The Meso-Neoarchaean Belomorian eclogite province: Tectonic position and geodynamic evolution

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The Meso-Neoarchaean Belomorian eclogite province: Tectonic position andgeodynamic evolution

Michael V. Mints ⁎, Ksenia A. Dokukina, Alexander N. KonilovGeological Institute RAS, Pyzhevsky lane 7, Moscow, Russia

a b s t r a c ta r t i c l e i n f o

Article history:Received 21 April 2012Received in revised form 10 October 2012Accepted 7 November 2012Available online 8 December 2012

Keywords:Archaean eclogiteNortheastern Fennoscandian ShieldPetrologyHigh-pressure metamorphismSubductionP–T–t path

The aim of this paper is to review the main features of the Meso-Neoarchaean Belomorian eclogite province(BEP) in the northeastern Fennoscandian Shield, including regional and local geology, geochemistry, petrologyand geochronology and to compare the Belomorian eclogites with Precambrian eclogites elsewhere. Twoeclogite associations have been recognized within Belomorian TTG gneisses: (1) the subduction-type Salma as-sociation and (2) Gridino eclogitized mafic dykes. Protoliths of the Salma eclogites represent a sequence com-prising gabbro, Fe–Ti gabbro and troctolites, formed at ~2.9 Ga in a slow-spreading ridge setting (like theSouthwest Indian Ridge). The main subduction and eclogite-facies events occurred between ~2.87 and~2.82 Ga. Injection of mafic magma into an active continental margin setting, recorded by the Gridino dykeswarm, is attributed to subduction of a mid-ocean ridge, commencing at 2.87 Ga. Crustal delamination of the ac-tive margin and subsequent involvement of the lower crust in subduction between 2.87 and 2.82 Ga ago causedhigh-pressure metamorphism of the Gridino dykes, culminating in eclogite-facies conditions between 2.82 and2.78 Ga and accompanying amalgamation of the Karelia, Kola and Khetolamba blocks and formation of theMesoarchaean Belomorian accretionary–collisional orogen. The clockwise P–T paths of the Salma andGridino as-sociations cross the granulite-facies P–T field. Detailedmetamorphic studies indicate a complicated post-eclogitehistory with thermal events and fluid infiltration, related to plume activity at 2.72–2.70, ~2.4 and ~1.9 Ga. Theeclogite assemblageswere exhumed tomid-to-lower crustal depths at ~1.7 Ga,while erosion or younger tecton-ic events were responsible for final exhumation to the surface. Comparison of P–T–t paths and data for peakmetamorphic parameters demonstrates the general similarity of the Archaean and Palaeoproterozoic eclogitesworldwide and their associationwith anomalously “hot” environments. The occurrence of high-T conditions dur-ing eclogite-faciesmetamorphism can be attributed to either subduction of amid-ocean ridge (Archaean, BEP) orto interaction with mantle plumes (Proterozoic).

© 2012 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction

The purpose of this paper is firstly to present a reviewof the recentlydiscovered Archaean Belomorian eclogite province, in the eastern partof the Fennoscandian Shield (Fig. 1). Then we further consider the tec-tonic significance of Archaean eclogites and the important implicationsthis has for the timing of onset of modern-style plate tectonics.

It was long considered that eclogite facies metamorphic regimeswould not have existed during the early stages of the Earth history,due to the inferred higher geothermal gradients (Green, 1975; Baer,1977 and others). However, Mesoarchaean eclogite xenoliths havebeen found in kimberlite pipes from the Kaapvaal and Siberia cratons,where the oldest diamonds containing inclusions of eclogite-faciesmin-erals (omphacite and garnet) have been dated at ~2.9 Ga (Gurney et al.,

2010 and references therein). Palaeoproterozoic eclogites were recog-nized for the first time in the NW Highlands of Scotland in 1984(Sanders et al., 1984; Sanders, 1988) within the Glenelg–Attadale base-ment inlier within the Caledonian Orogen. Proterozoic eclogites havesubsequently been found in many places (see below).

Early Precambrian eclogites were first described from the rocks ofthe eastern Fennoscandian Shield by Sudovikov (1936), Batieva(1958) and Smirnova and Baboshin (1967) but it was only recentlythat their petrological and isotopic characteristics attracted further at-tention. The first detailedmetamorphic petrogenetic study and attemptto date the Belomorian eclogites was reported by Volodichev et al.(2004),whodescribed the small eclogitic bodies that occurwithin a tec-tonic mélange zone near the Karelian village of Gridino on the westernshore of the White Sea. These authors concluded that the origin of thenewly discovered eclogites was caused by Archaean subduction(Volodichev et al., 2004, 2005). Archaean eclogites were subsequentlyfound, in 2002, by Vsevolod Barzhitsky, in a roadside outcrop near theSalma Strait in Lake Imandra in the Kola Peninsula (Shchipansky et al.,2005a,b). The above-mentioned papers by Volodichev and Shchipansky

Gondwana Research 25 (2014) 561–584

⁎ Corresponding author. Tel.: +7 4959513020; fax: +7 4959510443.E-mail address: [email protected] (M.V. Mints).

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and co-authors provided the impetus for numerous detailed studies ofwhat turned out to be the multiple eclogite bodies (Volodichev et al.,2004; Dokukina et al., 2005; Travin and Kozlova, 2005; Volodichev etal., 2005; Slabunov et al., 2006; Kozlovsky and Aranovich, 2008; Rosenet al., 2008; Volodichev et al., 2008; Aranovich and Kozlovskii, 2009;Dokukina et al., 2009; Volodichev et al., 2009; Dokukina et al., 2010;Kaulina, 2010; Kozlovsky and Aranovich, 2010; Kozlovsky et al., 2010;Mints et al., 2010a,b,c; Skublov et al., 2010a,b; Dokukina and Konilov,2011; Konilov et al., 2011; Mints et al., 2011; Mints and Konilov,2011; Morgunova and Perchuk, 2011a,b; Perchuk and Morgunova,2011; Skublov et al., 2011a,b; Slabunov et al., 2011; Dokukina et al.,2012a,b; Shchipansky et al., 2012a,b). Investigations carried outby these various research groups have identified the main characteris-tics of eclogite bodies, highlighted the main features of the igneous,

metamorphic and tectonic evolution of the eclogites and the eclogite-bearing associations in general, and reconstructed the main aspects oftheir geodynamic evolution. However, the interpretation of the datingresults, the nature and sequence of the metamorphic transformationsand reconstructions of geodynamic settings proposed by different au-thors show significant and in some cases, radical differences.

The aim of this paper is to outline and examine the most impor-tant characteristics of the Belomorian eclogite province (BEP). Wewill review the regional tectonic setting and geological features ofthe BEP, the sequence of major magmatic and metamorphic eventsand the respective geodynamic settings of their appearance and willattempt, based on available information, to reconstruct the historyof the origin and transformation of the various rock types found with-in BEP. We also discuss the reasons behind different interpretations of

Fig. 1. Simplified geological maps of the northeastern Fennoscandian shield, showing the tectonic position of the Belomorian eclogite province (for detailed 3D image of theBelomorian eclogite province, see Mints et al., 2009).Modified after Mints et al. (2010a).

562 M.V. Mints et al. / Gondwana Research 25 (2014) 561–584


the nature and evolution of the BEP and attempt to explain our posi-tion with respect to the alternative conclusions drawn by some of ourcolleagues.

2. Geological background

The Archaean Karelia and Kola continents (consisting predominant-ly of granite–greenstone terrains) and the Belomorian accretionary-collision orogen represent the principal tectonic units of the easternFennoscandian Shield (Mints et al., 2010b) (Fig. 1). Integration of theexisting geological database (Glebovitsky, 2005; Slabunov et al., 2006and references therein) with the results of recent reflection seismic in-vestigations of the crust and upper mantle along the 1-EU and 4Bgeotraverses (with a combined length of about 1300 km), has allowedconstruction of a three-dimensional model of the deep structure ofthe eastern Fennoscandian Shield (Mints et al., 2009; Mints, 2011 andreferences therein).

Tonalite–trondjemite–granodiorite (TTG) granitoids and gneissesof 2.89–2.72 Ga age dominate within the Karelian granite–greenstoneterrains. In the central part of the Karelia continent within the Kuhmo–Segozero microcontinent (see Mints et al., 2011; Mints, in press formany details), the greenstone belts consist principally of meta-sediments, mafic, komatiitic and intermediate-felsic metavolcanicsand banded iron formation (BIF) datedmainly at 2.76–2.74 Ga (the east-ern branch of the Kostomuksha belt, where the base of the volcanic–sedimentary sequence was dated at 2.79 Ga (Bibikova et al., 2005a,b)). A little younger paragneiss belts (2.75–2.72 Ga (Vaasjoki et al.,1993; Kontinen et al., 2007)) consisting mainly of metasediments,paragneisses and BIF, were formed within an intracratonic setting. Dis-tinctive sanukitoid intrusions of 2.74–2.72 Ga age are also characteristicof the final stages of the Archaean evolution in Karelia (Halla, 2002;Bibikova et al., 2005a; Heilimo et al., 2007).

The Inari–Kola microcontinent (granite–greenstone terrain) oc-cupies the western, central and southern parts of the Kola Peninsula.Granite–gneiss complexes (mainly TTG-gneisses) were formed be-tween about 2.9 and 2.73 Ga while poorly dated metasediments,metavolcanics and BIF assemblages in greenstone belts lie within the2.72–2.66 Ga interval (Mints et al., 2010b).

The Belomorian tectonic province has been of particular interest togeologists for a long period of time (Shchiptsov, 2005; Slabunov et al.,2006 and references therein). This province is distinguished by repeat-ed episodes of intense deformation and high- and moderate-pressuremetamorphism during both the Archaean and Palaeoproterozoictimes. A number of researchers have suggested that the Belomoriantectonic province represents a long-lived Belomorian mobile belt thatformed along the eastern margin of the Karelia continent as a result ofa succession of the tectonic events: westward subduction (in present-day coordinates) of the Archaean oceanic lithosphere beneath theKarelian continent, accretion of island–arc complexes to the Karelianmargin and final collision at approximately 2.75–2.65 Ga (Bibikova etal., 1999; Miller et al., 2005; Slabunov et al., 2006; Slabunov, 2008).However, this model is geometrically inconsistent with the crustal

scale seismic reflection fabric imaged in the 1-EU seismic profile cross-ing the Belomorian province (Mints et al., 2009).

In our understanding, taken together Meso-Neoarchaean com-plexes of the Belomorian tectonic province (in traditional meaning)represent accretionary-collision orogen that formed between ~2.88 Gaand ~2.76 Ga (Mints et al., 2010b). The Khetolamba microcontinent(granite–greenstone terrain) is one of the main elements of theBelomorian orogen. It contains three generations of TTG gneisses. Theearliest Group 1 TTG gneisses have an age of 3.12–3.11 Ga (Kröner andCompston, 1990). The Group 2 TTG gneisses formed synchronouslywith greenstone belts at 2.88–2.83 Ga (possibly starting as early asfrom 2.90 Ga) (Juopperi and Vaasjoki, 2001; Glebovitsky, 2005;Slabunov, 2008 and references therein). The 2.83–2.79 Ga Group 3 TTGgneisses occur as small bodies and dykes. Younger subalkaline granitoids(monzonites, syenodiorites) were intruded from 2.74 to 2.64 (locally upto 2.58) Ga (Mints et al., 2010b and references therein).

Synformal Kovdozero thrust nappe overlaps the border betweenthe Khetolamba microcontinent and the Karelia continent at thepresent-day erosion level. It comprises TTG-granitoids and gneisses,together with metasediments and metavolcanics in several green-stone belts, including the Iringora greenstone belt, which have beenassigned to the Parandovo–Tiksheozero island arc that was activefrom ~2.81 to 2.78 Ga (Mints et al., 2010b and references therein).The Iringora greenstone belt includes an ophiolite complex with anage of 2.78 Ga (Shchipansky et al., 2004).

The mafic–ultramafic Central Belomorian greenstone belt, separat-ing the Keret' and Khetolamba units has been dated at 2.88–2.85 Ga(Bibikova et al., 1999). The available geological, isotopic, and geochem-ical data from themafic–ultramafic rocks of the greenstone complex arecompatible with interpretation as tectonically disrupted and metamor-phosed remnants of a Mesoarchaean ophiolitic association (Slabunovet al., 2006). In turn, the Keret' tectonic nappe, sandwiched betweenKhetolamba and Inari–Kola microcontinents, contains 3.00–2.70 GaTTG-gneisses and greenstones, as well as numerous eclogite bodies,which are the main subject of this paper (Mints et al., 2010b) (Fig. 2).

The overall structural setting, combined with the above-mentioneddata and evidence for emplacement of eclogite bodies within TTGgneisses, above the Central Belomorian Belt, the eclogite occurrencesdescribed below, permit us to interpret this belt as a Mesoarchaean col-lisional suture zone. The reflection seismic data indicate that theKhetolamba microcontinent continues at depth to the northeast, be-neath the Inari–Kola microcontinent, although the boundary zone be-tween the two units, which coincides with the Central-Belomorianbelt, is not strongly expressed in the seismic image (Mints et al., 2009,2010b;Mints, 2011). Based on geological mapping and the results of re-flection seismic profiling, the crustal segment covered by the Keret' tec-tonic nappe has been interpreted an active margin of the ArchaeanInari–Kola microcontinent (Figs. 1, 2). According to our understanding,these three units (the Khetolambamicrocontinent, the eclogite-bearingKeret' active margin of the Inari–Kola microcontinent and the Central-Belomorian suture) represent a coherent Meso-Neoarchaen tectonicpackage that we refer to as the “the Belomorian eclogite province”(Mints et al., 2010a,b,c).

Fig. 2. Geological cross-section along the 1-EU reflection seismic profile. See Fig. 1 for location and legend.Modified after Mints et al. (2009).

563M.V. Mints et al. / Gondwana Research 25 (2014) 561–584


Within the Keret' nappe two distinct eclogite-bearing associationshave been distinguished (Fig. 1). The Salma association is attributed tosubduction and high-pressure metamorphism of the oceanic litho-sphere (e.g., Mints et al., 2010a,b,c; Konilov et al., 2011), whereasthe Gridino association is considered to represent injection of maficdykes followed by metamorphism under eclogite-facies conditions(Volodichev et al., 2004, 2005; Mints et al., 2010b; Dokukina andKonilov, 2011). Eclogite bodies belonging to the Salma association,are concentrated in the southern part of the Kola Peninsula, at UzkayaSalma and Shirokaya Salma, Upolaksha, Chalma (Kura-Vaara quarry)and Hangaz-Varaka occurrences. In addition, boudin-shaped eclogiticbodies of the same type are known in the area of Gridino village(Volodichev et al., 2004, 2005; Slabunov et al., 2006). In contrast,the main and most typical and abundant components of the Gridinoeclogite association are partially eсlogitized dykes of the Gridinodyke swarm, which are well exposed along the White Sea coastlineand adjacent islands. There has been considerable discussion in re-cent years concerning the timing of eclogite-facies metamorphismin the Belomorian province and whether it was Archaean or early-or late-Palaeoproterozoic in age? We return to this problem in detailbelow.

The synformal thrust nappe formed by kyanite gneisses of the Chupaseries occupies the axial area of the Belomorian tectonic province(Figs. 1, 2), which experienced high-pressure (up to 10–12 kbar)amphibolite- to granulite-facies metamorphism occurred at 2.71–2.70 Ga (Bibikova et al., 2004). The synformal structure of the Chupathrust nappe together with the geochronological data indicate thatthis unit formed after the termination of the Archaean collision, and inconnection with Neoarchaean plume-related evolution (Mints et al.,2010b, 2011).

The Karelia and Kola continents together with the Belomorianaccretionary-collision orogen placed between them are considered torepresent a segment of suggested Meso-Neoarchaean supercontinent(Mints, in press) that was later tectonically disrupted and partlyburied beneath intracratonic Palaeoproterozoic sedimentary–volcanicbelts, which can be interpreted as deformed rifts and, locally, as sutures.Sedimentary and volcanic successions with a predominance of maficvolcanics were formed at 2.50–2.32 Ga in Karelia (Sokolov, 1987;Hanski et al., 2001; Glebovitsky, 2005 and references therein) andfrom ~2.45 to ~1.77 Ga within the Pechenga–Imandra–Varzuga belt inthe Kola province (Balashov, 1996; Skuf'in and Theart, 2005). EarlyPalaeoproterozoic layered mafic–ultramafic bodies intruded to uppercrustal levels of the Kola–Karelia continent between 2.51 and 2.39 Ga(Huhma et al., 1990; Balashov et al., 1993; Amelin et al., 1995; Amelinand Semenov, 1996; Hanski et al, 2001; Bayanova, 2004 and referencestherein). Small bodies and dykes of gabbro-norite and lherzolite com-position, typically showing corona textures (local name of these rocksis “druzites”) intruded extensively the Belomorian crust at 2.46–2.43 Ga (Lobach-Zhuchenko et al., 1998; Sharkov et al., 2004 and refer-ences therein). At the same time (2.47–2.45 Ga) gabbro-anorthositebodies were intruded within lower crust and underwent granulite-facies metamorphism (Mints et al., 2007; Kaulina, 2010; Mints et al.,2010b and references therein). Such gabbro-anorthosite bodies arealso present within the basal part of the Lapland granulite belt. Maficmagmatismwas accompanied by intrusion of charnockites and granitesand eruptions of quartz-phiric porphyritic lavas (possibly ash-flows) at2.45–2.36 Ga (e.g., Levchenkov et al., 1994; Zlobin et al., 2005; Lauriet al., 2012).

2.1. The Salma eclogite association

The strongly retrogressed eclogites of the Salma association havebeen investigated in details at Uzkaya Salma and Shirokaya Salma(Mints et al., 2010a,b; Konilov et al., 2011), at Chalma (Kuru-Vaaraquarry) (Shchipansky et al., 2012a,b) and on Stolbikha and other

islands in the White Sea in the vicinity of Gridino village (Volodichevet al., 2004).

The best studied large eclogite bodies are exposed along the south-ern shore of the Uzkaya Salma Strait in Lake Imandra in the KolaPeninsula. The federal highway between St. Petersburg–Murmansktrends across the strike of these bodies at distance of 1192 km fromSt Petersburg (Figs. 2, 3). An eclogite-bearing association 300–500 mwide can be traced northeastwards for more than 4 km along theshores of the Babinskaya-Imandra and Ekostrovskaya-Imandra lakes.The host TTG gneisses vary in composition from quartz diorite totrondjemite and contain a diverse range of mafic eclogites, layersand lenses of Fe–Ti eclogites and high-Mg eclogite-facies rocks(piclogites), layers and lenses of garnetites, and garnet-bearing andgarnet-free amphibolites. Banding within the gneisses and the con-tacts of the eclogite bodies dip steeply towards the north–northwest.The Archaean rocks are discordantly cut by vein- and lens-like bodiesof the late-Palaeoproterozoic granitoids and pegmatites intrudedunder amphibolite-facies conditions. Most of the eclogite bodies arecharacterized by strong retrogression and formation of amphibolites,with only locally preserved relict domains of high-pressure assem-blages. Retrogressed eclogites typically consist of poikilitic garnetporphyroblasts (4–5 mm) set in a fine-grained light-green matrix ofNa-poor clinopyroxene–plagioclase symplectite (after omphacite), andminor amphibole and quartz. Olive-green hornblende, plagioclase, andilmenite replace clinopyroxene–plagioclase symplectite, while garnetis replaced by kelyphitic rims of plagioclase with hornblende. Maficeclogites are the predominant type among the eclogite bodies. Themarginal zones of the larger eclogite bodies are characterized bygarnet-bearing or garnet-free amphibolites, while their interiors are insome cases amphibolized also. Amphibole growth is also common inthe exocontact zones of the late-Palaeoproterozoic quartz and pegma-tite veins andwithin local fracture zones. Mafic eclogites are often asso-ciated with high-Mg rocks (piclogites). Massive coarse- and medium-grained Fe–Ti eclogites can be readily distinguished owing to their“rusty” color. Garnetites consisting of garnet (up to 90%), plagioclase,pyroxene and quartz, with a conspicuous amounts of rutile, formlenticular bodies or extended and boudinaged “layers” and occasionallyintersecting bands ranging in thickness from a few centimeters to 1 mor more. In at least some cases, garnetites were formed by metamor-phism of Fe–Ti gabbro with titanomagnetite and ilmenite–magnetitemineralization. It is also possible that some garnetite layers representmetamorphosed veins that originally formed in the gabbroic protholithsduring hydrothermal–metasomatic processes in the ocean floor or sub-seafloor environment.

Mafic and ultramafic eclogite-facies rocks can be assigned geo-chemically to the tholeiite series. The mafic eclogites are resemblinglow-K gabbro-norite or tholeiitic basalts (Fig. 4). The high-Mg rockscontain b1% TiO2 and can be regarded as picritic (12%bMgOb18%)or komatiitic (>18% MgO) composition. Protholiths of the Fe–Tieclogitic rocks would have resembled olivine gabbro or troctolite,but have significantly elevated FeO* and TiO2 contents: 18–21% and1.5–2.5%, respectively.

The trace-element abundances point toward a mantle derivationfor the magmatic precursors, but are lower than those of N-MORB.Comparison of spider-diagrams describing all three Salma eclogitetypes demonstrates their unity and coherence of the geochemicalcharacteristics (Fig. 5). Negative Th anomalies and positive Nb anom-alies area are also characteristic for the Salma eclogites as a whole.Consistent positive U, Zr and Hf anomalies and negative Ba anomaliesalso emphasize the close relationship between piclogites and Fe–Tieclogites. On the Th/Yb versus Nb/Yb diagram of Gorton and Schandl(2000) (Fig. 6) the Salma eclogites are distributed around the meancomposition of modern MORB, whereas when plotted on the Nb/Y ver-sus Zr/Y projection of Fitton et al. (1997), the Salma eclogites fall intothe field of the mantle-plume array (Fig. 7). These features may implythe involvement of a mantle plume contribution in the source of the



eclogitic protholiths (Mints et al., 2010b). Layering formed by alternat-ing of “normal”, high-Mg and Fe–Ti eclogitic rocks together with theircompositional and structural features suggest that the original protolithassemblage might have consisted of intercalated “normal” gabbro-norite, Fe–Ti oxide gabbro, and olivine gabbro with local troctoliticrocks, resembling the gabbroic suite (seismic layer 3) from the modernoceanic crust of the slow-spreading Southwest Indian Ridge (Dick et al.,2000).

Several tens of eclogite blocks are randomly distributed within theTTG gneisses exposed in rock faces of the Kuru-Vaara quarry (theChalma eclogite locality). The kyanite- and orthopyroxene-bearingeclogite varieties known in the Kuru-Vaara quarry have not been

found at any of the other eclogite localities. Some of the eclogiteblocks bear evidence of the partial melting and initial segregation offelsic melt before or during eclogite-facies metamorphism, with theformation of veins and melt percolation channels (Shchipansky etal., 2012a,b).

2.2. The Gridino eclogitized dyke swarm

Eclogite-bearing mafic dykes of the Gridino dyke swarm are alongthe White Sea coastline and adjacent islands near the village ofGridino (Fig. 8). These dykes have been illustrated in detail in a num-ber of previous works (Sibelev et al., 2004; Volodichev et al., 2004,

Fig. 3. Location of the Uzkaya Salma eclogite assemblage: outcrop along the northeastern roadside along the St. Petersburg-Murmansk highway.Compiled by K.A. Dokukina, N.E. Kozlova and O.N. Platonova, photo: K.A. Dokukina.



2005; Slabunov et al., 2006; Sibelev, 2007; Travin and Kozlova, 2005;Stepanova and Stepanov, 2010 and references therein; Dokukina andKonilov, 2011). A tectonic mélange zone some 50 km long and 10 kmwide can be traced along the White Sea coast from north-west tosouth-east. Field observations and regional geological mapping haverevealed that the mélange zone is associated with a system of tectonicslices dipping northeastward. The mélange zone consists of a mixtureof migmatized granite-gneisses with abundant amphibolite enclavesas well as diverse ortho- and paragneisses with extensive lenses ofintensely deformed rocks. The dyke swarm, which has submeridionaltrend (Fig. 8) consists predominantly of gabbronorite dykes, withmetagabbro ones. The strain state of dykes varies considerably. Un-deformed dykes retain intrusive contacts and discordantly truncatebanding in the felsic gneisses. Deformed dykes vary in thicknessand degree of deformation, including folding, boudinage, andmigmatization. Extreme deformation of the dykes led to their break-up into pods and lenses concordant with the foliation of the hostgneisses (Fig. 9). Different dykes and various mafic pods and lensesrecord successive metamorphic events at eclogite, high-pressuregranulite, and amphibolite facies respectively.

Before addressing the key elements of the debate over the origin andsignificance of the eclogites, which is the main theme of this paper, weshould emphasize that, startingwith the paper by Stepanov (1981), theGridino eclogitized dyke swarmwas considered and treated as a part ofthe previously mentioned Palaeoproterozoic mafic intrusive bodies re-ferred to as “druzites” (Stepanova and Stepanov, 2010 and referencestherein), despite the obvious differences in metamorphic grade andthe spatial and structural isolation of the eclogitized dykes with respectto the area of druzite emplacement. It should also be noted that theoriginal assignment of the Palaeoproterozoic age of the dykes wasbased solely on the geological inferences, such as the distribution, ap-pearance and morphology of intrusive bodies, the relationships withthe enclosing environment and the bulk composition of the rocks. How-ever, this information alone is not sufficient for correlation between theeclogitic dykes and the druzites. Therefore, it is necessary to considertheir geochemical characteristics and affinities, as reviewed in detailby Mints et al. (2010b) and Dokukina and Konilov (2011). Themetagabbroic and metagabbronoritic dykes and the Gridino eclogitebodies, for which subduction origin is proposed, belong to thetholeiite series. A distinctive feature of the metagabbronoritic dykes istheir high-Mg and high-Cr (Mg#=0.65–0.78; Cr — 940–1960 ppm),combined with low-Ti (TiO2 — 0.41–0.85%). In contrast, many meta-gabbros are significantly enriched in FeO* (up to 16%) and TiO2 (up to

2.0–2.5%). For comparison, the AFM diagram in Fig. 4 also showsthe Palaeoproterozoic “druzite” complex. The “druzite” compositionsare located within two non-overlapping fields that correspond, respec-tively, to the gabbro-anorthosite and lherzolite–gabbronorite com-plexes. The compositional fields of the Gridino eclogitized dykeswarm and the Palaeoproterozoic “druzites” differ significantly, withthe eclogitized dyke compositions plotting preferentially within thefields for intraplate basalts and basalts from active continental margins.On the Nb/Y–Zr/Y diagram (see Fig. 7), the datapoints for the dykes arebroadly distributed across the fields for both primitive mantle (PRIMA)and continental crust. On the Th/Yb–Ta/Yb diagram (Fig. 6), the maficdykes also show a wide range in composition, overlapping with thefields for both the Gridino subduction-type eclogites as well as that ofcrustally contaminated active continental margin volcanics. It is there-fore deduced that there is a distinct relationship between the protolithof the subduction-type eclogites and the mafic magmas that formedmafic dykes. The primary geochemical data and a set of geochemicaldiscrimination diagrams can be seen inMints et al., 2010b). It seems ev-ident that for a “final decision”, this problem needs a strict geochrono-logical basis.

3. Pressure–temperature evolution

Thermo-barometrical studies of the Belomorian eclogites wereperformed mainly using a multicalibration approach, based on theconsistent system of geothermometers and geobarometers of theTPF program developed at the Institute of Experimental Mineralogyof RAS (Fonarev et al., 1991, 1994; Konilov, 1999; Maaskant, 2004).

For detailed descriptions of the thermobarometric investigations,refer to Volodichev et al. (2004), Konilov et al. (2011), Dokukinaand Konilov (2011) and Shchipansky et al. (2012a).


The main features of the metamorphic evolution of the Salmaeclogites are shown in Fig. 13.

3.1.1. Prograde evolutionDiaspore inclusions in spinel from the Uzkaya Salma piclogite sug-

gest that the protolith, prior to being subducted, may have beenabove sea level and exposed to weathering and laterization. A moremodern analog for such a rock can be seen in the metabauxites de-scribed from Samos and Naxos islands, Greece (Feenstra, 1997; Urai

Fig. 4. AFM diagram: petrochemical types of the rocks in the Salma and Gridino eclogite associations.After Kuno (1968).



and Feenstra, 2001). The formation of ancient laterite containing ox-ides or oxyhydroxides was previously known in the 2.2 Ga Hekpoortpaleosol of the Transvaal Supergroup in southern Africa (Beukes et al.,2002). It would imply an oxygenated atmosphere which is not widelyaccepted for the Earth at 2.8 Ga. Additional information about theearliest stages of metamorphism is provided by mineral inclusionspreserved in the prograde-zoned atoll-type garnets from the UzkayaSalma locality. These low P–T minerals are pumpellyite, actinolite,and albite, zoisite and chlorite, which occur either as isolated mineralphases or as distinct relict mineral assemblages armored by the atollgarnet. The textural evidence clearly suggests that the pumpellyite–actinolite assemblage was replaced by later hornblende, not the re-verse. Moreover, there is textural and compositional evidence indi-cating that these inclusions were entrapped during prograde growthof the atoll-type garnet. Thus, these phases can be interpreted as apart of the prograde mineralogy (see e.g., Page et al., 2003, 2007) andto indicate that a mafic protholith has passed through pumpellyite–actinolite metamorphic facies during its prograde path. The low-T alter-ation in the gabbroic suite of the slow-spreading Southwest Indian

Ridge is localized and typically confined to fractured regions where in-tense alteration of the host rocks can be observed adjacent to veins/veinlets filled with smectite, smectite–chlorite mixed layer minerals,or chlorite±calcite±zeolite±sulfide±Fe-oxyhydroxide (Dick et al.,2000; Bach et al., 2001).Most of theminerals listed above are common-ly considered to indicate sea-floor metamorphism (e.g., Nozaka et al.,2008). Several reports of similar low P–Tmineral assemblages includedin garnet fromPhanerozoic eclogites have been recently published (e.g.,Enami et al., 2004; Yang and Powell, 2008). A later prograde event hasbeen documented by the clinopyroxene that armors chlorite in thecores of garnets from the Salma eclogites. The maximum temperaturecalculated from the Grt–Cpx assemblage was 636 °C at 10 kbar refer-ence pressure. These aspects of the Salma eclogites have been recentlydescribed in greater detail by Mints et al. (2010b), Konilov et al.(2011) and Shchipansky et al. (2012a).

Fig. 5. Primitive mantle (Hofmann, 1988) normalized trace-element spectra in eclogites.After Mints et al. (2010b).

Fig. 6. Ta/Yb–Th/Yb plot. IOA — intraoceanic arcs; ACM — active continental margins;WPVZ — within-plate volcanic zones; and WPB — within-plate basalts.After Pearce (1983) and revised by Gorton and Schandl (2000). The Salma and Gridinoeclogite association based on data from Volodichev et al. (2004), Slabunov et al. (2008),Mints et al. (2010b) and Dokukina and Konilov (2011).

Fig. 7. Zr/Y–Nb/Y plot.After Fitton et al. (1997). The Salma and Gridino eclogite association based on data fromVolodichev et al. (2004), Slabunov et al. (2008), Mints et al. (2010b) and Dokukina andKonilov (2011). Primitive mantle and N-MORB after Hofmann (1988) and continentalcrust after Rudnick and Fountain (1995).



3.1.2. Eclogite eventAlthough the localities studied contain mainly retrogressed

eclogites with almost complete replacement of primary matrixomphacite by diopside-plagioclase symplectite, petrographic studysuggests that in all instances the HP stages were plagioclase-free.The minimum pressure for the eclogite stage was estimated as a func-tion of the Jd content in clinopyroxene coexisting with quartz and al-bite. Inclusions of omphacite were observed within garnet in severalthin sections, enabling us to restrict the estimates for the peak P–Tconditions to this assemblage. For the Shirokaya Salma eclogite, thepeak pressures are inferred to have been ~13 kbar at a temperatureof 700 °C. The samples from the Uzkaya Salma eclogite yield pres-sures of 13–14 kbar over the temperature range 700–750 °C. Thesamples from the Chalma (Kuru-Vaara quarry) eclogites give slightlyhigher temperatures of 750–775 °C and pressures of 14–14.3 kbar.

We consider that all these pressures should be regarded as minimagiven the absence of plagioclase equilibration with the assemblage.Relict kyanite occurs amongst the diopside-plagioclase symplectites,indicating that omphacite was abundant in the peak-pressure assem-blage, although the composition of the peak-pressure omphaciteformerly associated with kyanite remains unknown. Omphacitewith a jadeite component >32% has been found as a single inclusionwithin garnet. A further important observation is that relict matrixomphacite (Jd up to 28%) preserved in a sample from the Kuru-Vaara quarry shows evidence for quartz exsolution (Konilov et al.,2011; Shchipansky et al., 2012a). Similar exsolution microstructureshave been found in eclogite from several UHP metamorphic belts,including the Kokchetav Massif in Kazakhstan (Katayama andMaruyama, 2009), Pohorje, in the Eastern Alps (Janák et al., 2004),Alpe Arami, Switzerland (Dobrzhinetskaya et al., 2002) and theBlumenau eclogite, Erzgebirge (Chopin and Ferraris, 2003); most au-thors consider these features to be diagnostic of UHP metamorphism.

3.1.3. Retrograde evolutionThe eclogite decompression has promoted omphacite breakdown

and the production of plagioclase to form diopside-plagioclasesymplectite. A pressure range from 10 to nearly 13 kbar at 750 °C wasinferred for symplectitic mafic and Fe–Ti eclogites from the ShirokayaSalma and Uzkaya Salma localities. Lower values for these parameterswere obtained for garnet rims and Cpx–Pl–Qtz coronitic assemblages,within the range 7–10 kbar and 670–700 °C. Hornblende- and biotite-bearing assemblages record temperatures generally from 650 °C to700 °C at inferred pressures of 7–9 kbar. However, it remains unclearwhether these assemblages represent continual cooling with a slightdecrease in pressure or the result of later metamorphic events. Thus,the retrograde paths of the eclogite bodies studied should in generalpass through the garnet-amphibolite facies to upper amphibolite facies,thus representing near-isothermal decompression. Another kind ofevolution has been recognized for eclogites from the Shirokaya Salamaand Chalma (Kuru-Vaara quarry) localities, which experienced de-compression coupled with heating. The presence of orthopyroxeneintergrown with the plagioclase–diopside symplectite implies retro-gression through the granulite facies at temperatures of 730–780 °C.Small temperature increases up to medium-grade granulite faciesafter attainment of peak pressure has been reported from several Phan-erozoic subduction-related eclogites (e.g., Page et al., 2003, and refer-ences therein) and from almost all known Precambrian occurrences(e.g., Möller et al., 1995; Collins et al., 2004; Carlson et al., 2007;Mosher et al., 2008). We will discuss this evolutionary feature and itssignificance later.


The main features of the metamorphic evolution of the Gridinoeclogitized dykes are shown in Fig. 13. Studies of thin sections fromthe dykes clearly reveals that all of the dykes examined underwentsuccessive metamorphic transformations from eclogite throughHP-granulite to amphibolite facies. They also display a structural het-erogeneity, which is more pronounced in the larger bodies. In theundeformed thin dykes granulite assemblages predominate: garnet–orthopyroxene–clinopyroxene (not omphacite) with rare relics ofclinopyroxene–plagioclase symplectite. Within the larger dykes(25–30 m thick), however, all stages of the metamorphic evolutionmay be discerned (Dokukina and Konilov, 2011; Dokukina et al.,2012a; Dokukina et al., 2014–this issue; see also Volodichev et al.,2005, 2008).

3.2.1. Igneous stageIgneous structures and mineral assemblages are commonly pre-

served in the central parts of the dykes, typically comprising ortho-pyroxene and pyroxenes with inverted pigeonite and augite exsolution

Fig. 8. Sketch-map of the Gridino dyke swarm. Location of the image indicated by as-terisks under the numbers 6 and 8 in Fig. 1, right.After Mints et al. (2010b).

Fig. 9. The metagabbro dyke on Vorotnaya Luda island. Undeformed part of the dykehas intrusive contacts and crosscut the felsic gneiss. Extreme deformation of thedyke led to segmentation into pods and lenses concordant with the foliation of thehost gneisses.



textures. A crystallization temperature of about 1030 °C has been esti-mated for quartz-bearing dykes and 1200 °C for olivine-bearing dykes.A minimum temperature of 600 °C at 5 kbar for breakdown of augiteand pigeonite corresponds to a post-magmatic stage of isobaric cooling.

3.2.2. Incipient eclogitizationA garnet coronawith inclusions of kyanite, plagioclase, clinopyroxene,

and quartz formed at the contacts between igneous pyroxene and plagio-clase records this earliest stage of eclogitization.

3.2.3. Eclogite eventAs the eclogite assemblage evolved further, coronitic garnet with

kyanite inclusions completely replaced igneous plagioclase. The rockmatrix is an aggregate of orthopyroxene–clinopyroxene–plagioclasesymplectite, which is typical of strongly retrogressed eclogites. Inrare cases omphacite relics (with up to 36 mol% jadeite) are pre-served in the symplectite and as inclusions in garnet, together withkyanite and quartz. Some omphacite grains in the matrix contain ori-ented quartz needles. As mentioned above, similar exsolution micro-structure has been found in eclogites from several UHP metamorphicbelts.

3.2.4. Granulite eventOrthopyroxene also occurs within the clinopyroxene–plagioclase

symplectite, providing evidence for a retrograde metamorphic trendin granulite facies P–T space (Page et al., 2003; Groppo et al., 2007).Orthopyroxene also forms coronas around garnet and is found withingranoblastic orthopyroxene–clinopyroxene domains. In some placesdykes have been totally recrystallized under granulite-facies condi-tions, with the formation of garnet-two-pyroxene paragenesis.

3.2.5. Amphibolite-facies recrystallizationDyke contacts with host gneisses are usually characterized by an

amphibolite assemblage, in a zone up to 10 cm wide. Amphibolite fa-cies assemblages are also present in later fractures and along latefelsic pegmatite and carbonate veins. Temperature and pressure esti-mates for amphibolite grade metamorphism of mafic and felsic rocksare 530–660 °C and 7.9–9.6 kbar (Sibelev et al., 2004; Volodichev etal., 2004).

3.3. The Gridino host gneisses

The mineralogy of the host gneisses is generally monotonous: bi-otite, garnet–biotite, or garnet–hornblende plagiogneiss of tonalitecomposition. Typically, the gneisses are migmatized and eclogite-facies assemblages have not been found, which is a characteristic fea-ture of the BEP as well as some other eclogite-bearing terrains. Thequartz-rich rocks are more susceptible to dynamic recrystallizationand retrograde metamorphism than rocks of mafic composition(Koons and Thompson, 1985). In rare cases, a granulite-facies mineralassemblage is present in the gneisses, represented by biotite-garnet-or kyanite-garnet mineral parageneses. Generally, the temperaturesand pressures derived from the mineral assemblage are consistentwith those of eclogites retrogressed under amphibolite facies condi-tions. Granulite-facies mineral assemblages in the gneisses corre-spond to temperatures and pressures of 750–800 °C and 10–12 kbar(Dokukina and Konilov, 2011).

3.4. Granite leucosomes

Evidences for the development of successive metamorphic eventsare preserved in migmatitic felsic rocks that formed under high pres-sure conditions. A granitic leucosome developed at the continuationof a metagabbroic dyke at Cape Vargas, to the northwest of Gridinovillage, is of particular interest. Petrological studies of the leucosomeshowed relics of high-pressure igneous minerals: Ba-bearing phengite,

K-feldspar and K–Ba feldspar, myrmekite, and near-solidus symplectiticintergrowths of clinozoisite, phengite, and quartz (for more details seeDokukina and Konilov, 2011). Similar clinozoisite-quartz symplectiteswere described from the UHP diamond-bearing clinozoisite gneissesof Kokchetav massif in northern Kazakhstan (Korsakov et al., 2002,2006). The phengite geobarometer (Caddick and Thompson, 2008)yields high-pressure conditions for leucosome crystallization, over arange of 15 to 25 kbar and 650–750 °C. Evidences of the transitionfrom eclogitic to granulite-facies conditions during prograde decom-pression have also been recorded (750–800 °C and 10–12 kbar): biotitereplaces phengite; grossular garnet and clinopyroxene replaceclinozoisite-quartz symplectite; plagioclase breaks down, with the for-mation of antiperthite. The last metamorphic event recorded by theleucosome is a secondary amphibolitization, with amphibole replacingclinopyroxene and the crystallization of new biotite and epidote.

The eclogitized olivine gabbronorite dyke in the eastern outskirtsof the Gridino village, is crosscut by an enderbite vein, for which pres-sure and temperature estimates of about 750 °C and 12.6 kb wereobtained, indicative of high-pressure granulite conditions (Dokukinaand Konilov, 2011). In a detailed study of this dyke, a series ofveins, confined to a discrete zone of deformation was also found. Out-side these veins, the dyke consists of orthopyroxene eclogite withoutquartz. In contrast, the veins are quartz-bearing symplectic eclogiteswith the linear veinlets of quartz–biotite–plagioclase or pure quartz.Along the quartz veinlets, elongate clusters of garnet and rutile grainshave crystallized,with orthopyroxene–plagioclase reaction rims formingbetween quartz and garnet. Relics of omphacite (Jd up to 30%) arepreserved in the two-pyroxene–plagioclase symplectite, while garnetcontains inclusions of kyanite and omphacite. Crystallization P–Tparameters of the symplectitic mineral assemblage correspond to9–10ce:hsp sp="0.25"/>kbar at 700–750 °C. Within the symplectites,the Cl-apatite grains (Cl up to 6.37%) are ubiquitous, together withrutile, high-Ti biotite (TiO2 up to 7%) and linear chains of zircons(Dokukina et al., 2012b).

4. Geochronology

The BEP provides unique evidence of Meso-Neoarchaean subduc-tion and collision. However, even though the geochronological dataobtained by different researchers generally coincide, there havebeen some fundamental differences in interpretation of the results.

Recently, three different models for the nature and the age of themetamorphic events in the BEP, have been presented and discussed,particulary in some Russian journals. The first model suggests threedistinct episodes of metamorphism under eclogite-facies conditions:two events relating to subduction processes in the Archaean, at2.82–2.80 Ga and 2.72–2.70 Ga, followed by a distinct event thatcaused metamorphism of the Palaeoproterozoic Gridino mafic dykeswarm at ~2.45–2.40 Ga (Volodichev et al., 2004; Slabunov et al.,2011; Shchipansky et al., 2012b). The second model proposes a singleeclogite-facies metamorphic event within the Belomorian province at~1.9 Ga (Skublov et al., 2010a,b, 2011a,b). In the third model, devel-oped by the authors of this paper, considers the BEP as a combinationof the subduction-type eclogites and eclogitized mafic dykes areinterpreted as the product of a succession of closely interrelatedevents within a single tectonomagmatic cycle. The magmatic process-es and the major eclogite-facies events were restricted within thenarrow time span between ~2.9 and ~2.8 Ga.

Most of the geochronological data referred to below have alreadybeen published elsewhere. The U–Pb dating and Lu–Hf-isotope analy-ses of zircons from the various types of the Salma eclogites and somerocks from the Gridino area have been undertaken at the GEMOC KeyCenter, Macquarie University (Sydney, Australia) using the laser abla-tion technique LAM-ICPMS (Mints et al., 2010a,b,c, see also the paperDokukina et al., 2014–this issue). The zircons frommost other samplescollected from the eclogites of the Salma and Gridino associations,



host gneisses and crossing felsic veins were studied using the SHRIMPII system at VSEGEI (Saint-Petersburg, Russia) (Kaulina et al., 2007;Slabunov et al., 2008; Dokukina et al., 2009, 2010; Kaulina, 2010;Skublov et al., 2010a,b; Dokukina and Konilov, 2011; Konilov et al.,2011; Skublov et al., 2011a,b,c; Slabunov et al., 2011; Dokukina etal., 2012a,b; Shchipansky et al., 2012b). Some zircons were alsodated by NORDSIM at the National Museum of the Natural History(Stockholm, Sweden) (Volodichev et al., 2004) and at the GeologicalInstitute of the Kola Science Centre RAS (Apatity, Russia) (Kaulina etal., 2007; Dokukina and Konilov, 2011; Dokukina et al., 2012a,b). Inthis section we summarize the available geochronological data.


Recent investigations have established crystallization and recrys-tallization of specific zircon populations related to the successiveevents in the evolution of the Salma eclogite association (Fig. 10)(Kaulina, 2010; Kaulina et al., 2010; Mints et al., 2010a,b,c).

(1) Zircons with characteristic patchy textures and numerous voidsand mineral inclusions (quartz, plagioclase, clynopyroxene, rutile,calcite, muscovite, apatite, Al-titanite, orhtite, epidote, pyrite,galena) (Fig. 10a, b) and high (magmatic-type) concentrationsof trace element admixtures including HREE. The most reliableestimate for the age of crystallization age is ~2.89 (possibly

2.9) Ga. A concordant point at 2.82 Ga might represent theage of the earliest metamorphism, although there is as yet noevidence to preclude this age from relating to peak or retro-grade metamorphism. The initial 176Hf/177Hf isotopic ratios inzircons show negligible variations compared with the rangeof measured U–Pb ages, providing evidence that the U–Pb sys-tematics were disturbed, while the Hf isotope system remainedrobust. Estimates for the model TDM age are confined to withinthe 2.88–2.66 Ga time span, hence the emplacement of melt inthe crust immediately followed its separation from the mantlesource. The values of εHf vary widely from−16 to +8, but thesignificance of this is not yet understood (Fig. 11) (Mints et al.,2010a). Metamorphic transformations including those in theeclogite facies led to a gradual decrease in the abundance oftrace elements and reduction of the Th/U ratio from 2.5–3.0to 0.1–0.2. The crystal rims and recrystallized zones withinthe zircon crystals are related to later events dated at 2.5–2.4 Ga, 2.3–2.2 Ga and 2.0–1.9 Ga (Mints et al., 2010b). Zirconsof this type have been first extracted from Fe–Ti eclogites inves-tigated in the Uzkaya Salma locality. Subsequently, morphologi-cally similar zircons of the same age have been recognizedin the cores of crystals in the “normal” mafic eclogites fromthe Chalma (Kuru-Vaara quarry) locality (Shchipansky et al.,2012b). The above-described features related to themorphology,structure and geochemistry of zircons are considered to have

Fig. 10. Specific zircon populations that are related to the successive events in evolution of the Salma eclogite association. Refer to explanation in the text.After Mints et al. (2010a,b,c), Kaulina (2010), and Kaulina et al. (2007, 2010).



formed during the magmatic crystallization involving highlyfractionated residual melt or fluid in the shortage of space. Asconditions changed, the zircon crystals were partially recrys-tallized, although the zircon-bearing rock apparently did notundergo partial melting and fluid activity. Similar features ofrecrystallized late-magmatic zircon inmetamorphic rocks formedat high P and/or T in various regions (e.g., Hoskin and Black, 2000;Tomashek et al., 2003; Kaczmarek et al., 2008; and referencestherein).

(2) Almost homogeneousmultifaceted rounded “granulitic-type” zirconsthat show either an unzoned internal structure or an indistinct weakzoning (Fig. 10c, d). Low trace element concentrations in thesezircons suggest crystallization in the equilibrium with garnet.Two types of zircons have been recognized: colorless low-U zir-con with Th/U ranging from 0.16 to 0.47, and dark-brownhigh-U (up to 900 ppm) grains with lower Th/U ranging from0.12 to 0.33. Based on their morphology, these zircons appearto be metamorphic and could have crystallized under high-pressure granulite- or eclogite-facies conditions (e.g., Corfu etal., 2003; Bibikova et al., 2004). However, the close associationof zircons of both types is more consistent with crystallizationunder granulite-facies conditions (Kaulina, 2010). Age estimatesfor these zircons based on both LAM-ICPMS and SHRIMPmethods are coincident at 2.72–2.70 Ga. On the 176Hf/177Hf–U–Pb age diagram (Fig. 11), “granulitic” zircons are generally locat-ed between the CHUR and DM evolution lines (Mints et al.,2010a). Values of εHf vary from +1 to +4, and the model TDMage is 2.77–2.90 Ga. The significant difference between theLu–Hf model age and the U–Pb age of magmatic crystallization(up to180 Ma) provides evidence that “granulite” zircon crystal-lized a long time after the formation of the igneous host-rocks.The lower intersection of the discordia yields 1.91 Ga as the ageof the final metamorphic event.The 2.72–2.70 Ga zircons of this type were found in theShirokaya Salma (Kaulina et al., 2007; Kaulina, 2010; Kaulina etal., 2010; Mints et al., 2010a), Chalma (Kuru-Vaara quarry)(Shchipansky et al., 2012b) and Gridino (Volodichev et al.,2004) eclogites, all of which were reworked under granulite-facies conditions, including eclogitized mafic dykes (Dokukinaand Konilov, 2011). The corresponding varieties of rocks can bedenoted as “eclogite-granulites”. It is interesting and importantto note that the Type 1 and Type 2 zircons described here havenot been found together within the same samples. We inferfrom this that “granulitic” zircons crystallized utilizing Zr liberat-ed from the eclogite-forming minerals in rocks, which originallyhad no magmatic zircons.

(3) Zircons from the garnetite are irregular to tabular. Many crystalshave inclusion-filled cores similar to the Type 1 zircons, collect-ed from the Fe–Ti eclogite, surrounded by broad, transparent,structureless cracked rims (Fig. 10e, f). Some crystals havebeen completely altered. The geochemical features and isotopicsystematics of the zircon crystals, including their relict oldercores were substantially modified during Palaeoproterozoicmetamorphism at about 1.89 Ga. Unlike the other samples,this population shows a large range in εHf, from −14 to +14(Fig. 11). The lowest values are similar to those of the Fe–Tieclogite; the highest are well above the depleted mantle evolu-tion line (Mints et al., 2010a). The large range of εHf suggeststhat the zircons in the garnetite were generated by the localbreakdown of older Zr-bearing minerals with a range ofLu/Hf, as has been documented for zircons from garnetite xeno-liths in basalts of the Four Corners area of the United States(Smith and Griffin, 2005). Drastic modification of the isotopic–geochemical properties of zircons may be related to the forma-tion of zones of increased fluid permeability in the coarse-grained garnetites. They might play a role of distinctive ballbearingswithin theweakened zoneswhere the Palaeoproterozoicshear strain was localized.Similar processes of zircon recrystallization in eclogites can beseen in the description of the zircons from two samples fromthe Chalma locality (Kuru-Vaara quarry) that contain multiface-ted round and oval-shaped zircons that preserve the 2.88 Gacores with numerous inclusions (Skublov et al., 2011b). Datingof zircon rims from eclogite and zircons with similar morphologyfrom the eclogite-like garnet–amphibole–clinopyroxene rockyielded ages of 1.91 and 1.84 Ga, respectively. It is interesting tonote that the recrystallization event under granulite facies condi-tions in the samples of Skublov and co-workers can only be recog-nized due to the roundish to oval-shaped morphology of thezircons (resembling Type 2 crystals) but not in the isotope-geochronological characteristics of zircon. As mentioned above,Shchipansky and co-authors also dated Type 2 “granulite” zirconsin the samples from the same locality, these zircons andmorpho-logically similar zircons from other localities within theBelomorian province yielded exactly the same age, 2.72 Ga.Apparently, as in the case of the zircon from garnetites, theisotope-geochronological characteristics of the zircons in thesamples analyzed by Skublov and co-workers could have beencompletely reworked during the latest metamorphic event.


The zircons in the mafic eclogitized dykes of the Gridino swarmshow greater morphological diversity when compared with those ofthe Salma eclogites. This makes it more difficult to correlate individu-al zircon populations with specific magmatic and/or metamorphicevents and engenders considerable discussion on the provenance ofdifferent zircon populations in the eclogitized dykes and, hence, onthe age of the dyke intrusions and their sequential reworking underthe eclogite-, granulite- and amphibolite-facies conditions.

To date, the results of the detailed geochronological studies of threeeclogitized dykes have been published: from the cape on the northeast-ern outskirts of Gridino village (dyke B-16 (Volodichev et al., 2009;Slabunov et al., 2011)), at Cape Vargas (dyke D17 (Dokukina et al.,2009; Dokukina and Konilov, 2011; Dokukina et al., 2012a)), and inthe coastal outcrop to the southeast of the Gridino village (GridinCape, dyke D44 (Dokukina et al., 2014–this issue)). In additions, a gabbrointrusion, located on the Suprotivny island in the immediate vicinity oftheGridinodyke swarm, has also been studied (Slabunov et al., 2008). Be-sides, several felsic veins crossing the eclogitized dykes have also recentlybeen studied and dated (Dokukina et al., 2009, 2012a,b, 2014–thisissue). As there has been considerable debate in recent years as to

Fig. 11. Plot of U/Pb age vs. 176Hf/177, in which Hf shows relatively narrow horizontaldistribution of points that suggests resetting of U–Pb system, where zircons with youngerU–Pb ages have clearly been affected by non-zero Pb loss (~1.9 Ga). TDM — modal age.Modified after Mints et al. (2010a).



whether the eclogite-facies metamorphism in the Gridino dykeswarm is of Paleoproterozoic or Neoarchaean age, the discovery of apopulation of early-Palaeoproterozoic zircons in the metasomatic vein-lets (point 1111-08 in Dokukina et al., 2014–this issue) cutting anolivine metagabbro-norite dyke that had been described previously byVolodichev et al. (2009) and Slabunov et al. (2011) (dyke B-16) hasproven to be of particular interest.

Data relating to zircon populations in both the mafic eclogitizeddykes and the cross-cutting felsic veins (see Dokukina et al.,2014–this issue for more details) permit us to delineate the nextsequence of events.

(1) Subhedral oscillatory zoned zircons, consistent with igneous or-igin and magmatic crystallization (Fig. 12a, b), have been spo-radically recognized in both mafic dykes and crossing felsicveins and have yielded 206Pb/207Pb ages of 3.0–2.91 Ga. Theseages are greater than the maximum age suggested for theprotoliths of the Salma eclogites and most probably corre-sponds to the “embryos” or the earliest continental fragmentsfound in the Fennoscandian Shield (Mints et al., 2010b;Dokukina et al., 2014–this issue).

(2) Stubby prismatic to bipyramidal zircons dated at 2.87–2.82 Ga(Fig. 12c, d) have been found in some mafic dykes and in thegabbro intrusions in Suprotivny island (Slabunov et al., 2008).These zircons differ by their morphology, internal structureand age from the zircons in the country rock gneisses andgranite-gneisses (unpublished data of Belousova and Natapov,see also in (Mints et al., 2010b)) and might have crystallizedfrom mafic melt. Thus, these zircons, most probably, dategabbro intrusions in Suprotivny island and the Gridino dykeswarm.

(3) Elongate oval-shaped zircons mainly with large dark cores andlight or gray-colored rims in CL images dated at 2.86–2.83 Ga(Fig. 12e, f) and 2.82–2.78 Ga (Fig. 12g, h) occur in both theeclogitized mafic dykes and in crosscutting migmatite felsicveins. Geological relationships between the dykes and veins,together with the morphological and compositional similarityof the zircons are consistent with close interaction between in-trusion of mafic magma, generation of felsic melts and theirmutual crystallization and recrystallization under high-pressure(up to eclogite-facies) conditions. The existence of two zirconpopulations of the same type might indicate successive repeti-tion of the crystallization–recrystallization process. We concludethat these zircons correspond to the eclogite-facies event in themagmatic–metamorphic evolution of the Gridino dykes.

(4) The metagabbroic dyke at Cape Vargas contains characteristicroundish oval-shaped zircons similar to the “granulite-type” zirconsfrom the Salma eclogites. In both cases these zircons yielded thesame age of 2.72–2.70 Ga. Zircon crystals of similar age (2.72–2.71 Ga) from the enderbite vein crossing the metagabbro-norite dyke 1111 (on the northeastern outskirts of Gridinovillage, see Dokukina et al., 2014–this issue) are also oval-shaped,with wide gray zones in CL images (Fig. 12i). Petrological studyindicates that these zircons have a complexmagmatic andmeta-morphic history. In the metagabbro from the Suprotivny islandthe 2.74–2.73 Ga zircons are stubby prisms with only weaklydiscernible zoning. Finally, the phengite-bearing leucosome cut-ting metagabbro dyke at Cape Vargas contains subhedral zirconcrystals with a distinct core and rim structure. The cores displayrelict oscillatory zoning and some crystals have the “fir-tree”cores (Fig. 12j) typical of high-grade zircon (Corfu et al., 2003).A single cluster of near-concordant estimates yielded an upper-intercept age at 2.71 Ga (Dokukina et al., 2012a).

(5) Zircons extracted from the thin quartz veinlets that penetrate themetagabbro-norite dyke at the cape on the northeastern outskirtsof the Gridino village (the early-Palaeoproterozoic dyke B-16

Fig. 12. Specific zircon populations that are related to the successive events in evolu-tion of the Gridino eclogite association. Refer to explanation in the text.Based on data from Dokukina et al. (2014–this issue).



according to (Volodichev et al., 2009; Slabunov et al., 2011) be-long to a single zircon population with a concordia intercept at2.39 Ga. The zircons have a distinctive “lumpy” form withvoids representing the traces of the gas-liquid inclusions andno internal zonation is evident in CL images (Fig. 12k). Thesezircons contain very high concentrations of Th (0.2–1.7%), U(0.5–0.9%), Y (0.5–3.0%) and Hf (0.7–0.9%) and also contain nu-merous inclusions of orthopyroxene, clinopyroxene, Ti-rich bio-tite, rutile, quartz, and Cl-apatite, all of which are characteristicof symplectite assemblages and the minerals crystallized fromfluid (Dokukina et al., 2012a,b). It is important to note that sim-ilar zircons from the B-16 dyke have previously been interpretedas “typical magmatic” zircons that date intrusion of the maficmagma (Volodichev et al., 2009; Slabunov et al., 2011). We be-lieve that our results allow us to elaborate upon this under-standing; accordingly, we propose that these zircons, togetherwith the enclosing high-temperature zone were formed afterthe intrusion of the dyke and its transformation under eclogite-and later granulite-facies conditions (see Dokukina et al.,2014–this issue for more details).

(6) In the rocks reworked by later processes, where the eclogite-and granulite-facies mineral assemblages were obliterated, orwhere they are only present relic form, the Archaean magmaticand metamorphic zircon cores are usually surrounded by thelow-U colorless rims with ages of about 1.9 Ga. Dating of singlezircon grains from such rocks throughout the Belomorian Beltas a whole tend to yield the same age (e.g., Bibikova et al,2004; Slabunov et al, 2006; Mints et al., 2010b).

5. Discussion

5.1. The origin of the protoliths

Interpretation of geochemical data allows us to infer intraoceanicorigin for the mafic and ultramafic rocks of the Salma eclogite associ-ation with some degree of confidence. The “normal” mafic and Fe–Tieclogites and piclogites also share geochemical features that indicatethat their respective protoliths were indeed genetically related(Figs. 4, 5). Relatively high Nb concentrations are consistent with amantle-plume component having been involved in the petrogenesisof the protoliths. The combined geological, petrological, geochemicaland “zirconological” data show that protolith of the Salma eclogite as-semblage could have been an suite of predominantly plutonic mafic–ultramafic rocks rather like the third layer of the oceanic crust(“layered gabbro”) in the modern Southwestern Indian mid-oceanridge (Dick et al., 2000 and references therein). An oceanic crustalaffinity for the protoliths is confirmed by the abundant findings ofrelict mineral assemblages characteristic of early metamorphic to hy-drothermal reworking in the ocean floor environment (Mints et al.,2010b; Konilov et al., 2011). These geochemical similarities betweenthe Salma subduction-type eclogites and the Gridino eclogitizeddykes naturally leads to speculation concerning the possibility thatthe two types of eclogites formed within the interrelated geodynamicenvironments (Mints et al., 2010b; Dokukina and Konilov, 2011)(Figs. 4, 6, 7).

5.2. The evolution of magmatism and metamorphism

Integration of the petrological and geochronological data allows usto define the history of the Salma eclogites in the form of a loop-likeP–T–t path (Fig. 13). The proposed oceanic nature of the eclogiticprotholiths permits us to consider the prograde P–T–t evolution asdirectly reflecting the P–T–t trajectory of the subducting slab. Atpresent, the BEP is the first and the only natural example that permitsreconstruction of the transport of mafic crust down to the realm ofeclogite-facies conditions during Mesoarchean time. The prograde

branch of the P–T path, starting from the prehnite–pumpellyite facies,passes through a region of high temperatures, skirting the P–T rangeof the blueschist-facies metamorphism. The Archaean subductionrecorded by the Salma assemblage was significantly warmer thanthe known examples of the modern “warm” subduction (Carson etal., 1999; Peacock et al., 2002; Aoya et al., 2003; Page et al., 2003),and could be referred to as “hot subduction”. At a depth of 25 km,where estimates of 640–670 °C in the Salma eclogites are characteris-tic, the Cascadia “warm” subduction zone separating the Juan de Fucaand North America plates, has temperatures some 100–200° lower,in the range of 450–550 °C. At shallow depths within the garnet–amphibolite field, the Salma P–T path crosses the “basalt wet solidus”line. This suggests the possibility of early partial melting of thesubducting slab, before reaching eclogite-facies conditions.

The prograde metamorphic evolution of the Gridino mafic dykeswarm includes: (1) isobaric cooling of the magmatic rock down toamphibolite-facies conditions (~5 kbar at ~600 °C); (2) burial withcorresponding increase of pressure and temperature; and (3) eclogite-facies metamorphism with peak conditions of 17–18 kbar (probablyup to 22 kbar) and ~800 °C (Volodichev et al., 2004; Dokukina et al.,2014–this issue) (Fig. 13). The sequence of events relating to the retro-grade metamorphism included periods of cooling and decompression,punctuated by strong thermal pulses caused by the emplacement ofmantle plumes in the underlying mantle at 2.72–2.70, 2.5–2.4, and2.0–1.9 Ga. As we noted above, these events are characteristic of theEarly Precambrian history of the East European Craton as a whole.Within the BEP, they were accompanied by the corresponding temper-ature increases up to amphibolite- and granulite-facies conditions. Theintervening cooling period after 2.4 Ga, by some 150 °C has been in-ferred by the petrological observations (for more details see the paperof Dokukina et al., 2014–this issue). We assume that similar coolingepisodes preceded the plume impingement at 2.72–2.70 and 2.0–1.9 Ga, although we have not yet been able to identify or constrainthese petrologically. The proposed alternating periods of cooling-decompression and plume-related increase in temperature in the crustand underlying mantle are displayed in the P–T–t diagram (Fig. 13) inthe form of a zigzag P–T trajectory. The diagram also shows similar trajec-tory illustrating the retrograde evolution of the Salma eclogite association.

Evidences for high-grade metamorphism in the quartzo-feldspathicrocks (granitoids and gneisses) are very scarce. A similar situation iscommon in other eclogite-bearing complexes (e.g., Rubie, 1986; Ernstet al., 1998; Gilotti et al., 2004). However, eclogite-facies mineral assem-blages have been recently found in the granite leucosome crosscuttingthe metagabbro dyke at Cape Vargas. Granulite-facies mineral assem-blages in felsic rocks have, however, been recognized more widely, typ-icallymanifest as garnet-kyanite gneisses and enderbite veins (Dokukinaand Konilov, 2011; Dokukina et al., 2012a, 2014–this issue).

5.3. The relationship between the eclogite associations in Salma andGridino, and geodynamic environment and evolution of the Belomorianeclogite province

Eclogites of both the Salma and Gridino associations, are confined toand distributed throughout the migmatized TTG gneisses of the Keret'complex, which plunges northeastward beneath the Inari–Kolamicrocontinent and lies structurally above the mafic–ultramafic assem-blage of the Central-Belomorian suture zone (Slabunov et al., 2006;Mints et al., 2009, 2010b). This leads us to suggest that the Keret'eclogite-bearing TTG gneisses were originally derived from theMeso-Neoarchaean active margin of the Inari–Kola microcontinentand the Kola continent as a whole. A striking feature of the eclogites isthe general coincidence of the P–T–t paths for both associations(Fig. 13) (Mints et al., 2010b; Konilov et al., 2011; Dokukina andKonilov, 2011). As was shown above, the reconstructed protholits in-ferred for the Salma eclogitewere formed by intercalated gabbronorites,troctolits and Fe–Ti gabbros resembling the Layer Three of oceanic crust



formed at slow-spreading ridges. In turn, the magmatic source for theGridino mafic dykes could have been a subducting mid-ocean ridge.

A model of the overall evolution of the BEP integrating the fea-tures described below is presented in Fig. 14.

(1) The age of the oceanic-type protoliths of the Salma eclogiteshas been dated at ~2.9 Ga although the extent of oceaniclithosphere and relationship to other tectonic units remainsunknown.

(2) The precise timing and duration of the eclogite-facies meta-morphism of the Salma assemblage has not been directly esti-mated. The reconstructed thermal regime of the inferredsubduction zone is however consistent with subduction of aslow-spreading ridge (see below). On the other hand, the geo-chemistry of the eclogites from both the Salma and Gridino as-sociations allows us to interpret the formation of the dykeswarm from 2.87 to 2.82 Ga in terms of injections of themafic magmas from the ridge source into the overlying crustof the active. Two distinct groups of zircons, dated at 2.86–2.83 Ga and 2.82–2.78 Ga, respectively, can be distinguishedin both the eclogitized mafic dykes and the crosscuttingmigmatite felsic veins. When the geological relationships be-tween dykes and veins are considered as well, it is reasonableto deduce a close interaction between intrusion of mafic

magma, generation and segregation of felsic melts and theirrespective crystallization and recrystallization under high-pressure (up to eclogite-facies) conditions. We therefore con-sider it is probable that these zircons do relate directly to theeclogite-facies event in the overall magmatic and metamorphichistory of the Gridino dykes. Involvement of the lower crustin subduction between 2.87 and 2.82 Ga, which led to high-pressure metamorphism of Gridino dykes, may be explainedby the crustal delamination at the active margin and subse-quent subduction of the low-crustal slices, hosting maficdykes, together with subducting oceanic plate (see Fig. 14).The existence of two zircon population of similar type couldevidence the existence of multiple repeated episodes ofsubduction-driven crystallization–recrystallization processes.Given the similarity in ages of the oceanic-type protolith ofthe Salma eclogites, the emplacement of the Gridino maficdyke swarm, which we have attributed to ridge subduction,and the complex interaction between crystallization–recrystal-lization processes in both dykes and cross-cutting felsic veinsunder high-pressure conditions, it seems realistic to correlatethe time interval from ~2.87 to ~2.82 Ga with the main sub-duction event. It remains unclear however, whether subduc-tion was terminated this time. On the other hand, it isappropriate to associate the second event, dated by the 2.82–

Fig. 13. The P–T–t evolution of the Meso-Neoarchaean Salma and Gridino eclogite associations and comparison of the apparent P–T conditions along the Phanerozic subductionzones, P–T–t paths of some Phanerozic and Precambrian eclogite-bearing complexes. PP, prehnite–pumpellyite, EBS, epidote blueschist, LBS – lawsonite blueschist, GS, greenschist,EA, epidote amphibolite, A, amphibolite, PG, pyroxene granulite, and GG, garnet granulite. P–T–t paths of typical eclogite associations: (1) the Palaeoproterozoic associations(shown in black): Sharyzhalgai complex, SW Siberian craton (Ota et al., 2004); Usagaran orogen in Tanzania (Möller et al., 1995; Collins et al., 2004); Snowbird tectonic zone inthe North America (Baldwin et al., 2003, 2004); and eclogitized mafic dykes, Hengshan Complex of the North China Craton (Zhao et al., 2001; Kröner et al., 2006); (2) theMeso-Neoproterozoic associations (shown in green): Llano Uplift of central Texas (Carlson et al., 2007; Mosher et al., 2008); G-All, Grenville Province, the Allochthonous Belt(Indares and Rivers, 1995; Indares, 1997; Indares et al., 1998, 2000;Wodicka et al., 2000; Rivers, 2009); SN, Sveconorwegian orogen (Möller, 1998; Möller et al., 2007); and Zambezibelt in NW Zimbabwe (Dirks and Sithole, 1999); and (3) the Phanerozoic associations (shown in blue): New Caledonia (Carson et al., 1999; Spandler and Rubatto, 2005); BR, BlueRidge Mountains (Southern Appalachians, North America) (Page et al., 2003); and Moldanubian zone of the Bohemian massif, European Variscan Belt (Faryad, 2011).P–T–t paths of the Salma and Gridino eclogite associations are summarized from Kaulina et al. (2007, 2010), Dokukina et al. (2009, 2012a,b), Mints et al. (2010a,b,c), Kaulina (2010)and Dokukina and Konilov (2011). Calculated P–T paths and metamorphic conditions encountered by oceanic crust subducted beneath Cascadia, southern Mexico, southwest Japanand northeast Japan after Peacock et al. (2002). Metamorphic facies after Peacock et al. (1994).



2.78 Ga zircon population, with the formation of the Parandovo–Tiksheozero island arc and the final collision that resulted in theamalgamation of the Karelia, Kola and Khetolamba continentalblocks and the formation of the Meso-Neoarchaean Belomorianaccretionary-collisional orogen (Mints et al., 2010b). The defor-mation associated with collision may have initiated a furtherhigh-pressure metamorphic event.

(3) The Neoarchaean–Palaeoproterozoic history of the BEP includesa series of events marked by crystallization and recrystallizationof the specific zircon populations at 2.72–2.70, 2.39 and ~1.9 Gaand single grains of 2.3–2.2 Ga age. For a comprehensive under-standing of significance of these dates we need to consider theirglobal and regional correlations.

It has became widely accepted that zircon U–Pb age peaks ofglobal significance at 3.1–2.9, 2.8–2.7, 2.6–2.4, 2.2–2.1, 2.0–1.7,1.2–1.0 Ga, most of which also correlate with Lu–Hf, Sm–Ndand Re–Os depleted-mantle model ages, reflect juvenile crustproduction and related rapid crustal growth (e.g., Condie,1998; Abbott and Isley, 2002; Hawkesworth and Kemp, 2006;Voice et al., 2011; Bradley, 2011, and references therein). Themost prominent peak at 2.8–2.7 Ga suggests a sharp andsudden turning-point in crust-forming processes in the earlyNeoarchaean. The two most dominant superplume eras in theProterozoic occurred at approximately 2.0–1.7 and 2.6–2.4 Gaago (Abbott and Isley, 2002). These age peaks are usuallythought to be associated with subduction of oceanic plates and

Fig. 14. A schematic model for the magmatic and metamorphic evolution of the Belomorian eclogite province.



suprasubductionmagmatic processes (Condie et al., 2009; Lee etal., 2011). However, it is difficult to conceive how global pulsesof crustal growth might occur in response to conventionalWilson-style plate tectonics, particularly because there are noage peaks in the past billion years, from which we might sur-mise that subduction-related magmatismwas the globally dom-inant process in the generation of new crust during certainperiods of time. An alternative explanation is that periodicityin igneous activity reflects rapid crust formation during thermalpulses associated with the emplacement of mantle plumes(Hawkesworth and Kemp, 2006).

(3.1) Both eclogite-bearing associations were somehow emplacedat higher crustal levels, within the granulite-facies P–T pa-rameters regime by 2.72–2.70 Ga. The time interval ofmore than 60 Myr between the second pulse of the subduc-tion and the granulite-facies metamorphism leads us to con-clude that these two events were independent of oneanother. At 2.7 Ga the newly formed eastern FennoscandianShield records extensive sedimentation, magmatism andhigh-temperature metamorphism, which we regard as afundamentally new stage of the crustal development.These processes have been concentrated in the oval-shapedKarelia–Belomorian and Kola provinces, which reflects the in-fluence of separate local mantle plumes (Mints, in press). TheBEP occurred in the outer zone of the Karelia–Belomorianarea, characterized by sporadic manifestation of granulite-faciesmetamorphism (Mints et al., 2010b, 2011). We proposethat formation of the granulite-facies mineral assemblagesand the growth of specific zircon populations were linkeddirectly to the above processes.

(3.2) Similarly, the ~2.4 Ga thermal event accompanied by infiltra-tion of metamorphic fluids and hydrothermal alteration,with corresponding growth of the high-U, high-Th zirconsshould be related to the early-Palaeoproterozoic impulseof the mantle plume impingement. This event between 2.5and 2.4 Ga, represents a Large Igneous Province, withgabbro-anorthosites, layered mafic–ultramafic complexes,charnockites and bimodal volcanism, throughout the easternFennoscandian Shield (Sharkov, 2006). This event alsoresulted in origin of intracontinental depressions, with rapidsedimentation and granulite-facies metamorphism of the Ar-chean basement rocks together with the Palaeoproterozoicdepression-fill sequences (Mints, in press).

(3.3) Between 2.1 and 1.9 Ga, a relative lull was followed by areactivation of the mantle plume or, might be, impingementby a fresh plume. Sedimentary basins fill and their basementagain underwent granulite-facies metamorphism. The 1.93–1.86 Ga compression led to the formation synformal thrustand nappe ensembles of the granulite-gneiss belts. In therocks of the Belomorian province that were reworked duringthis time, the Archaean magmatic and metamorphic zirconssometimes rimmed by low-U colorless overgrowths withages of 1.9 Ga (Fig. 10e, f). Dating of such zircons from therocks within the Belomorian Belt as a whole yields the sameage (e.g., Bibikova et al, 2004; Slabunov et al, 2006; Mints etal., 2010b), while titanite ages from TTG gneisses are~1.87 Ga (Bibikova et al., 2001), indicating equilibration at atemperature of about 700 °C at this time.The Sm–Nd mineral — whole rock isochron dates obtainedfrom the metagabbro-norite and metagabbro dykes (1.92–1.91 Ga) and from the Salma eclogites (1.89–1.87 Ga, withone exception at 1.79 Ga) essentially record the same age.Hence the late-Palaeoproterozoic metamorphism led to isoto-pic reequilibration of the Sm–Nd system in minerals. The Lu–Hf isochrones for the whole rock, garnets and clinopyroxenes(WR–Grt–Cpx) from the Salma-type eclogites (Kuru-Vaara

quarry and islands in the vicinity of Gridino village) gaveage estimates within about the same range from 1.93 to1.89 Ga (Skublov et al., 2011c). In contrast, the Sm–Nd sys-tem in rocks proved to be quite robust by the age of theSalma eclogites (~3.0 Ga) and of the Gridino mafic dikes(3.4–3.0 Ga) (Mints et al., 2010b; Kaulina, 2010; Skublov etal., 2010b; Dokukina et al., 2010, 2012a). The closing temper-ature of the Sm–Nd system in garnet was estimated at 650–670 °C (Ganguly and Turone, 2001), which corresponds toamphibolite facies conditions. Because little is known aboutdiffusion of Hf in common minerals, it is not clear whetherLu–Hf ages of eclogitic garnets reflect a closure temperaturerather than the timing of crystallization (Duchêne et al.,1997). It is known that a Lu–Hf closure temperature is greaterthan or equal to that of Sm–Nd in the same garnet. The actualTC values for Lu–Hf in garnet appear to depend principally onthe size of the garnets, ranging from ~540 °C (0.24 mm radi-us) to more than 700 °C (Scherer et al., 2000). The Lu–Hf andSm–Nd agesmay record cooling ages, but not prograde garnetgrowth or peak metamorphic conditions (e.g., Kylander-Clarket al., 2007).Data on mineral isotope geochemistry in the eclogites havebeen used to estimate their cooling history (Mints etal., 2010b; Kaulina, 2010; Dokukina and Konilov, 2011;Dokukina et al., 2012a). The cooling of the rocks to tempera-tures below 550–500 °C was recorded by the 40Ar/39Ar meth-od on amphibole (Harrison, 1981). The dates obtained fromrutile using U–Pb method, 1.79–1.76 Ga are much youngerand reflect the cooling time of themetamorphic rocks to tem-peratures of 400–450 °C (Mezger et al., 1991). These valuesare confirmed by 40Ar/39Ar dating of muscovite and biotitein the leucosome, which record an age of 1.79–1.69 Ga(Dokukina et al., 2010, 2012a) for cooling of the systemdown to 425–280 °C (Harrison et al., 1985, 2009).

(3.4) The above data indicate that the BEP eclogite associationswereexhumed to mid-to-lower crustal depths by approximately1.77 Ga, after which erosion or younger tectonic events wereresponsible for the final exhumation to the surface.

5.4. Discussion on existing interpretations of the availablegeochronological data

There has been considerable discussion in recent years on the in-terpretation of the geochronological data from the BEP. Volodichev,Slabunov and Shchipansky with co-authors argued that two Archaean(~2.82 and ~2.72 Ga) and two or more Palaeoproterozoic eclogite-facies events are recorded in the rocks of the BEP. The arguments ofthese researchers have been published earlier (Volodichev et al.,2004, 2005; Slabunov, 2008; Slabunov et al., 2010; Shchipansky etal., 2012b).

A fundamentally different model was recently presented in a se-ries of papers by Skublov with co-authors (Skublov et al., 2010a,b,2011a,b,c). These researchers have made persistent attempts to“issue a final decision” on the late-Palaeoproterozoic “Svecofennian”(~1.9 Ga) age of eclogite-facies metamorphism in the BEP. Estimatesof U–Pb zircon ages obtained by these authors, are in agreement withthe results that were previously obtained by other researchers. How-ever, Skublov and co-workers have interpreted the 2.88 and 2.70 Gaages exclusively in terms of magmatic events. In contrast, the age es-timates in the range of 1.91–1.84 Ga obtained from the outer zirconrims, with low concentrations of trace elements (“typical geochemi-cal characteristics of zircons from eclogite” (Skublov et al., 2011b,c))were interpreted as corresponding to a late-Palaeoproterozoic agefor the high-pressure metamorphism throughout the entire BEP. In jus-tifying their model these authors ignored petrological and geologicaldata relating to the overall context, focusing solely on the geochemical



characteristics of the dated zircons. The similarities between thelate-Palaeoproterozoic overgrowths and replacements of the earliergenerations of zircons and “typical eclogite” zircons, as defined on thebasis of published data, has been adopted by these authors as an essen-tial and sufficient criterion for advocating a late-Palaeoproterozoic agefor eclogite-facies metamorphism in the BEP. However, individual zir-con grains as well as rims surrounding older zircon cores, sharing simi-lar geochemistry and morphology and of the same age are widelydistributed in various rock types throughout the Belomorian province,including rocks that have never undergone eclogite-facies metamor-phism (e.g., Bibikova et al., 2004; Serebryakov et al., 2007; Kaulina,2010). We also note that in the paper of Skublov et al. (2009), whichwas devoted to the dating of the Shueretskoe garnet deposit associatedwith late Palaeoproterozoic pegmatite bodies, the authors appear tomake a conclusion that is contrary to the proposed unique and diagnos-tic attributes of “eclogite” zircons: “Zircons formed coincidentally withgarnet are similar to so-called eclogitic zircons in their geochemicalcharacteristics (low concentrations of HREEs and LREEs, Th, and otherminor elements), even though the pressure during the formation ofthe garnet deposit was significantly lower than the parameters ofeclogite facies. … Metasomatism resulting in the formation of garnetdeposit occurred under the conditions of metamorphism peak at a tem-perature of 650–680 °C and a pressure of 7.8–8.5 kbar” (Skublov et al.,2009, p. 1547). The methodology used by Skublov and co-authorsmay well be applicable in reconstructing a relatively simple metamor-phic history (e.g., Herwartz et al., 2008), but not when attempting toidentify multiple metamorphic events under amphibolite and granulitefacies conditions. Such an approach to the study of the BEP rocks withtheir complex polymetamorphic history inevitably led to an erroneousinterpretation of the crystallization conditions for the late overgrowthssurrounding zircon crystals and untenable conclusions about themeta-morphic evolution and the age of eclogitic associations.

5.5. The role of the Belomorian eclogite province in the evolution of theEarly Precambrian crust in eastern Fennoscandian Shield

Discerning the nature and precise location of boundary between theArchaean Kola and Belomorian provinces in the eastern FennoscandianShield has been difficult to establish. For many years it was assumedthat this boundary was defined by the Palaeoproterozoic structures(e.g., Kratz et al., 1978; Mints et al., 1996; Glebovitsky, 2005;Balagansky et al., 2006; Daly et al., 2006). However, the recognition ofthe BEP and the results of investigations nowmakes it possible to delin-eate and characterize this Archaean collisional boundary (Fig. 1).

5.6. Specifics of the subduction processes and conditions of the eclogite-faciesmetamorphism in the Precambrian

As shown above, it is possible to correlate the eclogite-facies meta-morphism within the Meso-Neoarchaean BEP with “hot” subduction.However, was such a thermal regime specific to the Archaean? It canbe concluded from the review below that all known Proterozoiceclogites also belong to this high-T type. Analysis of the known occur-rences of eclogite metamorphism in the Precambrian (excludingdeep-seated xenoliths from kimberlite pipes from consideration)permits the recognition of three types of the high-T eclogites thatformed: (1) due to subduction of oceanic lithosphere (Salma, Usagaranorogen, Llano Uplift, NWHighlands of Scotland), (2) in connectionwiththe collision of continental blocks in which mafic dyke swarms intrudedTTG type assemblages (Gridino,mafic dykeswithinHengshan Complex),and (3) at the base of thick crustal sections subjected to granulite-faciesmetamorphism (Snowbird tectonic zone, Grenville–Sveconorwegianorogen).

The mid-Palaeoproterozoic crustal eclogites are known within theUsagaran orogen in Tanzania (Möller et al., 1995), located immediatelysoutheast of the Archean Tanzania Craton (Fritz et al., 2005). The

tectono-stratigraphic section of the Usagaran orogen includes theEclogite zone, which continues for about 35 km, consisting mainly ofArchaean (~2.7 Ga) gneisses, containing lense-shaped bodies of maficeclogites (locally also kyanite-bearing). The peak eclogite-facies condi-tions of 750 °C and 18 kbar were attained at 2.00 Ga. The retrogradenearly isothermal P–T–t path intersected the granulite-facies field(Fig. 13). The formation of symplectites and further recrystallizationunder amphibolite-facies conditions terminated not later than 1.99 Ga(Collins et al., 2004). Limited geochemical data indicate MORB-typebasaltic protoliths to the eclogites. The overall geological context is con-sistent with consideration of the Usagaran orogen as representingreunification of a craton, that had been disrupted by the formation ofa short-lived ocean basin ~2.00–1.99 Ga ago (Reddy et al., 2003). For-mation of the eclogite grade assemblage is attributed to deep burial ofoceanic crust.

Eclogites of late Palaeoproterozoic age (~1.9 Ga) are known fromthe Snowbird tectonic zone in the North America, where they wereemplaced at the base of the tectono-stratigraphic crustal section ofthe East Athabasca mylonite triangle (Baldwin et al., 2004), that is afragment of oval synform, which was called “Athabasca lozenge”(Hanmer et al., 1995). Formerly, these high-T and high-P myloniticquartzo-feldspathic gneisses and gneiss-hosted metabasites had beenconsidered to be Archaean ones (Snoeyenbos et al., 1995). Thecoarse-grained garnet-omphacite eclogite contains small amounts ofkyanite, pargasite, orthopyroxene and spinel. The metamorphic P–T es-timates of 920–1000 °C at a minimum pressure of 18–20 kbar exceedthe values of corresponding parameters in associated mafic granulites,which are 890–960 °C at and 13–19 kbar respectively (Fig. 13). Thesemetamorphic events were considered to occur in an intracratonic set-ting, so they do not define a paleosuture (Baldwin et al., 2003, 2004).

With respect to our study, the characteristics of late Palaeoproterozoicmafic dykes in the northern part of the Hengshan Complex of the NorthChina Craton are of particular interest. The North China Craton had along history from ~3.8 Ga prior to 1.8 Ga final high-grade metamorphicevent that is to be related to collision of the northernmargin of the cratonwith the Columbia supercontinent (Kusky, 2011 and references therein;Zhai and Santosh, 2011). The late Palaeoproterozoic mafic dyke swarmsare widespread in the North China Craton. Their geochemical features in-dicate a continental rift environment (Kusky and Santosh, 2009). Maficdykesweremetamorphosed successively under eclogite, high-P granuliteand amphibolite facies conditions during orogenic event. Magmatic andmetamorphic zircons have been dated from ductile deformed gabbroicdykes. These dykes occur as boudins and deformed sheets within TTGgneisses and contain the relict high-pressure granulite- or even formereclogite-facies assemblages. Within a single mafic body, a variety of mi-crostructures can be found ranging from undeformedmagmatic texturesstatically overprintedby granulite facies or locally amphibolite faciesmin-eral assemblages, to completely recrystallized amphibolite facies domainsentirely devoid of relictmagmatic textures (Kröner et al., 2006). Petrolog-ical studies have defined a near-isothermal decompressional clockwiseP–Tpath suggesting that theHengshanComplex underwent initial crustalthickening, subsequent exhumation, cooling and retrogression. The earli-est recognized metamorphic mineral assemblage (M1) is preserved onlyin the form of quartz and rutile inclusions within garnet porphyroblasts,and omphacite pseudomorphs represented by clinopyroxene–plagioclasesymplectites. High pressure (HP) and ultrahigh-temperature (UHT)metamorphic assemblages have been described from granulites in sever-al areaswithin the Khondalite belt in the northern part of theNorth ChinaCraton, with temperatures exceeding 1000 °C and pressures above12 kbar (Kusky and Santosh, 2009). Metamorphic eventM2 correspondsto high-pressure granulite-facies conditions of 770–840 °C and 13.4–15.5 kbar, while the next retrograde event M3 was characterized bylower grade conditions (750–830 °C and ~6.5–8.0 kbar) and the last rec-ognized M4 event corresponds to high-T amphibolite facies conditions(680–790 °C and 4.5–6.0 kbar) (Zhao et al., 2001) (Fig. 13). Ages formagmatic zircons are ~1.92 Ga, whereas metamorphic zircons are in



the range 1.89–1.85 Ga. These results show that crustal extension accom-panied bymaficmagma injections occurred in the late Palaeoproterozoic.This preceded a major high-pressure collision-type metamorphicevent in the central part of the North China Craton (Kröner et al.,2006 and references therein). On the other hand, the highly meta-morphosed extensive mafic dyke swarms might be interprerted asignal of a major anomalous thermal event correlated with astheno-spheric upwelling and emplacement of mafic magmas in response tothe breakup of the supercontinent Columbia, which provided the dryheat input to achieve generation of granulite facies assemblages(Kusky and Santosh, 2009).

The best known Meso-Neoproterozoic eclogites are within theGrenville–Sveconorwegian orogen (Wardle et al., 1986; Anderssonet al., 2008). There are two types of eclogites in these units:(1) subduction-related types and (2) high-T types, which are closelyassociated with granulites.

Subduction-related eclogites (Type 1) were discovered and inves-tigated at flanks of the orogen: in the Llano Uplift of central Texas andin the NW Highlands of Scotland. Mezoproterozoic eclogites of theLlano Uplift are in the extreme southwestern part of the GrenvilleOrogen (Carlson et al., 2007; Mosher et al., 2008 and references there-in). As in many collisional orogens, evidence for the highest-pressureportion of the Llano metamorphic history is restricted to boudins ofmafic eclogite encased in felsic gneisses. The host rock assemblagedated within the age range 1.33–1.23 Ga represent a mixture ofmetamorphosed terrigenous clastic, volcanic, and plutonic rocks,dominated by quartzo-feldspathic gneisses but containing associatedsemipelitic schist, amphibolite and marble. A large body of ser-pentinized harzburgite has been interpreted as an ophiolite fragment.Metamorphism comprised both an initial high-pressure phase (610–775 °C at 14–24 kbar) and a subsequent medium-pressure phase(~700 °C at ~7 kbar) (Fig. 13). The age of HP/MP metamorphism fallswithin the time span from ~1.15 to ~1.12 Ga.

Eclogites in the NW Highlands of Scotland are known within theGlenelg–Attadale basement inlier within the Caledonian Orogen(Sanders et al., 1984; Sanders, 1988). In the eastern part of the inlier,the TTG gneisses are associated with multiple inclusions of kyanite-and omphacite-bearing gneisses, mafic eclogites and manganiferousmetasediments. The source of the mafic protoliths was depletedmantle, suggesting an oceanic origin of the protoliths (unpublisheddata of Storey by Brewer et al., 2003). Peak conditions were estimatedas ~20 kbar and 750–780 °C. The eclogites subsequently followeda steep decompression path to ~13 kbar and 650–700 °C duringamphibolite-facies retrogression. Peak eclogite-facies metamorphismoccurred at ~1.08 Ga and retrogression at ~1.00 Ga (Storey et al.,2005 and references therein). The western unit also consists of TTGgneisses and granite–gneisses enclosing several Palaeoproterozoic(~1.75 Ga) mafic bodies, which have locally preserved eclogite-facies mineral assemblages. Mineral reactions also indicate thepresence of transient granulite-facies conditions during retrogression(Storey et al., 2010).

Related manifestations of Meso-Neoproterozoic high-pressureeclogite- and granulite-facies metamorphism (Type 2) were studiedin detail within both the Grenville and Sveconorwegian sectors ofthe orogen. Within the Grenville sector the high-pressure metamor-phic rocks are always closely associated with the base of the Alloch-thonous Belt (Indares and Rivers, 1995; Indares, 1997; Indares et al.,1998, 2000; Wodicka et al., 2000; Rivers, 2009). Eclogite-facies meta-morphism with peak P–T parameters of 850–920 °C and 12–14 kbarwere dated at 1.06–1.02 Ga. Eclogite-facies metamorphism has beenalso recognized more locally within the Parautochthonous Belt. TheSveconorwegian orogeny has been interpreted in terms of polyphaseimbrication of crust at the margin of Fennoscandia between 1.13 and0.96 Ga. Eastward thrusting resulted in successive metamorphicpeaks in different terranes and a corresponding eastward youngingof metamorphism from 1.13 to 1.10 Ga in the Bamble–Kongsberg

sector (e.g., Cosca et al., 1998) to 1.05–1.03 Ga in the Idefjorden ter-rane (Söderlund et al., 2008 and references therein) and finally0.98–0.96 Ga in the parautochthonous Eastern Segment (e.g. Bingenet al., 2005; Möller et al., 2007; Söderlund et al., 2008). Relict eclogitesand associated high-pressure rocks are also present in the EasternSegment of the SW Swedish gneiss region (the tectonic counterpartof the Parautochthonous Belt of the Canadian Grenville). The best-preserved eclogite relics suggest a clockwise P–T–t history, beginningin the amphibolite facies, progressing through the eclogite facies,decompressing and partially reequilibrating through the high- andmedium-pressure granulite facies, before cooling through the am-phibolite facies. Pressures of 9.5–12 kbar and temperatures of 705–795 °C for a stage of the granulite facies decompression have been es-timated (Möller, 1998, 1999 and references therein).

The eclogitic rocks within the Makuti Group gneisses of theZambezi Belt in north-western Zimbabwe form one more exampleof the complicated tectonic and metamorphic history. Within low-strain domains in the Makuti gneisses, undeformed metagabbroiclenses preserve eclogite and granulite facies assemblages, which re-cord a part of the metamorphic history that predates Pan-Africanevents. The eclogites preserve multi-staged symplectic reaction tex-tures that developed progressively as the rocks experienced loadingfollowed by decompression and heating. The decompression-heating textures reflect temperature increases during exhumation ofthe Makuti gneisses. The peak metamorphic conditions might bevery roughly estimated at ~20 kbar and ~720 °C. The peak assem-blage was overprinted by the multi-mineral corona assemblages, pos-sibly at ~19 kbar 760 °C, then at 15 kbar and 750 °C and still later at10 kbar and 760 °C. The eclogite facies rocks formed prior to 0.85 Gaand they may be related to a suture zone that developed along theaxis of the Zambezi Belt. The main deformation and metamorphismin the Makuti gneisses occurred around 0.80 Ga and involved exten-sion and exhumation of the high-P rocks (accompanying break-upof Rodinia), which experienced a high-T metamorphic overprint(Dirks and Sithole, 1999) (Fig. 13).

Phanerozoic eclogite complexes, in contrast to the Precambrianones, usually formed in connection with “warm” subduction zones.The prograde part of the P–T evolutionary path of these complexespasses through the blueschist-facies fields. This feature is equallycharacteristic for both HP and UHP complexes. A few examples aregiven below.

A surprisingly complete reconstruction of the P–T evolution of theUpper Ordovician (460 Ma) eclogites has been made from the easternBlue Ridge Mountains (Southern Appalachians, North America). Al-though prograde metamorphism of eclogites is typically obscured bychemical equilibration at peak conditions and by partial reequilibrationduring retrograde metamorphism, the Eastern Blue Ridge eclogites inNorth Carolina retain evidence of their prograde path in the form of in-clusions preserved in garnet (Page et al., 2003). The combined P–T datashow a clockwise loop from the amphibolite to eclogite to granulitefacies, all of which are overprinted by a texturally late amphibolitefacies assemblage. The biotite–hornblende–garnet–quartz assemblageis within the range of possible peak pressures and temperatures(13–16 kbar, 630–660 °C) for the eclogite assemblage. The 0.46 Gaage of this assemblage closely resembles to other Taconic metamorphicages in the Eastern Blue Ridge, including the 0.46 Ga highest-grade granulite-facies metasediments at Winding Stair Gap. Thegarnet-rich leucosomes crossing the granulite-facies metasedimentsare unfoliated, crosscut the dominant regional fabric in sillimanitegneisses, and contain kyanite as a late or subsolidus phase. Mineralogicthermobarometry, oxygen isotope thermometry, and experimentalstudies constrain melting to have occurred at significantly elevatedtemperature ~850 °C at 8 kbar (Moecher et al., 2004). The shallow“hot” prograde path and granulite facies overprint of these eclogites(Fig. 13), coupled with their formation during the peak of Taconicmetamorphism in the southern Appalachians, suggests that these



rocks are the product of deep underthrusting during continental colli-sion (Page et al., 2003).

The Variscan Belt in the Western Europe contains further sig-nificant examples of Phanerozoic eclogites, including Devonian–Carboniferous high-/ultrahigh-pressure (HP/UHP) metamorphic rocks,extending from the IberianMassif through the ArmoricanMassif, FrenchMassif Central, Schwarzwald, and Vosges, as far as the Bohemian Massif(Keppie et al., 2010). Eclogites are common in all massifs, but blueschistsand blueschist-facies rocks are exposed only in the Iberian, Armorican,and Bohemian Massifs. In most cases, eclogites were derived frommid-ocean ridge basalts (MORBs). The HP/UHPmetamorphic rocks usu-ally occur within nappe stacks overlying amphibolite-to-greenschist-facies units, or form lenses and boudins in amphibolite-facies rocks. Inthe Moldanubian Zone of the Bohemian Massif, the HP felsic granulites,garnet peridotites, and HT eclogites are the part of the Gföhl nappe,which overlies the amphibolite-facies assemblages. The presence ofeclogites suggests that the Moldanubian Zone was, during the Variscanorogeny, part of an accretionary wedge into which the HP/UHP weretectonically emplaced. The common occurrences of felsic granuliteswith lenses of garnet peridotites, garnet pyroxenites, and eclogites, andthe similarities of their metamorphic ages (365–350 Ma), suggest acommon metamorphic history, at least during their exhumation(Bröcker et al., 2009; Faryad, 2011 and references therein). Felsic- andmafic-granulites of the Saxon Granulite Massif have undergone aregional ultrahigh-temperature (UHT) metamorphic event under high-pressure conditions (~1000 °C at P~23 kbar) (Rötzler et al., 2004). Sev-eral attempts have been made using numerical simulations to establishand understand the causes and mechanisms by which such UHT condi-tions can develop in continent–continent collision zones. However,most of themodels failed to simulate the attainment of themetamorphictemperatures as high as 700–900 °C. Another way to supply sufficientheat for UHT metamorphism in collision zones could be a mantleplume (Arnold et al., 2001). The finding of microdiamond from thegarnet-bearing peridotite in the Bohemian Massif provoked the ideathat this peridotite was originally derived from the hot asthenosphereas mantle diapirs, which advected heat to relatively shallow depth(Medaris et al., 2003). Such a heat supply from the orogenic garnet peri-dotites could have played an important role to create the UHTmetamor-phic condition of the Variscan orogeny (Carswell and O'Brien, 1993;Medaris et al., 2003; Naemura et al., 2011 and references therein).

The Saramta massif in the Palaeoproterozoic Sharyzhalgai complex(SW Siberian craton) may be an ancient analog of the Variscan Belt inthe Western Europe. Mainly composed of spinel-peridotites withgarnet-websterites, it is enclosedwithin granite-gneisses andmigmatiteswith mafic intercalations of granulite-facies grade. The ultramafic rocksunderwent four-stage metamorphism: stage 1 (pre-peak), 9–15 kbar at640–780 °C; stage 2 (peak), 23–30 kbar at 920–1030 °C; and stage 3(post-peak), 5–9 kbar at 750–830 °C at. Finally, the garnet-websteritesshow veining by lower amphibolite- to greenschist-facies minerals(stage 4) (Fig. 13). Thus, the Saramta massif was carried to depths ofabout 100 km by subduction, and metamorphosed under eclogite-facies conditions in the Palaeoproterozoic, despite the commonly heldview that high geothermal gradients in those times would haveprevented such deep subduction (Ota et al., 2004). However, it must beacknowledged that this metamorphic complex has not been preciselydated.

The Eocene (44 Ma) high-P, medium-T Pouébo terrane of the PamPeninsula, northern New Caledonia is a typical example, includingbarroisite- and glaucophane-bearing eclogite and variably rehydratedequivalents. The eclogites experienced a clockwise P–T path thatreached P~19 kbar and T~600 °C (Fig. 13) (Carson et al., 1999;Spandler and Rubatto, 2005).

The best preserved Triassic high- and ultrahigh-pressure (HP andUHP) metamorphic terrane in the Qinling–Dabieshan–Sulu orogen,western Dabieshan, contains five eclogite-bearing zones. The HP/UHPmetamorphism took place at about 220–230 Ma, as observed in the

Dabie and Sulu terranes (Jahn et al., 2005). The peakmetamorphic con-ditions range from ~500° to 700 °C and from 14 to 17 for Zone I to 26–29 kbar for Zone III. It was inferred that the LT/HP, HP and UHP sliceswere broken up from the downgoing slab during subduction andreached different depths along different geothermal gradients. The pro-grade P–T–t path crossed greenschist–epidote amphibolite–blueschist–amphibole eclogite and zoisite eclogite facies; the almost isothermalretrograde path crossed high-pressure granulite and amphibolite faciesfields (Liu et al., 2004; Jahn et al., 2005 and references therein). Onemore example of the Permo-Triassic high-pressure rocks is knownfrom the northwestern Korean Peninsula, where HP eclogites occur asrelicts inmafic granulites in theHongsung area of thewestern Gyeonggimassif (Sajeev et al., 2010).

Integration of the above review with the BEP main features allowus to make some general inferences about the nature of eclogite-facies metamorphism in the Precambrian. A comparison of the set ofP–T–t paths and data on the peak P–T parameters of the above listedmetamorphic events demonstrates their similarity (Fig. 13). Thehigh-T origin of these eclogites has a clear and simple explanationin the case of high-T eclogites at the base of granulite complexes, inthe interpretation of which additional sources of heat are invoked, ei-ther due to crustal thickening during subduction and collision(Percival, 1994; Brown, 2007, 2009) or by the involvement of mantleplumes (Sandiford, 1989; Harley, 1989; Mints, 2007; Mints et al.,2007). The “non-subduction” and “subduction” type eclogites ofProterozoic age are more or less clearly linked with intracontinent en-vironments, in which a subsidiary or major role for mantle plumes isimplicated. The high-temperature character of Proterozoic eclogite-facies metamorphism is impossible to explain by invoking an anoma-lous thermal regime for the mantle. However, the P–T parameters ofthe eclogite metamorphism in Proterozoic and Meso-Neoarhaeancomplexes are remarkably similar to one another. Thus, the thermalregime of the Archaean eclogite-facies metamorphism representedby the P–T parameters of the Salma and Gridino eclogites cannot beconsidered as confirming the concept of a “warmer” Archaeanmantle. Specific features of the thermal regime in subduction zonesdepending on the rate of approach of a spreading ridge werehighlighted by the results of 2D numerical modeling results byUehara and Aoya (2005). In the absence of elevated mantle flowand shear heating, the ratio of the ridge approach rate, u, to subduc-tion rate, v, was found to be an important parameter for the thermalstructure of the lithosphere just before ridge subduction. The petro-logically derived P–T path for the Salma eclogite assemblage closelyresembles the “hottest” trajectory that was calculated for the condi-tions of the lowest u/v ratio (0.1) and corresponding lowest spreadingrate (0.5–1 cm/year). Bearing in mind the above-mentioned similari-ty of the suggested Salma protholith assemblage with the gabbroicsuite from Layer 3 of the modern oceanic crust of the slow-spreading Southwest Indian Ridge, we believe that the explanationof the hot-type subduction in the Salma case, which suggests involve-ment of the slow-spreading ridge subduction, seems to be quiteplausible.

6. Conclusion

In this paper, the main features of the BEP have been reviewedfrom various points of view: regional and local geology, geochemis-try, petrology and geochronology and by comparisons with datafrom other Precambrian and Phanerozoic eclogites known at the time.

1. Eclogite-facies mafic rocks occur within gray gneisses of TTGaffinity in the northeastern part of the Belomorian Province, KolaPeninsula. These are characterized by widespread omphacite-breakdown textures and locally preserved relics of omphacite.The province contains two types of eclogites: (1) the subduction-



type Salma association, and (2) eclogitized mafic dykes of theGridino association.

2. Protoliths of the Salma eclogites can be inferred as a sequence ofalternating normal gabbro, Fe–Ti gabbro and troctolites, formedabout 2.9 Ga in slow-spreading ridge setting (similar to the mod-ern Southwest Indian Ridge). The main subduction event andeclogite-facies metamorphism of the Salma association occurredwithin the time interval from ~2.87 to ~2.82 Ga.

3. Mafic magma injections into the deep crust of the active marginand formation of the Grigino dyke swarm were associated withthe arrival of the mid-ocean ridge at the subduction zone, startingfrom about 2.87 Ga. Crustal delamination at the active margin andsubsequent involvement of the lower crust in subduction between2.87 and 2.82 Ga led to high-pressure metamorphism of Gridinodykes, culminating eclogite-facies conditions during the proposedcollision event between 2.82 and 2.78 Ga. This collision resultedin amalgamation of the Karelia, Kola and Khetolamba blocksand the formation of the Mesoarchaean Belomorian accretionary–collisional orogen.

4. Thermobarometry indicates a clockwise P–T path for both theSalma and Gridino associations with both crossing the granulite-facies P–T field. Detailed studies of the succession of magmaticand metamorphic processes indicate the complicated post-eclogite history, with mantle plumes invoked to explain thermalreactivation and hydrothermal processes at 2.72–2.70 Ga, ~2.4 Gaand ~1.9 Ga. The eclogite assemblages of the Belomorian provincewere exhumed to mid-to-lower crustal depths at about 1.7 Gawhere erosion or younger tectonic events were responsible forfinal exhumation to the present erosion surface.

5. Comparison of the P–T–t paths and data on the peak P–T parame-ters of metamorphic events demonstrates the similarity of theArchaean and Proterozoic eclogites worldwide and their affinitywith “hot” environments. In contrast, Phanerozoic HP and UHPeclogite complexes formed in connection with “warm” subductionzones. The high-T character of the eclogite-facies metamorphismin the Proterozoic is impossible to explain through a distinctivelyanomalous thermal regime as is often suggested for the Archaeanmantle. Consequently, the thermal regime of the Archaean eclogite-facies metamorphism represented by P–T parameters of the Salmaand Gridino eclogite associations cannot be considered as confirmingthe concept that the Archaean mantle was warmer in comparisonwith later periods in Earth history. Instead, to explain the appearanceof high-T conditions during eclogite-facies metamorphism, we con-sider that a decisive role was played by processes associated withsubduction of mid-ocean ridge (Archaean, BEP) or bymantle plumes(Proterozoic).

Acknowledgments

We thank the reviewers, Tim Kusky and PSW (anonymous), forvery useful comments, which helped us to clarify a number of signif-icant positions of this paper. We are especially grateful to PSW for his(her) great and hard work that significantly improved English in ourtext.

We are grateful to the Russian Foundation for Basic Research, pro-jects 09-05-00926, 09-05-01006, 11-05-00492, and 12-05-00856 forsupport of this study. This work contributes to the Program no. 6 ofthe Earth Sciences Department of the Russian Academy of Sciences.

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The Meso-Neoarchaean Belomorian eclogite province: Tectonic position and geodynamic evolution

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