of basaltic composition in thc Amazon region (Almaraz, 1967).

Mafic intrusions control, to some extent, the physical boundaries of ore bodies, as disclosed by adits and drill holes at N-1 which have intersected both concordant and discordant mafic dikes and sills or have terminated in mafic rocks. The dikes and sills range from less than 1 m to several tens of metres in thickness. Aerial photographs reveal several north-east-striking dikes (presumably diabase), as much as 10 km long, that intrude the rocks between Serra Norte and Serra Sul. Whether these tabular intrusives correspond to the mafics encountered at depth is not known. Both appear to be basaltic in composition but they may be of different ages.

Structure

The geologic map of the Araguaia Project (Barbosa et al., 1966) shows a belt of north-trending Precambrian rocks parallel to the Araguaia and Tocantins Rivers, This structure was formerly recognized by Kegel (1965) who named it the ‘ Araguaia-Tocantins lineament’. About 150 km south of Marabá the strike of these rocks curves to the west and in the Carajás region they trend west to west- north-west. This major change in strike, also noted by Kegel (1965), coincides approximately with the eastern end of Serra Sul where the iron-formation turns southward (Fig. 2). This important tectonic feature may be related to the structural complexity of the Carajás rocks as well as responsible for the eastward diversion of the Araguaia and Tocantins Rivers east of Marabá (Fig. 1).

Deformation of the rocks in areas north and west of Serra dos Carajás produced simple, open folds with gentle dips and horizontal fold axes striking west to west- north-west. In contrast to this relatively mild type of deformation, the Carajás rocks are complexly folded with west-north-west fold axes that plunge both east and west. Tight, crenulated folds, isoclinal in places, with amplitudes varying from a few centimetres to several metres, are typical features of in situ hematite and itabirite. Plastic flowage has caused thickening and thinning of beds.

In addition to folding, block faulting and tilting have been instrumental in segmenting Serra Norte and Serra Sul into the several isolated plateaux which characterize those ranges (Fig. 2). Aerial photographs show a prominent fracture system oriented north-east and north to north-west; faults with apparent displacements of as much as several hundred metres have been discerned. A major lineament, interpreted as a fault, strikes west-north-west across the area between the two iron ranges and is tan- gential to the easternmost area of Serra Sul (Fig. 2).

The structure of Serra Sul is somewhat less complex than that of Serra Norte. The western part of Serra Sul is intersected, and in places displaced, by north-striking faults, but the itabirite beds are generally more continuous than those at Serra Norte. The central sector of Serra Sul appears to be offset to the north relative to the eastern

and western blocks (Fig. 2). The area between the western parts of Serras Norte and Sul, underlain by arenaceous sediments and metasediments, is intersected by a system of north-east- and north-west-trending fractures with a ‘herringbone’ pattern.

Iron deposits

ITABIRITE PROTORE

The Caraj ás iron-formation has been described generally in a preceding section. Its chemical and physical properties closely approximate oxide-facies iron-formation found in other parts of the region as well as in the Quadrilátero Ferrífero in Minas Gerais (Dorr, 1964, 1965). Comparison of the size of quartz grains in itabirite from the two regions, for the few data available, suggests that the coarser grain sizes are similar, but quartz in the Carajás itabirite tends to be finer-grained. Other facies of iron-formation have not been found except for one 1.4 m interval of banded carbonate rock, probably dolomite, intersected in a drill hole. The iron content of unenriched iron-formation in the few samples that have been analysed ranges from 17 to 41 per cent Fe (Table 1 ). The average of these few analyses is not considered to be representative. The original thickness of the iron-formation was probably a few hundred metres, but in some areas complex folding and faulting have multiplied the apparent thickness to as much as 1 km.

TABLE 1. Analyses of itabirite samples, Serra dos Carajás

Distances from Wt percentage, dry basis No. adit portal,

inmeters Fe0 FeZO:. Si03 A1,0, PcOa Fe

N1 T2 68 6.53 51.23 38.47 0.48 0.11 40.90 N1 T2 69 1.15 39.39 56.19 0.55 0.07 28.44 N1 T5 165 1.00 23.35 60.84 0.37 0.09 17.11 M A 51 Surface sample 0.18 5.45 70.39 0.68 0.06 20.26

The uniform composition and texture as well as the consistency in thickness of laminae provide evidence that this rock was a shallow water sediment consisting of alternating layers of silica and hydrous iron oxides. Sub- sequent burial and diagenesis converted silica to chert and removed water from the iron minerals. Metamorphism recrystallized chert to quartz and the iron minerals to magnetite and subordinate hematite. Most of the hematite observed today resulted from the oxidation of magnetite under surface or near-surface conditions.

FRIABLE ORE BODIES

Detailed mapping of the ore deposits has been completed only in the northern range (Fig. 3). Furthermore, the determination of the dimensions of ore bodies by subsurface

275


LEGEND

Contact High-grade hematite, in place

Strike and diD of Hematite canga bedding

+ He'rnatite-goethite c a n g a Vertical bedding

Fault

45 _I

SCALE 0 I 2 3 4 5 k m I

@ .. ... .. .. .. . . . . ..

FIG. 3. Geologic map of Serra Norte showing distribution of ore bodies.

exploration is limited to area N-1, because no drilling or excavation of adits has yet been started on other deposits. Therefore, the 'ore bodies' as shown on Figure 3, except for N-1, are delineated by surface mapping only. The major ore deposits underlie hills that stand above the surface of the plateaux. Although flanks of the hills are covered with canga, the crests are generally formed by hard, in situ hematite. Principal ore bodies are composed of friable hematite and occupy an irregular zone between outcropping hard hematite, or the canga capping, and unleached itabirite below (Fig. 4). The thickness of the ore bodies depends mainly on the extent to which meteoric water has been able to penetrate and leach the iron-formations. Friable ore, as observed in adits, is dark reddish- brown or metallic grey. Preliminary grain size analyses of a few samples show that 24 per cent of the ore is above one-quarter inch and 76 per cent is below this size.

Partial sampling of the ore bodies in area N-1, which may or may not be representative of other untested deposits, indicates an average grade of more than 65 per cent Fe (Table 2). The iron content in the N-1 ore bodies seems to be fairly constant to the depths attained by current

drilling (Fig. 4). Drill hole NlD14 shows a decrease in Fe and corresponding increases in alumina and silica in the 20-40 m interval where the hole intersects a diabase dike (Fig. 5). Owing to zones of partially leached itabirite, similar variations in grade appear between 120-200 in. A lower iron content accompanied by an increase in alumina and silica mark the presence of a dike between 20-30 m in drill hole NlDlO.

The following ore minerals were identified (confirmed by X-ray analyses): hematite, magnetite, goethite and martite. The texture of friable hematite is granular or platy with two common grain sizes, 0.03 mm and 0.2 111111. Xenomorphic grains of hematite commonly have an equidimensional or elongate habit. Tabular crystals of hematite or specularite, growing perprndicular to other hematite grains or filling cavities, probably represent a later stage of crystallization. Specularite also occurs in indurated zones along the walls of dikes.

Magnetite crystals are equidimensional, idiomorphic and vary in size from 0.02 mm to 0.05 mm. Although magnetite is generally oxidized to hematite, the converse has not been observed. Along grain boundaries and part-

276


Friable hematite I ta biri te

~- - - - - - - - Geologic contact,

Mafic intrusive d a s h e d where inferred

FIG. 4. Cross section of part of area N-1, Serra Norte, showing relations between1ore bodies, itabirite, canga and mafic intrusives.

ing planes magnetite is altered to martite. The ore is generally slightly magnetic and, in a few drill holes, short intervals of 1-2 m contain as much as 40 per cent magnetite; however, the average magnetite content of the ore is estimated to be only a few per cent.

Goethite, which is the most important secondary mineral, replaces hematite along grain boundaries and fractures. Accompanied by gibbsite, it frequently fills interstices and cavities.

HARD HEMATITE

Indurated hematite occurs on the crests of hills, where it commonly retains the laminated structure of the protore, and, in addition, forms lenses and tabular bodies within soft ore deposits. It is composed of metallic-grey hematite and specularite and generally has a tenor of 66 per cent Fe or higher (Table 2).

The surficial type of hard hematite forms a crust 10- 20 m thick and is usually more hydrous and higher in phosphorus than the underlying ores. In places it has been partially disintegrated and cemented by hematite or goethite.

TABLE 2. Chemical analyses of selected ores from Serra dos Carajás

Wt percentage, dry Sample1 Losson

Fe P SiO? Also3 M n Tio, ignition

A 63.47 0.228 0.97 2.64 0.04 0.44 0.019 7.05 B 68.13 0.079 0.47 1.13 0.04 0.20 0.020 2.52 C 66.58 0.068 0.77 3.72 0.62 - - 0.93 D 67.55 0.035 0.61 2.58 0.09 0.02 0.003 2.24 E 69.11 0.014 0.68 0.03 0.03 < 0.005 - 0.62

Fe+ + E 3.42 E (Spectrographic analysis) 0.001-0.01 Ca, Pb < 0.02 Zn, W, Sb, N a

0.0002-0.002 Cu < 0.01 Sr 0.0005-0.005 M g < 0.006 N b

< 0.003 Cr, S n < 0.001 V, Ba, Bi, Ni, Co, M o < 0.0005 A g

1. Sample A: Canga, from area N-4, Serra Norte. Sample B: Hard, in situ hematite ore from surface, N-1, Serra Norte. Sample C: Friable ore from adit N1 T1, area N-1, Serra Norte. Sample D: Friable ore from surface, western part of Serra Sul. Sample E Hematite fines from adit N1 T1, area N-1, Serra Norte.

277


Drill hole NfDfO

FIG. 5. Profiles of drill holes NlDlO and NlD14 in area N-1, Serra Norte, showing variations in iron, silica and aluminium with depth.

Lenses of hard ore below the surface range from a few centimetres to several metres in thickness and extend along strike from less than 1 m to tens of metres. Normally this material has a massive texture, however, it may also exhibit relict bedding; both concordant and discordant relationships are observed with the enclosing rock. Although few data exist, it is believed that most of the hard hematite, other than the surface material, is the product of metamorphism and redeposition of iron oxides; it does not appear to be a residual concentration resulting from the leaching o€ silica by meteoric water. An alternate interpretation, unsubstantiated by this study but possibly valid in some of tlie occurrences, advocates the primary deposition of iron minerals, with little or no silica, as lenses in banded iron sediments. Subsequent metamorphism recrystallized the iron minerals to iron oxides and deformation may or may not have displaced the hard iron lenses to discordant positions with respect to the enclosing rocks.

CANGA

Covering the slopes of the hills and the flat portions of plateaux is a ferriferous capping, or canga, that ranges

Drill hole NfDf4

from 1 to 20 m in thickness (Fig. 4). Canga consists of unoriented grains, pebbles or fragments, cemented by hydrated iron oxides and minor clay. Fragmental material is composed mostly of hematite, goethite, itabirite, some magnetite, pebbles and blocks of previously formed canga, and clay minerals, Rarely, other rocks, such as phyllite or schist, comprise the cemented material. Fragments may be of millimetric size or as large as several decimetres in diameter.

Canga varies from a hematite-rich variety found on the flanks of hills to progressively more hydrated types (goethite) near the plateau borders. Analyses of a few grab samples of this material contain above 60 per cent Fe and average 0.20 per cent phosphorus (Table 2). Surface water draining through fractures in the canga erodes the underlying soft ore, phyllite and schist, carving out large caverns especially under the rims of the plateaux. Canga is considered to have formed by the mechanical decomposition of bedrock and subsequent cementation of fragments by chemically precipitated mixtures of hydrated iron oxides, fine-grained iron oxides and minor clay. Erosion and transportation of surface rubble has resulted in the formation of canga in areas which are not necessarily underlain by itabirite or ore deposits.

278


Summary and origin of ore deposits

At this early stage in the study of the Carajás iron deposits not enough data are available to recapitulate in detail the events related to their origin. Thus the following summary is based partly on observable facts in the Serra dos Carajás and partly on comparisons with similar deposits. The evolution of the iron deposits began in early or middle Precambrian with the rhythmic deposition of iron-rich chemical precipitates, intercalated with siliceous and argillaceous sediments, in a shallow basin probably adjacent to or associated with a stable cratonic region. The basin of deposition was approximately 300 km long in an east-west direction and 50-60 km wide. Whether or not volcanic rocks formed part of the above sequence is unknown. The only rocks in the region with a possible volcanic origin deposited at this time are amphibolites and their exact stratigraphic relation to the iron-formations and metasediments is not clear.

Subsequently the unconsolidated iron sediments were buried and compacted. In middle Precambrian the basin

was affected by an orogeny accompanied by complex folding, faulting aiid metamorphism. Folding, and to a lesser extent, faulting, exerted an important influence on the later development of ore bodies by increasing the thickness of the protore. lion-rich sediments were metamorphosed to itabirite. Later faulting and fracturing increased the permeability of the iron-formations. The fact that Serra dos Carajás was the locus of intense deformation is significant because relatively undeformed iron- ïormations in outlying areas with similar relief did not form ore deposits.

The physiographic evolution of this region began with the development of an extensive erosion surface above which a few monadnocks of resistant iron-formation pro- truded. Epeirogenic movement was responsible for uplifting the peneplain approximately 700 m, where most of it was destroyed by erosion leaving knobs and ridges of resistant quartzite and itabirite. The protective canga capping prevented the erosion of soft ore, but was sufficiently permeable for meteoric water to leach silica from underlying itabirite resulting in the residual enrichment of iron oxides and the ore bodies observed today.

Géologie et dépôts de minerai de fer de lu Serra dos Carajás, Pará, Brésil (G. E. Tolbert, J. W . Tremaine, G. C. Melcher et C. B. Gomes)

En 1967, de nouveaux gisements de minerai de fer importants ont été découverts dans la région de la Serra dos Carajás dans la partie sud de l’État de Pará au Brésil (6” de latitude S, 51” de longitude O). Les recherches effectuées à ce jour comprennent l’établissement d’une carte photogéologique régionale et d‘une carte topographique d’une partie de la surface à l’échelle du 1/5 000, quelques forages, quelques tunnels dans certains dépôts et des études minéralogiques. Ce programme, auquel s’ajoute un levé magnétométrique de surface, se poursuit. Les gisements sont localisés dans une série de plateaux de direction géné- rale nord-ouest coiffés de canga qui apparemment sont les flancs d’un synclinal régional qui plonge rapidement vers le nord-ouest. Ces flancs sont distants de 30 km dans la région ouest de Carajás et ils convergent vers le nez du synclinal à 100 km à l’est. U n réseau de failles et de fractures pointant vers le nord-est et le nord-ouest constitue le caractère structural dominant de la région. Les plateaux sont discontinus et leur surface varie de un à quelques dizaines de kilomètres carrés. Ils s’élèvent approximati - vement à 500 mètres au-dessus des terres basses couvertes de forêts et s’apparentent à d‘autres chaînes de la région, ce qui fait penser qu’ils sont peut-être les vestiges d’une surface d’érosion étendue.

La croûte de canga, dont l’épaisseur varie de 1 à 30 mètres, est constituée de fragments d’hématite et d‘autres composants ferrugineux liés par des oxydes de fer hydratés. Le minerai primitif (protore) consiste en formations de fer métamorphosées (itabirite) de l’époque précambrienne qui ont été vigoureusement plissées. Le minerai est constitué de plaquettes d’hématite à grain fin, friables, avec en second lieu de la magnétite ; cependant, en quelques endroits, le minerai est dur et massif. Quelques analyses préliminaires indiquent que la teneur est équivalente à celle des autres dépôts d’hématite à haute teneur qu’on rencontre dans le monde entier.

O n observe des intrusions de filons mafiques dans les formations de fer ; leur épaisseur varie de quelques centi- mètres à plusieurs dizaines de mètres. Parmi les autres types de roche qu’on trouve dans la région figurent la phyllite, le quartzite, le grès et des roches cristallines telles que le granite et le gneiss. Les datations de ces roches cristallines par la méthode du potassium-argon indique un âge d‘environ 2 milliards d‘années.

La lixiviation supergène de la silice semble avoir été la cause de la formation de la plupart de ces gisements de fer. Parmi les autres facteurs qui ont contribué au déve- loppement des gisements figurent la préservation de la surface d‘érosion et de la croûte ferrugineuse, les plissements et failles complexes des formations de fer, ainsi que le climat tropical.

279


Bibliography / Bibliograghie

AMARAL, G. 1969. Nota previa sôbre o reconhecimento geocronológico do Pre-cambriano da região amazônica. XXIII Congresso Brasileiro de Geologia, Bol. Especial, no, 1,

AMARAL, G. (In preparation .) Reconhecimento geocronológico do Pre-cambriano da região amazônica.

ALMARAZ, J. S. U. 1967. Determinações K-Ar na região do curso médio do Tocantins. Bol. Soc. bras. Geol., vol. 16, no. 1, p. 121-6.

ALMEIDA, F. F. M. de. 1967. Origem e evolução da plataforma brasileira. Bol. Div. Geol. min. Rio de J., no. 241, 36 p.

ALMEIDA, F. F. M. de; MELCHER, G. C.; CORDANI, U. G.; KAWASHITA, K.; VANDOROS, P. 1968. Radiometric age determinations from northern Brazil. Bol. Soc. byas. Geol.,

BARBOSA, O.; ANDRADE RAMOS, J. R.; GOMES, F. A. de;

p. 81-2.

vol. 17, p. 3-14.

HELMBOLD, R. 1966. Geologia estratigrafia, estrutural e economica da área do projeto Araguaia. Monogr. Div. Geol. min., Rio de J., no. XIX, 94 p.

DORR, J. Van N. II 1964. Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., vol. 59, no. 7, p. 1203-40.

-. 1965. Nature and origin of the high-grade hematite ores of Minas Gerais, Brazil. Econ. Geol., vol. 60, no. 1, p. 1-64.

KEGEL, W. 1965. Lineament-tektonik in Nordwest-Brasilien. Geol. Rdsch., vol. 54, no. 2, p. 1240-60.

PARADA, J. M.; FORMAN, J. M. A.; FERREIRA, J. P. R.; LEAL, J. F. 1966. Pesquisas minerais no Estado do Pará. Bol. Div. Geol. min., Rio de J., no. 235, 24 p.

TOLBERT, G. E.; SANTOS, B. A. dos; ALMEIDA, E. B. de; RITTER, J. E. 1968. Recente descoberta de ocorrências de minério de ferro no Estado do Pará. Mineraç. e Metall., vol. 48, no. 288, p. 253-6.

Discussion

M. V. MITKEYEV. In your collection you have a specimen of hard hematite. How do you account for the origin of this ore?

G . E. TOLBERT. The specimen of hard hematite exhibited is not a typical or representative type of ore from the Serra

dos Carajás. Hard hematite is rather rare. Although we have very little information to date, we tend to attribute the origin of the hard hematite lenses to metamorphic differentiation, thas is, an origin similar to that advocated for similar ores in Minas Gerais (Brazil) by Professor J. Van N. Dorr.

280

Enrichment of banded iron ore, Kedia d’Idji1, Mauritania

F. G. Percival Sadlers End, Haslemere, Surrey (United Kingdom)

Introduction The chief purpose of this paper is to record certain evidence which has a bearing on the age of leaching and enrichment of the iron ores of the Kedia d‘Idji1, Mauritania. In order to assist in evaluating the evidence a brief description of the ores and rock types is necessary.

The Kedia d’Idjil (Fig. 1) is an inselberg of Pre- cambrian rocks with quartzites, schists and banded ironstones along its northern flanks and heights, running mainly east-west for some 24 km, and dipping steeply to

the south. To the west it narrows and descends to plain level near the village and fort of the former Fort Gouraud, now called F’Derik. Eastwards the range widens to a maximum north-south development of about 10 km. It rises to elevations of 300-500 m above the Saharan plain, which here consists mainly of orthogneisses covered with ‘reg’-a loosely consolitated blanket of sand, gravel and boulders. South of the belt of regularly bedded rocks a so-called ‘breccia’ blankets the hill, as shown in Figure 1. Towards both extremities of the range the strike has a marked twist to a NW.-SE. direction, and at these

L E G E N D

1-1 Reg ml Breccia

B.H.Q. &

Schists

m Bedded Breccia Canga

Cambrian Sandstone ml Conglomerate v)I Visible Ore ..

e-

r

/3 22.45’ 11115111P

ZO U E R A T E MIHE TOWNSHIP

KEDIA D‘IDJIL

I n i z 3 4 5 6 7 a 9 i p

K I L O M E T R E S

FIG. 1. Geological map of the Kedia d’Idjil, somewhat simplified.

281 Unesco, 1973. Genesis of Precumbriun iron and manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.)

F. G. Percival

distortions the largest enriched ore bodies are located. The banded ironstones are formed of largely recrystallized bands of heinatite and quartzite. The bands are of varying thickness, but normally do not exceed 5 m m . These rocks are similar to the itabirites of Brazil and this term will be used for convenience, though at the mine they are commonly called BHQ (banded hematite quartzites). A series of approximately N.-S. faults crosses the range, marked by deep canyons. These have not been shown on the simplified map (Fig. 1).

Blondel (1952) gave a valuable short account of the Kedia, and Blanchot (1955) gave more details of the deposits, their covering of breccia, and their relationship to the sub-horizontal Palaeozoic sandstones to the east, He also described a Conglomerate containing boulders of this breccia-boulders whose size frequently attained that of a camel’s head-exposed at the ravine of O u m el Hbel on the southern flank of the range (Fig. 1). Both the breccia and the conglomerate will be more fully discussed lierein. Lethbridge and Percival (1954) wrote a general description of the occurrences of iron ore, and of the prospecting work in hand up to 1953, and papers on special aspects of the deposits have been published by Huvelin (1963) and by Baldwin and Gross (1967). A description of the mining development of the ore bodies was given by Audibert et al. (1964).

To the east of the Kedia lie the Cambrian sandstones, with some dolomite, and with a very gentle easterly dip. They develop a NE.-SW. trending escarpment, a feature that continues for some 300 km south-westerly to the oasis of Atar. These sandstone overlie the Precambrian rocks (both the bedded rocks and the breccia) at the eastern end of the Kedia, but only to a moderate height, well below the crest of the range. Blanchot considers that

probably the higher levels were never covered and the Kedia stood out as an island in the Cambrian sea. At O u m el Hbel these Palaeozoic sandstones pass into the conglomerate facies mentioned above, containing large boulders of breccia. Blanchot suggests that this is a local beach development, and describes the passage eastwards, between Oumel Hbel and Chig, as marked by recurrences of sandstone 1-1 O m inextent interstratified in theconglomerate.

The folding of the itabirite series of the Kedia was completed long before the Cambrian was laid down and there is no evidence on the hill to suggest that the itabirites along the crest were covered at any time by rocks later than the Cambrian other than localized surface canga and laterite. The tectonic conditions at F’Derik and Tazadit were evidently favourable for leaching throughout a period of Precambrian time, whose duration is at present unknown, and on to the present time.

The Breccia d’ldjil

Various theories of origin have been suggested for this unusual type of rock. It forms a blanket stretching from the southern outcrop limites of the steeply dipping itabirites southwards for distances up to 7 km or more, to form ravine walls more than 100 m high along the southern boundary of the Kedia. Earlier observers who first met it exposed along its northern borders unhesitatingly called it a tectonic breccia. Those who first saw it in the southern ravines were equally certain it was a conglomerate, having rounded boulders of itabirite with a mainly siliceous cement. East of Rouessa the itabirite in situ is in places preserved in the process of crumpling up to form breccia, undoubtedly tectonic (Fig. 2) and if one assumes, as

FIG. 2. Photograph

282

of itabirite, east of Rouessa, crumpling up to form breccia (‘banded breccia’).

Enrichment of banded iron ore, Kedia d‘Idjil (Mauritania)

seems probable, that this movement was a part of the general orogeny of the Kedia, then the formation of the breccia was contemporaneous with the uplift of the range. In the deep ravines of the south the ‘breccia’ boulders are well-rounded and obviously water-worn, with occasional lenticles of sand. These cemented rolled boulders have been transported and cannot be matched against each other. One may conclude that the whole breccia/conglomerate was formed on a steeply sloping, foundering shoreline, with the boulders becoming rolled and worn in the off-shore depths. The ‘crumpled’ breccia just mentioned has been mapped as an E.-W. zone of ‘bedded breccia’ between the unbrecciated itabirite (with localized enriched ore bodies) and the completely brecciated area. This transition zone has been included with the itabirite in Figure 1.

At a number of places minor enrichments in hematite of the breccia and bedded breccia occur.

Types of enriched ore

As previously stated, the banded ironstones are similar to the Brazilian itabirites. They have been locally enriched in iron by the leaching-out of the siliceous bands, with or without the introduction of secondary iron oxide. Where secondary iron oxide has completely taken the place of the leached silica, massive hard hematite ores of great purity have been formed, with an iron content of 67 per cent or more. The deposit of F’Derik, at the western end of the range, is mainly of this type. The primary bands are fine-grained and massive. The secondary bands are less compact, and in some cases specular, and their secondary nature is shown by local stringers from band to band, crossing the primary bands. In other deposits along the Kedia, where the leached-out silica bands have not been replaced by secondary iron oxide the residual hematite bands may still be held in place by local transverse developments, but generally they tend to collapse, forming biscuity fragments locally called ‘plaquette ores’. In some parts the siliceous bands originally contained dispersed tiny grains of iron oxide, and this insoluble hematite remains in some places after the leaching of the silica, as a fine-grained powdery ore. In the Rouessa deposits, located about the middle of the Kedia Range, both plaquette and powdery ores occur, and also a moderate amount of hard ore in which the silica bands have been replaced by secondary iron oxide (Fig. 3; Percival, 1967).

The foregoing ore types are all formed directly in place of bedded itabirite, but comparatively small occurrences of ore, formed by the enrichment of the breccia in iron, are seen locally, e.g. at Azouazil, and on a small peak 500 m south of eastern Rouessa. They are also reported in the upper part of the O u m el Hbel valley. These occurrences are south of the alignment of the main high-iron enrichments of bedded itabirite. Their textures are those of the breccia, but the eiirichment in iron may be partly an impregnation rather than a replacement of the itabirite fragments; the secondary iron oxide fills up crev-

I I I m m

FIG. 3. Rouessa hard ore, borehole RBI, depth 16.4 m; micro- photo, incident light, showing bands of primary (massive) hematite and bands of secondary specularite, with a stringer crossing the massive bands.

ices and other cavities, and moulds itself round the fragments. Recognizable itabirite is largely doubly recrystallized, and has the appearance of having been under a distorting pressure that has caused a coarse recrystallization of the quartz, with a concomitant flow of the more mobile hematite. Fragments of enriched bedded hematite in these occurrences are rather rare, but such fragments do exist.

The Oum el Hbel conglomerate

In the paper on Fort Gouraud by Lethbridge and Percival (1954) it was suggested that it was not unlikely that the bulk of the leaching of silica from the banded rocks took place in Precambrian times. This surmise was prompted by the fact that certain Precambrian conglomerates with which the present author had been familiar in India contained

283

F. G. Percival

FIG. 4. Oum el Hbel conglomerate, filling the Precambrian valley.

pebbles of enriched hematite ore of a type very like that of F’Derik. Consequently when Blanchot (1955) reported the existence of a Cambrian conglomerate, on the south- east flank of the Kedia, the writer arranged to visit this deposit to search for possible boulders of enriched hematite. The cement in the Cambrian conglomerate is mainly siliceous. The conglomerate in situ proved to be infilling a Precambrian valley and spreading over on to its flanks (Figs. 4 and 5). Boulders of hematite were so numerous in the lower levels of this valley as to suggest that there had been some kind of placer segregation by gravity. The first impression was that the problem of age of enrichment was solved, and the hematite boulders were taken to be of massive ore of the F’Derik type. Later examination of the specimens collected shows that this is far too simple a solution. The hematite boulders are well-rounded, and there is no doubt whatever that the enrichment was completed before incorporation in the O u m el Hbel conglomerate. These boulders are intensely hard, and more difficult to break than the massive hematite of F’Derik. Most of them have a lower density (3.9-4.4) than that of F’Derik hematite (5.0). The normal itabirite of the Kedia has a density of about 3.4. The density of the unenriched breccia is, of course, much the same as that of the itabirite. Microscopic examination reveals the presence of a good amount of silica in most of the hematite boulders, and the textures of nearly all of them examined to date are those of enriched brecchia, and not of normal leached and enriched itabirite. Certain exceptions will be described later.

Small areas of enriched breccia may have been sources of the O u m el Hbel hematite boulders and, if so, the enrichment of the breccia was Precambrian. One may also concede that a small number of non-brecciated enriched hematite FIG. 5. Oum el Hbel conglomerate, close view.

284

Enrichment of banded iron ore, Kedia d’Idjil (Mauritania)

boulders and fragments from the Rouessa hard hematite outcrops further north could have travelled southwards down the valley, but the distance of the O u m el Hbel conglomerate from the outcrop of the belt of itabirite that lends itself readily to enrichment is at least 7.5 km.

Textures of the hematite boulders of the Oum el Hbel conglomerate

Before discussing the make-up of the hematite boulders it should be emphasized that whether one has an almost pure hematite boulder or one that is only moderately enriched, each is an isolated boulder embedded in the siliceous matrix of the conglomerate, and the majority of the neighbouring boulders are rounded lumps of breccia with no suggestion of hematite enrichment, i.e. there is no indication whatever of any enrichment in hematite after the incorporation of the boulders in the conglomerate. What is in question is whether or not some occurrences of hematite within the enriched boulders are fragments of leached and enriched itabirite comparable with the enriched ores of F’Derik, Rouessa or Tazadit. In any event, accepting the Cambrian age of the conglomerate, the enrichment in hematite was Precambrian.

To date the author has found only one conglomerate boulder completely made up of hard hematite, with only a very small content of silica. It has been sectioned in two planes at right angles to each other. Some indication of bedding or banding is seen, but it is not well marked, and here is no close resemblance between this hematite and the

enriched bedded ores. It is not brecciated. A few cavities within the hematite have allowed larger crystals of hematite to develop, and numerous tiny cavities occur, some of which contain small amounts of recrystallized quartz. These cavities show some alignment, but it is not so well marked as the banding of the bedded ores, yet it is possible that this specimen is a boulder of high-iron bedded hematite that has suffered some recrystallization.

A second hematite boulder from the conglomerate contains two portions that seem to be derived from enriched bedded ore of the Rouessa type. One (Fig. 6) shows alternating bands OS dense and less-dense hematite, but with a small amount of silica remaining as scattered quartz. The other (Fig. 7) shows hematite bands that are undoubtedly original hematite bands from itabirite. They are still almost in their original relative positions, but the intervening silica bands have been leached away and little, if any, secondary iron has taken their place. As a result, we get a group of plaquettes.

Thus, of a number of hematite boulders examined, only two have been Sound which contain textures resembling the enriched ores of Rouessa. This may seem a meagre collection, but considering the distance of the O u m el Hbel conglomerate from the Rouessa-Tazadit ore bodies, it is surprising that any such textures at all have survived. Further search may produce more examples.

Occasional doubts have been raised as to the exact age of the O u m el Hbel conglomerate. Blanchot (1955) does not claim that it is of basal Cambrian age, but that it is a local facies of the Cambrian sandstone, and the author sees no adequate reason to question this judgement.

I J Imrn

FIG. 6. Oum el Hbel conglomerate hematite boulder; inicro- photo, incident light, showing alternating bands of primary

(massive) and secondary hematite, the latter less dense, with some residual silica (grey).

285

F. G. Percival

I mm

FIG. 7. Oum el Hbel conglomerate hematite boulder; micro- photo, incident light. Most of the silica has been leached from

Established conclusions regarding age of enrichment

The chief conclusions may now be summarized as follows. 1.

2.

3.

4.

5.

Although the hematite boulders of the conglomerate have not been analysed, it is quite obvious from their density and from the appearance of the sections and polished specimens examined that the hematite content is much higher than that of the local unaltered itabirites or of the local breccia. The breccia from which these boulders were derived had been enriched in iron, and this enrichment was markedly earlier than the age of the conglomerate, i.e. it was Precanibrian. One of the boulders contains plaquette fragments that are evidence of leaching of silica from itabirite. Thus leaching of silica had certainly commenced in Precam- brian times. One boulder shows a fragment with coarser and finer hematite banding, but this is not completely convincing because some of the silica remains. It would in any event be only rarely that enriched hematite pieces from Rouessa would be carried southwards over a distance of 7.5 km, but as only a small number of boulders have been examined further work may reveal more specimens of Rouessa hard ore type. The apparent absence of fragments of hematite in the breccia in general certainly suggests that iron enrichment was not well developed when the breccia was formed, but the age of breccia formation may have been

this fragment, leaving primary hematite bands collapsed as ‘plaquettes’. (hematite =white, quartz =grey, cavities=black.)

6.

coeval with the orogeny of the Kedia as a wholelong before Cambrian times. Thus the commencement of leaching and iron enrichment of the itabirite of the Kedia (and locally of the breccia) was at some time between the age of breccia formation and the age of deposition of the Cambrian conglomerate of O u m el Hbel. O n general grounds, as the uplift, folding and distortion of the itabirites of the Kedia were Precambrian, it is reasonable to expect that Precambrian surface waters would initiate the leaching of silica and the residual enrichment in iron, and boulders of the O u m el confirm this.

tce evidence of the hematite Hbel conglomerate tends to

Age of enrichment of deposits elsewhere

Evidence of the age of leaching and enrichment of itabirite is in some cases inconclusive and in other cases the age may only be indicated within wide limits.

Veriezrtela. Ruckmick (1963) published a study of the silica content of spring waters emerging along the lower flanks of the iron ore bodies of Cerro Bolivar, Venezuala, from which he deduced that ‘if the assumption is made that present climatic conditions have prevailed in the past’ the rate of removal of SiO, ‘suggest that the Cerro Bolivar ores have been developing for approximately 24 million years, or since the Oligocene’. Ruckmick emphasized that his

286


calculations ‘do not represent a precise dating method because of the assumptions involved, and because of several factors, such as rates of physical and chemical erosion at the surface, which are difficult to assess’. In spite of this warning, it may be considered that there is proof that leaching and enrichment at Cerro Bolivar date back only to the Oligocene. One should not assume, however, that the whole of the silica content of the Cerro Bolivar waters has come from the siliceous bands of the itabirite. As to the assumption that present climatic conditions have prevailed in the past, such an assumption would certainly not be justified in the area of our present survey in Mauritania for example. This is now a desert region, but deposits of laterite and canga on the Kedia indicate a former humid monsoon type of climate, and the very large numbers of neolithic and earlier artefacts found in the dried-up kale bed of the Sebkha d’Erguya, 20 ltm east of the Kedia, suggest that this humid climate prevailed within human times.

South Africa anù South- West Africa. The age of enrichment of the ores of Postmasburg is given as late Precam- brian by Boardnian (1952). Strauss (1952) is less certain about the ore of Thabazimbi, but considers it as possible that the ore may be Pre-Waterburg (late Precambrian) in age.

India. In 1931 the author (Percival, 1931) published photographs of an Indian conglomerate composed of pebbles of enriched hematite-a conglomerate :hat Dunn (1940) took as the basal conglomerate of his (Precambrian) Kolhan series. Occurrences of this conglomerate, locally with enriched heinatite pebbles, are widespread in the Singhbhum-Orissa iron ore field, as recorded by Percival and Spencer (1940), and there is no doubt that leaching of the banded hematite jasper, and enrichment in iron, had commenced in Precambrian times in this region. Locally the conglomerate forms a high-grade direct-shipping ore. Dunn (1940) states that enrichment ‘took place at any stage in the geological history of this area when conditions were suitable’, and concludes that ‘whenever and wherever these rocks were exposed at the surface, and subjected to circu- lating waters during their geological history, re-arrangement of Fe,O, was in progress’.

Western Aiistualiu. Discussing the ores of the Hamers- ley Province of Western Australia, Campana (1966) states that ‘the high grade iron mineralization must have been favoured by the Tertiary climatic cycle’ and adds more emphatically that this Tertiary cycle ‘brought about the Hamersley iron field’, but MacLeod (1966), discussing the same area, writes that though this process has operated on a limited scale, and may indeed be proceeding at the present time, the restriction of iron-formation to such a limited

time poses severe difficulties. ‘It would have been impossible to produce ore bodies of such grade and dimensions within these residuals after maturity of the surface for the reason that therewouldnot havebeenenoughiron available.’ MacLeod suggests as an alternative theory of origin that ‘the ores are of great antiquity and that their formation has been going on concurrently with the degradation of the land surface. . . . The ore bodies are, in effect, residues rich in iron, much of which has been derived from overlying iron-formation now removed by erosion’.

From this brief review of the published data on the ages of leaching and enrichment of itabirite to produce high-grade natural ores, it appears that further investigations are desirable to give greater precision than has yet been achieved.

Comment

The results of the present investigation are rather disappointing, as they do not establish the existence of masses or lenses of bedded enriched hematite at the time of formation of the Cambrian conglomerate, although proof that such masses existed may yet be found. All that is firmly established is that the processes of leaching and enrichment had commenced prior to the deposition of the O u m el Hbel conglomerate, and we may deduce that these processes have been continuously in operation since then, to form ore bodies that have a vertical depth below peak outcrop at Tazadit of more than 400 m, with diminishing thickness at this depth. At F’Derik, where the deposit appears to be entirely of the hard massive type, the ore has been proved by drilling to about 250 m below peak outcrop and, at this depth, the base of the enriched ore had not been reached.

The iron for the secondary enrichment was presumably derived from the weathering of overlying itabirite, as suggested by MacLeod for the Hamersley area, and with such depths of enriched ore a considerable cover of itabirite has been removed. These deep ore bodies are compatible with a long period of leaching, but the rate of this leaching would vary with climatic variation, both directly through temperature and humidity, and indirectly through their effect on vegetation and resultant humic acids.

Acknow1edgerneii.t

The author wishes to express his thanks to the S.A. des Mines de Fer de Mauritanie for their hospitality and assistance in his visits to the O u m el Hbel area, and for their consent to the publication of this material.

287

F. G. Percival

Résumé

Enrichissement des minerais zoné,s de fer de la Kedia d’Ic$ìZ, en Mauritunie (F. G. Percival)

La Kedia d’Idjil est un inselberg de roches précambriennes qui s’élève à quelque 500 mètres au-dessus de la plaine saha- rienne. Les minerais de fer de la Kedia s’étendent le long d’une zone d‘enrichissement de l’hématite dans les itabirites, qui s’allonge dans une direction sensiblement est-ouest le long de la crête de la chaîne sur une distance d’environ 24 km. Au sud de cette zone, la chaîne a une couverture de brèches formée de blocs fracturés (ou partiellement roulés) d‘itabirite. Des enrichissements mineurs de la brèche, d’âge incertain, se rencontrent localement.

Les itabirites ont des pendages raides vers le sud. A l’est, et recouvrant la bordure orientale de la Kedia, des grès cambriens ont un pendage très peu accentué vers l’est. A Oum el Hbel, sur la bordure sud-sud-est de la Kedia, on

rencontre un conglomérat formé esseniiellement de blocs roulés de la brèche, ce qu’on interprète comme un faciès local du Cambrien. Ce conglomérat contient un nombre modéré de blocs d‘hématite enrichie, mais à ce jour l’examen de la plupart d‘entre eux a révélé qu’ils avaient la contexture de la brèche et non celle régulièrement veinée, lixiviée et enrichie des masses d’hématites de la zone prin- cipale du minerai.

Cependant, il est évident que ces blocs ont été enrichis en fer avant d‘être incorporés dans le conglomérat cambrien, et certains des fragments qu’il contient mettent en évidence une lixiviation de la silice, avec formation d’hématite rési- duelle, ce qui conduit à une date précambrienne au moins pour le commencement de ces processus.

L’âge de la formation d‘hématites enrichies similaires qu’on rencontre ailleurs fait l’objet d‘une courte discussion.


AUDIBERT, J.; CARUEL, P.; CHOUBERSKY, A. 1964. Development of the Kedia d’Id$ orebodies (S.A. des mines de fer de Mauritanie) MIFERMA, Islamic Republic of Mauritania. Symposium on Opencast Mining, Quarrying andAlluvial Mining. London, Institution of Mining and Metallurgy (Paper no. 20), 34 p.

BALDWIN, A. B.; GROSS, W . H. 1967. Possible explanations for the localization of residual hematite ore on a Precambrian iron-formation. Econ. Geol., vol. 62, p. 95-108.

BLANCHOT, A. 1955. Le Précambrien de Mauritanie occidentale. Bulletin de la Direction fédérale des mines et de la géologie, no. 17. Dakar.

BLONDEL, F. 1952. Les gisements de fer de l’Afrique Occidentale française. Symposium sur les gisements de fer du monde. XX Congr. géol. int., Alger, t. I, p. 7-9.

BOARDMAN, L. G. 1952. Short description of the Postmasburg iron ore deposits. Symposium sur les gisements de fer du Monde. XIX Congr. géol. int., Alger, t. 1, p. 252-6.

CAMPANA, B. 1966. Stratigraphic-structural-palaeoclimatic controls of the newly discovered iron ore deposits of Western Australia. Mineralium Deposita, Berlin, vol. 1, no. 1, p. 53-9.

Dum, J. A. 1940. The stratigraphy of South Singhbhum. M e m . geol. Surv. India, LXIII, part 3, p. 303-69. GROSS, G. A. 1968. Geology of iron deposits in Canada.VoI. III Iron Ranges of the Labrador GeosyncZine. Ottawa, Geological Survey of Canada. 179 p. (Economic Geology Report No. 22.)

GRUNER, J. W. 1946. The minera ogy andgeology of the Taconites and iron ores of the Mesabi Range, Minnesota. p. 1-127. St. Paul, Minn., Commissioner, Iron Range Resources and Rehabilitation.

GUILD, P. W. 1953. Iron Deposits of the Congonhas District, Minas Gerais, Brazil. Econ. Geol., vol. 48, p. 639-76.

- , 1957. Geology and mineral resources of the Congonhas District, Minas Gerais, Brazil, Prof. Pap. U.S. geol. Suvv., no. 290, p. 1-90. HUVELIN, P. 1963. Précisions sur la genèse de la brèche d‘Idjil (Fort Gouraud, Mauritanie). Chronologie de la formation des minerais de fer et de la brèche. Bull. Soc. géol. Fu., t. IV, no. 2, p. 322-8.

LETHBRJDGE,’ R. F.; PERCNAL, F. G. 1954. Iron deposits at Fort Gouraud, Mauritania, French West Africa. Trans. Instn. Min. Metall., Lond., vol. 63, p. 285-98.

MACLEOD, W. N. 1966. The geology and iron deposits of the Hamersley Range area, Western Australia, Bull. geol. Surv. W. Airst., no. 117, p. 1-170.

PERCIVAL, F. G, 1931. The iron-ores of Noamundi. Trans. Min. geol. Inst. India, vol. 26, p. 169-271.

-. 1967. Possible explanations for the localization of residual hematite ore on a Precambrian iron-formation. Discussion. Econ. Geol., vol. 62, p. 739-42.

PERCIVAL, F. G.; SPENCER, E. 1940. Conglomerates and lavas in the Singhbhum-Orissa iron ore series. Trans. Min. geol. Inst. India, vol. 35, p. 343-63.

ROYCE, S. 1948. Discussion on paper by Roberts, H. M. and Bertley, M . W.-Replacement hematite deposits, Steep Rock Lake, Ontario, Trans. Anier. Inst. min. (metall.) Engrs., vol. 178, Mining Geology, p. 387.

RUCKMICK, J. C. 1963. The iron ores of Cerro Bolivar, Vene- zuela, Econ. Geol., vol. 58, p. 218-36.

STRAUSS, C. A. 1952. The deposits mined by the South African Iron and Steel Industrial Corporation Limited. Symposium sur les gisements de fer du monde, XIX Congr. géol. int., Alger, t. 1, p. 241-52.

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Discussion

R. P. PETROV. Are the terms ‘magnetite-hematite quartzite’ and ‘itabirite’ synonymous? What is the difference between itabirite and taconite?

F. G. PERCIVAL. There is great need for agreed definitions of itabirite, jaspilite, taconite, and also canga. I shall be very glad if some agreed definitions can be an outcome of this symposium.

S. J. SIMS. Is there any structural control of the localization of the high-grade hard hematite?

F. G. PERCIVAL. There is no separate structural control for the hard high-gracie hematite. All three types-hard, plaquette and powdery-occur together at Rouessa and at Tazadit.

S. J. SIMS. D o you consider specularite to be of supergene origin?

G. CHOUBERT. If I remember correctly, the sandstones around the southern part of Kedia d’Idjil are not Cambrian, but older, approximately Upper Precambrian. They are older than the stromatolite limestones of Atar-Hanck, which are about 700-900 m.y. old (by determination of M m e Bertrand and M m e Raber).

F. G. PERCIVAL. Yes, I wrote that there is some doubt as to the exact age, but this still leaves the date of leaching as Precambrian.

G. CHOUBERT. What is the relationship betwzen the Kedia d‘Idjil quartzites and the ancient granites on the north? D o the granites break through the quartzites, or do the quartzites overlie the transgressive granites?

F. G. PERCIVAL. N o boreholes have been made to reveal the junction of the Kedia iron-formation and the granites and I have no certain information on this.

F. G. PERCIVAL. Yes, but the supergene waters may locally have been heated by intrusive veins of dolerite that cut the ore bodies.

289

Iron ores of the Hamersley Iron Province, West ern Australia

W. N. MacLeod Carpentaria Exploration Company Pty Ltd, West Perth

Introduction Systematic exploration of the iron ore potential of the Hamersley Range area of Western Australia commenced early in 1961. The region had long been recognized as one in which iron deposits were likely to occur as banded iron- formations of great thickness and lateral extent had been recorded during the early geological reconnaissance surveys of the region.

Production and export of iron ore commenced in 1966 from the Mount Tom Price deposit, operated by Hamersley Iron Pty Ltd. This was followed by the initiation of production from the Mt Whaleback deposit in 1969 by the Mount Newman Mining Co. The productive capacity of the operating mines is now in excess of 30 million tons per annum. Reserves of ore containing more than 50 per cent iron are estimated to amount to 18,000 million tons of which over half is comprised of hematite-goethite ore containing between 58 per cent and 65 per cent iron.

Physical features

The Hamersley Iron Province is situated in the North-West Division of Western Australia, about 1,000 km north of Perth. The region is mountainous and arid and was virtually uninhabited prior to the commencement of iron ore mining.

The iron province is defiled by the extent of the Pre- cambrian Hamersley Group of sedimentary and volcanic rocks. This stratigraphic unit was originally deposited in a discrete sedimentary basin at least 85,000 km2 in area. The distinctive lithology of the group is matched by an equally distinctive topography. The resistant banded iron-formations, totalling over 1,000 m thick and separated by softer dolomite and shale beds, weather out to bold sinuous ridges and plateaux rising up to 500 m above the intervening valleys.

The majority of hill summits in the region are gently domed and concordant in level. These domes and plateaux are remnants of an older land surface, possibly of early

Tertiary age. Stream rejuvenation, with a new erosion cycle commenced in later Tertiary time and has continued to the present day. This appears to have been initiated by a major regional upward of the entire Hamersley block. Most of the present rivers flow in valleys with steep-walled gorges in the lower parts and gentle upper slopes upon which are preserved deposits of older detrital material.

This ancient profile is of considreable economic significance. Many of the large hematite deposits are found in erosion residuals of this older surface at higher levels. In the valleys conglomeratic ores and pisolitic limonite deposits form part of the lower ancient profile and also remain as erosion residuals.

Stratigraphy and lithology

The Hamersley Group of Lower Proterozoic sediments and lavas was deposited in an ovoid basin about 500 km long and 250 km wide. The group forms part of a major stratigraphic unit referred to as the Mount Bruce Supergroup which is a conformable succession of sediments and lavas about 10,000 m thick. Radiometric age dating indicates that the Mt Bruce Supergroup was deposited between 2,200 and 1,800 m.y. ago. It is subdivided as follows.

TOP Wyloo Group (3,100 m) Hamersley Group (2,700 m) Fortescue Group (4,200 m)

Banded iron-formation, chert and dolomite are the principal constituents of the Hamersley Group. No coarse clastic sediments have been recorded from the succession. The chemical sediments have been intruded by thick dolerite sills and are interbedded with acid lavas in the upper part. Banded iron-formations constitute over one-third of the total thickness of the group and these are the host and source rocks of the iron ore deposits.

The sedimentary units of the Hamersley Group are remarkably persistent in lithology and thickness throughout the entire iron province.

Unesco, 1973. Genesis of Precainbriun iron und ~nuizguiiese deposits. Proc. Kiev Synzp., 1970. (Earth sciences, 9.) 291

W. N. MacLeod

Table 1 illustrates the stratigraphical subdivision of into a broad synclinorium with average dips of less than 5". the group. There are local zones within this broad structure with

strong,flexuring and faulting, but such features die out over TABLE 1. Subdivision of the Hamersley Group , short distances.

The central zone of the basin is much more strongly folded and the topographic forms faithfully reflect the

Formation Lithology (metres) major structural units. Two superimposed fold trends can be recognized, lying almost at right angles, and the combi-

TOP nation of these has produced a striking echelon pattern of Boolgeeda Iron Formation Iron-formation, domes and basins. The limbs of the major structures in this

part of the province dip between 30" and GO", but complex

Thickness

ferruginous shale 220 - Woongarra Volcanics Rhyolite flows, tuffs,

Weeli Wolli Formation Iron-formation, shale,

Brockman Iron Formation Iron-formation, chert,

Mt McRae Shale Shale, siltstone,

Mt Sylvia Formation Iron-formation, shale, chert 40

Wittenoom Dolomite Dolomite, shale, chert 150 Marra Mamba Iron Iron-formation, chert, Formation shale 200 Base

iron-formation 600

dolerite sills 550

, shale, dolomite 675

dolomite, chert 1 O0

The iron ore deposits are associated with, and derived from, the three principal iron-formations, the Marra Mamba, Brockman and Boolgeeda. Of these, the Brockman Iron Formation is by far the most important and is the host and source rock for all the largest known deposits in the province.

The Brockman Iron Formation has been subdivided into four members as follows (Trendall and Blockley, 1970).

TOP Yandicoogina Shale Member (100 m) Joffre Member (360 m) Whaleback Shale Member (30 m) Dales Gorge Member (180 m)

It was early recognized during the initial exploration of the province that most of the major hematite deposits occurred within the Dales Gorge Member. Further work has shown that the Joffre Member can also be an important host for ore under optimum structural conditions.

Structure of the Hamersley Iron Province

The Hamersley Group sediments occupy an intermediate position between the generally flat-lying rocks of the For- tescue Group and the quite strongly folded Wyloo Group on the southern side. The northern half of the Hamersley Basin is virtually undisturbed with the rocks gently warped

secondary and tertiary fold patterns have been superimposed on these major structures, particularly in fold axial zones and adjacent to major fault zones. It is these minor structures that provide the most important loci for iron ore deposits within the province. Practically all the major ore bodies occur within this zone of moderate to strong folding in or adjacent to synclinal troughs in the parent banded iron-formation. In the extreme southern and western parts of the province tectonic forces have been at their strongest, with the development of major faults and tectonic slides which disrupt the normal stratigraphy.

Within the basin there is a complete range from virtually flat-lying, undisturbed rocks to those which have been overturned and isoclinally folded. Despite the variations in intensity of folding the grade of metamorphism remains very low, and these iron-formations provide one of the best examples in the world of unmetamorphosed banded iron-formation.

Iron ore deposits

Four principal types of iron ore have been recognized in the Hamersley Iron Province. In a sense these ore types are transitional and reflect differing degrees of the protracted processes of enrichment of the parent banded iron-formation and the differing environments in which the enrichment occurred. The ore types are summarized as follows.

Hard mussive hematite oye (blue ore). Iron content usually greater than 64 per cent rising as high as 69 per cent. Phosphorus generally less than 0.05 per cent and silica and alumina less than 1 per cent. Low content of combined water. Hematite in randomly oriented small bladed crystals is the dominant mineral with only a minor content of goethite.

Bunded hematite-goethite oye. Laminated texture of the original iron-formation is well preserved with alternating bands of variable hematite-goethite ratio. Sometimes porous with interlaminate cavities. Iron content within the range of 58-64 per cent iron, phosphorus up to 0.15 per cent, but averaging about 0.10 per cent. Silica and alumina up to 3 per cent and combined water content as high as G per cent. This is the commonest ore type in the province and occurs in intimate admixture with, and is transitional to, the hard massive hematite.

Coriglomevutìc ores. A cemented scree of fragments of massive hard hematite in a matrix of friable goethite and

292

Iron ores of the Hainersley Iron Province, Western Australia

limonite. Usually occurs in association with the types described above and is intermediate in composition.

Pisolific limonite oye. Confined to older drainage channels as terraces and mesas. Complex mineralogy with hematite, maghemite, goethite and limonite. Contains 50-60 per cent iron with low phosphorus and high water content of 10-12 per cent.

Hematite and hematite-goethite ores

GENERAL CHARACTERISTICS

The hematite and hematite-goethite ores occur within the banded iron-formation of the Boolgeeda, Brockman and Marra Mamba Iron Formations and minor local enrichments have been recorded in the thinner iron-formations of the Weeli Wolli Formation. U p to the present time, all deposits of economic significance occur in the Brockman Iron Forniation and particularly, although not exclusively, in the Dales Gorge Member of this formation.

The three principal iron-formations are generally similar in over-all mineralogical composition (Table 2), with an original iron content of about 30 per cent, with the iron mainly in the form of magnetite. Silica and carbonate minerals comprise most of the remainder of the rocks with accompanying minor amounts of iron-bearing phyllosilicates, notably stilpnomelane and amphiboles, including riebeckite and crocidolite.

The hematite-goethite ore bodies have originated by processes of supergene enrichment of this primary iron- formation material and, in the province as a whole, all stages of this enrichinent can be observed. Probably all iron- formation exposed at or near the present surface is altered and enriched to some degree. The essentially transitional character of the process is observable both from the examination of outcrops and from the chemical data provided by the analysis of drill core. It is postulated that ground and meteoric waters are the principal agents of enrichment of the parent iron-formation and that the principal processes involved are the selective removal in solution of most of the silica and carbonate and the concurrent redistribution and recrystallization of iron oxides within the system.

Hematite-goethite ore zones occur in practically all areas where the Brockman Iron Formation is exposed. These range in extent from a few hectares to several square kilo-

metres and the thickness of the ore from a thin surface crust of a few metres thick to deep trough-like occurrences between 100 and 300 m thick.

DISTRIBUTION A N D M O D E OF OCCURRENCE

The largest hematite ore bodies are found along the major central axis of the elliptical Hamersley Basin and to the south of this axis. This distribution is clearly related to the stronger folding and faulting of the iron-formations in the southern half of the basin. In the northern half, that is in the main Hamersley Range Synclinorium, ore occurrences are restricted in size and grade in unfolded areas, but appear in areas of localized flexuring along the flanks and troughs of synclines. In some areas there is an extensive development of platy hematite-goethite ore on remnants of the older surface even where the rocks are unfolded. Such ore zones may contain a very substantial tonnage of enriched material but the grade is variable and such deposits would present more mining difficulties than the deeper and more compact ore bodies now being worked.

The principal areas of ore deposits are in the Mt Brock- man and Mt Turner Synclines, the Weeli Wolli Anticline, the Parraburdoo Range and the Ophthalmia Range and its satellite hills. These areas account for over 80 per cent oî the immense reserves of hematite ore known in the province and will doubtless continue as the main productive areas for many decades to come.

Genesis of the hematite ores

The hematite-goethite ore bodies of the Hamersley Iron Province appear to have originated as a result of the enrichment of the parent banded iron-formation under supergene conditions. This is common to all hematite deposits in the province. Such variations as do exist in grade, ore texture and size of the deposits can be attributed to variations in the degree to which these supergene processes have operated.

There is no convincing evidence to suggest that the ores represent primary depositional concentrations. A hypogene origin seems equally improbable as the iron-formations surrounding the deposits are unmetamorphosed and have not been affected by igneous intrusions to any degree. 'The

TABLE 2. Compositional range of iron ores from the Hamersley Iron Province (in percentages)

Brockman Iron Platy leached Compact banded Massive blue Conglomerate Pisolitic limonite Formation ore ore ore ore ore

Fe 25-40 55-60 58-63 66-69 60-64 52-60 P 0.05 0.05-0.15 0.05-0.1 5 0.02-0.05 0.05-0.10 0.02-0.06 SiO, 40-55 6-8 2-5 0.1-0.5 2-4 4-10 A120, 1.5-3.5 2-3 0.1-0.3 2-3 2-3 1-3 S Less than 0.5 0.01-0.05 0.01-0.05 0.01-0.05 0.01-0.05 0.01-0.05

293

W. N. MacLeod

Hamersley Group rocks are not intruded by granite and although there are many doleritic sills and dykes in the sediments, these have limited contact metamorphic effects only and cannot be regarded as agents for the large scale metasomatic transformations involved in the formations of these immense iron ore bodies. Meteoric and ground waters, operating over a long period of time, are believed to be the principal agents for the transformation and enrichment of the parent iron-formation. These agents have operated most effectively under certain repetitive structural controls.

From the textural and mineralogical characteristics of the ores it can be inferred that the following processes have operated during enrichment. 1. Leaching of silica and carbonates from the protore, leav-

ing an enriched residual product containing between 50 per cent and 60 per cent iron. The original magnetite is oxidized to hematite and some is hydrated to produce goethite and other hydrous iron oxides.

2. Infilling of the cavities and planar voids in the leached ore by reprecipitated iron oxides, principally goethite, to produce a more compact ore type in which remnants of the original bedding are still clearly discernible. The compact ore commonly contains between 60 per cent and 64 per cent iron; silica is lower than the leached cavernous ore, but there is little difference in phosphorus content.

3. Recrystallization of the iron oxides and dehydration to produce a blue-grey massive hematite ore which is virtually structureless. The material approaches the theoretical composition of pure hematite (69.94 per cent). Silica and alumina together amount to less than 1 per cent and the content of combined water has dropped to less than 0.5 per cent. The massive blue ore appears to represent the culmi-

nation of the process of enrichment. It is less commonly seen close to the surface and seems to be most abundant in the deeper parts of well-defined synclinal troughs, as at Mt Whaleback.

The observed top of the ore is usually a remnant of the old surface and commonly the ore extends downwards from this surface to the base of the Brockman Iron For- mation against the underlying Mt McRae Shale.

This consistent relationship between the ore bodies and the old surface led earlier workers in the province to the belief that the hematite deposits were developed after, or in a late stage of, the attainment of maturity of this surface (Campana et al., 1964; MacLeod, 1966). That this process has operated on a limited scale seems undeniable. Many of the shallow hematite crusts on all the major iron-formations probably have resulted from late stage surface enrichment in relation to the old land surface and could be classed essentially as laterites. However, to restrict the main period of ore formation to a limited period during the maturity of the old surface poses some severe difficulties when applied to the formation of all major ore deposits in the province.

During the earlier erosion cycle, dissection of the Hamersley Group sediments had proceeded far enough to leave many thin and isolated residuals of the Brockman

Iron Formation as erosion residuals in synclinal cores. It is clear that segmentation of the iron-formations by erosion was well advanced prior to the attainment of maturity of the Hamersley surface and well before the initiation of the present erosion cycle.

Many of these isolated residuals provide some of the largest hematite deposits known in the province and are the sites of an almost wholesale transformation of Brockman Iron Formation into hematite-goethite ore. It would have been impossible to produce ore bodies of such grade and dimensions within these residuals after maturity of the surface as there would not have been enough iron available. The erosion residuals would have become closed systems and the scale of iron enrichment and concentration that has occurred could be expected to leave profound and observable changes in the surrounding iron-formation. Such changes are not seen and, in fact, most of the iron-formation existing beside the ore bodies in the residuals is itself enriched in iron to some degree.

An alternative theory of origin, which accords better with the observed characteristics and distribution of the ore bodies, is based on the proposition that the ores are of great antiquity and that their formation has been going on concurrently with the degradation of the land surface, and particularly with the continuing erosion of the Brockman Iron Formation. In short, the present ore zones represent the end products of a long continued process of physical breakdown by erosion and chemical concentration of iron in the original iron-formation. The ore bodies are, in effect, residues rich in iron much of which has been derived from formerly overlying iron-formation, now removed by erosion.

The progressive downward enrichment of the iron- formation has been effected by the percolation of waters to the water table from the time when the iron-formation was ñrst exposed by removal of the overlying cover of younger rocks. Rainwaters falling on the exposed iron- formation dissolve minute quantities of iron and silica, which are carried down to near the water table. At the water table the iron is oxidized and precipitated in the ferric state but silica remains in solution. The zone of seasonal fluctuation of the water table thus becomes secondarily enriched in iroii. In folded areas and where there are strong fault zones, the movements of the groundwaters are canalized in impounding structures such as synclinal troughs or zones of intense drag folding on the limbs of major folds. It is in such sections that the process of enrichment has culminated and these are the loci of the largest and highest grade ore bodies in the province.

In Brazil Dorr (1964) observed that the maximum degree of supergene enrichment of itabirite was to a product containing 63 per cent iron. From the analysis of a large tonnage of ore he placed an upper limit of 65.5 per cent iron for enrichment by purely supergene processes. H e concluded that hematite cannot have formed directly by supergene enrichment and that an appreciable amount of goethite must remain.

These conclusions are applicable to many ore zones

294

Iron ores of the Hamersley Iron Province, Western Australia

in the Hamersley Iron Province, but in the two large ore bodies now being mined the process of enrichment has proceeded well beyond the upper limit set by Dorr. Both contain a substantial proportion of a massive very pure hematite close to the theoretical composition and containing minimal amounts of goethite, silica, alumina and combined water.

The pure hematite ore is closely intermingled with, and often transitional into, lower grade hematite-goethite ores so that it is difficult to dissociate the pure ore from the same pervasive cycle of enrichment that has produced the hematite-goethite ore and, at an earlier stage, the partially enriched iron-formation.

To achieve such purity it seems necessary to invoke metasomatic replacement of pre-existing ore and direct crystallization of hematite. Microscopic examination provides evidence that this has actually occurred, as the purer massive ore is seen to consist of plates and needles of hematite to the virtual exclusion of the martite octahedra which are so abundant in the hematite-goethite ores. These plates and needles appear to be a late stage crystallization and it is perhaps of some signiñcance that the pure massive ore mainly occurs at depth. In the case of the Mt Whaleback and Mt Tom Price deposits these depths are of the order of between 100 and 200 m and, of course, these depths may have been much greater at the time when the material was formed. The recrystallization of hematite may be a temperature effect dependent on depth.

A n implication of the theory of genesis by progressive downward percolation of iron rich groundwaters and concurrent erosional reduction of thickness of the iron- formation, is that there may be substantial deposits of hematite in synclinal troughs which have not yet been eroded sufficiently deeply to expose the full extent of the zone of basal enrichment. The occurrence of very high grade ore at depths of over 300 m in the Mt Whaleback deposit is probably not a unique occurrence, although in that particular case the folding is strong and very deep troughs exist in the parent iron-formation.

Enrichment of the iron-formation is probably influenced by variations in permeability. Folded and faulted areas, in addition to providing definite structures for the concentration and channelling of water movement, probably offer increased freedom of water movement through a host of minute fractures in the minor folds and faults. Minor folding extends down to microscopic amplitudes and such violently crenulated zones are common sites of ore formation. There is a frequent occurrence of major ore bodies near strong faults, as at Mt Whaleback and Parraburdoo.

Pisolitic limonite ores

DISTRIBUTION

The pisolitic limonite ore occurs as residual cappings of rnesaform hills and flat-topped ridges aligned along the

courses of fossil drainage channels. In some areas it occurs as broad valley-fill deposits at the confluence of internal drainage systems.

This ore type is represented in practically every drainage system in the province which in the past has drained extensive areas of the Brockman and Marra Mamba Iron Formations. There can be little question that these highly ferruginous sediments are the ultimate source of the iron in the secondary riverine and lacustrine accumulations of the pisolitic limonite. The maximum development of this ore type is on the western side of the iron province in the Robe River and Duck Creek drainage systems.

The grade of the pisolitic limonite (Table 2) ranges within the limits of 40-60 per cent iron, with the bulk of the material falling within the range of 52-58 per cent. Drilling has shown that there are mineable sections within many of the deposits with an average grade of 55-58 per cent iron. This richer ore has a total silica and alumina content well below 10 per cent and a combined water content of about 10-12 per cent. Phosphorus and sulphur are normally less than 0.05 per cent and titania less than 0.20 per cent.

The resources of this ore type in the Hamersley Iron Province are estimated to be at least 6,000 million tons. U p to the present it has not been utilized, although it has been established that the ore can be pelletized to a product containing between 63 per cent and 65 per cent iron. It constitutes an important reserve for the future, but at present suffers in competition with direct shipping higher grade lump hematitic ores from elsewhere in the province.

TEXTURE AND MINERALOGY

The iron oxides present in the pisolitic ore include amorphous isotropic limonite, goethite, hematite and maghemite. These minerals occur as components of the pisoliths and in the matrix between the pisoliths. The pisolitic texture of the ore is its most striking feature, both in hand specimens and under the microscope. The majority of the pisoliths range between 1 and 3 mm in diameter, exceptionally attaining a size of 5 mm, and ranging down to diameters of 0.1 mm or less. Many of the larger pisoliths are made up of aggregates of smaller pisoliths.

The matrix between the pisoliths usually consist of colloform isotropic limonite which is yellow to brown and often soft and ochrous, or black lustrous goethite. In some zones the limonitic matrix is continuous and imparts to the ore a vitreous lustre and sub-conchoidal fracture; in other zones there is a high porosity due to discontinuity of the matrix and such material is less coherent. Cavities and pore spaces are often partially filled with opaline silica and travertine.

The pisoliths themselves are greatly diversified in colour, texture and mineralogical composition. Many are soft and ochrous and can be scratched and disintegrated

295

W. N. MacLeod

with the fingernail whereas others are hard and metallic. Under the microscope most pisoliths are seen to be composite bodies with pronounced concentric layering which reflects successive stages of accretion of iron oxides. The majority of pisoliths are cored with hematite and around these cores there is a succession of thin, alternating shells of goethite, hematite and maghemite. Harms and Mor- gan (1964) state that the proportion of hematite in the pisoliths decreases with increasing distance from the source iron-formations, aiid in the Robe River deposits the over- all hematite content ranges from nil to 20 per cent.

Some zones of the ore contain a high content of fossil wood fragments, now completely replaced by hematite or limonite with good preservation of the original cellular structure.

Clay beds and lenses are commonly interbedded with the ore and a common feature is the appearance of clay- filled joints and pipes. Many of the deposits have a pronounced vertical jointing which is attributed to shrinkage during diagenesis and these joints have permitted the ingress of fine clays from above.

GENESIS OF THE PISOLITIC LIMONITE ORES

It seems a reasonable assumption that the deposits have a common mode of origin and all were formed at or about the same time. The occurrence of identical deposits in association with Archaean iron-formations elsewhere in the region, which are of similar age and geomorphological relationship to those derived from the Proterozoic sediments suggests that conditions favouring this type of iron ore accumulation were widespread.

From the consistent distribution pattern of these deposits in relation to long established drainage systems emanating from plateaux and ranges of iron-formation, there can be little doubt that these rocks comprise the ultimate source of the iron in these secondary riverine deposits. Theories differ, however, as to the mechanisms involved in the transport of the iron and of the processes which have effected the deposition of the limonite.

These concepts can be summarized as follows. 1. Direct clzemical precipitation as bog iron ore. This

theory suggests that the iron accumulations are the product of chemical precipitation of iron directly from iron-charged solutions which have moved into the drainage channels, Chemical conditions in the channels have been such as to favour rapid deposition of iron.

2. Replacement aiid desilicatiotz of iron-formation detritus in the river channels by iron charged solution. This theory favours the transport of some iron into solution into the drainage channels but suggests that the limonite formation has been effected within a matrix of iron- rich detritus which itself makes the major contribution to the final volume of the deposit.

3. Clastic accumulation of ferruginous detritus derived by weathering oJ the enriched iron-jormution. This theory is a derivation of the previous one, but differs in its sugges-

tion that much of the iron enrichment and hydration of the iron oxides has occurred as a result of weathering of the iron-formations in situ. Erosion has stripped weathered profile of ferruginous materials and deposited them in low gradient stretches of the drainage channels in the similar fashion to the deposition of a placer deposit.

All three theories are in part complementary, and all have some measure of credibility as an explanation for all or some of the features of the deposits. However, it is felt that no single process can account for all the features of the limonite deposits as observed in the province as a whole. In all likelihood, all three processes outlined above, and possibly others not yet visualized, have operated to different degrees in different parts of the region.

It is felt that the most important clue as to the genesis of the pisolitic limonite is provided by the geomorphological relationships of the deposits. As discussed previously, the most outstanding geomorphological features of the iron province are the widespread remnants of an earlier mature landscape of gently domed hills separated by broad valleys in which there are thick accumulations of colluvium. Iron-formation fragments and chert are practically the sole constituents of these colluvial deposits.

In many localities in the headwater sections of the drainages an intimate and apparently transitional relationship between the pisolitic limonite and iron-formation scree can be observed. Development of the limonite only seems to occur on a major scale where there has been a long persistent drainage system. The limonite is less abundant where the drainage has been meandering or diffuse.

The intimate relationship between the iron-formation scree and the pisolitic limonite is observable in the terraced deposits of the older colluvium in the headwater sections of the rivers. The gently sloping terraces merge upwards with the gently domed hills of Brockman Iron Formation on the one hand, and on the other merge outwards with mesaform remnants of pisolitic limonite in the drainage channels. In the upper sections of the terraces, away from the river, there is an abundance of cemented iron-formation scree locally transitional into hematite conglomerate or canga ore. On passing towards the centre zone of the drainage channel, there is commonly a zone of complex intermingling of pisolitic and conglomerate materials which are physically contiguous with, and directly gradational into, pisolitic limonite. This limonite is essentially the same material as that found well downstream in the mesa deposits beyond the hills.

It can be postulated that the pisolitic limonite in the river channels, and the intermingled zones of pisolite and conglomerate in the montane areas, are both manifestations of the same protracted cycle of transformation of ironstone scree to pisolitic limonite that has proceeded concurrently with the evolution of the old land surface.

As the landscape has been progressively degraded, the iron-formation detritus has been continually fed into persisting drainage systems. The movement of ground and surface waters towards and along these channels has

29 6

Iron ores of the Hainersley Iron Province, Western Australia

effected a progressive transformation and enrichment of the detritus resulting in an equilibrium end product of pisolitic limonite.

The transformation of the detritus would involve the introduction of iron and the selective removal of the bulk of the silica, and is essentially the same chemical transformation as is envisaged for the formation of the hematite ore zones. The drainage channels can be regarded as fulfilling the same canalizing roles as synclinal troughs in the formation of in situ hematite deposits. This concept implies a cogenesis of the hematite and pisolitic ores; both have originated by the enrichment of banded iron-formation and the agents of enrichment have been ground and surface waters operating over a long period of time within restricted spatial limits.

The process is not regarded as being initiated or dominated by a particular set of climatic or geomorphic controls. Rather it is a long continued progressive cycle of erosion, enrichment and downstream movement of the products of enrichment. For this reason, and for the lack of stratigraphic evidence, no definite age can be offered for the pisolitic limonite deposits. The process probably commenced when there were sufficiently large areas of iron-formation exposed to erosional processes to provide a predominance of iron-rich scree to the drainage channels and continued until the rejuvenation of drainage in the current erosion cycle. Most rivers since then have established new base levels leaving the pisolitic ores, for the most part, well above water table.

Conglomerate ores The conglomerate ores commonly occur in association with the large hematite-goethite deposits as scree mantles on the lower slopes of the hills. Some conglomerates are simply cemented iron ore scree shed from the iron ore zones, whereas others appear to have a more complex origin and appear to be the product of enrichment of iron- formation scree by essentially the same processes of leaching of silica and iron migration as has occurred in the in situ hematite deposits.

In the upper terraces of many of the river valleys there is an almost complete transition from pisolitic limonite in the lower part of the valley to conglomerate ore consisting of angular fragments of almost pure hematite set in a pisolitic limonite matrix. Most of the conglomerate ores occur in this terraced fashion as erosion residuals on mature valley terraces.

The conglomerate ores have been assessed in some areas and appear to be intermediate in composition between the hematite-goethite ores and the pisolitic limonite (Table 2). Where they occur in large quantity they are an attractive mining proposition as crushing costs are low and the material is upgraded several per cent on crushing with loss of limonite fines.

Résumé

Les minerais de jeu d’Harneusley , en Acwtualie occidentale (W. N. MacLeod)

La province de Hamersley, en Australie occidentale, contient des réserves estimées à 17 milliards de tonnes de minerai d’hématite-goethite réparties entre quelques centaines de zones de minerai sur une surface qui couvre approximativement 85 O00 m z. La capacité de production actuelle de la province est de l’ordre de 35 millions de tonnes par an et pourra dépasser 50 millions de tonnes vers 1972.

Les minerais de fer sont rattachés à trois formations de fer, rubanées, épaisses et étendues, totalisant 1 O00 mètres en épaisseur à l’intérieur du groupe d’Hamersley du Pro- térozoïque inférieur. Parmi les gisements de minerais, celui qui est le plus important et dont la teneur est la plus élevée se rencontre dans la formation de fer de Brockman, d’une épaisseur d‘environ 650 mètres, et surtout dans la partie inférieure de cette formation appelée Dales Gorge Member et épaisse d’environ 180 mètres.

On pense que les minerais ont eu leur origine dans l’enrichissement supergène de ces formations de fer sous l’influence des eaux souterraines et météoriques. Les for-

mations de fer sont fortement plissées localement, mais ne sont pas métamorphosées et n’ont pas été affectées par les intrusions ignées de façon perceptible.

Les formations de fer apparentées contiennent entre 30 et 40 % d‘oxyde de fer - surtout de la magnétite, le reste de la roche étant constitué par de la silice sous forme de silex - et des carbonates minéraux, en particulier de la dolomite ferreuse. L’enrichissement s’est produit à la suite de processus complexes mettant en jeu l’oxydation et l’hydratation des oxydes de fer, la lixiviation sélective et le déplacement de la silice et des carbonates et le dépla- cement et la redéposition d’oxyde de fer à des niveaux inférieurs dans le système. Les dépôts de minerai de fer sont considérés comme des résidus altérés chimiquement, en équilibre avec l’environnement présent, qui représentent le produit final de la décomposition physique et chimique prolongée des couches de la formation de fer originale. On considère que l’enrichissement s’est produit depuis que les formations de fer sont entrées pour la première fois dans le circuit des agents météoriques, et, bien qu’influencé par les changements de climat, le processus général n’a aucune relation avec aucun cycle climatique.

Dans certains dépôts, il y a quelque évidence d’une

297

Iron ores of the Hamersley Iron Province, Western Australia

cristallisation ultérieure du matériel enrichi aboutissant à un produit ñnal d‘une composition proche de la composition théorique de l’hématite pure. Certains sont en faveur d‘une origine épigénétique pour ce type de minerai.

Les minerais d’hématite-goethite se sont plus favo- rablement développés dans les régions de plissement mo- déré ou violent ou au voisinage des zones de grandes failles. Les fosses synclinales sont un moyen de contrôle de la structure.

De plus, la province contient de grands dépôts d’un minerai secondaire limonitique et pisolitique qu’on trouve

essentiellement dans les anciens canaux de drainage. On n’est pas fixé sur l’origine de ce type de minerai, mais il y a quelque raison de penser qu’il a une double origine : dépôt direct des oxydes de fer en solution et enrichissement par les eaux souterraines des dépôts considérables de détritus riches en fer accumulés dans l’ancien système hydrographique descendant des monts Hamersley.

Des minerais de conglomérat formés par la recimen- tation des talus d’hématite sur les pentes en dessous des grands dépôts d’hématite constituent une réserve appré- ciable de minerai.


CAMPANA, B. et al. 1964. Discovery of the Hamersley Iron

DORR, J. VAN N. 1964. Supergene iron ores of Minas Gerais,

HARMS. J. E.; MORGAN, B. D. 1964. Pisolitic limonite deposits

MACLEOD, W. N. 1966. The geology and iron deposits of the deposits. Puoc. Aust. Inst. Min. Engrs., no. 210. Hamersley Range area, Western Australia. Bull. W. Aust.

geol. Surv., no. 117. Brazil. Econ. Geol., vol. 59, no. 7. TRENDALL, A. F.; BLOCKLEY, J. G. 1970. Iron-formations of

the Precambrian Hamersley Group of Western Australia with in Northwest Australia. Proc. Aust. Inst. Min. Engvs., special reference to crocidolite. BuII. W. Aust. geol. Surv., no. 212. no. 119.

298

Significance of carbon isotope variations in carbonates from the Biwabik Iron Formation, Minnesota E. C. Perry Jr and F. C. Tan Department of Geology, University of Minnesota

Introduction Most carbonates in a suite of 3-1 x lo9 year-old rocks from southern Africa and the Canadian shield have SC13 values1 within J. 2 per mil of PDB, an isotopic standard described by Urey et cil. (1951) which has a value typical of Phanerozoic marine carbonates (Perry and Tan, 1970 and in preparation). Carbon isotopic determinations on Precambrian carbonates by Hoering (1967) also give a range of values that is normal for Phanerozoic marine sediments.

Carbonates from iron-formation have E l 3 values lower than PDB by as much as 18 per mil. The anomaly was first reported by Becker and Clayton (1970), who found that SC13 of limestones and dolomites stratigraphically above and below the Dales Gorge Member of the 1.9 x lo9 year-old Brockman Iron Formation of Western Australia is similar to that of normal Phanerozoic carbonates, but that õC13 of iron carbonates within the iron- formation is about 10-15 per mil lower. They concluded that : ‘(1) the carbon isotope ratio in the world oceans has been nearly constant for at least 2 x lo9 years; (2) the banded iron-formation of West Australia was deposited in a basin distinct from but with some connection to the ocean; (3) there probably was organic activity involved in either the transport of iron to the basin or in the precipitation of iron and silica. However, the possibility of juvenile os recycled carbon from volcanic sources sup- plying the lighter carbon, rather than organic activity, cannot be ruled out .’

Evidence about the contribution of carbon from volcanic sources is difficult to obtain and interpret. One clue may come from carbonates from iron-formations within the 3 x lo9 year-old Onverwacht and Fig Tree Formations of South Africa which also have anomalously low ôC13 values (Perry and Tan, 1970 and in preparation). Lime- stones and dolomites, sometimes associated with volcanics, and even carbonates from tuffaceous sediments in these formations have normal õC1” values, suggesting that there is no volcanic carbon contribution.

Experiment al pro cedures Composite samples and a few individual specimens from four continuous drill cores of Biwabik Iron Formation were analysed in this study. Composite samples were chosen because partial analyses were available (Pfleider et ul., 1968) and because there is no a priori way of determining equilibration volume within the iron-formation.

Carbon and oxygen isotopic data on carbonates were obtained from CO, liberated by 100 per cent phosphoric acid by the technique of McCrea (1950). The Biwabik Iron Formation contains coexisting ankerite and siderite (French, 1968), both of which react slowly with phosphoric acid, and successive gas fractions from iron-formation samples contain increasing proportions of gas from the carbonate with the slowest reaction rate. Analyses of such successive CO, fractions from mixed CaCO,-CaMg(CO,), assemblages is discussed by Epstein, Degens and Graf (1964). Pending accurate determination of the carbonate phases involved and of the phosphoric acid-carbonate oxygen isotope fractionation factors for the appropriate carbonates, we have arbitrarily used the oxygen isotope fractionation factor for calcite.,

Oxygen from SiO, (chert) was liberated with BrF, and converted to CO, by the technique of Clayton and Mayeda (1963). ôû18 for chert is reported with respect to Standard Mean Ocean Water (SMOW) (Craig, 1961).

Isotopic analyses of CO, gas were performed with a mass spectrometer having a 15 cm radius analyser and equipped with double inlet and double collector systems (McKinney et al., 1950).

1 (C1z/C’2)sample i. SC'^ = - i x i 000. Similarly,

1 (O*R/O*a)sample (0,8,0,0)standard - 1 x 1000. The standard used for

carbonates is PDB. 2. Since this paper was prepared a phosphoric acid-siderite fractionation

factor has been reported by Fritz et RI. (1910). Using this factor would decrease our values of carbonates by about 1.6 per cent, but would have no significant effect on the relative values or on any of the conclusions presented here.

Unesco, 1973. Genesis of Precamhriun iron and manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 299


Discussion of data Samples for this study were taken at intervals throughout the stratigraphic section of the Biwabik Iron Formation and the data are given in Tables 1 and 2. Stratigraphic logs of percentage iron from magnetite (Pfleider et al., 1968), sol8 and 6CI3 are shown for 3 cores in Figure 1. It is apparent from this figure that isotope ratios show a

% $AGNE,JiTE 15 FE o 40 35 30

I

s O,’& (per m;i) -26 -25 -24 -23 -22 -Li -20 -19 -18 -17 -16 -15 -14 -13 -12 -11 -10 -9

S o s,!,& per mil) e--- LOWER SLATY

FIG. 1. Core log showing 6 0 1 8 and 8Cl3 of carbonates and percentage magnetite iron for three continuous core samples of Biwabik Iron Formation.

strong stratigraphic trend. W e shall explore whether this is a primary trend or the result of isotopic reactions between minerals and other components of the system during diagenesis and metamorphism.

The iron carbonate-rich top of the Upper Slaty Member, the Intermediate Slate Member, and the uppermost part of the Lower Cherty Member of the Biwabik Iron Formation are iron-formations in the sense of containing about 20 per cent or more Fe in the form of silicates and carbonates, but they contain little or no magnetite. A plot of 8 0 1 8 versus 6C13 (Fig. 2) shows that these units occupy a field different from that occupied by magnetite-bearing iron-formation units. In particular, most samples of magnetite-free iron-formations have a 6C1s greater than - 4 per mil, whereas almost all magnetite- bearing iron-formations have a less than - 7 per mil.

The Intermediate Slate contains as much as 3 per cent reduced carbon (Gruner, 1946). In sample EP 2-66 this reduced carbon has a XC13 of - 33.2 per mil, a value which agrees well with Hoering’s (1967) value of SC13 PDB = - 30.5 per mil for carbon from the Thomson Formation, stratigraphically equivalent to the overlying Virginia Forma- tion, and 6Cl8 PDB = 30.3 per mil for carbon from the correlative Gunflint Iron Formation. It also agrees well with the value of - 33.1 per cent reported by Smith et al. (1970) for the HF residue in chert from the Gunflint Iron Formation. As shown in Figure 2, this large quan-

O

bearing horizons , .I

-4 1 ‘. , - - _ _ _ - ~ ’ -.

-8 -

42-

/-

o \ I I

/

A Magnetite bearing A I horizons o 1-

/

I I

\ 0 \ - 1 -

-20 -20 -4 6 -i 2

0 D Z 0 D 5 A 07

Y 1258 EPZ-66

-0 -4 SO,,‘& per mil

FIG. 2. versus of carbonates from Biwabik Iron Formation showing fields of magnetite-bearing and magnetite- free units.

300

Significance of carbon isotope variations in carbonates from the Biwabik Iron Formation, Minnesota

TABLE 1. Isotopic and related data for core samples of Biwabik Iron Formation

Reaction time SC13 (PDB) SOIn (PDB) Percentage Percentage Stratigraphic Sample number' per mil per mil Fe2 magnetite Fe* unit

D2 1620-1624.5 D 2 1697-1705

D2 1882.4-1886.2 D2 2010-2045.8

D 2 2085-2090

D2 2170-2175

D 5 524.1-530

D5 692.5-700 D 5 755-760 D 5 829-835 D5 965-970

D5 975-985

D5 1070-1075 D7 1083

D7 1094

D7 1175-1180 0 7 1200-1205 D7 1337

EP 2-66

- 1.08 - 7.75 - 6.19

- 17.01 - 12.20 - 10.16

- 8.44 - 6.47 - 13.82 - 13.81 - 4.58 - 1.45

- 10.00 - 13.38 - 7.99 - 3.87 - 3.51 - 4.00 - 3.47

- 17.48 - 2.39 - 2.14 - 1.46 - 2.88

- 2.48 - 2.50 - 10.39 - 11.03 - 9.17 - 9.20

- 8.56

- 3.71

- 12.18 - 13.88 - 14.28

- 15.14 - 14.19 - 14.33

- 13.19 - 13.40

- 15.73 - 16.69

- 8.99 - 8.70

- 16.16 - 14.86 - 12.18 - 11.82 - 12.53

- 11.15 - 10.61 - 12.86 - 12.18 - 11.98 - 12.13 - 10.19

- 10.02 - 10.00 - 13.37 - 13.40 - 14.94 - 15.21

- 16.69 - 5.84

6.23 16.30

29.72 27.57

37.87

29.40

17.42

25.60 29.05 26.05 13.98

19.43

27.73 10.91$

10.773

32.31 30.59 27.3P

20-30

O 10.8

18.0 16.6

33.8

25.5

O

10.9 12.2 14.8 O

O

7.9 03

03

27.1 24.5 6.13

O

Upper Slaty Upper Cherty

Upper Cherty Lower Cherty

Lower Cherty

Lower Cherty

Upper Slaty

Upper Cherty Upper Cherty Lower Slaty Lower Cherty

Lower Cherty

Lower Cherty Lower Cherty

Lower Cherty

Lower Cherty Lower Cherty Lower Cherty

Intermediate Slate

0-2 hours 0-2 hours 2 hours-

1 day 0-1 day O-2.hours 2 hours-

14 days 0-2 hours

3 days 0-1 hour

2 hours-

1 hour- 21 days

0-1 hour

7 days 0-3 days 0-7 days 0-4 days 0-1 hour

3 days 0-1 hour

3 days 0-7 days 0-2 hours

4 days 4-7 days 0-1 hour

1 hour-

1 hour-

1 hour-

2 hours-

1 hour- 10 days

10-17 days 0-1 day 0-4 days 0-2 hours

4 days 4-7 days

0-1 hour

2 hours-

1. Core 2 was taken near Biwabik, Minnesota (Sec. 22, T58N, R16W); core 5 was taken near Buhl, Minnesota (Sec. 36, TSSN, R20W); core 7 was

2. Taken from Píìeider et al. (1968). 3. Single hand specimen, but analysis is for five foot composite.

taken near Keewatin, Minnesota (Sec. 36, T57N, R22W).

tity of isotopically light carbon has not produced a Sig- nificant SC13 shift in the carbonate.

The Intermediate Slate is composed dominantly of siderite and chamosite (French, 1968) and is considered to have a volcanic ash component (Morey, this volume, page 193). W e conclude from the essentially normal 8Cl3 values in this unit that it is unlikely that the S O 3 anomaly

in iron-formations is caused by a direct contribution of volcanic carbon to the sedimentary basin.

In a number of samples in which we have analysed successive fractions of CO, gas, these fractions shift according to a variety of trends (Tables 1 and 2, points connected by lines in Fig. 2). Because these shifts show no consistent trend, we reject the trivial interpretation

301


TABLE 2. Isotopic data and iron oxide percentage for composite samples of Biwabik Iron Formation (Hanna y 1258), Cooky, Minnesota

SCL3 (PDB) SO1B (PDB) Percentage Percentage Stratigraphic Reaction time per mil per mil magnetite hematite unit Sample number

47-74 74-93 93-121 121-1 58

158-1 72 172-194 194-208 208-241

- 11.95 - 12.64 - 12.78 - 12.36 - 12.43 - 11.92 - 12.82 - 12.33 - 13.15

- 14.97 - 12.40 - 12.92 - 14.21 - 17.66 - 13.25 - 15.85 - 14.01 - 14.78

41.9 35.9 30.4 30.5

28.7 27.7 34.1 27.3

1.2 O 4.2 8.3

2.3 1.3 1.8 3.2

Lower Cherty Lower Cherty Lower Cherty Lower Cherty Lower Cherty Lower Cherty Lower Cherty Lower Cherty Lower Cherty

0-1 day 0-1 day 0-1 day 0-1 day 1-3 days 0-1 day 0-1 day 0-1 day 0-1 day

that they are the result of surface exchanges with atmospheric CO,. Instead, they probably represent carbonate components with differing reaction rates and different isotopic compositions and fractionation factors. Since the shifts are small in short core specimens (D 7 samples with single footage numbers, Table i), we attribute the larger shifts to sampling intervals that exceed the volume of local equilibration. Although the shifts are a nuisance in the interpretation of data, they suggest that diagenesis and metamorphism in the Biwabik Iron Formation involved reactions in a number of small subsystems with little interaction between them.

All but one of the samples from core Y 1258 contain hematite and magnetite. W e note, without explanation, that in this buffered system SC13 varies only from - 12.0 to

. - 13.2 per mil, whereas Sol8 ranges from -12.4 to - 17.7 per mil.

Interpretation

At least three interpretations of the data of Figures 1 and 2 are possible: Both Fe for magnetite and carbon for carbonate are derived from a volcanic source characterized by low SC13. W e regard this as unlikely because the SC13 anomaly is not evident in the Intermediate Slate, the only unit of the Biwabik Iron Formation in which there is direct evidence (shards) of a volcanic contribution.

Iron is transported or precipitated by some organic process which produces carbonate with a low S O 3 value. This is the model of Becker and Clayton (1970); its validity rests on whether the isotopic patterns in Figure 1 are primary stratigraphic features or are the diagenetic and metamorphic result of primary differences. Normal E 1 3 values for carbonates from zones containing about 20 per cent iron in the form of iron silicates and carbonates, but no magnetite, indicate that the SC13 anomaly is most probably not related to transportation of iron.

W e propose here a model in which oxidation-reduction reactions during diagenesis and metamorphism produce shifts of carbon from an abundant reduced reser-

voir of organic carbon (less than - 30 per mil) to the carbonate reservoir (initially near O per mil); similar shifts may transfer oxygen from hematite to carbonate (a shift of over 20 per mil). Although this model is tentative, it can be extended to explain some of the perplexing features of deposition and diagenesis of iron- formation.

W e postulate that much of the magnetite in iron-formation was precipitated as ferric-oxide-hydroxide simultaneously with a variable amount of organic matter. During diagenesis and low grade metamorphism the following reactions (or similar ones) occurred:

(1) (2) (3)

6Fe,03 -t C(orgUnic) + 4Fe30, + CO, C1202 + FeCl3O3 3 C130, + FeC.120, 2Fe30, + 6C0, f-r GFeCO, $- O,

A more complete set of possible reactions is given by Yui (1966). After exchange according to reaction (2), CO, is assumed to have escaped from the immediate system; because of the sandwich-like nature of the iron- formation, gaseous products may continue to react in other zones, but large local variations in isotope ratios discussed previously suggest that interaction between zones was limited.

W e are not aware of CO,-FeCO, isotope exchange experiments or calculations but, by analogy with dolomite (O'Neil and Epstein, 1966), the CO, produced by reaction (1) could be out of equilibrium with FeCO, by about 40 per mil for carbon. This is probably sufficient to produce the shifts of approximately 10 per mil that we observe in the carbonate values.

A similar shift occurs in 6018, but this is complicated by possible exchange reactions in the system. Figure 1 shows the correlation in core 5 of Sol8 between carbonates and coexisting chert. The large spread in carbonate values compared to the relative constancy of coexisting cherts suggests that oxygen exchange reactions associated with oxidation and reduction may not be completely masked. Clayton (1970) reports 5 per mil shifts iii SOI8 for magnetite compared to 1 per mil maximum shifts for chert in the Dales Gorge Member of the Brockman Iron For- mation. H e attributes this to late exchange and homogen- ization of the chert, but one might speculate that the large

302


variations result from oxidation reduction reactions that do not affect the chert. Table 3 shows $OIS for selected chert samples from core 5.

TABLE 3. 60ls (relative to SMOW) of selected chert (SiO,) samples from core 5

Feet ô015 (per mil)

524-530 18.29 k .50 755-760 20.02 I. .o2 829-835 20.80 1070-1075 19.68 11 15-1 124 19.83

Implications

Our model can readily be extended to explain some of the intriguing features of iron-formation deposition. In this limited discussion we make no effort to deal extensively with the geochemical problems associated with iron- formations. A summary of the literature and an extensive bibliography are given by James (1966). Thermodyn- amics of diagenesis and metamorphism are discussed by Yui (1966).

One question which has concerned many people (Gruner, 1922; Lepp and Goldich, 1964) is how the large quantities of iron found in Precambrian iron-formations was transported. Our model is compatible with a generally reducing atmosphere in which ferrous iron could be transported to the site of deposition. Here it could be oxidized in the photo-synthetic zone by primitive organisms using the Fe++-+Fe+++ reaction as an electron donor, as suggested by Cloud (1968). Ferric oxide-hydroxide might then be precipitated and incorporated in the sediment together with dead organisms. The present day Fe,O,/ Fe,O, ratio in iron-formations would then depend on the depositional ratio of Fe,O,/C. Less than i per cent would be sufficient for reaction (1) to go to completion in a rock containing 30 per cent magnetite and would produce CO, for exchange in a mole ratio of CO, (reduction)/FeCOy

concentration of reduced carbon in iron-formations is more than adequate for this mechanism.

This model is consistent with the facies concept of James (1954) since organic matter in shallow water deposits would be rapidly oxidized by water of high Eh, whereas in deeper water organisms and Fe,O, would be deposited together. The model also provides an explanation for the puzzling coexistence of Fe,O, and Fe,O, in many iron- formations (Eugster, 1957). This coexistence merely implies that reaction (1) stopped when the local system became depleted in carbon.

- - ~ . 4 in a sediment containing 10 per cent FeCO,. Present

Iron-formation of the Transvaal System in South Africa and the .Brochan Iron Formation in Western Australia typically contain fine laminae that Trendall (1968) has suggested may represent annual accumulations. This is, of course, consistent with a mechanism for precipitation of iron oxides which depends on seasonal maxima in organic activity.

Magnetite and iron carbonates often display complex interrelated textures which have led several authors to postulate diagenetic or metamorphic conversion of FeCO, to Fe,O, (LaBerge, 1964; French, 1968). Reaction (1) is probably an early reaction that stabilizes siderite by increasing CO, pressure (reaction (3)). Higher temperature can decom- pose siderite to produce magnetite (Yui, 1966). This complicated system, coupled with variability in the size of the equilibrating system, probably explains why we cannot demonstrate a direct relationship between 8C13 and Fe,O,/ FeCO, .

Summary and conclusions

Figures 1 and 2 provide good circumstantial evidence that a negative 6C1, shift in iron-formation carbonates is related to oxidation-reduction reactions involving magnetite which permit exchange between two carbon reservoirs differing by about 30 per mil in their carbon isotope composition. As an example of one possible reaction, we have postulated a model in which primary hematite could be converted to magnetite by oxidation of organic carbon; and we have suggested some implications of this model. Four approaches may be useful in testing this model: (a) more detailed correlations such as those shown in Figure 1; (b) a tracer study of hydrogen-deuterium ratios in chamosite and other hydrous iron-formation silicates; (c) a study of end member assemblages such as hematite-carbonate; and (d) a close correlation of textural relations to isotopic data. W e are currently making a much more detailed study of carbon and oxygen isotope variations in carbonates and of oxygen isotope variations in silicates and oxides of the Biwabik Iron Formation (test 1). Our model may not be unique, but we emphasize that carbon isotopes provide an important tracer that can be used to help understand the petrology of iron-formations .

Acknowledgments

This research was sponsored by the Minnesota Geological Survey and the National Science Foundation (GP 10855). W e wish to acknowledge the help of L. A. Mattson, J. S. Owens and R. D. Scamfer of the Hanna Mining Co. and R. Bleifuss, G. B. Morey and P. K. Sims of the University of Minnesota.

303


Résumé

Signifrcatiori des variations cles proportions cles isotopes clu carborze datis les carbonates des gisements de jev de Biwabik, duns le Minnesota (E. C. Perry Jr. et F. C. Tan)

Les rapports C13/C12 dans les carbonates des formations de fer du Précambrien inférieur et moyen sont plus faibles que ceux des autres carbonates précambriens et des carbonates marins phanérozoïques. Dans la formation de fer de Biwabik, dans le Minnesota, qui appartient au Précambrien moyen, la perte de C13 varie de 7 à 8 o/oe par rapport à la calcite crétacée normale (PDB). D e faibles rapports C13/C12 ne caractérisent que les zones qui contiennent de la magné- tite et dans lesquelles est absente l'ardoise intermédiaire, matière sidéritique finement laminée, sans magnétite, qui contient certains fragments volcaniques. Cependant, dans cette matière l'abondant carbone réduit voit le CI3 diminué dans la proportion de 33 'Ioo.

L'auteur propose une réaction diagénétique d'oxydation-réduction avec production de magnétite à partir d'hé- matite, qui permettrait un échange entre le carbone organique et le carbone des carbonates :

Une conversion réversible de magnétite en carbonate est aussi possible :

2Fe,0, + 6CO,* 6FeC0, + O, r31

Ce modèle propose le transport de fer comme Fe++, l'oxydation par des organismes photo-synthétiques en hydroxyde-oxyde (comme cela a été proposé par Cloud en 1968), la précipitation de l'oxyde ferrique hydraté avec des organismes morts, et une réaction diagénétique entre l'oxyde ferrique et les organismes morts, sauf quand ces organismes sont oxydés dans des eaux peu profondes à Eh élevé. Ce modèle peut expliquer les feuillets à caractère saisonnier, semblables à des varves, avec coexistence de magnétite-hématite et certaines structures de remplacement mettant en jeu de la magnétite et du carbonate. Cela s'accorde avec le transport de fer Fe+, et n'introduit que de minimes modifications au concept de faciès de James (1 954).


BECKER, R. H.; CLAYTON, R. N. 1970. C13/C12 ratios in a Pre- cambrian banded iron-formation and their implications, Ab- stract in Trans. Arner. geophys. Un., no. 51, p. 452.

CLAYTON, R. N. 1970. Oxygen isotopes in ancient sediments. Abstract in the 14th annual report on research, The Petroleum Research Fzind, American Chemical Society, Washington, D .C.

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CLOUD, P. E. Jr. 1968. Atmospheric and hydrospheric evolution on the primitive earth. Science, vol. 160, p. 729-36.

CRAIG, H. 1961. Standards for reporting concentrations of Deu- terium and oxygen-18 in natural waters. Science, vol. 133, p. 3467.

EPSTEIN, S.; GRAF, D. L.; DEGENS, E. T. 1964. Oxygen isotope studies on the origin of dolomites. In: H. CRAIG, S. L. MILLER, G. J. WASSERBURG (eds.), Cosmic and isotopic chemistry, p. 169-180. Amsterdam, North Holland.

EUGSTER, H. 1957. Reduction and oxidation in metamorphism. In: P. H. ABELSON (ed.), Researches in geochemistry. p. 397- 426, New York, Wiley.

FRENCH, B. M. 1968. Progressive contact metamorphism of the Biwabik Iron-formation, Mesabi Range, Minnesota. Bull. Minn. geol. Surv., vol. 45.

FRITZ, P.; BINDA, P. L.; FOLINSBEE, F. E.; KROUSE, H. R. 1970. Isotope composition of diagenetic siderites from Cretaceous sediments of Western Canada. J. sediment. Petrol. (In press.)

GRUNER, J. W. 1922. The origin of sedimentary iron-formations: the Biwabik forinaiion of the Mesabi Range. Econ. Geol., vol. 17, p. 407-60.

- . 1946. Mineralogy and geology of the Mesabi Range, St Paul, Minn., Minnesota Office of the Commissioner of the Iron Range Resources and Rehabilitation.

HOERING, T. C. 1967. The organic geochemistry of Precambrian rocks. In: P. H. ABELSON (ed.), Researches in geochemistry, vol. 2. New York, Wiley.

JAMES, H. L. 1954. Sedimentary facies of iron-formation. Econ. Geol., vol. 49, p. 235-93. - . 1966. Chemistry of iron rich sedimentary rocks. Prof. Pap. US. geol. Surv., 440-W.

LABERGE, G. 1964. Development of Magnetite in iron-formations, The Lake Superior Region. Econ. Geol., vol. 59, p. 1313-42.

LEPP, H.; GOLDICH, S. S. 1964. Origin of Precambrian iron- formations. Econ. Geol., vol. 59, p. 1025-60.

MCCREA, J. M . 1950. On the isotopic chemistry of carbonates and a paleotemperature scale. J. chem. Phys., vol. 18, p. 849- 57.

MCKINNEY, C. R.; MCCREA, J. M.; EPSTEIN, S.; ALLEN, H. A.; UREY, H. C. 1950. Improvements in mass spectrometers for the measurement of small differences in isotope abundance ratios. Rev. sei. Instrum., vol. 21, p. 724-30.

O'NEIL, J. R.; EPSTEIN, S. 1966. Oxygen isotope fractionation in the system dolomite-calcite-carbon dioxide. Science,

PERRY, E. C. Jr, TAN, F. C. 1970. Carbon and oxygen isotope ratios in 3,000 m.y. old rocks of southern Africa. Abstract in Geological Society of America Annual Meeting Program, Milwaukee.

PFLEIDER, E. P.; MOREY, G. B.; BLEIFUSS, R. L. 1968. Mesabi Deep Drilling Project, Progress Report no. I, Minnesota

vol. 152, p. 198-201.

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Section, AIME Forty-fist Annual Meeting, Minneapolis, University of Minnesota.

SMITH, J. W.; SCHOPF, J. W.; KAPLAN, I. R. 1970. Extractable organic matter in Precambrian cherts. Geochn. et cosmoch.

TRENDALL, A. F. 1968. Three great basins of Precambrian banded iron-formation deposition: A systematic comparison. Bull. geol. Soc. Amer., vol. 79, p. 1527-44.

Acta, vol. 34, p. 659-75.

UREY, H. C.; LOWENSTAM, H. A.; EPSTEIN, S.; MCKINNEY, C. R. 1951. Measurement of paleotemperatures and temperatures of the Upper Cretaceous of England, Denmark, and the south- eastern United States. Bull. geol. Soc. Ameu., vol. 62, p. 399- 41 6.

Yur, S. 1966. Decomposition of siderite to magnetite at lower oxygen fugacities: A thermochemical interpretation and geological implications. Econ. Geol., vol. 61, p. 768-76.

305

Genesis and supergene evolution of the Precambrian sedimentary manganese at Moanda (Gabon) F. Weber Laboratoire de Géologie et de Paléontologie, Université de Strasbourg, France

Introduction The manganese deposit at Moanda lies in a region occupied by the Precambrian series of the Francevillian; this series derives its name from Franceville, the prefecture of the Haut-Ogooué.

The occurrence of manganese in the region of France- ville was first described by Barrat in 1895. Later Babet, Choubert, Bergé, Nicault and Briot reported manganese several times. After these discoveries, systematic prospecting was commenced in 1951 by the Bureau Minier de la France d’outre-Mer (BUMIFOM) in collaboration with the United States Steel Corporation. In September 1953, the Compa- gnie Minière de l’Ogooué (COMILOG) was founded using French and American capital to continue investigating the deposit and to commence its exploitation. A total of 96 million United States dollars were required to start exploitation in 1962, most of this being needed for transport facilities (Vigia, 1963). The mine is an open-cut, the ore, enriched in situ by scouring, is carried by cable and railway to the harbour at Pointe Noire. The mean annual production is about 1,600,000 metric tons of washed high- grade ore (48-52 per cent Mil). At this rate the reserves should last for more than 150 years.

The geological environment

The Francevillian is a Precambrian non-metamorphic series lying unconformably on the metamorphic basement of the Du Chaillu and North Gabon massives (Fig. 1). Towards the east it disappears under the overlying Cretaceous and Tertiary beds of the Batéké Plateaux. According to radiometric dating, the Francevillian is 1,740&20 m.y. old (Bonhomme et ul., 1965; Vidal, 1968; Bonhomme and Weber, 1969). The measurements have been made on three series of samples by different methods; pelites (Rb-Sr), interstratified cinerites (K-Ar) and intrusive syenites (Rb- Sr). The results of these different methods are in good agreement.

deposit

The Francevillian sediments are distributed on both sides of a median northwest-southeast ridge penetrated by large inliers of basement which are the moles of Asséo, Amieni and Ondili (Fig. 2). This ridge was a submarine rise at the time of deposition, the thickness of the sediments being weak there and the series involving stratigraphic gaps. This ridge separates two palaeogeographic domains. In the north-east, i.e. in the ‘Okondja deep’, the series is very thick (more than 3,000 m) and folded in the southwest there is an epicontinental domain where the series is less thick (1,500 m) and has only weak undulations and faults (Fig. 3). This epicontinental domain includes several basins separated by sills; in one of these basins, the Franceville basin, lies the Moanda ore deposit.

From recent studies of the Commissariat à l’finergie Atomique and the Bureau de Recherches Géologiques et Minières, a lithographic scale has been established (Donnot and Weber, 1969). Five formations have been described, which are found throughout the series, with some variation (Fig. 4).

The basal sandstone formation (FA of Figs 3 and 4) is an arkosic gritty encroachment of fluvio-deltaic character which spread all over the Francevillian area. The basal sandstone formation is very thick (3-400 m) on the border of the Du Chaillu Massif, vhere it contains thoriferous conglomerates; elsewhere this formation fills depressions in the basement. It is not very thick on the median ridge and is lacking in parts of it. The facies become thinner from south- west to northeast.

The lower pelitic and volcano-sedimentary formation (FB of Figs 3 and 4). In the Okondja deep, where subsidence was the strongest, more than 1,500 m of argillo-sandy sediments were deposited. A basic submarine fissure volcanism occurred in the northern part of the deep. Hyaloclastites were mixed with the terrigenous sediments in which sills and lavaflows with spilitized pillow lavas were intercalated.

On the submarine rises, where terrigenous supply was nonexistent, silicification took place. In the epicontinental basins, confined conditions prevailed and carbonate sedimentation took place. In the Lastoursville basin massive

Unesco, 1973. Gemsis of Precambrian iron und manganese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 307

O o

-h ?+

h Ri0 A t b

Mc

C A M E R O O N

MUN¡

EKÉ A T E A U

L E G E NO,

u ûuoternary Tertiary

N-Noya system [IIIIIITIIDD Upper P r e c a m b r i a n C-Western Congo system

S e c o n d a r y

F- Froncevillian series Middle Precambrion S I-Intermediary system c B - B a m b a system [ml Lo w e r Precambrian

o 50 100, 160 aYJtI I

according to t h e geological m a p of A.EF. ~/2000.000, by G.GÉAARD 1958

FIG. 1. Geological draft of Gabon, location of the Francevillian series.

Genesis and supergene evolution of the Precambrian sedimentary manganese deposit at Moanda (Gabon)

r 1 Oe'*" N o r t h G a b o n M a s s i f

+ + + + + +

N'GOUTOU REGION + + i

+ + + ABE1 LLES REGION

i - + + + +

t + + t + t

+ + + + + t + P L A T E A U D U C H A I L L U M A S S I F

+ + + + + + + + + + 20 + I20

I 50 K m I

FIG. 2. Schematic m a p of the principal domains of the Francevillian.

309

F. Weber

O K O N D J A BASIN - D U CHAILLU M A S S I F - F R A N C E V I L L E BASIN -MOLE OF ONDILI - AKIENI BASIN - Syncline &nticline Syncline Anticline Of YCYé of Okondia of Oyabi of Ambinda

N E

1mm- >OO.,. * ..t**++ om- 1 I * *

FIG. 3. Schematic section across the Francevillian from Moanda to Okondja.

No rt h - we st ern reg ion South region North-eastern region I I I

I LASTOURSVILLE BASIN I i FRANCEVILLE BASI! I i ! I

O K O N D J A B A S I N Olcunga

i L.dzOmdnl i

dolomlte5 basic tuffs

ocid tufts cherts

black shales coarse-grained Volcanic racks basement a peiites

FIG. 4. Diagrammatic scheme showing the correlations and variations of facies between the different domains of France- villian. The sections were based on the top of the jasper formation FC.

310


dolomites were deposited, while in the Franceville basin the subsidence was more marked and an argillo-sandy sedimentation in which dolomites began to deposit, gradually passed to an ampelitic sedimentation. The latter was suddenly stopped by a renewal of erosion leading to the deposition of a sandstone stratum. In this basin the following succession can be observed downwards:

b 30-40 m pelites and ampelites (Djoumou

a 30-100 m isogranular quartzy sandstones

c 50-150 m ampelites with dolomitic and

b 20-100 m sandstone pelites alternating -T-. with dolomitic sandstones; near

River pelites).

(Poubara sandstones).

manganiferous layers.

FB2 1 rJ% ) the bottom, intraformational

breccias. a 10-20 m greenish lustrous pelites; at the

bottom, a conglomerate with quartzitic boulders and pelitic fragments.

The Jasper formation (FC of Figs 3 and 4). Through a thickness of about 40 m , chert layers alternate with ampelites and cinerites which are the first signs of a new volcanic phase which developed in the FD formation.

The upper ampelitic and volcano-sedimentary formation (FD of Figs 3 and 4) is a thick deposit of ampelites mixed with pyroclasts. An acid volcanism of ignimbritic type induced the deposition of wide-spread layers of vitroclastic tuffs. At that time, the Francevillian was one basin only, open towards the Okondja deep; the thickness of the FD formation, which is about 150 m in the region of Franceville, increases towards the north-east and reaches more than 1,000 m in the region of Okondja.

The upper sandstone formation (FE of Figs 3 and 4) includes alternate layers of pelites and micaceous grey- wackes .

After their deposition the Francevillian formations were deformed with a mild folding in the region of Okondja and undulations, flexures and faults due to differential movements of compartments of the basement in the other regions. A last, undated, volcanic phase caused the emplacement of dolerite dikes.

(Bangombé pelites)

Summary description of the manganese and iron deposits

THE MANGANIFEROUS PLATEAUX OF THE FRANCEVILLE BASIN

The manganese deposits occur in the form of a superficial layer covering, at an altitude of 600 m, several plateaux (Fig. 5) the most important of which are the Okouma plateau (mineralized surface: 13 km2) and the Bangombé plateau (mineralized surface: 19 km2) where present-day

exploitation is carried out. The schematic section of the mineralized formation comprises downwards (Fig. 6): 0.10-0.40 m. (5) The argillo-sandy humic horizon. This horizon is leached of manganese and contains some pisolites.

5-6 m. (4) The loose pisolitic layer. The pisolites, 3 to 6 mm in diameter, are almost perfectly spherical and inserted in a yellow ochrous earth mainly composed of goethite, gibbsite and some kaolinite. They consist of a core, which is usually an ore fragment, around which concentric layers of gibbsite, goethite and, more rarely, lithiophorite alternate. This bed (15 per cent Mn) is not exploited.

0.5-1 m . (3) The transition horizon. This more or less cuirassed horizon contains fragments of mineralized plates, aggregates of pisolites cemented by concretionary cryptomelane, and big blocks of a coarse-grained cavernous feldspathic sandstone.

3-9 m (average thickness: 5 m). (2) The platy horizon. This horizon is the main part of the productive bed. Its chemical composition is given in Table 1, columns 1 and 2. Plates of ore, one or more centimetres thick, as well as massive fragments, are inserted in an ochrous matrix containing small fragments of ore around which small pisolites occasionally developed. The plates generally have a layered structure and occur in almost horizontal beds, but show in detail many undulations. In places sink-hole depressions are formed together with vertical plates and elements originating from upper horizons (fragments of the transition horizon and pisolites). Here and there in the platy horizon massive concretionary blocks are found and at the bottom special facies occur, named from their appearance ‘heavy layered ore’,

TABLE 1. Chemical composition of manganese aiid iron ores (yo)

SiO, Alzo, Ca0 MgO Na,O K2O Tio, P Fe M n

1 2

2-3.5 7.0 6-7 3.2

0.10 0.10

0.41 0.10-0.13 0.17

3-4 4.4 50-52 44

Loss on heating

3

10.4 3.2 0.35 0.28 o. 1 0.35 0.10

0.7 35.2

4

23 6.3 8.6 4.3 0.13 1.3 0.17 0.14 2.5 15

5

35.8 0.2 2.0 1.8 0.06 0.07 0.04 0.42

0.1s

~

31.2

30.2 29.1 13.6

1. M e a n composition of the ‘marketable ore’ of the Bangombé plateau during the first months of exploitation; 2. Composition of the ore straight from the mine in the pit P 39 on the Bangombé plateau. These values give a good idea of the mean composition of the ore in situ, in the zone of initial exploitation of the Bangombé plateau; 3. Analysis of a sample of layered ore with rhodochrosite collected on the Bangombé plateau at the bottom of the mineralized bed; 4. Mean composition of manganiferous ampelites from the Bangombé borehole (10 analyses); 5. Average of two chemical analyses effected in the silicate facies with greenalite of the Okouma-Bafoula iron-formation.

311

F. Weber

’ * \\A Abouka

+

o IO Km - Z o n e in w h i c h m o n g a n i f e r o u s f o r m a t i o n s u b s i s t s u n d e r c a p - r o c k

P l a t e a u x with highgrode. .oxidized ore

Z o n e in w h i c h the m a n g a n i t e r o u s f o r m a t i o n has b e e n e r o d e d

P l a t e a u x with low g r a d e oxidized a r e

0 D e e p borehole

FIG. 5. Schematic section across the mineralized horizon (after Bouladon et al., 1965).

312


1

P L A T Y

HORIZON

I

\ \ COMPACT-. -

L A Y E R - -

-like ore

:h shales

I 1 O 4m

FIG. 6. Location of the mineralized plateaux and probable extension of the manganiferous formation in the Franceville basin.

‘ polypary-like ore’ and black scoriaceous ore’. At different levels rather dislocated thin beds of sandstones with manganiferous cement and of ferruginous red shales are interbedded in the plates. The main constituents of the ore are amorphous manganese hydroxides in which can develop polianite, lithiophorite, nsutite and cryptomelane and in the matrix iron and aluminium hydroxides.

0.20-0.50 m. (1) The compact basal layer. Generally a thin band (2-5 cm) of pyrolusite (a pseudomorphism of manganite into polianite) lies on the substratum. On top is a massive layered ore composed mainly of amorphous hydroxides, manganite, groutite, lithiophorite and nsutite. Rhodochrosite appears, either as beautiful pink crystals lining geodes, or in a less spectacular form at the lowest part of the deposit as the principal constituent of a greyish shaly ore which is epigenized into manganite and pyrolusite. The analysis of this ore is given in Table 1, column 3.

The substratum is composed of subhorizontal or slightly wavy ampelites with rare intercalations of fine-grained

sandstones and dolomites which belong to the upper third of the FBI formation. Diaclases are sometimes filled with pyrolusite or rhodochrosite but, except for these accidental concentrations, the content of manganese in the ampelites of the substratum seems to be very low; according to the few analyses which have been effected, from 0.2 per cent to 0.7 per cent MnO.

True cuirasses are observed in the lower zones of the plateau. The transition horizon is hardened by the development of a cement made of bluish concretionary cryptomelane. This hardening attains progressively the underlying mineralized layer, cementing small plates and pisolites, but it stops at the shales of the substratum. These cuirasses have often been incised by the brooks draining the plateau. Thus ‘cliffs’ of massive ore and enormous boulders have been formed, which are found on the slopes and which signal the presence of the ore deposits. These brooks remove manganese which is deposited today in the form of an often rather thick wad coating.

313

F. Weber

THE CARBONATE MANGANIFEROUS FORMATION OF THE DEEP BANGOMBE BOREHOLE

In the centre of the Bangombé plateau a small plating of Poubara sandstones (FB,,) has been protected from erosion in a downfault compartment (Fig. 7). A deep borehole carried out by the COMILOG passed through these sandstones, then the pelites of Bangombé (FB,) and reached the basal sandstones (FA). This borehole showed that in the upper third of FB, is a very thick but low-grade carbonate manganiferous formation with around 13 per cent M n through 75 m.

Figure 8 shows the changes in manganese content in the deep borehole of Bangombé. In the lower two-thirds of FB,, the fluctuations of the amount of MnO (which is always below 1 per cent) parallel those of C a 0 and MgO; manganese is related to the dolomitic facies. At a depth of about 130 m in the borehole, the content of M n O suddenly increases and reaches values between 20 and 30 per cent, whereas the contents of M g O and Ca0 remain unchanged. The manganese content suddenly decreases near 55 m depth, 25 m under the bottom of the FB,, sandstones. In the last 25 m of the FB, formation, the manganese contents are about the same as in the lower two thirds.

Figure 9 shows in more detail the changes in manganese content in the manganiferous formation, compared with the variations in iron and phosphorus content. The manganiferous formation is preceded and accompanied at its bottom by an increase in the phosphorus and iron contents. At a first approximation, the three elements achieve their maximum concentrations in the following order: phosphorus, iron, then manganese.

The manganiferous formation is mainly composed of ampelites with a few intercalated sandstones and dolomites, the total thickness of which does not exceed 10-15 per cent of the formation. Dolomites are more frequent towards the bottom, sandstones towards the top. The ampelites are very rich in carbonates, which occur most frequently in the form of small radiate fibrous concretions scattered in the matrix

N W- Explalted zone

I Leconi Rlv.

1 I

which is made opaque by organic matter and pyrite; detrital elements (quartz and degraded micas between 20 and 50 p) are rare. The clay minerals are illite and chlorite, illite being largely predominant. Small aggregates and lenses of secondary silica (chalcedony) are occasionally observed.

The mean chemical composition of the manganiferous ampelites is given in Table 1, column 4.

Manganese is associated with calcium and magnesium in carbonates: a manganiferous dolomite and a calcic rhodochrosite, the average formulae of which are approximately (Mgn.8, MnOn.2). Ca(CO3)z and (Mno.9, Cao.,) CO,. Iron occurs essentially in the form of pyrite. The average mineralogical composition is approximately the following: Quartz, 11 per cent; Illite (+ chlo lori te), 23 per cent; Carbonates, 56 per cent (MnCO,, 31 per cent; Cacoa, 16 per cent; MgCO,, 9 per cent); Pyrite, 4 per cefit; Organic matter, 6 per cent.

The manganese content of the dolomitic and sandstone layers is lower than that of the ampelites (3-4 per cent Mn). This explains the rapid variations in the M n content in the profile of Figure 9, and the slightly lower mean content of M n in the manganiferous formation compared with the ampelites (13 per cent instead of 15 per cent).

THE IRON FORMATION OF OKOUMA-BAFOULA

In the Okouma and Bafoula plateaux in the periphery of the zone of ore deposits, the manganese ore lies upon a banded iron-formation about 10 m thick. The following three facies have been described from the base: sulphide facies, carbonate facies and silicate facies.

The sulphide facies is characterized by a high percentage of pyrite within a microcrystalline quartz-chalcedonious matrix containing apatite, chlorite, degraded micas and organic matter.

The carbonate facies contains alternating siliceous and carbonate beds. The siliceous beds are composed of microcrystalline quartz, the elements of which (5-20 IJ,) are

S E iecounou ~ l v .

I borehole I

3001

* I I . . . . . . 'I. . . . . . . . . . . . ..... I

01~~ 20 BFB~ DF~ Enrlched oxldixed ore Lor-gr'ade carbonritid ore . . . . .. . .

FIG. 7. Section across the Bangombé plateau through the deep borehole and the exploited zone.

314

F/! .... ,....a. ......a .....e. ......a . . . . . . .

350m

Conglomerates.with rolled a Micaceous finegrained . Dolomite quartzy boulders sandstones

Alternation of finegrained Manganiferous carbonated a Polygenic brecclas B pelites or black shaies' . black shales

Ankeritic dolomite

Mn

1 -

=?

5 10 15 20 25 %

O

Fe

O 2 4 6 8 10 12%

Sandstone Gritty breccia Dolomite Black shale Carbonated black shale

Diagrams of Fe and P must be compared with the peaked line of the Mn diagram

FIG. 9. Contents of manganese, iron and phosphorus in the manganiferous formation of the Bangombé borehole.

moulded one against the other. Isolated crystals and small aggregates of siderite are scattered in the quartz matrix. A little pyrite is also observed. In the carbonate beds the crystals of siderite are closely packed and occupy almost ali the rock; quartz, generally finer-grained than in the siliceous beds, remains in the interstitial space. Pyrite is rarer and a discrete green clay mineral (chlorite or greenalite) appears.

The silicate facies generally contains greenalite as the principal iron-bearing mineral. A much rarer facies with ferristilpnomelane occurs at the top of the formation. Siderite is always present, more or less abundant, and some beds contain a little pyrite. These facies present a regular rhythmic interbedding of 0.5-5 mm thick silicated beds of alternate laminated or spherolitic texture. Siliceous beds analogous to those of the carbonate facies, but finer-grained, are irregularly distributed, In the silicate beds with laminar texture the greenalite fibres are perpendicular to the shaly structure, which is marked by a discontinuous line of organic matter. In the silicate beds of spherolitic texture greenalite occurs in the form of sheaf-like structures. These beds are enriched in siderite, pyrite and a phosphate belonging to the apatite group. Sometimes, at the top of the silicate beds of spherolitic texture, a phosphate bed appears which is composed of large spherolites of apatite (150-300 p) within a microcrystalline matrix of quartz, siderite, apatite and silicates.

The silicate facies is characterized by the greatest content of iron; its chemical composition is given in Table 1, column 5. Note the low contents of manganese, calcium,

P

O 0.1 0.2 0.3 0.4 0.5 0.6 %

aluminium and alkali metals and the high content of silica. This chemical composition is similar to that of the other Precambrian iron-formations.

The role of supergene weathering in the genesis of the Moanda ore deposit

HYPOTHESES AND DISCUSSION

In the first descriptions of the Moanda ore deposit published by Baud (1954, 1956), the author regarded it as a ‘residual deposit’ originating from a process analogous to that of the formation of laterites and bauxites, the manganese having originated in the Franceviliian rocks where its content does not exceed the ‘ clarke’. Varentsov (1964) and Thienhaus (1967) argued from this that sediments poor in manganese (less than 1 per cent) were the original rock of high-grade supergene ore deposits.

In contrast, Bouladon et al. (1965) considered the Moanda ore deposit as a sedimentary deposit subsequently enriched by lateritization. From a metallogenic study, Bouladon described a ‘primary ore’ with preserved layered structure; the main constituents of this ore are amorphous hydroxides and, in the basal bed, manganite and rhodochrosite. From these constituents, cryptomelane, nsutite, lithiophorite and polianite could have developed during lateritic weathering.

31 6


The discovery in the Bangombé borehole, at the top of the FBI formation, of a manganiferous formation whose stratigraphic position is identical to that occurring in the Moanda ore deposit reopens the problem of the origin of the deposit. It is likely that the ‘primary’ ore of Bouladon resulted from the transformation in situ of a primary carbonate ore, analogous to that which is locally protected by faults in the small caved-in compartments of the borehole. This transformation would be the result of a first phase of supergene weathering earlier than that which produced the cuirasses and pisolites. The principal arguments in favour of this hypothesis are the following. Residues of layered ore with rhodochrosite occur at the bottom of the ore deposit. The replacement of rhodochrosite by manganite, which is itself transformed into pyrolusite, has been observed on several occasions (Bouladon, 1963; Bouladon et al., 1965; Weber, 1969). Note, however, that the layered ore with rhodochrosite differs from the manganiferous ampelites in that it contains three times the amount of Mn, in the nature of the carbonates (rhodochrosite without substitution of M n by Ca or Mg) and in its structure. Recrystallization of secondary rhodochrosite partly hides the primary structure of the sediment. It should be considered as a transitional facies, hardly evolved, but already transformed by supergene factors, rather than as aii intact residue of the primary carbonated ore.

The microscopic structure of the manganiferous ampelites is partially preserved in the parts of the oxidized ore which have been the least transformed during recent secondary reworking. In fact, the amorphous hydroxides which are considered by Bouladon (1963) as the principal ‘primary’ constituents of the ore, sometimes show a spotty structure (‘fish spawn’) imitating that of the carbonates in the manganiferous ampelites.

The presence of gibbsite and kaolinite shows that the ore bed must have undergone a strong lateritic weathering. Indeed these minerals are lacking in the unweathered sediments of the Francevillian. Therefore, it is not surprising that residues of carbonate ore are so scarce on the plateau and found only at the very bottom of the deposit. Such a strong weathering necessarily destroyed most of the carbonate in the mineralized formation.

N o borehole has ever met an interstratified bed of oxidized manganese below the weathering zone in the Francevillian format ion.

The ore deposits occur at the top of the plateaux apparently coinciding with an ancient peneplain (Chatelin, 1964). Thus, this carbonate formation is evidence of a for-

mation which was originally much more extensive. After outcropping it was subjected to weathering and gave rise to the ore deposit.

GEOCHEMICAL BALANCE O F THE SUPERGENE TRANSFORMATIONS OF THE ORE DEPOSIT

On the basis of the hypothesis that the parent rock of the ore had a composition very near that of the carbonated manganiferous ampelites of the Bangombé borehole, using the isovolumetric method (Millot and Bonifas, 1955), an approximate balance of the transformations undergone by the ore by weathering can be drawn up. In Table 2 the average composition of manganiferous ampelites per unit volume is compared with that of the ore in situ in the deposit of the Bangombé plateau. Since the sedimentary structure of the ore was preserved in most of the mineralized bed, probably the volume did not vary much during weathering.

TABLE 2. Isovolumetric balance of the supergene enrichment of the ore

Weight of 100 cms

M n SiO, A1203 Fe P M g O Ca0

Manganiferous ampelites of the Bangombé borehole

282 g

42 65 (quartz 31) 18 7.0 0.39 12 24

Ore of the Bangombé plateau, in situ

213 g

94 15 19 9.4 0.36 0.21 0.21

The weathering oí‘ the ore resulted in an important enrichment in manganese, the weight of which per unit volume more than doubled. On the other hand calcium, magnesium and a high percentage of the silica were removed. The contents of alumina, iron and phosphorus remained relatively constant with a weak enrichment in iron.

For the other elements the existing data are not complete enough to draw up an exact balance. Note, however, that the sulphur originally combined with iron as pyrite has been almost completely removed; in the ‘marketable ore’ the amount of sulphur is generally below 0.05 per cent. The same is observed for the carbon of organic matter and for carbonates. On the other hand potassium and barium fixed in cryptoinelane have been only partially removed.

The manganese which concentrated in the ore deposit probably originated in the eroded upper part of the carbonate manganiferous ampelitic formation. An approximate balance shows that only 20 per cent of the manganese originally contained in the 75 m thick ampelites did concentrate at the bottom of the bed in a 5 m thick layer. A small part (5 per cent) remains in the superficial pisolitic horizon, but most of it (75 per cent) has been lost. Ap- plying the same balance horizontally, the percentage of the original manganese recoverable today for exploitation is no longer 20 per cent but only 1-2 per cent, since the manganiferous formation has been completely eroded over more than 90 per cent of its surface (Fig. 5).

317

F. Weber

ATTEMPT TO RECONSTRUCT THE ENRICHMENT PROCESS IN THE ORE

Superficial waters rich in oxygen and CO, percolated through the manganiferous formation after erosion. In the top horizon of the weathering profile, oxidation of pyrite produced sulphuric acid which reacted with carbonates and gave rise to sulphates and CO,. In the deep horizons, the presence of sulphates favoured the action of sulpho- reducing bacteria which developed in this sediment rich in organic matter, oxidizing it to CO2 and H,S. Acid and very corrosive waters percolated through the formation, but their oxidizing character decreased rapidly because of the formation of H,S. Carbonates were attacked and manganese dissolved in a bivalent state in the form of ions and manganous complexes.

While percolating, the waters were enriched in bicar- bonates and the p H increased, lowering the solubility of manganese considerably. The increased p H favoured the oxidation of manganous ions and complexes, resulting in the precipitation of manganese hydroxides, but in a sufficiently reducing medium, manganese also reprecipitated as the carbonate.

Thus, at the bottom of the weathering profile, a manganese accumulation horizon could form by epigenesis of complex carbonates to manganese hydroxides and/or rhodochrosite.

Iron did not show any tendency to follow manganese in its migration. It occurred principally in the form of pyrite which was attacked only under the highly oxidizing surface conditions. In the pyritic zone the dissolution of iron in the bivalent state would have required a considerable decrease of pH, incompatible with the buffering capacity of the carbonates. The mineralogical forms of iron and manganese differ in the parent rock, and this explains why these two elements, despite their similar chemical properties, had different destinies during the first phase of weathering of the ore, and why only manganese migrated and concentrated in the lower horizons of the weathering profile.

THE SECONDARY TRANSFORMATIONS O F THE ORE DEPOSIT

After this first phase, the ore was subjected to other transformations which have been shown by Bouladon (1963) to be more directly related to lateritization; at this time the cuirasses and pisolites were formed and the bed of enriched ore was completely oxidized and dislocated.

The cuirasses resulted from horizontal migrations of manganese towards the depressed zones of the plateau within the mineralized horizon which was being dislocated. The pisolitic overlap is probably the deeply transformed residue of the upper horizons, which have been leached of manganese and relatively enriched in iron and alumina. Pisolites formed by concretion of these elements around small fragments of ore; this is the beginning of iron and

alumina incrustation. Since the pisolitic horizon is very homogeneous, it is clear that, unlike manganese, iron and aluminium hardly migrated. It is likely that the presence of manganese hindered the migration of iron; colloidal solutions of manganese hydroxide (Mn(OH),) are weakly acid and flocculate iron hydroxide (Fe(OH),) which is weakly basic. The excess manganese can then migrate forming almost pure manganese cuirasses, iron having been fixed in sitic. Thus the formation of pisolites seems to be complementary to that of cuirasses.

In the mineralized bed, a redistribution of the elements is observed. In the fragments of massive layered ore, stratified solution cavities appeared in which the removal of manganese oxides left a limonitic residue having the same nature as the ochrous sterile matrix of the ore. The composition of this matrix-gibbsite, goethite and traces of kaolinite-shows that it is the result of strong lateritic weathering. Iron and alumina were fixed again in small pisolites analogous to those found in the pisolitic overlap. Redistribution of manganese resulted in an enrichment of the small plates at the expense of the intercalated beds, which are more or less completely leached; the auto- catalytic power of MnO, explains this redistribution. Mas- sive cuirassed boulders developed locally within the mineralized horizon.

EVOLUTION OF THE MORPHOLOGICAL CONDITIONS DURING THE SUPERGENE TRANSFORMATIONS OF THE ORE DEPOSIT

The phase of ore enrichment took place beneath the groundwater level. The base level, slightly lower than the bottom of the manganiferous formation, probably ensured sufficient drainage and a continuous circulation of water through the formation, but it was probably higher than it is today. The present-day morphological disposition does not allow the presence of a permanent water table in the mineralized bed.

Lateritization and encrustation occurred in the zone of water table fluctuation. Lowering of the base level exposed the enriched horizon to weathering. It was dislocated and locally invaded by cuirasses. Then the process of enrichment stopped, except in the basal bed, where a permanent water table remained in the decimetres overlying the impermeable shales of the substratum. Here the rhodochrosite which formed initially in a lower horizon was epigenized to manganite.

Finally, after another lowering of the base level, the ‘cuirasses’ themselves were notched by erosion and partially dislocated.

The manganiferous plateaux of the Franceville region, like other plateaux in that region settled on manganese- devoid formations, give evidence of a peneplain which could be related to the ‘inner peneplain’ of Cameroon, as defined by Segalen (1907). This peneplain would correspond to a cycle of erosion which started at the end of the Cretaceous or at the beginning of the Eocene; it would

318


have been notched by erosion from the beginning of the Miocene, after epirogenetic movements which affected the African continent during the Late Tertiary and the Quaternary. Supposing these correlations are exact, the major phases (enrichment and cuirassement) of the supergene evolution of the Moanda ores can thus be referred to the Eocene cycle of erosion and peneplanation.

Genesis of the primary carbonate ore deposit

ORIGIN OF MANGANESE

The existence of an important volcanic activity during the deposition of the Francevillian sediments and the chert- manganese association in some deposits suggest a volcanic origin; volcano-sedimentary manganese deposits are common all over the world (Chatsky, 1954; Routhier, 1963) and generally related to siliceous rocks.

Despite their location in the domain of the continental shelf, it seems that the Moanda manganese deposits must be related to the spilitic volcanism of Okondja rather than to an ignimbritic volcanism. The ignimbritic volcanism of the Francevillian is particularly manifest in the FD formation, after the deposition of manganese, whereas the spilitic volcanism of Okondja occurred at the same time as the deposition.

The acid tuffs of the Francevillian have very low contents of manganese (always below 0.15 per cent Mn). In contrast, the basic lavas and spilites of Okondja and the associated hyaloclastites have high contents of manganese, reaching sometimes 0.8 per cent (average: 0.35 per cent Mn).

The spilitic volcanism of Okondja helps to explain the supply of important quantities of manganese to the basin during the deposition of the FBI sediments.

THE MODE OF INTRODUCTION O F MANGANESE IN THE SEDIMENTARY BASIN

As emphasized by Bernard (1968), the hypothesis of ‘ trans- vaporization’ (Brousse, 1968) throws new light on the problem of mineralizations related to spilitic complexes. The rise of the magma through a great depth of unconsolidated sediments impregnated with salt brine resulted in the formation of large quaniities of hydrothermal fluid. These hydrothermal solutions, leaching the sediments of their most mobile elements, led to a considerable enrichment of the sea-waters in heavy metals which were redis- tributed according to the rules of sedimentary mechanisms.

In the case of the volcano-sedimentary formation of Okondja, the sediments in which the magma intruded were principally hyaloclastites originating from anterior eruptions. These hyaloclastites were composed of lava rich in

manganese and therefore the mechanism of transvapor- ization could have enriched the sedimentary basin in this element.

THE SEDIMENTARY MECHANISMS O F MANGANESE CONCENTRATION

Other elements were dissolved at the same time as manganese, especially iron, calcium, magnesium and silica, which are much more abundant than manganese in the lavas. The sedimentary mechanisms of deposition sorted these elements and concentrated manganese in certain parts of the sedimentary basin. W e shall consider here the mechanisms responsible for the enrichment of manganese vis-à- vis iron and calcium, whose geochemical behaviour is normally rather similar to that of manganese in a sedimentary environment.

The partition of iron and manganese during sedimentary and volcanosedimentary processes has been studied by several authors (Marchandise, 1956; Krauskopf, 1956; Michard, 1969). Krauskopf (1956) showed that one could not expect an important enrichment in the solutions compared with the lava since iron and manganese are leached in similar proportions. The partition of iron and manganese occurs during deposition by an early precipitation of iron, most of the iron compounds, especially sulphides and oxides, being less soluble than those of manganese. The manganiferous formation of Bangombé is preceded by iron enrichment, but in the formation of Okouma-Bafoula it is laterally that the associated iron deposit must be found. The suppliyng waters of the manganiferous basin deposited a part of their iron content on its periphery before flowing into it, while the conditions of precipitation of the manganese carbonate had not been reached. Thus a ferriferous deposit, almost devoid of manganese, formed laterally to the manganiferous deposit, but before it. When the manganese precipitated in its turn, the iron which had remained in solution also precipitated, but the solution had been previously impoverished in iron which appears only in small proportions in the manganiferous deposit (Mn/Fe = 6).

The solubility product of manganese carbonate is higher than that of calcium carbonate, but Michard (1968, 1969) showed that ‘the direct precipitation of manganous carbonate is generally impossible in marine environments because of the high proportion of calcium; in reducing environments one observes a coprecipitation resulting in a rather weak enrichment’. Michard showed, however, that calcareous beds can be enriched in manganese by a mechanism in which diffusion phenomena intervene in sediments between an oxidized superficial zone and a reducing deep zone. In the formation of the Bangombé borehole, manganese was fixed in the form of a mixed precipitate with calcium and magnesium carbonates. However, the rate of enrichment in manganese is higher than in the models established by Michard with the used parameters (concentration of sea-water, sedimentation velocity, etc.). In the

319

F. Weber

Franceville basin the waters must have been rather stroiigly enriched in manganese by volcanic exhalation, and during the Precambrian the calcium concentration of the sea was probably lower than today. Moreover, the diagenetic enrichment must have been favoured by a very low sedimentation velocity, which agrees with the ampelitic nature of the sediment.

to the mechanism proposed above; lastly, most of calcium and magnesium constituted the dolomitic deposits in the nner part of the gulf (basin of Lastoursville).

Conc1usion

THE PALAEOGEOGRAPHIC SCHEME

The deposition of manganese took place in a coastal basin barred by submarine rises at a rather great distance (100 km or so) from the source of exhalative supplies related to spilitic volcanism. The reducing character of the sedimentary environment allowed manganese to remain in solution and to migrate over great distances.

The introduction of manganese into coastal basins and the distribution of the different deposits in these basins can be interpreted by a scheme analogous to that proposed by Brongersma-Sanders (1965) for the Kupferschiefer where sedimentation occurred at the boundary of the euxinic and evaporitic domains. The scheme was inspired by the mode of circulation of currents in present-day bays and estuaries, especially in the Cariaco gulf (Venezuela). A superficial current runs toward the margin of the basin, whereas a deep current runs toward the centre where water rises to the surface. This model explains thoroughly the distribution of the chemical deposits in the Francevillian. The elements carried by the deep currents coming from the open sea were deposited in the order of their increasing solubility (Fig. 10): iron, silica and phosphorus were deposited first on the submarine rises barring the gulf inouth; in the centre of the gulf (Franceville basin) calcium and magnesium began to deposit, fixing dissolved manganese according

The concentration of manganese in the Moanda ore deposit was accomplished in three stages: the first was magmatic (the spilitic volcanism of Okondja supplied lavas with a manganese concentration higher than the ‘clarke’ of the earth’s crust); the second was sedimentary (and diagenetic) (manganese originating from lavas was concentrated in sediments); the third was supergene (manganese of sediments was concentrated in the present-day ore deposits). Let us consider the ‘rate of concentration’ corresponding to these three stages, i.e. the ratio of the manganese content of one stage to the manganese content of the previous one, the starting point being the mean ‘clarke’ of the earth‘s crust.

TABLE 3.

Mean Magmatic Sedimentary Supergene ‘Clarke’ stage stage stage

Mn content

Rate of concentration - 3 50 3

(%I 0.1 0.3 15 45

The highest rate of concentration occurred during the sedimentary stage, but the other stages were still necessary in order to make a high-grade ore deposit out of what would otherwise only have been a geochemical anomaly.

Ca , ‘Mg Mn (Ca, Ms) Si, Fe, P (Mn)

FIG. 10. Rongersma-Sander’s (1965) scheme applied to the Francevillian.

Résumé

Genèse et évolution siqwgène du gisement sédimentaire précambrien de manganèse de Moandu, au Gabon (F. Weber)

Le gisement de manganèse de Moanda, mis en exploitation en 1962 par la Compagnie minière de l’Ogooué (COMILOG), produit annuellement environ 1 600 O00 tonnes de minerai à haute teneur. La couche exploitée, d‘une puissance moyenne de 5 mètres, forme l’entablement du

plateau de Bangombé, où elle couvre une superficie de plus de 19 km2. D’autres plateaux minéralisés existent dans la région.

Le minerai est formé pour l’essentiel de plaquettes d‘oxydes et d‘hydroxydes de manganèse dans une matrice argileuse d’hydroxydes de fer et d’alumine, avec un peu de kaolinite. Un recouvrement pisolitique stérile de 5 k 6 mètres d‘épaisseur surmonte la couche minéralisée. Le substratum

320


est constitué de schistes noirs stériles du Francevillien. Le Francevillien est une série sédimentaire précam-

brienne, non métamorphique, dont l’âge est de 1740 20 millions d‘années, selon les datations radiomé-

triques. La série stratigraphique, tronquée par l’érosion dans la zone minéralisée, est conservée dans un petit compartiment effondré, situé au centre du plateau de Bangonibé. Un sondage profond y a mis en évidence une formation manganésifère, à faible teneur (13 % Mn) niais très puissante (75 mètres). La base de cette formation correspond stratigraphiquement à la base de la couche exploitée. Cette formation est constituée de schistes noirs carbonatés légèrement pyriteux, pauvres en éléments détri- tiques. Le manganèse est associé au calcium et au magné- sium dans des carbonates complexes. Le minerai en plaquettes dérive probablement de l’évolution sur place de ces schistes noirs carbonatés manganésifères, sous l’action des agents supergenes. L’enrichissement et l’oxydation du minerai se sont produits lorsque la formation manganésifère originelle affleurait sur une ancienne surface, actuellement entaillée par l’érosion; de nombreux témoins de cette surface subsistent dans la région.

Des bilans isovolumétriques montrent qu’il n’y a pas eu simplement oxydation du manganèse et lessivage des

cations solubles (Ca et Mg). Le minerai se serait formé dans un horizon profond du profil d’altération. La partie infé- rieure de la formation manganésifère a été enrichie par un apport en manganèse provenant du lessivage des horizons supérieurs. Les carbonates complexes de Ca, M g , et M n ont ainsi été épigénisés par des oxydes et des hydroxydes de manganèse. Une étape intermédiaire comportant épi- génie des carbonates complexes par de la rhodochrosite doit sans doute être envisagée, au moins dans les horizons les plus profonds. Le recouvrement pisolitique proviendrait des horizons supérieurs partiellement lessivés en manganèse. Ultérieurement, par suite d‘un abaissement du niveau de la nappe, le minerai a subi une altération latéritique intense qui l’a démantelé et qui est responsable de la formation de cuirassements latéraux.

L’origine du gisement carbonate peut être mise en relation avec un volcanisme spilitique qui se manifeste à l’époque du dépôt dans une fosse située à 100 km au nord-est du gisement. Le dépôt de manganèse s’est effectué en bordure de cette fosse dans des bassins épicontinentaux isolés du large par des barrières de hauts-fonds. Une formation ferrifère rubanée siliceuse à sidérose pyrite et greenalite se rencontre autour du dépôt de manganèse et lui est antérieure.


BAUD, L. 1954. Notice explicative de la feuille Franceville-Est, Carte géologique de reconnaissance au 11500 000. Brazzaville, Direction des mines et de la géologie de l’A.-E.F., 34 p., et Chron. min. colon., no. 221, p. 260-61. __ . 1956. Les gisements et indices de manganèse de l’A.-E.F. XXe Congr. géol. int., Mexico. Colloque sur les gisements de manganèse, vol. II, p. 21-30.

BERNARD, A. 1968. Introduction pétrographique et métallogé- nique sur le cycle géosynclinal et la métallogenèse cratonique. Conférences et séminaires de recyclage-Métallogénie, 1, III, p. 1624, Nancy, 10-14 juin 1968 (inédit).

BONHOMME, M.; WEBER, F. 1969. Compléments à la géochrono- logie du bassin de Franceville et de son environnement. 5O Col- loque de géologie africaine, Clermont, 1969, à paraître dans Ann. Fac. Sci. Clermont, fasc. Géol. Miner. -- - ; FAVRE-MERCURET, R. 1965. Age par la méthode rubidium-strontium des sédiments du bassin de Franceville (République gabonaise). Bull. S. Cartegéol. Als. Lorr., no. 18, fasc. 4, p. 243-52.

BOULADON, J. 1963. Le gisement de manganèse de Moanda (Gabon). Étude de la zone de première exploitation. Rapport BRGM, no. 5313/MPMG (janvier 1963) (inédit). __ ; WEBER, F.; VEYSSET, C.; FAVRE-MERCURET, R. 1965. Sur la situation géologique et le type métallogénique du gisement de manganèse de Moanda, près de Franceville (République gabonaise). BuII. S. Carte géol. Als. Lorr., vol. 18, fasc. 4,

BRONGERSMA-SANDERS, M. 1965. Metals of Kupferschiefer supplied by normal sea water, Geol. Rdsch., Ed. 55, p. 365-75.

BROUSSE, R. 1968. In: AUBOUIN, J.; BROUSSE, R.; LEHMANN, J. P. 1968. Précis de géologie, VOI. I, 711 p., Paris, Dunod.

p. 253-76.

CHATELIN, Y. 1964. Notes de pédologie gabonaise. Cah. ORSTOM, vol. II, fasc. 4, p. 3-28.

CHATSKY, N. S. 1954. Sur les formations manganésifères et la métallogénie du manganèse. I: Les formations manganésifères volcanogènes-sédimentaires. Bull. Acad. Sci. URSS (Moscou), Série géologie, no. 4, p. 3-37, [English translation in Int. geol. Rev., vol. 6, no. 6, p. 1030-56 (1964).]

DONNOT, M.; WEBER, F. 1969. Carte géologique de reconnaissance au 11500 000. Franceville-Ouest, avec notice explicative. Paris, BRGM. (A paraître.) -~ . 19696. Carte géologique de reconnaissance au 11500 000. Franceville-Est, avec notice explicative. Paris, B R G M . (A paraître.)

KRAUSKOPF, K. 1956. Separation of manganese from iron in the formation of manganese deposits in volcanic associations. XXc Congr. géol. int., Mexico, 1956, Colloque sur les gisements de manganèse, vol. I, p. 119-31.

MARCHANDISE, H. 1956. Contribution à l’étude des gisements de manganèse sédimentaire. XXe Congr. géol. int., Mexico, 1956, Colloque sur les gisements de manganèse, vol. I, p. 107-18.

MICHARD, A. 1968. Coprécipitation de l’ion manganeux avec le carbonate de calcium. C.R. Acad. Sei., Paris, no. 267, p. 1685-8.

-. 1969. Contribution à M u d e du comportement du man- ganèse dans la sédimentation chimique. Thèse Faculté des sciences de Paris, 194 p.

MILLOT, G.; BONIFAS, M. 1955. Transformations isovolumé- triques dans les phénomènes de latéritisation et bauxitisation. Bull. Carte géol. Als. Lorr., vol. 8, p. 3-10.

ROUTHIER, P. 1963. Les gisements niéfall$ères. Paris, Masson. 1282 p.

321

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SEGALEN, P. 1967. Les sols et la géomorphologie du Cameroun. Cuh. ORSTOM, Série pédologie, no. 2, p. 137-87.

THIENHAW, R. 1967. Montangeologische Probleme lateritischer Mangaiierz-Lagersttäten. Mineralihm Deposita, vol. 2, no. 4, p. 253-70.

VARENTSOV, I. M . 1964. Sedimentary manganese ores. Amster- dam, Elsevier. 119 p.

VIDAL, P. 1968. La méthode potassium-argon dans la datation-.

des séries sédimentaires. Application aux sédiments du bassin de Franceville. Thèse 3" cycle, Faculté des sciences de Stras- bourg. 55 p.

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WEBER, F. 1969. Une série précambrienne du Gabon, le France- villien; sédimentologie géochimie, relations avec les gîtes miné- raux associés. Thèse Faculté des sciences de Strasbourg. 367 p.

322

The Belinga iron ore deposit (Gabon)

S. J. Sims Bethlehem Steel Corporation, Pennsylvania (United States)

Introduction The Belinga iron ore deposit is the largest deposit in the Mekambo district, a vast, isolated, and largely unexplored area in the north-eastern part of the Gabonese Republic in West Equatorial Africa (Fig. i). The Mekambo district includes at least five deposits: Boka-Boka, Batouala, Belinga, Minkebe and Kokomeguel (Fig. 1). All of the deposits are similar in type and origin of iron ore; all have been derived from Precambrian iron-formation. It is estimated that within the Mekambo district there could be 1,000 million tons of iron ore averaging about 64 per cent Fe. Further exploration could well enlarge this figure. It is clear that this is an important area of undeveloped iron ore and will surely gain importance as known world supplies of iron ore are steadily consumed.

The occurrence of iron ore in the Mekambo district has been recognized for many years and is briefly mentioned in several early reports (Barrot, 1895; Launay, 1903; Periquet, 1911; Choubert, 1937; Chochine, 1938; Rouquette, 1938; Chochine, 1950; Devigne and Plegat, 1954; Aubague, 1955, 1956). Because of the remoteness of the district, it is only relatively recently that exploration has taken place. In 1954 the French Direction des Mines mapped the Boka- Boka deposit and collected samples (Devigne and Plegat, 1954). In 1955 the Bethlehem Steel Corporation, in conjunction with the French Bureau Minier de la France d'outre-Mer, undertook a reconnaissance examination of the Boka-Boka deposit, and subsequently formed the Syndicat de Mekambo in order to study this promising deposit as well as the nearby Batouala deposit. In 1958, as it became obvious that the Mekambo district had indeed a large potential, the Société des Mines de Fer de Mekambo (SOMIFER) was formed in order to explore the much larger Belinga deposit. Accordingly, this area was explored from 1958 through 1962, and this paper is based on the results of that exploration effort. Since 1962 very little work has been done on the area.

The Belinga area is 65 km north-east of the town of Makokou and is accessible by boat on the Ivindo River

o IO 20 30 40 50 km , I

FIG. 1. Map of north-eastern Gabon showing the Mekambo iron district. Iron ore deposits are hachured and the area of Figure 2 is outlined.

Unesco, 1973. Genesis of Precambrian iron and ìnaizg'onese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 323

S. J. Sims

from Makokou or by air to a small landing strip near the deposit (Fig. i). The deposit is located in an area roughly 20 km north-south by 5 km east-west. The latitude and longitude of the centre of the deposit are 13'14'E and 1'6' N.

The Belinga deposit is in the equatorial rainforest, where rainfall averages about 2.5 m per year. Several north- south trending ranges are in the area and rise up to 550 m above a gently rolling plain which has an average elevation of 500 m above sea level. Maximum relief in the area is about 550 m, and ranges from a low elevation at the Ivindo River of 450 m to a high elevation of 1,000 m at the highest peak in the area. Topography ranges from moderate to rugged, with many oversteepened slopes. Vegetation is dense and ubiquitous, and exposures of rock are limited to scattered outcrops along the crests of the ranges and in a few stream courses.

The Belinga deposit consists of six explored and four unexplored ore bodies situated along the crests of the ranges, the distribution of which is shown in Figure 2.

BELINGA A R E A

:=----; I R O N FORMATION CONTACT. '9 \,'\

I '\'

I ' I I

ORE BODY, EXPLORED I I :I ORE 00DY. UNEXPLORED

I o I 2 3 4 skm - J ,

I I

I ' \I

FIG. 2. Map of the Belinga area showing the explored and unexplored ore zones and the distribution of iron-formation.

Exploration of the ore bodies was by means of adits and by surface geological and topographic mapping. Chame1 samples were taken at 2 m intervals along the adit walls. A total of 8,027 m in fifty-nine adits was driven, proving at least 515 million metric tons of iron ore averaging 64.2 per cent Fe, 2.2 per cent SiO,, 3.5 per cent Alzo,, 0.122 per cent P, and 3.8 per cent ignition loss (in these rocks this can be considered as equivalent to H,O +). An additional 50 million metric tons of iron ore of similar analysis are estimated in the four unexplored ore bodies. Previous exploration at Boka-Boka and Batouala yielded about 300 million metric tons of iron ore. The over-all probable tonnage for the Mekambo district is, therefore, 865 million metric tons. Within the Mekambo district additional but unknown tonnages of ore exist in the Minkebe and Kokomeguel deposits.

The north-eastern part of Gabon is a vast plateau of Lower Precambrian basement rocks consisting mainly of quartz diorites with scattered areas of amphibolites and iron-formation (Hudeley and Belmonte, 1966). Almost nothing is known of the structure and stratigraphy in this region. Included within íhis basement complex is the iron-formation, a regionally metamorphosed layered quartz- iron oxide rock. Known exposures in Gabon extend over an arc-shaped area from Boka-Boka on the south-east through Batouala and Belinga to Minkebe on the north and Kokomeguel on the north-east. It is presumed that this iron-formation is a single unit or series, but this has not yet been established.

Almost everywhere the basement complex has been weathered to laterite and lateritic clay from which almost all of the main mineral components, except alumina, iron oxide, and silica, have been leached. However, original textures and structures are preserved in many places allowing tentative identification of the parent rock. The iron-formation is much more resistant to weathering and erosion and consequently forms distinct ridges throughout the region.

Rock types

IRON-FORMATION

The iron ores are derived from, and are gradational to, the iron-formation. An arbitrary limit of 60 per cent Fe is defined as the boundary between ore and iron-formation in this paper. The iron-formation at Belinga is a regionally metamorphosed layered rock unit consisting of thin alternating layers, lamellae, and lenses of quartz and iron oxide and also clay (mainly as kaolinite). Because of structural complicatioiis and lack of data, no thickness for the iron- formation has been measured. It is estimated to be between 100 m and 200 m thick. In places the iron-formation is almost entirely clay and hematite. The presence of clay in the iron-formation at Belinga has led to a subdivision of the iron-formation into three rock types based on the relative amounts of quartz, iron oxide and clay. These are: ítabirite, argillaceous itabirite and hematitic phyllite. The

324


term itabirite is used herein as defined by Dorr and Barbosa (1963) in Brazil. If the iron-formation contains over 5 per cent clay, it is termed argillaceous itabirite, and if it contains less than 5 per cent quartz, it is termed hematitic phyllite. Figure 3 is a triangular diagram of the three mineral components showing the compositional fields of each type of iron-formation. This is an empirical diagram and represents arbitrary limits based on field identification of the three types of iron-formation during mapping of the exploration adits. The chemical analyses of channel samples for each type of iron-formation as mapped in the adits were converted to percentages of mineral components, and the mineral components for each type of iron-formation were then plotted on a triangular diagram. The lines separating the three types of iron-formation in Figure 3 are straightened and separate the fields of maximum concentration of each type of iron-formation. This diagram was made from 725 points.

IRON OXIDE

ARGILLACEOUS ITABI RITE

CLAY 8

QUARTZ BAUXITE

FIG. 3. Triangular diagram showing the compositional fields of types of iron-formation at Belinga. Based on 725 points.

The relative amounts of each type of iron-formation based on the intercept-distance in the adits are: itabirite, 47.1 per cent; argillaceous itabirite, 22.7 per cent; hematitic phyllite, 30.2 per cent. N o stratigraphic relationships have been worked out among these three types of iron-formation.

Itabirite. This rock type consists of interlayered quartz and iron oxide. In general the layers are discontinuous, range in thickness from about 0.05 mm to 10 mm, and consist mainly of either one or the other component. Quartz layers are composed of a granoblastic mosaic of grains which typically show undulatory extinction and locally have strain lamellae. Quartz grains range in diameter from 0.01 mm to 0.2 m m . Contacts between grains vary from sharp to indistinct where very fine-grained

impurities are concentrated between the grains. In some samples quartz grains are elongate within the layering and are oriented parallel to isoclinal fold hinges. Iron oxide layers are composed of grains of hematite (typically with relict traces of magnetite in the cores of the hematite grains) and in places limonite (as partial replacement of hematite). In many layers the grains of hematite appear to have grown together forming an anhedral tabular mass of hematite. Iron oxide layers are approximately the same thickness as quartz layers. Contacts between layers are relatively sharp. Itabirite ranges from very friable to hard and massive, depending on the degree of weathering. Weathering of itabirite causes a break-down in intergranular contacts between quartz grains due to leaching of silica.

The average chemical analyses (wt. per cent) of 446 samples of itabirite at Belinga is: Fe, 46.9; Mag. Fe, 5.9; SO,, 30.3; P, 0.047; Alzo3, 1.2; loss on ignition, 1.1. These analyses include both fresh and weathered itabirites. For comparison, the average analysis (wt. per cent) for fifty-four samples of fresh itabirite is: Fe, 38.8; Mag. Fe, 5.8; Sioz, 42.7; P, 0.035; Alzo,, 0.8; loss on ignition, 0.7.

In two separate adits, itabirite rich in a prismatic mineral altered to limonite was noted. The prismatic form of the mineral strongly suggests amphibole, but because of the high degree of alteration, no positive identification was made. Amphiboles were identified in itabirite at Boka- Boka as hornblende and riebeckite (Mekambo Syndicat, 1959), and these may well be present at Belinga too. This is not a widespread type of iron-formation at Belinga.

Argillaceous itabirite. This rock type is an itabirite with over 5 per cent clay mineral (kaolinite) interlayered with quartz and iron oxide. The clay occurs in distinct layers of about the same size as quartz and iron oxide layers and is also intermixed with iron oxide forming a groundmass for hematite grains. All samples of argillaceous itabirite observed were weathered and very friable. The average chemical analysis for 214 samples of argillaceous itabirite is (in wt. per cent): Fe, 49.4; Mag. Fe, 6.5; Sioz, 21.9; P, 0.082; A1,0,, 4.1; loss on ignition, 3.0.

Hematitic plzyllite. This type of iron-formation is composed of clay and hematite with scattered layers of granular quartz. The rock is very fine-grained and thinly layered, with layers alternating between hematite-rich and clay-rich, but on a scale such that megascopically the rock appears nearly homogenous. The average grain size of hematite is about 0.02 mm. In some places distinct layers of white clay up to 10 mm thick are present. Quartz layers range up to 10 mm thick and are unevenly distributed. Hematitic phyllite is friable in all observed occurrences and is typically reddish. The average analysis of 235 samples of hematitic phyllite (in wt. per cent) is: Fe, 52.4; Mag. Fe, 8.1; Sioz, 9.3; P, 0.090; A1203, 9.2; loss on ignition, 6.0.

It should be noted that in the hematitic phyllites, Alzo, is in excess of the SiO, necessary to form clay mineral. In this case it is assumed that the excess Alzo, occurs as bauxite. No fresh samples of hematitic phyllite were noted at Belinza.

325

S. J. Sims

OTHER ROCKS

Other rock types present in the area are almost completely altered to clay mineral (kaolinite). Relict textures and structures suggest the following types are represented: phyllite, schist, cataclastic rock and intrusive igneous rock. These rock types are interlayered with the iron-formation, range in thickness from about 1 m to several tens of metres, and are typically in sharp contact with iron-formation. Lineations and drag folds are present in these rocks as in the iron-formation, indicating they are concordant. Scat- tered lenses of quartz occur throughout these rocks but are not typical, suggesting the original rocks were mainly quartz-free. The main valleys of the Belinga area are probably underlain by these rocks which were less resistant to erosion.

Quartz veins cut all the rock and ore types (with the exception of hard massive ore) and occur as irregular masses and as true veins, mainly discordant, and frequently with associated coarse crystals of specularite. Quartz veins are always deformed and the quartz in the veins is splintery and friable. Some of the quartz masses have spots of white clay suggesting altered feldspars and the possible presence of pegmatites.

ORIGIN

Only a brief statement is given concerning the origin of the iron-formation at Belinga. The itabirite is believed to have formed from ferruginous cherts by recrystallization during regional metamorphism. The argillaceous components of argillaceous itabirite and hematitic phyllite represent clastic interruptions during the predominantly chemical sedimentation of ferruginous cherts when shaly material was deposited.

Structure

Two generations of folding are evident in the Belinga rocks. The earlier folding was isoclinal and was formed in response to metamorphic deformation and recrystallization. These folds have attenuated limbs, thickened, sharp crests (Fig. 4), and are present throughout the area. They range in size from microscopic to at least several metres across. The hinge lines of these folds define a lineation throughout the area which is illustrated in Figure 5, a stereogram of 165 measured lineations. This shows a maximum concentration of points plunging 70" S, 50" E and a rotation of points about a horizontal axis trending about N 17" E. The horizontal rotation is caused by the second generation of folding.

The second generation of folds has nearly horizontal axes and open and irregularly shaped crests with many open cavities parallel to the layering. These folds are characteristic of brittle folding, and in places have an almost

lrn

FIG. 4. Isoclinally folded itabirite. Adit 121, 87 m, Mombo Range, looking parallel to fold axes.

N

s FIG. 5. Stereogram of lineations in the Belinga area. Equal area net, lower hemisphere, 165 points, contours at 3, 6, 9, 12, 15 and 18 per cent.

326


chevron structure. Figure 6 is a stereogram of 203 fold axes showing a concentration of axes nearly horizontal and trending N 18" E. These folds occur exclusively in leached iron-formation and iron ores (with the exception of hard massive ore) and are attributed to collapse due to leaching of silica. Figure 7 illustrates this type of folding.

All of the rocks at Belinga are layered to varying degrees. The layering is considered parallel to the original bedding in most places, modified by recrystallizatioii, but nevertheless reflecting bedding, In some exposures, however,

N

s FIG. 6. Stereogram of secoiid generation fold axes in the Belinga area. Equal area net, lower hemisphere, 203 points, contours at 3, 6, 9 and 12 per cent.

the layering is at an angle to bedding as illustrated, for example, in a hematitic phyllite where a granular quartz bed is cut by layering and elsewhere where a contact between schist and soft platy ore is at an angle to layering. In these places layering is a foliation. Figure 8 is a stereogram of poles to layering for 1,496 points in the Belinga area and shows that the layering (bedding) is concentrated at about a strike of N 15" E and a dip of 30" S-E. Figure 8 also illustrates the second generation of folding by a scattering of points rotated about a nearly horizontal axis trending N 15" E. Isoclinal folding would not be illustrated because both limbs of isoclinal folds have nearly the same attitude.

Based mainly on the idea that small scale structures reflect large scale structures and that the attitude of layers in the Belinga area is relatively consistent, it is suggested that the rocks of the area are isoclinally folded on a large scale and are thereby repeatedly exposed throughout the area. Also, as a result of this folding, the iron-formation is locally thickened, thereby providing favourable zones for iron ore development.

No major fractures were encountered in the adits, and consequently no faults are shown on Figure 2, although it is highly possible many faults will be uncovered when mining begins. A stereogram of poles to shear fractures measured in the adits shows a bimodal concentration of vertical planes trending about N-S and N 20" E, or nearly parallel to the second generation fold axes, suggesting that the shears may also have formed.in response to collapse of the iron-formation and iron ores.

In the adits intraformational breccias were observed in places mainly in argillaceous itabirite. These breccias are

lm

FIG. 7. Second generation collapse folds in high grade soft platy ore. Adit 116E, 50 rn, Bakota South Range.

3 27

S. J. Sims

N

O

,s FIG. 8. Stereogram of poles to layering in the Belinga area. Equal area net, lower hemisphere, 1,496 points, contours at 2, 4, 6, 8 and 10 per cent.

typically less than 1 m thick and are concordant with the layering, at least where observed in the adits. These zones represent places where metamorphic folding exceeded the ability of the rocks to accommodate plastically to the deformation. In some thin sections the crests of microfolds are fractured along axial planes forming an axial plane cleavage.

metres

W

800 825 I . -

..... : .. . .

. . . ,.,,, . ..;,,,., +&$. T-117 VfA . ....

The preceding types of fracturing may well have influenced the permeability of the iron-formation and consequently may have been a control for ore formation.

Iron ores

The iron ores of Belinga occur on and beneath the crests and upper flanks of the ranges in ten ore zones (Fig. 2). The ore grades down dip to iron-formation, the bottom contact ranging from less than 1 m to over 100 m below the surface. The bottom contact is irregular because it interfingers with iron-formation. The typical shape of the ore bodies is, therefore, crudely tabular with the length parallel to the range, the width perpendicular to the range, and the thickness perpendicular to the upper surface of the range. With the possible exception of the hard massive ore, exploration results show that almost all of the ore at Belinga occurs within 100 m of the surface. A n example of typical ore occurrence at Belinga is illustrated in Figure 9, which shows the surficial nature of the ore grading downwards to iron-formation.

The iron ores at Belinga are classified according to grade, texture, and structure. High grade ore contains over 66 per cent Fe, intermediate grade ore ranges from 60 per cent to 66 per cent Fe, and low grade ore ranges from 45 per cent to 60 per cent Fe. Only intermediate and high

metres

CZZl GANGA E%Zi CLAY SOIL !XZl INDURATED, HYDRATED, PLATY U SOFT, PLATY ORE [

HEMATITIC PHYLLITE O CLAY (ALTERE0 SCHIST)

o I O0 200 D INTRUStVE

metres

825

800

775

750

'725

FIG. 9. Structure section, Section 117, Bakota South Range, showing position of supergene ore and repetition of units by isoclinal folding.

328


grade ore are economic, so that the low grade ore can be considered only as enriched iron-formation. Ore textures range from soft (friable) to indurated (due to secondary limonite) to hard (compact) and structures range from granular to platy to massive. A given ore is then referred to by grade, texture and structure as for example, intermediate grade indurated platy ore.

The Belinga ores (both high and intermediate grades) fall mainly into the following classifications: hard massive, indurated platy, soft platy, and soft granular. Other types, such as indurated granular and hard granular, are present, but are. The distribution of the ore types (both high and intermediate grades) at Belinga in the several ore bodies is shown in Table 1.

TABLE 1. Volume per cent distribution

Ore body Indurated Soft platy Soft granular Hard massive

Bakota North Bakota South Bakouele Mombo Babiel North Babiel South Kombi Total, Belinga

50.0 45.3 65.8 40.5 13.9 57.2 88.9 46.1

7.8 38.7 3.5 12.1 8.5 34.1 .11.9 20.3 4.0 6.1 53.4 O 2.5 83.6 O 9.5 33.3 O 8.5 O 2.6 9.4 33.4 11.1

It is noteworthy that soft platy and indurated platy ores are most prevalent and that hard massive ore is mainly restricted to one ore body. Soft granular ore is widely distributed, but does not make up a large percentage of the total.

Typically, indurated platy ore occurs at the surface and grades downward to soft platy and soft granular ores. However, there are places where soft platy ore is directly beneath a thin soil cover. Hard massive ore on Bakota South Range occurs at the surface and continues to depths of at least 100 m. Contacts between ore types are characteristically gradational over a distance of several metres and are not necessarily defined by layering.

Canga, a separate and unimportant type of ore at Belinga, is a rock composed of detrital material derived from iron ore and iron-formation and cemented by limonite. Canga is not widespread, occurs at the surface, and is rarely more than 2 m thick. The occurrence of canga at Belinga is in sharp contrast to that in Brazil (Dorr, 1964) where it makes up a considerable percentage of the iron ore. This is probably due to differences in erosion rates between the two areas. The average analysis of canga at Belinga is 61.3 per cent Fe, 0.158 per cent P, 0.7 per cent SO,, 5.3 per cent Alzo, and 5.2 per cent loss on ignition.

The iron ores are composed mainly of hematite, with varying amounts of limonite and minor amounts of quartz and clay. Magnetite occurs only as remnants in hematite grains and makes up less than 10 per cent of the iron oxide. Limonite occurs as rims around hematite, as linings in cavi-

ties, and as a ground mass for hematite grains. Limonite may also replace individual hematite grains, generally along a given granular layer, forming indurated plates. Hematite occurs in both granular and specularitic forms, the latter being present mainly in the hard massive ore but also in vugs in hard plates in soft platy ore. Granular hematite ranges from 0.005 mm to 0.5 mm in diameter and averages about 0.1 mm . Specularite blades occur mainly as out- growths from granular hematite and in places specularite forms concordant layers composed of intergrown blades. In many samples of hard massive ore and some samples of soft platy ore, finely crystalline specularite lines open cavities in the ore. Specularite crystals have grown at the expense of quartz and in ores of high specularite content (hard massive ore) replacement relations between specularite and quartz are noteworthy.

Quartz is present in most of the ores in amounts less than 5 per cent, where it occurs mainly as loose grains in pores. Quartz grains tend to be concentrated along layers parallel to the layering in the ore. Clay minerals are present in some of the ores, both as a primary compound and as coatings on fractures as the result of infiltrations from the surface.

The iron ores are layered to various degrees. In hard massive ore layers are not well defined, but on close inspec- tion contrasting grain size and layers of slightly more porous material define layering and isoclinal folding the same as observed in itabirite. In the soft layering is more distinct and is defined by plates of harder hematite in granular hematite, and by layers of contrasting grain size and porosity. Small scale isoclinal folds are, for the most part, not preserved, having been obliterated by the second generation folding. In the indurated platy ores, plates of hematite cemented and partly replaced by limonite define layering.

The ores show a range of porosity from a low of 8 per cent for some samples of hard massive ore to 55 per cent for some samples of soft granular ore. The porosity is reflected by the average in-place density for each ore type as shown in Table 2.

TABLE 2.

Ore type Density, in-place, tons/m3

Hard massive Indurated platy Soft platy Soft granular

4.2 3.6 2.9 2.6

The following table (Table 3) summarizes the chemistry of the various ore types as shown by averages of channel samples for each type.

No analyses are available specifically for high grade soft granular ore, although experience shows that this type of ore is chemically similar to high grade soft platy ore. It is noteworthy that the A1,0, and loss on ignition analyses are relatively high for all ore types, especially the intermediate grades. This reflects the presence of limonite and clay

329

S. J. Sims

TABLE 3.

Inteurnediate guade ore Fe 63.6 63.4 63.9 62.2 Mag. Fe 0.3 5.1 7.9 9.2 P 0.128 0.165 0.101 0.072 SiO, 1.3 1.3 2.2 6.4

4.1 3.2 2.3 1.4 Loss on ignition 2.6 4.0 3.2 2.4 Number of analyses (96) (415) (514) (144)

High guade ore Fe 67.3 67.2 67.5 Mag. Fe 0.4 6.8 5.9 No P 0.082 0.116 0.078 analyses SiO, 0.7 0.5 1.3 available

Loss on ignition 1.2 1.4 0.9 Number of analyses (105) (205) (399)

A1203 1.7 1.4 1.2

in the ore, but mainly it reflects the presence of surficial lateritic clay which has infiltrated by means of meteoric water. The phosphorus content is also relatively high for those ores with high Alzo, and loss on ignition and is believed to be due mainly to the association of phosphorus and limonite. Experience at Belinga has shown a close relationship between limonite content and high phosphorus analyses. It should be noted that the above analyses are from channel samples that were not washed because the in-place analyses were needed. Consequently, infiltrated surface material along joints in the ore accounts in part for the unusually high Al,O, contents of the ores as shown in the channel samples.

Granulometric studies on bulk samples of indurated platy and soft platy ores were made. They show that, for indurated platy ore, the average size analysis is 36 per cent plus 3/8” and 12 per cent minus 100 mesh, and for soft platy ore the analysis is 28 per cent plus 3/8” and 14 per cent minus 100 mesh. These tests also showed that, in general, there is little chemical variation between size fractions for a given sample, but that Alzo3 is slightly higher in the coarser fractions and Si02 slightly higher in the finer fractions. Phosphorus content is nearly equal for all size fractions.

Origin of the ores

The hypothesis presented herein for the origin of the ores at Belinga involves two generations of concentration of iron oxide diflering widely in time of concentration, method of concentration and type of ore produced, but with one process superimposed on the other.

The first generation of iron oxide enrichment formed hard massive ore by metasomatic replacement of quartz in iron-formation by specular hematite. It is believed this

replacement took place after, or near the end of, Precam- brian metamorphic deformation, probably as a result of hydrothermal activity associated with emplacement of igneous rocks. The evidence that suggests this origin of hard massive ore is as follows: 1. Structural details, the same as observed in iron-

formation, are preserved in hard massive ore. This includes the small isoclinal folds and a consistent linear direction.

2. Layers of intergrown specularite were noted only in hard massive ore.

3. Replacement relations between specularite and quartz were noted only in hard massive ore.

4. Hard massive ore is found in one ore body, is not widely distributed as are the other types, and is not obviously related to the present-day surface.

5. Lenses of hard massive ore up to 1 m thick were observed in unleached itabirite showing discordant contacts. A lens of this type is clearly metasomatic because small isoclinal folds are preserved in it, and is not supergene because the itabirite is not highly leached.

6. Polished sections of hard massive ore show no evidence of hydration or replacement of hematite by limonite. Hard ores tend to show the least amount of relict magnetite in hematite, which indicates a higher degree of replacement of magnetite by hematite than in the other ore types.

7. The presence of quartz veins with coarse specularite throughout the area attests to a period of hydrothermal activity. Small vugs of specularite in the ores are thought to represent lenses of quartz which were incompletely replaced and later leached out.

It is thought that certain zones of the iron-formation were more permeable to hydrothermal fluids, perhaps due to favourable structures such as fold crests fractured during late-stage metamorphism. In these zones quartz was replaced and hard massive ore was formed, the degree of replacement determining high or intermediate grades. It is also thought that a source of the iron could have been the oxidation of magnetite to hematite, evident in the iron- formation throughout the area. This oxidation, for equal volumes, yields a slight amount of excess iron which could easily account for the enrichment. It should be clearly noted that the above suggestion is speculative and much more information is needed to coníìrm it.

At a much later geological time, when the iron- formation was exposed to surface weathering, silica, predominantly quartz, was leached by percolating meteoric waters above the water table leaving a porous hematite-rich rock. Associated with the leaching of quartz was the alteration of the other rocks to clay. At and near the surface, iron oxide was partly hydrated forming the indurated platy ores. Hydration is seemingly a near surface phenomenon and as such may be related to vegetation, as pointed out by Ruckmick (1963) at Cerro Bolivar in Venezuela.

The type of ore formed depended on the degree of leaching, degree of hydration and the nature of the hematite- rich layers and the distributioii of the quartz and hematite

330


Y. Fe

in the original iron-formation. For instance, argillaceous itabirite would yield alumina-rich ore, itabirite with dense hematite layers would yield platy ore, itabirite with less well-defined hematite layers would yield granular ore, and itabirite with partial leaching would yield intermediate grade ore. The leaching of quartz is undoubtedly continuing under present day conditions.

The following observations suggest that the iron ores (other than hard massive ore) formed by ground water leaching of quartz from the iron-formation:

i

1. 2.

3.

4.

5.

6.

- - Iron ore is restricted to near the surface. Iron ore grades to iron-formation at depths generally less than 100 m. Density and porosity measurements show that if pore spaces in soft platy and soft granular ores were refilled with quartz, the density would correspond to an itabirite. Collapse structures in the ores indicate removal of a large volume of material. In two adits a down-dip gradation from iron ore to iron- formation can actually be observed. Silica content measurements in springs and streams were measured by Park in 1958 and showed a range of from near zero in adits in the ore zone to 14 ppm in springs at the base of the ranges. The Ivindo Rover measured 9 ppm. These measurements show that SiO, is soluble under present day conditions.

Figure 10 shows analyses taken from a typical exploration adit. This shows a decrease in Fe and loss on ignition (IL) and increase of SiO, with depth, reflecting a decrease in leaching of SiO, and hydration as distance below the surface is increased.

Densitymeasurements inplaceweremade on 100 samples each of the four main ore types. Using only the densities obtained for soft platy and soft granular ores, and assuming the pore spaces were once filled with quartz and that the iron oxide mineral is hematite, a density was calculated for the assumed unleached quartz-hematite rock. From a total of 114 samples, the following results were

r70 7

obtained: range of porosities: 30-55 per cent; average porosity: 36 per cent; range of calculated densities: 3.1-4.2; average calculated density: 3.7.

From the density, a composition of Fe and SiO, was calculated assuming only hematite and quartz. Using the above data, the range in Fe for the unleached parent itabirite would be 20-52 per cent and the average would be 40.5 per cent Fe. It is interesting to note that the average Fe percentage for mainly fresh itabirite is 38.8 per cent Fe. Therefore, the calculated Fe content of pre-leached itabirite (40.6 per cent Fe) compares well with the actual average of fresh itabirite (38.8 per cent Fe) and strongly suggests that leaching alone can account for the soft platy and granular ores. It should be noted that the density measurements in these ores represent maximum values because the ores have been collapsed in part due to leaching of silica. Conse- quently, the Fe content as calculated would also be a maximum.

In two adits the down dip gradation from ore to iron- formation could be observed. In one of these adits a sample was taken in vertically dipping layers at the top of the adit in platy ore aiid at the bottom in itabirite, a separation down the dip of about 2 m. The results are shown in Table 4.

In the other adit no comparative analyses are available,

TABLE 4.

Fe Mag. Fe P SiO, A1,0, Loss on ignition

62.3 55.8 3.3 1.3 0.032 0.040 7.8 19.2 1.3 3.0 1.2 1.0

7. cio2 r

.... ........ ............ ........ ............... si02 . . ............................................ ...

O O 10 eo 30 40 50 60 70 80 90 96

HORIZONTAL DISTANCE FROM PORTAL

FIG. 10. Graph showing chemical analyses of channel samples v. horizontal distance in metres from the portal. Adit 124, Bakouele Rangs.

331

S. S. Sims

but a striking contrast from quartz-free platy ore grading down dip at 60" to quartz-bearing soft itabirite was noted and is illustrated in Figure 11.

It has been suggested by Park (1959) that hard massive ore in French Equatorial Africa may have formed by supergene replacement of quartz by hematite. It is possible that locally, on a small scale, some hematite has formed at the surface, but little evidence of this was observed at Belinga. For reasons cited above, the large mass of hard massive ore at Bakota South is not considered supergene.

It is apparent that, because iron ore is not everywhere formed over iron-formation, there are some controls to supergene enrichment. Grain size of parent iron-formation may have been a factor, but more likely structural deformation was more important. In zones of greater metamor-

phic deformation, the iron-formation was probably more permeable to meteoric waters, as it would have been to hydrothermal fluids. Thus the association of hypogene and supergene ore may be more than coincidence.

In summary, then, the Belinga iron deposit is believed to have formed as the result of supergene leaching of silica from an iron-formation that had been structurally thickened by isoclinal folding and enriched locally by hydrothermal replacement of quartz by hematite. The fortuitous combination of metasomatic replacement and structural thickening of the iron-formation af Belinga during Precambrian deformation prepared a favourable locale for surficial leaching when the area was exposed to prolonged weathering under the stable geologic environment of the central African Shield.

FIG. 11. Soft platy ore grading down-dip to itabirite. Note appearance of white granular quartz layers in the centre of the picture. Adit 115W, 66.5 m, Bakota North Range.

Résumé

Les minerais de feu de Bélinga, au Gabon (S. J. Sims)

Le gisement de minerai de fer de Bélinga, dans la partie nord-est du Gabon, est le plus important gisement du district de Mekambo qui est, pour la plus grande partie, encore inexploré. Il a été découvert en 1955 et contient plus de 550 millions de tonnes de minerai avec une teneur en fer de 64 %. I1 se répartit entre six massifs de minerai explorés et au moins quatre encore inexplorés le long de crêtes de direction générale nord-sud. I1 occupe une surface totale d'environ 5 km sur 20 km dans la forêt équatoriale humide.

En raison des conditions tropicales extrêmes et de la couverture forestière, la géologie de la région est peu connue, et celle de la région de Bélinga a été interprétée par extrapolation entre les sections accessibles et d'après les carac- tères topographiques.

Les minerais de fer proviennent de la formation de fer de Bélinga, à des degrés différents, et cette formation peut être subdivisée en trois types de roches ou de faciès : itabirite argileuse et phyllite hématitique. La distinction entre chaque type est faite d'après la proportion relative des trois composants principaux : hématite, quartz et argile (kaoli-

332


nite). L'itabirite se présente sous forme d'hématite discrè- tement rubanée et de lamelles de quartz. L'itabirite argileuse est une itabirite dans laquelle on reconnaît des lamelles d'argile et la phyllite hématitique est intercalée et mélangée à l'argile et à l'hématite. D'autres types de roches de la région de Bélinga sont ia phyllite, le schiste, les roches cataclastiques et les roches intrusives. Toutes sont décom- posées à l'état d'argile. Des filons de quartz coupant çà et là la formation de fer sont caractérisés par la présence de spécularite. Toutes les roches datent du Précambrien.

La formation de fer de Bélinga a une direction géné- rale nord-sud. Elle plonge rapidement surtout vers l'est. La formation de fer est plus résistante à la désagrégation et forme des chaînes de collines. Au point de vue structural, la région est interprétée comme une zone de plissements isoclinaux, ce qui a pour effet la répétition des formations de fer le long d'une série de crêtes parallèles, La déformation à l'intérieur de la formation de fer est marquée par une structure rubanée et l'on y observe des plis isoclinaux avec des flancs atténués et des crêtes serrées formant une structure linéaire qui plonge rapidement. En superposition à ces plis métamorphiques plus anciens, on note des plis horizon- taux de petites dimensions dans les formations altérées de fer et dans le minerai de fer. On en conclut que le plissement horizontal plus récent a eu pour cause l'effondrement dû à la lixiviation de la silice de la formation de fer.

Les minerais de fer ont été classés en fonction de leur teneur, de leur texture et de leur structure. Les minerais à haute teneur contiennent plus de 66 % de fer, les minerais intermédiaires entre 60 et 66 % de fer et les minerais à faible teneur de 45 à 60 % de fer. On estime que seuls les minerais à haute teneur et à teneur intermédiaire ont, pour le moment, une valeur économique. Les minerais à haute teneur se présentent sous différentes formes : compact dur, compact lamellaire, tendre, compact tendre. On passe des structures massives à des structures lamellaires puis granuleuses, et de textures tendres (friables) à des textures indurées puis dures. L'auteur présente les analyses chimiques des différents types de minerai. O n passe progres- sivement d'un type de minerai à l'autre.

A l'exception du minerai compact et dur à haute teneur, les minerais ont été formés par lixiviation de la silice de la formation de fer. A cette lixiviation s'est asso- ciée une addition d'oxyde de fer hydraté à certains types de minerai. Toutefois, on connaît de nombreux exemples de minerai tendre à haute teneur dans lesquels on ne relève aucune hydratation ou seulement une faible hydratation. Par endroits, la densité et la porosité ont été mesurées sur de nombreux prélèvements de minerai à haute teneur ou à teneur intermédiaire et allant de la structure lamellaire tendre au minerai granuleux. Ces mesures montrent qu'un minerai à haute teneur peut provenir d'une itabirite par simple lixiviation du quartz; l'hydratation est en effet secondaire et limitée essentiellement aux couches voisines de ia surface. L'article donne des exemples de passage du minerai à l'itabirite lixiviée lamellaire tendre et à l'itabirite renouvelée. L'effondrement du plissement , évident dans les minerais et la formation de fer lixivié, est la preuve d'un déplacement d'une importante quantité de sílice.

Le minerai dur compact à haute teneur est sans doute d'origine sinmétamorphique. II a été formé par le remplacement métasomatique hydrothermal du quartz par de l'hématite durant les derniers stades du métamorphisnie. Ce processus a eu, apparemment, à Bélinga, un développe- ment limité car ce type de minerai ne compte que pour 10 % de la totalité du minerai de fer. O n a reconnu de l'hématite compacte dure dans l'itabirite sous la forme de lentilles discordantes dans lesquelles la survivance de certaines structures de l'itabirite a été conservée. O n pense que durant les derniers stades du métamorphisme, après le plissement, des intrusions ignées accompagnées de courants hydrothermaux ont pénétré la formation de fer localement. Le quartz a été remplacé localement dans la formation de fer. Tout cela a probablement été favorisé par les structures existantes. L a combinaison fortuite du remplacement méta- somatique et de l'épaississement structurai de la formation de fer pendant la déformation précambrienne a favorisé la lixiviation superficielle lorsque la région a été exposée à une altération prolongée sous l'environnement tectoni - quement stable du bouclier africain central.


AUBAGUE, M. 1955. Les gisements de fer de la région Makokou- Mekambo. BdI. Div. Min. Géol. A.E.F., no. 7, p. 61-7. - .1956. Les gisements de fer de la région Makokou-Mekambo (Massif du Djaddie-Djouah et de l'lvindo). Bull. Div. Min. Géol. A.E.F., no. 8, p. 45-52.

BARROT, M. 1895. Sur la géologie du Congo français. Ann. Min., Paris, 9' Série, t. VII, p. 379-510.

CHOCHINE, N. 1938. Notes sur trois gisements de fer dans la zone F. Brazzaville, Gouvernement Général de I'AEF, Service des Mines (Unpublished.) __ . 1950. Notice explicative sur la feuille Malcolcou-Est. Brazzaville, Gouvernement Général de I'AEF. 16 p.

CHOUBERT, B. 1937. Étude géologique des terrains anciens du Gabon. Thèse, Paris, Rev. Géogr. Phys., 210 p.

DEVIGNE, J. P., PLEGAT, R. 1954. Le gisement de fer de Boka- Boka. Rap. Annu. Serv. Géol. A.E.F. 1954, p. 71-4.

DORR, J. VAN N. II 1964. Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., vol. 59, p. 1203-40. DORR, J. VAN N.; BARBOSA, A. L. M. 1963. Geology and ore deposits of the Itabira district, Brazil. Pyof. Pap. U.S. geol.

HUDELEY, H.; BELMONTE, Y. 1966. Carte géologique de la Réprr- blique gabonaise. Paris, Bureau de recherches géologiques et minières.

SWV., 341-C, 110 p.

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LAUNAY, L. de. 1903. Les Richesses minérules de l’Afrique. Paris, Béranger.

MEKAMno SYNDICAT 1959. Gabon, French Equatorial Africa, the Boka-Bolca iron deposits, March, 1959. Bethlehem Steel Cor- poration private report, 68 p.

PARK, C. F. Jr. 1959. Origin of hard hematite in itabirite. Econ. Geol., vol. 54, p. 573-87.

PERIQUES, L. 1911. Mission d’étirdes au Gabon: Chemin de fer du Nord et Mission Iiydrogruyhique, Paris.

ROUQUETTE, G. 1938. Étude des gisenients de fer de Boka-Bolca, Cocotiodie, Ivindo (Gabon), Brazzaville, Gouvernement Géné- ral de I’AEF, Service des Mines. 63 p.

RUCKMICK, J. C. 1963, The iron ores of Cerro Bolivar, Vene- zuela. Econ. Geol., vol. 58, p. 218-36.

Discussion

J. VAN N. DORR. Does the distribution of the supergene ore suggest that ore formation was related to particular elevations? In other words is ore formation related to a particular erosion surface or peneplain?

S. J. SIMS. Yes, formation of the supergene ores appears to be related to an ancient erosion surface. The elevation of ore bodies in the Mekambo District is roughly equal.

G. A. GROSS. What is the proportion of supergene ores to hypogene ores?

S. J. SIMS. Hypogene ore makes up about 10 per cent of the total ore.

B. CnouBERT. Is there any relationship between the Belinga and Boka-Boka deposits in the eastern part of the region?

S. 5. SIMS. Yes, I believe that Boka-Boka is related to Belinga, perhaps by the same iron-formation. However, the precise relationship has not yet been established.

334

Itabirite iron ores of the Liberia and Guyana shields

H. Gruss Gewerkschaft Exploration, Dusseldorf, Federal Republic of Germany

In 1969 world production of iron ores reached approximately 670 million tons, of which nearly one-third-about 210 million tons-was shipped from the producing countries to the consumers. Thus, Liberia, Sierra Leone and Vene- zuela nowadays may be considered to be the most important oversea’s iron ore suppliers for the United States of America and European industries.

All ores of these countries were mined in itabirite iron ore deposits. Similarities and analogous qualities of the ores, as well as a common geology which reaches back to the oldest Precambrian, justify a mutual study and description o€ itabirite iron-formations of both the continents. The first part of the study, therefore, is a summing up of the genesis of itabiritic iron ores and in the second part a short description is given of all mines producing at present.

It is no easy task to compile a summary on the sedimentation of itabirite iron-formations, their ages, degree of metamorphosis, orogenic modification, weathering and the result and formation of high grade ores on both shields. According to the political splitting of the areas on both sides of the ocean, each country started its own geological research. Thus, during the past decades, a multitude of conceptions on the geological structure of the various countries has been set up, and though they are valid for the country coiicerned, they often lack relationships to the neighbouring countries. The sedimentary and metamorphic change of facies of the Precambrian rocks, as well as their varying definitions and nomenclature, make it even more difficult to make a comparison.

In spite of this, the author collected all interesting details concerning itabirite iron-formations, and with this material he attempted to compile a kind of summary which, however, cannot be considered as complete or infallible. It only shows the present state of knowledge regarding itabirite iron-formations in those areas.

During a check-up of the details, it also became evident that the deposits’ geological investigations progressed at different speeds and so they are often incomplete. Thus, when drafting the summary, at first the well-known

areas were described and, based on these, the geology of the lesser-known areas was treated.

Stratigraphical situation and age of the Precambrian and its itabirites in the Liberia and Guyana shields The data acquired through radiometric dating methods place the sedimentation of itabirites between 2,500 and 3,000 1n.y. in the Liberian as well as in the Guyana shield. These data are mainly obtained from gneisses, which comprise the cores of both shields. This is the case with the so-called Kasila-schists of Sierra Leone (3,200 m.y.), large parts of the granitic or granulitic basements of this country (2,700-3,600 m.y.), as well as with the gneisses of the itabirite deposits Bong Range (2,910-3,280 m.y.), Mano River (2,660-3,350 m.y.) and Nimba (2,500 m.y.) in Liberia. W e also find similar ages in the Guyana shield, where the gneisses of the itabiritic Imataca-series are also dated from 2,700-2,900 m.y. (All age data are obtained by whole rock analysis based on the Rb/Sr method.)

Overlying this crystalline underground, there are more or less metamorphic sediments and igneous rocks. Contrary to most other Precambrian shields, they are not, however, separated by an unconformity from the footwall gneiss, but grade into each other depending on their degree of metamorphism. Thus, at least in the itabirite provinces of both the shields, up to now no real basement has been found.

The metasediments overlying the gneiss consist of quartzites, quartz mica schists, amphibolites and igneous rocks, interlain by ítabirites. Though in Africa as well as in South America they bear different names, for instance Kambui-series (Sierra Leone), Nimba-series (Liberia), Siniandou-series (Guinea), Imataca-series (Venezuela), they may generally be considered to have a similar age (Fig. i).

As metamorphosis of these Precambrian sediments resulted in the formation of the basal gneisses-as ascer- tained in Liberia (Bong Range) or Venezuela (1mataca)-

Unesco, 1913. Genesis of Precarribrinii iron and iitnizgnizese deposits. Proc. Kiev Syi>ip., 1970. (Earth sciences, 9.) 335

H. Gruss

LI BERIA SHIELD GUYANA SHIELD orogenesis - folding and metamorphism

700 -800 m , y. (gneisses, NW-SE striking Liberia i Sierra Leone )

sedimen tat ion . _. . . .. , . . . Torkwaien (Ghana ) (Malasse- facies)

orogenesis - metatexis - anatexis 1800 -2000 m.y.

gran ¡tes and gneisses

Roraima -Farmation 1 675 m.y.

I 14 o I o s se -Fa c i es )

metatexis -anatexis 1 800 - 2 OCO m.y. "yo un ger gro n i t es " -r

OrOgWleS¡S - folding and regional metamorphism folding and regional metamorphism C

" o1 der granit es " O

O ._ A-.

72.500 - 73.000 m.y. 2 700 -2.900 m.y.

U

O ul C O U

._ - i ta bi ri tes metosediments Barama -Mazar uni -system

Imataca - Series I sedimentation ... . . . ... .__

Nimba -series Dahomeyen

unconformity 2

basement 2 kasila -schists (3 200m.y.) "basement"( 2700-3.600 m.y,)

FIG. 1. Precambrian of the Liberia and Guyana shields.

?

the measured age of all these gneisses ought to be the same for these different series of metasediments and their itabirites. Therefore, gneisses and metasediments are defined as Precambrian I.

Already, in the early Precambrian, the sediments were intensively folded, accompanied by a more or less vigorous regional metamorphosis (Liberia-green schist facies to amphibolite facies). For this reason the geothermic gradient of this metamorphism in the Bong Range itabirite deposit (amphibolite facies) reached 5.5 kb and a temperature of 570-630" C, corresponding to a modification in a depth of 20 km.

Sporadically, this regional metamorphism was more intense and led, in both shields, to the formation of gneisses and granite intrusions (Sula Mountains, Sierra Leone, Iwokrama-granite Guyana). The coastal areas of both the shields were especially affected, whereas the intensity of the regional metamorphosis seems to diminish towards the interior. Based on investigations in Liberia, Leo and White (1968) declared that the age of this orogenesis and metamorphosis is in the Precambrian I-more precisely, lying between 2,500 and 3,000 m.y. This is in accordance with the age of intruded granites in Guyana and in the Iwokrama-series (2,595 m.y.).

It seems that in both the shields the gradient of orogenesis tends from the present coast towards the interior. This fact is valid for the area of the Imataca-series in Venezuela (S-vergence) as well as for the Liberian deposits

(Bong Range, Bomi Hill with N-vergence). In the pasts of the shields situated near the coast, the degree of metamorphism is higher and a flat folding seems to prevail, whereas isocline-type folding is likely to be found in the inland areas.

This orogenesis did not, however, cause a consolidation of the two shields, as in later periods of the Pre- cambrian thick series of igneous rocks and sediments were deposed on them (Precambrian 11). Examples are the Pastora-series of Venezuela, with an age of 1,600- 1,800 m.y., the so-called Birrimien (which is supposed to form large parts of the Eastern Liberian shield near the Ivory Coast, Ghana and Upper Volta) the age of which is stated as 1,800-2,000 m.y. (Machens, 1966).

Up to now no itabirites have been found in these middle Precambrian series of the Liberian shield, whereas, in the Guyana shield, this age could be valid for the itabirites and gondites of the Amapa area in Brazil and probably even for the itabirites of Southern Surinam.

The next discernible period in both shields is a second orogenesis and metamorphosis. It affected the rock sequences of Precambrian I as well as those of Precambrian II. It caused extended granitizations and the formation of gneisses in the Ivory Coast and in Liberia. In the Ivory Coast the age of this metamorphism is stated as 1,800- 2,000 m.y., whereas in the Bong Range itabirite deposit in Liberia it is at least 1,600 m.y. Here, through detailed mapping, it could be proved that the younger gneiss

336


border cuts across the already existing fold system and rises in anticlinal regions, while sinking in synclinal regions. Contrary to the first metamorphism, the gradient was 3.5 kb, corresponding to a depth of 13 km and a temperature of 640 to 680" C. This orogenesis and metamorphism in the Liberian shield have their counterpart in the Guyana shield within the so-called younger granites aged between 1,800 and 2,000 m.y.

Thus, in both shields, a state of rigidness was reached, as younger Precambrian sediments belong to the molasses facies which is practically unfolded, not metamorphic and lies unconformably flatly on the crystalline underground. In the Guyana shield area this is called the Roraima- formation aged at least 1,675 m.y., and in the Liberian shield it is named Tarkwaien and is found in Ghana and Upper Volta (Fig. 1).

Sedimentation and facies of itabirite iron-formations

At the moment the only detailed mapping describing the sedimentary facies relationships of the itabirites with the country rock have been carried out in Liberia, i.e. Bong Range (Stobernack, 1968), Nimba (Berge, 1968) and Goe Range (Berge, 1965). Based on these mappings, the sedimentary sequence of the strata begins everywhere with quartzites, with a thickness of several hundred metres. These sometimes become coarse-grained and conglomerate- like and show some characteristics of an itabiritic sedimentation (Bong Range, Fig. 2; Goe Range). Overlying these is generally a series of quartz-muscovite (i.e. quartz- biotite) schists which, for instance in Nimba, reach a thickness of700 m. It is not certain whether the intercalated amphibolites may be considered as igneous rocks. Similar

W

sediments also form the footwall of the itabirites in the Imataca-series of Venezuela.

After the sedimentation of these rocks, in the area of both shields a deposition of itabirites took place. However, it is evident that this happened only sporadically and the thickness varies. Thus, from the Imataca-series of Venezuela, it is known that the thickness of the itabirites reaches only a few metres in places. It is only in larger deposits that the thickness increased to several tens or hundreds of metres and was often increased due to a later folding. The following sedimentary thicknesses of itabirite have been noted: Cerro Boliva, 200 m; Bong Range, 20-80 m; Nimba, 250-400 m.

It is certainly not coincidence that the areas with a relatively thick itabirite sedimentation were later transformed into synclinoria. It is quite evident that the synclinoria are syn-sedimentary, representing former areas of subsidence and troughs, thus accumulating larger sedimentation masses than tectonically more stable areas in the neighbourhood. In general, the area of sedimentation in which the itabirites were deposited probably resembled an epicontinental shelf. It is interesting that in the Bong Range deposit (Fig. 2), for instance, the itabirites are laterally intercalated with coarse-grained quartzites and finally grade into them. At the same time there is a similar change of facies in the footwall of the itabirites where quartz-muscovite-schist gradually grades into amphibole- biotite-schist. The amphibole-biotite-schist, as well as the coarse-grained quartzites, represent a kind of a syncline- facies, while quartz-muscovite-schist and itabirites are a marginal shelf facies. Similar facies relations are-though on a much larger scale-also valid for the itabirites of the Minas-series in Brazil (Eichler, 1967; Pflug, 1967).

In the area of the Guyana and the Liberian shields the itabirites generally correspond to the oxide facies of James (1954); itabirites of the carbonate or -sulphide facies

EN F

a coarse-groined quartzite

with itabirite indications hematite-magnelile- ___ ilobirile

pegmaloid

- 5 bonded gneiss 5

au

c quortz-muscwite-sChi*t -. -. -. -. -- -. -_ -.

c metadalerite E € - - - \;. . . . . . . . . . . . . . òg "",*$y*. . . . . . . . . . . . . . - . - with ore indications

. . . . . . . . . . %\. . . . . . . . . . . . . . . . iinely banded biotite -quartzites and . . . . . . . . . . . . . . . . finely bonded quartzites

x qgb.A ELS rquorlz -bonded amphibole- biotite-schist . . . . . . . . . . . . .

FIG. 2. Bong Range: Stratigraphy and Facies (Stobernack, 1968).

337

H. Gruss

arepractically unknown. However, the oxide facies may be of a varying mineral composition. In both shields magnetite- quartz- and magnetite-hematite-quartz-itabirites prevail; with admixtures of iron silicate of varying proportions (greenalite, grünerite, cummingtonite). Phyllite-banded itabirites, however, seem to be a special facies, where quartz is represented by chlorite, sericite and amphiboles. The itabirite deposits of Tonkolili (Sierra Leone), Mano River and Wologisi (Liberia) in the Liberian shield belong to this type.

The different itabirite facies seem to be deposited according to a certain rhythmic pattern. For instance mining operations at the Cerro Bolivar deposit show very clearly that the silicate itabirites mostly appeared at the borders of the synclines (i.e. at the footwall of the itabirite sequence) whereas in the core of the synclines (i.e. in the higher parts of the itabirite sequence) silicate-free itabirites prevail. The same observation could be made in the so- called Northern deposit of Bong Range (Liberia), a relatively flatly folded part of the deposit. There, the itabirites, which have a thickness of 73 m, are divided into three successive zones, in each of which the following change of facies gradually proceeds:

Footwall Hanging wall low grade Fe +- high grade Fe high grade Fe-silicate -+ Fe-silicate-free high grade magnetite + low grade magnetite low grade hematite -+ high grade hematite

Consequently, the degree of oxidation gradually increases from the base towards the hanging wall, where either an interruption of the sedimentation took place after which the deposition started anew, or the sedimentation suddenly encountered changed conditions which mark the beginning of every cycle.

Similar cycles axe known from itabirite series of other Precambrian shields, such as the Minas-series in Brazil (Eichler, 1968; Gruss, 1966) or from Canada (Goodwin, 1956), and they find their counterpart in the oolitic minette- ores of Lothringen (Bubenicek, 1960).

The stratigraphic hanging wall of the itabirites is again schists with varying degrees of metamorphism, as for instance, in the Bong Range and Nimba deposits. Because of the exposed position of the present itabirite outcrops, these strata have already been eroded, thus no exact data on their facies and thickness proportions can be obtained today. Nevertheless, being the youngest known rocks of the Precambrian I, they are of stratigraphical interest.

Formation of high grade ores through metamorphic differentiation

The economic significance of the itabirites in both shields for the world's iron ore industry lies in the occurrence of large deposits of high grade ores. These, however, have not been formed by sedimentation, but originate from

epigenetically modified itabirite iron-formations. Princi- pally, two types of high grade ore can be distinguished. The first was already of economical interest thirty years ago, and it was formed by metamorphic differentiation. As already explained, the sedimentary sequence of Pre- cambrian I with its itabirites underwent two orogeneses and periods of metamorphism in its history. The meta- morphisms were especially pronounced in the central part of the orogene which today forms the coastal areas of both shields. In these regions theitabirites sometimes came into direct contact with risinggneissic fronts and deep seated intrusions of basic plutons. In such cases a metamorphic conversion of the itabirites into high grade ores took place. This compositional change-as a result of increased pressure and temperature conditions-sometimes caused the mobilization and removal of silica, whilst the itabirite's content of magnetite-hematite was residually enriched and, after a recrystallization, formed massive, high grade ore bodies. Typical examples are the deposits of Bomi Hill (Liberia) and El Pao (Venezuela).

In Bomi Hill (Fig. 13) alternating sedimentation of itabirites and chlorite schists of approximately 450 m thickness form a flat, east-west striking and north-vergent syncline, which is lying directly on granite gneiss. Obser- vation has shown, that the removal of silica and enrichment of ore started metasomatically at the contact with the granite and continued along the stratification. It was combined with an extensive alkali-metasomatism and the dissolved silica was partly precipitated in overlying schists, The result of this metamorphic differentiation of an itabirite sequence, whose original thickness was about 80-100 m, is a 30 m thick layer of compact magnetite ore, which shows the same structure as the former itabirite. The overlying itabirite layers, which did not come into contact with the granite and were separated from the basement by about 20-50 m chlorite schist, however, were not influvnced by the metamorphic differentiation.

The same conditions can be found in the El Pao deposit of Venezuela (Figs. 7 and 8). There, an itabirite formation with a thickness of about 100 m is folded by a system of EW.-NS.-striking synclines into an underground of gneiss and granite. This direct contact of granite and itabirite did not cause the metamorphic differentiation of the latter, which is due to a later intrusion of a gabbroid magma. When rising, the magma nearly always followed the contacts of itabirites and granites and caused the same metamorphic differentiation as described for Bomi Hill. The geological map of El Pao (Fig. 7) clearly shows that the high grade ores are bound to the contacts of gabbro with itabirite and not to those of itabirite with granite. This is also demonstrated by the cross-section through the deposit (Fig. 8) showing a flat itabirite syncline, the core of which consists of gabbro. Here, the itabirites were not influenced at the footwall, only the superior part of the itabirite sequence next to the gabbro was transformed into high grade ore.

So far no reliable particulars can be given as to the age of the metamorphic formation of high grade ores.

338

Most likely, it was the second orogeny and metamorphism which is responsible for such metamorphic formation. Thus the gneiss formation of the Bong Range deposit which is next to Bomi Hill, can be estimated at about 1,600 m.y. Certain dolerites of the Guyana shield (Roraima Plateau) show about the same age, and it seems possible that there is a relation between them and deeper situated gabbro intrusions like that of El Pao.

Formation of supergene high grade ores

More frequently another type of high grade iron ore is found, which for several decades has been generally known as supergene for,mation. However, the first qualitative and quantitative study of this ore was carried out at Cerro Bolivar in Venezuela by Ruckmick (1963). Meanwhile, similar investigations were carried out in Minas Gerais (Eichler, 1967), which in general verify the results obtained by Ruckmick. Based on these results, the formation of high grade ores from itabirites is due mainly to a removal of the silica by rain or subsoil water and a relative upgrading of iron and alumina as residual formation.

In both shield areas tropical climatic conditions have been prevalent during the youngest periods of the geological history. Today, an annual rainfall of 3,000 mm is measured in the coastal region of the Liberian shield, which gradually diminished farther inland. Similar conditions are also encountered in the Guyana shield where, for instance, 1,700 mm are measured for the Cerro Bolivar area. These rainfalls also affect the outcrop of itabirite iron-formations. Here especially the silica is leached by way of hydrolysis because the rain-water, containing only small amounts of carbon dioxide, is able to dissolve a considerable quantity of silica. However, the portions of iron and alumina dissolved are relatively small (Table 1).

Thus the solubility of silica depends mainly on the solvent properties of the available quartz surface i.e. grain size. Of further importance is the period of time in which the rock becomes affected by water. Eichler's (1968) research in Minas Gerais, Brazil, show that it takes twenty-

TABLE i. Quantities of SiO,, Fe and Al leached by rain-water

Area Grain size of the itabirite mg/' mg/l mg/l

Fe AI (mm) SiO?

Cerro Bolivar (Ruckmick, 1963) 0.05-0.15 10-15 0.05-0.1 Unknown

Minas Gerais, Brazil (Eichler, 1967) 0.05-0.1 6.20 0.34 1.95

Minas Gerais, Brazil (Eichler, 1967) 0.5-1.5 1.60 0.14 0.84


(a) Rainfall and temperatures. Solubility of SiO? in ground water'

5 O0 rnm

40 o

300

200

100

max.OC -a9..i

I VI XII

(b) Solubility of itabirite in pure rain-water

3OoC -

2ooc

IOOC

1Lpprn SI02

12

10

8

6

4

2

O

. ,

28.1X.1965 O 6.75 + 260 14.2 30.1X.1965 2 6.45 260 16.1 11.X.1965 13 5.80 310 trace - - 20.0 12.X.1965 14 5.78 415 19.8 23.X.1965 25 5.65 442 5.60 0.24 0.38 20.1 1.111.1966 153 4.45 400 6.20 0.34 1.95 20.5

FIG. 3. Solubility of silica in subsoil waters of Minas Gerais, Brazil (Eichler, 1968).

five days for the waters to dissolve a considerable quantity of Sioz, and it is not during the period of maximum rainfall that subsoil waters contain most of the dissolved silica (0.5-2.0 mg/l), but towards the end of the dry period (8-10 mg/l) (Fig. 3).

The result of this leaching of itabirites is a weathering profile with a typical zonal structure (Fig. 4 and Table 2).

The description of the weathering profile shows that the composition of weathering residue of zones A and B is identical with the supergene high grade ores, and is mainly dependent on the composition of the primary rocks, i.e. on itabirite facies.

The supergene high grade ores of both the Guyana and Liberian shields rarely occur in the coastal areas, but are mostly found about 200 km inland. This is especially true for Nimba and Simandou deposits in Liberia and Guinea, as well as for Cerro Bolivar and San Isidro deposits in Venezuela.

Considering the dependence on the grain size of the itabirites of quartz leaching, this geographical distribution has been caused by the varying degrees of metamorphosis. High grade metamorphism resulted in coarse-grain sizes

339

H. Gruss

TABLE 2. Weathering profile

Zone On silicate-itabirites On oxide-itabirites

Crust of limonitic cemented tabular hematites, highly clayey, sometimes compact, about 58 per cent Fe

tabular hematite and limonite crusts, mainly clayey matrix, brown, 55 per cent Fe

itabirite, 40-50 per cent Fe

unweathered, hard, containing mica and amphibole, magnetitic, 35 per cent Fe

Loose admixture of

Clayey, brown detrital

Silicate-i tabirite,

Crust of limonitic-hydro- hematitic-cemented tabular hematites, compact, but porous, 62-67 per cent Fe

Loose admixture of tabular hematites, in powdery matrix brown, black, blue, 63-69 per cent Fe

Loose, quartz-rich detrital itabirite, 40-50 per cent Fe

unweathered hard, hematitic or magnetitic, 35-40 per cent Fe

Oxide-itabirite,

and small grain surface per unit volume in coastal itabirites. Corresponding to this, quartz leaching was not intensive but extensive, i.e. deep, and consequently only little upgraded cappings of weathered itabirites were formed there. However, these may be enriched to high grade concentrates at comparatively low cost (Bong Range, Liberia; Marampa, Sierra Leone; Maria Luisa and Piacoa, Venezuela). In the itabirite deposits located further inland and characterized by lower metamorphism and small grain sizes, intensive leaching and upgrading prevail, producing direct-shipping high grade ores (Nimba, Liberia-Guinea; Simandou, Guinea; Cerro Bolivar and San Isidro, Venezuela).

In spite of its proved efficiency, the solubility of the silica is small and, therefore, it is important to know in which period of time high grade ores were formed. Ruckmick (1963) states that the formation of the high grade ores at Cerro Bolivar began about 24 m.y. ago, in other words, it is younger than Oligocene, in any case not older than Cretaceous.

This date has been verified by other geological investigations: the high grade ores of Cerro Bolivar as well

cementation ore

ific

FIG. 4. Typical weathering profile of itabirite iron ores in tropical climates (Thienhaus, 1963).

340

Itabirite iron ores of the Liberia and Guyana shields %

Grands Rochers sw

Mt Piérre Richoud NE

I -

FIG. 5. Blue high grade ores of Nimba, Guinea (Gaertner, 1961).

as those of San Isidro are related to an old peneplain, relicts of which are still found on the ore mountains. This levelling corresponds to the old Gondwana Peneplain which, in the Cerro Bolivar area, still has an altitude of 700-750 m , but shows an incline of 1” towards the north and is covered with sediments of the Neocomian (lower Cretaceous) and younger sediments north of the Orinoco river.

This leads to the assumption that during the lower Cretaceous the Gondwana-Peneplain was lying horizontally near sea level and, therefore, the Neocomian sea could transgress over it. After this an elevation to the present level must have taken place, followed by erosion and intersection of the peneplain and lowering of the water table.

Most probably the ore formation started at the same time and, according to Ruckmick, might be Oligocene.

The accuracy of this dating might be verified by the studies of King (1957), according to whom the north- east Brazilian shield was elevated during the early Tertiary until the Miocene, causing an erosion cycle (Sulamaricano cycle) and according to Eichler (1967)-led to the formation of the supergene high grade ores in the ‘Iron Quadrangle’ of Minas Gerais, Brazil. After detailed calculations Eichler also came to the conclusion that the ore formation began during the Oligocene (26 n1.y. ago).

Similar conditions can be expected in the Liberian shield, where in the Nimba Mountains of Liberia and Guinea supergene high grade ores are levelling with an

I600 m

1500 m

1L00m

1300 m

altitude of 1,100-1,300 m, and may be considered as a relict of the Gondwana Peneplain. In the more coastal itabirites of Putu, Liberia, the sanie ores can be found at an altitude of 700 m . Thus, in the Liberian shield too, an inclination of the Gondwana Peneplain towards the coast is indicated.

However, this theory is complicated by another type of high grade ore which is similar in physical composition and grain size to parts B, C and D of the normal weathering profile (Fig. 4), but shows an Fe content which is generally 2 per cent higher than the one of zone B, with less alumina, and is marked by a high amount of secondary hematite. These ores are metallic blue and can easily be distinguished from the so-called brown and black supergene high grades ores. They rather resemble the so-called ‘hematite ores’ of the ‘Iron Quadrangle’ in Brazil whose formation is considered to be hypogene-metasomatic (Dorr, 1959; Eichler, 1968), or those of north-west Australia, representing pure weathering formations (MacLeod, 1966). This already shows the different opinions regarding the genesis of the blue high grade ores in general as well as those of Nimba and Simandou, Guinea.

Gaertner (1961), after geomorphological studies on the position of high grade ore bodies in the northern Nimba Mountains and Simandou chain in Guinea, came to the conclusion that the blue high grade ores are bound to a plateau with an altitude of 1,600-1,650 m and are dislocated by younger faulting (Fig. 5). The binding of the blue high grade ores to a higher and older plateau than

341

H. Gruss

1 O00 k m 1 I

Precambrian shields '0 Itabirites

0 Roraima -formation / Tarkwaien A high-grade ores of metamorphous differentiation

EEl striking of structures high-grncle ores of alteration

FIG. 6. Itabirite iron ores of the Liberia and Guyana shields (Gruss, 1966).

342


the Gondwana Peneplain, strengthens the argument that the blue high grade ores were formed by weathering and belong to a Precretaceous cycle. The tectonic dislocation can also be seen as a proof for the greater age of the blue high grade ores.

The question as to what caused the varying mineral content (recrystallization of hematite) and the low alumina content in the blue ores still remains. So far, at least in West Africa, no corresponding studies have been made. However, investigations by Eichler (1967) carried out in Brazil give some details regarding the alumina content in subsoil waters. H e states that by hydrolysis of itabirites not only silica can be dissolved, but iron and also considerable quantities of alumina. Thus, the blue ores of higher levelling might be considered as more mature, than the supergene high grade ores of the younger Gondwana Peneplain. This is also in accordance with the results of morphological studies. Regarding the recrystallization of hematite, MacLeod's (1966) investigations are interesting. In higher parts of the weathering section, the cementation of supergene high grade ores in north-west Australia is mainly limonitic, in lower parts, however, hematitic. The author considers the blue high grade ores of the Nimba Mountains and Simandou chain as weathering formation, and he is of the opinion that the recrystallization and cementation of the high grade ores with hematite does not necessarily prove hydrogenic-metasomatic procedures. With regard to the brown and black weathered high grade ores, it should be examined whether the physico-chemical conditions in the roots of blue high grade ores, which may have a depth of several hundred metres, permit the recrystallization of secondary hematites.

Geological relations between the itabirites of the Guyana and Liberian shield

By comparing interpretations of facts presented in the previous sections, it is clear that the geology of the iron- formations of both shields is nearly identical. It would be interesting to pay special attention to these relations, enabling the corresponding inferences to be drawn.

The itabirites (Fig. 6), as well as their associated formations, belong to Precambrian I. They were deposited in a shelf-like sea area 2,500-3,000 m.y. ago, and towards the end of this period were affected for the first time by an orogenic folding and metamorphosis. The cores of this orogene can today be found in the coastal areas of both shields. The vergence of this folding was directed towards the present inland areas. The second metamorphic modification affected both shields about 1,800 m.y. ago. In both shields the results of these transformations are the mostly coarse-grained itabirite deposits near the coast and the fine-grained itabirites farther inland. At the same time the formation of itabirite high grade ores took place by metamorphic differentiation in the central parts of the

orogene. As the metamorphosis in the marginal parts was less effective, no high grade ores were formed there. The finer-grained itabirites of these areas were predestined for the formation of supergene high grade ores, which are bound to the Gondwana Peneplain or older levellings of Precretaceous time. Besides the synchronous geological events, a remarkable symmetric structure for both shields can be observed, as for instance in the vergence of folding, zones of the same metamorphic grade and the distribution of ,different itabirite formations and their high grade ores. These facts and the argument that coastal as well as tectonic structures of both shields fit perfectly together, may, therefore, be considered as a proof that in Precretaceous times the Guyana and Liberian shields formed a single unit and that at least the itabiritic provinces of the shields belonged to the same sedimentation basin which later was developed as geosyncline. During two orogenics this geosyncline was folded into a mountain chain with a symmetric structure and a marked crest zone, along which the orogene was divided into the Liberian and Guyana shields when the Gondwana Continent disintegrated during lower Cre- taceous period.

El Pao (Venezuela)

The El Pao iron ore deposit (Figs. 7 and 8) has been known since 1926, but it was not until 1950 that Iron Mines Company of Venezuela was able to start full mining operations.

The average analysis of the reserves is almost the same as the analysis of the shipped ore: 62.6 per cent Fe; 1- 2.5 per cent SO,; 3.5-4 per cent Alzo,; 0.06 per cent P; 3.66 per cent ignition loss.

Contrary to the Cerro Bolivar and San Isidro deposits, El Pao contains high grade itabirite ores formed by metamorphic differentiation. They form two flat synclines, one striking N. 80" E. and covering an area of 1,000 X 500 m, and a maximum depth of 350 m . This east-west striking syncline is followed, towards the north, by another striking N. 20" E. and extending over an area of 700 x 500 m . Both synclines are parts of two main folding-directions, forming sort of a lattice (Fig. 7).

As shown by the geological map and section (Fig. 8), the metasomatic mineralization of the deposit is always bound to the contacts of itabirite and intrusive gabbro. Thus, the hard ore body reaches a thickness of 10-50 m , with underlying high grade metamorphic itabirites , while the hanging wall forms an intrusive, medium-grained gabbro (norite). The latter ñlls the whole trough circumscribed by the hard ore body. The metasomatic high grade ores consist of hematite and magnetite in varying proportions with grain sizes up to 10 mm . The following analysis is typical: 67.5-71.0 per cent Fe, 0.1-0.7 per cent Sioz, 0.1-4.0 per cent Alzo3, 0.01-0.1 per cent S, 0.01-0.03 per cent P.

All rocks are marked on the outcrop by a deep alteration, especially the itabirites. Due to high grade metamorphism and the coarse-grain sizes (1-5 mm), weathering was

343

H. Gruss

1- gabbrolnorite [-I itabirites 1 km = lump ore I ] gneiss

t I

FIG. 7. Geological m a p of El fao, Venezuela (Iron Mines of Venezuela, S.A.).

344


I ,pre-mining surfoce North

I ---___ South c _ c - - - - - - _ _ /

er t + +

' + I L + , * t l L C

+ + +

+ & +

+ A + + i .L + + L

+ + I + +

+ + + r i

+ + + + +

* * +

c -

+ I + + + + . *

lumo ore

itabirites.weathered

gabbro a: unweathered b:weathered

gneiss a: Unweathered b;weathered

FIG. 8. El Pao, Venezuela: cross-section L (Iron Mines of Venezuela, S.A.).

not intensive, but extensive, i.e. it had a deep reaching effect, Thus, there was no formation of 'genuine' supergene high grade ores, but only concentrations, which are typical for zone C of the alteration profile of itabirites, i.e. the formation of siliceous fine ores. For a cut-off grade of 56 per cent Fe these ores show the following analysis: 56-62 per cent Fe, 6-10 per cent Sioz, 2-4 per cent Alzo3, 2-5 per cent ignition loss.

O n the average 1.16 tons of overburden per ton of shipping ore are to be moved. During mining operations both types of ore are mined simultaneously, crushed and screened, thus producing a direct-shipping ore as described at the beginning and showing the following grain sizes: more than 51 mm (20.33 per cent), 13-51 mm (26.26 per cent), less than 13 mm (53.41 per cent).

Cerro Bolivar (Venezuela)

This deposit was discovered at the beginning of the forties. Since 1954 it has been exploited by the Orinoco Mining Company. At a cut-off grade of 55 per cent Fe, the average analysis is: 63.84 per cent Fe, 1.86 per cent Sioz, 1.44 per cent Alzo3, 0.10 per cent P, 5.11 per cent ignition loss.

These deposits represent the relicts of a synclinorium of itabirite-bearing metasediments, which reaches from Cerro Bolivar 80 km east to the Rio Caroni. The supergene high grade ores are bound to the old Gondwana Peneplain which cuts the island mountains with its itabirite outcrops at approximately 700 m above sea level.

The Cerro Bolivar deposit has a strike length of 20 km with outcrops up to 750 m wide. In this area an itabirite formation with 200 m of sedimentary thickness is isoclinally folded. The special synclines staggered to the

I l I 1200rn

- 600 m

- 200m

riglit can reach a depth of 200-250 m , divided by steeply rising anticlines of footwall-schists (Figs. 9 and 10). The iron ores belonging to the brown and black type of weathered high grade ores, show the typical profile already described.

The unweathered rock (zone D) consists of fine to coarse-banded (0.05-2.0 cmj itabirites with 39 per cent Fe and 42 per cent SiO, on average, and grain sizes of between 0.05 and 0.15 m m . The main iron mineral, besides magnetite, is specularite. However, the majority of the itabirites also contain Fe silicates as muscovite, sericite and, less frequently, amphiboles and pyroxenes.

The fresh rock is overlain by a zone (C) of soft itabirites , which often is no more than 10 m thick. Technically, two types are distinguished siliceous fine ores (50-62 per cent Fe, 6-10 per cent Sioz, 1 per cent Alzo, and 3 per cent ignition loss); soft itabirites (45-55 per cent Fe, approximately 30 per cent Sioz, 0.5 per cent Alzo3 and 1.5 per cent ignition loss).

However, with a maximum 100 m depth, zone B is much thicker, consisting of black and brown supergene high grade ores, the black ores resulting from mostly non-silicate itabirites, the brown ones from silicate-bearing itabirites. The following analyses are characteristic: brown fine ores (62-64 per cent Fe, 0-6 per cent Sioz, 1 per cent Alzo3 and 3 per cent ignition loss); black fine ores (66- 68 per cent Fe, 0-6 per cent SO,, 1 per cent Alzo3 and 0-3 per cent ignition loss).

Experience shows that brown fine ores mostly appear on the rims of a syncline, while the black fine ores predominate in the centre. About two-thirds of the reserves of zone B consist of brown ores, the rest of black ores.

At Cerro Bolivar the surface ores of zone A are 10-30 m thick. Depending on intensity of weathering and composition of the primary rock, the hard and lumpy material

345

H. Gruss

D laterite = high-grade ores itabirites

FIG. 9. Geological map of Cerro Bolivar, Venezuela (Orinoco Mining Company).

NW

7fim

SOOm - 500m -

SE

EEBI crustal ores - black fines brown fines

O laterite, a itabirites

FIG. 10. Cerro Bolivar, Venezuela: cross-section A-B (Orinoco Mining Company).

TABLE 3.

Crustal Fine ore black Inch/mesh mm ore crushed Fine Ore

-100 mm brown

1.050 26.6 6.72 0.742 18.85 12.53 0.525 13.33 23.59 0.371 9.42 36.05 3 6.68 44.47 6 3.23 54.39 10 1.65 63.31 20 0.83 71.25 35 0.42 79.99 65 0.21 89.70 100 O. 147 92.94 200 0.074 100.00 3

30.43 3.43 6.96

43.47 14.92 - 26.97 48.80 33.01 64.58 46.16 80.93 57.34 89.33 66.15 91.57 73.25 93.10 79.16 93.90 81.93

i 00.00 100.00

-

A

200m

shows the following cheniical analysis: crustal ores (62- 69 per cent Fe, 0.1-6 per cent Sioz, 0.1-1.5 per cent AI& and 0-5 per cent ignition loss).

In order to guarantee a sufficient grade control during mining operations, the four zones described above are subdivided into thirty-seven ore-types which differ more or less with regard to hardness, mineralogical composition, colour and chemical analysis.

Table 3 shows the grain size distribution after screening.

346


San Isidro (Venezuela) This mine, which is situated only 15 km south of the Cerro Bolivar deposit, was discovered in 1948 and since then has belonged to the Venezuelan Ministry of Mines and Hydro- carbons (State Reservation).

The deposits of supergene high grade ores are based on a synclinorium of itabirite-bearing metasediments, whose steep special folds mostly strike WSW .-ENE. and integrate with a north-south striking system. The individual deposits cover a total area of 50 kmz, and the relation to the Gond- wana Peneplain at 700 m above sea level is clearly evident. Again, analogous to Cerro Bolivar, the substance of the deposit consists of supergene high grade ore of the brown and black type. Thus, the fresh itabirite rock of zone D presents a hard, mostly fine-banded (millimetres), sometimes also unbanded hematite/magnetite-quartzite (H : M : Q = 38 : 22 : 39 weight per cent), which also contains some iron silicate. Grain sizes range between 0.03 and 0.2 111111. The unweathered itabirite contains approximately 42 per cent Fe and 39 per cent SiOz on average.

Due to the fine-grain of the itabirite, the overlying zone C is rather thin (10-20 mm). The following types of iron ore can be distinguished: siliceous fine ores (58 per

cent Fe, 4 per cent Sioz, 0.5 per cent Alzo3 and 5.0 per cent ignition loss); soft itabirites (50 per cent Fe, 30 per cent Sioz, 0.3 per cent Alzo, and 2.8 per cent ignition loss).

Zone B of San Isidro is much better developed than that of Cerro Bolivar, and reaches a maximum depth of 240 m (Figs. 11 and 12). Here, too-depending on the content of Fe silicate in the itabirites-black and brown fine ores can be distinguished, which, at a cut-off grade of 58 per cent Fe, show the following average analyses: brown fine ores (62 per cent Fe; 2.8 per cent Sioz, 0.5 per cent Alzo3 and 4.0 per cent ignition loss); black fine ores (67 per cent Fe, 0.8 per cent SiO,, 0.5 per cent Alzo, and 2.8 per cent ignition loss).

Contrary to Cerro Bolivar, the black fine ores prevail at San Isidro in a ratio of black to brown ores of 2 : 1.

At San Isidro the limonite crustal ores of zone A generally have a thickness of 10 m, with individual roots reaching to a depth of 30 m. At a cut-off grade of 58 per cent Fe they show the following analysis: crustal ores (62-67 per cent Fe, 0.6-1.3 per cent Sioz, 0.5-1.3 per cent Alzo3, 2.5-4.3 per cent ignition loss).

Based on the above analyses and the distribution of reserves the following average composition can be calculated for the main deposit of San Isidro: 58 per cent Fe

%

x

O 0.5 1.0 Km

%

FIG. 11. Sketch map of iron ore deposit San Isidro, Venezuela (Ministerio de Minas e hydrocarbones de Venezuela).

347

H. Gruss

P 3-5 NW SE

- 500 m

-4 crustal ores _. black fines

brown fines itabirites

III laterite

cut-off (65.14 per cent Fe, 1.23 per cent %Oe, 0.59 per cent Alzo3 and 3 .O5 per cent ignition loss); 55 per cent Fe cut-off (63.3 per cent Fe, 3.0 per cent SiO,, 0.6 per cent Alzo3 and 3.2 per cent ignition loss); 0.03 per cent Mn, 0.05 per cent Tio,, 0.03 per cent P, 0.01 per cent S.

Grain size distribution can be expected to be as in Table 4.

During mining operations 0.05 tons of overburden are to be moved for each ton of ore (20 : 1).

Present plans of the Ministry of Mines and Hydro- carbons provide for a large-scale development of the deposit

TABLE 4.

Inch/mesh mm Total percentage

1.050 0.742 0.525 0.371 3 6 10 20 35 65

1 O0 200

26.6 18.85 13.33 9.42 6.68 3.23 1.65 0.83 0.42 0.21 0.147 0.074

8.2

20.5

37.1 50.2 60.8 70.7 77.1 82.2 85.3 100.0

-

-

200 m

FIG. 12. San Isidro, Venezuela: cross-section 27

of San Isidro, so that beginning in 1972, 4.2 million tons per year will be mined; from 1973, 2.5 million tons per year of this tonnage are to be delivered as pellets.

Bomi Hill (Liberia)

The Bomi Hill iron ore deposit in Liberia has been known since the beginning of the thirties, when for the first time it was geologically investigated by a Dutch firm. After the second World War the Liberia Mining Company Ltd bought the mining concession for the deposit and starting mining in 1951.

The direct-shipping ore has the following chemical composition: 64.5 per cent Fe; 4.5 per cent SiOs; 1.5 per cent Alzo3; 0.13 per cent P; 0.12 per cent S.

Of these ores, 53 per cent is lump ore (11-37 mm) and 47 per cent fines (minus 11 mm). In addition, the mine disposes of larger reserves of itabiritic low grade ores. If weathered and suitable for grinding, they can be upgraded by dressing (Humphrey Spirals and magnetic separator) to sinterfeed concentrates. The concentrate has following analysis: 64.0 per cent Fe; 6.0 per cent SiO,; 1.0 per cent Alzo8; 0.04-0.05 per cent P; 0.08-0.12 per cent S.

The Bomi Hill deposit (Fig. 13) represents an east-west- striking syncline with a steeply dipping southern limb and a flatly dipping northern limb. The syncline extends over 500 x 1,000 m and has a depth of 180 m. Its core consists of a series of itabirite-bearing metasediments, the basement

348


- s-

-N- lgoo FT 800

700

600 . . . . . _ _ . - - soo

LOO

300

200

100

FIG. 13. Bomi Hill: North-south cross-section through central part of main deposit (Zigtema, 1968).

of which is bordered by younger granite-gneiss. Directly contacting the granite, there is an ore body averaging 40 m thick, composed of coarse, magnetitic high grade ore formed by a metamorphic differentiation. The hanging wall is formed by about 40 m of schist, 60 m of itabiritic low grade ores and again up to 60 m of schist, which are all removed as overburden and get only partly dressed. Besides this main deposit, in the continuation of the strike there are several smaller deposits, the main reserves of which are

also nined today. In the main deposit ore and overburden are in the ratio of 1 : 3.6.

Bon!? Range (Liberia)

The Bong Range itabirite deposit (Figs. 14 and 15) was discovered about the end of the thirties and has been worked since 1965 under the management of Bon Mining Company.

O low rn I

Upper Ouartz-Biolile -Schist

Itobirite Coarse-groined Ouorlzite

Quartzbonded Amphibole -Schist Sillimanite-Schist

[SJTI Granitoid Gneiss Lower Cuortz -eio:ite -Schist

Ouorlz-Muscovite-Schist

Banded Gneiss

Amphibole -8iotile -Schist

Gneiss Front -+-- Anticline

/ Syncline fine-banded Biotite -0umtzite

fine-bonded Ouortzile

FIG. 14. Bong Range: geological map (Stobernack, 1968).

349

H. Gruss

E SE

m / 1.1 PI Bong Peak

LOO 400

200 200.

O 0

D N

m/NN

400

200

O

B NNW

C

Eastern / /' '. Zoweoh I

Northern Deposit

____.

m/NN

LOO

200

O

Western Zaweoh I

A SSE

m/NN

LOO -400 INor thern

- 200

O,

FIG. 15. Bong Range: cross-sections (Stobernack, 1968).

350


TABLE 5.

Probable Proved (million tons) Possible

- - 1. Zaweah I 232 2. Zaweah II - 60 3. BongPeak - 128 - 4. Gomma 5. Northern deposit 98

-

15 - - - -

- __ - Bong Range 330 188 15

The Bong Range area comprises four individual deposits, with geological reserves calculated and estimated as in Table 5.

A total of 275 million tons of the reserves proved for Zaweah I and the Northern deposit, are mineable, with 235 million tons still to be exploited on 1 January 1970.

The four ore bodies (1-4 above) form a single itabirite syncline striking east-west and extending over 13 km. This syncline is steeply folded into foot-wall schists and reaches an outcrop width of up to 300 m, mainly in the western part. In front of the western end, and towards the north, the so-called Northern deposit is situated. It forms only a rather flat syncline. This ore body shows that the sedimentary thickness of the itabirite formation is not more than 80 m, and that larger widths of outcrop are due to steeply isoclinal folding.

The itabirites, which are mostly of the mesozonal- metamorphic type, have an average grain size of 0.1 mm, and belong to the oxide facies. Width of banding is, in general, between 1 and 10 mm. The primary mineral stock is formed by magnetite, hematite and quartz, in addition to varying proportions of iron silicates, as e.g. biotite, cummingtonite and grunerite. Due to their relative coarseness, the itabirites have undergone a deep weathering which caused an oxidation of the magnetite (martitization) and a loosening of the rock bond, which, however, did not result in the formation of high grade ores. Thus, only zones C and D of the Bong Range deposit represent the characteristic profile of weathering. From this and also from the operating point of view, the following types of ore can be distinguished, starting from the top of the profile: 1. Spiral ores: zone C, soft itabirite, weathered; 11.5 per

cent of the proved reserves; 42.6 per cent Fe, 7.1 per cent magnetite, 37.1 per cent Sioz, 0.5 per cent Alzo3, 0.05 per cent P, 0.008 per cent S; 80 per cent-0.25 m m .

2. Ttmsitional ores: zone C, medium hard itabirite, slightly weathered; 13.5 per cent of the proved reserves; 40.6 per cent Fe, 12,O per cent magnetite, 40.0 per cent SiO,, 0.6 par cent Alzo,, 0.03 per cent P, 0.01 per cent S; 90 per cent-0.1 m m .

3. Magfietic ores: zone D, hard itabirite, unweathered; 75 per cent of the proved reserves; 37.4 per cent Fe, 35.2 per cent magnetite, 42.0 per cent SiO,, 0.4 per cent Alzo3, 0.04 per cent P, 0.03 per cent S; 90 per cent -0.1 mm.

The crude ore is mined by modern open-pit methods (overburden ratio is 1 ton: 0.5-1.0 ton), crushed, ground to liberation size and upgraded by means of Humphreys spirals and magnetic separators to a high grade concentrate. For a weight recovery of 4244 per cent and an Fe recovery of 70-74 per cent, the average Bong Range concentrate analysis is: 65.16 per cent Fe, 9.64 per cent Feo, 7.00 per cent SO,, 0.28 per cent Al,O,, 0.034 per cent P, 0.022 per cent S, 0.05 per cent Mn, 0.05 per cent Cao, 0.06 per cent MgO, 0.00 per cent Cu, 0.60 pur cent ignition loss and 4.76 per cent moisture.

Nimba (Liberia)

The Nimba deposit (Figs. 16, 17 and is), considered to be the largest iron ore mine in Africa at present, is managed by Lamco Joint Venture Operating Company.

The high grade ores of the Nimba Mountains originate from itabirites of the oxide facies which, in general, are fine-banded (0.5-5 nim) and fine-grained (grain size 0.03- 0.1 mm). The sedimentary thickness of the itabirites ranges from 250 to 400 m, the width of outcrop being often increased by isoclinal folding. Ore minerals are almost exclusively magnetite and hematite, while iron silicates are negligible. The fine-grained itabirites have undergone a sometimes deep weathering during their geological history resulting in the formation of high grade ores. Thus, high grade ores of two weathering cycles can be distinguished, the ones bound to the Cretaceous Gondwana Peneplain (f 1,300 m above sea level) and others originating from older, higher situated levellings (+ 1,600 m above sea level).

The high grade ores of the Gondwana Peneplain usually form flat caps, reaching a depth of 75-100 m and representing thecomplete typical profile of supergene high grade ores.

The cementation ores of zone A are 2-5 m thick on average, but sometimes also maintain a depth of 15-20 m . The hard, porous, limonitic ore has the following composition: 63.5 per cent Fe, 0.8 per cent SiO,, 2.8 per cent Alzo3, 5.5 per cent ignition loss.

The brown’ ores of zone B contain a high proportion of fines and belong to the type of the brown (f black) fine ores; their thickness may reach as much as 100 m. The following chemical composition is typical: 65.5 per cent Fe, 1.5 per cent Sioz, 0.8 per cent Alzo,, 4.0 per cent ignition loss.

Here too, the brown varieties seem to originate from iron silicate-bearing itabirites, whilst the black fine ores stem from itabirites free of iron silicate.

The foot wall of the brown fine ores is formed by soft itabirites of zone C and is rather thin. These siliceous fine ores have the following composition: 50-60 per cent Fe, 10-20 per cent Sioz, 1 per cent Al,O,,

The fine-grained, hard itabirites of zone D give the following analysis: 38 per cent Fe, 42 per cent Sioz, 0.5 per cent Alzo3, 1.5 per cent ignition loss.

Compared with the brown high grade ores bound to the Gondwana Peneplain, the blue high grade ores of older

351

BLUE ORES BROWN ORES

-1 MT ALPHA PHYLLITE ['.='.:I NIMBA ITABIRITE 7 4 GBAHM RIDGE PHYLLITE DZi SEKA VALLEY AMPHIBOLE

HIGH -GRADE ORES

NIMBA SERIES

SCHIST YEKEPA SERIES

FIG. 16. Nimba area/Liberia (Lanco J. V. Co.) (Berge, 1968).

352

NW S E NIMBA SERIES

YEKEPA ! I

PROFILE 5 S O U T H CENTRAL N I M B A HILL

PROFILE 13 N O R T H GBAHM-NIMBA RIDGE

FIG. 17. Nimba area: geologic cross-sections (Berge, 1968).

-NW-

Bh

I / hard ore /

1..1..1 soft ore itabiriie

I__) schist


NW S E NIMBA SERIES

I I -3

\ /- - YEKEPA ; SERIES ‘\ /I \

I L’ ! 1300 rn

l500m

PROFILE 8 N O R T H C E N T R A L G B A H M G U E S T HOUSE HILL

MT. ALPHA PHYLLITE NIMBA ITABIRITE G B A H M RIDGE PHYLLITE S E K A VALLEY AMPHIBOLE SCHIST. YEKEPA SERIES

O 1000 2000 3000rn I 1

-CE- Bh.25

FIG. 18. Nimba: section across main ore body (Thienhaus, 1963).

100 rn

353

/- ,.--' g Contact ,P' Strike and direction of dip of foliotion , , ,' _______ Probable fault , ,4' Vertical foliotion t -__ Anticline

..:;.::.. . .. 1 K m L-- Syncline j2::v,..i! Limit of open pi1 workings

A. Radiometric age locality I I t O

FIG. 19. Geologic map of the Mano River Mine area, Grand Cape Mount County, Liberia (White and Baker, 1968).


z LT

4 U w !x .a

5 m ,z

TABLE 6.

100 per cent crude ore 37 per cent washed lump 43 per cent fine ore 20 per cent slimes

- 85 mm f 5 m m 0.25-5 mm - 0.25 mm 63.0 per cent Fe 64.5 per cent Fe 66.9 per cent Fe 6.17 per cent Siû, 4.0 per cent SiO, 3.1 per cent SO, 1.03 per cent A&O, 0.92 per cent Alzo, 0.73 per cent Alzo, 0.057 per cent P 0.07 per cent P

2.1 per cent ignition loss 1.G per cent ignition loss 0.048 per cent P

periods of weathering do not show the typical profile. Their areal extension is rather limited, but they go as deep as 600 m below surface. Even so, the blue ores of the Nimba Mountains represent only the deepest, non-eroded roots of larger ore bodies, which, for instance, on Guinean territory are bound to levellings at an altitude of 1,600-1,650 m above sea level. Although the entire weathering section is no longer preserved, the blue high grade ores may be considered as ores of zone B. The following chemical composition is characteristic: 67.8 per cent Fe, 1.5 per cent SO,, 0.5 per cent Al,O,, 1.5 per cent ignition loss.

The blue ores are mostly fine ores, but a secondary hematite mineralization, to which this type of ore owes its colour, sometimes resulted in a cementation (medium hard ores), thus lump ore production of the blue ores after mining and crushing amounts to approximately 10 per cent. There- fore, the blue ores correspond practically to zone B of the itabirites alteration profile.

For a weight recovery of 98 per cent, see Table 6. The slimes are enriched by flotation of the tailings to a concentrate, which is pelletized. The pellets give the following chemical analysis: 63.9 per cent Fe, 5.2 per cent SiO,, 1.99 per cent Alzo3, 0.065 per cent P, 0.76 per cent Cao, 0.40 per cent MgO.

Mano River (Liberia)

The Mano River deposit (Fig. 19) was discovered and geologically investigated at the end of the fifties, and in 1961 it was opened up by the National Iron Ore Company Ltd. Deposit A is already exhausted, whereas deposits H, I, No. 4 and J are mined; the ore bodies E, V, No. 5 and 6, however, have not yet been opened up.

The calculation of ore reserves, as well as mining operations, is based on a cut-off grade of 50 per cent Fe; in addition, so-called lean ores containing 45-50 per cent Fe are eliminated, which are separately mined and stocked. Ore and overburden are in the ratio of 1 : 0.4.

The crude ore is mined by modern open-pit methods, dressed by washing and at present gives the following analysis: 50-55 per cent Fe, 2-5 per cent SiO,, 4-7 per cent Alzo,, 0.02-0.06 per cent P, 10-13 per cent ignition loss.

During the dressing process, which includes crushing, washing and screening, the following qualities of direct- shipping ores are obtained (weight recovery 70-75 per cent):

lump ore 9.5-150 mm (56-59 per cent Fe, 3-4.5 per cent SO,, 6-7.5 per cent AlzoB, 0.05 per cent P and 8 per cent ignition loss); fine ore 0.3-9.5 mm (56-59 per cent Fe, 3.0-4.5 per cent SiO,, 4.5-6.5 per cent Alzo3, 0.05-0.06 per cent P and 7.0-9.0 per cent ignition loss).

The Mano River iron ore deposit consists of a series of metamorphic schists, amphibolites and itabirites (Fig. 20). This series reaches a total thickness of more than 300 m and forms flat, NE.-SW .-striking synclines, which on their footwall are bordered by younger gneisses. The metasedi- nients are marked by an abrupt change of facies, thus a stratigraphical subdivision of this series cannot be set up. According to James (1954), the intercalated itabirite horizons nearly always belong to the silicate facies, integrating and alternating with schists hozirontally as well as vertically over very short distances (Fig. 20).

Because of the intensive weathering, investigations up to the present rarely showed hard, unweathered itabirites

A Hill I Hill

H Hill 5 Hill

O 1000 METRES '-

EXPLANATION FOR MAP AND SECTIONS Vertical exaggerotion ZX

0 High -Grade Ores

EcI u 1 tromaf ic intrusives

Iron formation.schist. and amphibolite

Ea

3 55

I] schists

600 ~

LOO

2 O0

O

O 100 200 m

FIG. 21. Geological map of Mesaboin hill/Marampa, Sierra Leone (Sierra Leone Development Company Ltd).

356


of zone D . There are mostly banded magnetite/Fe silicate rocks (amphibolites).

The overlying zone C consists of a friable, limonite- coloured, clayey rock detritus with an Fe content of 40- 54 per cent, showing the specification of the so-called lean ores.

The ores of zone B towards the hanging wall are also mined. They are marked by intense leaching of the silica and enriched alumina, which caused the clayey composition of the material.

The outcrop ores of zone A consist of a 10 m thick limonite crust with high Al content and is yielding lump ore.

Marampa (Sierra Leone)

Managed by the Sierra Leone Development Company Ltd, the Marampa mine has been in production since 1933 without interruption.

At a cut-off grade of 37 per cent Fe the crude ore shows the following average composition: 40.0 per cent Fe, 32.0 per cent Sioz, 4,5 per cent Alzo3, 0.2 per cent M n .

The crude ore is mined by open-pit operation (ore: overburden = 1 : 0.85) and upgraded by means of Hum- phrey spirals to a high grade concentrate, bringing about a weight recovery of 42 per cent at present, which it is intended to increase to 50 per cent by improved dressing operations. The concentrate shows the following analysis and grain composition: 64.1 per cent Fe, 6.37 per cent Sioz, 0.84 per cent Alzo3, 0.23 per cent Mn, 0.008 per cent P and 0.65 per cent ignition loss; 0.32-3.1 mm = 19.5 per

cent, - 0.32 mm = 20.0 per cent, - 0.25 mm = 22.5 per cent, - 0.18 mm = 18.5 per cent, - 0.125 mm = 14.5 per cent,-O.O9mm = 1.5percentand-0.075mm = 3.5per cent.

In the area of the Marampa deposit (altitude approximately 250 m), we find a series of highly metamorphic hematite quartzites and hematitic mica schists (Marampa schists), which were formed by itabirites of the oxide and silicate facies. The strata are divided into the following horizons, starting from the top: upper hematite-quartzites, approximately 100 m; upper quartz-mica-schists, approximately 40 m; middle hematite-schists, approximately 75 m; middle quartz-mica-schists, approximately 60 m; lower hematite- schists, approximately 40 m; and lower quartz-mica-schists, > 100 m.

The metasediments form a flat, north-south-striking syncline extending to about 500 x 500 m, and with a rolling pitch (Fig. 21); this structure is based on granites and gneisses of the Kasila-series.

When unweathered (zone D) the hematite-quartzites are hard, distinctly slaty rocks, composed of quartz, hematite aiid little biotite. Because of its coarseness (liberation size approximately 0.5 mm), no high grade ores were formed near the surface, but only enriched, soft itabirites, which may be placed into zone C of the typical weathering section. Their Fe content averages 49 per cent. As there is no zone B (weathered high grade ores), the clayey-lateritic crustal ore of zone A is directly placed on it with a thickness of 5-9 m and depending on the aluminium content it may show 50-65 per cent Fe.

While in former years zones A and C were mined, the present reserves originate mostly from zone D.

Résumé

Les minerais de fer d’itabiuite d~i Libésia et du bouclieu guyanais (H. Gruss)

Le Précambrien du Libéria et du bouclier guyanais contient des dépôts de minerai de fer d’itabirite qui, pendant les deux dernières décennies, sont devenus de plus en plus importants particulièrement pour les États-Unis et l’Europe occidentale comme sources de matières premières, avec une production et une exportation qui se sont élevées à 37,8 millions de tonnes en 1968. Les similitudes de la structure des minerais de fer d‘itabirite des deux continents sont dues à leur histoire géologique commune, qui remonte au plus ancien Précambrien et qui a pris un cours analogue même après la séparation des deux continents la période mésozoïque.

Les itabirites des deux boucliers représentent les plus jeunes éléments des strates géosynclinales des métasédi- ments et des vulcanites, dont le substratum est ou bien connu ou ne peut être identifié. Ces roches furent plissées par des mouvements orogéniques, il y a 2,5 ou 3 milliards d’années et ont subi des altérations métainorphiques régio-

nales. Une autre métamorphose s’est produite il y a 1,8 à 2 milliards d’années avec des intrusions de gneiss et des intrusions acides ou alcalines. Le géosynclinal précambrien formé de cette façon a une structure symétrique avec UR noyau métamorphique mésozonal ou catazonal et des bordures métamorphosées épizonalement, la direction des plissements allant toujours du centre vers l’extérieur. En conséquence, les dépôts d’itabirite du centre sont caracté- risés par un haut degré de métamorphose, par la seule présence de plis aplatis, un grain grossier et, en partie, une différenciation métamorphique du minerai à haute teneur (Bomi Hill, El Pao), tandis que les dépôts d’itabirite périphériques sont, en général, faiblement métamorphiques, aux plis fortement redressés et à grains fins.

Après le plissement et le surhaussement, l’orogénie précambrienne a été nivelée à l’état de pénéplaine. Comme résultat de la dislocation du continent du Gondwana pendant le Crétacé supérieur, le synclinal s’est fendu le long de sa crête plongeante nord-ouest/sud-est et s’est séparé pour former les boucliers actuels de Libéria et de Guyane.

357

H. Gruss

L'érosion profonde et la désagrégation qui ont commencé à se produire après cette séparation ont conduit, sur les deux continents, à la formation d'itabirite enrichie (Ma- rampa, Bong Range) et des dépôts très importants de minerais désintégrés à haute teneur qui forment aujourd'hui la base des exportations de minerais de fer des pays concer- nés (Nimba, Mano, Cerro Bolivar, San Isidro).

Les itabirites des boucliers du Libéria et de la Guyane présentent en général des épaisseurs sédimentaires qui n'excèdent pas quelques mètres ; cependant dans les bassins de sédimentation qui ont évolué plus tard en synclinoriums, les épaisseurs ont augmenté pour atteindre 100 et même 250 mètres. I1 a été démontré que parfois les itabirites se sont déposées sur le bord des bassins et latéralement se sont entremêlées avec des sédiments clastiques à grains grossiers, A l'occasion, la sédimentation a eu lieu au cours de nombreux cycles, chacun commençant avec un faciès de silicate qui se transforme en un faciès de magnétite et se termine par un faciès d'oxyde d'hématite.

Dans les zones catamétamorphiques, il s'est produit une différenciation métamorphique des itabirites du faciès d'oxyde accompagnée essentiellement de la formation de magnétite et de minerais grumeleux à haute teneur (67 % de fer), qui n'existent pas dans les parties mésozonales et épi- zonales du synclinal.

Au contraire, les minerais à haute teneur désagrégés et décomposés se sont développés aux époques fossiles ou récentes dans un climat tropical humide, présentant une section verticale typique qui dépend de la structure de la roche

Zone A

originale (silicate ou faciès d'oxyde).

Faciès-silicate Faciès-oxyde Croûte de lamelies Croûte de lamelles d'hématite cimentées d'hématite cimentées de limonite, haute de limonite- concentration d'argile, hydro-hématite

B

C

D

Dans

partiellement compacte, 50 %Fe Mélange meuble et argileux de lamelles d'hématite et de croûtes de limonite, brun, 55 %Fe Argileux, désintégration de l'itabirite, brun, 40-50 % Fe Itabirite silicatée, dure, contenant du mica et de l'amphibole, avec magnétite, 35 % Fe

compacte, mais poreuse, 62-67 %Fe Mélange meuble de lamelles d'hématite martite, brun, noir et même bleu, 63-69 % Fe

Meuble, désintégration de l'itabirite en quartzite (haute teneur), 40-50 % Fe Itabirite oxydée dure, avec hématite et magnétite, 35-40 % Fe

les itabirites catamorphiques, c'est-à-dire à wains - grossiers, la désagrégation, en raison de la surface limitée du grain par unité de volume, est étendue et pénètre profon- dément ; il n'en résulte pas la formation de minerai à haute teneur, mais plutôt la formation d'itabirites molles qui peuvent aisément être concentrées (Maramba, Bong Range). D'autre part, dans la zone épizonale métamorphique, la désagrégation, en raison de la finesse du grain, a été très profonde, le résultat étant la formation des minerais désagrégés à haute teneur (Nimba, Mano, Cerro Bolivar, San Isidro).

Ces dépôts se rencontrent en liaison avec la péné- plaine crétacée (Gondwana) et sur des plans d'érosion plus récents. Le minerai désagrégé à haute teneur des péné- plaines plus anciennes et situées à un niveau plus élevé, pénètre plus profondément (à plus de 500 mètres) que celle des plans d'érosion plus jeunes (de 50 à 200 mètres). Tandis que les minerais à haute teneur de cette dernière sont de couleur brune et noire, les minerais à haute teneur des plus anciennes pénéplaines sont caractérisées par une couleur bleue et une pénétration zonale radiculaire, cimentée par l'hématite qui, de l'avis de l'auteur, est supergène.


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p. 1089-98.

vol. III, p. 16-27.

3 59

Structural control of the localization of rich iron ores of Krivoyrog

G. V. Tokhtuev Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrainian S.S.R.

FIG. 1. Distribution of Krivoyrog ore fields. 1. Illych mine; 2. Dzerzhinsky mine; 3. Kirov mine; 4. Karl Liebknecht mine; 5. Komintern mine; 6. Frunze mine; 7. XX Party Congress mine; 8. Rosa Luxembourg mine; 9. Lenin mine.

The Krivoyrog basin, one of the largest centres of iron ore mining, has been intensely studied and prospected.

The general structure of the basin consists of a group of conjugated second order folds which form a synclinorium 7 km wide and 70 km long. These folds are found in the following succession (from east to west): Saksagan syncline and anticline, Main Krivoyrog syncline and anticline, Main Krivoyrog syncline, Taranakholihman anticline, and Lih- man (Iiiguletz) syncline. Each of the above folds forms the structural foundation of a separate ore field, which differs from the others not only in its structure, but also in the type of rocks, degree of metamorphism, ore type, and in its degree of weathering.

Saksagan ore field This field (Fig. 1) is very important because of its scale of ore mineralization. The field is located within the limits of two conjugate second order folds: the Saksagan syncline and Saksagan anticline. Both of these folds are complicated by a longitudinal thrust fault and by a series of smaller faults. The larger ore bodies are localized in the Saksagan syncline (90 per cent of the ore deposit). The Saksagan anticline is not as important.

The Saksagan ore field is characterized by low grade metamorphic rocks of the green schist facies, and by a high intensity of oxidation that extends as deep as 2.5 km.

Structural control of mineralization is indicated for this ore field by the observation that the rich ores are concentrated in the trough of the Saksagan syncline and are associated with deformation of its east limb. Small deposits in the Saksagan anticline are controlled by longitudinal dislocations and loop-shaped foldings of layers.

STRUCTURAL TYPES OF DEPOSITS IN THE SAKSAGAN ORE FIELD

The Saksagan ore field consists of eight separate deposits (Fig. 2). Each deposit has a particular structure that

Unesco, 1973. Genesis of Precambrian iron and mangunese deposits. Proc. Kiev Symp., 1970. (Earth sciences, 9.) 361

G. V. Tokhtuev

Vertical longitudinal section

Saksagan syncline

FIG. 2. Structural types of Saksagan ore field deposits. 1. deposits associated with the Saksagan trough bend; 2. deposits related to knots of compression and shear deformations on the Saksagan syncline limbs, connected with flexure-type mineral-

controls ore mineralization. The Saksagan field deposits can be divided into three structural types: 1. Deposits associated with the trough of the Saksagan syn-

cline. For example, a section of the Communav mine at the south end of the Dzerzhinsky mining district, where the trough of the Saksagan syncline intercepts the surface.

2. Deposits associated with compressional knots and cross- cutting folded fractures joining at depth with ore mineralization of the trough type. They are illustrated in sections of the Gigant mine, the Dzerzhinsky Sak- sagan mine, the Kirov mine, and the Karl Liebknecht mine. Trough mineralization of the Saksagan syncline plunges north and forms ore at depths ranging from 300 to 2,000 m. Steeply plunging ore chimneys, belonging to the cross-cutting zones of pressure and deformation, extend from the top of the ore. The Komintern deposit belongs to this structure type.

3. Deposits associated with cross-cutting compressive zones and the development of folded-faulted deformations on the limb of the Saksagan syncline. To this type belong the deposits of the northern part, of the Saksagan ore field where the depth of the trough of the Saksagan syncline is 3-4 km. It is not yet known whether ore concentration in the trough extends farther to the north. The following mines are encountered: the Frunze, Twentieth Congress of the CPSU, Rosa Luxembourg and V. I. Lenin mines.

Cross-cutting zones of deformation on the east limb of the Saksagan syncline, which make up the basic structure of almost aíl deposits of the Saksagan ore field, are regularly distributed along the syncline strike, accompanied by small folds, breaks, depots, cleavages, boudinage, etc. In these zones, 1 or 2 km wide, numerous ore bodies are located associated with high order structures. The distance between the cross-cutting zones ranges from 2 to 3 km. Compression of the ferruginous strata of the fifth and sixth ferruginous horizons characteristically caused a decrease in thickness in the limits of the cross-cutting zones of shearing as compared to the thickness of the horizons in non-ore

ization at depth; 3. deposits confined to transverse zones of compression and development of fold-fault deformations on the Saksagan syncline limbs.

locations between mines. Especially, considerable decrease in the thickness of the ferruginous horizons is observed in places of intense metallization where boudinage structures were formed. The process of formation of boudinage structures consisted of the removal of quartz from the compressive zone, because quartz became unstable as the result of high stress, and was easily dissolved and removed by metamorphic solutions. Ore minerals were not mobile under such conditions, and they accumulated in the interboudine pinches, leading to the formation of ore bodies. The increase of quartz solubility under high pressure has been proved experimentally by Syromjatnikov.

Thus, in the formation of cross-cutting zones of shearing the development of boudinage along with the development of folding and fracturing played an important role. It should also be mentioned that cross-cutting zones of shearing of some ferruginous horizons resulted in the development of parallel beds. For example, in the Komintern mine, ore bodies are located in the ñrst, second, fifth and sixth ferruginous horizons. In some mines parallel beds are found in the fifth and sixth horizons, and in the Rosa Luxembourg, and V. I. Lenin mines there are five to six or more parailel chains of beds in the fifth and sixth ferruginous horizons. The beds in various strata are located strictly along the cross-cutting zone of shearing.

STRUCTURAL TYPES O F ORE BODIES IN THE SAKSAGAN ORE FIELD

Ore bodies are typically controlled by various high order folds or fractures, either belonging to the complex of the structure of the deposit, or being independent forms with their own characteristic orientation. The structures controlling ore bodies are different; they can be folds and flexures of various types and orders, zones of microfolding, various tectonic cross-cutting and diagonal ruptures, zones of interboudine pinches, zones of intensive development of jointing and cleavage, zones of breccia, etc.

362

Structural control of the localization of rich iron ores of Krivoyrog

The following are the structural types of ore bodies in

1. Ore bodies belonging to the trough of the Saksagan

2. Ore bodies belonging to arched-up high order folds. 3. Ore bodies controlled by cross-cutting small open folds.

T o this structural type belong both large and small ore bodies, where the relationship of ore mineralization and small folding is so close that the localization of the ore body is easily determined by the orientation of the folds.

4. Ore bodies controlled by the zone of isoclinal folds not undergoing the process of mineralization. The deposits of this structural type are located very close to and parallel with the zone of isoclinal folds, and plunge in the direction and angle of the hinges of the folds, which do not undergo the process of mineralization.

5. Ore bodies belonging to small flexural folds along the strike.

6. Ore bodies belonging to interboudine pinches in zones of macroboudinage development. This structural type is widely distributed not only in the Krivoyrog basin, but also in the Kremenchug and Belozersk areas as well. Interboudine ore bodies are controlled by the orientation of interboudine pinches, which either correspond to the direction of pods of included rocks, or to a diagonal direction which plunges south or north.

7. Ore bodies in zones or cross-cutting faults of high orders.

8. Ore bodies belonging to zones of longitudinal thrust faults.

9. Ore bodies in zones of intraformational and inter- formational breccias. Evidently, these zones are formed as the result of the release of points of tectonic strain where the plastic limit of the rocks was exceeded. These zones undergo the process of mineralization and make up small ore bodies of irregular wasted form.

10. Ore bodies in zones of thickening of cross-cutting shear joints. On the flanks of the Saksagan syncline is a system of closely spaced cross-cutting joints. The spacing of joints ranges from 10 to 20 cm and up to 50 cm. But in some intervals of 50-100 m occur zones of very closely spaced joints where the distances between joint planes are not more than 1-2 cm with simultaneous intensive development of cleavage. Such places are occasionally mineralized, making up small ore bodies controlled by the direction of the joint zones.

11. Ore bodies in zones of intensive development of two cross-cutting shear joints. This structural type is characterized by ore mineralization associated with the intersection of joints which control the position of the ore body in space.

The above typical structural types of ore bodies in the Saksagan ore field are the most important. However, there are other types, in which the relationship of ore mineralization to the structures is less evident.

the Saksagan ore field:

syncline.

Northern ore field The Northern ore field is situated along the continuation of the Saksagan ore field directly to its north. It begins with a large flexural bend of the east flank of the syncline, beyond which the extension of the structures changes from NNE. to N N W . The Northern ore field is also characterized by a higher degree of rock metamorphism than the Saksagan field, by the extensive development of the process of metasomatism, and by a magnetite-type of rich iron ore.

Structural types of deposits and beds of the Northern ore field are the following: 1. Deposits and beds belonging to faulted surface structures. 2. Deposits and beds associated with steeply-dipping syn-

clinal folds and flexures.

Central Krivoyrog ore field

The Central Krivoyrog ore field joins on its west to the southern part of the Saksagan ore field. The major structure of this field is the Main Krivoyrog syncline which makes up the central part of Krivoyrog synclinorium. It is complicated by complex folding of higher orders.

Ore bodies occur along the complexly folded contacts of the rocks of the middle ore suite and cover the upper schist suite of the Krivoyrog series. Characteristic features of the Central Krivoyrog ore field are complexly folded structures complicated by disturbances of various types and orders, a low degree of metamorphism (green schist facies) and the development of chlorite-magnetite and carbonate-magnetite ores which have been transformed in the zone of oxidation into martite ores up to a depth of 150- 200 m.

Deposits of the Central Krivoyrog ore field are controlled by one type of complicated steep folded structure of the third, fourth and higher orders in the zone of the contact of the iron ore suite with the overlying layers of the upper schist suite.

Tarapako-Lihman ore field

The Tarapako-Lihman ore field is situated to the west of the Central Krivoyrog ore field. It belongs to a large structure of the second order, the Tarapako-Lihman anticline, which forms the west flank of the Krivoyrog synclinorium. Ore zones are located on the contact between ferruginous beds and the upper suite, developing mainly on the flanks of the fold. They are complicated by numerous fractures and, more rarely, are distributed up to the crest of the anticline.

This field is characterized by a high degree of metamorphism (zone of garnet-cummingtonite schists), by the development of magnetite-rich ores, by a shallow depth of the zone of oxidation (60-100 m), and by small thicknesses (3-10 m) of bedding-plane ore bodies.

363

G. V. Tokhtuev

Deposits of the Tarapako-Lihman ore field belong to the following structural types: 1. Deposits and ore bodies on the flank of the anticline,

which have been complicated by post-ore cross-cutting and longitudinal faults.

2. Deposits and ore bodies associated with arched-up folds of the Tarapako-Lihman anticline and folds of higher orders.

Inguletz (Lihman) ore field

The Inguletz (Lihman) ore field is located along the western border of the Krivoyrog basin. The structural base for this ore field is the Lihman syncline, conjugated in its northern part with the Tarapako-Lihman anticline. It extends 30 km south of the other submeridional structures of the Krivoyrog synclinorium. The western limb of the Lihman synclines is in the main part and

is cut off by the Western Thrust. In the southern extremity of the fold, a part of the western limb is preserved in the limits of the Inguletz mine.

In ternis of the types of ores and their stratigraphic relationships, the Inguletz ore field is analogous to the Tarapako-Lihman and Central Krivoyrog ore fields, However, the southern extremity of the ore field (Inguletz mine) is characterized by a thick trough mineralization in the Lihman syncline and a highly developed zone of oxidation. According to these characteristics, it approaches the Saksagan type of trough mineralization. The ore controlling structures of the Inguletz ore field produce two structural types of ore deposits and beds. 1. Deposits and ore bodies in the trough of the syncline

(Inguletz mine). 2. Deposits and ore bodies in zones of shearing on the

eastern limb of the syncline (Pahmanovsky mine and small exhausted deposits north of the Inguletz River).

Résumé

Détermination structurale de la localisation des minerais de fer à haute teneur de Krivoyrog (G. V. Tokhtuev)

1. Les relations structurales qui conduisent à la localisation des minerais de fer de Krivoyrog sont déterminées par des études de structures de différents ordres : minérali- sation, screening, minerais.

2. La région de faciès structural de Krivorozhsky- Kremenchugsky a une structure qui est définie par un synclinorium composé de roches d’une formation de fer siliceux. Elle est compliquée par une fracture longitudinale et couvre une étendue de 400 à 500 km. Ici le contrôle structural a été utilisé pour planifier et exécuter un levé géophysique à petite échelle au sol et aéroporté.

3. La région de minerai de fer de Krivorozhsky (bassin) fait partie de la zone de faciès structural de Kri- vorozhsky-Kremenchugsky. Sa structure consiste en un groupe de larges plis conjugués qui s’étend sur 70 km dans la direction du gisement. Les minerais de fer sont ici déplacés. Les éléments structuraux servent alors à la prospection et à l’étude des perspectives.

4. Les gisements de minerai de fer qui composent le bassin de Krivorozhsky sont définis par de larges plis séparés du troisième ordre compliqués par des dislocations longitudinales et une rupture de continuité (synclinaux de Saksagansky, de Likhmanovsky, Krivorozhsky et anticlinal de Tarapaco-Likhmanovsky). La longueur du bassin de minerai est déterminée par les dix premiers kilomètres. Les facteurs structuraux sont utilisés ici pour une exploration préliminaire.

5. Les dépôts de minerai dans chacun des bassins de

Krivoyrog sont liés à des structures complexes, flancs de raccordements de quatrième ordre, courbes de larges plis. En général, c’est aux nœuds des plis transversaux, flexures et fractures, structures de microboudinage, etc., que la minéralisation est liée. La dimension des dépôts de minerai est mesurée par les premiers kilomètres. Le contrôle structural est utilisé comme base pour la prospection détaillée des différentes mines (dépôts).

6. Les gisements de minerai des différents dépôts sont liés à des structures à minerais des ordres les plus élevés (différents types de plissement, structure de boudinage, zones de jointement intensif et de clivage, zones de brèche et de cataclase, dislocations avec des ruptures de conti- nuité, différents types de déplacements, etc.). La relation structurale de localisation et de minéralisation apparaît ici d’une façon particulièrement claire et précise et déter- mine la morphologie et la localisation des gisements. Le contrôle structural a été effectué en vue d’une prospection détaillée et opérationnelle.

7. La morphologie des gisements de Krivoyrog est extrêmement différente. Elle dépend du type, de la forme et des dimensions des structures à minerai. Les grandes colonnes de minerai prédominant. Leur section varie ; elles pénètrent jusqu’à plus de 2 km. On trouve aussi des dépôts du type à large strate et des dépôts articulés complexes (limités aux coudes des grands plis). Des gisements de moyenne et grande taille prennent la forme de lentilles, de masses à configuration extrêmement irrégulière et de poches dans lesquelles il n’est pas toujours possible de prédire les structures à minerais qui servirant de base pour déterminer les formations de minerai.

3 64

Iron deposits of Michigan (United States of America)'

J. E. Gair U.S. Geological Survey, Washington D.C. 20242 (United States of America)

The iron deposits of Michigan are principally in four areas, the Gogebic, Iron River-Crystal Falls, Marquette and Menominee districts, all in the northern peninsula of Michigan (Fig. 1). Iron-bearing beds form part of a sequence of middle Precambrian metasedimentary rocks, perhaps 1,900-2,500 m.y. old. The metasedimentary sequence (Fig. 2) is generally considered to have been deposited in a marine environment.

Geologic structure in the Gogebic and Menominee districts is essentially ' one-sided'-monoclines or the drag- folded flanks of large regional uplifts; the Marquette district is in a narrow synclinorium and the Iron River- Crystal Falls district occupies a broad three-cornered structural basin. The trend of synclinal and basin axes and of monoclinal iron-formation is generally eastward. The Marquette and Menominee depositional basins of Pre- cambrian time probably were elongate, with long axes being roughly equivalent to the present tectonic axes. Clastic sediments in the iron-formation along the south side of the Marquette synclinorium indicate that the south side was closer to a shoreline than the north side. Little can be determined about the outlines of the Gogebic and Iron River-Crystal Falls depositional basins in Pre- cambrian time.

The iron-formation of the Gogebic, Marquette and Menominee districts is thought to be correlative, and is in the middle part of the middle Precambrian sequence. The iron-formation in the Iron River-Crystal Falls district is younger, in the upper part of that sequence. The correlative rock formations of the first three districts are correlated principally because of similarity of the rock sequence containing the iron-formation (Fig. 2). Iron-formation facies may change along strike in a given district and detailed iron-formation stratigraphy is markedly different in the three districts; whether the iron-formation or any of the associated rock units ever were entirely continuous from one district to the other is unknown. Basement rock for the middle Precambrian rock sequence in the Gogebic, Marquette and Menominee districts is gneissic and/or intrusive granite, amphibolite and/or volcanic greenstone,

all of lower Precambrian age, 2,600 m.y. old or more. The stratigraphic sequence containing the iron-formation in the Iron River-Crystal Falls district is underlain by middle Precambrian volcanic greenstone (Fig. 2).

In the Gogebic district, the thickness of the iron- formation is between 600 and 1,000 ft (180-300 m), and is 800-900 ft (244-274 m) in most places. In the Marquette district the thickness ranges from 450 to 3,500 ft (135- 1,060 m) or more, and commonly is about 1,000 ft (300 m). In the Menominee district, thicknesses range from 300 to 600 ft (91-180 m) and average about 450 ft (135 m). In the Iron River-Crystal Falls district, thicknesses range from 150 to 600 ft (46-180 m) at the west end of the district and from 500 to 800 ft (150-130 m) at the east end.

The iron-formations nearly everywhere have been recrystallized during regional metamorphism, and minerals possibly of diagenetic or low-grade metamorphic origin generally cannot be distinguished from recrystallized primary minerals.

Dominant primary minerals in the Gogebic and Marquette districts are siderite-chert and, locally in the stratigraphic section, hematite-chert (Fig. 3). Most magnetite is probably primary or diagenetic. The primary nature of chert and siderite is indicated by the widespread uniformity of beds, compositions and textures, which are not consistent with a replacement origin. Also, stylolites and preconsolidation slump structures that invqlve chert and siderite indicate that these minerals are primary or very early (Figs. 4 and 5). The draping of ferruginous laminae over chert beds and slump fragments is another indication of the presence of chert early in the history of the iron-formation (Fig, 6). Hematite is deduced to be primary where it occurs in oolites and granules and, in granules of probably organic origin, in relatively thick wavy or pod-shaped layers. Oolites or granules, particularly where they form lenticular beds, are interpreted as deposits that originated in shallow agitated water; the expected primary iron mineral is a ferric oxide. In places,

1. Publication authorized by the Director of the U.S. Geological Survey.

Unesco, 1973. Genesis ofPrecambrian iron and manganese deposits. Proc. Kiev Symp,, 1970. (Earth sciences, 9.) 365

J. E. Gair

47

46

90 o 89' 88" I I I

FIG. 1. Geologic sketch map of western part of northern peninsula of Michigan (United States), showing location of major iron-producing districts.

unoxidized interbeds of siderite, magnetite, iron silicate or greywacke in hematitic iron-formation indicate a lack of oxidation since deposition of the rock and virtually prove the primary nature of the adjacent hematite.

Magnetite is important in all districts except Iron River-Crystal Falls and may be primary, diagenetic or metamorphic. A primary or diagenetic origin is deduced for large amounts that are widely distributed in thin uniform laminations in iron-formation of low metamorphic grade. Uniformly alternating thin layers of magnetite and siderite or magnetite and hematitic chert are more readily explained by fluctuations in conditions during sedimentation than by post-depositional processes, but generally it has not been possible to distinguish primary from diagenetic magnetite by direct evidence. Evidence for the

diagenetic origin of magnetite by the reduction of ferric oxide has been shown by several workers in other regions. In the eastern part of the Marquette district, some magnetite, possibly a large amount, has formed diagenetically (Han, 1962), mainly by the oxidation or decarbonation of siderite. Small amounts also have formed by the oxidation of iron silicate. At the Empire taconite mine, small relict 'islands' of siderite iron-formation occur sporadically within a unit of magnetite-rich iron-formation for a strike distance of about f mile (about 850 m) and through a thickness of 300400 ft (about 125 m). Commonly, bedding is continuous from sideritic relicts into the magnetite- rich rock (Fig. 7). In a few places, marginal concentrations of magnetite occur in granules that consist dominantly of carbonate, iron silicate or chert (Fig. 8). The magnetite

366

MAROUETTE \

EX P LA N AT I O N pz4

Greywacke

Greenstone

pJ .. . ..._. :; :... Quartzite

0 Iron-Formation

IRO-N RIVER-CRYSTAL FALLS DISTRICT

Dolomite

Gneiss, granite

FEET

Columns broken where part of stratigraphic section omitted

FIG. 2. Correlation of major lithologies in Michigan iron-producing districts.

3 U7

J. E. Gair

-500 FEET

O

IRON RIVER-CRYSTAL FALLS DISTRICT

- I

N C

H

E S

-

-

-

Sid (Stilp)

IRON-FORMATION I I MARQUETTE DISTRICT

_/A.

GOGEBIC DISTRICT /-A--- /

NEGAUNEE IRON-'

FORMATION

IRON- FORMATION

Sid (H-Mt)

Cid (H-Mt) ________-- -.

2 rnrn O-

- H-Mt

\ Cid-Sil-Mt

Cid-Sil

H-Hematite Mt-Magnetite Py-Pyrite Sid-Siderite Sil-Silicate

Stilp -Stilpnomelane

Subordinate minerals shown in parenthesis

\ \ \ \ \ \ \ \ \,MENOMINEE DISTRICT

---_-__

IRON- FORMATION

FIG. 3, Primary-diagenetic iron minerals in Michigan iron- producing districts.

368

Iron deposits of Michigan (United States of America)

FIG. 5. Photograph, drill core; slump structure in cherty and carbonate layers.

1

, . J

FIG. 6. Photograph, drill core; laminae rich in silicate and magnetite draped over chert-rich fragment.

FIG. 7. Photograph, polished surfaces; replacement of sideritic layers by magnetite; relict sideritic ‘islands’ commonly bordered by reaction rim of secondary carbonate.

FIG. 8, Photomicrograph; marginal to complete replacement of minnesotaite granules by magnetite. Plane light.

3 69

J. E. Gair

is attributed to oxidation of the carbonate or silicate or to replacement of the chert during diagenesis. Evidently, magnetite replaced iron carbonate or iron silicate in newly deposited sediment in response to a change in the original neutral or reducing conditions to moderately oxidizing conditions, or possibly because of a change in p H from near neutral to alkaline (see results of experimental studies dealing with influence of Eh and p H on the deposition of iron minerals, reported by Garrels, 1960; Huber, 1958; and Krauskopf, 1957). Shallowing of the sea bottom could have increased the oxygen content of sea-water and adjacent interstitial water in bottom sediments, or by improving near-bottom circulation, may have lowered the acidity of sea-water. The experimental work cited above shows that siderite-stable conditions can change to magnetite-stable conditions by an increase in pH, with no change in Eh, or even with a decrease in Eh, although the actual geologic conditions that could produce a simultaneous increase in p H and drop in Eh are difficult to visualize.

Iron silicates, particularly minnesotaite and stilpnomelane, are abundant in parts of the Gogebic and Mar- quette districts, and stilpnomelane and iron chlorite are locally abundant in the Iron River-Crystal Falls district, but the absence of these minerals in unmetamorphosed iron-formation or in post-Precambrian ironstone of other regions suggests that they are not primary or diagenetic, but of low-grade metamorphic origin. On the basis of chemical composition (Deer, Howie and Zussmann, 1963, Winchell, 1951), minnesotaite probably does not require a silicate parent and may have been derived solely by diagenetic or metamorphic reactions between primary chert and siderite. The significant aluminium content of stilpnomelane, on the other hand, indicates a substantial increment $of aluminous silicate in the primary sediment from which that mineral was derived. The widespread lack of a siderite-chert reaction at low metamorphic grade has been cited as evidence that both minnesotaite and stilpnomelane developed from primary silicate material (James, 1954).

TABLE 1. Modes of typical silicate iron-formation, eastern part of Marquette district (in volume per cent)

1 2 3

Chert Siderite Magnetite S tilpn omelane Minnesotaite Mixed magnetite and

Gruneritel Secondary hematite

iron silicate

Trace 6.2 4.0

1 .o 2.5 22.0 90.0

18.4 49.5 3.5

1.2 0.5

54.4 20.9 20.6 5.1

1, Attributed to contact metamorphism by intrusion of mafic sill.

FIG. 9. Photomicrograph; granules of minnesotaite and minnesotaite-magnetite (granules marked M) surrounded by chert-rich matrix. Note marginal concentrations of secondary magnetite. Cross nicols.

In the eastern part of the Marquette district, some thinly laminated iron-formation rich in iron silicate has a low chert content, less than 10 per cent (Table 1); other varieties of thinly laminated iron-formation typically contain 15-50 per cent chert. The silica content of cherty iron-formation varies widely depending mainly on the amount of chert. Pure chert-siderite iron-formation, having about 61 per cent chert, contains about 45 per cent silica by weight, comparable with the percentage of silica in some of the silicate-rich, chert-poor iron-formation. The silicate iron-formation therefore differs from other thinly laminated iron-formation in the vicinity mainly in lacking chert laminae. This may be a result of the incorporation of original chert into iron silicate minerals formed after sedimentation.

Some layers of iron-formation consist largely of iron silicate granules. Gr anules and matrix commonly are similar in composition, as would be expected if granules formed by agitation of the original sediment. However, in some layers, iron silicate granules are surrounded by silicate-poor material, generally chert or siderite (Fig. 9), or silicate-poor granules may be surrounded by silicate. Such silicate granules or matrix seem to be best explained by selective replacement of cherty or sideritic material during diagenesis or by concretionary growth during diagenesis. A chert matrix for closely packed granules of silicate can be explained by infilling by silica, but this explanation does not seem adequate for widely scattered silicate granules in chert. An alternative explanation, that granules differentiated during sedimentation or agitation of bottom sediments, seems unlikely without an accompanying segregation of the minerals into layers. Such differentiated granules, therefore,

370


FIG. 10. Photomicrograph; riebeckite in chert-magnetite-carbonate iron-formation.

seem to provide clear evidence of the diagenetic or metamorphic growth of iron silicate.

Primary iron minerals in the Iron River-Crystal Falls district are principally siderite and pyrite. The association of bedded pyrite both with siderite and carbon-rich sediment is a strong indication that it originated as a primary sediment. Stilpnomelane is common in the east part of the district but, as in the Gogebic and Marquette districts, is considered to be of low-grade metamorphic origin. In the Menominee district, hematite and possibly magnetite, were important primary minerals.

Riebeckite and aegirinaugite are present in thin zones in the iron-formation in the eastern part of the Marquette district through a stratigraphic interval of 300-400 ft (about 125 m) and a distance along strike of about 2 miles (3 km) and down dip for at least mile (about 850 m) (Fig. 10). The soda content of such iron-formation ranges from 0.5 per cent to 6 per cent. Some of the riebeckite-bearing iron-formation contains, or is associated with, clastic sediment. I interpret the sedimentation of the riebeckite-bearing iron-formation as having taken place locally in shallow water under evaporite conditions. Soda-bearing parts .of the Wabush Iron Formation of Labrador are also considered to have originated in an environment both of high Eh and high salinity (Klein, 1966).

In the Iron River-Crystal Falls, Marquette and Meno- minee districts there is little or no evidence of contemporaneous volcanism in the iron-formation or in conformable rock below. Significant volcanism is known to have occurred during iron-formation deposition only in the eastern part of the Gogebic, district, but even there only for a limited part of the entire period of ferruginous sedimentation.

FIG. 11. Photomicrograph; jaspilite; thin finer-grained layer is of jasper (hematitic chert); coarser-grained layers are of martite and chert. Note flattened granules in jasper layer and clear chert granules in thicker martite layer.

Middle Precambrian sedimentation was brought to a close, or was followed closely by, the regionwide Penokean orogeny about 1,900 m. y. ago. Metamorphism caused recrystallization of iron-formation almost everywhere in the area. Although primary or diagenetic chert, hematite and magnetite commonly persist from lowest grades of metamorphism to the sillimanite grade, they are recrystallized to increasingly larger grains at higher grades of metamorphism. Jaspilite is a recrystallized variety of hematite-chert iron- formation of low to moderate metamorphic grade. Much jaspilite in the Marquette district, however, contains a large percentage of magnetite or martite; commonly the iron- rich laminae are mainly magnetite-martite, and the jasper laminae are hematitic chert with minor magnetite (Fig. 11). The sizes of chert and hematite grains, in particular, correspond closely to metamorphic grade. Siderite, on the other hand, recrystallized early at low metamorphic grade but, at higher metamorphic grades, reacted with chert to form grunerite, and some possibly was altered to magnetite.

The characteristic siderite in three of the four Michigan districts suggests deposition in basins isolated from the circulation of the open sea. The normal oxidizing condition of open seas was eliminated (changed to negative Eh or to higher than normal acidity) by stagnation, except possibly in the shallower parts of depositional basins, such as near shore, or more widely after uplift or infilling of basins of sedimentation. In shallow areas, fully oxidizing conditions and normal slightly alkaline conditions persisted or re- curred, marked by positive Eh and the deposition of hematite or ferric hydroxide-like precipitates, Magnetite may have precipitated in such an environment under marginal oxidizing conditions, at Eh near or a little less than zero. The

371

J. E. Gair

abrupt small-scale alternation of layers rich in hematite and magnetite, in jaspilite, suggests relatively rapid local fluctuations in oxygen activity during sedimentation, a condition most likely to be realized in shallow water near the base of wave action. In the eastern part of the Marquette district, however, magnetite that formed by the diagenetic replacement of siderite indicates a change toward positive Eh or greater alkalinity after sedimentation, as cited already. In the Iron River-Crystal Falls district, a combination of stagnation in a depositional basin and sufficiently deep water locally to prevent ‘ contamination’ by atmospheric oxygen permitted organic carbon to accumulate with chemically precipitated iron sulphide.

Between the time of deposition and diagenesis of the iron-formations and the time of their regional metamorphism at the close of the middle Precambrian, they were partly weathered and eroded in places, shortly after they were deposited, and intruded by mafic igneous rock. In the Marquette district, at least, mafic intrusions took place both before and after the middle Precambrian episode of weathering and erosion. Secondary iron oxide produced from siderite and magnetite during that episode in the Marquette district, and possibly elsewhere, was recrystallized during the Penokean metamorphism. This iron oxide in its recrystallized form is generally indistinguishable from recrystallized primary hematite. Thersfore, the principal recrystallized hematitic rock, jaspilite, may be either primary hematitic rock, or may have been derived by secondary oxidation prior to regional metamorphism. Criteria for distinguishing these two types are few and obscure. The widespread spatial association of jaspilite with the erosion

FIG. 12. Photomicrograph; retrograded porphyroblasts of grunerite in minnesotaite-magnetite-carbonate iron-formation. Pseudomorphs of grunerite are mainly of quartz, plus minor siderite and magnetite.

surface cutting into the iron-formation to different stratigraphic levels in different parts of the Marquette district is the strongest evidence for the derivation of jaspilite by premetamorphic weathering. The rare gradation of jaspilite into small, apparently relict ‘islands’ of sideritic iron- formation also is evidence for the secondary origin of jaspilite. The principal direct evidence of jaspilite of primary origin is the presence in places of hematitic oolites or granules. Jaspilite that appears to grade into sideritic iron- formation is thinly laminated, with most layers less than 0.5 inch thick (about 1.25 cm), whereas layers of granular jaspilite commonly are 1-3 inches thick (about 2.5-7.5 cm) and are pod-shaped.

In the eastern part of the Marquette district, the mafic intrusions prior to Penokean metamorphism locally modified the iron-formation by converting some siderite to magnetite or pyrite near dyke contacts, and by forming grunerite porphyroblasts in siderite-chert layers or in layers that are now mainly minnesotaite. During ensuing low-grade regional metamorphism in the area, the grunerite porphyroblasts commonly were altered retrogressively and replaced by quartz, quartz-siderite, or quartz-siderite-magnetite (Fig. 12).

Ore-grade concentrations of hematite that accumulated during the middle Precambrian weathering episode recrystallized during Penokean metamorphism to form what is called hard ore. Because the hard ore deposits typically average several per cent richer in iron than ore-grade concentrations of ferric oxide (soft ore) that developed after metamorphic recrystallization, it is likely that secondary oxidation alone cannot explain the concentration and that hydrothermal solutions aided the concentration of hard ore. Hydrothermal solutions alone have been invoked as the concentrating agent by some workers, but there is virtually no evidence for such solutions in iron-formation underlying hard ore bodies. The only known igneous source of such solutions (the mafic intrusive bodies) cannot explain the localization of many of the most important orebodies at the middle Precambrian erosion suiface. Small amounts of autogenous hydrothermal solutions might have been derived from the heating of connate water and the dehydration of chert during regional metamorphism and supplemented the concentration of hard ore.

After the Penokean metamorphism the recrystallized iron-formation evidently remained largely unchanged for hundreds of millions of years until the Keweenawan-early Palaeozoic interval about 600 to 900 m.y. ago. Then, parts of the iron-formation were oxidized and leached (weathered) by supergene or mixed supergene and hypogene ground-water solutions (James et al., 1968), producing sporadic concentrations of earthy hematite aiid goethite in the iron-formation-soft ore-localized to a large degree along the axes of synclines and in other upward-opening structural traps. Typical occurrences of hard and soft ore are shown in Figure 13.

372


Iron-formation ._____._-----.___I_-- _____- ---.

Sideritic iron-formation

I.. __-.' .-_--I .--. ,.I-------______

,__,

' iron-formation

\ \

SOFT ORE

iron-formation Sideritic u/ / HA R D...O R E /

FIG. 13. Cross-sections showing typical occurrences of iron ore, Michigan.

Acknowledgement

Preparation of this report was aided by published work on the Michigan iron districts by James, 1954 James et al., 1968; Bayley et al., 1966; Huber, 1959, and Prinz, 1967; and unpublished data for the Gogebic district supplied by R. G. Schmidt of the U.S. Geological Survey.

373

J. E. Gair

Résumé

Gisements de feu du Michiguii, aux États-Unis d'Amérique (J. E. Gair)

L'auteur présente un résumé des faciès de formation de fer, de l'environnement dans lequel les dépôts corrélatifs se sont formés ainsi que les modifications qui sont intervenues après le dépôt, y compris celles qui se rapportent à la genèse du minerai. Ces différents aspects sont comparés dans les quatre zones principales d'industrie minière du fer de l'État de Michigan aux États-Unis - les districts Gogebic, Iron- River Crystal Falls, Marquette et Menominee. Dans les districts Gogebic et Marquette les minerais primaires dominants sont du silex à sidérite, peut-être de la magnétite et, dans de petites parties de la zone stratigraphique, du silex à hématite ; dans le district $Iron-River Crystal Falls, le silex à siderite et du matériel pyritique ; dans le district Menominee, du silex à hématite et peut-être de la magnétite. La nature primaire du silex et de la sidérite est indiquée par la régularité des lits, leur composition et leur texture, et par des stylotites et des structures consistant en éboulements consolidés antérieurement. Le silex primaire est indiqué de plus par le drapage de feuillets ferrugineux superposés à lentilles de silex et des fragments d'éboulis. L'hématite primaire est suggérée par son association avec des granules et des colites d'eaux peu profondes et des granules ayant peut-Ctre une origine organique dans des couches épaisses onduleuses ou lenticulaires, et par l'absence d'une oxydation secondaire dans des lits intermédiaires de sidérite, ma- gnétite, silicate de fer et de sédiments clastiques. La magné- tite est importante dans tous les districts sauf dans celui #Iron-River Crystal Falls. Elle y est peut-être primaire, diagénétique ou les deux. Dans la partie est du district de Marquette la plus grande partie de la magnétite est formée par voie diagénétique à partir de sidérite à des profondeurs faibles, à peu près à l'époque où les lits intermédiaires minces de graywacke, de quartzite feldspathique et de formations de fer riches en soude se sont déposés immédia- tement au-dessus. L'origine diagénétique est indiquée par de nombreux flots rémanents de formation de fer sidéri- tique à l'intérieur d'une formation de fer magnétitique.

La sidérite et la pyrite sont attribuées à des conditions presque stagnantes dans les parties les plus profondes de lagunes ou de bassins isolés de la haute mer et l'hématite aux bordures peu profondes de bassins au voisinage du rivage et à la diminution de la profondeur du fait de la sédimentation ou des subsidences. Des formations de fer riches en soude sont associées à des formations d'évaporite.

Dans le district de Marquette certaines jaspilites peuvent être associées à des formations de fer à faciès d'oxyde primaire, et certaines à des formations de fer oxydées pendant la subsidence du Précambrien. moyen, la détérioration météorologique et l'érosion, mais beaucoup de jaspilites dans la partie est du district ne présentent aucune structure liée à des eaux peu profondes ni à une relation évidente avec une formation de fer sidéritique.

La recristallisation des formations de fer s'est produite pendant une orogénie régionale et au cours du niétamor- phisme à la fin du Précambrien moyen, c'est-à-dire il y a 1,7 milliard d'années, la dimension des grains de quartz (silex) variant en relation avec le degré de métamorphisme. La minnesotaite et (ou) le stilpnomelane se sont formés soit durant la déformation aux degrés inférieurs du méta- morphisme ou pendant une diagenèse plus ancienne. Des effets spécifiques diagénétiques ou métamorphiques ne peuvent généralement pas être reconnus. La présence dans la partie est du district de Marquette de formations de fer et de silicates pauvres en silex et de formations de fer sidé- ritique riches en silex, chacune avec à peu près 45 % de silice, fait penser que le contenu en silice a été fixé au cours du dépôt en grande partie sous la forme de silex, et que les minéraux silicatés postérieurs se sont développés par l'incorporation de silex. Les granules riches en silicate dans une matrice pauvre en silicate peuvent être plus facilement expliqués par le remplacement sélectif ou le développement de concrétions durant la diagenèse plutôt que par la diffé- renciation pendant la sédimentation. D e la riebeckite et de l'augite aegyrinique dans la partie est du district de Marquette furent à l'origine d'une formation de fer contenant du silex, du carbonate et de la magnétite riche en soude durant un métamorphisme régional peu avancé. La grunerite est un produit du métamorphisme régional avancé dans la partie ouest du district de Marquette et un produit de métamorphisme de contact local à l'est.

D u minerai d'hématite dure et tendre peut être diffé- rencié en concentrations prémétamorphiques (recristalli- sées) et postmétamorphiques. Le minerai tendre s'est concen- tré dans des structures s'ouvrant vers le haut dans des régions à faible degré de métamorphisme et en relation avec la circulation d'eau souterraine d'une surface d'érosion datant du Précambrien récent. Une origine essentiellement analogue pour le minerai dur du district de Marquette est indiquée par l'association du minerai dur avec une surface d'érosion datant du milieu du Précambrien, malgré l'évi- dence d'effets locaux hydrothermaux.

374



BAYLEY, R. W.; DUTTON, C. E.; LAMEY, C. A. 1966. Geology of the Menominee iron-bearing district, Dickinson County, Michigan, and Florence and Marinette Counties, Wisconsin. Prof. Pap. US. geol. Surv. 513.

DEER, W. A.; HOWIE, R. A.; ZUSSMAN, J. 1962. Rock-forming minerals. New York, Wiley. 4 vols.

GARRELS, R. M. 1960. Mineral equilibria at low temperature and pressure. New York, Harper & Bros. 254 p. HAN, T.-M. 1962 Diagenetic replacement of ore of the Empire mine of northern Michigan and its effects on metallurgical concentration (abs.). 8th Annual Meeting Institute on Lake Superior Geology, Houghton, Mich., Michigan Coll. Mining and Technology, p. 7.

HUBER, N. K. 1958. The environmental control of sedimentary iron minerals. Econ. Geol., vol. 53, p. 123-40.

__ . 1959. Some aspects of the origin of the Ironwood iron- formation of Michigan and Wisconsin. Econ. Geol., vol. 54, no. 1, p. 82-118.

JAMES, H. L. 1954. Sedimentary facies of iron-formation. Econ. Geol., vol. 49, no. 3, p. 235-90.

Geology and ore deposits of the Iron River-Crystal Falls district, Iron County, Michigan. Prof. Pap. U.S. geol. Surv. 570.

KLEIN, Jr. 1966. Mineralogy and petrology of the metamorphosed Wabush Iron Formation, south-western Labrador. J. Petrol., vol. 7, no. 2, p. 246-305.

KRAUSKOPP, K. B. 1957. Separation of manganese from iron in sedimentary processes. Geochim. et cosmoch. Acta, vol. 12,

PRINZ, W. C. 1967. Pre-Quaternary geologic and magnetic map and sections of part of the eastern Gobegic Iron Range, Michigan. Misc. geol. inv. M a p US. geol. Surv., 1-497. WINCHELL, A. N. 1951. Elements of optical mineralogy, 4th ed. New York, Wiley.

JAMES, H. L.; DUTTON, C. E.; PETTIJOHN, F. J.; WIER, K. L. 1968.

p. 61-84.

375


On an international basis, the nomenclature of banded ferruginous-cherty sedimentary rocks is unfortunately impre- cise. In different countries, different terms are used for the same rocks and identical terms are also used for different rocks. In order to come to a common understanding throughout the world as to exactly what particular type of iron-rich rock a particular name is meant to describe, clear concise definitions of terms should be presented in papers of more than local significance. In the published papers of the International Symposium on the Geology and Genesis of Precambrian Iron-Manganese Formations and Ore Deposits, held in Kiev, U.S.S.R., in1970, the following terms were used by Russian contributors: taconite (talconit), itabirite (itabivit), jaspilite (dzhespilit), ferruginous quartzite (zhelezisty Jcvartsit j, iron hornfels (zhelezisty rogovik), ferruginous chert (zhelezisfy Jcremen’), ferruginous jasper (zhe- lezistuya yashma), iron ore (zheleznaya rida).

Usage in the U.S.S.R. differs considerably from that in much of the Western world; usage in the Western world is also not consistent. In the Western world a greater emphasis is placed on conditions of sedimentation of iron- formations. The effects of widely differing degrees and types of metamorphism further complicate nomenclature.

In the United States of America, Canada, Australia and South America, the generic term for banded ferruginous-cherty rocks of sedimentary origin has come to be ‘ iron-formation’ (zhelezìstayajovrnafsiuj. Most geologists in those countries accept James’ definition (1954, p. 239) or that of Gross (1966, p. 41). Iron-formation, generally believed to be a dominantly chemical (or biochemical) precipitate, typically consists of chert (kremen‘) or jasper (yashma) interbanded with one or more iron-rich minerals: oxide, carbonate, silicate or sulphide. Very rarely, the rock does not contain chert. On both empirical and theoretical (Krumbein and Garrels, 1952; Kiauskopf, 1967) grounds, the primary iron-rich mineral is an indicator of the p H and Eh of the environment of deposition of the iron- formation. Therefore the dominant type of iron mineral was used by James to name primary ‘facies’ of iron- formation, Oxide-facies (hematitic) iron-formation indicates

a positive Eh, sulphide-facies a strongly negative Eh, and carbonate- and silicate-facies are intermediate. The facies are intergradational; the hematitic oxide- and the sulphide- facies are incompatible.

The term iron-formation in Western usage is strictly parallel to limestone, a generic lithologic name. Just as there are many different types of limestone, there are also different types of iron-formation. Thus, in naming formations in a stratigraphic sense, a formation in Australia may be called the ‘ Brockman Iron Formation’ for example, just as we may call an American limestone formation the ‘Niobrara Limestone’. To avoid confusion, the U.S. Geo- logical Survey hyphenates iron-formation when the words are used in a lithologic sense; the words are capitalized when used in a stratigraphic sense. This practice might be more widely adopted.

The term ‘jaspilite’ was ñrst applied in the Lake Superior area to oxide-facies iron-formation in which the silica is present as jasper (yashma) . Subsequently, jaspilite has achieved wide international usage for other oxide-facies iron-formations. Unfortunately, in some areas it has been applied to rocks that would not be called jaspilite in the Lake Superior area.

‘Taconite’ is another term for iron-formation that originated in the United States and has achieved some international currency, particularly in the U.S .S.R. The word is a general term, now used primarily by mining engineers and metallurgists, and is without exact niineral- ogical or environmental implications. Therefore, many geologists feel it should be dropped from scientific literature in favour of more specific terminology.

In South Africa, oxide-facies iron-formation has been called banded ironstone, although on the American continents and elsewhere ‘ironstone’ is reserved for the minette- type ores, generally not cherty or banded, commonly in part clastic and fossiliferous, and almost everywhere post- Cambrian in age. The distinction between iron-formation and ironstone seems worth preserving, because the environment and processes of deposition are different for the two rock types.

Unesco, 1973. Genesis of Precambrian iron and tnanganese deposits. Proc. Kiev Synrp., 1970. (Earth sciences, 9.) 377


Most, but not all, Precambrian iron-formations have been metamorphosed. James (1955) has shown that the grain size of the recrystallized chert or jasper varies directly with the degree of metamorphism. Susceptibility of oxide- facies iron-formation to supergene enrichment is closely controlled by grain size (Dorr, 1964).

‘Itabirite’ is a Brazilian term that has achieved wide usage in South America, West Africa, and elsewhere for oxide-facies iron-formation that has been metamorphosed to a degree that makes the individual crystals of the rock megascopically distinguishable (Dorr and Barbosa, 1963). The term has a specific and restricted meaning in both field and economic applications and may be worth preserving in the international nomenclature.

‘Banded hematite-quartzite’ (BHQ) is a term widely used in India and to a lesser extent in Australia and elsewhere for oxide-facies iron-formation. Some of this material has been highly enough metamorphosed to make it the equivalent of the itabirite of Brazil; much is of lower metamorphic grade and cannot be considered the equivalent of that rock type. The latter is, in part, an equivalent of the jaspilite of the Lake Superior region.

Although iron-formation may locally contain considerable amounts of detrital material such as interbedded or intermixed shale, tuff and even sand or pebbles, it is dominantly a chemical or biochemical precipitate. ‘Fer- ruginous quartzite’ is, in Western usage, reserved for rocks of dominantly detrital origin. Although the rock may have essentially the same chemical composition as iron-formation, the quartz, and in many cases the iron minerals, are clastic in origin. It may or may not be grossly banded. In the U.S.S.R., however, ferruginous quartzite is used, according to Semenenko (1956, 1959, 1967), in three different senses: ferruginous clastic quartzose rock, coarse-grained metamorphosed iron-formation of either oxide- or silicate-facies, and all ferruginous cherty rocks.

Hornfels’ is a term used with widely different meanings in the West and the U.S.S.R. In the West, hornfels is ‘a fine-grained nonschistose metamorphic rock resulting from contact metamorphism. Large crystals may be present and may represent either porphyroblasts or relict phenocrysts’.l To our knowledge, hornfels has never been applied to iron-formation in the West. In the U.S.S.R., hornfels is commonly used for fine-grained rocks including, but not restricted to, silicate- and oxide-facies iron-formation, that need have no relation to contact metamorphism. The material may be somewhat schistose or foliated by regional or dynamic metamorphism. ‘Iron hornfels’ in the U.S.S.R. literature is a coarse-banded iron silicate-chert rock with

fine-grained quartz; some authors also consider it as a synonym for ferruginous jasper.

‘ Jaspilite’ in the U.S.S.R. is a banded rock with iron present as hematite, magnetite, or martite and silica as ‘fine-grained quartz-jasper or hornfels’. The term ‘ itabirite’ has no usage in the U.S.S.R., but is implied for metamorphosed jaspilite.

‘Iron ore’ is used very loosely in the U.S.S.R., as it is by some authors in the West. In some cases the term has specific economic implication, in others it has no implication of economic viability. Similarly, the word ‘iron’ is very loosely used in the literature of the U.S.S.R.; it merely indicates the presence of the element in some form, without any implication as to quantity or oxidation state or chemical composition of the iron-bearing mineral. The usage is the Sam= as that in the term iron-formation; this can be disconcerting to Western readers when applied to ‘iron hornfels’, ‘iron quartzite’ or ‘iron chert’.

Western readers commonly approach geological literature of the U.S.S.R. via a translation. One often wonders how much of the terminological difficulty is caused by inept translation rather than original usage; certainly our colleagues in the U.S.S.R. must have the same problem with our literature.

Eventually it will be to the advantage of our science to adopt a nomenclature that can be used on a world-wide basis to describe these distinctive ferruginous rocks that are very common in the Precambrian sedimentary column and are also known in the early Palaeozoic. The more specific the meaning of the words used, the easier will be international scientific communication. It would be pre- sumptuous of the small group taking part in the symposium at Kiev to attempt to set up such international standards; for this season we asked each author in the symposium to define his terms. It is to be hoped that from this small beginning a coherent and internationally acceptable nomenclature for these rocks will eventually evolve. Until it does, clear definitions of rock terms used in papers for international audiences, if only by reference to standard accessible publications, will prevent obscurities and mis- understandings.

The Ad Hoc Committee on Nomenclature was composed of the following geologists participating in the Kiev symposium: R. T. Brandt (Australia); J. Van N. Dorr II (United States of America); G. A. Gross (Canada); H. Grüss (Federal Republic of Germany); and N. P. Seme- nenko (U.S.S.R.).

1. American Geological Institute, Glossary, 2nd ed., p. 140.

378


Russian terms English equivalent

General terms Ferruginous-cherty forination

Iron-formation = ferruginous formation

Iron ore

Sedimentary facies terms Iron-cherty-slate-keratophyrel

(Zhelezisto-kremnisto-slantsevokeratofivo Jaspilite-leptite iron-cherty-metabasitel (Dzhespilito-leptitovaya zhelezisto-kremnisfo-metabazitovaya)

Iron-cherty-ultrabasitel (Zhelezisto-lcremnisto-ul'trabazifovaya)

Iron-cher ty-slatel (Zhelezisto-kremnisto-slantsevaya)

Ferric ferrous (oxide-protoxide)-iron-cherty ( Okisno-zalcisnaya zhelezisto-lcremnistaya)

Iron-cherty carbonatel (Zhelezisto-kremnisto-karbonatnaya)

Iron-cherty silicate1 (Zhelezisto-kremnisto-silika friaya)

Metamorphic facies terms Slate stagel

Phyllite stagel

(Z~ielezisto-kremnistaya formatsia)

(Zhelezistaya formaisia)

(Zheleznaya ruda)

(Stupen' aspidnykh slantsev (pumpellitovaya fatsia))

(Filitovaya stupen' (zelenoslantsevaya fatsia))

(Rogovikovaya stupen')

(Gneisovaya stupen' (amnfibolitovaya i granulifovaya fatsia))

Hornfels stagel

Gneiss stage'

Petroguaphic terms Hornfels2 (Rogoviki)

Iron hornfels (Zhelezisfye rogoviki)

Ferruginous quartzite (Zhelezisty kvartsit)

Taconite (Takonit)

Itabirite (Ztabirit)

Jaspilite

Iron chert

Jasper

(Dzhespilit)

(Zhebzisty /cremen')

( Yashma)

Iron-formation (geological term)

Iron-formation (geological term)

Iron-formation (economic term)

Iron-formation, Algoma type, associated with keratophyres

Iron-formation, Algoma type, associated with leptites-metabasites

Iron-formation, Algoma type, associated with ultrabasites

Iron-formation, Lake Superior type, oxide facies

Iron-formation, Lake Superior type, ferric ferrous facies

Iron-formation, Lake Superior type, carbonate facies

Iron-formation, Lake Superior type, silicate facies

Epizonal metamorphic iron-formation (pumpellyite facies)

Mesozonal metamorphic iron-formation Greenschist facies

Hornfels facies

Katazonal metamorphic Amphibolite facies Granulite facies

Fine-grained metamorphic quartzite, including iron-formation

Iron silicate chert rock coarse bandedz Ferruginous chertz Ferruginous clastic quartzose rock1 Coarse-grained, metamorphosed iron-formation* Ali ferruginous cherty rocks' Itabirit2 Silicate ferruginous quartzite2 Silicate iron ' hornfeW2 Silicate 'itabirW2 Non-silicate ferruginous quartzite' Non-silicate iron 'hornfels' Ferruginous quartzite' Metamorphosed jaspilitel Ferruginous jasper Non-silicate iron 'hornfels', fine banded2 Chert

rocks

Jasper3

1. Used or defined by N. P. Semenenko. 2. Used or defined by R. R. Petrov. 3. Used or defined by V. M. Chernov.

379



DORR, J. V. N., II. 1964. Supergene iron ores of Minas Gerais, Brazil. Econ. Geol., vol. 59, p. 1203-40.

DORR, J. V. N.; BARBOSA, A. L. M., 1963. Geology and ore deposits of the Itabira district, Minas Gerais, Brazil. Prof. Pap. U.S. geol. Scirv. 341-c, 110 p.

GROSS, G. A. 1966. Principal types of iron-formation and derived ores. BuII. Canad. Inst. Min. vol. 59, no. 648, p. 150-3.

JAMES, H. L. 1954. Sedimentary facies of iron-formation. Econ. Geol., vol. 49, p. 235-93.

-. 1955. Zones of regional metamorphism in the Precambrian of northern Michigan. Bull.geol. Soc. Amer., vol. 66, p. 14-56-87,

KRUMBEIN, W. C.; CARRELS, R. M. 1952. Origin and classification of chemical sediments in terms of p H and oxidation- reduction potentials. J. Geol., vol. 60, p. 1-33.

KRAUSKOPF, K. B. 1967. Introduction to geochemistry. New York, McGraw-Hill. 721 p.

SEMENENKO, N. P. et al., 1946-1953. Struktura rudnykh poky Krivorozhskikh zhelezorudnykh mestoruzhdeniy [Structure of ore fields of Krivoy-rog iron ore deposits]. BuIl. Acad. Sei. U.R.S.S., Kiev, vol. 1, 1946; vol. II, 1953. - . 1956. Petrography of iron-cherty formations of Ukrainian S.S.X. Kiev, Ukrainian Academy of Sciences. (In Russian.) - . 1959. Geologiya zhelezisto-kremnistykh formatsiy Ukrainy [Geology of iron-cherty formations of Ukrainian S.S.R.]. BuII. Acad. Sci. U.R.S.S., Kiev. - . 1967. Geology of sedimentary-volcanogenic formations of Ukrainian Shield. Kiev, Naukova Dumka. (In Russian.)

3 80

List of participants / Liste des participants

BELEVTSEV, D r Y. N., Institut of Geochemistry and Physics o Metals, Academy of Sciences of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

BELYAEV, D r M. V., Geological Service, Ministry of Ferrous Metais. Industry of the Ukrainian S.S .R., Dniepropetrousk (Ukrainian S.S.R.).

BEYGULENKO, D r I. P., Geological Service, Ministry of Fer- rous Metals Industry of the U.S.S.R., Moscow (U.S.S.R.).

BORISENKO, D r S. T., Geological Prospecting Service, Minis- try of Geology of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

BRANDT, D r R. T., Goldsworthy Mining Limited, P.O. Box 84, Port Hedland, Western Australia 6721 (Australia).

CAMBEL, D r B., Slovak Geological Institute, Obrancov mieru 41, Bratislava (Czechoslovakia).

CHERNOV, D r V. M., Institut of Geochemistry, Karelian Branch of the Academy of Sciences of the U.S.S.R., Pe- trozavodsk (US3 .R.).

CHOUBERT, Dr Boris, Résidence Bernard-Palissy, 77 Avon- Fontainebleau (France).

CHOUBERT, D r Georges, Directeur de Recherches, Bureau de Cartographie Géologique Internationale, Muséum Na- tional d‘Histoire Naturelle, 36 Rue Geoffroy Saint-Hilaire, 75005 Paris (France).

DORR, Dr John vanN., II, U.S. Geological Survey, Washington, D.C. 20242 (United States of America).

DZHEZDALOV, D r A. T., Trust ‘Lenruda’, Ministry of Ferrous Metals Industry of the Ukrainian S.S.R., Krivoyrog (Ukrai- nian S.S.R.).

EGOROV, D r E. V., Far East Geological Service, Ministry of Geology of the Russian Soviet Federated Socialist Republic, Khabarovsk (R.S .F.S .R.).

FAURE-MURET, Miss A., Muséum d‘Histoire Naturelle, 36 Rue Geoffroy Saint-Hilaire, 75005 Paris (France).

FRIETSCH, D r Rudyard, Geological Survey of Sweden, 104 05 Stockholm 50 (Sweden).

CAIR, D r Jacob E., U.S. Geological Survey, Washington, D.C. 20242 (United States of America).

GAVELYA, D r A. P., Trust ‘Krivbassgeologiya’, Ministry of Geology of the Ukrainian S.S.R., Krivoyrog (Ukrainian S .S .R .) .

GOODWIN, Professor A. M., Department of Geology, Univer- sity of Toronto, Toronto 5 (Canada).

GORYAINOV, Dr M. V., The Kola Branch of the Academy

of Science of the U.S.S.R., Apatity, Murmansk Region (U3 .S .R.).

GROSS, D r G. A., Head Geology of Mineral Deposits Section, Geological Survey of Canada, 601 Booth Street, Ottawa 4, Ontario (Canada).

GROSSI SAD, Geol. J. H ., Dept. Engenharia de Minas, Univer- sidade de Minas Gerais, Escola de Engenharia, Rua Espirito Santo, 35-7”, Belo Horizonte, M. G. (Brazil).

GRUSS, D r Hans, Gewerkschaft Exploration, Steinstrasse 20, Postfach 3526, Düsseldorf (Federal Republic of Germany).

INGERSON, Professor Earl, Department of Geological Sciences, University of Texas at Austin, Austin, Texas 78712 (United States of America).

KALUGIN, D r A. S., Siberia Research Institute of Geology, Geophysics and Mineral Resources, Ministry of Geology of the US .S .R., Novosibirsk (U3 .S .R.).

KOBZAR, Dr V. N., Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

KRAVCHENKO, D r V. M., Yakut Thematic Expedition, Ministry of Geology of the Russian Soviet Federated Socialist Re- public, Yakutsk (R.S.F.S.R.).

KRISHNAN, D r M. S., Hyderabad (India). MACLEOD, D r W. N., 6 Airlie Street, Peppermint Grove, Western Australia GOO5 (Australia). MALYLITIN, D r E. I., Ministry of Ferrous Metals Industry of the U.S.S.R., Moscow (U.S.S.R.).

MITKEEV, D r M. B., Trust ‘Dnieprogeologiya’, Ministry of Geology of the Ukrainian S.S.R., Dniepropetrovsk (Ukrai- nian S.S.R.). MOMDZHI, Dr Y. S., All-Union Research Institute of Mineral Resources, Ministry of Geology of the U.S.S.R., Moscow (U.S.S.R.).

MOREY, D r G . B., Minnesota Geological Survey, University of Minnesota, Minneapolis, Minnesota 55455 (United States of America).

NIKIFOROV, D r M . S., Trust ‘Dzerzhinskruda’, Ministry of Ferrous Metal Industry of the Ukrainian S.S.R., Krivoyrog (Ukrainian S.S.R.).

NOVOKHATSKY, D r I. P., Institute of Geological Sciences, Academy of Sciences of the Kazakh S.S.R., Alma-Ata (Kazakh S.S.R.).

PERCIVAL, D r F. G., Sadlers End, Haslemere, Surrey (United Kingdom).

381

List of participants

PERRY, D r Eugene C., Jr, Department of Geology, University of Minnesota, Minneapolis, Minnesota 55455 (United States of America).

PLAKSENKO, Dr N. A., Voronezh State University, Voronezh (U .S .S .R.).

POLUNOVSKIY, Dr R. M., Azov Expedition, Ministry of Geology of the Ukrainian S.S.R., Volnovakha (Ukrainian S.S.R.).

RIBEIRO FILHO, Professor Evaristo, Instituto de Geociencias e Astronomia, Universidade de Sao Paulo, Cidade Universi- taria, São Paulo (Brazil).

ROY, Dr Supriya, Department of Geological Sciences, Ja- davpur University, Calcutta-32 (India).

SCARPELLI, D r Wilson, c/o ICOMI, Caixa Postal 396, Belem do Para (Brazil).

SEMENENKO, Academy Professor N. P., Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrai- nian S.S.R., Kiev (Ukrainian S.S.R.).

SHKOLNIK, D r E. P., Far East Geological Service, Ministry of Geology of the Russian Soviet Federated Socialist Republic, Khabarovsk (US .S .R.).

SHKUTA, D r E. I., Geological Service, Ministry of Ferrous Metals Industry of the Ukrainian S .s.R., Dniepropetrovsk (Ukrainian S.S.R.).

SHTSHERBAK, D r V. M., Institute of Geological Sciences, Academy of Sciences of the Kazakh S.S.R., Alma-Ata (Kazakh S.S.R.).

SHTSHERBAKOV, D r B. D., Ministry of Geology of the U.S.S.R., Moscow (U.S.S.R.).

SIMS, Dr Samuel J., Bethlehem Steel Corporation, Bethlehem, Pennsylvania 18016 (United States of America).

SIROSHTAN, D r R. I., Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

STRUEV, Dr M . I., Ministry of Geology of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

TOKHTUEV, D r G. V., Institute of Geochemistry and Physics of Minerals, Academy of Sciences of the Ukrainian S.S.R., Kiev (Ukrainian S.S.R.).

TOLBERT, Dr G. E., Cia Vale do Rio Doce Div. de Desenvol- vimento Av. Graça Aranha, 26-8 andar Rio de Janeiro(Brazi1).

TRENDALL, Dr A. F., Geological Survey of Western Australia, 26 Francis Street, Perth, Western Australia (Australia).

TUGARINOV, Professor A. I., Institute of Geochemistry and Analytical Chemistry, Academy of Sciences of the U.S.S.R., Moscow (U.S.S.R.).

VERIGIN, Dr M. I., Trust ‘Skrivbassgeologiya’, Ministry of Ge- ology of the Ukrainian S.S.R., Krivoyrog (Ukrainian S.S.R.). WEBER, D r F., Laboratoire de géologie et de paléontologie, Université de Strasbourg, 67 Strasbourg (France).

ZAITSEV, Dr Y. S., Voronezh Geological Prospecting Ex- pedition, Ministry of Geology of the Russian Soviet Feder- ated Socialist Republic, Voronezh (U.S.S.R.).

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Documents

Genesis of Precambrian iron and manganese deposits