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oxygenation of the atmosphere
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www.elsevier.com/locate/earscirev
Earth-Science Reviews
Archaean atmospheric evolution: evidence from the
Witwatersrand gold fields, South Africa
Hartwig E. Frimmel*
Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa
Received 6 July 2004; accepted 12 October 2004
Abstract
The Witwatersrand gold fields in South Africa, the world’s largest gold-producing province, play a pivotal role in the
reconstruction of the Archaean atmosphere and hydrosphere. Past uncertainties on the genetic model for the gold caused
confusion in the debate on Archaean palaeoenvironmental conditions. The majority of Witwatersrand gold occurs together with
pyrite, uraninite and locally bitumen, on degradational surfaces of fluvial conglomerates that were laid down between 2.90 and
2.84 Ga in the Central Rand Basin. Although most of the gold appears as a precipitate within, or associated with, post-
depositional hydrothermal phases and along microfractures, available microtextural, mineralogical, geochemical and isotopic
data all indicate that this hydrothermal gold, analogous to some pyrite and uraninite, was derived from the local mobilisation of
detrital particles. Some of the key pieces of evidence are a significant correlation of the gold, pyrite and uraninite with other
heavy minerals as well as sedimentary lithofacies, local preservation of in-situ gold micronuggets and abundant rounded forms
of pyrite and uraninite, compositional heterogeneity on a microscale of the gold as well as the rounded pyrite and uraninte, and
radiometric age data that indicate an age of the gold, pyrite and uraninite that is older than the maximum age of deposition for
the host sediment. None of these observations/data is compatible with any of the suggested hydrothermal models, in which
auriferous fluids were introduced from an external source into the host rock succession after sediment deposition. In contrast,
those arguments, used in favour of hydrothermal models, emphasise the microtextural position of most of the gold, which
highlights the undisputed hydrothermal nature of that gold in its present position, but does not explain its ultimate source.
Furthermore, the macro-scale setting of the stratiform ore deposits is in stark contrast to any known type of epigenetic,
hydrothermal gold deposit. Consequently, the best-fit genetic model involves post-depositional textural and mineralogical
modification of original fluvial placer deposits.
On the Kaapvaal Craton, Witwatersrand-type mineralisation is recorded over an extended period of time from 3074 to 2642
Ma. Rounded pyrite is common in the coarser grained fractions of the siliciclastic basin fill. A lack of sulphur isotope
fractionation and typical magmatic d34S values support its detrital origin. Together with rounded uraninite, which is particularly
abundant in the older beds, it provides important constraints on the redox potential of the Meso- to Neoarchaean (3.1–2.6 Ga)
atmosphere and hydrosphere. In combination with eukaryotic steroids documented from the Pilbara Craton, Australia, the
0012-8252/$ - s
doi:10.1016/j.ea
* Current addr
fax: +49 931 88
E-mail addr
70 (2005) 1–46
ee front matter D 2004 Elsevier B.V. All rights reserved.
rscirev.2004.10.003
ess: Institute of Mineralogy, University of Wqrzburg, Am Hubland, D-97074 Wqrzburg, Germany. Tel.: +49 931 888 5420;
8 4620.
esses: [email protected], [email protected].
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–462
ambient Neoarchaean oxygen fugacity is calculated as having been approximately 10�3, in equilibrium with a relatively acidic
hydrosphere (pH=6). This is in agreement with the preservation of mass-independent S isotope fractionation, which provides
independent support for an anoxic atmosphere and which has so far been recorded predominantly from sediments older than 2.3
Ga. An acidic meteoric palaeoenvironment is supported by intense chemical weathering below erosional unconformity surfaces
in the Witwatersrand Basin. In contrast to the pyrite-bearing fluvial and near-shore shallow marine deposits, marine shale
deposits contain magnetite. This supports the postulated reducing environment but also highlights total sulphur concentrations
in the ancient ocean that were orders of magnitude lower than in modern ocean water.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Witwatersrand; Palaeoplacer deposits; Anoxic Archaean atmosphere; Gold genesis
Fig. 1. Contrasting models for the evolution of atmospheric
chemistry: (A) Models involving a reducing Archaean atmosphere
pO2 evolution curve proposed by (a) Kasting (1987, 2001) and (b)
Rye and Holland (2000), remaining curves after Pavlov et al
(2001a) and Kasting (2001); (B) model involving an oxidising
atmosphere (Ohmoto, 2004).
1. Introduction
The basic conditions for the most important aspects
of modern life on Earth have been shaped as early as
in Precambrian times. These include most likely the
origin of life and definitely the evolution from the
origin of the eukaryote cell to hard bodied animals,
the growth of continental crust, the oxygenation of the
atmosphere as well as the formation of mineral
deposits without which modern civilisation would be
unthinkable. All of these aspects appear to be
interrelated but the various feed-back mechanisms
that controlled the relationships between life’s evolu-
tion, plate tectonic processes, palaeoclimate and
distribution of metals between hydrosphere and geo-
sphere throughout the critical periods of the Precam-
brian remain highly speculative. In fact, the
determination of the above relationships can be
regarded as a fundamental problem in earth sciences.
One of the more important questions concerns the
timing and cause of the rise of atmospheric O2.
Photosynthesis was the principal process by which
free O2 was produced, though some of it may also be
ascribed to hydrogen escape to space after CH4
photolysis (Catling et al., 2001). Consequently, the
evolution of the Archaean (3.8–2.5 Ga) to Proterozoic
(2.5–0.54 Ga) atmosphere must have been strongly
influenced by organisms. Production and decomposi-
tion of organic matter, in combination with the
volcanic degassing of the planet, also influenced the
atmospheric CO2 and CH4 concentrations through a
carbon cycle that involves chemical weathering and
formation of silicates and carbonates. As both CO2
and CH4 are greenhouse gases, the evolution of life
and that of Archaean to Eoproterozoic (2.5–2.05 Ga)
palaeoclimate are intimately interlinked.
General agreement seems to exist on substantial O2
levels from around 2.3 Ga onwards, but the preceding
atmospheric evolution is a matter of intense debate
(Fig. 1). Two competing schools of thought try to
explain the environments and mechanisms that con-
trolled the emergence of life: one assumes that life
originated under a reducing atmosphere (O2b1 ppm)
and that the formation of an oxic atmosphere triggered
the emergence of eukarya (Kasting and Siefert, 2002;
Knoll, 1992; Rye and Holland, 2000), whereas the
other proposes the emergence of oxygenic photo-
;
.
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 3
synthetic organisms as early as 4 Ga with essentially
constant atmospheric O2 levels (N10%) since then
(Lasaga and Ohmoto, 2002). Towe (1990) argued for
a low (0.2–0.4%) but stable O2 concentration in the
Archaean that was maintained by aerobic respiration.
This is in contrast to anaerobic carbon cycles that are
implicit in any model of a reducing Archaean
atmosphere.
The debate extends to the question of Archaean
palaeoclimate. As the Sun’s luminosity is estimated to
have been about 30% less than today (Newman and
Rood, 1977), Earth’s hydrosphere should have been
completely frozen. No geological support exists for
such a scenario, however, and elevated greenhouse
gas concentrations have therefore been proposed to
offset this so-called Faint Young Sun Paradox (Sagan
and Mullen, 1972). A dense Archaean atmosphere is
thought to have consisted largely of CO2, and the
stronger greenhouse gases H2O and CH4. For the time
around 3.0 Ga, atmospheric CO2, CH4 and O2
concentrations of approximately 3000, 1000 and
b0.1 ppm, respectively, have been suggested by those
advocating a reduced Archaean atmosphere (Kasting,
2001; Kasting and Siefert, 2002; Rye and Holland,
2000), which implies that CH4 was the dominant
greenhouse gas. In contrast, some workers (Lasaga
and Ohmoto, 2002; Ohmoto et al., 1999; Ohmoto,
2004) argue for CO2 having been the principal
greenhouse gas at that time at a level of about 10%.
As the continents tied up vast amounts of carbon in
the form of organic matter, hydrocarbons, and
carbonaceous as well as carbonate rocks, their growth
must have led to the removal of considerable amounts
of carbon from active circulation. In addition to burial
of biogenic carbon in sediments, the consumption of
CO2 by biota and the oxidation of CH4 by biogenic O2
must have caused a further decrease in the concen-
tration of greenhouse gases throughout the Neo-
archaean (2.8–2.5 Ga) and Proterozoic. This
decrease was further exacerbated by a decrease in
radiogenic heat production, volcanic activity, and thus
a decrease in the emission rate of potential greenhouse
gases. The consequential cooling of the Earth’s
surface triggered repeated global glaciations. Owing
to their old age, the preservation potential of Archaean
glacial deposits is relatively small, but glaciogenic
deposits have been reported from Mesoarchaean (3.2–
2.8 Ga) sedimentary successions (Young et al., 1998).
They are more common in Proterozoic successions, in
which major glaciations are recorded around 2.35 Ga
and again repeatedly throughout the Neoproterozoic
(1.0–0.54 Ga). Their close stratigraphic association
with marine carbonate successions is suggestive of
severe climate fluctuations, which highlight that no
equilibrium had been achieved between solar lumi-
nosity, greenhouse gas concentrations and global
climate.
It was by no means only carbon that was removed
from active circulation in the biosphere during the
Precambrian. Apart from gravitation towards the
planet’s core during the very early stages of Earth’s
history, most heavy metals were removed from
biogeochemical circulation by both the growth of
continental crust and mineralisation processes. Indi-
cations of the latter exist in the form of huge
sedimentary Fe, Cu, Pb–Zn, Mn and possibly also U
and Au deposits that are characteristic of the Archaean
and Proterozoic Aeons (Lambert et al., 1992). To use
Fe as an example, the distribution of iron formations,
which provide the best direct evidence of Fe-depletion
of the ocean water, appears particularly intriguing.
Their occurrence in Archaean and Eoproterozoic
strata has been explained by lower atmospheric O2
concentrations at those times that made possible a
stratified ocean with Fe-rich bottom waters (Trendall,
2002). According to this model, a steady increase in
ocean oxygenation towards the end of the Eoproter-
ozoic, as living organisms and photosynthesis became
more abundant and effective, would have lowered the
interface between Fe-rich anoxic bottom waters and
oxygenated waters down to effectively seafloor level
and thereby prohibited any further deposition of iron
formations. Global ice ages seem to be particularly
favourable for the deposition of iron formations when
a largely, or possibly completely, ice-covered Earth
would have shut down ocean water circulation and led
to the development of anoxic, Fe-rich ocean bottom
waters (Kirschvink, 1992; Klein and Beukes, 1993).
Although a relationship between iron formation,
palaeoclimate, palaeolatitude and volcanism is indi-
cated, the nature of this relationship remains elusive
and is the subject of on-going debate in the current
literature.
It appears clear from the above that considerable
uncertainty exists regarding the palaeoenvironmental
conditions prior to 2.2 Ga. One of the best natural
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–464
laboratories to study these conditions is the Kaapvaal
Craton in South Africa. It hosts, similar to the Pilbara
Craton in Australia, an undisturbed sedimentary cover
(Transvaal Supergroup) that spans almost the entire
Eoproterozoic Era. This sedimentary succession con-
tains prime examples of banded iron formations and
glaciogenic diamictite deposits, as well as a number of
paleosols that have been studied extensively by
numerous researchers and provided most useful
constraints on the atmospheric evolution in the
Eoproterozoic. One of the more recent results that
emanated from these studies is the recognition that
already at 2.2 Ga weathering was comparable with
modern tropical laterites under an oxic atmosphere
(Beukes et al., 2002).
The Kaapvaal Craton is also host to some of the
world’s finest examples of Palaeoarchaean (3.6–3.2
Ga) sedimentary successions. These occur in the 3.5
to 3.2 Ga Barberton Supergroup of the Barberton
Greenstone Belt as the Fig Tree and Moodies Groups.
There the rocks contain evidence of Fe-rich hydro-
thermal vents forming iron formations on the seafloor
(de Ronde et al., 1994), extensive silicification that
signals ocean water saturated in SiO2, as well as
remnants of early stromatolites (Byerly et al., 1986)
and microfossils (Walsh, 1992).
Whereas these earliest sedimentary rocks provide
minimum constraints on the age of the oldest life
forms, and the younger Eoproterozoic successions
provide minimum constraints on the timing of an oxic
atmosphere covering the planet, the time in between,
i.e. the Meso- and Neoarchaean Eras, holds the key
for the principal question of atmospheric evolution
and its bearing on the further evolution of life
following its emergence sometime in the early
Archaean. The best available geological record to
study the palaeoenvironmental conditions during that
crucial time span is the siliciclastic sediment fill of the
Mesoarchaean Witwatersrand Basin on the Kaapvaal
Craton. It represents the world’s best preserved
Archaean sedimentary succession and it contains
redox-sensitive minerals that might hold the key for
our understanding of the Mesoarchaean atmosphere
and hydrosphere.
Iron oxides are conspicuously lacking in fluvial
sediments of the Witwatersrand, with the principal Fe-
bearing phase being pyrite. The coarse-grained
fraction of these pyrite-rich, fluvial deposits (con-
glomerate) hosts the world’s largest known accumu-
lation of gold, but also represents its largest unmined
inferred uranium resource. More than 49,400 metric
tonnes (t) of gold have been produced from these
conglomerate beds (reefs) between 1886 and 2003,
amounting to almost 40% of all the gold ever mined
during recorded history (Frimmel and Minter, 2002;
Sanders et al., 1994). South Africa is still the world’s
number one gold producer with a share of approx-
imately 16% and, according to the South African
Chamber of Mines, the remaining reserves in the
Witwatersrand Basin, estimated around 38,000 t
(Frimmel and Minter, 2002), amount to 46% of
known world reserves. Between 1952 and 1975, as
much as 1.5 t of U3O8 were produced from
Witwatesrand conglomerates at an average grade of
271 ppm (Frimmel et al., in press). Although poorly
exposed, the Witwatersrand is one of the best-
documented basins of its kind in the world thanks to
more than 100 years of underground mining and
exploration. Yet, in spite of the enormous economic
significance of the Witwatersrand gold deposits, the
genesis of these deposits is still a matter of con-
troversy. Comparing the extensive older with the more
recent literature (for a review of the older and more
recent literature see Pretorius, 1975 and Frimmel and
Minter, 2002, respectively), it appears fair to say that
the debate around this controversy has not lost
anything of its intensity since it had been described
by Davidson (1965) as bthe most disputed issue in the
history of economic geologyQ.Depending on the preferred genetic model for the
Witwatersrand gold, the pyrite and uraninite, both of
which occur predominantly as rounded particles, are
interpreted either as sedimentary heavy sands or as
hydrothermal precipitates. In the latter case, the
abundant rounded pyrite is explained as pseudomor-
phic replacement of detrital Fe–Ti oxides, Fe-pisolite,
ferricrete, banded iron formation and Fe-rich shale
(Phillips and Law, 2000) and/or product of post-
depositional dissolution and re-precipitation mecha-
nisms (Barnicoat et al., 1997; Phillips and Myers,
1989). The different genetic models have equally
different ramifications for the inferred redox state of
the atmosphere at the time of sediment deposition as
they imply either Fe-sulphides or Fe-oxides as having
been stable under the Archaean atmosphere. One of
the major goals of this paper is therefore, to review
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 5
and assess the various genetic models that have been
proposed for the Witwatersrand gold deposits, and by
implication also for the associated pyrite and uranin-
ite. It will be shown that the best-fit genetic model is
that of modified palaeoplacer deposits. The signifi-
cance thereof for the reconstruction of the Meso- to
Neoarchaean atmosphere will then be discussed.
The siliciclastic successions of the Witwatersrand
are characterised by an abundance of erosional
unconformity surfaces. Theoretically, geochemical
studies across these unconformities should make it
possible to gain insight into the extent of chemical
weathering at the time. Similarly as with the debate
around the genesis of the gold and the associated
pyrite and uraninite, a difference in opinion exists
with regard to the cause of widespread acidic
alteration of the siliciclastic mineral assemblages that
has been reported from throughout the Witwatersrand
Basin (Barnicoat et al., 1997; Phillips and Law, 1994).
Weathering under an aggressive, CO2 and/or CH4-
rich, acidic atmosphere should lead to considerable
chemical change in the weathered rock (loss of
alkalies and alkaline earths). Unfortunately, such a
chemical change would be difficult to distinguish
from acid leaching by post-depositional magmatic or
metamorphic fluids. Thus the question arises whether
systematic chemical changes observed across the
unconformities reflect paleosols or post-depositional
hydrothermal infiltration. Those workers who prefer a
hydrothermal model for the gold, pyrite and uraninite,
postulate basin-wide H+-metasomatism after sediment
deposition and link this with the formation of the gold,
pyrite and uraninite during hydrothermal infiltration
(Barnicoat et al., 1997). Others concluded from
geochemical and mineralogical studies of profiles
across stratigraphic units that metamorphism was
essentially isochemical, except for potassium (Sutton
et al., 1990). A further aim of this contribution is,
therefore, to assess this apparent discrepancy by
providing alteration profiles across various siliciclas-
tic Witwatersrand units and discussing the signifi-
cance of trends in the calculated chemical index of
alteration.
Finally, the distribution of redox-sensitive and pH-
sensitive minerals, such as Fe-sulphides versus Fe-
oxides, uraninite verus brannerite or feldspars versus
pyrophyllite will be examined across the various
stratigraphic units of the Witwatersrand Basin. Trends
in the distribution of these minerals along stratigraphic
directions as well as between different lithofacies will
be discussed in terms of their implications for the
palaeoenvironmental conditions. In summary, the
overall goal of this paper is to provide: (1) an up-to-
date review of the evolution of the Witwatersrand
Basin fill; (2) a best-fit genetic model for the world’s
largest gold province and thus also of for the redox-
sensitive monitor phases pyrite and uraninite; and (3)
constraints on the likely Mesoarchaean to Eoproter-
ozoic atmospheric and hydrospheric conditions based
on the spatial and temporal distribution of redox-
sensitive minerals.
2. Geological setting
The Witwatersrand Basin occupies a central
position on the Archaean Kaapvaal Craton (Fig. 2).
The causes of its development and early evolution are
linked with tectonic processes in and around this
craton, whose history is subdivided into two main
periods (de Wit et al., 1992). The first period (3.64–
3.08 Ga) saw the initial formation of the continental
lithosphere of the craton, with a major pulse of
accretion around 3.2 Ga. The second period (3.08–
2.64 Ga) was dominated by the development of
intracontinental basins and likely subduction-related
magmatism along the edge of the Kaapvaal Craton.
By the end of the Archaean Eon, the Kaapvaal Craton
had amalgamated with the Zimbabwe Craton along
the Limpopo Belt (Fig. 2).
2.1. Pre-Witwatersrand basement
The oldest known crustal fragment of the Kaapvaal
Craton is the ~3.64 Ga old Ancient Gneiss Complex
in Swaziland (Kroner and Tegtmeyer, 1994). Products
of Palaeoarchaean crust formation are well preserved
within the Barberton greenstone belt, where domi-
nantly basic–ultrabasic magmatism between 3.49 and
3.42 Ga followed tonalite emplacement between 3.55
and 3.52 Ga (de Ronde and de Wit, 1994). The mafic
to ultramafic rocks have been interpreted as remnants
of ocean-like lithosphere (de Wit et al., 1987). The
upper clastic part of the greenstone belt is interpreted
to comprise synorogenic deposits (Moodies Group)
related to 3.2 Ga accretion. A change from compres-
Fig. 2. Distribution of the main Archaean stratigraphic units of the Kaapvaal Craton. The Witwatersrand Basin fill comprises the West Rand and
Central Rand Groups; also shown is the outline of the three crustal blocks that are believed to have amalgamated by 2.8 Ga to form a single
craton (modified from Schmitz et al., 2004).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–466
sional to transtensional tectonic activity around 3.1 Ga
was accompanied by widespread orogenic gold
mineralisation along late shear zones. This inversion
in the overall stress field marked the beginning of
intracontinental basin formation that eventually led to
the formation of the Witwatersrand Basin.
2.2. Witwatersrand basin development
The first stage of sediment deposition in what
eventually became the Witwatersrand Basin is recog-
nised in the Dominion Group. This group comprises
an up to 2250 m thick bimodal volcanic succession
with a thin basal siliciclastic unit for which a
continental rift basin of unknown extent is assumed.
Maximum and minimum age constraints are given by
U–Pb zircon ages of 3086F3 (the youngest age of
pre-Dominion basement: Robb et al., 1992) and
3074F6 Ma (the age of volcanism: Armstrong et
al., 1991), respectively. The basal siliciclastic unit
includes a conglomerate bed with abundant uraninite
and pyrite but relatively low gold content (Dominion
Reef). Palaeocurrent data consistently point to a
source area to the north or northeast (Frimmel and
Minter, 2002). Thus, a continuation of the original
basin to the south of the present distribution of
Dominion Group rocks is likely (Fig. 3).
The development of the proper Witwatersrand
Basin followed with a hiatus of almost 100 million
years. This great time gap implies that the tectonic
regime for the Dominion Basin is unrelated to that of
the subsequent Witwatersrand Basin. No agreement
Fig. 3. Simplified surface and subsurface geological map of the Witwatersrand Basin, also showing the distribution of Archaean granitoid
domes, the location of the gold fields, major faults and palaeocurrent directions of reefs in the Central Rand Group (from Frimmel and Minter,
2002).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 7
exists on the detailed lithostratigraphic correlation
between the various gold fields across the Witwaters-
rand Basin, but a broad subdivision of the basin fill
into the West Rand and Central Rand Groups, both
constituting the Witwatersrand Supergroup, has long
been recognised (Fig. 4).
2.2.1. West Rand Group
The metasedimentary rocks of the West Rand
Group (Fig. 3) rest with angular unconformity above
the Dominion Group volcanic rocks. The group
attains a maximum thickness of 5150 m in the
Klerksdorp gold field and thins to the northeast. No
information is available from the Welkom gold field.
In the most distal section south of the Vredefort
Dome, about 2000 m of West Rand Group rocks have
been intersected in a borehole (Stevens and Preston,
1999). Sediment input throughout the West Rand
Group was consistently from the north or northeast
(Frimmel and Minter, 2002). Uranium–Pb data from
detrital zircon grains provide a maximum age of
2985F14 Ma for West Rand Group sedimentation
(Kositcin and Krapez, 2004). A minimum age for
most of the group is given by the Crown Formation
lava, the only volcanic unit within the succession, for
which a U–Pb single zircon age of 2914F8 Ma has
been obtained (Armstrong et al., 1991).
Three subgroups are distinguished based on vary-
ing shale/sandstone ratios and basin-wide disconform-
ities (Fig. 4). Within the basal Hospital Hill Subgroup,
the shale/sandstone ratio decreases up-section, with
the sandstone being predominantly quartz arenite,
Fig. 4. Generalised stratigraphic column for the Witwatersrand Supergroup (from Frimmel and Minter, 2002); also shown are the stratigraphic
positions of the main auriferous conglomerate beds (reefs) and their relative significance as gold producers (insert); average gold grade typical of
mined reefs from Frimmel and Minter (2002); SHRIMP U–Pb or Pb–Pb age data from (1) zircon (Armstrong et al., 1991), (2) zircon and
xenotime (Kositcin and Krapez, 2004) and (3) authigenic xenotime (Kositcin et al., 2003); Chemical Index of Alteration from Gartz and
Frimmel (1999), Sutton et al. (1990) and H.E. Frimmel (unpubl. data).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–468
which is interpreted as subtidal deposits (Eriksson et
al., 1981). Evidence of tidal deposits has been
reported from several units. Thick–thin pairs of
siltstone–shale couplets within the upper Coronation
Formation (Fig. 4) are a particularly good example
(Eriksson and Simpson, 2004). Higher up in the
succession, feldspathic sandstone and quartz wacke
become more abundant (Law et al., 1990). The
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 9
sediment record of the West Rand Group reflects
fluctuations between distal fluvio-deltaic and shore-
face to offshore environments, ascribed by some
workers to eustatic sea level changes (Stanistreet
and McCarthy, 1991). Indirect evidence of eustatic sea
level changes comes from the presence of two
diamictite beds within the Government Subgroup,
which are possibly correlatives of the oldest known
glacial deposit on Earth in the Pongola Supergroup
(Young et al., 1998).
Of particular significance is also the presence of
several magnetic shale beds, because they contain
magnetite as principal Fe-phase and not Fe-sulphides
as in the coarser grained units. These shale beds have
been described from the northern and eastern parts of
the basin, as well as from the Klerksdorp gold field
(Tankard et al., 1982), where they occur at a number
of stratigraphic levels throughout the West Rand
Group (Fig. 4). Geophysical maps showing aeromag-
netic anomalies (Corner and Wilshire, 1989) that are
caused by the strong response of up-turned magnetite-
rich shale beds along the basin margin indicate a
basin-wide distribution of this lithotype.
2.2.2. Central Rand Group
The Central Rand Group lies unconformably above
the West Rand Group and reaches a maximum
thickness of 2880 m near the centre of the basin.
Similar to the West Rand Group, a series of cycles can
be distinguished, each of which comprises fluvially
dominated coarse-grained siliciclastic metasedimen-
tary rocks above an erosion surface, but in contrast to
the West Rand Group, shale (mudstone) plays a
subordinate role. Fluvio-deltaic processes dominated
sediment deposition, but tidal reworking has been
suggested for the interfaces between fluvial and
shallow marine systems (Els, 1998). A particularly
fine example of siltstone-shale couplets, that have
been interpreted as dominant and subordinate semi-
diurnal tidal currents, respectively (Eriksson and
Simpson, 2004), occur in the uppermost Randfontein
Formation (Fig. 4). During deposition of the Central
Rand Group, the palaeoslope direction changed from
a consistent dip to the south or southwest to east and
northeast along the western and southwestern margins
and to the southeast and south along the northwestern
and northern margins (Minter and Loen, 1991). Two
minor, locally developed, amygdaloidal basalt units
(Bird lava) in the Krugersdorp Formation provide the
only evidence of volcanism in the Central Rand
Group.
The Central Rand Group is subdivided into the
lower Johannesburg and the upper Turffontein Sub-
groups. The maximum age of deposition for the
sediments of the group is constrained by the youngest
age obtained on detrital zircon from the bottom of the
Johannesburg Subgroup (i.e. 2902F13 Ma: Kositcin
and Krapez, 2004). The youngest detrital zircon age
of 2872F6 Ma from the Krugersdorp Formation sets
the best available constraint for the age of the top of
this subgroup, whereas 2849F18 Ma based on detrital
zircon, or 2840F3 Ma based on detrital xenotime,
represents the maximum age for the top of the Central
Rand Group (Kositcin and Krapez, 2004). A mini-
mum age constraint on deposition of the Central Rand
Group is 2780F3 Ma, which is the oldest age
obtained on any authigenic mineral (xenotime) to
date (Kositcin et al., 2003).
2.2.3. Tectonic setting of the Witwatersrand Basin
General agreement seems to exist on the lower
part of the West Rand Group having been deposited
in a passive continental margin setting, facing an
open ocean to the south. Continental rift-related
rhyolite at the bottom of the Pongola Supergroup
(Nsuze Formation, Fig. 1), considered a correlative
of the Witwatersrand Supergroup (Beukes and
Cairncross, 1991), has an age of 2985F11 Ma
(Hegner et al., 1994), equivalent to that of the lower
West Rand Group. The inferred passive margin
setting is therefore explained by post-rift thermal
subsidence.
A change from upward-deepening to upward-
shallowing at the Hospital Hill/Government Subgroup
boundary has been used to suggest a change from a
passive margin to a foreland basin setting (Coward et
al., 1995). Others prefer a passive margin setting for
the entire group (de Wit et al., 1992). In a recent
SHRIMP detrital zircon study, a general decrease in
the complexity of zircon provenance age spectra is
recorded up-section through the West Rand Group
(Kositcin and Krapez, 2004). This is interpreted as
reflecting an increasing maturity of the hinterland as is
expected for an evolving passive margin, thus
supporting a thermal subsidence origin for the entire
West Rand Group.
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4610
The variation in palaeoslope directions in the lower
Central Rand Group, together with the increase in
continental to marine sediment ratio up-section,
indicates a change to a progressively shrinking
continental basin. This is well reflected by the detrital
zircon age spectra, which increase in complexity up-
section through the Central Rand Group. A progres-
sively greater variety of source rocks must have been
eroded through Central Rand Group times, reflecting
tectonic loading and thus a foreland basin setting for
this group. The same study also suggested that
granitoids were contemporaneously unroofed. Such
a provenance is more consistent with a back-arc fold-
thrust belt than with a foreland fold-thrust belt, which
led Kositcin and Krapez (2004) to postulate a retroarc
basin setting for the Central Rand Group.
With considerably more and better age data
available, more reliable integrated sedimentation rates
can be calculated. Interestingly, these indicate a
minimum of 63 m/my for the West Rand Group,
considerably higher than 18 m/my calculated for the
Central Rand Group. Such a difference seems to be
inconsistent with a simple retroarc foreland basin and
would point to a significant additional strike-slip
component, as suggested by Maynard and Klein
(1995). Stanistreet and McCarthy (1991) have already
emphasised both sinistral and dextral strike-slip along
the northern and western margins, respectively, and
postulated a southeast-directed tectonic escape basin.
However, some of the faults referred to by these
authors are post-Witwatersrand in age. Alternatively,
the lower integrated sedimentation rate calculated for
the Central Rand Group might simply reflect a higher
degree of sediment re-working, and thus less accom-
modation space, and not necessarily lower actual
sedimentation rates.
By analogy with the Pongola Supergroup, a major
folding event in upper Central Rand Group times
between 2837F5 and 2824F6 Ma, defined by pre-
and post-folding felsic intrusive rocks, has been
suggested (Gutzmer et al., 1999). Along the western
basin margin, the angles of unconformity increase up-
section in the Central Rand Group, which reflects
progressive uplift of the hinterland to the west
(Frimmel and Minter, 2002). From the above tectonic
synthesis, it becomes apparent that the generally used,
deeply entrenched term bWitwatersrand BasinQ can be
misleading, because it embraces at least two different
basin types, with a major sequence boundary between
the West Rand and Central Rand Groups. Therefore
the Witwatersrand Basin is best described as a
successor basin that contains erosional remnants of
tectonically stacked sediments originally deposited in
at least two fundamentally different tectonic settings.
The two mafic volcanic units within the Witwa-
tersrand Supergroup remain somewhat enigmatic, as
neither passive margins nor foreland basins are
particularly favourable settings for such volcanism.
Basaltic volcanism on passive margins can be caused
by reactivation of fundamental faults as a far-field
response to changes in the rate or direction of plate
tectonic processes or plume magmatism. By implica-
tion the Crown Formation lavas might record a
change from thermal to reactivated subsidence. In
contrast, the Bird volcanic interval in the Upper
Central Rand Group could be impactogenic, implying
a change from a subduction- to collision-related,
retroarc foreland domain (Stanistreet and McCarthy,
1991).
Based on the palaeoslope directions, the main
tectonic domains in the hinterland controlling the style
of sedimentation in the Witwatersrand Basin are to the
north and west, although palaeoslope directions in the
Evander gold field are more locally controlled. Along
the northern margin of the Mesoarchean Kaapvaal
Craton, east- to northeast-trending greenstone belts
(Murchison, Pietersberg and Giyani greenstone belts,
Fig. 2), surrounded by granitoids, yielded single
zircon age data that overlap with the age of
Witwatersrand sediment deposition (Brandl et al.,
1996; Kroner et al., 2000; Poujol, 2001; Poujol and
Robb, 1999; Poujol et al., 1996). All of these belts
seem to have a 3.2–3.3 Ga basement. In the
Murchison granitoid–greenstone terrain, metamor-
phosed granite, tonalite and rhyolite are dated between
3.02 and 3.09 Ga, overlapping with deposition of the
Dominion Group rocks. Felsic volcanic rocks and
granite from the same terrain, dated between 2971 and
2966 Ma, might reflect crustal thinning related to
early Witwatersrand rifting. Younger felsic intrusions
in the Giyani and Pietersberg granitoid–greenstone
terrains are dated at 2874 Ma, whereas, in the
Murchison terrain, granitoid emplacement is indicated
around 2901 and 2820 Ma. Shear-zone hosted gold
deposits occur in the mafic to ultramafic rocks of
these greenstone belts. Only the Pietersberg Belt hosts
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 11
a sequence of predominantly coarse-grained silici-
clastic metasedimentary rocks with minor palaeo-
placer gold occurrences (Uitkyk Formation) that are
possibly correlatives of the Witwatersrand Supergroup
(de Wit et al., 1992).
A series of intrusive events has also been identified
in the Amalia–Kraaipan granitoid–greenstone terrain
along the western margin of the craton (Anhaeusser
and Walraven, 1999; Poujol et al., 2002; Robb et al.,
1992; Schmitz et al., 2004). A minimum age for the
supracrustal successions is given by a U–Pb zircon
age of 3033F1 Ma obtained for a mafic intrusive
body. Ages between 2932 and 2926 Ma for meta-
morphic and anatectic zircon date both the time of
accretion of a ca. 2930 Ma volcanic arc and
continental collision between the so-called Kimberley
and Witwatersrand crustal blocks (Schmitz et al.,
2004). Deep seismic reflection profiles through the
west-central Kaapvaal Craton indicate a westward-
dipping subduction (de Wit and Tinker, 2004).
Subsequent crustal thickening and uplift led to
variable exhumation of, and decompression melting
in, the Kimberley block, with resulting high-level
granitoids as young as 2727 Ma, while the Witwa-
tersrand block underwent subsidence with sedimenta-
tion in the Central Rand Basin.
The age of the youngest granitoids in the Kraaipan
Belt (around 2790 Ma) is similar to the age of the
Gaborone Suite rapakivi-type granite and related
volcanic rocks in southern Botswana (2783 Ma;
Grobler and Walraven, 1993; Moore et al., 1993).
Considering the inferred retroarc foreland setting and
above age constraints for the Central Rand Group, the
younger granitoids in the hinterland to the north,
northwest and west might reflect the corresponding
magmatic arc, possibly related to a southward-dipping
subduction zone. The ocean that was closing at that
time probably separated the amalgamated Witwaters-
rand–Kimberley block from the Pietersburg block
(Fig. 2). Such a combination of foreland basin, related
to westward subduction, and retroarc basin, related to
the closure of an ocean further north, would help
explain the apparent inconsistency between integrated
sedimentation rates and a simple retroarc basin
mentioned above. The Limpopo orogeny, often used
as an explanation for the inferred foreland position of
the Central Rand Basin in the past (e.g. Burke et al.,
1986), took place more than 100 million years after
the Witwatersrand Supergroup rocks had been laid
down. Apart from great ambiguity about the existence
of a Himalayan-style Limpopo orogeny, the available
age data rule out any relationship between tectonic
events that shaped the Limpopo Belt and the
Witwatersrand Basin.
2.3. Post-Witwatersrand evolution of the Kaapvaal
Craton
Mild deformation in the form of block faulting and
subordinate folding and thrusting along the western
and northern basin margin, and low-grade regional
metamorphism of the Witwatersrand Basin fill, are
testimony to post-depositional alteration related to a
series of tectono-thermal events initiated during, and
succeeding, Witwatersrand sediment deposition.
Stratigraphically above the Witwatersrand Super-
group follows the Ventersdorp Supergroup, which
attains a thickness of more than 3600 m in places. In
the northern parts of the basin, the two supergroups
are separated by a regional angular unconformity that
is overlain by a thin fluvial conglomerate (Venterspost
Conglomerate Formation). In places, the conglomer-
ates are highly auriferous and form important ore
bodies (Ventersdorp Contact Reef) that are similar to
other Witwatersrand reefs. Further south, in the
Welkom gold field, the Ventersdorp Supergroup rests
paraconformably above the Witwatersrand Super-
group and the Ventersdorp Contact Reef is not
developed (Minter et al., 1986). This is explained by
a lack of re-working of the older, auriferous Witwa-
tersrand sediments along that contact there.
The thin siliciclastic sequence of the Venterspost
Conglomerate Formation is conformably overlain by a
thick succession of flood basalt (Klipriviersberg
Group), though localised erosional scours exist and
have been explained by contemporaneous eruption of
channelised lava flows over unconsolidated sediment
(Hall, 1997). A U–Pb zircon age of 2714F8 Ma is so
far the best constraint on the timing of eruption
(Armstrong et al., 1991), which implies a hiatus of at
least 50 million years for the underlying unconform-
ity. This is supported by the youngest detrital
xenotime age obtained on the Ventersdorp Contact
Reef, i.e. 2729F19 Ma (Kositcin and Krapez, 2004).
The Klipriviersberg Group basalt layers are over-
lain by siliciclastic and volcanic, predominantly felsic
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4612
and minor mafic, deposits of the Platberg Group. The
siliciclastic sediments reflect alluvial fan deposits that
prograded into lacustrine environments. A U–Pb
zircon age of 2709F4 Ma for Platberg Group basalt
(Armstrong et al., 1991) shows that the bimodal
volcanic activity recorded in the Ventersdorp Super-
group was short-lived. Deposition in a continental rift
that evolved into an Atlantic-type continental margin
is inferred from seismic profiles that indicate consid-
erable thickening of the Ventersdorp Supergroup
towards the west of the craton (Tinker et al., 2002).
A link between Ventersdorp rifting and collision
between the Kaapvaal and Zimbabwe Cratons, as
suggested previously (Burke et al., 1985), is unlikely.
Granulite facies metamorphism associated with south-
ward thrusting of the Southern Marginal Zone of the
Limpopo Belt on to the Kaapvaal Craton is dated at
2691F7 Ma Ma (Kreissig et al., 2001) and syn-
tectonic granites from the Central Zone of that belt
range in age from 2664 to 2572 Ma (McCourt and
Armstrong, 1998), later than onset of Ventersdorp
flood basalt extrusion. However, the age of the
Klipriviersberg Group overlaps with an earlier north-
ward thrusting phase, dated at 2729F19 Ma (Passer-
aub et al., 1999), that affected greenstone belts along
the northern Kaapvaal Craton. The Klipriviersberg
extension could thus be explained by southward-
directed subduction beneath the Kaapvaal Craton
prior to continental collision. Alternatively, the
ultimate cause of Ventersdorp rifting might not reside
in crustal plate tectonics but as a result of a mantle
plume (Hatton, 1995).
A second thrust event in the Witwatersrand Basin,
post-Platberg and pre-Transvaal in age (Roering,
1990), is most likely a distant expression of the
collisional tectonic processes in the Southern Mar-
ginal Zone of the Limpopo Belt. At that stage, lower
greenschist facies metamorphic conditions were
attained for the first time in the Witwatersrand rock
column, at least along the northern margin of the basin
(Coetzee et al., 1995). This phase of compression and
uplift was followed by regional peneplanation. The
corresponding unconformity represents a major
sequence boundary, separating the Ventersdorp from
the overlying Transvaal Supergroup.
Of significance to the debate on the genesis of the
Witwatersrand gold and of Archaean atmospheric
evolution is the presence of a thin, but laterally
extensive, basal conglomerate and sandstone unit
(Black Reef Quartzite Formation) that rests on the
post-Ventersdorp erosion surface. This basal silici-
clastic unit contains a conglomerate (Black Reef) that
is almost indistinguishable from the auriferous,
uraniferous and pyrite-rich Witwatersrand reefs,
except for a lower metamorphic grade. A syn-
Witwatersrand or syn-Ventersdorp age for the Black
Reef, as suggested by Phillips et al. (1997) is
untenable, because of field relationships (Els et al.,
1995) and geochronological data. The Black Reef
Quartzite Formation reflects fluvial sedimentation in
channels on a locally deeply incised palaeosurface,
braided fluvial and braid delta deposits that grade into
shallow marine deposits. A locally transitional and
conformable upper contact with the overlying pre-
dominantly chemical, marine sedimentary successions
of the Chuniespoort Group is unequivocal evidence of
an early Transvaal Supergroup age for this genetically
important auriferous formation (SACS, 1980). An
indirect age constraint for the Black Reef Quartzite
Formation exists in the form of a Pb–Pb single zircon
age of 2642F2 Ma (Walraven and Martini, 1995),
obtained for a lava in a stratigraphic correlative
(Vryberg Formation). In agreement with the above
timing of tectonic activity in the Limpopo Belt, the
sediment sources were located to the north (Barton
and Hallbauer, 1996; Els et al., 1995).
The overlying Chuniespoort Group starts with a
transgressive black shale, followed by thick platform
carbonates (Malmani Subgroup), banded iron forma-
tion (Penge Formation) and, erosive into the older
strata and reflecting a eustatic sea level fall, lacustrine
deposits including a glaciogenic diamictite (Duitsch-
land Formation). An upper age constraint is given by a
Pb–Pb single zircon age of 2550F3 Ma for a tuff layer
in the lowermost formation of the Malmani Subgroup
(Walraven and Martini, 1995), whereas U–Pb single
zircon ages of 2432F31, 2465F7 and 2480F6 Ma
reported for the stratigraphic equivalent of the Penge
Formation (Nelson et al., 1999; Walraven and Martini,
1995) set some constraints on the minimum age of the
group. Subsidence rates calculated from the thickness
of the group are in good agreement with thermal
subsidence due to lithospheric cooling after the
Ventersdorp thermal anomaly (Tinker et al., 2002).
A further major sequence boundary separates the
Chuniespoort Group from the overlying Pretoria
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 13
Group. These two groups, though unified as Transvaal
Supergroup in the literature, have little in common.
The boundary is marked by a strongly weathered
erosional unconformity reflecting a hiatus of ~80
million years, above which lies a volcano-sedimentary
succession with alluvial fan and fan delta deposits
with minor marine influence. This hiatus is likely to
reflect a glacio-eustatic sea level fall, because the
unconformity is overlain by glaciogenic diamictite.
Alluvial to lacustrine deposits grade into shallow to
deep marine deposits before a further erosional
unconformity at the top of the Timeball Hill For-
mation marks another sequence boundary. The uncon-
formity above that formation contains the paleosol
examples that have been used to constrain the rise in
atmospheric O2 during the Eoproterozoic (Beukes
et al., 2002).
Deposition of the Pretoria Group was followed
after another hiatus by the contemporaneous extrusion
of felsic volcanic rocks (Rooiberg Group), and
emplacement of the mafic to ultramafic 2059F1 Ma
(Buick et al., 2001) Rustenberg Suite and 2054F2 Ma
(Walraven and Hattingh, 1993) granitic Lebowa Suite,
both of the Bushveld Igneous Complex, to the north
of the Witwatersrand Basin. An elevated crustal
geotherm related to large-scale magmatic underplating
and intraplating also probably resulted in the intrusion
of alkali granite and mafic plutons within the
Witwatersrand Basin, some dated at 2078F12 Ma
(Moser, 1997). The last major disturbance experi-
enced by the Witwatersrand strata is reflected by a 30-
km-wide domal structure, the Vredefort Dome (Fig.
3), that could well represent the oldest (2023F2 Ma;
Kamo et al., 1996; Moser, 1997) and largest (250 to
280 km diameter; Henkel and Reimold, 1998) known
terrestrial impact structure.
2.4. Post-depositional alteration of the Witwatersrand
Supergroup
The multitude of tectono-thermal events that
affected the Kaapvaal Craton after the deposition of
the Witwatersrand Supergroup sediments is reflected
by recrystallisation, formation of metamorphic min-
eral assemblages, and likely also metasomatic reac-
tions, thus masking geochemical and mineralogical
characteristics of the sedimentary protolith. A series of
detailed petrological, geochronological and fluid
inclusion studies made it possible to distinguish
between several stages of post-depositional alteration
throughout the basin.
Initial fluid–rock interaction most probably
involved leaching by basin-wide penetration of
meteoric water shortly after sediment deposition. A
prime Phanerozoic analogue of this process has been
documented for the Parana Basin of Brazil (Franca et
al., 2003) and a similar scenario is likely to have
affected the Witwatesrand Basin fill (Phillips et al.,
1990), when uplift of at least one basin margin
provided a steep hydraulic gradient for groundwater to
flow towards the basin centre. A decrease in pH with
increasing burial is predicted from the maturation of
organic material and the release of H+ from dehy-
dration reactions. This, in turn, could have led to the
dissolution of detrital feldspars, thus generating a
secondary porosity for further diagenetic fluid flow.
Following diagenesis, dated between 2776 and
2780 Ma (Kositcin and Krapez, 2004), a first stage of
lower greenschist facies metamorphism was attained
along the northern basin margin, coeval with high-
grade metamorphism and tectonism in the Southern
Marginal Zone of the Limpopo Belt. In most parts of
the basin, regional peak metamorphic conditions of
300 to 350 8C at 3 kbar were likely achieved during
deposition of the Pretoria Group (Frimmel and Minter,
2002). Only around the Vredefort Dome were
metamorphic grades up to amphibolite facies reached
(Gibson and Wallmach, 1995). There the peak of
metamorphism is ascribed to the emplacement of the
Bushveld Igneous Complex, which is supported by an
anomalously high geothermal gradient (Frimmel,
1997).
Several stages of enhanced fluid circulation
through the Witwatersrand Basin are recognised.
These range from diagenetic basin dewatering to a
number of hydrothermal fluid infiltration events.
Hydrothermal rutile, zircon and xenotime age data
are considered most robust, whereas Rb–Sr, U–Pb and
Pb–Pb ages of various hydrothermal precipitates have
only limited geochronological meaning (Zartman and
Frimmel, 1999). The most reliable of these age data
(Armstrong et al., 1995; Kositcin and Krapez, 2004;
Robb et al., 1990) cluster around 2720 Ma (Ven-
tersdorp extension), 2580 Ma (early Transvaal thermal
subsidence), possibly around 2200 Ma (Pretoria
extension), and 2060 Ma (Bushveld event). A further
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4614
hydrothermal infiltration event has been ascribed to
the Vredefort impact, based on cross-cutting relation-
ships with dated pseudotachylyte, fluid inclusion and
gold chemistry data (Frimmel et al., 1999). Although
there is some evidence for post-2.0 Ga hydrothermal
alteration from K–Ar age spectra on very fine-grained
white mica (Zhao et al., 1999), none of these younger
events, which are far field effects of various stages of
accretion along the craton margin, are of great
significance for the post-depositional alteration his-
tory of the Witwatersrand sediments.
3. Sedimentological and mineralogical
characteristics of the orebodies
3.1. Host lithofacies
The Witwatersrand gold orebodies typically occur
in conglomerate beds (reefs) on unconformities within
major stratigraphic sequences (Pretorius, 1981). Gold
has been mined from at least 30 such reefs, the most
important of which are shown in Fig. 4, with those
from the Central Rand Group accounting for 90% of
total production. In addition, the base of the younger
Ventersdorp and Transvaal Supergroups also host, in
places, economic to subeconomic, orebodies of
comparable lithology and mineralogy. The orebodies
comprise various fluvial lithofacies that range from
clast-supported oligomictic conglomerate to loosely
packed conglomerate, pebbly arenite, or just pebble
lag surfaces associated with trough cross-bedded
quartz arenite. In exceptional circumstances, ore is
associated with immature debris flow lithofacies.
These include black argillite in the Beatrix Reef
(Minter et al., 1988), where overlying playa facies
covered eroded older ores that occurred as eroded
subcrops, and risers between terraces of the Venters-
dorp Contact Reef (Henning et al., 1994), where
Fig. 5. (A) Oligomictic pebbly quartz arenite (reef), Vaal Reef, Stilfontein
unconformity and 3 cm above the base on a bedding plane defining the to
ventifact; scale bar=1 cm. (B) same as A, but highlighting the position of
pyrite pebble lags between quartz pebble lags, Basal Reef, Welkom gold
Ventersdorp Contact Reef, Tau Lekoa mine, Klerksdorp gold field, showi
contact with Klipriviersberg Group mafic lava flow is erosive into originally
rounded, oolitic pyrite (arrow) within pyrite pebble lag shown in C (white r
cross beds in pebbly quartz arenite at the base of the Basal Reef, Welkom g
(see Minter et al., 1993); scale bar=1 cm.
underlying immature sedimentary rocks were eroded
and incorporated into the alluvium during entrench-
ment.
The Witwatersrand reef rocks are generally
described as quartz pebble conglomerates. Exceptions
exist, however, such as the Steyn Reef in the Welkom
gold field, which is polymictic with as much as 30
vol.% yellow quartz porphyry in both the pebble and
sand fractions. The Kimberley Reef in the Evander
gold field contains as much as 50% chert pebbles
(Tweedie, 1986). The average pebble assemblage is
85 vol.% vein quartz, 12 vol.% chert, 2 vol.% quartz
porphyry, and 1 vol.% metamorphic clasts (Frimmel
and Minter, 2002).
The orebodies range in thickness from decimetres
to a few metres and are confined between a basal
degradation surface, usually an angular unconformity,
and an upper planar bedding surface that separates it
from overlying quartz wacke or siltstone (Fig. 5A,B;
Minter, 1991). They have a lenslike geometry and
define fluvial bar-and-channel bedforms that display
unimodal palaeocurrent directions (Smith and Minter,
1980). Multichannel sequences of conglomerate and
quartz arenite, deposited by repeated flood- and
waning-stage flows, make up the thicker orebodies
(Fig. 5D).
Depositional environments that have been recon-
structed from the geometry of the basal degradation
surface, coarseness of the sediment, and the distribu-
tion of lithofacies, range from proximal alluvial fan
(e.g. EA Reefs in the Eldorado Formation), to terraced
fluvial deposits (e.g. Ventersdorp Contact Reef;
Henning et al., 1994) to braid plains (Composite
Reef; Tucker, 1980), to braid deltas that merge with
shoreline environments. In some places, aeolian
deflation has been suggested to account for the planar
top of the orebodies (Minter, 1999), whereas in others,
shoreline encroachment and associated wave action,
followed by burial beneath fine-grained sediment is
mine, Klerksdorp gold field; note the veneer of bitumen on the basal
p of pebbly layer; cross-bedding (so) in hanging wall, and an in-situ
fine- to coarse-grained rounded pyrite. (C) Upward fining imbricate
field; scale bar=1 cm. (D) Underground exposure of upper half of
ng multichannel sequences of conglomerate and quartz arenite; top
unconsolidated conglomerate; scale bar=50 cm. (E) Truncated, well
ectangle), scale bar=0.5 cm. (F) Concentration of gold particles along
old field. Three foresets converge with the bottomset toward the right
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 15
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4616
indicated as in the Basal Reef (Buck, 1983) and
Carbon Leader Reef (Buck and Minter, 1985).
3.2. Textural, chemical and isotopic features of the
main ore components
The bulk of gold, uraninite and pyrite, whose
origin is debated, occur together with undoubted
allogenic detrital minerals, such as mechanically
abraded zircon and chromite grains, and more rarely
PGE-minerals and diamond (for a complete list of
minerals identified to date in Witwatesrand reefs see
Phillips and Law, 2000), on degradation surfaces
marked by pebble lags or the base of clast-supported
conglomerate beds, cross-bedded foresets, bottomsets,
and coset boundaries (Fig. 5F). In thick (1–2 m)
conglomerate units, representing multichannel
sequences, the allogenic minerals are concentrated
along the basal degradation surface of each graded
bed. The highest concentrations of allogenic minerals
are found above unconformity surfaces, which reflect
a direct relationship to the amount of erosion. Cut-
and-fill structures, such a trough cross beds, contain
greater allogenic mineral concentrations than aggrada-
tional structures, such as planar cross beds. Experi-
ments (James and Minter, 1999) have confirmed that
the dominant concentration mechanism was by
selective entrainment of coarser, less dense particles
during bedload transport (Slingerland and Smith,
1986).
Average gold grades of mined reefs range from 3 to
25 g/t (Fig. 4). Case studies on the element
distribution within reefs, such as the Kimberley Reef
(Rasmussen and Fesq, 1973), the Steyn Reef (Frim-
mel and Minter, 2002), the Vaal Reef (Fox, 2002), and
the Ventersdorp Contact Reef (H.E. Frimmel, unpubl.
data) illustrate a relatively good positive correlation
between U and Au, but only poor correlation between
Au and Zr, and Cr—elements that are controlled by
detrital phases (zircon, chromite) that were not
susceptible to hydrothermal mobilisation. All data
sets show gold enrichment being linked with Zr
enrichment, but samples rich in Zr typically do not
contain elevated Au contents. A detailed study of
different lithofacies in a channel sidebar of the Steyn
Reef, Welkom gold field, where 13 aggradation events
can be separated (Frimmel and Minter, 2002) showed
that degradation surfaces were preferentially mineral-
ised with an average of 38 ppm Au, 410 ppm Zr, 1750
ppm U, and 300 ppm Cr. Higher concentrations of all
these elements were recorded in the coarser sediment
fraction, reflecting higher flow velocities. Similar
observations were also made quantitatively on other
reefs, such as the Leader Reef (Smith and Minter,
1980) and the Ventersdorp Contact Reef (H.E.
Frimmel, unpubl. data).
The Witwatersrand reefs are not only extremely
rich in gold but also are one of the world’s largest
uranium depositories. Between 1952 and 1975, as
much as 1.5�106 t U3O8 (Frimmel et al., in press) was
produced at an average grade of 271 ppm (Camisani-
Calzolari et al., 1984). One of the richest reefs was the
Monarch Reef, a small-pebble distal placer, mined at a
mean grade of 2860 ppm U3O8. The principal U-
minerals are uraninite, brannerite and leucoxene, with
a systematic decrease in uraninite/brannerite ratio up-
section. For example, in the Welkom gold field, the
uraninite/brannerite ratio is 8.7 in the Steyn Reef,
whereas in the younger Beatrix Reef it is zero (Minter
et al., 1988). Similarly, all the U in the Black Reef,
from which a specimen with as much as 3350 ppm
U3O8 has been reported (Bourret, 1975), occurs as
brannerite.
There is a broad systematic trend in the U/Au ratio
up-section. In contrast to the Central Rand Group
reefs, the Dominion Reef is highly uraniferous but
does not contain significant amounts of gold. Within a
given stratigraphic unit, both Au and U concentrations
decrease from the basin margin towards its centre, but
at different rates. A systematic increase in the U/Au
ratio down the paleoslope from 10�3 to 10 was noted
in the Welkom gold field, with uraninite being
enriched in the more distal facies, probably as a result
of hydraulic mineral sorting (Minter et al., 1986). This
is also illustrated by the coarse-grained Main Reef,
which has an average pebble size of 37 mm and a U/
Au ratio of less than 25 as opposed to the relatively
finer grained, more distal Monarch Reef, whose
average pebble size is 16 mm and U/Au ratio is 128
(Vennemann et al., 1995).
It has been noted in many studies that Witwaters-
rand gold is intimately associated with carbonaceous
matter, which occurs as stratiform seams and as
spherical, glassy globules. Both forms have been
recognised as metamorphosed solidified hydrocar-
bons, i.e. pyrobitumen (Gray et al., 1998). The
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 17
bcarbon seamsQ occur preferentially in deposits
reflecting distal environments, but are conspicuously
absent from proximal, high-energy deposits. Hydro-
carbon derivation from original algal mats is sug-
gested by the distribution of the bcarbon seamsQ onpalaeosurfaces, on sedimentary accumulation surfaces
and on trough cross-beds and ripple surfaces (Buck,
1983; Minter, 1981). In those reefs that contain such
carbonaceous matter, such as the Carbon Leader,
Basal and Vaal Reefs, the highest Au and U
concentrations are located in the zones that are
particularly rich in pyrobitumen. Nagy (1993) esti-
mated that about 40% of all mined Witwatersrand
gold was hosted by such bitumen seams. On a micro-
scale, the pyrobitumen-filled microfractures typically
contain some gold. Oil-migration, and thus by
implication gold transport, has been suggested both
prior (England et al., 2002a) and during (Jolley et al.,
2004) fracturing. It should be noted, however, that
many highly auriferous reefs contain little or no
carbonaceous matter; a number of economic reefs,
such as the Main and Kimberley reefs in the Central
and West Rand gold fields do not contain any
noteworthy amounts of bitumen.
The bitumen is interpreted to have formed by the
polymerisation and crosslinking of liquid hydrocar-
bons around irradiating grains, predominantly uranin-
ite, in the host sedimentary rock (Schidlowski, 1981).
Oil-bearing fluid inclusions provide direct evidence of
oil migration through the Witwatersrand sedimentary
rocks (Drennan et al., 1999; England et al., 2002a).
Bitumen derivation from a variety of biomass in a
reducing environment, with subsequent short-range
hydrothermal mobilisation is indicated by organo-
geochemical, bulk and molecular C isotopic studies
(Spangenberg and Frimmel, 2001).
The phases associated with the gold, which are
most important for the topic of this paper are pyrite
and uraninite. The textural, mineral chemical and
isotopic characteristics of these phases will, therefore,
be discussed in somewhat greater detail below.
3.2.1. Pyrite
Pyrite is the most common heavy mineral in all of
the fluvial deposits of the Witwatersrand. Only in
marine sedimentary rocks, such as shales in the West
Rand Group, are Fe-oxides (predominantly magnetite)
found instead of pyrite (Frimmel, 1996). Locally, such
as in parts of the Ventersdorp Contact Reef, Klerks-
dorp gold field, pyrrhotite is the stable Fe-sulphide
instead of pyrite. There the distribution of the two Fe-
sulphides seems to be controlled by locally variable
oxygen fugacity of post-depositional fluids. Gener-
ally, the pyrite occurs in a number of different textural
forms (England et al., 2002b; Hallbauer, 1986;
Ramdohr, 1958) that are grouped into (1) rounded,
compact, (2) rounded, porous, and (3) euhedral.
The rounded compact variety is by far the most
abundant form of pyrite in all reefs (Fig. 5C) except
for the Ventersdorp Contact Reef, where the pyrite
grains are predominantly euhedral. However, etching
of these euhedral grains reveals that most have one or
more rounded cores, with the euhedral outline being
an artifact of secondary, authigenic/hydrothermal
overgrowth around pre-existing rounded, compact
pyrite cores (Fig. 6I; England et al., 2002b; Frimmel
and Minter, 2002). Evidence of mechanical abrasion
of the rounded grains is given by truncation of
oscillatory growth zonation, defined by variable As
contents at grain boundaries (McLean and Fleet,
1989). A crystallographic study (Fleet, 1998)
revealed that some of these rounded grains are single
crystals and not the polycrystalline or twinned
crystals one would expect if they were pseudomorphs
after bblack sandsQ as suggested by Phillips and
Myers (1989) and Phillips and Law (2000). A further
argument against basin-wide sulphidation of bblacksandsQ is that the latter typically comprise titanomag-
netite. Pyritisation of such a precursor characteristi-
cally leads to intergrowths of minute rutile needles.
Such rutile-bearing pyrite pseudomorphs are the
exception and not the rule in the Witwatersrand reefs
(Ramdohr, 1958). The rounded pyrite grains from the
Witwatersrand are, however, devoid of Ti. Associated
with this form of pyrite are similarly rounded,
compact arsenopyrite and cobaltite particles (Saager
and Oberthur, 1984; England et al., 2002b). Genet-
ically significant mineral inclusions in rounded pyrite
are feldspar (Fig. 6F), calcite, corundum and spes-
sartine, which are absent in the metamorphic mineral
assemblage of the metasedimentary host rocks.
Examples of several centimetre-thick, almost mono-
mineralic, fining-upward pyrite beds, displaying
imbrication with the same orientation as intercalated
quartz pebble lags, occur within the Basal Reef of the
Welkom gold field (Fig. 5C).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4618
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 19
The porous pyrite displays a variety of internal
textures, ranging from laminated aggregates, and
rounded concretions to oolitic-colloform and dendritic
forms. Many concretionary and colloform varieties
are fragmented, broken and have their internal
structures truncated, from which mechanical transport
is inferred (Fig. 5E).
Post-sedimentary, euhedral to subhedral pyrite
occurs preferentially adjacent to zones of hydro-
thermal alteration, such as veins and faults. This form
is typically associated with other authigenic/hydro-
thermal sulphides (chalcopyrite, cobaltite–gersdorf-
fite, pyrrhotite, galena and arsenopyrite) and
pyrobitumen (Gartz and Frimmel, 1999). Euhedral
pyrite overgrowths are common and, in places, are
contiguous with pyrite that fills fracture or pore
spaces.
A laser-ablation sulphur isotope study (England et
al., 2002b) revealed that the rounded pyrite forms
have a wide range in d34S values (�5.0x to +6.7x),
not only at the mine and stope-face scale but even at
the sample scale over less than 1.5 cm2. In a previous
SHRIMP study (Eldrige et al., 1993), large hetero-
geneities in d34S (�7x to +32x) were noted in
single pyrite grains and between different morpho-
logical types. Such heterogeneity is difficult to
reconcile with precipitation from a geochemically
homogeneous hydrothermal fluid, and more likely
reflects variation in pyrite from the eroded source rock
and/or microbial sulphate reduction in the depositio-
nal environment. The heterogeneity in rounded pyrite
is in contrast with the narrow range in d34S values
obtained for authigenic/euhedral pyrite (�0.5x to
+2.5x), which is also distinguished from rounded
pyrite by higher Ni and As contents as well as gold
inclusions. Of particular significance are two textur-
ally adjacent but isotopically contrasting ooid-like
Fig. 6. Photomicrographs illustrating morphological and textural features o
light, scale bars=0.2 mm, except for scanning electron microscope image
particles occurring together on a mm-scale. (B) Gold micro-nugget (Au) w
shown in A. (C) Spheroidal gold micro-nugget (Au I) with secondary,
secondary gold filling a fracture within a detrital zircon grain that pierce
overfolded rims next to rounded pyrite; all of the above from the same han
Sericite pseudomorphs after K-feldspar (Fsp–Psm) as inclusions within rou
Rand gold field; matrix silicates are chloritoid (Ctd) and muscovite (M
fragment and quartz pebble, Ventersdorp Contact Reef, Klerksdorp gold f
II), B-Reef, Free State Geduld mine, Welkom gold field. (I) Euhedral,
Ventersdorp Contact Reef, Klerksdorp gold field.
pyrite grains from the Ventersdorp Contact Reef,
which display a strong isotopic zonation but of
opposite signs (England et al., 2002b). The d34S
ratios in one of the two grains increase systematically
from �4.1x in the core to +1.4x in the rim, whereas
those in the other grain decrease from �0.8x in the
core to �4.5x in the rim. The former has been
interpreted by these authors as indicative of sulphi-
dation of an original sulphate grain. The latter grain
reflects a different provenance and highlights that the
isotopic differences must be pre-depositional and
cannot be due to fluctuations in redox potential of a
hydrothermal, sulphidising fluid that mixed with local
meteoric formation water during diagenesis as pro-
posed by Phillips and Law (2000).
Attempts to date the various forms of pyrite using
the U–Th–Pb isotope systems (Barton and Hallbauer,
1996; Poujol et al., 1999; Zartman and Frimmel,
1999) were met with mixed success. Authigenic and
hydrothermal pyrite is typically enriched in urano-
genic Pb, whereas the rounded forms have a less
radiogenic isotopic signature. However, absolute Pb–
Pb ages need to be viewed with caution as the 238U
and 235U decay schemes were likely decoupled,
presumably by the selective diffusion of 222Rn from
uraninite and its subsequent capture in hydrothermal
precipitates, leading to erroneous ages (Zartman and
Frimmel, 1999). More reliable are Re–Os data
obtained on rounded, compact pyrite from the Vaal
Reef, which yielded an age of 2.99F0.11 Ga (Kirk et
al., 2001). This is older than the time of sediment
deposition, which provides a strong argument for the
detrital origin of much of the pyrite.
3.2.2. Uraninite and leucoxene
Most of the uraninite particles are well rounded
and enclosed, or partially replaced, by bitumen, thus
f Witwatersrand gold orebodies (combined transmitted and reflected
s C and D: 0.1 mm): (A) Contrasting morphological types of gold
ith overfolded rims next to rounded pyrite (Py) from same sample as
well crystallised gold overgrowth (Au II). (D) Same as in C, but
s a rounded gold grain. (E) In-situ gold micro-nugget (Au I) with
d specimen of Basal Reef, Welkom gold field, shown in Fig. 5B. (F)
nded pyrite from the Kimberley Reef, South Roodepoort mine, West
us). (G) Hydrothermal gold with chlorite in matrix between lithic
ield. (H) Gold inclusions within secondary, hydrothermal pyrite (Py
secondary pyrite overgrowths around older, rounded cores (Py I),
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4620
explaining the good correlation between bitumen
seams and U content (Minter, 1978). Considering
the relationship between bitumen formation and
locally available radioactivity, at least some of the
uraninite must be older than the hydrocarbons, and
hence detrital or early diagenetic. Secondary, post-
depositional uraninite and other U-bearing minerals,
mainly brannerite and uraniferous leucoxene, are
considered to be products of partial mobilisation of
the earlier, rounded uraninite (England et al.,
2001a). A detrital origin of the rounded uraninite
grains is suggested by their mineral chemistry and
U–Pb geochronology. Even on a thin section scale,
adjacent grains show a great variation in Th/U ratio,
which reflects provenance from a variety of source
rocks. Many uraninite grains are rich in Th (average
3.9 wt.%), which is inconsistent with a low-temper-
ature hydrothermal origin but indicative of granitic
to pegmatitic sources (Feather and Glathaar, 1987;
Grandstaff, 1981). Direct dating of uraninite grains
from the Dominion Reef yielded an U–Pb age of
3050F50 Ma (Rundle and Snelling, 1977). This age
overlaps with that of Dominion Group sedimenta-
tion but is distinctly older than the Witwatersrand
sediments.
Another important U-mineral in the Witwatersrand
assemblages is brannerite with a composition close to
UTi2O6. It is characteristically of secondary origin,
derived from the oxidation of uraninite in the presence
of rutile, which, in turn, is an alteration product
derived from original detrital ilmenite and minor
titanomagnetite particles. Form relics of rounded
ilmenite occur concentrated on all scour surfaces
throughout the Witwatersrand Supergroup. Most of it
is, however, altered to leucoxene that is in many cases
highly uraniferous. Typical TiO2 concentrations in the
siliciclastic metasedimentary rocks are between 0.1
and 1 wt.%. The very fine-grained nature of the
rounded leucoxene pseudomorphs after ilmenite
points towards them being the result from weathering.
This is in contrast to the presence of distinct, euhedral
to subhedral rutile and brannerite grains, both of
which are related to the hydrothermal oxidation of
original ilmenite and uraninite. Direct dating of such
rutile from the West Rand Group (Robb et al., 1990)
confirms such an age relationship as it yielded a U–Pb
age of 2578F34 Ma, which is markedly younger than
the age of sediment deposition.
3.2.3. Gold
It has been noted by many workers that the
Witwatersrand gold appears late in the paragenetic
sequence (Ramdohr, 1958), with most of it occurring
in textural association with bitumen, hydrothermal/
metamorphic chlorite (Gartz and Frimmel, 1999) or
pyrophyllite (Barnicoat et al., 1997), along micro-
fractures (Jolley et al., 2004), and as inclusion within
euhedral, secondary pyrite (Fig. 6H). In contrast to the
latter, rounded pyrite is typically devoid of gold
inclusions. This generation of gold displays either
euhedral crystals or dendritic or otherwise irregularly
shaped habit. In a few instances, however, gold
particles have a completely different morphology. In
contrast to the above, they display rounded, spher-
oidal, disc-like and toroidal forms. Of particular
importance is that both morphological types, the
rounded to torroidal and the dendritic to euhedral,
secondary gold, can occur together on a micro-scale
(Fig. 6A,C,D), i.e. within millimetres in the same thin
section (Minter et al., 1993). This provides strong
evidence for a polyphase gold entrapment history,
with the rounded particles derived by mechanical
(fluvial) transport and secondary gold by precipitation
from a hydrothermal fluid.
Considerable compositional variability with
respect of Au:Ag:Hg ratios in gold particles exists
between reefs, within a given reef, and in some cases
even on a micro-scale within a given thin section
(Frimmel and Gartz, 1997; Reid et al., 1988).
Individual gold particles are, however, homogeneous,
which is readily explained by the diffusion rates of Ag
and Hg through gold at the temperatures to which the
Witwatersrand rocks were subjected during burial and
metamorphism (Frimmel et al., 1993). Only gold
particles in quartz veins that have been ascribed to the
Vredefort impact, based on field relationships with
impact-related pseudotachylite, have internal compo-
sitional variability. This is readily explained by the
short duration of that event that did not permit
diffusional homogenisation (Frimmel and Gartz,
1997).
First attempts to date the gold directly by the Re–
Os method yielded ages that are older than that of
sediment deposition (Kirk et al., 2002). Four gold
samples from the Vaal Reef define an isochron that
corresponds to an age of 3016F110 Ma and, when
combined with rounded pyrite from the same hand-
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 21
sample, define an age of 3033F21 Ma. Exceedingly
high Os concentrations, of as much as 4.16 ppm,
reported for gold from the Vaal Reef (Kirk et al.,
2001), are possibly an artifact of contamination by
minute inclusions of platinum-group element miner-
als. However, a series of subsequent analyses of gold
from the Vaal Reef (Kirk et al., 2002) and from the
Basal Reef, both of toroidal micro-nuggets and
secondary, hydrothermal gold crystals (Frimmel et
al., in press) show consistent Re concentrations
between 4 and 37 ppb and Os concentrations between
2 and 15 ppb. These values are one to four orders of
magnitude greater than those for younger gold
deposits as well as for average continental crust (Kirk
et al., 2002).
3.3. Sediment provenance
Derivation of the pebbles from mainly Archaean
granite and pegmatite (55%) as well as mesothermal
quartz veins and marine chert (45%) is indicated by
oxygen isotope data (Vennemann et al., 1992, 1995).
A granitic and/or pegmatitic source is further indi-
cated by rare grains of detrital cassiterite, molybdenite
and columbite (Feather and Koen, 1975) and by bright
cathodoluminescence of many detrital quartz grains
(Gartz, 1996). Furthermore, the concentrations of
granitophile elements, such as Zr, Ta, Th, and rare
earth elements, show a very good correlation (rN0.9)
with each other in the conglomerates.
By comparison with Archaean greenstone terrains
in the Kaapvaal Craton, which typically contain, apart
from the dominant mafic to ultramafic volcanic rocks,
also highly siliceous chemical sedimentary rocks and
felsic volcanic rocks, an equivalent to such terrains is
inferred as source region from the abundance of chert
and locally quartz porphyry pebbles in the Witwa-
tersrand conglomerates. A mafic to ultramafic com-
ponent in the source area, as expected for an Archaean
granitoid–greenstone terrane, is indicated by the
abundance of detrital chromite and subordinate
platinum group elements (PGE)-bearing minerals in
the conglomerates, but also by elevated Cr, Co and Ni
concentrations in all Witwatersrand shale units
(Wronkiewicz and Condie, 1987). The ratios between
the different PGE is surprisingly consistent through-
out the Witwatersrand gold mines with (Os+Ir)/
(Os+Ir+Pt+Ru) around 0.7, but significantly different
from younger deposits, such as the Rustenberg
Layered Suite, dunite and kimberlite, for which this
ratio is around 0.1 (De Waal, 1982). This difference
may reflect a high maturity of the placer (Cousins,
1973). However, the consistency in both the PGE
mineralogy and the PGE ratios along strike and down-
slope, and the relative proximal position of the
Witwatersrand placer deposits, may be an indication
of a specific source area characteristic, i.e. that of a
chondritic to subchondritic mantle, as recently pro-
posed for PGE alloys in the Evander gold field
(Malitch and Merkle, 2004). Furthermore, the prior
existence of a relatively stable cratonic block is
implied from the rare presence of diamond in some
reefs (Feather and Koen, 1975; Ramdohr, 1958),
probably related to the presence of kimberlite pipes in
the source area.
Detrital zircon age spectra (Kositcin and Krapez,
2004) indicate the following ages of significant felsic
rocks in the source area: 3310–3300, 3090–3060,
2990–2980, 2950–2940, and 2920–2910 Ma. The
zircon provenance age spectrum for the Central Rand
Group is considerably more complex and spans a
wider range (3450 to 2870 Ma) than that for the West
Rand Group (3300 to 2960 Ma). This confirms the
inferred tectonic setting of a passive margin for the
West Rand Group, with sediment supply from fewer
sources and no tectonic rejuvenation, and that of a
foreland basin for the Central Rand Group, with
increasingly more varied source rocks, continuous
tectonic rejuvenation and erosion to older stratigraphic
levels. Corresponding counterparts for all of the
observed detrital zircon age modes are known from
the surroundings of the Witwatersrand Basin (Frim-
mel et al., in press). Detrital xenotime ages ranging
from 2840 to 2813 Ma are not represented among the
detrital zircon ages, most likely because they were
derived from high-U granitoids, whose zircon grains
would have been metamict (Kositcin and Krapez,
2004). High-U granite and pegmatite bodies of
comparable age are known from the southern Murch-
ison Belt (Poujol, 2001; Poujol and Robb, 1999), near
the Giyani Belt (Kroner et al., 2000) and the
Barberton Belt (Meyer et al., 1994).
Almost 3 billion years of erosion would make any
of the currently exposed tectonic units improbable
source areas for the Witwatersrand sediments. A
nearly complete overlap of detrital zircon and
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4622
xenotime age spectra with ages from Palaeo- to
Neoarchaean granitoid–greenstone terranes surround-
ing the Witwatersrand Basin is, however, evident and
source areas that correspond to at least some of these
terranes are therefore possible. Ages that correspond
to the oldest detrital zircons from the Central Rand
Group are reported only from the Barberton Belt.
Equivalents to the detrital zircon age mode of 3310–
3300 Ma (West and Central Rand Groups) are known
from the Giyani and Barberton Belts. The minor
detrital zircon age modes between 3210 and 3090 Ma
reflect various granitoid bodies that form the basement
to the Witwatesrand Basin. Most of the detrital zircon
grains in both the West Rand and Central Rand
Groups represent the age modes between 3060 and
3080 Ma, which correspond to the time of felsic
volcanism in the Dominion rift. Comparable ages are
also known from felsic volcanic rocks and granitoids
in the Murchison Belt (Poujol and Robb, 1999; Poujol
et al., 1996). All younger detrital zircon grains could
have been sourced, based on sediment transport
directions and age correlations, from higher crustal
level equivalents of the Mesoarchaean Amalia–Kraai-
pan, Murchison and Giyani granitoid–greenstone
terranes.
4. Neoarchaean weathering
One of the pillars, on which recent hydrothermal
models rest, is the apparently large-scale, acidic
hydrothermal alteration of the Witwatersrand Basin
fill (Barnicoat et al., 1997). As weathering under an
acidic Archaean atmosphere would lead to similar
bulk rock chemical changes as acid leaching by post-
depositional fluids, the question arises whether
systematic chemical changes observed in the silici-
clastic successions across unconformities reflect pale-
osol horizons or whether they are related entirely to
post-depositional fluid–rock interaction.
Most of the Witwatersrand siliciclastic rocks did
not achieve thermodynamic equilibrium during post-
depositional alteration, including burial and regional
metamorphism. This is indicated by the widespread
survival of detrital clasts that are now embedded in a
metamorphic matrix. Detrital quartz is distinguished
from secondary quartz by displaying highly variable
cathodoluminescence (Gartz and Frimmel, 1999),
variable degrees of strain and different fluid inclusion
populations (Frimmel et al., 1993). Similarly, detrital
white mica can be distinguished from metamorphic
mica both on textural and compositional grounds
(Frimmel et al., 1993; Sutton et al., 1990). The most
common metamorphic silicates in these rocks are
muscovite, pyrophyllite, chlorite, sudoite and chlor-
itoid. In the coarser grained rocks these are randomly
orientated, whereas in the argillitic units they define a
slaty cleavage. Based on the silicate equilibrium
assemblages, the peak metamorphic temperatures
achieved throughout the Witwatersrand Basin range
between 300 and 400 8C (Phillips and Law, 1994),
except for the area around the Vredefort dome, where
up to amphibolite facies conditions are recorded in the
upturned, lowermost parts of the basin fill, i.e. the
West Rand Group (Gibson and Wallmach, 1995).
Away from the Vredefort dome, the presence of
kyanite and the mineral assemblage chlorite+sudoi-
te+muscovite (and/or pyrophyllite) in the middle
Central Rand Group constrain peak metamorphic
conditions at approximately 300 8C and 3 kbar
(Frimmel, 1997). Sericite pseudomorphs after anda-
lusite at the bottom of the Transvaal Supergroup and
in the Ventersdorp Supergoup (McCarthy et al., 1986)
reflect a lower pressure, i.e. an expected lower
overburden for the higher stratigraphic levels.
Metamorphic biotite is rare but has been described
from the bottom of the Central Rand Group (Phillips
et al., 1988) and also from the West Rand Group,
together with stilpnomelane, relics of K-feldspar and
chlorite (Frimmel, 1994). Feldspar is extremely rare in
the Central Rand Group (Fig 6F), but few units,
particularly within the West Rand Group contain as
much as 30 vol.% detrital K-feldspar that can be
related to a granitic source due to its perthitic textures
(Fuller, 1958). Albite is very rare in the metasedi-
mentary rocks but has been reported as metamorphic
phase from an argillite in the uppermost Johannesburg
Subgroup (Booysens Formation; Frimmel, 1994).
Pyrophyllite is widespread and becomes particu-
larly important towards the basin margin. Cross-
cutting relationships between pyrophyllite distribution
and stratigraphic boundaries have led Barnicoat et al.
(1997) to postulate basin-wide acidic hydrothermal
alteration. These workers reported intense pyrophyl-
litic alteration within mineralised conglomerate hori-
zons, as well as for more than 500 m above and below
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 23
surrounding quartzite. A postulated correlation of gold
distribution with this alteration (Phillips and Law,
1994) that postdates the hydrothermal/metamorphic
phyllosilicates forms the fundamental argument for a
hydrothermal origin of the gold.
Considering the multi-stage tectono-metamorphic
history of the Witwatersrand Basin, it appears unlikely
that all the post-depositional alteration phenomena in
the basin are ascribable to a single hydrothermal
infiltration event. Although there is little doubt that
H+-metasomatism has caused the formation of some of
the pyrophyllite at the expense of mica, as evidenced
by cross-cutting pyrophyllite veinlets, it might be
difficult, in other places, to distinguish such hydro-
thermal acid leaching from the effects of weathering
under acidic conditions. The latter can be expected to
occur along erosional unconformities, where kaolinite
would have been the starting material for pyrophyllite
formation during prograde metamorphism.
The findings of Barnicoat et al. (1997) are at
variance with those of Sutton et al. (1990) who
found, by studying the compositional and minera-
logical changes across stratigraphic units, a strong
stratigraphic control on the chemistry and mineral-
ogy of Witwatersrand arenites and concluded that,
except for potassium, metamorphism was essentially
isochemical.
4.1. Geochemical alteration profiles across
stratigraphic units
The shape of geochemical alteration profiles across
individual unconformities can be used to test whether
observed alteration patterns are related to palaeo-
weathering or to post-depositional metasomatism. In
the former case, a systematic change in bulk rock
chemistry is expected towards the top of the footwall,
but not in the hanging wall. In the latter case, both
footwall and hanging wall should show an alteration
halo around the unconformity, which is presumed to be
the principal post-depositional fluid pathway, with
dispersion causing alteration to comparable extent both
above and below that pathway. To this effect, the
chemical index of alteration (Nesbitt and Young,
1982),
CIA ¼ ½Al2O3=ðAl2O3 þ CaO4þ Na2Oþ K2OÞ�
�100
in which the oxides are expressed as molar propor-
tions and CaO* is CaO in silicates, as opposed to
carbonates and phosphates, was applied to analyses of
siliciclastic rocks across a number of reefs throughout
the Central Rand Group and the upper West Rand
Group (Figs. 4 and 7). The concentration of CaO in
most samples studied is less than 0.1 wt.% but
typically above the lower limit of detection (0.004%).
Small amounts of Na (Na2O contents are typically
around 0.2 wt.%) are bound largely to white mica as
paragonite intergrowths (Frimmel, 1994), except for
albite-bearing shale in the West Rand Group. Con-
sequently, the CIA reflects essentially the distribution
of K-feldspar, muscovite and pyrophyllite, i.e. the
extent to which K was leached out of the rock during
lateritic weathering and/or post-depositional reaction
with an acidic fluid.
Comparison of average CIA values across the
stratigraphic units (Fig. 4) shows significantly lower
indices for the few analyses available from the West
Rand Group. Analyses for most of the West Rand
Group are lacking, but data from a likely stratigraphic
equivalent, the Mozaan Group, a ca. 5000 m thick
siliciclastic succession within the Pongola Supergroup
(Beukes and Cairncross, 1991), may serve as proxies.
There, the matrix of diamictite units and associated
mudstones have significantly lower CIA values (on
average 66) than other mudstones in the Pongola
Supergroup, and these have been interpreted as
indicative of a glacial origin for the diamictite beds
(Young et al., 1998). Most of the available data for the
West Rand Group (Sutton et al., 1990) cluster between
50 and 60, whereby 50 is typical of unweathered
material, representing the composition of fresh feld-
spar. It should be noted that these data were obtained
on arenitic samples and thus reflect sediment that has
been transported over shorter distance. They are
therefore not directly comparable with those obtained
on mudstone for which the calculated CIA values
across the Witwatersrand Supergroup are variable
(70–98; Wronkiewicz and Condie, 1987). Never-
theless, a marked difference in the chemical weath-
ering intensity between West Rand and most of the
Central Rand Group sediments seems to be indicated,
similarly as noted for the Pongola Supergroup. This is
well shown by a systematic increase in the CIA to as
much as 85 towards the sequence boundary with the
Central Rand Group.
Fig. 7. Variation in Chemical Index of Alteration (CIA) and Fe/Al ratio with distance from the Denny’s, Crystalkop (Frimmel and Minter, 2002; note that the length scale in the
original source paper is incorrect) and Ventersdorp Contact Reefs (Gartz and Frimmel, 1999) at various mines in the Klerksdorp gold field (A–G) and the Welkom gold field (H.E.
Frimmel, unpubl. data; H); data are for argillitic to arenitic siliciclastic rocks, except for the hanging wall of the Ventersdorp Contact Reef, which consists of metabasalt and thus has a
lower CIA.
H.E.Frim
mel
/Earth
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ceReview
s70(2005)1–46
24
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 25
For most of the Central Rand Goup, the CIA is
around 80, which reflects above average (Nesbitt and
Young, 1996) extent of chemical weathering. Only in
the uppermost formation do the CIA values decrease
to around 70, which is in agreement with average
chemical weathering. The internal variability in CIA
within the group, particularly across unconformities,
is, however substantial (Fig. 4). A closer inspection of
this variability, using examples from the footwall and
hanging wall of the Crystalkop and Denny’s Reefs at
Vaal Reefs mine, Klerksdorp gold field (Frimmel and
Minter, 2002) and the B-Reef at Freestate No. 3 mine,
Welkom gold field (H.E. Frimmel, unpubl. data)
reveals a gradual upward increase in CIA in the
footwall towards the reef contacts (Fig. 7). This trend
is exemplified by two profiles through the Crystalkop
Reef (Fig. 7A,B). In some areas, this trend can be
traced over more than 10 m (Fig. 7A,B), whereas in
others, a systematic increase in CIA towards the reef
was found over a distance of only a few metres below
the reef (Fig. 7C,D,E,F). In the hanging wall, this
trend is reversed with an abrupt decrease in the CIA
away from the reefs (Fig. 7A,B,C,D,E,H). Note that
the very low CIA in the hanging wall of the
Ventersdorp Contact Reef (Fig. 7G) is an artifact of
the lithology as it consists of metabasalt and not of
arenitic metasedimentary rocks as in all other cases.
Superimposed on the above relatively large-scale
trends, with the highest CIA values reaching a
maximum of 95 in the footwall near the respective
reef contact, is a decrease in CIA on a smaller scale
along the reef horizons. This is best exemplified by
two profiles through the Crystalkop Reef (Fig. 7B,D),
which show a marked drop in the CIA within a few
metres both above and below the reef. Note that this
trend affects both the footwall and the hanging wall,
whereas the former, larger-scale trend appears asym-
metrical below and above the reef horizons. Similarly,
corresponding chemical changes, both above and
below the reef, have also been observed at the
Ventersdorp Contact Reef (Fig. 7G). There an earlier
potassic alteration (Gartz and Frimmel, 1999) affected
the pyrophyllite-bearing footwall arenite as well as the
metabasaltic hanging wall, causing a decrease in CIA
over the top few metres in the footwall and an increase
in CIA over the bottom few metres in the hanging-
wall. This was followed by chloritisation only along
the reef and its immediate contact zones over a few
centi- to decimetres, which is reflected by a sharp but
very local increase in CIA.
4.2. Interpretation of trends in CIA
Trends in CIA are observed on different scales, i.e.
on the hundred metres, few metres and decimetre-
scale. An example for the former is the systematic
increase in CIA towards the West Rand/Central Rand
Group unconformity over more than 100 m. This is
too long to be explained by palaeoweathering along
that unconformity, but may point to a systematic
change in the environmental conditions (increase in
temperature and/or acidity) towards that major strati-
graphic boundary.
The upwardly increasing CIA trends on the scale of
a few metres are typical of, but more extreme than,
those of both modern and Eoproterozoic paleosols
(Nesbitt and Markovics, 1997). With one exception
(Fig. 7F), the very high CIA values are confined to the
footwall, which is not the expected result if they were
related to reef-parallel acidic hydrothermal infiltration.
The very high CIA values are, therefore, ascribed to
intense chemical weathering.
Whereas Ca and Na are removed during the initial
stages of development of a weathering profile, K is
removed only during the latest stages (Nesbitt and
Markovics, 1997; Nesbitt and Young, 1984). Con-
sequently, in Al2O3–(CaO+Na2O)–K2O (A–CN–K)
space (Fig. 8), a typical weathering trend will follow a
line parallel to the A–CN join until it reaches the A–K
join from where it will continue towards the A apex.
Plotting a large data set of arenitic bulk rock analyses,
predominantly from the Central Rand Group with a
few data from the West Rand Group, in that space
(Fig. 8) reveals that the bulk of the non-mineralised
arenitic rocks from various stratrigraphic levels
between individual reef (unconformity) horizons
follows a weathering trend of progressive removal
of CaO and Na2O prior to removal of K2O, with a
starting point close to the average composition of
Neoarchaean upper continental crust. A considerable
number of samples are displaced towards the K apex,
thus indicating variable and, in places, considerable
K-metasomatism. Comparison between footwall and
hanging wall analyses (from a distance of up to 1 m
below or above a given reef), shows that especially
hanging wall samples show evidence of this K-
Fig. 8. A–CN–K (Al2O3–[CaO+Na2O]–K2O) diagram showing
alteration trends in siliciclastic metasedimentary rocks from
auriferous reefs (conglomerate), immediate footwall and hanging-
wall of reefs, and from arenite of various stratigraphic positions in
the Witwatersrand Supergroup (n=226); data from Frimmel and
Minter (2002), Gartz and Frimmel (1999), Sutton et al. (1990), Q.
Hawes (unpubl. data), and H.E. Frimmel (unpubl. data); small
triangular diagram in top left corner shows general alteration
trends: 1—chemical weathering of average Neoarchaean upper
continental crust (��������) from Condie (1993), 2—Ca/Na-metasomatism,
3—K-metasomatism.
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4626
enrichment, which consequently has to be inferred as
post-depositional. Such metasomatism would explain
the local negative excursion in CIA along the
Crystalkop Reef as shown in Fig. 7D. In contrast,
most footwall samples plot very close to the A apex
(very high CIA values)—a feature that is absent in the
hanging wall samples and therefore regarded as
reflecting palaeo-weathering.
Analyses from within reef beds display a wide
spread close to the A–CN line (Fig. 8). Such a trend
cannot be explained by any weathering process but
clearly illustrates the effects of post-depositional
alteration. As most of the plotted analyses come from
the Ventersdorp Contact Reef, the enrichment in Ca,
and to a lesser extent Na, reflected by the trend
towards the CN apex, can be explained by interaction
of a hydrothermal fluid with the overlying Ca- and
Na-rich metabasalt of the Klipriviersberg Group.
Alteration patterns around this reef (Fig. 7G) also
illustrate well the effects of K-metasomatism as
described in more detail by Gartz and Frimmel
(1999). Using Ti as the least mobile reference
element, a sharp increase in K is observed with
increasing proximity to the reef, which reflects
sericitisation. Only in the immediate reef environ-
ment, within less than one metre distance, the rocks
are depleted in K and enriched in Fe, reflecting
chloritisation. As the potassic alteration over several
metres and the smaller-scale ferric alteration over a
few decimetres affected both the footwall and the
hanging wall, they cannot be related to weathering on
a palaeo-surface but must be explained by reef-
parallel post-depositional hydrothermal fluid flow as
postulated on structural grounds by Jolley et al.
(1999). This type of alteration might also be respon-
sible for the halo of elevated CIA values observed
around the Denny’s Reef (Fig. 7F) and a sericitisation
similar to that in the VCR might explain the local
sharp drop in the CIA values around the Crystalkop
Reef as shown in Fig. 7D.
In summary, large-scale trends in CIA may be the
result of an overall change in climate and/or reflect
different degrees of sediment re-working. Systematic
increases in CIA to very high values in the footwall
beneath erosional unconformities on the scale of
several metres is ascribed to deep chemical weath-
ering along these palaeo-surfaces and thus provide
indirect information on the contemporaneous atmos-
pheric composition. Small-scale variations in CIA in
the immediate vicinity of reefs, typically over centi- to
decimetres, reflect dispersive metasomatism caused
by reef-parallel fluid flow.
5. Pre- or post-depositional age of the ore?
A sedimentary model for the Witwatersrand gold
deposits, originally proposed by Mellor (1916), was
first challenged by Graton (1930), who suggested a
magmatic–hydrothermal model. Since then, workers
who have studied the ore and host rocks on a
microscopic scale have emphasised a hydrothermal
model, because gold is typically late in the para-
genetic sequence as micro-fracture fills and inclusions
in secondary, clearly hydrothermal mineral grains
(Feather and Koen, 1975; Ramdohr, 1958). In con-
trast, those who have studied the rocks on a macro-
scale noted a strong sedimentological control on gold
grade, which has been used highly successfully
throughout the history of Witwatersrand exploration
and day-to-day mining, and prompted them to
advocate a sedimentary palaeoplacer model (Minter,
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 27
1978). Following the generally accepted recognition
of regional metamorphism and post-depositional fluid
flow throughout the basin (Phillips and Law, 1994), a
palaeoplacer model sensu stricto, in which all gold
particles are perceived as detrital, has been aban-
doned. Today there is general agreement that the
majority of gold particles that appear late in the
paragenetic sequence are indeed hydrothermal precip-
itates. Thus, the debate has shifted to the question of
whether the source of that hydrothermal gold was
proximal, fluvially deposited, detrital gold within the
conglomerate beds (modified palaeoplacer model), or
external to the host rocks (hydrothermal model).
These two possibilities represent end-member models,
which are, each with variations, the focus of current
debate:
(1) In the modified palaeoplacer model, transport
of detrital gold particles into the host sediments is
assumed to have taken place by fluvial processes
with subsequent short-range mobilisation of the
gold by infiltrating hydrothermal fluids and/or
degradation of in situ hydrocarbon or hydrous
phases. Gold mobilisation and recrystallisation
induced by hydrothermal fluid infiltration has been
ascribed to the emplacement of the Ventersdorp
Supergroup lavas (Pretorius, 1991), to burial meta-
morphism in lower Transvaal Supergroup times and
the Vredefort impact event (Frimmel et al., 1999),
as well as to Pretoria Group deposition and the
emplacement of the Bushveld Igneous Complex
(Robb et al., 1997).
(2) Hydrothermal models explain the presence of
gold as the result of post-depositional hydrothermal
fluids from an external source. The presence of gold
in the conglomeratic host rocks is inferred as
consequence of long-range, basin-wide fluid flow
combined with chemical and structural controls
(Phillips and Law, 2000). One hydrothermal model
infers a separate origin for gold, uraninite and
hydrocarbons (Phillips and Law, 2000), whereas
another seeks to explain all of them as cogenetic
(Barnicoat et al., 1997). The infiltration of the
inferred auriferous hydrothermal fluids has been
linked with several different events that range from
Ventersdorp volcanism (Phillips et al., 1997),
regional metamorphism (Phillips and Myers, 1989),
or the emplacement of the Bushveld Igneous Com-
plex (Stevens et al., 1997).
All of the currently debated genetic models
include the presence of gold derived from the
movement of hydrothermal fluids, but they differ
principally in the inferred distances of hydrothermal
gold transport and the composition of that gold-
transporting fluid. The composition of the gold-
bearing fluid has been suggested to be similar to that
inferred for Archaean orogenic gold deposits (Phil-
lips and Law, 2000), in which gold is assumed to
have been transported as bisulphide complex in an
H2O–CO2 dominated, relatively reducing, low-sul-
phur, low salinity fluid. In that model, all rounded
pyrite is assumed to be the product of post-deposi-
tional hydrothermal sulphidation of black sands that
existed originally of various Fe–Ti oxides and Fe-
oxyhydroxide pisolites (Phillips and Myers, 1989).
This contrasts with models that suggest gold transport
in highly acidic, oxidising fluids (Barnicoat et al.,
1997), and models which prefer gold transport as a
hydroxy-complex (Gray et al., 1998). Reduction of
aqueous gold species to elemental gold by interaction
with pre-existing bitumen plays an important role in
all hydrothermal models and is used to explain the
strong association of gold and bitumen and the
textural position of gold grains within microfractures
in bitumen.
A major argument in favour of an external,
hydrothermal source of gold is based on mineral and
chemical zonation patterns at the deposit to hand-
specimen scale as well as elemental correlations (Fox,
2002). The recorded zonation patterns clearly dem-
onstrate hydrothermal fluid–rock interaction, but do
not contribute to solving the question of the gold
source. Hydrothermal versus detrital element correla-
tions are more crucial in this regard. Excellent
correlations between Au, U and Ag, a poor but
positive correlation of Au with Cr and (Co+Ni), and
only a very weak correlation of Au with Zr have been
noted in several studies on the Kimberley Reef
(Rasmussen and Fesq, 1973), the Vaal Reef (Fox,
2002), the Steyn Reef (Frimmel and Minter, 2002),
and the B- as well as the Ventersdorp Contact Reef
(H.E. Frimmel, unpubl. data).
The good correlation between Au and Ag is
expected, because gold is the principal sink for Ag,
with Au and Ag being alloyed in variable proportions
with each other and Hg (Frimmel and Gartz, 1997;
Reid et al., 1988; Utter, 1979). Similarly, the positive
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4628
correlation with Co and Ni is readily explained by
precipitation of cobaltite–gersdorffite together with
hydrothermal gold particles (Fox, 2002; Frimmel et
al., 1993) and Co as well as Ni enrichment in co-
existing pyrite. The good correlation between Au and
U as well as Cr can be ascribed to a spatial association
with heavy sands that contained detrital uraninite and
chromite particles. As the formation of bitumen is
genetically and spatially related to detrital uraninite,
hydrothermal gold precipitation by a locally available
reductant would have taken place preferentially near,
or within, fractures of uraninite, thus further exacer-
bating the positive Au–U correlation.
The poor correlation between Au and Zr has been
used as evidence against a detrital origin of the gold,
because the distribution of Zr is controlled by zircon,
almost all of which is detrital (Fox, 2002). Apart from
gold and zircon being derived from different source
rock types, the poor correlation could be a function of
differences in the hydraulic behaviour during repeated
sediment re-working (Smith and Minter, 1980). It is
noteworthy that many Zr-rich samples are devoid of
Au, simply indicating a gold-poor source area,
whereas only very few samples show elevated Au
but low Zr contents and these typically contain
auriferous veinlets indicative of local gold mobilisa-
tion. If the gold had been introduced into the host
sedimentary rocks by hydrothermal fluids, both
zircon-rich and zircon-poor domains should have
been affected to a similar degree. This is not the case.
A less than perfect correlation between Au and other
elements concentrated in detrital minerals, such as Zr,
is likely the result of dispersion of the gold by short-
range hydrothermal mobilisation.
Other arguments for a hydrothermal origin include
the observation that the bulk of gold grains are located
within or near micro-fractures, filled with bitumen and
uraninite/brannerite, which post-date early, bedding-
parallel pyrite–pyrrhotite–quartz filled fractures that
contain no gold (Fox, 2002). While this observation
demonstrates the undisputed hydrothermal nature of
the bulk of the gold particles, it does not clarify the
distance of hydrothermal gold transport. The same
applies to effectively all other arguments that have
been brought forward in favour of a hydrothermal
model (Table 1). In contrast, a number of observations
and analytical data can hardly be explained other than
in terms of a modified palaeoplacer model.
Some of the most important lines of evidence for a
modified placer model are the observation of gold
micro-nuggets that are spatially associated with
secondary, locally remobilised, hydrothermal gold.
Although these spheroidal to torroidal micro-nuggets
are very rare and have so far only been found in
samples from the Basal Reef, Vaal Reef, B-Reef and
Crystalkop Reef (Frimmel and Minter, 2002; Minter et
al., 1993), their existence gives a clear clue as to the
primary process of gold enrichment in the Witwaters-
rand sediments. Strong support for such a sedimentary
gold enrichment process also comes from the Re–Os
isotope data, which indicate an age for the gold that is
clearly older than that of host sediment deposition, but
the same as for rounded pyrite (3033F21 Ma; Kirk et
al., 2001, 2002). Notwithstanding local evidence of
sulphidation of Fe–Ti oxides (Ramdohr, 1958), chert
and iron formation pebbles (Hallbauer, 1986; Hirdes
and Saager, 1983), the Re–Os data imply a detrital
origin of the most abundant, rounded form of pyrite.
All of the above examples of sulphidation can be
related to the same fluids that caused the formation of
secondary pyrite at various stages throughout the
complex post-depositional alteration history of the
Witwatersrand sedimentary rocks. The noted textural
association of gold as inclusions within secondary
pyrite and the lack of gold inclusions in rounded pyrite
further indicate that hydrothermal gold transport
cannot be linked with a postulated sulphidation event
that supposedly formed all the pyrite. A similar
argument can be applied to uraninite. The U–Pb age
of 3050F50 Ma obtained for uraninite from the
Dominion Reef (Rundle and Snelling, 1977) is older
than the age of Witwatersrand sediment deposition
and, though subject to a considerable uncertainty,
effectively rules out a post-depositional introduction of
significant amounts of U into the basin fill.
In particular, the interpretation of the Re–Os data
appears to be very robust as the only alternative to the
interpretation given above would be that of isotope
mixing between two or more different ad-hoc end
members. Any mixing model would be difficult to
reconcile with the excellent agreement between the
Re–Os isochron ages of gold and rounded pyrite from
various localities, the high precision of the isochron
ages, and the initial Os isotopic compositions that are
identical within error to the Os isotopic composition
of the mantle at 3 Ga.
Table 1
Main arguments for a hydrothermal and modified palaeoplacer model, respectively, for the Witwatersrand gold
Hydrothermal model Modified palaeoplacer model
Gold is late in paragenetic sequence (Barnicoat et al., 1997; Feather
and Koen, 1975)
Rare co-existence of rounded gold micro-nuggets with secondary,
hydrothermal gold on mm-scale (Minter et al., 1993)
Gold is associated with acid metamorphic alteration (Phillips and
Myers, 1989; Barnicoat et al., 1997)
Composition of fluid inclusions in auriferous hydrothermal quartz
indicate neutral to basic pH (Frimmel et al., 1999)
Basin-wide distribution of pyrophyllite related to large-scale
H+-metasomatism (Barnicoat et al., 1997)
Increase in chemical index of alteration in footwall towards reef
related to chemical weathering under acid atmosphere (Sutton et al.,
1990; Frimmel and Minter, 2002)
Abundant rounded pyrite and uraninite particles associated with the
gold ore are of post-depositional hydrothermal origin (Barnicoat
et al., 1997)
Isotopic data for rounded sulphides and uraninite yield ages older
than time of sedimentation (Rundle and Snelling, 1977; Kirk et al.,
2001)
Rounded pyrite derived from basin-wide sulphidation of dblacksandsT (Phillips and Myers, 1989)
Pyrite morphology, cyrstallography, and truncated growth zonation
patterns indicate detrital nature of rounded grains (McLean and Fleet,
1989; Fleet, 1998; England et al., 2002b)
Strong correlation between gold and hydrothermal pyrobitumen
(Nagy, 1993; Gray et al., 1998)
Derivation of hydrothermal pyrobitumen from local intrinsic oils,
based on bulk and molecular y13C data (Spangenberg and Frimmel,
2001) and fluid inclusion studies (England et al., 2001a)
Local variability in gold composition within a given reef reflects
differences in source areas (Frimmel and Gartz, 1997)
Conglomerate beds were preferred channels for infiltration of
gold-bearing hydrothermal fluids (Barnicoat et al., 1997)
Sedimentological control on gold distribution (Minter, 1978);
Negative correlation between authigenic/hydrothermal xenotime and
ore bodies (Kositcin et al., 2003)
Gold was introduced into the Witwatersrand Basin after sediment
deposition (Phillips and Myers, 1989; Barnicoat et al., 1997)
Re–Os ages of the gold are older than time of sedimentation (Kirk
et al., 2002)
Hydrothermal introduction of gold into the basin during peak
metamorphism, coeval with 2.06 Ga Bushveld event (Phillips and
Law, 1994)
Lack of significant secondary permeability after N600 my of burial,
diagenesis and low-grade metamorphism (Frimmel, 1997)
Hydrothermal introduction of gold into the Witwatersrand basin
during global 2.7–2.6 Ga gold-forming thermal event, coeval with
Ventersdorp Supergroup volcanism (Phillips et al., 1997)
Sedimentary reworking of Witwatersrand gold ore in
late-Ventersdorp diamictite was followed by Witwatersrand-style
gold mineralization in the post-Ventersdorp Black Reef
Lack of suitable source area for placer gold (Phillips and Myers,
1989)
Calculated background value for eroded source area corresponds to
mean Au content of Archaean granitoid–greenstone crust (Loen,
1992)
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 29
If the gold were brought into the Witwatersrand
Basin during post-depositional fluid infiltration, large
fluid/rock ratios would be expected. The auriferous
fluids would have had to flow preferentially along the
conglomerate beds in order to explain the apparent
sedimentological control on the basin-wide ore dis-
tribution. Although some evidence exists for bedding-
parallel fluid flow (Jolley et al., 1999), mass balance
calculations (Gartz and Frimmel, 1999) point to rather
limited external fluid infiltration into the reefs. Only if
all the pyrophyllite is explained by post-depositional
H+-metasomatism (Barnicoat et al., 1997), can a case
be made for large-scale fluid infiltration. Bearing in
mind that the loss of alkalies on the scale of tens of
metres can be attributed to palaeoweathering, as
outline above, it is more likely that a large proportion
of pyrophyllite in the Witwatersrand metasedimentary
rocks is derived from the prograde metamorphism of a
kaolinite-bearing protolith, thus revoking the neces-
sity for significant post-depositional fluid infiltration.
However, even if all the pyrophyllite were hydro-
thermal–metasomatic, it has been shown from fluid
inclusion analyses that in those cases studied, hydro-
thermally mobilised gold was transported by a fluid of
a composition that is incompatible with the stability of
pyrophyllite (Frimmel et al., 1999). Only limited
interaction between potentially auriferous fluids and
the host conglomerate beds is also supported by
studies on the chemistry and age distribution of
xenotime, which is a particularly useful monitor for
hydrothermal infiltration (England et al., 2001b;
Kositcin et al., 2003). These studies revealed that
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4630
diagenetic xenotime is relatively abundant in the reefs
compared to their surroundings, but hydrothermal
xenotime is present in much lesser amounts within the
reefs. This implies that the reefs were more permeable
during diagenesis but less permeable during post-
diagenetic hydrothermal events, which is in disagree-
ment with proposed hydrothermal models.
Apart from the textural, geochemical and isotopic
evidence, there is also important geological evidence
that speaks against a hydrothermal model not only
for the Witwatersrand gold, but also for the
associated pyrite and uraninite. The Witwatersrand
ore bodies must have formed prior to deposition of
the Platberg Group sediments as the latter contain
pebbles of mineralised Witwatersrand reef material
(Phillips et al., 1997). Yet, Witwatersrand-style
mineralisation is evident in the Black Reef at the
base of the Transvaal Supergroup that is some 70
million years younger than the Platberg Group.
Formation of a placer deposit under similar environ-
mental conditions repeatedly through time is to be
expected, but the same style of hydrothermal metal
introduction into the basin at different stages of basin
development in different tectonic settings is unlikely.
If there had been a second major hydrothermal ore-
forming event in the central Kaapvaal Craton,
including the conglomerates at the base of the
Transvaal Supergroup, it would have affected the
Witwatersrand rocks after they had undergone dia-
genesis and first low-grade metamorphism. Very
little permeability would have remained at that stage
in the Witwatersrand strata, and basin-wide, large-
scale, predominantly bedding-parallel fluid flow, as
postulated by the various hydrothermal models,
would have been effectively impossible.
Problems with the hydrothermal models extend
from the micro- to the macroscale. On a tectonic
scale, the Witwatersrand deposits have often been
compared with orogenic gold deposits by those who
favour a hydrothermal model. This is mainly for the
similarity in the paragenetic sequence and in the
common gold–pyrite–hydrocarbon association (Phil-
lips and Myers, 1989). There are, however, a number
of significant differences between the two styles of
deposit (Frimmel et al., in press; Groves et al., 2003).
The rounded, sub-spherical morphology of most of
the Witwatersrand pyrite and its highly variably
geochemical and S isotopic composition (England
et al., 2002b) are in contrast to the typically subhedral
to euhedral, compositionally restricted pyrite found in
orogenic deposits. Wallrock alteration, inferred to
have taken place over several hundreds of metres
across stratigraphic boundaries throughout the Wit-
watersrand Basin (Barnicoat et al., 1997), is orders of
magnitude more extensive than known from any
orogenic gold deposit. The latter are characterised
by the abundance of auriferous quartz veins, but the
basin that hosts the by far greatest known concen-
tration of Au is characteristically devoid of a
plentitude of such veins. A foreland/retroarc basin
setting is indicated for the bulk of Witwatersrand
deposits. This is analogous with modern placer gold
deposits but in stark contrast to orogenic gold
deposits, which typically occur in near-arc or arc
settings. Finally, the geometry of the Witwatersrand
orebodies (gently dipping decimetre to metre thick,
laterally extensive sheets) is unlike the overall shape
of most known epigenetic or orogenic deposit.
5.1. Best-fit genetic model
A modified palaeoplacer models accounts best for
all the available data and observations. Clastic sedi-
ments, first laid down in the 3074 Ma Dominion rift
graben, were largely derived from felsic sources,
which led to enrichment in detrital uraninite but only
low gold contents. During subsequent West Rand
Group sedimentation (2985–2914 Ma), progressively
more mafic rocks from Mesoarchaen greenstone belts
in the hinterland were eroded, but only during Central
Rand Group times (2902–2780 Ma) were the high
levels of gold concentrations in the placer sediments
reached, for which the Witwatersrand has become so
famous. These extraordinary gold concentrations are
explained by a combination of factors that range from
fertile source regions, to tectonic setting and to
palaeoenvironmental conditions. Initially, the Central
Rand Basin took a foreland position relative to the
overriding Kimberley block in the west, with the
intervening Amalia–Kraaipan greenstone belt repre-
senting an obducted slice of former oceanic crust. A
subsequent change in the continental stress field led to
the accretion of the Murchison greenstone belt to the
north, with the Central Rand Basin taking a retroarc
position (Frimmel et al., in press). Gold and chromite,
predominantly from the surrounding greenstone belts,
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 31
together with uraninite and zircon from the associated
granitoids, were brought into this basin by fluvial
transport and concentrated by mineral sorting. Aeolian
transport and re-sedimentation on a number of ero-
sional unconformities led to further up-grading of
gold along the basal degradation surfaces. Sedimen-
tary re-working in braided stream systems and
effective wind sorting were particularly vigorous
because of intense acid weathering and a lack of
vegetation and widespread organisms.
The complex post-Witwatersrand tectono-thermal
evolution of the central Kaapvaal Craton affected the
gold to a variable extent. The bulk of fluid flow
through the Witwatersrand Basin must have taken
place due to diagenetic dewatering. Different parts of
the basin experienced further fluid flow to various
degrees as consequence of the following events: (1)
low-grade burial metamorphism; (2) the syn-Venters-
dorp thermal anomaly, including syn-Platberg rifting;
(3) dynamic metamorphism in a compressional stress
field during thrusting of the Southern Marginal Zone
of the Limpopo Belt on to the Kaapvaal Craton; (4)
lower Transvaal thermal subsidence-induced exten-
sion; (5) the thermal anomaly related to the Bushveld
magmatic event; and (6) pervasive fracturing due to
the Vredefort impact. Some of these events caused the
mobilisation of the gold, together with other detrital
phases, such as pyrite and uraninite. Hydrocarbons,
derived from oil/bitumen, played an important role in
the re-precipitation of the mobilised gold by acting as
reductants. The distances over which gold, pyrite and
uraninite were mobilised, in general were in the order
of millimetres to centimetres. Locally, fracture-con-
trolled fluid flow allowed transport of Au over longer
distances. Most of the hydrocarbons, and thus by
implication gold, was first mobilised during dia-
genesis, but meteoric waters that percolated through
Vredefort impact-related secondary interconnected
(micro-) fracture space were also capable of trans-
porting hydrocarbons, sulphides and gold.
6. Neoarchaean palaeoenvironment
6.1. Neoarchaean atmosphere
It has long been recognised that attempts to
constrain the Archaean atmospheric composition
hinge essentially on four lines of evidence: (1)
presence of detrital uraninite; (2) presence of detrital
pyrite; (3) composition of detrital gold particles; and
(4) the presence and composition of paleosols (Hol-
land, 1984). Since then the debate has shifted from a
previously favoured palaeoplacer model to various
hydrothermal models for the Witwatersrand gold, thus
invalidating the reliability of the above pieces of
evidence (Holland, 1994). The recognition that the
best-fit model for the Witwatersrand deposits is that of
a modified palaeoplacer, reaffirms however the use-
fulness of the above pieces of evidence for the debate
on the Archaean atmospheric composition.
The significance of the Witwatersrand in this
regard is obvious, considering that all four lines of
evidence can be tested there. However, other Meso-
archaean to Eoproterozoic deposits that bear strong
similarities with those of the Witwatersrand outside
the Kaapvaal Craton should not be ignored and can
contribute useful additional information. These
include the 2.13 to 2.10 Ga Tarkwaian System
(Ghana), the 2.09 to 1.88 Ga Jacobina and the poorly
dated 2.8 to 2.2 Ga Moeda deposits (Brazil), as well as
the 1.90 Ga Roraima Supergroup (northern South
America). Similar styles of mineralisation with
abundant detrital pyrite and uraninite, but a conspic-
uous lack of gold, are known from the 2.9 to 2.6 Ga
Bababudan Group (India) and the 2.45 Ga lower Elliot
Lake Group (Huronian Supergoup) in Canada. All of
these deposits have in common that a case for a
palaeoplacer origin can be made (Frimmel et al., in
press), and that the fluvial siliciclastic host sediments
were laid down in foreland/retroarc basins. An
Archaean to Palaeoproterozoic greenstone terrain is
suggested as the most likely source area for all of the
above gold palaeoplacer deposits. In addition, detrital
pyrite, uraninite, and signficantly also siderite, have
been reported from effectively unmetamorphosed
fluvial siliciclastic sedimentary rocks from the Pilbara
Craton (Rasmussen and Buick, 1999).
The first piece of evidence to be assessed is the
presence of detrital uraninite. Based on thermody-
namic grounds U4+-minerals, such as uraninite, are
not stable under modern atmospheric conditions, as
they would oxidise rapidly (Fig. 9). The dependence
of the uraninite stability on oxygen fugacity is almost
independent of pH and the fugacities of other critical
species, such as CO2 and CH4. Consequently, these
Fig. 9. Oxygen fugacity versus temperature diagram showing the
conditions at which oxidation of pyrite to goethite (solid lines) takes
place at variable pH, as well as that of uraninite to a dissolved
oxyhyroxide (dashed line); calculated using PHREEQC (Parkhurst
and Appelo, 1999).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4632
reduced U-minerals are unlikely to survive mechan-
ical transport in the fluvial environment under an
oxidising atmosphere (e.g. Holland, 1994). Localised
modern occurrences of detrital uraninite, such as in
sand from the Indus river, are confined to domains
that have experienced virtually no chemical weath-
ering and can, therefore, not be used as analogues for
the Witwatersrand occurrences (Maynard et al., 1991),
for which intense chemical weathering is indicated.
Detrital uraninite is abundant from the oldest
siliciclastic rocks, the Dominion Reef, throughout
the entire Witwatersrand Supergroup, to the base of
the Ventersdorp Supergroup. Its occurrence thus spans
a period from 3074 to 2714 Ma. The uraninite/
brannerite ratio varies greatly between reefs. Branner-
ite is not important in the Dominion Reef. A system-
atic decrease in the uraninite/brannerite ratio towards
younger stratigraphic levels has been noted in the
Welkom gold field, where this ratio changes from 8.7
in the Steyn Reef to zero in the Beatrix Reef (Minter
et al., 1988). In the Klerksdorp gold field, the younger
Ventersdorp Contact Reef does contain uraninite, but
its textural relationships (typically as inclusions within
bitumen) do not permit to distinguish between a
detrital and a hydrothermal derivation. No uraninite is
known from the 2642 Ma Black Reef, in which all U
occurs as brannerite. This trend towards lower
uraninite/brannerite ratios up-section might be inter-
preted as reflecting repeated sediment re-working
under an atmosphere that became slightly more
oxidising in the course of the Neoarchaean Aera. As
the brannerite is secondary, such an interpretation is,
however, not imperative and this trend might equally
reflect a more oxidising hydrothermal fluid composi-
tion at shallower crustal levels. The latter explanation
is preferred, because of the abundance of detrital
uraninite in the Elliot Lake Group, which is signifi-
cantly younger than the Black Reef.
The second piece of evidence concerns rounded
pyrite, whose detrital nature has been established. The
stability of pyrite requires even lower oxygen fugacity
than that of uraninite (Fig. 9) thus supporting a
reducing environment. Kinetic limitations to the
solubility of pyrite are unlikely to explain the
preservation of pyrite in the fluvial to fluvio-deltaic
sedimentary rocks of the Witwatersrand. In particular,
along those unconformities that reflect repeated re-
working of the underlying sediment, pyrite must have
been exposed to the meteoric environment over
sufficient lengths of time to equilibrate with its
immediate environment.
The rounded pyrite type includes a variety of
textural forms, ranging from compact rounded to
ooid-like particles, all of which have formed prior to
sediment deposition. However, as the mentioned
petrographic and S isotope study (England et al.,
2002b) revealed, some varieties represent pseudo-
morphs after other minerals, including sulphates, with
replacement having taken place before erosion of the
source rocks. It must be emphasised that rounded,
detrital pyrite is not restricted to some small, localised
occurrences, but is a characteristic feature of all fluvial
deposits that range in time from the Dominion Group
to the bottom of the Transvaal Supergroup (3074–
2642 Ma), and in space over several hundred square
kilometres. The common occurrence of detrital pyrite
in fluvial sediments is not a peculiarity of the
Kaapvaal Craton, but also typical of all other known
Neoarchaean to early Palaeoproterozoic fluvial depos-
its, such as those of the Pilbara Craton in Australia,
the Elliot Lake Group in Canada, the Bababudan
Group in India and the Moeda deposits of Brazil, all
of which are older than 2.2 Ga.
The third piece of evidence, the composition of
detrital gold particles is more problematic and less
conclusive. Modern placer gold tends to be depleted
in Ag, owing to an increased mobility of Ag relative
to Au in oxidising waters (Morrison et al., 1991).
Modern placer gold typically shows compositional
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 33
zonation with an increase in Au/Ag ratio towards the
rim of an individual particle (Groen et al., 1990). Such
a zonation pattern is absent in all of the thousands of
studied gold particles from the Witwatersrand. This is
readily explained by diffusional homogenisation with
respect to Ag (and Hg) in gold particles at the
temperatures to which the Witwatersrand gold was
exposed (Frimmel et al., 1993). In spite of this intra-
grain homogenisation, inter-grain homogenisation
was not achieved as evidenced by considerable
differences in gold composition between reefs (Utter,
1979), between different domains within a given reef
(Frimmel and Gartz, 1997) and even between indi-
vidual grains on a hand-specimen scale (Reid et al.,
1988). These differences are most likely a function of
provenance and reflect gold sources of variable
composition.
Mobilisation of detrital gold particles during post-
depositional hydrothermal alteration may have led to
some modification of the composition, because of
different mobility of Au- and Ag-bearing dissolved
species. Consequently, individual gold particles that
formed as hydrothermal precipitates should not be
used as reference for deciphering contemporaneous
atmospheric oxidation potential. However, the com-
positions of preserved detrital gold particles may be
more illuminating in this regard. The only example of
well studied detrital gold particles from the Witwa-
tersrand concerns the Basal Reef in the Welkom gold
field, for which average Ag and Hg concentrations of
8.9 and 1.2 wt.%, respectively, have been obtained
(Frimmel et al., 1993). Such a composition is in very
good agreement with Archaean greenstone-hosted
gold (for compilation see Morrison et al., 1991).
The elevated Ag contents could thus be used as
argument against an oxidising environment during
fluvial transport, but it may be argued that Ag-
depleted rims that had developed were subsequently
mechanically eroded during fluvial transport. Further-
more, Au-rich rim formation in placer gold is
probably related to a combination of self-electro-
refining and cementation instead of selective leaching
of Ag and intra-grain diffusion (Groen et al., 1990), in
which case little inferences can be made regarding
atmospheric conditions.
Last but not least, paleosols can provide some of
the most reliable information on the composition of
the atmosphere at the time of surface exposure. In an
extensive review of paleosols described from the time
period of interest here (Rye and Holland, 1998), it was
concluded that all paleosols older than 2.2 Ga are
characterised by significant Fe-loss. This includes
examples from the Witwatersrand, but they are all
problematic because of their complex post-depositio-
nal alteration history. Metamorphic chloritisation or
formation of pyrophyllite at the expense of mica
would cause enrichment or depletion in Fe, respec-
tively. In many cases, it is not clear, whether an
observed loss in Fe is related to palaeo-weathering or
hydrothermal alteration. At least some of the inferred
paleosols from the Witwatersrand, previously used to
make inferences regarding Archaean atmospheric
conditions, appear to reflect hydrothermally altered
zones (Palmer, 1986), as documented, for example,
for the so-called Deelkraal paleosol below the
Ventersdorp Contact Reef (Jolley et al., 1999).
From the geochemical changes across stratigraphic
boundaries, outlined above, chemical weathering over
several metres below unconformity surfaces can be
deduced and distinguished from hydrothermal alter-
ation that took place, in many cases, over only a very
limited distance away from a given reef. Assuming
that Al behaved conservatively, the extent of Fe-
enrichment or depletion may be illustrated by the total
Fe/Al ratio. A sympathetic relationship between Fe/Al
ratio and CIA can reflect either hydrothermal chlori-
tisation, as exemplified by the Ventersdorp Contact
Reef (Fig. 7G) or a stratigraphic difference in
sediment grain size and thus clay content of the
protolith, as is the case in the footwall of the B-Reef
(Fig. 7H). The former is typically found over very
short distances (Fig. 7F,G), whereas the latter is found
over longer distances across stratigraphic boundaries
(Fig. 7H). In contrast, an antipathetic relationship may
be explained by weathering under an acidic, reducing
atmosphere. Indeed, some Witwatersrand reefs show a
distinct depletion in Fe in their immediate footwall
that is not linked to a corresponding decrease in CIA
(Fig. 7B,C,D,F). Even in profiles in which no distinct
trend is recognisable, the overall Fe/Al ratios tend to
be very low, i.e. less than 0.2. This compares well
with numerous examples of Meso- to Neoarchaean
paleosols from other areas, for which Fe depletion has
been described (Rye and Holland, 1998).
The exact acidity of the contemporaneous atmos-
phere is difficult to constrain, but an upper limit on pH
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4634
can be placed from the lack of detrital feldspar in
otherwise relatively immature siliciclastic sediment.
The principal chemical weathering product was
probably kaolinite, which subsequently gave rise to
the abundant metamorphic pyrophyllite in the succes-
sion. In a number of samples, detrital white mica can
still be recognised (Frimmel et al., 1993; Sutton et al.,
1990). A pH close to the boundary between the
kaolinite and muscovite (illite) stability fields, calcu-
lated as 6.2 at a temperature of 25 8C (for aK+=0.1), is
therefore likely. Such acid weathering agrees well with
the very high CIA values obtained for most footwall
sections beneath unconformities, where they exceed
80 (for comparison, the CIA of modern tropical stream
sediments is around 75, Maynard et al., 1991).
Further evidence for a reducing sedimentary
environment during upper Witwatersrand times comes
from the bulk and molecular isotopic composition of
the indigenous kerogen component in bitumen (Span-
genberg and Frimmel, 2001). The C isotopic compo-
sition and distribution of n-alkane in stratiform
bcarbon seamsQ within reefs point to considerable
input from autochthonous algal-bacterial lipids. It may
be argued that the spatial association between these
hydrocarbons and all the other evidence provided
above for a reducing Mesoarchaean atmosphere may
render this evidence inconclusive. The argument
could be that localised areas that were covered by
terrestrial biomass, provided islands in which reduc-
ing conditions prevailed under an overall oxidising
atmosphere. However, the distribution of detrital
pyrite and uraninite, and even gold, is not at all
restricted to domains rich in hydrocarbons. The bulk
of siliciclastic fluvial to fluvio-deltaic sedimentary
rocks of the Witwatersrand do not contain bcarbonseamsQ (they are restricted to localised lithofacies on
unconformities) but contain abundant rounded pyrite
as well as elevated uraninite concentrations.
A reducing Meso- to Neoarchaean atmosphere,
inferred here from multiple aspects of the nature of the
Witwatesrand placer deposits, is also in agreement
with totally different types of data from mass-
independent fractionation of S isotopes. This phenom-
enon, which refers to the deviation from the mass-
dependent relationship between S isotopes typical of
most processes in aqueous solution or solid phase
(d33Si0.515d34S, d36Si1.91d34S), has so far been
reported from Archean to Eoproterozoic sedimentary
sulphate and sulphide minerals, volcanic beds in ice
cores and modern sulphate aerosols (Farquhar and
Wing, 2003). A large scatter in d33S, which does not
obey mass-dependent relationships between S iso-
topes, is evident in samples older than 2.45 Ga, with a
transition for the loss of this peculiar isotopic
behaviour spnning from 2.45 to 2.09 Ga (Farquhar
et al., 2000). The phenomenon, which has since been
verified with data from a number of older cratons
(Farquhar and Wing, 2003; Mojzsis et al., 2003), is
explained by photochemical reactions, such as SO2
photolysis. As SO2 and SO photolysis are caused by
intense ultraviolet radiation, the existence of such
photolytic reactions in the Archaean atmosphere
implies a lack of ozone and oxygen, which are the
principal atmospheric components that absorb ultra-
violet radiation. According to the photochemical
model of Pavlov and Kasting (2002), the preservation
of the observed mass-independent S isotope fractio-
nation is only possible in an atmosphere with O2
concentrations less than 10�5 times the present
atmospheric level. Furthermore, the preservation of
the mass-independent S isotopic signatures is only
possible in the absence of large-scale, homogenising,
mass-dependent bacterial S processing in a marine,
sulphate-rich reservoir (Farquhar et al., 2000). Thus,
the mass-independent S isotope fractionation that
characterises Archaean to Eoproterozoic sediments
provides a very strong, independent argument for an
anoxic atmosphere at those times.
Reducing atmospheric conditions must have pre-
vailed until at least 2.64 Ga as indicated by the pyrite-
rich Black Reef at the bottom of the Transvaal
Supergroup and probably lasted until at least 2.45
Ga, taking into consideration the uraninite-bearing
conglomerates of the Elliot Lake Group. A minimum
age for a reducing atmosphere is given by the 2.2 Ga
Hekpoort paleosol, for which lateritic weathering with
Fe-enrichment in the top zone has been documented
(Beukes et al., 2002)—in agreement with paleosols of
similar age in other areas, for which weathering under
a highly oxidising atmosphere is indicated (Holland
and Rye, 1997).
6.2. Neoarchaean hydrosphere
Having established that the Archaean atmosphere
was most likely reducing, one might intuitively jump
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 35
to the conclusion that the contemporaneous hydro-
sphere was equally reduced. Such a conclusion seems
not to be justified, however, because syngenetic barite
deposits as old as Palaeoarchaean, with good exam-
ples known from the 3.5–3.2 Ga Barberton Super-
group in the Barberton greenstone belt (for a
compilation of all known occurrences see Huston
and Logan, 2004), clearly evidence that sulphur was
present in its oxidised state in the form of sulphate as
dominant hydrous S-species at least in some parts of
the Archaean ocean. The main reductants supplied to
the Archaean hydrosphere by hydrothermal discharge
as well as metamorphic and weathering fluxes involve
Fe2+, H2, CO, H2S, and SO2. In this context it is
particularly interesting to scrutinise the distribution of
Fe-oxides, principally magnetite, Fe-sulphides, essen-
tially pyrite, and sulphates in the various sedimentary
environments. The Witwatersrand rock record pro-
vides some pivotal information to this effect that can
be summarised as follows: (1) Pyrite is stable in all
fluvial to fluvio-deltaic deposits, even in those that
experienced extensive sedimentary re-working and
thus prolonged exposure to the meteoric environment;
(2) various textural forms of evidently detrital pyrite
show a complex S isotopic composition; and (3)
magnetite is present instead of pyrite in most marine
shale deposits (Fig. 4).
The significance of the first criterion has already
been discussed in the previous section and it suffices
to state here that the above constraints on the
Neoarchaean atmosphere are equally applicable to
the meteoric hydrosphere and even the oceanic top
waters. Sulphur isotopic composition has been used
repeatedly to constrain Archaean O2 distribution (for
review of available data see Strauss, 2003). Palae-
oarchaean barite, at least some of which is supposedly
sedimentary, has a fairly uniform S isotopic compo-
sition (d34S=+2.7x to +8.7x) that matches the
composition of Palaeoarchaean pyrite from the Bar-
berton Supergroup (d34S=�3.1x to 8.8x), which, in
turn, corresponds to that of magmatic sulphur. Larger
variations in, and deviation from 0x, of d34S is
typical of younger sedimentary sulphides and is
usually ascribed to enhanced S isotope fractionation
between seawater sulphate and reduced sulphide,
accomplished by biological S recycling. However,
some authors pointed out relatively large variations in
d34S already in Archaean sediments (e.g. Ohmoto et
al., 1993; Shen et al., 2001), which prompted them to
postulate microbial sulphate-reduction to have taken
place as early as in Palaeoarchaean times. The S
isotopic composition of the rounded, evidently detrital
pyrite from the Witwatersrand corresponds to the
limited range typical of Palaeoarchaean pyrite. This
does not seem to support a model of extensive
microbial sulphate reduction, but in the absence of a
good control on the difference between the isotopic
composition of the S source (marine sulphate) and the
product (pyrite), no definitive conclusions on the role
of microbial sulphate reduction in the Archaean
hydrosphere can be drawn.
The third of the above criteria highlights an
apparent stratification with regard to deep ocean
waters and near surface or freshwater environments
during the Neoarchaean. Although the magnetite in its
current textural position is clearly metamorphic, it is
implausible to derive it from the oxidation of original
sulphide, bearing in mind the comparatively over-
whelming proportion of sulphide in the entire strati-
graphic column. Any post-depositional fluid that
percolated through the Witwatersrand Basin fill is
more likely to have been enriched in S-species from
the partial dissolution of the abundant pyrite in the
stratigraphic succession than to have been capable of
selectively oxidising a hypothetical primary sulphide
in the intercalated shale beds. This is clearly evidenced
by the common observation of secondary pyrite
overgrowths at many stratigraphic levels. Furthermore,
the shale beds were most likely those with the lowest
permeability and thus least likely to have been affected
by chemical change due to fluid circulation across
stratigraphic boundaries. Derivation of the magnetite
in these marine shale deposits from an oxide precursor
is therefore assumed (Frimmel, 1996).
As illustrated by Huston and Logan (2004), Fe
solubility in the system Fe–S–O is highest and lowest
in the magnetite and pyrite stability fields, respec-
tively (Fig. 10). Magnetite is only stable under
reduced conditions (ASO4/AH2Sb10�2.5) and very
low total sulphur concentrations, whereas pyrite can
be stable even under oxidising conditions (ASO4/
AH2Sb108), provided the total sulphur concentration
is high, i.e. close to that of modern seawater. In the
presence of Ba, the pyrite stability field decreases to
reducing conditions (ASO4bAH2S) as barite precip-
itates under relatively oxidising conditions (ASO4/
Fig. 10. Phase relationships and Fe solubilities in the system Fe–
Ba–S–O as a function of redox potential (shown as ASO4/AH2S)
and total sulphur concentration normalised to modern ocean water
composition at a temperature of 25 8C and pH=7.8 (from Huston
and Logan, 2004).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4636
AH2SN10�2). In contrast to barite, Ca-sulphates are
highly soluble and a considerable degree of evapo-
ration is needed to precipitate gypsum. No occurrence
of sedimentary sulphate is known to date from the
Witwatersrand. However, based on S isotopic evi-
dence, England et al. (2002b) concluded that some of
the rare ooid-like pyrite grains found in the Venters-
dorp Contact Reef represent the product of microbial
sulphate reduction. This might reflect sulphate in the
source area, such as barite that has been described
from the Palaeoarchaean Barberton greenstone belt
(see compilation by Huston and Logan, 2004).
The apparent lack of sulphates in the Witwaters-
rand rock record may be explained in various ways.
From a thermodynamic standpoint it should reflect a
decrease in redox potential and/or total S concen-
tration, or a lack of Ba. The latter possibility is not
considered further as the principal Ba source is
hydrothermal discharge, which is unlikely to have
been shut down over several hundred million years
during Witwatersrand sediment deposition. Whether
the lack of sulphate in the Witwatersrand rocks in
particular, and in Meso- to Neoarchaean rocks in
general, reflects a combined increase in oceanic Fe
and decrease in oceanic sulphate as well as total S
concentrations relative to the Palaeoarchaean ocean,
remains contentious and dependent on the interpreta-
tion of Palaeoarchaean sulphate deposits. An abrupt
change in ocean water chemistry from a stratified
ocean with sulphate-bearing top waters to a sulphate-
free, Fe-rich ocean has been suggested to have
occurred around 3.2 Ga (Huston and Logan, 2004)
and has been ascribed to heavy bombardment of
Earth’s surface by meteorites (Glikson, 2001). Palae-
oarchaean sulphate precipitation may well have been
restricted to isolated oases of marine evaporative
ponds (Pavlov and Kasting, 2002; Shen et al., 2001)
and not representative of the world ocean. The lack of
comparable deposits in the Witwatersrand succession
could be merely a function of a lack of suitable
environments for such oases in the Witwatersrand
Basin at that time, and of preservation as evaporite
deposits would be prone to erosion especially in a
tectonically active foreland/retroarc setting.
The presence of magnetite in many of the
Witwatersrand shales mirrors a global surge in iron
formation occurrences during the Meso- to Neo-
archaean (Trendall, 2002), which reflects a reduced
ocean, in which high concentrations of Fe2+ were
possible in the bottom waters. The total S concen-
tration in that ocean must have been less than 10�5
that of modern ocean water (Fig. 10). In contrast, the
prevalence of pyrite in fluvial to shallow marine
Witwatersrand deposits points to relatively higher
total S concentration and/or higher O2 levels (Fig. 10)
in the oceanic top waters and the meteoric realm.
One of the magnetite-rich shale beds in the West
Rand Group rests above diamictite deposits (Corona-
tion Formation, Fig. 4). A causal link between iron
formation and global ice age, as suggested for the
younger Neoproterozoic iron formations, which are
viewed as result of isolation of the oceans from the
atmosphere by global, or near-global, ice cover
(Kirschvink, 1992; Klein and Beukes, 1993), may
therefore be applicable also to the magnetite-rich shale
beds of the Witwatersrand. According to that model,
melting of the ice cover would have triggered Fe-
oxide precipitation following hydrothermal Fe-enrich-
ment during glaciation. It should be noted, however,
that a number of magnetite shale beds in the West
Rand Group are not associated with diamictite.
Furthermore, the above model is not universally
accepted, not even for the Neoproterozoic deposits
(for critical assessment of existing evidence see
Young, 2004). Consequently, glaciation might have
further facilitated the deposition of iron formation
(magnetite shale) in the Meso- to Neoarchaean, but
was probably not the primary cause of it.
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 37
6.3. Quantifying Neoarchaean oxygen levels
Apart from the mentioned preservation of mass-
independent S isotope fractionation in pre-2.3 Ga
sediments, the two main constraints on the O2
concentration in the palaeo-atmosphere between 3.0
and 2.6 Ga are based on the detrital mineralogy of
placer deposits and on biochemical data on early
microfossils. As pyrite requires even lower oxygen
fugacity to be stable than uraninite, it is the
preferred phase for setting an upper limit on ancient
O2 levels. A lower limit is given by lipid biomarker
data from sedimentary rocks of the Pilbara Craton
in Western Australia. They provide evidence of
oxygenic photosynthesis at least as early as 2.7 Ga
(Brocks et al., 1999), and a lower limit of at least
1% of present atmospheric O2 is set by the
presence of steranes, found in these rocks, because
eukaryotic steroids require free oxygen (Jahnke and
Klein, 1983). Although some workers have sug-
gested the presence of aerobic bacteria already in
Palaeoarchaean times, the evidence for that is
problematic and a question of debate (Canfield
and Raiswell, 1999).
Fig. 11. Rate of oxidation of Fe2+ to Fe3+ in a solution of modern seawat
oxygen fugacity ( fO2), pH and temperature (T): (A) pH of modern seawat
modern atmospheric O2 pressure corresponds to fO2=10�0.67; calculated u
In order to quantify the redox conditions that are
required to explain the presence of pyrite in the
given environments, the physico-chemical conditions
for, and the rate of, oxidation of divalent to trivalent
iron were calculated (Fig. 11). Considering that
pyrite was also stable in Meso- to Neoarchaean
shallow marine environments, a modern seawater
composition (pH=8.22) was used for the initial
calculations. Under these conditions, oxidation of
Fe2+ to Fe3+ would be effectively instantaneous at
fO2 of 10�3, very slow at fO2 of 10�7 and
impossible at fO2 of less than 10�8 at a temperature
of 25 8C, with two orders of magnitude lower fO2
required for comparable reaction rates at a temper-
ature of 50 8C (Fig. 11A). This result is incompatible
with the above biochemical limit. As indicated by
the geochemical data, the atmosphere at the time of
interest was likely acidic. A lower pH of the
contemporaneous seawater would be more in line
with the combined evidence from detrital pyrite and
eukaryotic steroids. At a pH of 6.0, which would
correspond to the inferred silicate weathering to
kaolinite and the partial survival of detrital musco-
vite, oxidation of Fe2+ to Fe3+ would be very rapid
er composition as a function of time, in dependence of atmospheric
er, (B) pH=6; left panel—T=25 8C, right panel—T=50 8C; note thatsing data generated from PHREEQC (Parkhurst and Appelo, 1999).
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4638
at fO2 of 10�0.67—the modern atmospheric O2
level—and extremely slow at fO2 of 10�3—the
biochemical limit (Fig. 11B). Consequently, the
biochemical fO2 limit of 1% present atmospheric
level cannot have been exceeded significantly.
7. Conclusions
The Witwatersrand gold fields in South Africa,
which are the world’s largest gold producing
province, hold important keys for understanding
Archaean atmospheric and hydrospheric evolution.
Crucial to the debate around atmospheric O2 levels
at that time is the genesis of redox-sensitive
minerals that are associated with the gold.
Although most of the gold appears as a precipitate
within, or associated with, post-depositional hydro-
thermal phases and along microfractures, available
microtextural, mineralogical, geochemical and iso-
topic data, as well as the macro-scale stratiform
distribution of the ore bodies and its strong
sedimentological control, all indicate that this
hydrothermal gold, analogous to the associated
pyrite and uraninite, was derived from the local
mobilisation of detrital particles. Some of the key
pieces of evidence are a significant correlation
between gold and other heavy minerals as well as
sedimentary lithofacies, local preservation of in-situ
micro-nuggets with well preserved delicate textures
indicative of aeolian abrasion, compositional heter-
ogeneity on a microscale, and radiometric age data
that indicate an age of the gold, pyrite and
uraninite (3.03 Ga) that is older than the maximum
sedimentation age for the host sediment (2.90 Ga).
None of these observations/data is compatible with
a hydrothermal model, in which auriferous, poten-
tially sulphidising fluids were introduced from an
external source into the host rocks after sediment
deposition. In contrast, those arguments, used in
favour of hydrothermal models, emphasise the
microtextural position of most of the gold, which
highlights the undisputed hydrothermal nature of
much of the gold in its present position, but does
not explain the ultimate source of that gold.
Similarly, microtextural features, S isotopic
heterogeneity within and between grains, as well
as direct dating by the Re–Os method indicate that
the by far most abundant morphological variety of
pyrite in the Witwatersrand deposits, i.e. rounded
pyrite, is detrital. The same applies to rounded
uraninite, which is responsible for the Witwaters-
rand, together with the siliciclastic deposits of the
2.45 Ga Elliot Lake Group, representing the
world’s largest inferred U resource. Mineral chem-
ical characteristics and particularly variability
between grains, together with direct dating by the
U–Pb method, provide independent evidence of its
detrital nature. As with the gold, both pyrite and
uraninite were mobilised during post-depositional
fluid–rock interaction to variable degree, whereby
partial to complete oxidation to brannerite affected
the detrital uraninite, especially at higher strati-
graphic levels.
The currently available data point to mechanical
transport of gold, pyrite and uraninite from eroded
source regions that bear similarities to granitoid–
greenstone terranes currently exposed to the north
and west of the Witwatersrand Basin. The propor-
tion between felsic and mafic/ultramafic source
rocks varies strongly between and within reefs,
and this explains the only poor correlation between
Au and elements that are concentrated in detrital
minerals. Sediment deposition took place initially in
a continental rift (Dominion Group), followed by a
passive margin (West Rand Group), and then in a
foreland setting relative to the collision between the
Witwatersrand and Kimberley crustal blocks, with a
subsequent change to a retroarc position in con-
sequence of oceanic basin closure to the north of
the craton (Central Rand Group). Most of the gold
accumulated in the foreland/retroarc setting as that
setting favoured particularly intense sediment-
reworking on a series of erosional unconformities.
The recognition that rounded pyrite and uranin-
ite, both of which are characteristic of 3.07 to 2.64
Ga fluvial to fluvio-deltaic siliciclastic sediments on
the Kaapvaal Craton are not the products of post-
depositional hydrothermal fluid infiltration, as sug-
gested repeatedly in the past (e.g. Barnicoat et al.,
1997; Phillips and Law, 2000; Phillips and Myers,
1989), but detrital components of numerous placer
deposits re-affirms their pivotal importance for
constraining the evolution of O2 concentrations in
the Archaean atmosphere and hydrosphere. The lack
of Fe-oxides/hydroxides and the survival of pyrite
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–46 39
and uraninite during fluvial transport strongly
support a reducing Meso- to Neoarchaean atmos-
phere. In contrast to the fluvial to near-shore coarser
grained deposits, marine shale deposits contain
magnetite instead of pyrite. Comparably reducing
conditions are therefore inferred also for the oceanic
bottom waters, but with orders of magnitude lower
total sulphur concentrations than in modern ocean
water. Thus the observed distribution of pyrite and
magnetite in the Witwatersrand strata are in support
of the hypothesis, largely based on sedimentological
and S isotopic evidence, that Fe2+ was the principal
oceanic redox buffer prior to 2.4 Ga, whereas after
1.8 Ga, following an intervening transition period,
sulphate took over that role.
During the Meso- to Neoarchaean Aeras, total S
levels in the oceans must have been extremely low.
Sulphate-reducing bacteria have been inferred from as
early as Palaeoarchaean times (Shen et al., 2001). Any
microbial sulphate reduction, together with hydro-
thermal inorganic sulphate reduction, would have
removed sulphate rapidly to precipitate Fe-sulphides
from Fe-rich ocean waters. Consequently, the oceans
during Witwatersrand times would have been essen-
tially free of sulphate. Furthermore, the lack of
oxidative weathering of terrestrial pyrite would have
prevented the supply of sulphate to the oceans at a rate
that was greater than sulphate removal by Fe-sulphide
precipitation.
Most other evidence used for constraining
Archaean atmospheric O2 concentrations, such as
mass-dependent S isotope fractionation between
Archaean sulphate and sulphide, chemical compo-
sition of gold grains, and geochemical character-
istics of inferred paleosols, are not as conclusive as
the abundant occurrence of detrital pyrite and
uraninite in sediments that were laid down over
extensive areas on several cratons. Independent
support for a reducing Archaean atmosphere comes
from mass-independent S isotope fractionation. An
acid atmosphere, inferred from geochemical data,
was probably in equilibrium with a correspondingly
acid ocean. Kinetic calculations of the oxidation
from Fe2+ to Fe3+ show that a pH of 6 is required
in order to explain both the survival of detrital
pyrite and the presence of eukaryotic steroids in
Neoarchaean sediments. The calculated fO2 of 10�3
is in good agreement with the atmospheric evolu-
tion suggested previously by Kasting (1987, 2001;
Fig. 1A: curve a).
In summary, the available data endorse an acidic
hydrosphere beneath a reducing atmosphere during
the Archaean and early Palaeoproterozoic. Such an
acid environment requires elevated concentrations of
greenhouse gases, predominantly CO2. In addition,
the postulated anoxic atmosphere would have
favoured methanogenic bacteria that could contribute
to elevated atmospheric CH4 concentrations, in
agreement with the available C isotope record for
the Archaean (Pavlov et al., 2001b). Thus the
Archaean atmosphere was likely to be enriched in
effective greenhouse gases that would have effi-
ciently offset the lower solar luminosity in the early
history of Earth as suggested by Walker et al.
(1983). Such a palaeoclimate model explains the
mineralogy of ancient placer deposits. Variably
modified palaeoplacer deposits are known from the
Mesoarchaean to the Palaeoproterozoic from a
number of cratons, but only those older than about
2.4 Ga contain detrital pyrite and uraninite, whereas
in the younger deposits, detrital sulphides and
uraninite are conspicuously lacking and Fe-oxides
occur instead.
Acknowledgements
The author is indebted to Lawrie Minter who
has never hesitated in sharing his enormous
experience on the sedimentology and economic
geology of the Witwatersrand and related deposits
and who has given unrestricted access to his
valuable sample collection, some of which has
historic value as it contains material from mined
out areas that are not accessible anymore. The work
presented is based on numerous visits to under-
ground mines and core yards, which would have
been impossible without the cooperation of a
number of mining companies. Of particular impor-
tance for the conclusions presented in this paper
was the logistic support granted by Anglogold and
its staff. J.B. Maynard is thanked for a constructive
review of the manuscript. Parts of the work were
funded through grants from the South African
National Research Foundation and the University
of Cape Town.
H.E. Frimmel / Earth-Science Reviews 70 (2005) 1–4640
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Hartwig E. Frimmel, PhD, is Professor at
the University of Wqrzburg, Germany,
where he heads the Institute of Mineralogy.
He is also Professor at the Department of
Geological Sciences, University of Cape
Town, where for the past 15 years he has
worked inter alia on the genesis of the
Witwatersrand gold deposits. Other major
research interests include the relationship
between plate tectonics, palaeo-climate and
syn-sedimentary ore-forming processes,
with particular focus on the Neoproterozoic Aera, and the geo-
dynamic evolution of Precambrian supercontinents, with regional
emphasis on Antarctica, Africa and South America. Since 1998 he
has been the leader of the Earth Science subprogramme within the
South African National Antarctic Programme. He has served on
several editorial boards, supervised numerous post-graduate stu-
dents and has over 80 publications in international journals and
books to his credit.