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Archaean to Palaeoproterozoic high-grade evolution of the Belomorian eclog-ite province in the Gridino area, Fennoscandian Shield: Geochronologicalevidence
Ksenia A. Dokukina, Tatiana V. Kaulina, Alexander N. Konilov, Michael V.Mints, Konstantin V. Van, Lev Natapov, Elena Belousova, Sergey G. Simakin,Elena N. Lepekhina
PII: S1342-937X(13)00105-6DOI: doi: 10.1016/j.gr.2013.02.014Reference: GR 1017
To appear in: Gondwana Research
Received date: 19 April 2012Revised date: 2 February 2013Accepted date: 27 February 2013
Please cite this article as: Dokukina, Ksenia A., Kaulina, Tatiana V., Konilov, AlexanderN., Mints, Michael V., Van, Konstantin V., Natapov, Lev, Belousova, Elena, Simakin,Sergey G., Lepekhina, Elena N., Archaean to Palaeoproterozoic high-grade evolution ofthe Belomorian eclogite province in the Gridino area, Fennoscandian Shield: Geochrono-logical evidence, Gondwana Research (2013), doi: 10.1016/j.gr.2013.02.014
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Archaean to Palaeoproterozoic high-grade evolution of the Belomorian
eclogite province in the Gridino area, Fennoscandian Shield:
Geochronological evidence
Authors:
Ksenia A. Dokukina1,2, Tatiana V. Kaulina3, Alexander N. Konilov1,4, Michael V. Mints1,
Konstantin V. Van4, Lev Natapov5, Elena Belousova5, Sergey G. Simakin6, Elena N. Lepekhina7 1 - Geological Institute RAS, Pyzhevsky lane 7, Moscow, Russia; 2 - Lomonosov Moscow State University,
Moscow, Russia; 3- Geological Institute of the Kola Scientific Centre RAS, Apatity, Russia; 4 – Institute of Experimental Mineralogy RAS, Chernogolovka, Moscow Region, Russia ; 5 - GEMOC ARC National Key Centre, Department of Earth and Planetary Sciences, Macquarie University, NSW 2109, Australia; 6 - Yaroslavl’ branch of the Physical Technological Institute RAS, Yaroslavl, Russia; 7 - Center for Isotopic Research, Karpinskii All-Russia Research Institute of Geology, St. Petersburg, Russia
ABSTRACT
The Belomorian eclogite province was repeatedly affected by multiple deformation
episodes and metamorphism under moderate to high pressure. Within the Gridino area, high
pressure processes developed in a continental crust of tonalite-trondhjemite-granodiorite (TTG)
affinity that contains mafic pods and dykes, in which products of these processes are evident
most clearly. New petrological, geochemical and geochronological data on mafic and felsic
rocks, including PT-estimates, mineral chemistry, bulk rock chemistries, REE composition of the
rocks and zircons and U-Pb and Lu-Hf geochronology presented in the paper make it possible to
reproduce the magmatic and high-grade metamorphic evolution in the study area. In the
framework of the extremely long-lasting geologic history recorded in the Belomorian province
(3-1.7 Ga), new geochronological data enabled us to define the succession of events that includes
mafic dyke emplacement between 2.87 and 2.82 Ga and eclogite facies metamorphism of the
mafic dykes between ~2.82 and ~2.72 Ga (most probably in the time span of 2.79-2.73 Ga). The
clockwise PT path of the Gridino association crosses the granulite- and amphibolite-facies PT
fields during the time period of 2.72 Ga to 2.64 Ga. A special aspect of this work concerns the
superposed subisobaric heating (thermal impact) with an increase in the temperature to granulite
facies conditions at 2.4 Ga. Later amphibolite facies metamorphism occurred at 2.0-1.9 Ga. Our
detailed geochronological and petrological studies reveal a complicated Mesoarchaean-
Palaeoproterozoic history that involved deep subduction of the continental crust and a succession
of plume-related events.
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Keywords:
Mesoarchaean-Palaeoproterozoic Belomorian province
eclogite
granulite
thermal impact
continental subduction
1. Introduction
It has been long recognized that eclogite facies metamorphic regimes could not existed in
the Archaean due to the inferred higher geothermal gradients (Green, 1975; Baer, 1977; Labrosse
and Jaupart, 2007 and others). Palaeoproterozoic crustal eclogites dated at ~2.0 Ga were found in
the Usagaran belt of Tanzania (Müller et al., 1995; Collins et al., 2004) and were until recently
considered to be the oldest. At present, the time of the start of subduction-style processes is
widely discussed in the literature (e.g., Zhang et al., 2006; Kröner et al., 2006, Saraiva et al.,
2009; Brown, 2009, 2010). Summarizing evidence from subduction proxies, such as ophiolites,
blueschists, and ultrahigh-pressure metamorphic terrains, Stern (2005) suggested that the modern
style of subduction tectonics began in Neoproterozoic time. However, Archaean eclogite
xenoliths have been found in kimberlites (Pearson et al., 2003). Geodynamical and geochemical
features of the Achaean assemblages provide evidence that episodic steep subduction could
operate in the Archaean (van Hunen and Moyen, 2012 and references therein).
The first attempt to date the Belomorian eclogites was reported by Volodichev et al.
(2004), who described small eclogitic bodies near the Karelian village of Gridino on the western
shore of the White Sea. Based on geological mapping and the results of reflection seismic
profiling, this area was interpreted as the South Kola active margin adjacent to the northeastern
boundary of the Belomorian accretion orogen (Fig. 1, inset). Two Meso-Neoarchaean eclogite
associations were recognized within Belomorian TTG gneisses: (1) the subduction-type Salma
association and (2) Gridino eclogitized mafic dykes (collision-related eclogites) crosscutting the
crust of the South Kola active margin (Mints et al., 2010a, 2010c; Dokukina and Konilov, 2011,
see also Mints et al. in this issue). The HP/UHP processes in the Gridino area developed within
continental crust of TTG affinity and are most readily evident in mafic enclaves and dykes
(Dokukina and Konilov, 2011). These data imply that plate tectonic processes operated, at least
locally, in late Mesoarchaean time (Mints et al., 2010a).
Many recent works were aimed at defining certain key issues in Early Precambrian
geodynamics through the study of the Belomorian eclogite province: the ages of the eclogite
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facies metamorphic events (Berezin et al., 2012; Dokukina et al., 2009; 2010, 2012; Kaulina et
al., 2010; Mints et al., 2010a, 2010b, 2010c; Shchipansky et al., 2012b; Skublov et al., 2010a,
2010b, 2011a, 2011b, 2012; Slabunov et al., 2011; Volodichev et al., 2004, 2009, 2012); the
petrology and geochemistry of eclogite facies rocks and the nature of the original protolith
assemblages (Dokukina and Konilov, 2011; Konilov and Dokukina, 2011; Konilov et al., 2011;
Kozlovskii, 2010; Morgunova and Perchuk, 2012a, 2012b; Perchuk and Morgunova, this
volume; Shchipansky et al., 2012a; Volodichev et al., 2004, 2008) and some others. A number of
non-subduction models of the “autonomous” eclogite-facies metamorphism of the Belomorian
assemblages has been suggested: overpressure at deformation under amphibolite facies
conditions (Travin and Kozlova, 2005, 2009), autoclave effect within the magmatic camera
under amphibolite facies conditions (Volodichev et al., 2008), igneous crystallization of
omphacite (Sibelev, 2007), and a metasomatic scenario involving low silica activity in the
amphibolite facies metamorphic event (Kozlovskii and Aranovich, 2010). General problems of
Archaean geodynamics (Mints et al., 2010b, 2010c and in this volume; Shchipansky et al.,
2012b; Slabunov et al., 2006) and evidence of Archaean ultrahigh-pressure metamorphic
conditions (Dokukina and Konilov, 2011; Konilov and Dokukina, 2011; Konilov et al., 2011;
Morgunova and Perchuk, 2012b; Perchuk and Morgunova, this volume; Shchipansky et al.,
2012a, 2012b) were discussed based on data from the Belomorian eclogite province.
Currently one of the main problems regarding these rocks is the number and age of
eclogite facies metamorphic events recorded in the rocks of the Gridino assemblage. Some
authors consider two stages of eclogite facies metamorphism: (1) the Archaean stage (~2.72 Ga)
associated with the subduction of an oceanic slab is recorded in mafic lenses and pods included
in the TTG gneiss matrix (symplectitic eclogite boudins from Stolbikha Island (Volodichev et
al., 2004, 2012 and Slabunov et al., 2006)) and (2) the Palaeoproterozoic stage of “autonomous”
eclogitization of Palaeoproterozoic mafic dykes that occurred at ~2.4 Ga (various scenarios were
suggested by Volodichev et al. (2008), Sibelev (2007), Travin and Kozlova (2005, 2009),
Kozlovskii and Aranovich (2008)). The hypothesis of single late Palaeoproterozoic eclogite
facies event (~1.9-1.8 Ga) has been put forth by Skublov et al. (2010a, 2010b, 2011a, 2011b,
2012). A more detailed description of these models, their justification and criticism are presented
by in Mints et al. (this issue).
This article is centered; first and foremost, on present the principal results of
geochronological, geological, geochemical and petrological studies of the Gridino
eclogite/granulite facies assemblage. The geochronological data permit us to identify the
Archaean time of mafic dyke intrusions and to constrain the Archaean time span of eclogite
facies metamorphism of Gridino association. We have not yet found any zircon that
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unambiguously corresponds to the eclogite facies event. We consider regard the time span
between the magmatic and retrograde events as brackets for the age of eclogite facies
metamorphism. The second purpose of this paper is to point out the phenomenon of 2.4 Ga age
zircons in the eclogitized mafic dykes.
2. Geology
The Archaean Karelia and Kola continents, consisting predominantly of granite-
greenstone terrains and the Belomorian accretionary-collision orogen, are the major tectonic
units of the eastern Fennoscandian Shield (Fig. 1, inset). The Belomorian tectonic province is a
NW-trending segment of the Archaean nucleus of the Fennoscandian Shield, which is noted for
repeated episodes of intense deformation and high- and moderate-pressure metamorphism during
both Archaean and Palaeoproterozoic times. It has been suggested that the Belomorian tectonic
province represents a long-lived mobile belt that formed along the eastern margin of the Karelia
continent as a result of westward subduction (in present day coordinates) of the Archaean
oceanic lithosphere beneath the Karelian continent, accretion of island-arc complexes to the
Karelian margin and final collision at approximately 2.75–2.65 Ga (Bibikova et al., 1999; Miller
et al., 2005; Slabunov et al., 2006; Slabunov, 2008). The Keret' and Khetolamba granite-
greenstone units, which are the main components of the Belomorian province, are separated by
mafic–ultramafic Central Belomorian greenstone belt dated at 2.88–2.85 Ga (Bibikova et al.,
1999).
In contradiction with the above mentioned evolutionary model, the reflection seismic data
indicate that the Khetolamba microcontinent continues, at a certain depth, to the northeast,
beneath the Kola continent, although the boundary zone between two units, which coincides with
the Central-Belomorian belt, is not clearly pronounced in seismic images (Mints et al., 2009,
2010b; see also Mints et al. in this issue). Available geological, isotopic, and geochemical data
on mafic–ultramafic rocks of the Central-Belomorian greenstone complex are compatible with
its interpretation as tectonically disrupted and metamorphosed remnants of a Mesoarchaean
ophiolitic association (Slabunov et al., 2006). The Keret' tectonic nappe that thrust over the
Khetolamba unit and contains 3.00–2.70 Ga TTG-gneisses and greenstones, as well as numerous
eclogite bodies, which are the subject of this paper. Based on geological mapping and the results
of reflection seismic profiling, this nappe was interpreted as an active margin of the Archaean
Kola continent (Fig. 1, inset).
According to our understanding, these three units (the Khetolamba microcontinent, the
eclogite-bearing Keret' active margin of the Kola continent and the Central-Belomorian suture)
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compose a coherent Meso-Neoarchaen tectonic package that we refer to as the Belomorian
eclogite province. The Belomorian eclogite province is a NW-trending strip over 500 km long
and 50-60 km wide. It includes two eclogite associations: Salma in the southern Kola Peninsula
and Gridino in northeastern Karelia (Fig. 1, inset) (for more details, see Mints et al. in this issue
and references therein). The eclogite bodies are distributed within TTG gneisses of the South-
Kola active margin. Compositional and structural features of the subduction type Salma eclogites
suggest that the protoliths of the Salma eclogites encompass an association of gabbro, Fe-Ti
gabbro and troctolites, formed at ~2.9 Ga in a slow-spreading ridge setting (such as the
Southwest Indian Ridge) (Konilov et al., 2011; Mints et al., this volume and references therein).
The high-pressure processes in the Gridino area developed in a continental crust of TTG
composition and are inferred most certainly from the mafic enclaves and dykes (collision-related
eclogites) (Dokukina and Konilov, 2011).
The strongly retrogressed eclogites of the Salma association have been studied in detail at
Uzkaya Salma and Shirokaya Salma (Mints et al., 2010a, 2010b, 2010c; Konilov et al., 2011)
and at the Chalma locality (Kuru-Vaara quarry) (Shchipansky et al., 2012a, 2012b). The host
TTG gneisses vary in composition from quartz diorite to trondhjemite and contain a diverse
range of the mafic eclogites, layers and lenses of Fe-Ti eclogites and high-Mg eclogite facies
rocks (piclogites), layers and lenses of garnetites, and garnet-bearing and garnet-free
amphibolites. The banding of the gneisses and the contacts of the eclogite bodies dip steeply
towards the north-northwest. The Palaeoproterozoic granitoid veins and pegmatite dykes cut
across the Archaean rocks. The retrogressed eclogites typically consist of poikilitic garnet
porphyroblasts (4-5 mm) set in a fine-grained pale green matrix of pyroxene-plagioclase
symplectite (pseudomorphs after omphacite), and minor amphibole and quartz. Olive-green
hornblende, plagioclase, and ilmenite replace clinopyroxene-plagioclase symplectite, while
garnet is replaced by kelyphitic rims of plagioclase with hornblende.
Eclogitized mafic dykes of the submeridional Gridino dyke swarm occur in a coastal
zone and in a archipelago of the White Sea near the village of Gridino (Fig. 1). They were
illustrated in detail in earlier papers (Volodichev et al., 2004; Slabunov et al., 2006; Sibelev et
al., 2004, Travin and Kozlova, 2005; Stepanova and Stepanov, 2010). A tectonic mélange zone
some 50 km long and 10 km wide extends from north-west to south-east along the White Sea
coast. Field observations and geological mapping show that the mélange zone is associated with
a system of tectonic slices dipping northeastward. The mélange zone consists of a mixture of
migmatized granite gneisses with abundant amphibolite enclaves and of diverse ortho- and
subordinate paragneisses with numerous lenses of intensely deformed rocks. The «Archaean
eclogites» (according to Volodichev et al. (2004)) occur in a felsic gneissic matrix as strongly
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deformed pods and layers ranging from tens of centimeters to several meters. These eclogites
were interpreted as formed during the "warm" subduction of oceanic crust (Volodichev et al.,
2004).
The Gridino dyke swarm includes predominant gabbronorite and metagabbro dykes of
several types that differ in geochemistry and mineral assemblages: ferriferous gabbro, and
chromium-rich, magnesian gabbro and quartz- or olivine-bearing gabbronorite. Dykes vary in
thickness and range from undeformed to strongly deformed. Undeformed dykes crosscut the
felsic gneiss and have distinct intrusive contacts. Deformed dykes vary in degree of deformation,
including folding, boudinage and migmatization. Extreme deformation of the dykes led to their
breakup into pods and lenses conformable with the foliation of the host gneisses.
All mafic and felsic rocks of the Gridino area preserve evidence of eclogite facies
metamorphism at a minimum pressure of 16-17.5 kbar that was followed by near-isothermal
decompression. The retrograde P-T paths pass through high-pressure granulite (14-10 kbar and
800-750 °C) and amphibolite (7.9-9.6 kbar, 530-700 °C) facies fields (Dokukina et al., 2009;
Konilov and Dokukina, 2011). Petrological observations provided circumstantial evidence of a
much higher pressure of Belomorian eclogite metamorphism, which could reach ultrahigh
pressure condition (Dokukina and Konilov, 2011). This hypothesis was recently confirmed by
Morgunova and Perchuk (Morgunova and Perchuk, 2012; Perchuk and Morgunova, this issue),
who demonstrate that the PT metamorphic conditions in the Gridino area could reached the
coesite stability field.
The U-Pb ages of zircon extracted from a boudin of symplectitic eclogite in Stolbikha
Island (VGS-84: N 65º53’, E 34º51’) are 2720±6 Ma and 1920 Ma (Bibikova et al., 2003).
Zircon from a plagiogranite vein crosscutting eclogite-bearing gneiss in Stolbikha Island yields
an age of 2701±8 Ma. A Mesoarchaean U-Pb age of 2822±39 Ma for igneous zircons was
obtained from an eclogitized metagabbro dyke in Vargas Cape (VGS-84: N 65º56’, E 34º40’)
(Dokukina et al., 2009). An age of 2713±6 Ma was obtained for zircon from a high-pressure
felsic leucosome that penetrates the metagabbro dyke (Dokukina et al., 2012). Some of the dated
mafic dykes and one retrogressed eclogite boudin were found to contain a subordinate population
of zircons with an age of ca 2.4 Ga (Dokukina et al., 2009; Slabunov et al., 2011; Volodichev et
al., 2009, 2012), together with older Archaean zircons. Archaean zircon (from rocks containing
Palaeoproterozoic 2.4 Ga zircon) was traditionally interpreted as inherited from host rocks
(Slabunov et al., 2009; Volodichev et al., 2009, 2012). Palaeoproterozoic U-Pb, Sm-Nd, 40Ar/39Ar age values of ca 1.9-1.8 Ga are often obtained for the Gridino area (Bibikova et al.,
2003; Slabunov et al., 2006; 2011; Skublov et al., 2010 a; 2010b; 2011 a; 2011b; Dokukina et al.,
2012).
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Figure 1 shows a geological map of a representative key area close to the Gridino village.
Here the metamorphic basement is constituted by felsic gneiss that includes elongated eclogite
bodies, strongly retrogressed symplectite-bearing eclogites, mafic and felsic granulites and
amphibolites, garnet-clinopyroxene amphibolites, metaperidotites, metapyroxenites, and
microcline granites. The mafic enclaves have an elongate morphology and are cut by migmatite
veins. Mafic dykes of various composition crosscut the gneiss foliation (Figs. 1-3). All dykes
underwent high-pressure metamorphism to the eclogite and granulite facies; and the narrow
contact zones between the dykes and host gneiss show overprinted amphibolization.
3. Methods
Geochemistry. Major oxides of whole rock samples were analyzed, using ARL 9800 X-
ray fluorescence (XRF) spectrometer, in prepared Li2B4O7 glass pellets at the Karpinskii All-
Russian Geological Research Institute (VSEGEI), St. Petersburg. The XRF data were calibrated
against internationally certified standards. The accuracy and precision are better than ±2% for
major oxides.
Trace elements were determined by ICP-MS after digestion of the fused beads with
HF+HNO3. Pure solution external standards were used for calibration, and geological (USGS)
standards were used to monitor the analytical accuracy. The measurements were performed on
an ELAN-DRC-6100l ICP-MS at the VSEGEI. The precision for REEs was better than 5%,
whereas the precision for Rb, Sr, Ba, Nb, Ta, Zr, Hf, U and Th was better than 10%.
Mineralogy. Chemical analyses of minerals were made using a Tescan VEGA II xmu
scanning electron microscope equipped with both an Oxford Instruments WDS Energy 700 and
an EDS Energy 450 X-ray detection systems, at the Institute of Experimental Mineralogy of the
Russian Academy of Sciences, Chernogolovka, Moscow Region. Microprobe analyses of
coexisting minerals were carried out in polished thin sections after their petrographic
examination. The operating conditions were 20 kV accelerating voltage, beam current ~350 pA
on element (Co) for quant optimization, and a beam of 1–5 µm in diameter or, when it was
necessary to determine the composition of a cryptocrystalline mineral intergrowth, analysis in
area 20-30 µm in raster mode was performed. Acquisition time was 70 s.
Geochronology. The U-Pb isotopic system of zircon was analyzed on a SHRIMP II
secondary ion mass spectrometer at the Center for Isotopic Research, Karpinskii All-Russian
Research Institute of Geology. The intensity of the primary beam of negative molecular oxygen
ions was 5 nA, and the diameter of the spot (crater) was 25 µm. The data were processed with
the SQUID program (Ludwig, 2000). The U–Pb ratios were normalized to 0.668 in the
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TEMORA standard zircon corresponding to an age of 416.75 Ma (Black et al., 2003). The
uncertainty of individual analyses (ratios and ages) was given at 1σ level, the uncertainties of the
calculated concordant ages and concordia intercepts are given at 2σ. Concordia diagrams were
plotted using the ISOPLOT/EX programs (Ludwig, 1999). LA-(MC)-ICPMS U-Pb dating and
Lu-Hf isotope analyses are described in Supplementary methods.
Rare-earth and trace elements in zircon (at the same analytical spots as the U–Pb isotope
measurements) were analyzed with the Cameca IMS-4f ion microprobe at the Yaroslavl branch
of the Physical Technological Institute, Russian Academy of Sciences. Details of the
measurement technique are described in Supplementary methods.
4. Sample description
4.1. Petrology and mineralogy
4.1.1. Cape Gridin
Several large mafic metamorphosed dykes cut Cape Gridin (Figs. 1, 2). The thick dyke of
olivine metagabbronorite (Fig. 2a) examined in the course of our studies preserves relic igneous
textures and minerals and was metamorphosed but displays clearly visible chilled margins with
the host gneiss. A detailed petrographic description of the olivine metagabbronorite will be
presented below. The olivine metagabbronorite dyke is cut by a younger completely
metamorphosed Fe-metagabbro dyke. Relations between these dykes in Vorotnaya Luda Island
are shown in Figs.1 and 2b. The Fe-metagabbro dyke has a pinch-and-swell morphology with a
granite leucosome visible in a thin zone along its contact (Fig. 2a). The leucosome extensively
penetrates the Fe-metagabbro dyke from its contact with the metagabbronorite dyke (Fig. 2b). To
constrain the age of the dykes, we collected samples for geochronological studies from Fe-
metagabbro dyke (sample d44-4) and from the granite leucosome vein crosscutting this dyke
(sample d44-1) (Figs. 2c and 2d).
Fe-metagabbro dyke (sample d44-4) crosscuts the olivine metagabbronorite dyke (Fig. 2)
and invariably contains quartz. Neither relict igneous textures nor minerals have ever been
found. Generally, the Fe-metagabbro displays equilibrium of the garnet-clinopyroxene-
plagioclase high-pressure granulite assemblage. However, the eclogite stage mineral assemblage
is sometimes preserved. The rock matrix is an aggregate of clinopyroxene-plagioclase
symplectite typical of strongly retrogressed eclogites. Occasionally omphacite relics (with up to
36-42 mole % jadeite, (Supplementary Table S1, sample d54a) are preserved within the
symplectite and as inclusions in garnet. Omphacite in the matrix is of two types: (1) free of
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inclusions and (2) with oriented quartz needles (rods). Garnet is homogeneous, except only for a
mild increase in the iron content in edges, which can be explained by the development of
kelyphitic rims during the late metamorphic evolution.
The Fe-metagabbro dyke of Gridin Cape has a granoblastic texture composed by garnet,
clinopyroxene, plagioclase and quartz (± amphibole and biotite) (Fig. 3a). Some clinopyroxene
grains have a relict clinopyroxene core with an elevated content of the Ca-Tschermak end-
member (up to CaTs 9-14 mole % at 13 Jd mole %, Supplementary Table S1, sample d44/2).
Decreasing Al2O3 in clinopyroxene in the Pl-Qtz assemblage testifies that the pressure and the
temperature decreased (McCarthy and Patiño Douce, 1998).
Host gneiss. The mineralogy of quartzofeldspathic rocks is generally monotonous: these
are biotite, garnet-biotite or garnet-hornblende plagiogneiss of tonalite composition and are
typically migmatized (Dokukina and Konilov, 2011).
Granite leucosome. Sample d44-1 was collected from the granite leucosome vein
crosscutting the Fe-metagabbro and the olivine metagabbronorite dykes (Fig. 2c). The
mineralogy of the granite leucosome (Fig. 3b) includes garnet, biotite, plagioclase, K-feldspar,
quartz with minor epidote and scapolite. Scapolite develops next to aggregates of pyrite,
chalcopyrite and pentlandite. Garnet porphyroblasts commonly contain inclusions of quartz,
biotite, epidote, plagioclase and potassium feldspar. In one sample, two inclusions of titanium-
rich phengite in the garnet were found. The garnet is surrounded by a clinopyroxene corona.
Representative mineral compositions are presented in Supplementary Tables S1 and S2. The
clinopyroxene corona was thus formed under high-pressure granulite-facies conditions (see
Supplementary Tables S6) after the emplacement of the granitoid vein.
4.1.2. Northeast of the Gridino village
The northeastern termination of the village of Gridino occupies a cape consisting of
tonalitic granite-gneisses, which contain boudins of retrograde eclogites and are cut by
eclogitized gabbro and gabbronorite dykes (Fig. 4). The structure of the cape and its rocks are
described in detail by Volodichev et al. (2008, 2012).
Olivine metagabbronorite. Igneous structures and mineral associations are preserved in
the central part of thick dykes. Igneous mineral relics are olivine, augite, pigeonite,
orthopyroxene, chromite, biotite, plagioclase with minor spinel, corundum and potassium-
sodium feldspar. Olivine inclusions are preserved in augite and igneous orthopyroxene.
Individual olivine grains are generally rimmed by orthopyroxene coronas, or olivine is often
pseudomorphically replaced by a fine-grained granoblastic orthopyroxene-clinopyroxene±garnet
aggregate.
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Coronas of omphacite with a progressively increasing jadeite content developed around
igneous pyroxenes (Fig. 5a). The omphacite corona changes to a symplectite of omphacite with
plagioclase including corundum and/or spinel. The jadeite content in coronitic omphacite reaches
55 mole % (Supplementary Table S3). Secondary garnet coronas with inclusions of omphacite,
plagioclase and corundum intersect primary coronas of omphacite and symplectitic
pseudomorphs (Fig. 5a). In some of our samples, garnet of this type contains inclusions of
pargasitic amphibole. Analogous coronas around igneous minerals were also documented in
(ultra)high-pressure metamorphic complexes elsewhere (e.g., Lang and Gilotti, 2001).
The metagabbronorite is locally composed of the quartz-free rutile-garnet-orthopyroxene-
omphacite (Jd up to 24 mole %) assemblage (Fig. 5b). Granulite facies metamorphism is evident
from equilibrium in the garnet-orthopyroxene-clinopyroxene assemblage. It should be noted that
this assemblage can be distinguished from magmatic and syn-eclogitic ones mostly only based
on mineral compositions and PT estimates. An amphibolite assemblage within a zone up to ten
centimeters thick marks the contact of the metagabbronorite dyke with the host gneiss. Garnet-
rich millimeter-thick zones developed at direct contact of the dyke and wall rock. Overprinted
late amphibolization is a characteristic feature of all Gridino dykes.
A long-lived stripe of later magmatic recycling and deformation is localized in
northwestern part of the dyke (Fig. 4), has a NW-SE trend, is about 1-2 m thick and contains an
enderbite vein and metasomatic veinlets that were dated (see below). Thin amphibolite linear
zones postdate there the eclogite mineral assemblage of the metagabbronorite dyke and are made
up of garnet with low-temperature amphibole, orthopyroxene, clinopyroxene, plagioclase and
quartz (Fig. 5c, Supplementary Table S4d).
We distinguish two generations of orthopyroxenes: (1) relatively large orthopyroxene
grains included in amphibole, which probably crystallized before low-temperature amphibole,
and (2) small orthopyroxene grains in equilibrium with clinopyroxene and plagioclase, which
formed along margins between individual amphibole grains (Fig. 5c). In this situation and also in
contacts with garnet, the amphibole compositions change from actinolitic to low-Al hornblende
(Fig. 5c, d, Supplementary Table S4d). It follows that second generation of orthopyroxene grew
at a temperature increase (Supplementary Table S6).
Enderbite vein (samples 1111-06 and 1111-09) that intruded the eclogitized olivine
gabbronorite dyke in northeast of the Gridino village is an example of granitoid rock formed
under high-pressure granulite facies conditions. The boundary between eclogitized gabbronorite
and the enderbite vein is pronounced as a narrow zone of garnet-orthopyroxene-plagioclase
amphibolite (Fig. 6a). The marginal zone of the enderbite vein is about 1 cm thick, bears more
mafic minerals and is composed of an equilibrium garnet-clinopyroxene-orthopyroxene-biotite-
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plagioclase-quartz assemblage (sample 1111-06, Fig. 6a) or this assemblage without
orthopyroxene (sample 1111-09). The central part of the enderbite vein is kyanite-garnet gneiss
with rare clinopyroxene and orthopyroxene. Representative mineral compositions from
amphibolite, boundary and central zones of enderbite vein are presented in Supplementary Table
S5a, b.
Garnet from the central part of enderbite vein shows strong prograde zoning (Fig. 6b, c)
with a higher content of iron and calcium in grain cores (Fig. 6c, Supplementary Table S5a) and
contains inclusions of kyanite and omphacite (up to 21 mole % Jd) in its magnesian rims.
Kyanite was also found in plagioclase. Garnets from the amphibolite and enderbite boundary
zones are homogeneous except only a mild increase in the iron content in rims (Supplementary
Table S5a).
Clinopyroxene is contained in the enderbite vein as (1) symplectite-like structures with
plagioclase and (2) fine-grained granoblastic assemblage with orthopyroxene, garnet and biotite
in quartzofeldspathic matrix. Biotite contains up to 5.8 wt. % TiO2. Prograde rutile overgrows
Mg-bearing ilmenite, and an analogous corona of rutile develops around Mg-bearing ilmenite in
the amphibolized olivine metagabbronorite (mineral compositions are presented in
Supplementary Table S5a).
We collected two samples for geochronological studies: one from the root part of the vein
(sample 1111-06, Fig. 4b) and another from the continuation of the vein continuation in the
metagabbronorite body (sample 1111-09 was taken at a 10-m distance from sample 1111-06,
Fig. 4c).
Metasomatic veinlets. Sample 1111-08 for zircon dating was collected from veinlets
within a superimposed deformation zone that cuts quartz-free orthopyroxene-bearing eclogitized
olivine gabbronorite (Fig. 4d). The veinlets are composed of symplectitic eclogite with linear
quartz-biotite-plagioclase and quartz streaks (Fig. 7). Linear accumulations of garnet and rutile
developed along the quartz streaks, and orthopyroxene-plagioclase coronas were formed at
boundaries between quartz and garnet. Also, orthopyroxene composes clinopyroxene-plagioclase
symplectite and occurs in contact with potassium feldspar. Relics of omphacite (Jd up to 30 mole
%) remain in orthopyroxene-clinopyroxene-plagioclase symplectite; the garnet contains
inclusions of kyanite and omphacite (up to 34 Jd mole %). Representative mineral compositions
are given in supplementary table S4b. It is important to note that the symplectite contains Cl-
apatite (up to 6.4 wt. % Cl, see supplementary table S4c for composition), rutile, Ti-rich biotite
(up to 7 wt.% TiO2) and chains of zircon grains (Fig. 7). Veinlets most likely resulted from the
percolation of brine or melt through the eclogitic rock (see below).
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4.2. Pressure–temperature evolution
4.2.1. Igneous stage
The emplacement depth and temperature of the magma that formed the olivine- and
quartz-bearing gabbronorite dykes (Dokukina and Konilov, 2011), were estimated from the
composition of relict Ca-rich plagioclase (for the estimation of the pressure during igneous
crystallization) by the CaTs-An-Qtz geobarometer (McCarthy and Patiño Douce, 1998),
pigeonite geothermometer and orthopyroxene-clinopyroxene geothermometer (Fonarev et al.,
1991), and the graphic diagram in (Lindsley, 1983) for the equilibrium mineral assemblage of
igneous Opx+Pgt+Aug (see Dokukina and Konilov (2011) for details). Hot mafic melt (1030-
1200 °C) intruded an upper level of the Mesoarchaean crust under amphibolite-facies conditions
of approximately 5 kbar and 600 °C. Our results are consistent with the PT estimates by Egorova
and Stepanova (2011) for gabbronorite magma intrusions: 3.5-5.5 kbar and 1050-1200 °C.
4.2.2. Eclogite stage
Oriented quartz needles in omphacite are usually interpreting as an exsolution texture of
supersilicic clinopyroxene during the decompression of ultrahigh-pressure eclogite (Katayama et
al., 2000; Tsai and Liou, 2000). The integrated composition of the omphacite and quartz needles
is identical with that of dendritic clinopyroxene-plagioclase symplectite with “single grain”
structure and has 5-7 mole % of the Ca-Eskola end-member (Dokukina and Konilov, 2011). The
possibility of UHP conditions in Gridino rocks were discussed in much detail in (Morgunova and
Perchuk, 2012; Perchuk and Morgunova, this issue). The minimum pressure of 16-17.5 kbar was
obtained using the jadeite geobarometers (Holland, 1980). The quartz-free assemblage rutile-
garnet-orthopyroxene-omphacite (Jd up to 24 mole %) in the metagabbronorites (Fig. 5b) was in
equilibrium at a pressure of 22 kbar (Supplementary Table S6) according to the Grt-Opx
geobarometer of Harley (1984). No eclogite facies assemblages have ever been found in the
quartzofeldspathic rocks. This scarcity of eclogite facies assemblages in quartzofeldspathic host
rocks is typical of eclogite terrains. Quartz-rich rocks are more susceptible to dynamic
recrystallization and retrograde metamorphism than mafic rocks (Koons and Thompson, 1985).
4.2.3. Granulite stages
Orthopyroxene constitutes part of the clinopyroxene-plagioclase symplectite in mafic
rocks, which furnishes evidence of a retrograde metamorphic PT path of the granulite facies
(Page et al., 2003; Groppo et al., 2007). Orthopyroxene also forms a corona around garnet.
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Locally, orthopyroxene occurs as part of granoblastic orthopyroxene-clinopyroxene domains in
metagabbronorite and garnet-clinopyroxene-plagioclase assemblage (Fig. 3a) in Fe-metagabbro.
Detailed study of thin sections of mafic rocks led us to suggest that there were at least
two granulite-facies events in the history of Belomorian eclogite province: (1) a post-eclogite
high-pressure decompression stage at widely varying estimated PT conditions (10-13 kbar at
750-800 °C, see Supplementary Table S6), and (2) a stage of a subsequent heating overprint of
the rocks at a middle level of the crust. The second stage of the granulite-facies reworking can be
inferred from, e.g., the metagabbronorite dyke in the northwestern of Gridino village (Fig. 5c, d):
the prograde growth of orthopyroxene around actinolitic amphibole is evidence of a temperature
increase from 550-600 °C to 750 °C under nearly isobaric conditions of about 10 kbar (Fig. 5c,
Supplementary Table S6). The PT conditions of the origin of the metasomatic veinlets
correspond also to ~700-750 °C at 9-10 kbar (Supplementary Table S6).
In rare instances, granulite facies biotite-garnet or kyanite-garnet mineral assemblages
remain preserved in the quartzofeldspathic rocks. The PT metamorphic parameters of both the
metagabbro and the leucosome from Cape Gridin are estimated at ~700-750 °C and 10-12 kbar
(Supplementary Table S6), and those for the enderbite vein in the northwestern village of Gridino
are ~750 °C and 12.6-9.5 kbar (Supplementary Table S6), thus being indicative of high-pressure
granulites. In general the temperature and pressure characteristic of the mineral assemblages in
the quartzofeldspathic rocks correspond to retrograde amphibolite facies conditions.
4.2.4. Amphibolite stage
The contacts of the dyke with host gneiss are made up of an amphibolite-facies
assemblage within zones up to ten centimeters thick. Discernible garnetite zones developed
along boundaries between the amphibolite and felsic gneiss. Amphibolite-facies assemblages
were also formed in younger fractures, felsic pegmatite and carbonate veins. Overprinted
amphibolitization is a distinctive feature of all Gridino dykes. The temperature and pressure of
the amphibolite-facies reworking of the mafic and felsic rocks are 530-660 ºС and 7.9-9.6 kbar
(Volodichev et al., 2004; Sibilev et al., 2004, Dokukina and Konilov, 2011) (Supplementary
Table S6).
4.3. Geochemistry
4.3.1. Mafic dykes
Metagabbronorite dykes are subdivided into two types: silica-poor and quartz-bearing
rocks. The silica-poor gabbronorites are generally mesocratic rocks without quartz in all of the
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igneous and most of the metamorphic assemblages. The silica-poor mafic dykes and the quartz-
bearing gabbronorites have high contents of magnesium (Mg # 0.65-0.78), chromium (Cr 940-
2050 ppm) and titanium (TiO2 0.41-1.04 wt. %) (Tables 1, 2). The quartz-bearing gabbronorites
are richer in Al2O3 and CaO and poorer in MgO than the silica-poor mafic dykes. The silica-poor
dykes are enriched in Ni, Co, Ti, Cu, Sr, Zr, Nd, and depleted in V, Cr Zn, Pb, Ba, Y, Li, Sc
(Volodichev et al., 2005) relative to the quartz-bearing gabbronorite (Fig. 8a). Both types of
gabbronorite show LREE enrichment (La/Lu)N=2.5-12.34 without Eu anomaly Eu/Eu*=0.87-
1.05.
Fe-metagabbro composes both the oldest and the youngest mafic dykes, which have
similar geochemical and petrological characteristics. We studied only the younger Fe-
metagabbro dyke crosscutting the older olivine gabbronorite dykes. These geological
relationships are readily visible in Vorotnaya Luda Island and Gridino Cape (Fig. 2a, c, see also
(Berezin et al., 2012, Volodichev et al., 2012)). The whole-rock composition of the Fe-
metagabbro exhibits relatively narrow variations in the concentrations of major and trace
elements (sample d44-4, Tables 1, 2). These dykes of Fe-metagabbro composition have unusual
flat REE patterns similar to N-MORB (La/Lu)N=0.59-1.49, Eu/Eu*=0.90-1.09).
4.3.2. Quartzofeldspathic rocks
Host gneisses (Fig. 8b). The migmatized and non-migmatized plagiogneisses of tonalite-
trondhjemite composition underlie the bulk of the Gridino area. They contain 61-69 wt.% SiO2
and have #Mg in the range of 0.40-0.67; Na prevails over K (Na2O+K2O 5.12-6.35, Na2O 3.46-
4.77, K2O 0.75-1.93 wt.%), display a negatively sloped REE pattern with relatively high total
REE contents (63-162 ppm) and a positive Eu anomaly Eu/Eu*=1.03-2.46. Representative
analyses of the host gneisses (samples d44e, d44c4) are listed in Tables 1, 2.
Granite leucosome. Sample D44-1 has a typical granite composition and #Mg 0.55
(Tables 1, 2). The leucosome is strongly enriched in LREE ((La/Lu)N=39 relative to both the
gneisses and the dykes of Gridin Cape and has a positive Eu anomaly (Eu/Eu*=1.39) (Fig. 8b).
Enderbite vein. Samples 1111-06 and 1111-09 have a tonalite composition (Table 1) at
#Mg = 0.49-0.58. The chondrite-normalized REE pattern is characterized by an enrichment in
LREE (La/Lu)N=9-13. Sample 1111-06 (vein root) displays a tonalite composition depleted in
REE and LILE, enriched in HFSE and showing a positive Eu anomaly (Eu/Eu* = 1.21) relative
to sample 1111-09 (Eu/Eu* = 0.55) (Fig. 8c, Table 2).
Metasomatic veinlets, sample 1111-08. The compositions of the veinlets and olivine
gabbronorite are notably different. The veinlets typically have elevated contents of silica,
alumina and potassium and a lesser content of magnesium relative to olivine gabbronorite (Table
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1). The veinlets are significantly enriched in U, Th and LREE relative to the host orthopyroxene
eclogite rock that is cut by the veinlets (Fig. 8d, Table 2). Orthopyroxene eclogite and its olivine
gabbronoritic precursor are very rich in siderophiles (Cr, Ni, Zn up to several thousand ppm) are
mildly enriched in Ba, Rb, Sr relative to metasomatic veinlets (Table 2).
4.4. Geochronology
4.4.1. Gridin Cape
Fe-metagabbro (sample d44-4) is very poor in Zr. Nevertheless, we managed to extract
numerous zircon grains from this sample, and fourteen of them were analyzed for U-Pb on SHRIMP
II in their cores and rims. The zircons are brownish or colorless, elongate, subhedral, up to 150 µm in
width and up to 500 µm in length grains. Zircon grains show core and rim structure in CL. Some
grains have thin oscillatory zoning in CL and PPL images Fig. 9a) suggesting their igneous origin.
They have medium to high Th/U ratios (0.38-1.39), positive Ce (Ce/Ce* = 2.35-11) and negative Eu
(Eu/Eu* = 0.53-0.77) anomalies, and are enriched in HREE relative to LREE (LuN/LaN = 213-941)
(Fig. 10a, Table 4). Some zircon cores are darker in CL and display a higher U content (351-1321
ppm) at a relatively low Th content (46-91 ppm) and low Th/U ratio (0.05-0.18) (Tables 3, 4). The
chondrite-normalized REE pattern of these cores exhibits positive Ce (Ce/Ce* = 2-30) and negative
Eu (Eu/Eu* = 0.06-0.77) anomalies and enrichment in HREE relative to LREE (LuN/LaN = 309-
6702), which also suggest an igneous origin (see, e.g., Hoskin and Schaltegger, 2003) (Fig. 10b,
Table 4). Nine points define a discordia with the upper intercept at 2869±41 Ma and the lower
intercept at 1801±260 Ma (probability of fit = 0.96). The concordia age for five points is 2846 ± 7 Ma
(Fig. 9h).
The low-thorium, colorless rims (spots 1.2, 7.1, 8.1, 9.4, Fig. 9a-b) with low Th/U ratios 0.01-
0.38 are depleted in all trace elements except Hf (Table 4). They are characterized by positive Ce
(Ce/Ce* = 1.55-7) and negative Eu (Eu/Eu* = 0.28-0.64) anomalies and by a relatively flat chondrite-
normalized REE pattern (LuN/LaN = 63-940; LuN/SmN = 12-48) (Fig. 10c, Table 4). Four points
define a discordia with an upper intercept at 2777 ± 67 Ma (lower intercept – 1349+/-1100 Ga,
probability of fit =0.94). The concordia age for two points is 2780 ±20 Ma (points 1.2, 8.1, Fig. 9b, c,
h).
Granite leucosome, sample d44-1. Zircons from the leucosome were dated by SHRIMP II
(VSEGEI, St. Petersburg) and by LA-ICPMS (GEMOC ARC National Key Centre, Sydney,
Australia). SHRIMP data show spreading of the points along the concordia line with 207Pb/206Pb ages
from 2926 to 2764 Ma (8 points) (see Supplementary Figure 1). The histogram of the LA-ICPMS 207Pb/206Pb ages (23 points) (Fig. 9i) shows age peaks at 2845 ± 14, 2797 ± 11 and 2720 ± 28 Ma.
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Three zircon populations may be distinguished on the basis of their morphology and structure
in CL (Fig. 9d-g). Some of the zircons are colorless subhedral grains 100 x 250 µm, pale in CL with
sector or oscillatory zoning (for example, spots 1.1, 5.1 Fig. 9d, Supplementary Fig. S1). This zircon
fraction has high Th/U ratios (0.55-1.01) and ages about 3.0 Ga (SHRIMP II, Tables 3, 4). The
chondrite-normalized REE pattern shows positive Ce (Ce/Ce* = 8.67) and negative Eu (Eu/Eu* =
0.27) anomalies, enrichment in HREE relative to MREE (LuN/SmN = 116) and relative to LREE
(LuN/LaN = 4602) (Fig. 10d), and correspond to magmatic zircons according to Hoskin and
Schaltegger (2003).
Zircons of the second group are similar to the prevalent zircons of the Fe-metagabbro (sample
d44-4). These are brownish, elongate and short-prismatic subhedral grains 50-100 x 200-300 µm in
size, grey in CL, with coarse oscillatory zoning (Fig. 9e). They typically have a moderate Th/U ratio
(0.20-0.46) and positive Ce (Ce/Ce* = 2.04-3.23) and lower negative Eu (Eu/Eu* = 0.55) anomalies
on the relatively flat chondrite-normalized REE pattern (LuN/SmN = 38-41; LuN/LaN = 152-203) (Fig.
10d, Table 4). Some grains have dark cores (for example, spots 2.1, 4.1, Fig. 9f), which are rich in U
(377-1653 ppm) at a variable Th/U ratio (0.04-0.45) (Table 3), positive Ce (Ce/Ce* = 1.36-2.92) and
negative Eu (Eu/Eu* = 0.29-0.77) anomalies and mild enrichment in HREE relative to MREE
(LuN/SmN = 9-41) and LREE (LuN/LaN = 97-364) (Fig. 10e, Table 4). Age estimations for zircon
grains of this group correspond to an age peak at 2845 ± 14 Ma and are coeval with igneous zircons
from the Fe-metagabbro dyke.
The low-uranium, colorless rims (like spot 4.2, Fig. 9g) yielded an age of 2.78-2.79 Ga (Fig.
9i, Table 5). These exhibit variable Th/U ratios of 0.01-0.48, a positive Ce anomaly (Ce/Ce* = 2),
negative Eu anomaly (Eu/Eu* = 0.49), a relatively flat chondrite-normalized REE pattern (LuN/SmN =
25, LuN/LaN = 47) (Fig. 10f, Table 4), and they can be correlated with young rims from sample D44-
4.
The histogram of the 207Pb/206Pb ages demonstrates one more peak at 2720 ± 28 Ma (Fig. 9i,
Table 5). The meaning of this age will be discussed in the next section.
All of these grains have similar Hf isotope compositions (Fig. 11, Table 5) and show a
relatively narrow stripe of points with εHf close to the CHUR evolution line. Different 207Pb/206Pb
ages at similar Hf isotope ratios suggest that the U-Pb system was reset without disturbance of the Hf-
isotope composition.
4.4.2. Northeast of Gridino village
In order to ascertain whether eclogite facies metamorphism took place before 2.71 Ga, we
dated zircon from an enderbite vein crosscutting the metagabbronorite dyke (samples 1111-06, 1111-
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09) (Fig. 3) and zircon from the metasomatic veinlets (sample 1111-08) penetrated the same body in
superimposed deformation zone of the dyke (Fig. 3).
Garnet-orthopyroxene-clinopyroxene-biotite-kyanite enderbite, sample 1111-06. Zircon grains
from the enderbite can be subdivided into several types. The first fraction consists of subhedral 100 x
200 µm grains with thin oscillatory zoning (Fig. 12a), rather high Th/U ratios (0.55-0.85), which may
be characteristic of igneous zircon from intermediate to felsic rocks. The chondrite-normalized REE
pattern displays positive Ce (Ce/Ce* = 19-29) and negative Eu (Eu/Eu* = 0.33-0.52) anomalies,
enrichment in HREE relative to MREE (LuN/SmN = 73-155), and strong enrichment in HREE relative
to LREE (LuN/LaN = 1313-4153) (Fig. 13a). These zircons yielded a discordant age of 3014 ± 53 Ga,
N = 3 (Fig. 12f, Table 7). These grains were likely inherited from the tonalite.
The grains of the second type are subhedral, 100 x 300 µm, with distinct core-rim structures in
CL (Fig. 12b). Zircon cores display poorly pronounced oscillatory zoning, while the 5-30 µm wide
rims are black and homogeneous. Cores and rims exhibit indications of resorbtion during continuing
formation of the enderbite. The cores have relatively low Th/U ratio (0.15-0.16) and are enriched in
REE (ΣREE = 1497 ppm) (Table 4). The chondrite-normalized REE pattern is flat (LuN/SmN = 1.7,
LuN/LaN = 17) and has poorly pronounced anomalies of Eu and Ce (Eu/Eu* = 0.39, Ce/Ce* = 1.6)
(Fig. 13a). The zircon-core 207Pb/206Pb ages are 2830-2879 Ma. The black rims have a low Th/U ratio
(0.03-0.34). The chondrite-normalized REE pattern of this zircon shows a positive Ce anomaly
(Ce/Ce* = 3.5-4.2), negative Eu anomaly (Eu/Eu* = 0.34) and insignificant enrichment in HREE
relative to MREE (LuN/SmN = 23-29, LuN/LaN = 270-336) (Fig. 13a, Table 4). The concordia age of
black rims is 2743±56 Ma, N = 3 (Fig. 12f).
Only one small rim bright in CL (spot 7.1) gave a 207Pb/206Pb age of 1984±22 Ma (Fig. 12f).
We investigated trace element distribution in garnets from this sample in the “amphibolite”
and “enderbite” zones (Fig. 6). Garnet from the “amphibolite” zone is enriched in HREE (SmN/LaN =
254, YbN/LaN = 1191), has a low positive Eu anomaly Eu/Eu*= 1.13 and Th/U = 0.67 (Fig. 13b). The
core of garnet from the “enderbite” zone has an elevated content of LREE (SmN/LaN = 4, YbN/LaN =
104), negative Eu anomaly Eu/Eu*= 0.82, high Th/U ratio = 1.49. The “enderbite” garnet rim has a
lower LREE content (SmN/LaN = 36, YbN/LaN = 431), positive Eu anomaly Eu/Eu*= 1.38, Th/U =
0.53 (Fig. 13b). Garnet in the “boundary zone” is relatively enriched in LREE (SmN/LaN = 65,
YbN/LaN = 201), has a negative Eu anomaly Eu/Eu*= 0.78 and Th/U = 0.39 (Fig. 13b). All garnets
are depleted in LREE and HREE and do not demonstrate equilibrium with zircon in the same sample,
because DHREE(Zrn/Grt>>1), while it must be near 1 at equilibrium (for example, Rubatto, 2002)
Garnet-clinopyroxene-biotite-kyanite enderbite, sample 1111-09. The main zircon population
in this sample consists of rounded or ellipsoidal colorless or pale brownish grains of 100x100 or
100x200 µm with sector, fir-tree and diffuse oscillatory zoning as it is seen in CL (Fig. 12c). These
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zircons yielded an U-Pb near-concordant age of 2717±13 Ma, N = 7 (Fig. 12g, Table 7). The
chondrite-normalized REE pattern of these zircon grains has a positive Ce anomaly (Ce/Ce* = 17-
106) and negative Eu anomaly (Eu/Eu* = 0.18-0.36) and shows mild enrichment in HREE with
respect to MREE (LuN/SmN = 61-83, LuN/LaN = 1049-4415) (Fig. 13c, Table 4). The zircon contains
polycrystalline inclusions of omphacite-tschermak solid-solution (20 Jd mole % at CaTs 15 mole %),
phengite (3.21 Si pfu), biotite and silica (Fig. 14, Table 8).
One white in CL core (spot 9.1) gave a concordia age of 2824 ±31 Ma (Fig. 12g). This core
yielded an igneous-type of REE pattern: a positive Ce anomaly (Ce/Ce* = 21), a negative Eu anomaly
(Eu/Eu* = 0.18), weak enrichment in HREE relative to MREE (LuN/SmN = 40) and strong enrichment
in HREE relative to LREE (LuN/LaN = 1818) (Fig. 13c).
During a SHRIMP analytical session, we spent much effort to find youngest zircons, but only
one thin bright in CL overgrowth on a sector zoning (БЫЛО ПРАВИЛЬНО) zircon (spot 8.1) gave a
concordia age of 1916±33 Ma (Fig. 12e, g).
Metasomatic veinlets, geochronological sample 1111-08. A number of veinlets were sampled
(Fig. 3d), and their material was separated from the host orthopyroxene eclogite (metamorphosed
olivine gabbronorite) by hand picking. Fifty zircon grains belonging to a single population were
separated from about 430 g of the rock. The zircon grains have an unusual “clumpy” morphology
with cavities, which are traces of fluid inclusions (Fig. 15a). The zircon shows no CL emission.
Cross-sections of the zircon grains are split and have pole-like, rounded and amoeboid morphologies,
which suggest grown in solid rock in limited-growth condition. The zircon grains contain numerous
inclusions of minerals of previous crystallization and reworking stages of the metagabbronorite.
These are orthopyroxene, clinopyroxene, Ti-rich biotite, rutile, quartz, Cl-apatite – symplectite
minerals and those quenched from fluid (Fig. 15b, c, Table 8). Two types of domains can be
distinguished in the zircon grains: homogeneous high-Th domains and domains of overprinted
hydrous recrystallization.
Homogeneous areas have abnormally high concentrations of Th (2338-17700 ppm), U (5027-
8500 ppm), Y (5027-30000 ppm), Hf (7704-9750 ppm) and REE (up to 14500 ppm) (Table 3), high
Th/U ratio (1.0-2.8), an elevated content of Ti and incompatible elements, such Al, Fe Ca and P
(Tables 4, 9). The chondrite-normalized REE pattern shows positive Ce (Ce/Ce* = 21-152) and
negative Eu (Eu/Eu* = 0.06-0.32) anomalies, low enrichment in HREE relative to MREE (LuN/SmN =
7-20) and high enrichment in HREE relative to LREE (LuN/LaN = 405-11709) (Fig. 16a). The U-Pb
SHRIMP II concordant age of these domains is 2394±6 Ma, N = 4 (Fig. 16c, Table 7).
The recrystallized domains are mostly restricted to the peripheries of zircon grains. Zircon
within these domains is hydrated and has a reduced Th/U ratio (0.91-0.04) and concentrations of Th
(18-3810 ppm), U (188-3263 ppm), Y (874-4864 ppm) and REE (1307-4234 ppm) (Table 4), contains
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up to 10 wt.% H2O, and has higher content of Na, K, Ca, Fe (Table 9), sometimes a very high content
of Ti, and numerous micron-sized inclusion of thorite. The chondrite-normalized REE pattern is flat
(LuN/SmN = 5-327, LuN/LaN = 22-657), with a negative Eu anomaly and poorly pronounced Ce
anomaly (Eu/Eu* = 0.49-0.92, Ce/Ce* = 1.7-1.9) (Fig. 16b). The concordia age (SHRIMP II) of these
domains is 1886±10 Ma, N = 3 (Fig. 16c, Table 7).
5. Discussion
5.1. Geological and geochronological consequences
The Belomorian tectonic province was repeatedly affected by deformation and high- to
moderate-pressure metamorphism. Geochronological data for Gridino rocks point to both Archaean
and Palaeoproterozoic ages. Deciphering the relative and absolute ages of distinct tectono-thermal
events is a difficult task in this area and in the Belomorian province as a whole. Figure 17 displays a
histogram that shows all most concordant (discordance of 0-2%) of our geochronological samples and
the typical zircon grains of each age peak.
5.1.1. Dyke and leucosome of the Gridin Cape
The igneous 2.9 Ga zircon grain from the granite leucosome (sample D-44-1) is the oldest in
our sample. This grain displays indications of its igneous genesis, was likely inherited, and yields the
crystallization age of the host tonalite. It is important that 2.9 Ga zircon was found only in the
leucosome and never in the Fe-metagabbro (Fig. 9d). We will discuss oldest zircons from our samples
the section devoted to enderbite (samples 1111-06, 1111-09).
The compositions of the Gridino mafic igneous rocks generally fall in the fields of intraplate
and active continental margin magmas in tectono-magmatic discrimination diagrams (such as Ta/Yb
vs Th/Yb diagram by Gorton and Schandl (2000) and others). The normalized patterns of REE and
other trace elements in most of the different mafic dykes suggest their strong crustal contamination.
However, the Fe-metagabbro dykes studied in the Gridin Cape (sample D44-4) and those in Luda
Vorotnaya Island (Stepanov and Stepanova 2009, 2010), which show flat REE patterns similar to
those of N-MORB, are the only important exceptions. This led us to believe that it is hardly possible
to collect a representative amount of zircons inherited from the country rocks in these dykes.
Igneous elongate zoned zircon grains with lower U and Th contents from the Fe-metagabbro
(D44-4) and the granite leucosome (D44-1) yield 207Pb/206Pb ages of 2.87-2.85 Ga (Fig. 9a). The
concordant age is 2846 ± 7 Ma and likely sets the upper limit for igneous zircon crystallization. The
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crystallization temperatures for these zircons, calculated on the basis of Ti contents, are within the
interval of 667-819 ºC and are very low for mafic melt (Supplementary Table S7, Fig. 18). Earlier we
have estimated the emplacement level of the intrusion of mafic magma in the Gridino mafic dyke
swarm. The hot mafic melt (1030-1200°C) intruded an upper level of the crust under amphibolite
facies conditions (5 kbar and 600 °C), and the depth was never greater (Dokukina and Konilov,
2011). Thus 2.87-2.85 Ga zircons most likely reflect the metagabbro dyke intrusion age and were
formed in interstitials late during mafic melt crystallization, at the equilibration of the temperatures of
the mafic melt and host rock.
Igneous zircon grains of this type are widely spread in the granite leucosome D44-1 (Figs. 9e,
f). The leucosome develops only along the boundary between the tonalite gneiss and the metagabbro
dyke (Fig. 2a) and extensively penetrates the dyke outside its contact with the metagabbronorite dyke
(Fig. 2c). The boundary between the gneiss and dyke was obviously favorable for fluid migration,
which facilitated the partial melting of the tonalite gneiss. Thereby zircon from the mafic dyke could
be captured by the newly formed granite leucosome.
Low-thorium 2.79-2.78 Ga rims that overgrew the older igneous zircons (Figs. 9c, g) are
negligibly depleted in HREE and were probably formed from partial felsic melt (Hokada and Harley,
2004). The calculated crystallization temperatures of these zircon rims are 700-720 ºC
(Supplementary Table S7) and can probably characterize the formation temperature of the granite
leucosome under high-grade metamorphic conditions. The histogram (Fig. 17) shows an age peak at
2.7 Ga for sample D44-1. The effect of the 1.9 Ga event is also always evident during age
calculations and in the U–Pb concordia plots.
5.1.2. Enderbite vein of northeast of the Gridino Village
Enderbite sampled from the continuation of the vein (sample 1111-09) is negligibly more
ferrous and potassic than the root part of the enderbite vein (sample 1111-06) and is enriched in REE,
yields an Eu anomaly and is probably a product of a higher degree of the tonalite gneiss melting. The
difference between the compositions of the samples may be due to the degree of saturation of the
unmelted residual gneiss material in tonalite melt. In other words, the root of the enderbite veins
contained a sufficient amount of unmelted parent material and inherited its geochemical properties.
The additional mature liquid tonalite melt with a small proportion of residual material that penetrated
deep into the olivine gabbronorite body has already the properties of the neosome. Older oscillatory
3.0 Ga zircons are predominant in sample 1111-06 (the root part of the enderbite vein). These grains
have igneous properties and probably correspond to the igneous crystallization age of the country
metatonalite. Similar igneous zircons with an age of ca 3.0 Ga were separated from the metagabbro
dyke in Cape Vargas, which has an igneous age of 2.82 Ga (Dokukina et al., 2010, Dokukuna et al.,
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2012). The crystallization temperatures of these zircons gave a significant dispersion (for zircons 1.1
and 1.2 respectively 640 and 1050 ºC) (Fig. 12a, Supplementary Table S7). Evidently, these zircons
correspond to zircons from the host gneiss, and the age of 3.0 Ga corresponds to that of the igneous
tonalite protolith. The model Sm-Nd ages of the oldest Gridino gneissic rocks are also 3089-2973 Ma
(Dokukina et al., 2010).
We can only hypothesize about the significance of LREE-enriched resorbed zircon cores with
an age of 2.88-2.83 Ga. These ages may reflect a response of the tonalite gneiss to the thermal effect
of the emplaced gabbronorite magma, because these ages are similar to the age of the mafic dyke in
Gridin Cape. The estimated crystallization temperature of this zircon is 800-815ºC (Fig. 18, Table
S7).
The rims black in CL and having the 207Pb/206Pb ages of 2.72-2.69 Ga and/or a U-Pb
concordia age of 2743 ±5.6 Ma are poorer in HREE and generally have a low Th/U ratio (0.03-0.04,
except one 0.34), which is related to uranium enrichment . We assume that they correspond to the
initial melting of the country metatonalite and the derivation of the granitoid melt which intruded the
metagabbronorite dyke.
The predominant group of zircon grains from sample 1111-09 with an U-Pb age of 2717 ± 13
Ma (Fig. 12g, Table 7) is similar to igneous zircon in terms of their high Y content, strongly enriched
REE pattern, positive Ce and negative Eu anomalies. Zircon has soccer-ball morphology and sectorial
and fir-three zoning in CL, which are typical of granulite-facies zircon in the presence of partial melt
(Vavra et al., 1996; Rubatto, 2002; Corfu et al., 2003, Whitehouse and Kamber, 2003). It means that
magmatic and metamorphic events are coeval, which is common for granulite complexes, where the
formation of synmetamorphic granitoids is pronounced. And magmatic zircons in crystallizing rock
will be of the same age as metamorphic zircons in already existed rock. The age of 2717±13 Ma
corresponds to the age of high-pressure granulite stage processes in this region (see Mints et al., this
issue and references therein, Dokukina and Konilov, 2011, Dokukina et al., 2012). Zircon from the
enderbite vein contains polycrystalline inclusions of omphacite-tschermak solid-solution with
phengite, biotite and silica (Fig. 14, Table 8) and also marks the upper age limit for the eclogite
facies metamorphism (see, for instance, Kröner et al., 2006).
5.1.3. Metasomatic veinlets with age ca 2.4 Ga of northeast of the Gridino Village
Volodichev with coauthors (Volodichev et al., 2009, 2012; Slabunov et al., 2011, their
samples V16-65, V16-66 in a given papers) studied zircons from metagabbronorite rock by SHRIMP
II and SIMS and obtained four populations of U-Pb ages: (1) 2.84-2.72 Ga; (2) 2.72-2.70 Ga, (3) ca
2.4 Ga; (4) ca 1.9 Ga. These authors interpreted the Palaeoproterozoic age of 2393±13 Ma for high
Th, U, Pb and REE zircon as the time of gabbronorite magma emplacement and subsynchronous
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eclogitization. This zircon is characterized by a high Ti content and bears crystalline inclusion of
orthopyroxene, clinopyroxene, amphibole and biotite. These features suggest that the zircon
crystallized during the late magmatic stage of the gabbronorite (Volodichev et al., 2009, 2012;
Slabunov et al., 2011). The Archaean ages (2.84-2.70 Ga) were interpreted as inherited, and the
Palaeoproterozoic age of 1.9 Ga is interpreted as that of overprinted amphibolite facies
metamorphism.
Zircons in the veinlets described above show unusual properties that were produced under
very specific conditions. Igneous zircon generally contains less that 1 wt. % REE and Y (Hoskin and
Schaltegger, 2003). Zircon from the metasomatic veinlets has anomalously high concentrations of
REE and Y (REE+Y up to 4.5 wt. %) and high contents of Th (up to 1.8 wt. %), U, LILE (Ca, Ba,
Sr), Ti, Nb and other incompatible elements (Tables 4, 9). Such zircon generally occurs in the alkaline
intrusive rocks and forms in hydrothermal-metasomatic process (Hoskin and Schaltegger, 2003;
Hoskin, 2005). Poikilitic zircon of clumpy morphology contains mineral inclusions corresponding to
the host rock mineral assemblage that provides evidence that the zircon grew in a solid rock. The
veinlets were undoubtedly formed after eclogite-facies metamorphism because (1) their mineral
composition correspond to the post-eclogitic assemblage (Fig. 7, Suplementary Tables S3, S4); and
(2) the zircon contains a mineral inclusion of the post-eclogitic stage (Fig 15). The veinlets are
reached in Si, Al, K, P, Th, U, Zr, Hf, REE and Y relative to the host orthopyroxene-bearing eclogite
(Tables 1, 2).
The 2.4 Ga zircon has an REE pattern similar to that of igneous zircons (Hoskin and
Schaltegger, 2003; Rubatto, 2002). The best positive correlation in zircon composition is between
REE and Y (determination coefficient R2= 0.998). The comparison of our data and data on the same
zircons published by Slabunov et al. (2011) and Volodichev et al. (2012) shows that zircons from
local metasomatic veinlets have the highest concentrations of components among zircons sampled
from the material without any discernible traces of metasomatism. Away from metasomatic veinlets
in the scattered metasomatic zones inside the dyke, the concentration of REE decreases, Ce and Eu
anomalies are flattened, and the zircon acquires features of a typical metasomatic zircon. The diagram
of Hoskin (2005) shows that the zircon composition data point plot between the magmatic and
hydrothermal fields (Supplementary Fig. S2).
The 2.4 Ga zircon has a very high Ti content, which widely varies from 18 to 287 ppm and
suggests a high temperature within the range of 750-1160 ºC (Fig. 18). This is not consistent with the
results of mineralogical geothermometry. However, the titanium content in the zircon has no
significant correlation with the contents of other structural and nonstructural elements, and the zircons
are extremely rich in mineral inclusions, including those of biotite and rutile. We suggest that the high
titanium content may be explained by micron- or submicron-sized domains of Ti-rich phases included
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in zircon and responsible for high-Ti contents at the analytical spots. This possibility is discussed by
Hoffman et al. (2009). To proof this hypothesis, we prepared an element distribution map using time-
of-flight mass spectrometer for one zircon grain 1111-08-9, which was previously dated by SHRIMP
and studied in two spots by ion microprobe. Micron-scale Ti hot-spot were found in analyzed points
with high Ti content (9.1 – SHRIMP and SIMS, 9.2 – SIMS, Table 3, Supplementary Fig. 3). Thus,
temperature estimates exceeding 800 ºC for such zircons seem to be overestimates and are probably
explained by the presence of Ti-bearing phases at the analytical spots.
We believe that the 2.4 Ga zircon was formed during the percolation of aggressive a hot
metasomatic agent through the eclogitized metagabbronorite. This metasomatic agent enriched the
rock in Si, Al, Na, P and Cl, as well as HFSE and REE. The veinlets and their zircon abound in
chlorine-apatite (Supplementary Tables S4c, Fig. 15b). This metasomatic agent was probably
represented water-poor felsic REE, Y, Th, U - rich chlorine-bearing brine-melt that percolated as
weak flux through the metagabbronorite dyke and modified the bulk composition of the dyke in
narrow bands.
2.4 Ga zircons from the metasomatic veinlets underwent alteration and hydration at ca 1.9 Ga.
The REE pattern of the 1.9 Ga domains is typical of hydrothermal zircon that crystallized in presence
of aqueous fluid. The zircon did not change its morphology in the hydrothermal process, but the
composition of the mineral was modified: it was saturated in Ca, Na, Fe, Sr and other nonstructural
elements, and Li, Al, Y, Th (with growth of thorite nanoinclusions). The Ti content in the
recrystallization zone varies from 15 to 4251 ppm. We think that the anomaly high dispersion is
related to nonequilibrium zircon recrystallization in presence of aqueous fluid (see Geisler et al.,
2007) and cannot be interpreted as an indicator of new zircon crystallization temperature.
5.2. P-T-t-evolution of Gridino association
The metamorphic evolution of the Gridino rocks involves at least seven Archaean and
Palaeoproterozoic events (Dokukina et al., 2012): (1) metamorphism of crustal rocks at moderate
pressures and temperatures, which preceded or accompanied the emplacement of the mafic
magmas; (2) metamorphism related to the intrusion of mafic dykes (syn- and post-intrusion
stages) with a possible increase in the crustal temperature that was followed by isobaric cooling
and solidification of mafic magma; (3) a burial stage at increasing pressure and temperature; (4)
peak stage of HP/UHP eclogite-facies metamorphism (Volodichev et al., 2004; Dokukina and
Konilov, 2011; Morgunova and Perchuk, 2012; Perchuk and Morgunova, this issue); (5) a
decompression stage with a P-T-t path passing through the HP-granulite and HP-amphibolite
facies fields; (6) overprinted subisobaric heating with a temperature increase to granulite facies
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conditions; (7) final retrogression under amphibolite facies conditions (Volodichev et al., 2004;
Dokukina and Konilov, 2011).
U-Pb isotope studies of mafic and felsic metamorphic rocks of the Gridino area displayed
some age peaks (Fig. 17) related to successive events and different morphological types of
zircons: (1) 3.0-2.9 Ga zircons corresponding to igneous crystallization of the tonalitic protolith
of the Belomorian gneisses (this study; Dokukina et al., 2010); (2) 2.87-2.82 Ga zircons of
magmatic origin, which date the mafic melt intrusion (this study; Dokukina et al., 2009, 2012);
(3) 2.79-2.78 Ga, 2.74 – the ages of partial melting of crustal rocks and anatectic granite
formation; (4) 2.72-2.64 Ga – the age of high-pressure granulite metamorphism during
decompression accompanied by anatectic granite formation (this work; Dokukina et al., 2009;
2010, 2012; Dokukina and Konilov, 2011); (5) 2.4 Ga thermal impact (this work, Dokukina et
al., 2012); (6) 2.0-1.9 Ga amphibolite-facies metamorphism (this work; Slabunov et al., 2006;
2011; Dokukina et al., 2010, 2012; Dokukina and Konilov, 2011). The P-T-t path of the
metamorphic evolution of the Gridino rocks is shown in Figure 19.
5.2.1. Timing of the eclogite-facies event
We have found out that mafic dykes of the Gridino swarm were emplaced within a
certain time interval between 2.87 and 2.82 Ga. Granulite type zircons with an age of 2.71-2.72
Ga pinpoint the time of high-pressure granulite-facies metamorphism. No zircon that crystallized
under eclogite facies conditions significant for our study has been found as of yet. Coronitic and
symplectitic rocks are notoriously difficult to work with because of their disequilibrium
assemblages. Zircon may or may not grow during these transient metamorphic events, therefore
magmatic and retrograde events are used in this work to bracket the age of the eclogite facies
metamorphism. A single reference to the eclogite metamorphism age at 2721 ± 8 Ma has been
presented by Volodichev et al. (2004). This age coincides with our estimation of age of the
granulite facies metamorphic event. Thus the time of eclogite facies metamorphism is within the
interval from 2.82 to 2.72 Ga. The same interval was suggested for the eclogite facies
metamorphism of oceanic rocks in the Salma eclogite association (Mints et al., 2010a, 2010b and
in this issue). We suggest that the Gridino dyke swarm might result from rifting that preceded the
origin of the Archaean Salma oceanic crust, as is the case, for example, with the Lanzo area in
the Italian Alps (Kaczmarek et al., 2008). However, the compositions of eclogitized mafic dykes
of the Gridino area intersect the composition fields of eclogitized gabbroids of the Salma area
(Mints et al., 2010c, this issue). The ages of the Salma oceanic protolith (2.9-2.82 Ga) and
Gridino dykes (2.87-2.82) are identical within error of the determination. It was hypothesized
that the Gridino eclogitized dykes could be derivatives of slow-spreading mid-oceanic ridge that
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plunged under the Kola active margin (see details in Mints et al., 2010b, 2010c and in this issue).
In this instance, the eclogite facies metamorphism should not have been separated too much in
time from the emplacement of the Gridino dykes and origin of the Salma oceanic crust. The
dykes were emplaced at shallow enough depths (not deeper 15-17 km), as was established by
Dokukina and Konilov (2011) and Egorova and Stepanova (2011) (Fig. 19), and the dilatation in
the plunging mid-oceanic ridge was still in progress for magmatic activity. Before the ridge
reached the high-pressure conditions, its magmatic activity terminated. The youngest age of
~2.82 Ga of the mafic dykes can thus be interpreted as marking the initial subsidence of the
Gridino continental crust or, in other words, the beginning of subduction of the rocks of the Kola
continental margins. The peak high-pressure metamorphic conditions (or ultrahigh-pressure
conditions, as was mentioned above) and eclogite facies metamorphism took place within the
time span of 2.82 to 2.72 Ga. The age of 2.72 Ga is unambiguously interpreted for the Gridino
area as the timing of the decompression event under high-pressure granulite facies condition
(Dokukina and Konilov, 2011, Dokukina et al., 2012). For this age range, we have two values
describing the derivation of anatectic granitoids at 2.8-2.78 and ca 2.74 Ga. These ages can
correspond to the time of dehydration melting during the subsidence of the rock in the
subduction zone or the time of decompression melting during ascent to an upper crustal level.
5.2.2. ~2.72-2.71 Ga granulite facies metamorphism
A granulite facies event is reliably identified in felsic and mafic rocks of the Gridino area
(this work; Dokukina and Konilov, 2011; Konilov and Dokukina, 2011) and is characterized by
granulitic zircon grown in equilibrium with partial melt. The high-pressure decompression
conditions of the origin of the phengite-bearing leucosome (Vargas Cape) were previously dated
by U-Pb SHRIMP II and ID-TIMS methods (Dokukina et al., 2009, 2012). This leucosome cuts
and penetrates a post-eclogite Fe-metagabbro dyke of granulite grade (Cape Vargas locality) and
contains several igneous and metamorphic zircon populations corresponding to various
conditions of host rock recrystallization. The ages of all zircon grains are within a narrow range
(2.71-2.72 Ga, the most concordant age is 2713±6 Ma – SHRIMP II, 2634±5 Ma – ID-TIMS),
and we assumed these dates as the upper time limit for the emplacement of the mafic dyke and
eclogite facies metamorphism in the Gridino area (Dokukina et al., 2009, Dokukina et al., 2012).
Two zircon grains from the high-pressure granulite zone within eclogitized metagabbro dyke
gave an age of 2715±19 Ma (Dokukina et al., 2009). These values are indistinguishable, within
error, from the zircon age of 2721 ± 8 Ma of the eclogite lens in Stolbikha Island (Volodichev et
al., 2004; Slabunov et al., 2006).
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Thus the age of 2.72-2.71 Ga appears to be statistically confirmed. The morphology and
geochemical features of ~2.7 Ga zircon suggest that this age corresponds to the metamorphic
growth of the zircon during post-eclogite decompressional granulite facies metamorphism, and
this sets the upper age limit for eclogite facies metamorphism in the Gridino area.
5.2.3. ~2.4 Ga thermal impact
We found that zircon with an age of ~2.4 Ga from veinlets is of metasomatic origin and
might be related to the hot water-poor felsic chlorine-bearing brine-melt that percolated through
the eclogitized metagabbronorite dyke (see above) at increasing temperature. Further studies are
required to elucidate the nature of the metasomatic agents. Obvious petrological indicators of
increasing temperature are the prograde growth of orthopyroxene around low-temperature
amphiboles in an amphibolite band in the olivine gabbronorite dyke (Fig. 5c). The possibility of
such re-heating in polymetamorphic high-grade complexes was discussed in the literature
(Fonarev et al., 2003; Goncalves et al., 2004; Mints and Konilov, 1998; Perchuk, 2005). The
retrograde PT path (Fig. 19) has a zigzag pulsating trend with a staggered declination in the
direction of higher temperatures ~2.4 Ga ago.
This event of plume nature widely operated throughout the whole Belomorian orogen. A
superplume was formed at 2.5-2.4 Ga the in the continental mantle. The superplume
predetermined the origin of a large igneous province (LIP) in the Kola-Karelia region, the partial
melting of the crustal rocks, the derivation of anatectic granite and granulite facies
metamorphism (Mints et al., 2010c and references therein, also in this issue).
5.2.4. Palaeoproterozoic zircon overgrowth with an age of 2.0-1.8 Ga
Detailed mapping in the Gridino area has shown that the Palaeoproterozoic U-Pb ages of
2.0-1.9 Ga are related to later zones of deformation and amphibolization, in which no high-
pressure mineral assemblages are generally preserved. Low-uranium colorless rims with ages of
~1.9-1.8 Ga around Archaean igneous and metamorphic zircons correspond to amphibolite-grade
overprinting during the Svecofennian orogeny, when imbricate thrusting brought deep-seated
rocks of the Belomorian belt to shallower levels (Bibikova et al., 2004; Slabunov et al., 2006;
Mints et al., 2010c). The zircon rim’s ages correspond to the 40Ar/39Ar age of amphibole from
the amphibolitized Gridino dykes, indicative of amphibole cooling to about 550°C and closing of
the 40Ar/39Ar isotopic system 2.0–1.9 Ga ago (Dokukina et al., 2010, 2012). Thus the
metamorphic transformation of the Gridino rock ensemble under eclogite and granulite facies
conditions occurred prior to 2.0–1.9 Ga.
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6. Conclusion
1) Our data allow us to regard the Meso-Neoarchaean Belomorian eclogite province as a
paleosubduction zone with the modern style of tectonics. HP/UHP processes in the Gridino area
developed in the continental crust, in contrast to the oceanic protolith of eclogites in the Salma
area.
2) The duration of the evolution of the Gridino rocks was more than one billion years
(3.0-1.7 Ga) and involved the origin of the continental crust, its subduction related to
delamination of the active margin crust and several superposed plume-related events.
3) Petrological and geochronological data permit us to determine the age of the major
magmatic and metamorphic events in the Gridino area: (1) continental crust forming – 3.0-2.87
Ga; (2) the emplacement of the Gridino mafic dyke swarm – 2.87-2.82 Ga; (3) HP/UHP
metamorphism within the time span of 2.82-2.72 Ga; (4) “retrograde” post-eclogite granulite
facies metamorphism accompanied by partial melting of the crust at 2.72-2.64 Ga; (5) the age of
~2.4 Ga corresponds to the overprinted thermal impact related to Palaeoproterozoic superplume;
(6) the age of 2.0-1.9 Ga corresponds to the amphibolite-grade overprinting during the
Svecofennian orogeny.
4) We consider the time span between the magmatic and retrograde events to be brackets
for the age of eclogite facies metamorphism. The time of eclogite-facies metamorphism of the
Gridino mafic dykes is between 2.82 and 2.72 Ga, most probably in the span of 2.79-2.73 Ga.
5) The retrograde PT path has a zigzag pulsating trend with a staggered declination in the
direction of higher temperatures at ~2.4 Ga.
Acknowledgements
We are grateful to Dr O.I. Volodichev (Institute of Geology, Karelian Branch, Russian
Academy of Sciences, Petrzavodsk), who showed us to the enderbite vein in the northeastern
vicinity of the Gridino village. We are also grateful to M. Scambelluri and anonymous reviewers
for help in revising and improving the English text of the manuscript. This work was supported
by grants of the Russian Foundation for Basic Research, projects 09-05-00926, 09-05-01006, 11-
05-000492, 12-05-00856 and Program 6 of the Earth Sciences Department of the Russian
Academy of Sciences.
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Figure captions
Figure 1. Simplified geological map of the study area around the village of Gridino
(modified after Slabunov et al., 2007) showing dike swarm and studied sampling sites (red box).
1 – Cover; 2–5 – metagabbroid eclogitized dykes: 2 – late metagabbronorite; 3 – late
metagabbro; 4 – early quartz-bearing and olivine metagabbronorite; 5 – early metagabbro; 6 –
early ferriferous metagabbro; 7 – granite veins; 8 – migmatized areas; 9 – mélange with a high
(up to 20–35% by volume) and low (less 20% by volume) percentage of mafic rock fragments
(eclogite and amphibolite); 10 – fractures; 11 – foliation.
Inset: Schematic geological map of the northwestern Fennoscandian Shield showing
eclogite locations (modified from Mints et al., 2010b, 2010c). Localities of subduction-related
eclogites: 1 - Uzkaya Salma, 2 - Shirokaya Salma, 3 - Chalma (Kuru-Vaara quarry), 4 -
Stolbikha Island. Localities of collision-related eclogites: 5 – Gridino dyke swarm, 6 – Krasnaya
bay (Kozlovskii and Aranovich, 2010).
Figure 2. (a) Geological map of Gridin Cape with sampling sites. (b) Field photo of
intrusive interaction between the eclogitized dykes of later ferriferous metagabbro (thin) and
quartz-bearing metagabbronorite (Vorotnaya Luda Island). (c) Granite leucosome penetrating a
Fe-metagabbro dyke. Red arrow shows a Fe-metagabbro apophysis cutting a host rocks foliation
(d) Field photo of interaction between a granite leucosome vein (sample d44-1), late Fe-
metagabbro dyke (sample d44-4) and olivine metagabbronorite.
Figure 3. PPL photomicrographs of thin sections of Fe-metagabbro dyke, Gridin Cape.
(a) Garnet-clinopyroxene-plagioclase assemblage in a Fe-metagabbro dyke, sample d44-4 (b) A
contact between the late Fe-metagabbro and granite leucosome (sample d44-1).
Figure 4. (a) Schematic geological map sketch of the northeast of the Gridino village
with sampling sites (modified after Volodichev et al., 2012). (b) Exposure of the root of an
enderbite vein from which geochronological sample 1111-06 was collected. (c) Exposure of the
continuation of the enderbite vein from which geochronological sample 1111-09 was collected.
(d) Exposure of metasomatic veinlets penetrating olivine metagabbronorite dyke from which
geochronological sample 1111-08 was collected. Mineral abbreviations here and below are
according to Whitney and Evans (2010).
Figure 5. Olivine metagabbronorite 1111, northeast of the Gridino village. PPL
photomicrograph (a) of omphacite corona around igneous orthopyroxene with systematically
increasing jadeite content toward the corona boundary with Omp + Pl + Crn/Spl symplectite.
Secondary corona of garnet with inclusions of corundum and omphacite cuts the omphacite
corona. The black line A-A’ marks the compositional profile via pyroxenes shown in the inset.
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(b) XPL photomicrograph of the equilibrium assemblage of garnet-omphacite-orthopyroxene-
rutile. (c) BSE image of amphibolite bands in orthopyroxene eclogite with the prograde growth
of secondary orthopyroxene with clinopyroxene, plagioclase, and quartz along amphibole grain
rims. The relationships between the amphibole and omphacite are uncertain, i.e. omphacite could
form after the low-temperature amphibole, or amphibole could be formed after omphacite. In the
former situation, the amphibole is relics of pre-eclogite facies metamorphism. In the latter
instance, the growth of new prograde orthopyroxene around amphibole provides evidence of a
temperature increase after eclogite metamorphism and after the ascent of the eclogitized rocks to
the level of amphibolite facies environment. (d) Enlarged fragment showing relations between
garnet and low-temperature amphibole marked by red box in (c). Numerals refer to analytical
spots presented in Supplementary Table S4d.
Figure 6. Enderbite vein, sample 1111-06, northeast of the Gridino village. (a) BSE
image of the boundary between amphibolite (metagabbronorite) and the enderbite vein. (b) BSE
image of garnet from enderbite. (c) Composition profile along line AA’ in (b). (d) Compositional
diagram showing the chemistry of the garnet from the enderbite vein and the amphibolite of the
boundary.
Figure 7. Metasomatic veinlets, samples 1111-02, 1111-07, 1111-08, northeast of the
Gridino village. (a) BSE image of a boundary between orthopyroxene eclogite and symplectite
eclogite in superimposed veinlets, sample 1111-07. Numerals refer to the mole percentage of
jadeite in omphacite. (b) BSE image of the symplectite eclogite in superimposed veinlets, sample
1111-07. Note orthopyroxene in the structure of the Cpx-Pl symplectite and the enlarged
fragment in the insert, in which orthopyroxene is in direct contacts with potassium feldspar. This
testifies that the symplectite developed at a granulite-facies temperature. (c) PPL
photomicrograph of Qtz-Bt-Pl streaks crosscutting the Opx eclogite in a zone of deformation and
symplectite retrogression. (d) BSE image of zircon grains in Opx-Cpx-Pl symplectite in
metasomatic veinlets, sample 1111-08. Note the high abundance of zircon.
Figure 8. Chondrite-normalized (Sun and McDonough, 1989) REE patterns (left-hand
column) and primitive mantle-normalized (Hofmann, 1988) trace element spider diagrams (right-
hand column) for the rocks. (a-d) See text for explanations.
Figure 9. (a-g) Cathodoluminescence images of the dated zircon from metagabbro,
sample d44-4 (a-c) and granite leucosome, sample d44-1, (d-g), Gridin Cape locality. See text
for explanations. White circles indicate the spots of SHRIMP U–Pb analyses and SIMS trace
element analyses. (h) U-Pb zircon concordia diagrams for the metagabbro, sample d44-4 dated
by SHRIMP II. 1 – igneous cores and rims, 2 - low-thorium rims. (i) Histogram (with default bin
widths) of U-Pb ages of zircon in sample d44-1, dated by LA-ICPMS.
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Figure 10. Chondrite-normalized REE patterns for different zircons from studied
samples. Concentrations were measured with SIMS overlapping the SHRIMP spots. (a-c)
Metagabbro, sample d44-4, see text. (d-f) Granite leucosome, sample d44-1, see text for
explanation.
Figure 11. U-Pb age vs 176Hf/177Hf plot showing a relatively narrow horizontal band of
points that suggests the reset of the U-Pb system, where zircons with younger TDM C—crustal
model age.
Figure 12. Cathodoluminescence images of dated zircon from an enderbite vein,
northeast of the Gridino village. (a-b) Sample 1111-06. (c-e) Sample 1111-09. See text for
explanations. White circles indicate the spots of SHRIMP U–Pb analyses and SIMS trace
element analyses. (f-g) U-Pb zircon concordia diagrams for the enderbite vein: (f) Sample 1111-
06. (g) Sample 1111-09. 1 – oldest grains with oscillatory zoning, 2 – LREE-rich core, 3 - black
structureless rims, 4 - granulite zircon, 5 - pale rim.
Figure 13. REE chondrite normalized patterns for various zircons and garnet in the
studied samples. The concentrations were measured by SIMS overlapping the SHRIMP spots.
(a-b) Sample 1111-06: (a) zircon, (b) garnet, see text for explanations. (c) Sample 1111-09,
zircon, see text for explanation.
Figure 14. BSE image of polycrystalline “nanoenderbite” inclusion with phengite and
omphacite in 2.72-Ga zircon from the enderbite vein, sample 1111-09.
Figure 15. Zircon from metasomatic veinlets, sample 1111-08, northeast of the Gridino
village (modified after Dokukina et al., 2012). (a) SE image of zircon grains. (b) BSE image of
dated zircons. Circles indicate spots of SHRIMP II and SIMS analyses; in spot number 9.2 only
a SIMS analysis was made. (c) AlIV+Na vs Fe/(Fe+Mg) composition diagram for pyroxenes from
olivine metagabbronorite, metasomatic veinlets and inclusions in zircon.
Figure 16. (a-b) Chondrite-normalized REE patterns of zircons from the metasomatic
veinlets, sample 1111-06. Concentrations were measured by SIMS overlapping the SHRIMP
spots. (a) Zircon dated at ca 2.4. Ga, using REE data of Slabunov et al. (2011) for comparison.
(b) Hydrated domains in zircon with an age of ca 1.9. Ga. (c) U-Pb concordia diagram for zircon
from the metasomatic veinlets, sample 1111-08 (after Dokukina et al., 2012).
Figure 17. Integrated histograms for all of geochronological samples. The histograms
constructed for concordant and near concordant points with discordance of 0-2% (Tables 3, 4, 6).
Figure 18. Age (Ma) vs temperature (ºC) diagram for all of the studied zircon grains with
using Ti-in-Zrn geothermometer (Ferry and Watson, 2007). Samples d17-1 and d17-7 from
(Dokukina et al., 2012); samples V-16-65 and V-16-66 from (Slabunov et al., 2011). Open
symbols are the data clearly affected of Ti contamination by ion beam from inclusions.
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Figure 19. PT paths calculated for Gridino rocks using data of Dokukina and Konilov
(2011) and Konilov and Dokukina (2011). 1 - PT path of metamorphic evolution; 2 -
metasomatic veinlets, sample 1111-08, northeast of the Gridino village ; 3 - enderbite vein,
samples 1111-06, 1111-09, northeast of the Gridino village; 4 - amphibolite layers with prograde
orthopyroxene growth in olivine gabbronorite, northeast of the Gridino village ; 5 - granite
leucosome d17, Vargas Cape (from Dokukina and Konilov, 2011); 6 - granite leucosome d44-1,
Gridin Cape; 7 - ferriferous metagabbro; 8 - prograde trend in quartz-bearing metagabbronorite;
9 - quartz-bearing metagabbronorite; 10 - olivine metagabbronorite; 11 - eclogite pods and post-
eclogite retrograde PT path (from Volodichev et al., 2004); 12 - PT evolution of metagabbro
metamorphism calculated using the TWQ software (from Morgunova and Perchuk, 2012); 13 -
PT conditions of gabbronorite dykes (from Egorova and Stepanova, 2011); 14 - field of
amphibole dehydration melting.
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Table captions
Table 1. Whole rock compositions (wt.%) of the studied samples.
Table 2. Trace element compositions (ppm) of the studied samples.
Table 3. SIMS trace element composition of zircon from the studied samples (complete
dataset).
Table 4. SHRIMP II U-Th-Pb isotope data for zircon from samples d44-4 and d44-1,
Cape Gridin locality.
Table 5. LAM-ICPMS U-Th-Pb isotope data for zircon from sample d44-1, Cape Gridin
locality.
Table 6. Lu-Hf isotope data for zircon from sample d44-1, Cape Gridin locality.
Table 7. SHRIMP II U-Th-Pb isotope data for zircon from samples 1111-06, 1111-09
and 1111-08, northeast of the Gridino village.
Table 8. Representative compositions of mineral inclusions in zircons.
Table 9. Representative EPMA analyses of zircon grain 1111-08-9 from the
geochronological sample 1111-08.
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Table 1. Whole rock composition (wt.%) of the studied rocks. Locality Cape Gridin Northeastern outskirt of Gridino Villag e
Rock type Gneiss Olivine
gabbronoriteFerriferous metagabbro
Granite leucosome
Olivine gabbronorite
Enderbite vein Metasomatic
veinlets Sample d44e d44c/4 d44b/1 d44/4 d44-1 1111-05 1111-06 1111-09 1111-07se
SiO2 67.90 66.40 49.10 49.50 66.31 48.50 64.20 60.64 51.40
TiO2 0.48 0.21 0.68 0.98 0.43 0.57 0.42 0.72 0.50 Al 2O3 16.28 12.30 9.56 12.60 13.79 8.92 16.10 15.21 12.90 FeO 3.54 4.00 11.43 12.15 5.76 11.34 4.54 7.51 12.06 MnO 0.06 0.09 0.18 0.18 0.12 0.17 0.08 0.12 0.18 MgO 1.35 4.65 16.90 7.27 3.88 17.80 3.54 4.04 8.19 CaO 4.02 6.35 8.06 11.20 2.59 9.04 5.01 5.24 7.91 Na2O 4.77 3.80 1.56 2.73 2.83 1.34 4.23 3.82 2.61 K2O 0.75 0.83 0.61 0.02 2.69 0.36 0.92 1.52 0.16 P2O5 0.19 0.05 0.09 0.08 0.06 0.07 0.06 0.07 0.17
C n.d. n.d. n.d. 0.13 n.d. n.d. n.d. n.d. n.d. F <0.01 n.d. n.d. n.d. 0.10 n.d. n.d. n.d. n.d. Cl n.d. 0.022 <0.005 0.012 0.100 0.005 n.d. 0.045 0.029 S <0.01 0.006 0.023 0.025 0.060 0.022 n.d. 0.018 0.035
LOI 0.16 n.d. n.d. n.d. 0.28 n.d. 0.33 0.13 n.d.
Total 99.50 98.70 98.19 96.88 99.00 98.13 99.42 99.07 96.14
#Mg 0.40 0.67 0.72 0.52 0.55 0.74 0.58 0.49 0.55
Note: n.d. – not detected. #Mg = Mg/(Fe+Mg).
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Table 2. Trace element compositions (ppm) of the studied rock types. Locality Cape Gridin Northeastern outskirt of Gridino Village
Rock Olivine
gabbronorite Ferriferous metagabbro
Granite leucosome
Gneiss Olivine
gabbronorite Enderbite vein
Metasomatic veinlets
Sample d44b/1 d44/4 d44-1 d44e d44e/4 1111-05 1111-06 1111-09 1111-07se Li 4.2 15.1 11.4 13.6 9.7 7.0 na 8.7 5.3 Be 0.33 0.40 0.11 0.73 0.93 0.26 na 0.68 0.21 Sc 29.1 46.6 6.3 5.4 16.7 29.0 na 24.3 28.9 V 182 278 118.4 56.4 59.6 183 85 126 140 Cr 2051 172 7.7 10.7 140 1599 74 41.5 205 Co 77.8 47.5 20.0 8.2 17.0 77.5 18.9 26.8 54.8 Ni 545 69.0 49.2 15.9 74.6 528 30.1 116 127 Cu 70.2 72.4 63.5 21.7 19.6 85.4 na 73.6 254 Zn 84.3 86.9 31.2 55.6 80.1 731.0 na 91.4 39.0 Ga 11.8 16.0 18.5 18.1 15.9 11.2 na 19.3 9.9 As - 0.16 0.50 - - 0.87 na 0.36 1.5 Se < 0.8 < 0.9 < 0.8 - - - na - - Rb 11.4 0.2 54.3 13.3 10.1 7.4 19.4 30.7 1.2 Sr 93 49 254 336 387 106 429 208 81.3 Y 11.5 17.1 7.0 3.1 10.7 10.1 8.13 14.4 12.1 Zr 35.6 9.6 87.3 127 49.8 39.7 102 68.7 49.5 Nb 2.3 1.6 1.9 3.2 5.3 2.2 4.51 7.6 2.7 Mo 0.33 - 0.79 0.20 0.21 1.2 na 0.50 1.1 Ag 0.045 - - - 0.04 0.11 na 0.043 0.16 Cd 0.074 0.071 - - 0.08 - na 0.10 0.060 Sn 0.86 0.56 3.6 0.48 - 0.39 na 1.3 0.63 Sb - - - 0.06 - 0.90 na 0.091 0.57 Cs 0.21 0.007 0.45 0.079 0.37 0.15 na 0.18 0.011 Ba 152 3.7 899 374 267 96.2 310 484 45.3 La 8.6 1.3 27.1 17.0 10.5 4.0 14 15.9 12.2 Ce 18.7 3.0 66.3 31.8 25.8 10.1 28.2 41.0 25.3 Pr 2.27 0.51 9.46 3.4 3.0 1.3 3.27 5.4 2.5 Nd 9.7 3.1 40.6 12.9 12.9 6.03 11.8 22.8 8.71 Sm 2.2 1.4 7.7 1.9 2.7 1.5 2.15 4.8 1.6 Eu 0.65 0.57 2.2 0.71 0.83 0.47 0.86 0.80 0.51 Gd 2.4 2.6 3.3 1.4 2.3 1.8 2.21 4.2 2.2 Tb 0.39 0.47 0.30 0.17 0.33 0.28 0.27 0.59 0.36 Dy 2.4 3.3 1.4 0.77 1.8 1.7 1.62 3.1 2.1 Ho 0.49 0.71 0.23 0.13 0.33 0.38 0.28 0.62 0.45 Er 1.5 2.2 0.62 0.38 0.93 1.1 0.67 1.6 1.3 Tm 0.20 0.32 0.082 0.047 0.13 0.15 0.091 0.22 0.18 Yb 1.3 2.1 0.52 0.34 0.86 1.1 0.66 1.3 1.2 Lu 0.19 0.31 0.074 0.052 0.12 0.15 0.12 0.19 0.16 Hf 1.0 0.44 2.2 3.5 1.4 1.0 na 1.9 1.2 Ta 0.16 0.11 0.11 0.12 0.17 0.23 0.12 0.26 0.22 W 0.10 - 0.10 0.084 0.14 0.066 na 0.11 0.50 Tl 0.079 - 0.14 0.046 0.047 0.039 na 0.14 0.020 Pb 3.6 0.47 1.9 2.2 6.2 1.7 na 3.1 2.4 Bi - - - - 0.038 - 2.33 - 0.065 Th 1.9 0.12 4.5 0.29 1.4 0.89 0.25 0.90 3.4 U 0.37 0.012 0.73 0.08 0.37 0.19 na 0.26 0.41
Note: na – not analysed.
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Table 3. SHRIMP II U-Th-Pb isotope data for zircon from samples d44-4 and d44-1, Cape Gridin locality.
Spot % 206Pbc U Th
232Th/238U206Pb* (1)
206Pb/238UAge
(1) 207Pb /206Pb Age
% Dis- cor- dant
(1) 207Pb*/206Pb* ±% (1)207Pb*/235U±%
(1) 206Pb* /238U
±%err corr
ppm
Late ferriferous metagabbro, sample d44-4 d44-4.1.1 0.04 723 54 0.08 309 2603 ±30 2722 ±13 5 0.1877 0.81 12.88 1.6 0.49751.4 0.87 d44-4.1.2 0.25 74 6 0.09 34.4 2784 ±35 2775 ±32 0 0.1939 2.00 14.44 2.5 0.54011.6 0.62
d44-4.2.1 0.66 37 50 1.39 17.8 2832 ±41 2840 ±29 0 0.2017 1.80 15.34 2.5 0.55171.8 0.70 d44-4.3.1 0.01 1321 67 0.05 603 2747 ±30 2790±6.5 2 0.19559 0.40 14.33 1.4 0.53131.3 0.96
d44-4.3.2 0.35 37 46 1.27 18.1 2878 ±47 2853 ±25 -1 0.2033 1.60 15.77 2.5 0.5630 2.0 0.79 d44-4.4.1 0.05 526 91 0.18 239 2730 ±31 2809±8.8 3 0.1979 0.54 14.39 1.5 0.52731.4 0.93
d44-4.5.1 0.06 393 274 0.72 184 2805 ±33 2850±9.6 2 0.2029 0.59 15.25 1.5 0.54511.4 0.92 d44-4.6.1 0.43 79 17 0.22 37.7 2830 ±43 2863 ±24 1 0.2046 1.50 15.55 2.4 0.5510 1.9 0.78
d44-4.7.1 0.11 112 1 0.01 49.9 2687 ±36 2753 ±16 2 0.1912 0.98 13.64 1.9 0.51711.7 0.86 d44-4.8.1 0.68 18 6 0.38 8.32 2793 ±50 2771 ±57 -1 0.1934 3.40 14.46 4.1 0.5420 2.2 0.54
d44-4.9.1 0.14 82 31 0.38 38.9 2831 ±34 2845 ±15 0 0.2024 0.93 15.39 1.7 0.55151.5 0.85 d44-4.9.2 0.11 351 46 0.14 144 2512 ±31 2691 ±12 7 0.1842 0.73 12.10 1.6 0.47661.5 0.90
d44-4.9.3 0.21 177 45 0.26 76.8 2624 ±34 2733 ±16 4 0.1890 0.98 13.09 1.9 0.50241.6 0.85 d44-4.9.4 0.20 134 6 0.05 57.6 2607 ±36 2717 ±18 4 0.1871 1.10 12.86 2.0 0.49831.7 0.84 Granite leucosome, sample d44-1 d44-1.1.1 -- 14 4 6.29 0.32 2665 ±81 2764 ±51 +4 0.193 3.1 13.6 4.9 0.512 3.7 0.8 d44-1.2.1 -- 377 83 171 0.23 2732 ±35 2763 ±9 +1 0.192 0.5 14.0 1.6 0.528 1.6 0.9 d44-1.3.1 0.05 181 35 84.7 0.20 2810 ±39 2808 ±11 -0 0.198 0.7 14.9 1.8 0.546 1.7 0.9 d44-1.4.1 0.02 654 25 307 0.04 2813 ±34 2840 ±7 +1 0.202 0.4 15.2 1.6 0.547 1.5 1.0 d44-1.4.2 -- 30 14 13.8 0.48 2751 ±59 2815 ±30 +3 0.199 1.9 14.6 3.2 0.532 2.6 0.8 d44-1.5.1 0.61 15 8 7.15 0.55 2843 ±86 2926 ±39 +3 0.213 2.4 16.3 4.5 0.554 3.7 0.8 d44-1.5.2 0.00 24 23 10.5 1.01 2685 ±70 2906 ±54 +9 0.210 3.4 15.0 4.6 0.517 3.2 0.7 d44-1.6.1 0.11 77 34 36.6 0.46 2851 ±46 2908 ±48 +2 0.210 3.0 16.1 3.6 0.556 2.0 0.6 Errors are 1σ; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in Standard calibration was 0.56% (not included in above errors but required when comparing data from different mounts).
(1) Common Pb corrected using measured 204Pb.
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Table 4. SIMS trace element composition of zircon from investigated samples (complete dataset). Sample P Ca Ti Li Sr Y Nb Ba La Ce Pr Nd Sm Eu Gd Dy Er Yb Lu Hf Th U Late ferriferous metagabbro, sample d44-4, Cape Gridin locality ppm d44-4.1.1 544 4 10.6 88 1.04 989 19 0.95 0.11 5 0.08 0.75 1.62 0.14 11 170 453 64 80 14602 73 1630 d44-4.1.2 351 23 6.1 14 0.74 443 17 1.47 0.27 9 0.35 3.00 3.46 0.55 11 78 159 36 27 10505 26 96 d44-4.3.1 1444124414.01074.86119921 0.58 1.53 10 1.09 7.95 5.37 2.15 14 252 655 79 10918702 88 2106 d44-4.3.2 275 5 6.7 11 0.48 506 12 0.54 0.27 9 0.13 1.11 2.45 0.12 14 84 153 46 27 10049 52 62 d44-4.4.1 372 1 6.4 63 0.64 982 9 0.39 0.11 10 0.05 0.97 2.65 0.13 19 151 247 84 40 11464 116 847 d44-4.5.1 558 6 9.0 3.5 0.881733 16 1.45 0.34 41 0.45 5.02 9.11 2.83 50 366 612 193 1008809 332 633 d44-4.7.1 238 27 7.5 22 0.81 178 23 1.49 2.05 10 1.10 8.26 7.32 1.86 11 31 71 16 14 12444 3.2 239 d44-4.9.1 161 17 5.2 7.3 0.60 543 31 1.33 0.93 14 1.1910.4312.07 3.01 23 89 175 51 31 10478 42 157 d44-4.9.2 286 26 8.9 54 1.06 811 30 1.32 1.34 16 2.1919.2020.27 3.73 31 127 273 61 44 12131 49 457 d44-4.9.3 178 20 10.6 39 0.80 648 31 1.94 0.86 11 1.01 9.34 10.20 2.50 19 110 224 50 36 11592 41 419 d44-4.9.4 17 30 14.8 13 1.05 257 22 29.430.82 7 0.66 4.51 4.72 1.10 9 45 103 23 19 10506 14 172
Granite leucosome, sample d44-1, Cape Gridin locality d44-1.1.1 4 87.40.3 2.23127017 18.780.15 5 0.11 1.89 3.74 0.81 23 220 411 97 72 6291 4.71 20.95d44-1.2.1 11 6.9 45 0.761244 9 1.96 1.55 14 0.92 7.08 6.62 1.26 27 196 369 94 60 9881 88.27 529 d44-1.3.1 9 4.0 17 0.85 469 12 1.93 1.73 11 1.08 7.03 4.17 1.18 11 76 161 33 28 9003 30.08 210 d44-1.4.1 32 12.2 55 1.98 767 17 3.30 5.39 29 5.2149.8138.4912.35 63 117 320 75 56 13230 27.29 844 d44-1.4.2 15 6.7 6.0 0.59 280 5.2 1.91 3.05 16 1.26 7.72 3.71 0.95 10 47 86 24 15 8118 12.55 41.6 d44-1.6.1 63 9.8 2.9 0.94 876 6.1 2.18 2.38 21 1.05 7.96 8.32 2.47 23 143 299 63 52 6850 37.83 120 d44-1.8.1 564 55 20.6 46 2.18263144 2.06 2.28 28 0.39 3.00 6.29 1.89 45 430 794 203 12712945504.72 1742 d44-1.8.2 124 8 6.6 17 0.78 497 20 1.10 0.89 12 0.40 2.74 2.38 0.53 10 83 164 34 28 9009 32.22 98.9 To be continued.
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Table 4 (Continued) Sample P Ca Ti Li Sr Y Nb Ba La Ce Pr Nd Sm Eu Gd Dy Er Yb Lu Hf Th U Enderbite vein, sample 1111-06, Northeastern outskirt of the Gridino Village locality 1111-06.1.1 1.96 6.7 2.4 0.59 729 12 0.79 0.11 10 0.06 0.91 1.87 0.48 10 49 126 269 48 7382 26.21 73.6 1111-06.2.1 7 119.3 4.9 0.82 877 8 19.28 0.36 26 0.31 2.95 4.12 1.52 19 65 141 281 50 6255 65.22 112 1111-06.3.1 349721.9 19 92.68 1026 18 11.1827.99 199 33.1 247.9 177.4 23.2 185 111 143 299 50 11084 76.73 550 1111-06.3.2 10 11.6 45 0.63 231 10 3.53 0.46 9 0.56 4.20 3.53 0.55 7 18 36 76 13 8552 11.08 402 1111-06.5.1 55 0.01 5.2 30 0.44 405 60 1.98 0.38 11 0.34 2.62 2.68 0.43 8 68 167 29 29 8681 23.04 240 1111-06.5.2 303 43 19.5 43 2.80 1082 12 6.04 10.99 85 11.8699.18101.15 15.35 147 174 330 112 55 9803 68.90 529
1111-06.8.1 4 4.6 71 0.57 296 14 2.03 0.58 8 0.60 4.73 4.33 0.66 8 20 48 108 20 8731 16.34 698
1111-06.9.1 53 6 6.2 33 0.52 377 37 2.07 0.30 11 0.31 2.60 3.27 0.77 11 61 132 29 24 9531 60.78 377
Enderbite vein, sample 1111-09, Northeastern outskirt of the Gridino Village locality 1111-09.1.1 82 0.42 5.1 22 0.33 297 30 0.77 0.04 13 0.08 0.95 1.78 0.43 8 47 94 25 18 9421 105.92 239 1111-09.3.1 128 0.89 4.6 26 0.52 353 14 0.32 0.06 14 0.08 1.08 1.79 0.44 9 59 126 30 24 10246120.53303.241111-09.4.1 161 0.27 5.8 14 0.58 360 6 0.68 0.04 21 0.05 0.77 1.52 0.21 9 64 123 30 21 9480 71.13 97.00 1111-09.6.1 131 0.70 5.7 10 0.45 374 12 0.54 0.08 20 0.06 0.71 1.46 0.16 8 65 123 30 22 9802 52.92 73.15 1111-09.6.2 159 0.34 3.1 1.5 0.77 167 6 0.42 0.02 16 0.05 0.45 1.54 0.18 7 58 124 21 16 7934 37.84 42.13 1111-09.7.1 0.68 19 8.1 8.3 0.49 265 19 1.89 1.40 23 1.24 9.79 10.84 1.47 17 43 97 24 17 9156 35.47 73.38 1111-09.7.1D 1.27 21 8.4 10.8 0.71 273 15 1.81 1.76 25 1.48 11.36 10.25 1.66 17 48 105 25 17 9074 44.09 89.18 1111-09.9.1 229 20 6.8 5.7 1.01 473 10 1.21 0.13 14 0.21 3.01 3.70 0.42 15 80 138 42 25 8318 49.61 33.78 1111-09.9.2 64 12 6.0 20.8 0.45 195 9 1.06 0.13 10 0.16 1.35 1.36 0.30 5 33 77 16 14 9324 68.51 169.64
Metasomatic veins, sample 1111-08, Northeastern outskirt of the Gridino Village locality 1111-08.1.1 493 172 213 62 1.5 11812328 12 5.6 1429 10 106 143 29 565 167923251457 298 8603 11940 4262 1111-08.2.1 587 21 48 10.2 874 21 16 20.0 143 21 155 93 27 89 176 439 72.7372.61 6527 188 1165 1111-08.3.1 1193545 273 37 9.2 25156282 18 2.6 1150 11 166 390 16 1572 275533363180 468 7705 2700 3028 1111-08.4.1 2130 26 157 30 2.4 30177548 5 0.5 808 4 62 168 10 1024 361045073690 562 9218 15361 8495 1111-08.5.1 1129 156 288 22 4.6 20338206 16 9.8 702 7 97 230 10 1019 238332172388 427 9740 2635 2778 1111-08.8.1 163 15 54 4.7 4558 76 6 3.5 19 1.9 11 4.5 2.1 24.06 745 1690 246 247 8686 692 2729 1111-08.9.1 503 755 113 51 4.3 5027 75 12 1.0 126 2.1 41 100 18 359 793 1498 647 221 8616 2001 2355 1111-08.9.2 5981 111 15 349 4865 230 310 105 366 24 127 69 17 169 833 1879 399 247 6851 10843 7149 1111-08.10.1RE 56784252 58 25.3 3696 90 107 56 213 13 62 16 6 52.9 652 1556 237 240 6478 3810 3264
Note: blank fields – concentrations below detection limits.
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Table 5. LAM-ICPMS U-Th-Pb isotope data for zircon from sample d44-1, Cape Gridin locality.
Spot
206Pbc, %
Concentration, ppm Corrected ages, Ma
Discor-dance, %
Corrected ratios
U Th 206Pb/238U 207Pb/206Pb
207Pb/206Pb
1σ 207Pb/235U
1σ
206Pb/238U
1σ
d44-1-02 88 116 2772 ±23 2845 ±18 3.2 0.20230 0.00217 14.98575 0.15956 0.53732 0.00550
d44-1-03 395 94 2682 ±25 2704 ±18 1 0.18569 0.00202 13.20727 0.15545 0.51593 0.00584
d44-1-04 175 80 2883 ±26 2864 ±19 -0.8 0.20471 0.00238 15.91320 0.19376 0.56392 0.00630
d44-1-06 1000 453 2734 ±26 2729 ±21 -0.2 0.18843 0.00239 13.72330 0.18235 0.52830 0.00607
d44-1-07 499 158 2776 ±24 2802 ±18 1.1 0.19703 0.00217 14.62286 0.16331 0.53831 0.00565
d44-1-08 1653 318 2786 ±24 2784 ±18 -0.1 0.19489 0.00208 14.52300 0.16058 0.54051 0.00574 d44-1-09 0.04 1578 131 2712 ±23 2798 ±18 3.7 0.19653 0.00211 14.17307 0.15443 0.52308 0.00544 d44-1-10 254 92 2755 ±25 2817 ±19 2.7 0.19887 0.00228 14.62014 0.17401 0.53323 0.00583 d44-1-11 29 9 2778 ±29 2834 ±26 2.4 0.20093 0.00313 14.92336 0.24066 0.53872 0.00704 d44-1-12 27 18 2786 ±24 2800 ±19 0.6 0.19683 0.00227 14.66895 0.16840 0.54055 0.00571 d44-1-13 415 54 2809 ±24 2830 ±18 0.9 0.20046 0.00211 15.09206 0.16590 0.54607 0.00583 d44-1-14 614 158 2764 ±27 2785 ±19 1 0.19506 0.00217 14.39223 0.17913 0.53531 0.00636 d44-1-15C 628 260 2735 ±24 2787 ±17 2.3 0.19528 0.00201 14.22894 0.15445 0.52850 0.00568 d44-1-15R 383 92 2764 ±25 2787 ±18 1 0.19522 0.00214 14.40829 0.16842 0.53534 0.00594 d44-1-19C 180 41 2876 ±27 2851 ±18 -1 0.20310 0.00221 15.73824 0.19102 0.56216 0.00658 d44-1-19R 404 6 2662 ±25 2712 ±19 2.3 0.18651 0.00209 13.14179 0.16255 0.51124 0.00596 d44-1-22 0.15 105 53 2618 ±24 2700 ±18 3,7 0.18519 0.00198 12.78774 0.14737 0.50091 0.00560 d44-1-23 427 135 2844 ±25 2848 ±18 0,2 0.20271 0.00215 15.49927 0.17580 0.55461 0.00611 d44-1-24C 171 74 2770 ±24 2801 ±18 1,4 0.19693 0.00215 14.57091 0.16483 0.53668 0.00576 d44-1-24R 181 45 2683 ±26 2755 ±20 3,3 0.19149 0.00232 13.61806 0.18181 0.51616 0.00623 d44-1-25 0.17 110 84 2709 ±25 2810 ±18 4,4 0.19808 0.00211 14.26454 0.16301 0.52240 0.00579 d44-1-26 176 53 2815 ±24 2838 ±18 1 0.20142 0.00214 15.20578 0.16369 0.54750 0.00568 d44-1-27 0.36 238 68 2711 ±26 2789 ±30 3,4 0.19551 0.00347 14.09118 0.18732 0.52274 0.00614 Note: C – core, R – rim.
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Table 6. Lu-Hf-Yb isotope data for zircon from sample d44-1, Cape Gridin locality.
Spot 176Hf/177Hf 1 σ 176Lu/177Hf 176Yb/177Hf U-Pb Age,
Ma 2 σ Hfi epsilon 1se
T(DM) (Ga)
T(DM) (crustal)
Hf Chur (t)
Hf DM (t)
d44-1-02 0.281024 0.000012 0.000521 0.016755 2845 36 0.280996 1.5 0.4 3.06 3.22 0.280954 0.281159 d44-1-03 0.280998 0.000009 0.000838 0.024718 2704 36 0.280955 -3.3 0.3 3.12 3.41 0.281047 0.281265 d44-1-06 0.281067 0.000013 0.001534 0.039520 2729 42 0.280987 -1.6 0.5 3.08 3.32 0.281031 0.281246 d44-1-10 0.281025 0.000009 0.000500 0.018468 2817 38 0.280998 0.9 0.3 3.06 3.23 0.280973 0.281180 d44-1-11 0.281064 0.000014 0.000634 0.020841 2834 52 0.281030 2.4 0.5 3.02 3.15 0.280961 0.281167 d44-1-14 0.281047 0.000076 0.000327 0.009236 2785 38 0.281030 1.3 2.7 3.02 3.18 0.280994 0.281204 d44-1-15 0.281026 0.000013 0.000730 0.022213 2787 34 0.280987 -0.2 0.5 3.08 3.28 0.280992 0.281202 d44-1-19c 0.280977 0.000010 0.000632 0.020160 2851 36 0.280942 -0.3 0.3 3.13 3.33 0.280950 0.281154 d44-1-19r 0.281057 0.000033 0.000150 0.004234 2712 38 0.281049 0.3 1.2 2.99 3.19 0.281042 0.281259 d44-1-24c 0.280987 0.000008 0.000360 0.010070 2801 36 0.280968 -0.5 0.3 3.10 3.31 0.280983 0.281192 d44-1-24r 0.281048 0.000027 0.000470 0.012094 2755 40 0.281023 0.3 1.0 3.03 3.22 0.281013 0.281226 d44-1-25 0.280998 0.000012 0.000391 0.013006 2810 36 0.280977 0.0 0.4 3.09 3.28 0.280977 0.281185
Note: c – core, r – rim. Scherer et al., 2001 - 176Lu decay constant (1.865x10-11)
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Table 7. SHRIMP II U-Th-Pb isotope data for zircon from samples 1111-06, 1111-09 and 1111-08, Northeastern outskirt of the Gridino
Village.
Spot % 206Pbcppm U ppm Th
ppm 206P* 232Th/238U(1)
206Pb/238UAge
(1) 207Pb/206Pb
Age
% Discor-
dant
(1) 207Pb* /206Pb*
±% (1)
207Pb*/235U±%
(1) 206Pb*/238U
±% Err corr ppm
Enderbite vein, sample 1111-06 1111-06.1.1 -- 68 36 35.3 0.55 3042 ±55 3013 ±19 -1 0.224 1.2 18.7 2.5 0.603 2.3 0.9 1111-06.2.1 0.26 94 77 45.4 0.85 2869 ±44 2952 ±14 +3 0.216 0.9 16.7 2.1 0.561 1.9 0.9 1111-06.3.1 0.04 431 61 196 0.15 2740 ±41 2830 ±8 +4 0.200 0.5 14.6 1.9 0.530 1.8 1.0 1111-06.3.2 0.01 296 12 133 0.04 2715 ±35 2732 ±9 +1 0.189 0.6 13.6 1.7 0.524 1.6 0.9 1111-06.4.1 -- 220 73 99.4 0.34 2722 ±36 2735 ±10 +1 0.189 0.6 13.7 1.7 0.525 1.6 0.9 1111-06.5.1 0.03 163 25 73.1 0.16 2712 ±38 2879 ±11 +7 0.207 0.7 14.9 1.9 0.523 1.7 0.9 1111-06.5.2 -- 293 10 127 0.03 2635 ±35 2680 ±10 +2 0.183 0.6 12.7 1.7 0.505 1.6 0.9 1111-06.6.1 -- 37 69 15.8 1.94 2605 ±54 2785 ±24 +8 0.195 1.4 13.4 2.9 0.498 2.5 0.9 1111-06.7.1 -- 103 27 30.1 0.27 1890 ±31 1984 ±22 +5 0.122 1.2 5.7 2.3 0.341 1.9 0.8 1111-06.8.1 0.02 566 20 255 0.04 2719 ±33 2759 ±7 +2 0.192 0.4 13.9 1.6 0.525 1.5 1.0
Enderbite vein, sample 1111-09 1111-09.1.1 0.33 143 85 0.62 64.2 2695 ±34 2724 ±14 1 0.1879 0.84 13.45 1.8 0.5190 1.6 0.88 1111-09.2.1 0.02 596 64 0.11 254 2597 ±30 2691 ±8.9 4 0.1842 0.54 12.60 1.5 0.4962 1.4 0.93 1111-09.3.1 0.01 192 97 0.52 84.5 2664 ±33 2691 ±14 1 0.1842 0.83 13.00 1.7 0.5118 1.5 0.88 1111-09.4.1 0.27 67 65 1.01 29.1 2632 ±41 2704 ±23 3 0.1857 1.4 12.91 2.3 0.5041 1.9 0.81 1111-09.5.1 0.18 56 32 0.59 24.3 2643 ±40 2737 ±20 4 0.1894 1.2 13.23 2.2 0.5067 1.9 0.84 1111-09.6.1 0.50 21 21 1.07 9.19 2668 ±55 2736 ±35 3 0.1893 2.1 13.38 3.3 0.5130 2.5 0.76 1111-09.6.2 0.51 50 48 0.98 21.4 2586 ±41 2718 ±25 5 0.1873 1.5 12.74 2.5 0.4935 1.9 0.78 1111-09.7.1 0.50 50 34 0.70 22.1 2643 ±43 2685 ±27 2 0.1835 1.6 12.83 2.6 0.5069 2.0 0.77 1111-09.8.1 1.17 60 4 0.07 18.1 1912 ±37 1930 ±70 1 0.1182 3.9 5.63 4.5 0.3453 2.2 0.50 1111-09.9.1 0.58 20 42 2.13 9.78 2859 ±59 2810 ±36 -2 0.1981 2.2 15.24 3.4 0.5580 2.5 0.76 1111-09.9.2 0.04 108 60 0.57 47.8 2668 ±35 2731 ±13 2 0.1887 0.79 13.33 1.8 0.5126 1.6 0.90
Metasomatic veins, sample 1111-08 1111-08.1.1 -- 2455 6757 959.0 2.84 2417 ±23 2392 ±6 -1 0.1541 0.3 9.66 1.2 0.455 1.1 1.0 1111-08.2.1 0.35 739 93 217.0 0.13 1898 ±40 1893 ±15 -0 0.1159 0.8 5.47 2.6 0.342 2.5 0.9 1111-08.3.1 -- 2599 3717 1063.7 1.48 2512 ±22 2396 ±6 -6 0.1545 0.4 10.15 1.1 0.476 1.0 0.9 1111-08.4.1 -- 6881 17715 2902.4 2.66 2575 ±21 2402 ±5 -9 0.1550 0.3 10.50 1.0 0.491 1.0 1.0 1111-08.5.1 0.16 1960 2338 753.3 1.23 2384 ±20 2397 ±6 +1 0.1545 0.4 9.53 1.1 0.448 1.0 0.9 1111-08.6.1 -- 1320 1311 523.7 1.03 2448 ±21 2394 ±7 -3 0.1543 0.4 9.83 1.1 0.462 1.0 0.9 1111-08.8.1 0.08 1759 508 509.1 0.30 1871 ±17 1880 ±8 +1 0.1150 0.5 5.34 1.1 0.337 1.0 0.9 1111-08.9.1 -- 2342 4892 912.8 2.16 2411 ±21 2387 ±5 -1 0.1537 0.3 9.61 1.1 0.454 1.0 1.0 1111-08.10.1 -- 429 18 126.0 0.04 1896 ±8 1848 ±28 -3 0.1130 1.5 5.33 1.6 0.342 0.5 0.3
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1111-08.10.1RE 0.26 2667 2346 894.0 0.91 2124 ±20 2149 ±7 +1 0.1338 0.4 7.20 1.2 0.390 1.1 0.9 Errors are 1 σ; Pbc and Pb* indicate the common and radiogenic portions, respectively. Error in Standard calibration was 0.56% (not included in above errors but required when comparing data from different mounts).
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Table 8. Representative compositions of mineral inclusions in zircons.
Locality Gridin Cape Northeastern outskirt of the Gridino Village Rock Granite leucosome Enderbite vein Metasomatic veinlets Sample d44-1 1111-06 1111-09 1111-08 Mineral Bt 4 Ms 2 Bt 3 Hbl 1 Kfs 2 Cpx11 Bt 10 Ph 9 Cpx22 Cpx23 Cpx29 Opx31 Bt 21 Bt 26 Hbl24 Location Incl Incl Incl Incl Incl Incl Incl Incl Incl Incl (?) Incl (?) Incl Incl Incl Contacts Zrn Zrn Zrn Zrn Zrn, Bt Zrn, Ph Zrn, Qtz Zrn Zrn, Hbl Zrn, Opx Zrn, Cpx Zrn Zrn Zrn, Cpx Oxides, wt.% SiO2 34.90 45.32 36.95 43.83 66.38 50.62 37.47 49.73 53.31 53.85 53.73 53.64 37.09 37.38 44.20 TiO2 3.54 0.18 2.56 1.14 0.05 0.48 3.56 0.87 0.17 0.33 0.21 0.13 5.97 5.70 1.97 Al2O3 15.09 31.93 18.32 16.7 19.42 9.38 19.48 32.54 4.30 4.30 1.88 1.45 15.10 15.28 12.12 Cr2O3 0.13 0 0.01 0.15 0.10 0.21 0.06 0.04 0.32 0.10 0 0.04 0.25 0.87 0.45 FeO* 18.76 2.15 10.02 10.84 0.43 6.03 12.15 1.87 4.04 4.27 5.70 18.22 9.64 8.76 7.47 MnO 0.02 0 0 0 0.13 0.22 0.01 0 0.05 0.04 0.26 0.07 0 0 0.09 MgO 12.18 1.08 15.32 12.88 0 11.17 13.41 2.10 14.39 14.99 15.37 26.19 17.40 17.27 15.27 CaO 0.14 0.14 0.04 11.51 0.18 19.02 0.17 0.31 21.35 20.14 22.19 0.26 0 0.19 13.09 Na2O 0.02 0.47 0.08 2.17 0.06 2.83 0.16 0.63 2.07 1.97 0.64 0 0.57 0.53 2.43 K2O 8.96 10.76 9.83 1.43 15.87 0.04 9.54 9.96 0 0 0.03 0 8.98 9.02 0.91 BaO - - 0 0 0.99 0 0 0 0 0 0 0 0 0 0 ZrO2 1.21 1.07 0 0 0 0 0 0 0 0 0 0 0 0 0 Total 94.95 93.10 93.13 100.65 103.61 100 96.01 98.05 100 100 100 100 95 95 98 Cations Si 2.686 3.135 2.751 6.196 2.975 1.854 2.723 3.212 1.943 1.954 1.973 1.954 2.712 2.725 6.378 Al IV 1.314 0.865 1.249 1.804 1.026 0.146 1.277 0.788 0.057 0.046 0.027 0.046 1.288 1.275 1.622 Al VI 0.078 1.768 0.358 0.979 0 0.259 0.391 1.688 0.128 0.138 0.054 0.016 0.014 0.037 0.439 Ti 0.208 0.009 0.143 0.121 0.002 0.013 0.195 0.042 0.005 0.009 0.006 0.004 0.328 0.312 0.214 Cr 0.008 0 0.001 0.017 0.004 0.006 0.003 0.002 0.009 0.003 0 0.001 0.015 0.050 0.051 Fe'' 1.228 0.126 0.624 1.282 0.016 0.185 0.738 0.101 0.123 0.130 0.175 0.555 0.589 0.534 0.901 Mn 0.001 0 0 0 0.005 0.007 0.001 0 0.001 0.001 0.008 0.002 0 0 0.011 Mg 1.421 0.113 1.700 2.715 0 0.610 1.453 0.202 0.782 0.811 0.841 1.422 1.896 1.876 3.286 Ca 0.012 0.010 0.003 1.743 0.009 0.746 0.013 0.021 0.834 0.783 0.873 0.010 0 0.015 2.024 Na 0.003 0.064 0.012 0.595 0.005 0.201 0.023 0.079 0.146 0.138 0.046 0 0.081 0.075 0.680 K 0.895 0.961 0.934 0.258 0.907 0.002 0.884 0.821 0 0 0.001 0 0.838 0.839 0.167 Ba 0 0 0 0 0.017 0 0 0 0 0 0 0 0 0 0 Total 7.855 7.051 7.775 15.709 4.965 4.029 7.700 6.957 4.028 4.013 4.004 4.011 7.761 7.739 15.775 O 11 11 11 23 8 6 11 11 6 6 6 6 11 11 23 X' 0.464 0.528 0.268 0.321 0.028 0.232 0.337 0.333 0.136 0.138 0.172 0.281 0.237 0.222 0.215
Notes: FeO* - all iron as FeO. At formula calculation the amount of Zr and equivalent quantity of Si were rejected. X' is mole ratio of (Ca+Ba)/(Ca+Na+K+Ba) for K-feldspar, X' - iron number Fe''/(Fe''+Mg) for the remaining minerals.
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Table 9. Representative EPMA analyses of zircon grain from the geochronological sample 1111-08.
Zrn33 Zrn35 Zrn38 Zrn39 Zrn40 Zrn41 Location Rim Core Middle Middle Rim Rim Comment Area with the age ca 2.4 Ga Area with the age ca 1.9 Ga Oxides, wt.% SiO2 32.85 33.87 32.94 32.67 27.18 28.14 TiO2 0.06 0 0.03 0.02 0 0.03 Al2O3 0.16 0.06 0 0.08 0.20 0 Cr2O3 0 0 0 0 0 0.24 FeO* 0.28 0.13 0.39 0.03 1.09 0.76 MnO 0.36 0 0.02 0 0.33 0 MgO 0 0 0 0 0.12 0.13 CaO 0 0.22 0.04 0.02 1.26 1.13 Na2O 0 0 0 0 1.47 1.22 K2O 0 0 0.21 0 0.17 0.05 ZrO2 62.34 63.93 64.57 65.73 54.64 55.72 HfO2 1.57 1.71 1.54 1.68 0.53 1.43 ThO2 0 0 0.53 0.64 0.59 0 UO2 0 0.31 0.53 0.92 0.73 0.56 H2O** 0 0 0 0 10.13 9.26 Total 97.62 100.23 100.80 101.79 98.44 98.67
Cations Si 1.024 1.029 1.008 0.995 0.925 0.963 Ti 0.001 0 0.001 0 0 0.001 Al 0.006 0.002 0 0.003 0.008 0 Cr 0 0 0 0 0 0.006 Fe'' 0.007 0.003 0.010 0.001 0.031 0.022 Mn 0.010 0 0.001 0 0.010 0 Mg 0 0 0 0 0.006 0.007 Ca 0 0.007 0.001 0.001 0.046 0.041 Na 0 0 0 0 0.097 0.081 K 0 0 0.008 0 0.007 0.002 Zr 0.948 0.947 0.963 0.976 0.907 0.930 Hf 0.014 0.015 0.013 0.015 0.005 0.014 Th 0 0 0.004 0.004 0.005 0 U 0 0.002 0.004 0.006 0.006 0.004 Total 2.010 2.006 2.012 2.001 2.052 2.071 O 4 4 4 4 3.850 3.943 H2O*** 0 0 0 0 1.150 1.057
Notes: Zircon grain was studied also by SHRIMP, SIMS and ToF SIMS. FeO* - all iron as FeO. ** Water content was calculated in accordance of excess oxygen determined by EDS-EPMA. *** Structural formulae were calculated on the ZrSiO4·H2O crystallochemical basis.
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Graphical Abstract
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HIGHLIGHTS
• Meso-Neoarchaean HP/UHP Belomorian eclogite province. • More than one billion years evolution (3.0-1.7 Ga). • The HP/UHP processes in the Gridino area at continental subduction. • The time span of eclogite-facies metamorphism is 2.82-2.72 Ga • Modern style of plate-tectonics in Archaean
