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Precambrian Research 259 (2015) 5–33

Contents lists available at ScienceDirect

Precambrian Research

jo ur nal homep ag e: www.elsev ier .com/ locate /precamres

rans-Baltic Palaeoproterozoic correlations towards theeconstruction of supercontinent Columbia/Nuna

. Bogdanovaa,∗, R. Gorbatscheva, G. Skridlaiteb, A. Soesooc, L. Tarand, D. Kurloviche

Department of Geology, University of Lund, SwedenInstitute of Geology and Geography, Nature Research Centre, Vilnius, LithuaniaInstitute of Geology, Tallinn University of Technology, Tallinn, EstoniaRepublican Unitary Enterprise “Research and Production for Geology”, Minsk, BelarusDepartment of Soil Science and Land Informational Systems, Belarusian State University, Minsk, Belarus

r t i c l e i n f o

rticle history:eceived 11 June 2014eceived in revised form 29 October 2014ccepted 25 November 2014vailable online 4 December 2014

eywords:alaeoproterozoicupercontinent–Pb zircon agesvecofennian accretionary orogenrans-Baltic correlations

a b s t r a c t

A comparative study of the central and southern parts of the Palaeoproterozoic Svecofennian orogen inthe Baltic/Fennoscandian Shield and the platform area to the east and south of the Baltic Sea indicatesthat at least these parts of the orogen are built up of several NW-SE trending, 100–300 km wide tectonicmegadomains separated from each other and complicated by major zones of mostly dextral shearing.The generation of these zones occurred successively between 1.86 and 1.75 Ga, concomitantly with con-tinuing crustal accretion younging towards the southwest. Even considering the distorting presence of anumber of microcontinents, this indicates the one-time existence and repeated episodic activity of a mas-ter subduction zone stepwise falling back to the present south-southwest. At 1.82–1.80 Ga, the obliquecollision of protocontinents Volgo-Sarmatia and Fennoscandia interfered with the accretionary growthof the crust in the Svecofennian orogen. In the west, the system of Svecofennian tectonic domains andshear zones is delimited by 1.70–1.55 Ga orogenic belts marking the Laurentia-Greenland-Baltica marginof Columbia. Altogether, the available U–Pb zircon datings and studies of key rocks and structures in

the South Baltic region allow more detailed Trans-Baltic correlation and the creation of new integratedmodels of the structural and tectonic evolution of the Svecofennian orogen in particular and northernEurope in general. The new findings will be important also in the continuing study of supercontinentformation and supercontinent cycles, and the drifting of Palaeoproterozoic protocontinents during theassembly of Columbia/Nuna.

© 2014 Elsevier B.V. All rights reserved.

. Introduction

New insight gained during recent reconstructions of Pre-ambrian supercontinents Rodinia (e.g. Li et al., 2008) andolumbia/Nuna (e.g. Zhao et al., 2004; Santosh et al., 2009; Evansnd Mitchell, 2011; Zhang et al., 2012) has been material to assesshe duration and specifics of Archaean and Proterozoic supercon-inent cycles and thus also of the changes in the Earth systemith time. While the configuration of Rodinia and the history of

ts assembly and break-up appear reasonably well established atresent (Li et al., 2008; Evans, 2013), in the case of Columbia/Nuna

here remain numerous problems related to the correlation ofithotectonic complexes and structures between various conti-ents and across terrane boundaries, which allows very different

∗ Corresponding author. Tel.: +46 706575132; fax: +46 46 2224419.E-mail address: Svetlana.Bogdanova@geol.lu.se (S. Bogdanova).

ttp://dx.doi.org/10.1016/j.precamres.2014.11.023301-9268/© 2014 Elsevier B.V. All rights reserved.

reconstructions. Due to the persisting shortage of isotopic age con-straints and geochemical data in a number of critical contexts, alsothe recognition and spatial connection of Palaeoproterozoic col-lisional and accretionary orogenic belts is still far from complete.Similarly, rather little is known about the scales, extents and natureof the oceans that must have been consumed during the amalgama-tion of Columbia at ca. 1.8 Ga. Some of these were internal, othersexternal (Windley, 1993) in correspondence with the formation of“introvert” and “extravert” supercontinents in the sense of Murphyand Nance (2003) and Nance et al. (2014), but it is not clear howthey were distributed across the planet during the Palaeoprotero-zoic.

In the present paper, we consider some aspects of the assem-bly of Columbia, focusing in particular on the correlation of the

part of the ca 2.0–1.75 Ga Svecofennian orogen exposed in the cen-tral and southern Baltic/Fennoscandian Shield with its continuationbeneath the Phanerozoic platform cover to the southeast of theBaltic Sea (Fig. 1).

6 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Fig. 1. (a) Major Palaeoproterozoic tectonic domains in the Baltic Sea area. The white diagonal ruling indicates Svecofennian sedimentary basins. The right-corner insetshows the three-segment subdivision of the East European Craton acc. to Bogdanova (1993) and Gorbatschev and Bogdanova (1993) and the left-corner inset the distributionof granulite- and amphibolite facies metamorphic rocks. The abbreviations for tectonic domains are: AL – Alutaguse, BB – Bothnian (Bothnia microcontinent), BS – Bergslagen(Bergslagen microcontinent), BPG – Belarus-Podlasie granulite belt, CE – Ciechanow, CFAS – Central Finland Arc Complex; CFGC – Central Finland Granitoid Complex (Keitelemicrocontinent), DO – Dobrzyn, ESL – East Småland, JO – Jõhvi, KB – Keitele microcontinent, KZ – Kaszuby, LA – Latgalia, LEL – Latvian-East Lithuanian, LGB – LaplandGranulite Belt, LKO – Lapland-Kola orogen, LS – Ljusdal, MD – Mazowsze, MLD – Mid-Lithuanian domain, NB – Norrbotten, NO – Novgorod, OKL – Okolovo, OMIB – Osnitsk-Mikashevichi Igneous Belt, PM – Pomorze, SEG – South Estonian granulite domain, T – Tapa, TN – Tallinn, Uu – Uusimaa, VV – Västervik, WE – West Estonian domain, WLG– West Lithuanian granulite domain. The abbreviations for deformation zones are: HGZ-GR – Hagsta-Gävle-Rättvik Zone, HSZ – Hassela Shear Zone, KSZ – Karlskrona Shearzone; LLSZ – Linköping-Loftahammar Shear Zone, MEFZ – Middle Estonian Fault Zone, PPDZ – Paldiski-Pskov Deformation Zone, SFSZ – South Finland Shear Zone, VNSZ –V HämeT P - Rim (a).

FbteoGPe1spfnac

ingåker-Nyköping Shear Zone and for volcanic belts and sedimentary basins: H –

m – Tampere; for major AMCG and A-type granitoid intrusions: MZ - Mazury, Ragnetic anomalies (modified after Wybraniec, 1999). The abbreviations are like in

Ever since Anna Hietanen (1975), the Svecofennian orogen in theennoscandian segment of the East European Craton (Baltica) haseen known as a prime instance of Palaeoproterozoic accretionaryectonics (Gaál and Gorbatschev, 1987; Nironen, 1997; Lahtinent al., 2005, 2009b), which more recently was interpreted as partf the active margin of Columbia along the edges of Laurentia,reenland and Baltica. While the evolution of this margin in the latealaeoproterozoic and the Mesoproterozoic has been well mod-lled (Gower et al., 1990, 2008; Rivers, 1997; Åhäll and Gower,997; Karlstrom et al., 2001), its early Palaeoproterozoic historytill appears fragmentary and possible variations of palaeogeogra-hy, tectonic regime and style of deformation must be accounted

or. This is particularly true of the Svecofennian orogen, which isow widely accepted to have been formed and developed by theccretion of microcontinents and island arcs scattered in a “Sve-ofennian ocean”. The existence of a 2.0 Ga old ocean (or oceans)

, Mk – Monki, O-J – Oskarshamn-Jönköping, P – Pirkanmaa, Pc – Poceai, Sc – Salcia,ga pluton. (b) Boundaries of the tectonic domains superimposed onto the map of

along the present-day SW margin of the (Kola-) Karelian and theNorrbotten cratons is demonstrated by relics of oceanic crust inthe 2.0 Ga Kittilä volcanic belt and by the ca. 1.95 Ga Jormua andother finds of ophiolites in Finland (Nironen, 1997; Peltonen andKontinen, 2004; Lahtinen et al., 2005; Melezhik and Hanski, 2012).At what time that ocean was formed is not wholly clear, however.From the distribution of mafic dyke swarms across the boundariesof the Superior and Kola-Karelia cratons, Bleeker and Ernst (2006)suggested that the break-up of the Archaean Superia superconti-nent and the final separation of Kola-Karelia (and Norrbotten) tookplace between 2.1 and 1.98 Ga. During that period, several internaloceans were opened, which divided the Kola-Karelia craton into

a number of minor continental blocks (e.g. Melezhik and Hanski,2012). The first signs of rift-to-drift and accretionary processesalong the edge of the northwestern part of the Karelia protocra-ton were the formation of passive margins and ca. 1.96–1.92 Ga

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rimitive island arcs like the Knaften arc (Wasström, 2005) and theavo oceanic arc (Lahtinen, 2000).

To explain this period and the subsequent formation of the Sve-ofennian orogen (or several separate, independent orogens in itslace) at 2.0–1.75 Ga, there exist two essential and in a way com-eting groups of plate tectonic interpretations:

(a) One of these envisages microplate tectonics and rapid “arc-accretion” with associated multiple subduction of the oceaniccrust in various directions, formation of arcs and their accre-tion, and collision of microcontinents with each other andwith the arcs, and also with the Archaean Karelia and Nor-rbotten cratons. Five orogenies or principal orogenic eventsbetween 1.92 and 1.79 Ga and four stages of orogenic evo-lution overlapping each other in space and time have beenproposed (Nironen, 1997; Lahtinen et al., 2005, 2009b). Thefour stages comprise microcontinent accretion (1.92–1.87 Ga),continent extension (1.86–1.84 Ga), continent–continent col-lision (1.84–1.79 Ga) and orogenic collapse and stabilization(1.79–1.77 Ga). This group of interpretations obviously refutesthe previous concepts of a “semi-continuous” Svecofennianevolution. In order to explain the “equidimensional” patterns ofthe continental crust in the Svecofennian orogen, this model hasrecently been modified by involving orocline tectonics (Beunkand Kuipers, 2012), as well as rotations of primarily linear ter-ranes and their “tectonic wrecks” (Lahtinen et al., 2014).

b) The other group of interpretations, which proposes marginalbasin accretion governed by a master subduction zone, has beenemployed to explain, among other things, the wide distributionof similar populations of detrital zircon and basin formation inthe northern and central parts of the Svecofennian orogen inSweden and Finland at 2.1–1.98, 1.97–1.92, and 1.91–1.86 Ga,while short events of accretion and compression interruptedbasin formation at 1.98, 1.92 and 1.88–1.86 Ga (Rutland et al.,2001, 2004; Williams et al., 2008). A related model, namedthe “migratory tectonic switching”, emphasizes the existenceof cycles of alternating extension and compression during thedevelopment of the Svecofennian orogen in central Sweden andalso elsewhere (Hermansson et al., 2008). It appears similarto the concept of “retreating-advancing” accretionary orogens(Collins, 2002; Cawood et al., 2009).

Various versions of both groups of interpretation challenge theiew that the Svecofennian orogeny and particularly the timingf the involved subduction were related directly to the formationf the coherent Laurentia – Greenland – Fennoscandia (Baltica)argin of assembling Columbia. Whether the Svecofennian oro-

en was formed by the consumption of an ocean along the marginf a large continent in the “West Pacific” style or by consumingeveral mini-oceans with interspersed microcontinents and arcs isnother still unresolved problem that also concerns the positionf Palaeoproterozoic Fennoscandia in relation to other continentallocks within the general Columbia framework at ca. 1.9–1.8 Ga.lso the dimensions of the Palaeoproterozoic ocean that separatedrotocontinents Fennoscandia and Volgo-Sarmatia, the two prin-ipal lithospheric components of Baltica, and their distances fromach other have not yet been determined by palaeomagnetic dataor the period between 2.0 and 1.8 Ga. Thus there remain manyncertainties in regard to their relationships and the precise timesf various stages of assembly (Bogdanova et al., 2013; Pisarevskyt al., 2014).

Completing the palaeogeography between ca. 2.0 and 1.8 Ga,

hich is known mostly from the Baltic/Fennoscandian Shield, weresent in this paper new data on the ages, compositions and ori-ins of the lithotectonic complexes that constitute a large partf the Palaeoproterozoic crust between the Baltic/Fennoscandian

Research 259 (2015) 5–33 7

and Ukrainian shields to the southeast of the Baltic Sea (Fig. 1).These shields characterize much of the Palaeoproterozoic evolutionwithin Fennoscandia and Volgo-Sarmatia, the two distinct crustalsegments of the East European Craton (Baltica) that belonged to twodifferent lithospheric plates after ca. 2.0 Ga (e.g. Bogdanova et al.,2008). Even though the crystalline crust there is hidden beneath aLate Proterozoic to Phanerozoic sedimentary cover reaching thick-nesses of up to 2 km, it has been disclosed by nearly a thousanddrillings, largely with still available core materials (e.g. Bogdanovaet al., 2006 and references therein). This allows Palaeoproterozoiccorrelation between the exposed lithotectonic complexes in theshield area and their covered counterparts across the Baltic Sea. Inits turn, trans-Baltic correlations make it possible to evaluate theextent and scales of the Baltica Palaeoproterozoic crustal structuresto be used for global Precambrian reconstructions.

To improve our knowledge of the Palaeoproterozoic evolution inwestern Fennoscandia, several new key-objects were dated in thepresent study. Particular attention was devoted to mylonites andother fault-related rocks along the deformation zones that separatethe various crustal units. We also revisited the existing geochrono-logical and geochemical data to determine terrane boundaries andto reconstruct the accretionary evolution of the South Baltic regionin relation to the Palaeoproterozoic tectonics in the Baltic Shield.

2. Principal tectonic subdivisions of the crust in the SouthBaltic region

All the way since the nineteen-thirties, it has been known thatthe Precambrian crust in the Baltic/Fennoscandian Shield con-tinued beneath the platform cover to the south and east of theBaltic Sea. However, it took tedious studies of geophysical patternsand numerous age determinations from drilling wells to realizethat most of that crust was not Archaean but rather belongedto the Palaeoproterozoic Svecofennian orogen (Gorbatschev andBogdanova, 1993; Bogdanova et al., 1994, 2006; Claesson et al.,2001). Actually, there exist no zircon-age indications of any pres-ence of Archaean crust or Archaean mantle-melt sources in theSouth Baltic region (SBR), all the Archaean ages in that area hav-ing been derived from detrital zircons in metasedimentary rocks(Puura and Huhma, 1993; Bogdanova et al., 1994; Bibikova et al.,1995; Claesson and Ryka, 1999; Claesson et al., 2001; Mansfeld,2001; Wiszniewska et al., 2007; Williams et al., 2009; Krzeminskaet al., 2011; this paper). As different from most of the shield area,however, the patterns of the Palaeoproterozoic crust beneath theplatform cover in Belarus, Estonia, Latvia, Lithuania and Polandare defined by a conspicuous array of alternating belt-shaped tec-tonic domains made up of rocks either of granulite or amphibolitefacies (Fig. 1a, inset). Many of these belts that were formed rela-tively late in the development history of the Svecofennian orogenhave been interpreted to represent crust stacked tectonically dur-ing the collision of Fennoscandia and Volgo-Sarmatia (hereafterthe “Fenno-Sarmatian” collision) at ca. 1.82–1.80 Ga (Bogdanovaet al., 2006). Commonly, the belts characterized by various grades ofmetamorphism overlap structurally defined tectonic domains andterranes, which each had their own pre-collisional Palaeoprotero-zoic histories and mostly are bounded by zones of major shearingand faulting (Skridlaite and Motuza, 2001; Taran and Bogdanova,2001; Soesoo et al., 2004; Bogdanova et al., 2006). The 40Ar/39Arages of amphibole from mylonites along the fault boundaries inLithuania and Belarus are all 1.69–1.67 Ga regardless of the agesand metamorphic grades of the rocks within the belts (Bogdanova

et al., 2001). This suggests that the ultimate arrangement of theSouth Baltic belts took place as late as ca.1.7 Ga ago.

A number of other features of the present-day structure inthe region and in its deep crust are largely due to extensional

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ectonics with recurrent events of high-grade metamorphismnd reactivation of the lithosphere, including major AMCG=Anorthosite-Mangerite-Charnockite-Granite) and A-type mag-

atism at ca. 1.65–1.62, 1.60–1.58, 1.53–1.50 and 1.50–1.45 GaBogatikov and Birkis, 1973; Rämö et al., 1996; Bogdanova et al.,001, 2006; Wiszniewska et al., 2002, 2007; Skridlaite et al., 2003b,008, 2014; Motuza et al., 2006; Duchesne et al., 2010; Vejelytet al., 2010). These were most probably far-field effects of remoteate Palaeoproterozoic and Mesoproterozoic orogenic processeslong the south-western extreme of the East European CratonÅhäll et al., 2000; Bogdanova et al., 2008; Skridlaite et al., 2014).

system of curving faults around the large AMCG Riga pluton isistinctly superposed upon the domain boundaries, while transten-ion at 1.53–1.50 Ga created a system of E-W and NW trendingones of faulting partly reactivating older shear zones. Subse-uently, at 1.47–1.44 Ga, these zones became transpressional andccommodated numerous small syntectonic intrusions and vari-us migmatites. Bogdanova (2001) referred these two events tohe Danopolonian orogeny. In Lithuania and in western Belarusanopolonian tectonism fundamentally rearranged and compli-ated the older Palaeoproterozoic crustal structures by displacingoundaries and creating bends, which much altered the early struc-ural trends (Skridlaite and Motuza, 2001; Motuza, 2005).

Altogether, the post-1.8 Ga processes modified the pre-ollisional palaeogeography, altering the spatial distribution ofithotectonic complexes and the tectonic grain of the South Balticegion, which enhanced the differences between the exposed shieldnd the hidden platform parts of Fennoscandia separated by thealtic Sea. This created problems of Trans-Baltic correlations andeconstruction of Palaeoproterozoic Fennoscandia (and Baltica),ome of which still remain unsolved.

In the following sections, we consider the most important char-cteristics of the various Palaeoproterozoic tectonic domains andelts in the South Baltic region. Many of these are distinct inegard to the compositions of their rock complexes, metamor-hism, and petrophysical properties. The various lithotectonic unitsll have their own geophysical patterns and conspicuous magneticnd gravity lineaments extending all the way to their boundariesFig. 1). As established by the EUROBRIDGE and POLONAISE seis-

ic profiling (Fig. 2b, for particulars see Bogdanova et al., 2006),hey are also different with regard to the thickness and inter-al structure of the crust. During the latest decade, many traitsf the compositions and tectonothermal evolutions of the vari-us domains and belts in the South Baltic region were assessed,ut some of these are now being reconsidered in regard to theirges, configurations, boundaries and bounding relationships. Par-icular attention is devoted to the origins and ages of the granuliteelts and domains which are a hallmark of the entire region (Fig. 1,

nset).All the currently available age determinations and Sm–Nd iso-

opic data for the South Baltic region are listed in Appendix A.

.1. Crustal domains in Estonia

The structure of the crystalline basement in Estonia andorthern Latvia is characterized by several NW and E-W trend-

ng large deformation zones constraining six different tectonicomains/structural zones (Koistinen, 1994; Soesoo et al., 2004).he ca. 30 km wide, NW-trending Paldiski-Pskov deformation zonePPDZ in Figs. 1 and 2) separates the structural domains in northernnd north-eastern Estonia from those in the west and the south. Itomprises two principal sub-parallel shear zones dipping between

5◦ and 75◦ SSW, features strong deformation, and contains sliv-rs of both granulite- and amphibolite-facies rocks. Characteristicre high gradients of the potential fields and a displacement ofhe seismic layering and the Moho boundary as documented along

Research 259 (2015) 5–33

the Sovetsk–Kohtla Järve deep seismic refraction profile (All et al.,2004).

Among the domains in NNE Estonia (Fig. 1), the Tallinn (TN),Tapa (T) and Alutaguse (Al) differ markedly from the other Esto-nian tectonic units by their dominantly amphibolite facies ofmetamorphism that passes gradually into the granulite facies atca. 3–5 kbar (Puura et al., 1983; Soesoo et al., 2006). They aremade up of metasedimentary and metavolcanic rocks ca. 1.90 Gain age (Appendix A), which resemble closely the rocks of adja-cent belts in southern and western Finland, and those in centralSweden. The Tallinn domain appears to represents a volcanicarc that continues in the Uusimaa belt in Finland (Kähkönen,2005). The metasedimentary Alutaguse domain, however, maybe part of the large Kalevian-age marginal basin that extends tothe vicinity of St. Petersburg in Russia and farther east to LakeLadoga. The deposition of turbidites in that basin has been esti-mated to have taken place between 1.96 and 1.90 Ga (Lahtinenet al., 2010; Melezhik and Hanski, 2012). However, the ages ofdeposition and metamorphism of the Alutaguse metasedimentarysequence are still unknown and must be obtained before its tec-tonic setting can be set up. Similar turbidites may also occupythe Novgorod domain in Russia (NO in Figs. 1 and 2), which ischaracterized by a particularly irregular mosaic field of low mag-netic anomalies, resembling that of the Bothnian domain in theBaltic/Fennoscandian Shield (BB in Fig. 1). The Novgorod domainhas also been interpreted as an ancient microcontinent (Peive et al.,1979).

The separate Jõhvi belt in NE Estonia (JO in Figs. 1 and 2)harbours a thick sequence of alternating Fe- and S-rich garnet-pyroxene-bearing quartzites, high-Al garnet-cordierite-sillimanitegneisses, and Ca-rich and Ca-poor pyroxene-, amphibole- andbiotite gneisses. The migmatization of these rocks and their pen-etration by granite occurred at ca. 1.8 Ga (Appendix A), which issimilar to the ages of late-and/or post-orogenic granites in southernFinland (Nironen, 2005; Kurhila et al., 2011).

The West Estonian domain (WE) that is bounded by the NW-trending PPDZ and the E-W striking Middle Estonian (MEFZ)deformation zones is dominated by wide areas of metasedimentaryrocks in the amphibolite to granulite facies, resembling somewhatthe rocks in the southern part of the Bergslagen tectonic domainin Sweden. Such rocks are less common farther SE in the SouthEstonian granulite domain (SEG in Figs. 1 and 2), which mostlycomprises mafic, intermediate and felsic meta-igneous granulitesrecorded geophysically by a series of E-W to NW-trending linearmagnetic and gravity anomalies. Chemically, some charnockitescan be attributed to a calc-alkaline I-type series, likely generateddue to the melting of lower crust under granulite facies condi-tions along with the emplacement of mafic dykes at ca. 1.77 Ga(Soesoo et al., 2006). This is somewhat younger than the 1.81 Gaage of granulite facies metamorphism in northern Latvia (Mansfeld,2001), into which the SEG domain continues. In the southeast, thatcontinuation incorporates the large Latgalian (also East Latvian)block (LA in Figs. 1 and 2) of mosaic-type magnetic anomalies. Itis surrounded by linear anomaly belts and has previously beeninterpreted as an “Archaean” microcontinent (Bogatikov and Birkis,1973) similar to the Novgorod domain. Here, it should be notedthat in our current interpretation the SEG domain is no longerthe northern extension of the Belarus-Podlasie granulite belt (BPG)but is truncated by the latter, which is now considered to asso-ciate with Sarmatia rather than Fennoscandia (cf. Sections 2.6 and4).

2.2. The Latvian-East Lithuanian domain (LEL)

Rocks of various origins and metamorphic facies can bedistinguished in the LEL, which features pronounced N-S

S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33 9

Fig. 2. Structure of the crust in the South Baltic region (modified after Bogdanova et al. (2006)); (a): major tectonic subdivisions and sites of dated drill cores (the numberscorrespond to those in Appendix A). Black diamond symbols mark zircon sampling sites of the present study and white diamonds the sites of previously dated rocks, whitecircles show the sampling sites for Sm–Nd isotope data. Thin lines show structural trends following linear magnetic anomalies. (b): EUROBRIDGE’95 and 96 integrated seismicrefraction profile and its revised interpretation (modified after Bogdanova et al., 2006). Density values are from the density–velocity modelling by Kozlovskaya et al. (2001,2002). Note that the geological interpretation of the geophysical gravity-seismic model has been revised in accordance with the tectonic subdivisions of the crust proposedin the present study. Abbreviations as in Fig. 1.

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rending magnetic and gravity anomalies (Fig. 1). In central Latvia,orphyritic mafic, intermediate and felsic metavolcanics, and asso-iated variously metamorphosed Fe-Mn-rich quartzites, pelites andarly rocks participate in a prograde metamorphic zoning extend-

ng from low to high amphibolite facies at moderate pressuresBogatikov and Birkis, 1973; Vetrennikov, 1991). All these metavol-anic and metasedimentary rocks appear juvenile with �Nd up to3–4 (Appendix A). A meta-rhyodacitic rock from northern Latviaith a age of 1870 ± 13 Ma was metamorphosed in granulite facies

t 1810 ± 2 Ma (Mansfeld, 2001).In Lithuania, farther to the south, felsic and mafic metavol-

anics as well as sedimentary rocks, particularly metagreywackesnd marbles, have been recovered by some drillings (Skridlaitend Motuza, 2001; Motuza, 2005). Mg- and Fe-rich skarns asso-iate particularly with dolomite-rich crystalline carbonates andeta-igneous rock successions with iron-ore deposits (Skridlaite

nd Motuza, 2001). While the metagreywackes in the southern LELontain some Archaean detrital zircon, Palaeoproterozoic zircon isominant and Sm–Nd isotopic modelling indicates Palaeoprotero-oic ages of deposition (Mansfeld, 2001).

The peak conditions of metamorphism in these metasedimen-ary rocks were 650–700 ◦C at 6–7 kbar, but most of these rocksield ca. 500 ◦C at 3–4 kbar. Migmatites as well as granite, dior-te and gabbro intrusions are common throughout the area. In theresent study, we dated a 1.89 Ga metadioritic-granodioritic rockrom the southern part of the West Lithuanian granulite domainWLG).

.3. The Mid-Lithuanian domain

The Mid-Lithuanian domain as defined in the present paperMLD in Figs. 1 and 2) has the shape of a curved, eastwards convexelt that ranges in width between 20 km in the SE and more than00 km in the NW, where part of it is separated from the northernart of the WLG by the tens of kilometres wide Telsiai deforma-ion zone (TDZ; Vejelyte, 2011) within which WLG and MLD rocksccur together, intercalated with each other (Figs. 1 and 2a). Theesults of the present study necessitate a fundamental reconsider-tion of the previous concepts with regard to the nature, size andhape of what was depicted as the Mid- (or Central-) Lithuanianuture Zone in Skridlaite and Motuza (2001) and Motuza (2005),hich model had also been adopted in Bogdanova et al. (2006),

ven though the southern and western boundaries of the MLD,s shown in Figs. 1 and 2, may indicate the presence of a sutureetween the SE-most part of the MLD and the generally youngerLG.The MLD consists dominantly of magmatic rocks (Claesson et al.,

001; Rimsa et al., 2001; Motuza, 2005; Motuza et al., 2008;kridlaite et al., 2011) and may possibly represent a one-time activeontinental margin, but sedimentary rocks occur only locally. Theorthern and western parts of the MLD comprise various charnock-

tes (sensu lato) and mafic granulites. Their ages vary between ca..86 and 1.82 Ga (Claesson et al., 2001; Motuza et al., 2008; Vejelytet al., 2012). The southern and central parts, however, are dom-nated by granite, granodiorite, diorite and gabbro suites with aresence also of their volcanic counterparts. Geochemically, allelong to the TTG calc-alkaline series and are interpreted as rem-ants of volcanic arcs (Rimsa et al., 2001; Motuza, 2005). The weaklyetamorphosed ca. 1.86–1.84 Ga gabbros and diorites in the south-

rn MLD (Skridlaite et al., 2011) resemble chemically the maficranulites in its northern and central parts, and so does also the

a. 1.84 Ga metagabbro dated in this study (sample Gl99, site 32 inig. 2a). A cordierite-bearing granite in the north, a strongly shearedranite, and porphyritic pyroxene-bearing meta-andesites all haveimilar ages close to 1.84 Ga (Motuza et al., 2008).

Research 259 (2015) 5–33

2.4. The West Lithuanian granulite domain

The West Lithuanian granulite domain (WLG) is a compositecrustal unit dominated by felsic and intermediate metavolcanicand metasedimentary rocks (Figs. 1 and 2a). Metasedimentarygranulites occur scattered throughout the WLG, however in thenorthwest they occupy a coherent area of several tens of squarekilometres, here named the Poceai (Pc) basin. In some drilling wells,the thicknesses of their sequences reach several hundred metres.Farthest west, the metapelites (sites 44 and 45 in Fig. 2a) containrare metapsammitic layers and numerous granitic veins and lensesindicating extensive migmatization. Some coherent bodies of gran-odiorites, granites and rare charnockites can also be distinguishedamong the metasedimentary rocks. Farther southeast, anothersequence of metapelites (Bl150, Fig. 2a), which are interbeddedwith metavolcanics, is interpreted to belong to an island arc simi-lar to 1.83–1.81 Ga volcanic belts in Poland (cf. below). These WLGmetavolcanics still need to be properly dated, which is true alsoof associated porhyritic felsic and intermediate volcanic and plu-tonic rocks of still uncertain ages that dominate the southern WLG.Generally, metamorphism decreases towards the south, and onlylocally granulites are present.

Some rocks in the southern WLG appear to be equivalents ofthose in the adjacent LEL and MLD (Skridlaite and Motuza, 2001;Skridlaite et al., 2003a,b), and it has been considered that the WLGand the adjacent part of the MLD were overthrusted onto the LEL(cf. Fig. 2b and Section 4).

2.5. The Mazowsze, Dobrzyn and Pomorze domains

Geophysically and tectonically, the crystalline basement in NEPoland can be subdivided into the Mazowsze (MD), Dobrzyn (DD)and Pomorze (PM) domains separated by roughly N-S trend-ing fault zones, which were named the Ciechanow and Kaszubymagmatic-metamorphic belts by Kubicki and Ryka (1982). Whilethese domains are characterized by mosaic, irregular magneticand gravity anomalies of low to medium intensity, linear posi-tive anomalies record zones of faulting (Wybraniec, 1999). ThePOLONAISE P3, P4 and P5 refraction seismic profiling (Sroda,1999; Czuba et al., 2001, 2002) recognized these tectonic subdi-visions in the seismic structure of the crust, in which the MazuryAMCG intrusions define high Vp-velocity bodies in the upper crustand corresponding upwellings of the mantle. The Ciechanow andKaszuby zones of faulting accommodate Mesoproterozoic Mazuryplutons as well as Carboniferous intrusions (Krzeminska et al.,2014), which explains the occurrence of high gravity and mag-netic anomalies along these zones. As far as the three domainsare concerned there exist some differences of their crustal seismicstructures and the ages of their granitoids, which, however, yieldsimilar Nd model ages of ca. 2.1–2.0 Ga throughout (Claesson andRyka, 1999). The best studied Mazowsze domain features volcanicrocks with subduction-related geochemistries and granitoids withcrystallization ages of 1.83 and 1.80 Ga (Valverde-Vaquero et al.,2000; Krzeminska et al., 2005; Wiszniewska et al., 2007). Recon-naissance work on the detrital zircons in the metasedimentaryrocks of the Mazowsze basins (Williams et al., 2009) has yieldedmaximum deposition ages between 1.86 and 1.83 Ga, the latteralso being the age of metamorphism in the area (Appendix A). Anoverlying sequence of weakly metamorphosed quartzites containsdetrital zircon as young as 1.76–1.75 Ga (Krzeminska et al., 2011). Incombination with the results of current comprehensive geochrono-logical work on the basement rocks of the region (Krzeminska

et al., 2014), this age pattern resembles that of the Late Svecofen-nian evolution in SE Sweden, where the formation of the juvenile1.83–1.82 Ga Oskarshamn-Jönköping volcanic arc (Mansfeld et al.,2005) was followed by the voluminous bimodal magmatism that

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ormed the older, 1.81–1.76 Ga generation of rocks referred to theransscandinavian Igneous Belt (TIB-1 in Högdahl et al., 2004).

.6. The Sarmatia-related Belarus-Podlasie granulite and Okolovoelts

As different from previous work on the tectonic subdivisionsf the South Baltic region (Bogdanova et al., 2006), the Belarus-odlasie granulite belt (BPG) and the Okolovo belt (OKL) adjacento the Fennoscandia-Sarmatia suture zone (Figs. 1 and 2a, b) areonsidered “exotic” in the present paper and not related directlyo the Svecofennian terranes proper. The reasons for that and fur-her considerations of these issues will be found in the “Discussion”ection 4.1.3.

The BPG is made up of several large lens-shaped bodies of gran-lite, separated from each other by zones of shearing and faultingarked by mylonites and granulites retrograded to the amphibolite

acies (Aksamentova and Naydenkov, 1990; Bogdanova et al., 1994).he BPG rocks proper are mostly Palaeoproterozoic granulites ofafic, enderbitic and charnockitic compositions, which belong to

wo different igneous suites. The older of these suites has an age ofa. 1.88 Ga (Claesson et al., 2001) and is calc-alkaline in composi-ion, while the younger is ca. 1.80 Ga, chemically more variable,lkali-calcic, and bimodal. Metapelitic gneisses and migmatitesre subordinate. Their Sm–Nd isotopic characteristics (TDM 2.3,Nd(1.9) −1.3) (Appendix A), as well as those of the intrusiveocks (TDM 2.3 Ga, �Nd(1.85) −0.5), indicate some contributionsf older materials. A few zircon ages from the BPG metapeliticigmatites plot between 1935 and 1910 Ma (detrital cores), while

natectic melting and migmatization in the prograde amphiboliteacies took place at 1880–1860 Ma. Granulite facies metamorphism750 ◦C at 7–9 kbar) overprinted the amphibolite facies migmatitest ca. 1.79 Ga (Bibikova et al., 1995; Claesson et al., 2001; Tarannd Bogdanova, 2003; Skridlaite et al., 2014). The 40Ar/39Ar ages ofmphiboles from mafic granulites and metagabbros within the BPGary between 1773 and 1710 Ma, while those from mylonites andtrongly sheared gneisses and amphibolites in deformation zonesostly range between 1700 and 1663 Ma. Along a few fault zones,

hese ages are even younger, plotting at 1550–1530 Ma (Bogdanovat al., 2001). Altogether, these data suggest that at ca.1.9 Ga thePG was a relatively mature magmatic arc, formed concomitantlyith the subduction of oceanic crust beneath the margin of Volgo-

armatia, which is indicated by the existence of the 2.0–1.95 Gakolovo (-Rudma) oceanic terrane and the 1.98–1.95 Ga Osnitsk-ikashevichi igneous belt (OMIB in Figs. 1 and 2a).The Okolovo (-Rudma) belt (OKL) forms a WNW-dipping, tec-

onically delimited slab of rock, nearly 10 km thick (Aksamentovat al., 1994). Along its borders towards the Belarus-Podlasie beltBPG) in the present NW, the rocks have been metamorphosed inhe granulite facies, forming the continuous Rudma metamorphicone abutting against the SE edge of the BPG. An enderbite fromhis zone has yielded magmatic crystallization ages of ca. 1.95 Gand ages of metamorphism of ca. 1.75 Ga (Appendix A).

The OKL is built up of basalts, andesites, dacites and rhyo-ites, and their tuffs intercalated with metasedimentary rocks. The

etadacites have yielded a U–Pb zircon age of 1998 ± 10 Ma andre juvenile with �Nd (T) +3.3 (Bibikova et al., 1995; Claesson et al.,001). Among the metasedimentary rocks that form the Okolovo-udma sequence, greywackes, black shales and ferrous as wells siliceous-carbonatic and volcanogenic chemical deposits areery conspicuous. The isotopic characteristics of carbon in graphite�13C −15.8 ÷ −23.2‰) and sulphur in pyrite (�34S = +9.6 − +28.2‰)

rom the black shales, as well as �13C values between −6 and 0n calcite from the carbonate rocks indicate a substantial involve-

ent of organic materials in the sedimentation (Taran et al., 2006).his agrees with the global increase of organic matter during

Research 259 (2015) 5–33 11

the so-called “Shunga” event at ca. 2.0 Ga (Melezhik et al., 1999).Lithotectonic reconstructions of the Okolovo-Rudma palaeobasin(Aksamentova et al., 2006) indicate that this basin was oceanic,which together with the chemistries of the metavolcanic rockssuggests an oceanic-arc tectonic setting and, in consequence, theexistence there of an ocean at ca. 2.0 Ga.

Prograde metamorphism in the amphibolite to granulite facies(650–725 ◦C at 7–8 kbar) occurred at ca. 1.9 Ga (Taran andBogdanova, 2001), which coincides with the timing of dioriteand tonalite intrusions featuring mantle-related �Nd (T) of +2.1(Claesson et al., 2001). Most probably, these events were relatedto the propagated development of an arc system. An accretionof the Okolovo terrane to a continental block or another arc (thesubsequent BPG?) at ca. 1.9 Ga cannot be excluded since majorcalc-alkaline magmatism and early metamorphism occurred in theadjacent BPG at the same time.

3. U–Pb zircon geochronology

In this paper, we report new U–Pb zircon datings of seven rocks,which were selected to promote the correlation of the Palaeo-proterozoic crust-forming and tectonic events in the South Balticregion with those known from the Baltic/Fennoscandian Shield. Wesampled:

1. Two rocks on either side of the Middle Estonian deforma-tion zone (MEFZ), one of these a high-grade metasedimentarymylonite at the southern edge of the West Estonian domain, theother a granitoid granulite in the NW-most part of the SouthEstonian domain (sites 7 and 10 in Fig. 2a, respectively). Theaim was to date the deformation and metamorphism of thesegranulites and also the source rocks of the detrital zircons in themetasedimentary granulites of the West Estonian domain.

2. A mafic granulite in the Mid-Lithuanian domain and a graniticmylonite along its boundary towards the Belarus-Podlasie gran-ulite belt (sites 32 and 34 in Fig. 2a), in order to assess the ageof the collision between Fennoscandia and Sarmatia-related ter-ranes.

3. Metasedimentary rocks and metagranitoids in the West Lithua-nian domain (sites 44, 46 and 29 in Fig. 2a, respectively) for amore reliable definition of its boundary and to date the depo-sition of its sedimentary rocks, their metamorphism and theirsources.

3.1. U–Pb SIMS datings

The U–Pb systematics of magmatic and detrital zircon from thestudied rocks were investigated using the Cameca ims1280 SIMSinstrument of the Nordic high-resolution ion-microprobe facility(NORDSIM) at the Swedish Museum of Natural History in Stock-holm. The analytical procedures closely followed those describedby Whitehouse et al. (Whitehouse et al., 1999; Whitehouse andKamber, 2005). The U/Pb ratios were calibrated against a 1065 Mareference zircon (Geostandards 91500; Wiedenbeck et al., 1995),which was analyzed repeatedly during each session. Reduction ofmeasured ratios was made using software developed at the NORD-SIM laboratory. Where the amount of 204Pb indicated the likelypresence of common Pb, it was corrected using the present-daycommon lead Pb composition (Stacey and Kramers, 1975) on theassumption that it represents surface contamination. Calculationsof ages were made using the Isoplot 3.75 software (Ludwig, 2012)

with results presented at the 95% confidence level. The errors inthe 207Pb/206Pb ratio were about 0.5%, in the U/Pb isotopic ratiosabout 1% (at the 1� level). Analyses of metamict and resorbedzircon grains and analyses with large errors (>25%), and those

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ith discordances of more than ±10% were excluded from the agealculations. Table 1 summarizes all the obtained results.

.2. Sample descriptions and dating results

.2.1. Strongly mylonitized metasedimentary migmatite to theorth of the Middle Estonian fault zone (sample Valgu 99)

This sample represents a high grade mylonite from sillimanite-arnet-cordierite migmatites (Fig. 3a) that dominates the Valgu9 drilling (site 7 in Fig. 2a), and was sampled from a core depthetween 415 and 504 m. The mineral composition and high Alontents of 17–21% suggest a metasedimentary origin of theseigmatites, which occupy a large area in western Estonia. The sam-

led Valgu migmatites reached a peak temperature of 700 ◦C at kbar and had a retrograde evolution involving intense deforma-ion during near-isobaric cooling down to 480 ◦C at 2 kbar (Kikas,001; Puura et al., 2004).

Zircon is mostly enclosed in feldspars, garnet and biotite. Itsrains are small (50–120 �m), rounded, and spherical to elongatedn shape; 40 spots were analyzed in 38 zircon grains (Table 1;ig. 4a). Eighteen subhedral moderately rounded grains, whichave cores displaying variable internal luminescence and mostlyscillatory zoning (OZ) patterns, have yielded 207Pb/206Pb agesanging between 1.98 and 1.89 Ga with frequency maxima at 1.94nd 1.90 Ga (Table 1, Fig. 5a). Ages of 1945 ± 3 and 1897 ± 2 Ma,btained using the unmix multicomponent test (Ludwig, 2012), areompatible with weighted average ages of 1943 ± 7 (MSWD = 6.6)nd 1897 ± 4 Ma (MSWD = 2.7), respectively. Two magmatic OZores were dated at 1973 ± 5 and 1930 ± 5 Ma. These groups ofetrital magmatic zircon showing low proportions of rounded andbraded grains, the absence of large age scatter and the tightge clusters suggest that the Valgu rock derives predominantlyrom a first generation relatively proximal sediment. A youngest07Pb/206Pb age of 1889 ± 12 Ma stemming from a detrital coreTable 1) bounds the maximum age of sedimentation.

Two small (50–80 �m) rounded featureless zircon grainsnd two CL-dark overgrowths recorded 1842 ± 1 Ma agesFigs. 4a and 5a). The Th/U ratios are low, mostly ranging between.06 and 0.3, which suggests metamorphic reworking. A group ofwo small (50–60 �m) featureless grains and three with mostlyL-dark rims, zones and embayments (Fig. 4a) has extremely lowh/U ratios ranging between 0.004 and 0.03 and displays an ageange between 1819 and 1779 Ma. The weighted average age of798 ± 8 Ma (MSWD = 6.2) is very similar to the 1798 ± 2 Ma unmixulticomponent data age (Fig. 5a) and is interpreted as the age of

second event of metamorphism.

.2.2. Deformed high-T granodioritic granulite in the Southstonian granulite domain (sample Kõnnu 300)

The Kõnnu 300 well penetrates garnet-biotite-ordierite ± orthopyroxene ± sillimanite-bearing granulites inhe South Estonian granulite domain (site 10 in Fig. 2a), whichave been intruded by 1.77 Ga metagabbro-norites (Soesoo et al.,006). The dated granulite, sampled from a core depth of 515 m, haseen derived from a coarse-grained S-type granodiorite, formedy the high-T melting of mixed sources, probably comprisingocks of various compositions, like greywacke, semipelites, andven mafic rocks. This is indicated by positive �Nd (1.9 Ga) of +0.4Puura and Huhma, 1993). The rock consists of fairly melanocraticarts, made up of medium-grained biotite, feldspar, quartz andpaques, alternating with coarser-grained lighter layers consistingf plagioclase, perthitic K-feldspar, quartz, cordierite, orthopyrox-

ne and garnet. All these rocks have been intensely mylonitized,ut the lighter layers still preserve igneous textures and high-Alrthopyroxene (Fig. 3b). Hölttä and Klein (1991), Kikas (2001) anduura et al. (2004) described a series of reactions that had taken

Research 259 (2015) 5–33

place in these rocks during their metamorphism. Peak conditionswere attained at 800–850 ◦C and 5–6 kbar, while retrogression andmylonitization occurred along a near-isobaric P-T path down to530 ◦C and 2 kbar.

The zircon grains are coarse (100–250 �m), slightly rounded,mostly elongated, and display variable internal luminescenceand zoning patterns, most commonly of an oscillatory (OZ) type(Fig. 4b). Thirteen grains were analyzed (Table 1). Two rounded,light, oscillatorily zoned cores yielded the oldest 207Pb/206Pb ages of2064 ± 18 and 1997 ± 12 Ma. Seven other oscillatorily zoned, rarelyblurred cores, in contrast, display a close range of 207Pb/206Pb agesbetween 1.91 and 1.88 Ga (Table 1) with a 1891 ± 3 Ma age obtainedby Ludwig’s Isoplot (Ludwig, 2012) “unmix” test of multiple agedistributions (Fig. 5b), and a weighted mean age of 1891 ± 9 Ma(MSWD = 8.8). These prismatic grains with well-defined OZ andTh/U ratios between 0.3 and 0.6 are magmatic in origin.

An age of 1858 ± 4 Ma (and a mean age of 1857 ± 9 Ma,MSWD = 2.5; Fig. 5b) was obtained from three faintly oscillatorilyzoned (OZ), in places blurred cores of very elongated (L/W = 3/1)grains and a thick CL-dark metamorphic overgrowth on one of theOZ cores (Fig. 4b). This age can be interpreted as a metamorphicone, even though neither the CL-dark overgrowths nor the feature-less cores of the elongated grains have Th/U ratios much differentfrom those of the magmatic 1.89 Ga zircon (Table 1). It may indicatea ca. 1.86 Ga metamorphic overprint event during the overall retro-grade evolution, when zircon growth was related to partial meltingand migmatization (Corfu et al., 2003).

3.2.3. Mafic granulite from the Mid-Lithuanian domain (sampleGl99) and granitic mylonite from the southernmost tip of thatdomain close to its tectonic boundary with the Belarus-Podlasiegranulite belt (sample Gr7)

Sample Gl99 from a log depth of 1217 metres in the drillingwell Geluva 99 (site 32 in Fig. 2a) is a mafic granulite intrudedby veins of plagioclase-rich rock and a 1.45 Ga unfoliated granite(Skridlaite et al., 2007). It is made up of plagioclase, biotite, orthopy-roxene, opaques (magnetite and ilmenite), apatite and zircon. Thestructure, texture (Fig. 3c) and chemical composition are those ofan igneous gabbro-diorite. The rock has been deformed and meta-morphosed in the granulite facies as indicated by the high aluminacontents of its hypersthene (Skridlaite et al., 2007).

The zircon grains are mostly coarse (100–350 �m) and pris-matic, commonly with some well-preserved crystal faces, but thereare also rounded, oval and irregularly shaped grains (Fig. 4c). CL-imaging demonstrated that virtually all the prismatic grains featurevariable euhedral zoning and distinct, thick (20–50 �m) CL-brightovergrowths. A few have well delimited cores (Fig. 4c). The roundedand irregular shapes of the grains with remnants of oscillatory zon-ing (OZ) in their central parts are defined by the overgrowths. Someprismatic and short-prismatic grains have irregular OZ shaped byrecrystallization.

The ages obtained from thirteen points in eight grains plot in twodistinct groups (Table 1, Fig. 5c). Five points from OZ cores indicate arather narrow range of concordant to moderately discordant U–Pbisotopic compositions (Table 1, Fig. 5c) yielding similar radiogenic207Pb/206Pb ages. A concordia age of 1839 ± 15 Ma (MSWD = 1.08),close to the weighted mean age of 1847 ± 9 Ma (MSWD = 1.04), isinterpreted as the age of igneous crystallization. Some magmaticgrains have apparent cores (Fig. 5c) which, within error lim-its, display similar 1.86–1.83 Ga (1857 ± 15; 1855 ± 7; 1853 ± 13;1840 ± 9, 1834 ± 10) ages. These may have been produced duringprotracted magmatic differentiation.

Three thick CL-bright overgrowths on OZ cores, a prismaticrecrystallized grain and a recrystallized part of an OZ core (Fig. 5c)yielded concordant and nearly concordant ages of ca. 1.82–1.79 Ga(1818 ± 8, 1811 ± 7, 1798 ± 3, 1795 ± 8, 1785 ± 5). Rounded and

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Table 1U–Pb isotopic SIMS data for zircons from rocks in the South Baltic region.

Sample/session/spot # [U] (ppm) [Th] (ppm) [Pb] (ppm) Th/U (calc) Th/U (meas) f206%a 206Pb/238U ±s (%) 206Pb/238U ±s 207Pb/206Pb ±s (%) 207Pb/206Pb ±s Disc. % 2s lim.

West Estonian domainValgu 99n1628-1a 677 12 234 0.013 0.017 0.21 0.31022 2.32 1741.8 35.6 0.11021 0.27 1802.8 5.0 0n1628-2 518 44 175 0.070 0.084 0.05 0.29746 2.31 1678.7 34.2 0.11377 0.30 1860.5 5.4 −7n1628-2a 846 63 310 0.071 0.074 0.02 0.32260 2.30 1802.4 36.3 0.11283 0.26 1845.5 4.7 0n1628-2b 244 101 102 0.393 0.416 0.06 0.33877 2.33 1880.7 38.2 0.11759 0.39 1919.9 6.9 0n1628-3a 397 303 169 0.674 0.762 0.02 0.31960 2.35 1787.8 36.7 0.11831 0.30 1930.8 5.4 −4n1628-3b 829 48 288 0.058 0.058 0.02 0.30754 2.32 1728.6 35.2 0.11165 0.26 1826.5 4.6 −2n1628-3c 557 166 191 0.224 0.297 0.05 0.28878 2.31 1635.5 33.5 0.11399 0.26 1864.0 4.7 −10n1628-5 464 84 166 0.150 0.181 0.03 0.30828 2.32 1732.2 35.4 0.11268 0.38 1843.0 6.9 −2n1628-5a 152 60 59 0.335 0.397 0.36 0.31738 2.36 1776.9 36.7 0.11918 0.53 1943.9 9.5 −5n1628-6 429 67 174 0.155 0.156 0.17 0.34681 2.33 1919.4 38.8 0.12146 0.33 1977.7 5.9 0n1628-7a 1158 331 379 0.247 0.286 0.03 0.27100 2.31 1545.9 31.9 0.11630 0.19 1900.1 3.4 −17n1628-6a 532 187 204 0.313 0.351 0.02 0.31612 2.31 1770.7 35.9 0.11586 0.30 1893.3 5.4 −3n1628-6b 513 192 212 0.345 0.374 0.07 0.33787 2.31 1876.4 37.6 0.11882 0.28 1938.5 4.9 0n1628-6c 169 93 74 0.503 0.552 0.11 0.34202 2.35 1896.4 38.7 0.12043 0.46 1962.6 8.2 0n1628-7 474 189 167 0.336 0.398 0.01 0.28704 2.32 1626.7 33.4 0.11250 0.33 1840.2 6.0 −9n1628-8 710 68 246 0.086 0.095 0.04 0.30490 2.30 1715.6 34.8 0.10940 0.24 1789.5 4.4 0n1628-9 219 56 85 0.242 0.256 0.19 0.32835 2.37 1830.4 37.9 0.11627 0.43 1899.6 7.7 0n1628-9a 506 127 196 0.241 0.252 0.2 0.32602 2.32 1819.1 36.8 0.11635 0.30 1900.8 5.4 0n1628-11 644 43 236 0.059 0.066 0.05 0.32278 2.31 1803.3 36.4 0.11594 0.23 1894.5 4.2 −1n1628-11a 1124 13 418 0.012 0.011 0.03 0.33419 2.30 1858.6 37.3 0.11002 0.25 1799.7 4.5 0n1628-11b 414 183 177 0.404 0.442 0.15 0.34619 2.31 1916.4 38.4 0.11766 0.32 1921.0 5.7 0n1628-10 315 202 120 0.498 0.644 0.26 0.29448 2.37 1663.9 34.9 0.11961 0.36 1950.4 6.4 −12n1628-10a 505 160 192 0.274 0.316 4.48 0.31672 2.31 1773.7 35.9 0.11548 1.07 1887.4 19.1 0n1628-12 307 138 122 0.377 0.449 0.08 0.32411 2.32 1809.8 36.7 0.11615 0.34 1897.9 6.2 −1n1628-12a 299 182 132 0.562 0.610 0.15 0.34279 2.33 1900.1 38.4 0.11930 0.42 1945.8 7.5 0n1628-12b 464 11 159 0.021 0.024 0.05 0.30625 2.32 1722.2 35.1 0.11032 0.32 1804.7 5.8 −1n1628-13 716 246 256 0.294 0.344 0.28 0.29476 2.32 1665.3 34.2 0.11556 0.32 1888.7 5.7 −9n1628-13b 781 30 239 0.039 0.038 2.2 0.27186 2.83 1550.2 39.2 0.10989 0.66 1797.6 12.0 −10n1628-13a 520 15 173 0.030 0.030 2.45 0.29737 2.31 1678.3 34.2 0.11125 0.48 1819.9 8.7 −4n1628-14 1133 35 430 0.033 0.031 0.03 0.33838 2.30 1878.9 37.7 0.11123 0.20 1819.6 3.6 0n1628-14a 807 223 311 0.261 0.276 0.11 0.32325 2.31 1805.6 36.5 0.11333 0.22 1853.4 4.1 0n1628-14b 266 124 101 0.365 0.466 0.25 0.30664 2.35 1724.1 35.7 0.11667 0.42 1905.9 7.5 −6n1628-15 451 210 188 0.453 0.466 0.2 0.33563 2.33 1865.6 37.8 0.11236 0.33 1837.9 5.9 0n1628-15a 460 110 183 0.225 0.239 0.27 0.33725 2.32 1873.4 37.8 0.11634 0.42 1900.7 7.5 0n1628-16 602 192 262 0.301 0.319 0.32 0.36258 2.32 1994.4 39.9 0.11784 0.27 1923.7 4.8 0n1628-16a 1360 5 490 0.003 0.004 0.02 0.32420 2.30 1810.2 36.4 0.11076 0.20 1812.0 3.5 0n1628-16b 488 203 192 0.366 0.415 0.25 0.31850 2.32 1782.4 36.3 0.11820 0.34 1929.2 6.1 −4n1628-18 204 67 85 0.303 0.330 0.12 0.34367 2.34 1904.3 38.8 0.11926 0.45 1945.2 8.1 0n1628-19 516 45 191 0.087 0.088 0.05 0.32579 2.31 1817.9 36.7 0.10849 0.28 1774.2 5.1 0

South Estonian Granulite beltKõnnu 300n1627-1 197 188 101 0.895 0.950 0.06 0.36698 2.34 2015.2 40.6 0.12751 0.53 2063.9 9.3 0n1627-2 648 187 259 0.271 0.288 0.01 0.33482 2.31 1861.7 37.4 0.11620 0.27 1898.6 4.8 0n1627-3 352 222 151 0.584 0.630 0.08 0.33408 2.32 1858.1 37.5 0.11573 0.34 1891.3 6.1 0n1627-4 491 204 195 0.389 0.414 0.03 0.32258 2.32 1802.3 36.5 0.11324 0.42 1852.0 7.6 0n1627-5 271 82 87 0.221 0.304 0.61 0.27091 2.42 1545.4 33.4 0.11278 0.75 1844.6 13.5 −13n1627-6 353 175 158 0.470 0.497 0.02 0.35515 2.33 1959.1 39.4 0.12278 0.36 1997.0 6.3 0n1627-7 426 129 168 0.281 0.304 0.02 0.32825 2.32 1829.9 37.1 0.11628 0.31 1899.8 5.5 0n1627-8 324 45 124 0.127 0.138 0.12 0.33168 2.33 1846.5 37.5 0.11553 0.37 1888.1 6.6 0

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Table 1 (Continued)

Sample/session/spot # [U] (ppm) [Th] (ppm) [Pb] (ppm) Th/U (calc) Th/U (meas) f206%a 206Pb/238U ±s (%) 206Pb/238U ±s 207Pb/206Pb ±s (%) 207Pb/206Pb ±s Disc. % 2s lim.

n1627-9 2046 445 588 0.162 0.217 0.11 0.24432 2.30 1409.1 29.1 0.11243 0.17 1839.1 3.1 −23n1627-12 597 148 222 0.226 0.248 0.11 0.31447 2.32 1762.7 35.9 0.11369 0.29 1859.2 5.2 −1n1627-13 413 129 150 0.267 0.314 0.03 0.30359 2.32 1709.1 34.9 0.11380 0.34 1861.0 6.1 −5n1627-15 236 74 98 0.304 0.312 0.09 0.34454 3.03 1908.5 50.3 0.11487 0.41 1877.8 7.4 0n1627-16 189 55 75 0.278 0.292 0.04 0.33331 2.33 1854.4 37.7 0.11483 0.46 1877.2 8.3 0n1627-17 284 132 87 0.292 0.463 0.04 0.24969 3.14 1436.9 40.6 0.11709 0.38 1912.3 6.8 −23

Mid-Lithuanian domainGeluva 99 (Gl99)n1261-4a 3042 990 1031 0.270 0.326 0.08 0.28275 1.46 1605.2 20.8 0.10990 0.16 1797.8 2.8 −9n1261-5a 340 80 134 0.232 0.236 0.03 0.33628 1.50 1868.8 24.4 0.11070 0.41 1811.0 7.3 0n1261-5b 124 63 50 0.482 0.506 0.03 0.32187 1.47 1798.9 23.1 0.11355 0.81 1857.0 14.6 0n1261-5c 271 119 106 0.416 0.441 0.08 0.31812 1.47 1780.5 22.9 0.11214 0.54 1834.4 9.8 0n1261-9a 214 88 86 0.397 0.411 0.02 0.32884 1.46 1832.8 23.3 0.11246 0.49 1839.5 8.9 0n1261-9b 236 105 93 0.424 0.446 0.03 0.31929 1.46 1786.3 22.8 0.11110 0.47 1817.6 8.4 0n1261-10 341 117 128 0.310 0.344 0.07 0.31328 1.46 1756.8 22.5 0.10971 0.42 1794.6 7.6 0n1261-19 1068 110 318 0.086 0.103 0.1 0.26111 1.49 1495.5 19.9 0.10915 0.30 1785.2 5.4 −15n1261-21 328 86 124 0.251 0.263 0.02 0.31952 1.46 1787.4 22.8 0.11340 0.40 1854.6 7.3 −1n1261-27a 475 141 162 0.255 0.297 0.02 0.28448 1.46 1613.9 20.9 0.11164 0.34 1826.4 6.2 −10n1261-27b 155 82 61 0.492 0.530 0.06 0.30857 1.49 1733.7 22.7 0.11329 0.70 1852.9 12.5 −3n1261-28a 69 47 27 0.658 0.682 0.03 0.30025 1.47 1692.5 22.0 0.10638 0.98 1738.3 17.9 0n1261-28b 298 73 87 0.209 0.245 0.17 0.24805 1.46 1428.4 18.7 0.10295 0.55 1677.9 10.2 −13Grodno 7 (Gr7)n2192-1 207 173 95 0.868 0.836 {0.03} 0.3377 0.88 1875.7 14.4 0.11260 0.42 1841.8 7.6 0n2192-2a 393 67 153 0.175 0.171 0.17 0.3338 0.88 1856.6 14.2 0.11254 0.34 1840.9 6.2 0n2192-2b 463 75 178 0.164 0.163 0.03 0.3327 0.88 1851.6 14.1 0.11245 0.29 1839.4 5.2 0n2192-3-1b 213 29 81 0.137 0.135 {0.01} 0.3320 0.88 1847.9 14.1 0.11172 0.42 1827.6 7.5 0n2192-3-1a 198 24 74 0.125 0.122 {0.02} 0.3262 0.88 1819.9 14.0 0.11102 0.44 1816.2 7.9 0n2192-3-2 127 57 51 0.431 0.447 0.06 0.3221 0.89 1800.1 14.0 0.11292 0.55 1846.9 10.0 0n2192-5a 112 36 45 0.330 0.321 {0.05} 0.3288 0.88 1832.5 14.1 0.11379 0.57 1860.7 10.3 0n2192-5b 196 39 76 0.206 0.200 {0.03} 0.3307 0.88 1841.7 14.2 0.11152 0.44 1824.3 7.9 0n2192-6-1 145 43 57 0.292 0.296 {0.02} 0.3281 0.89 1829.1 14.1 0.11216 0.50 1834.7 9.1 0n2192-6-2 206 47 81 0.226 0.228 {0.03} 0.3342 0.88 1858.7 14.3 0.11174 0.49 1827.9 8.8 0n2192-7 622 408 277 0.691 0.657 0.06 0.3410 0.88 1891.6 14.4 0.11233 0.25 1837.4 4.5 1n2192-8 130 46 53 0.351 0.355 {0.03} 0.3336 0.89 1855.6 14.3 0.11263 0.53 1842.3 9.5 0n2192-10-1a 235 41 92 0.181 0.175 {0.01} 0.3377 0.88 1875.6 14.4 0.11202 0.41 1832.5 7.4 0n2192-10-1b 189 39 75 0.210 0.203 {0.02} 0.3357 0.89 1865.7 14.5 0.11326 0.44 1852.3 7.8 0n2192-10-2 289 30 112 0.104 0.104 0.03 0.3382 0.88 1878.0 14.3 0.11251 0.38 1840.4 6.8 0

West Lithuanian Granulite domainLauksargiai (Lk2)n1260-1 967 659 445 0.675 0.682 0.11 0.35057 1.48 1937.3 24.7 0.11845 0.23 1933.0 4.0 0n1260-13 114 57 49 0.478 0.499 {0.02} 0.34253 1.49 1898.8 24.6 0.11947 0.84 1948.3 15.0 0n1260-15 47 66 22 1.287 1.383 {0.09} 0.30447 1.46 1713.4 22.0 0.11012 1.17 1801.3 21.1 0n1260-24 367 151 159 0.391 0.412 {0.01} 0.34904 1.48 1930.0 24.7 0.12227 0.34 1989.6 6.1 0n1260-25b 402 108 164 0.281 0.270 0.07 0.33809 1.46 1877.5 23.9 0.12048 0.37 1963.4 6.6 −2n1260-33 121 59 52 0.461 0.487 {0.00} 0.33954 1.46 1884.5 23.9 0.11713 0.70 1912.9 12.5 0n1260-34 282 224 126 0.737 0.794 {0.02} 0.33357 1.46 1855.7 23.6 0.11878 0.41 1938.0 7.3 −2n1260-39 1191 769 553 0.597 0.646 {0.01} 0.35685 1.49 1967.2 25.3 0.12463 0.56 2023.6 9.9 0n1260-41 45 28 28 0.603 0.623 {0.07} 0.47316 1.46 2497.4 30.3 0.16847 1.05 2542.5 17.5 0

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15

Pociai 3 (Pc3)n2033-2 150 65 68 0.438 0.433 {0.02} 0.3645 1.51 2003.5 26.1 0.12392 0.71 2013.5 12.5 0n2033-4-1 310 214 156 0.961 0.689 0.08 0.3633 1.51 1997.6 26.0 0.11938 0.81 1946.9 14.4 0n2033-4-2 799 1031 405 1.278 1.290 0.03 0.3403 1.51 1888.1 24.8 0.11717 0.38 1913.4 6.8 0n2033-5a 767 313 615 0.379 0.409 0.01 0.5849 1.51 2968.9 36.0 0.25782 0.24 3233.4 3.8 −7n2033-5b 176 92 150 0.515 0.525 {0.01} 0.5962 1.55 3014.5 37.4 0.27607 0.48 3340.9 7.4 −9n2033-7a 371 163 155 0.432 0.439 {0.01} 0.3369 1.51 1871.8 24.6 0.11415 0.49 1866.5 8.8 0n2033-7b 2031 674 399 0.151 0.332 0.23 0.1676 1.51 999.0 14.0 0.09817 0.39 1589.7 7.3 −37n2033-8 230 92 93 0.393 0.399 {0.00} 0.3292 1.51 1834.6 24.1 0.11552 0.64 1888.0 11.4 0n2033-10a 473 11 168 0.026 0.023 {0.01} 0.3159 1.51 1769.8 23.4 0.11530 0.45 1884.6 8.0 −4n2033-10b 233 179 198 0.805 0.768 0.02 0.5957 1.52 3012.3 36.6 0.21423 0.46 2937.8 7.4 0n2033-11 353 83 140 0.243 0.236 {0.01} 0.3336 1.51 1855.7 24.3 0.11452 0.49 1872.4 8.8 0n2033-12a 5156 1576 352 0.121 0.306 2.47 0.0593 1.51 371.7 5.5 0.06601 0.84 806.6 17.6 −48n2033-12b 3248 464 355 0.061 0.143 2.37 0.0973 1.52 598.6 8.7 0.08267 0.82 1261.5 15.9 −50n2033-14 765 68 252 0.076 0.089 0.17 0.2901 1.51 1641.9 22.0 0.11313 0.37 1850.3 6.6 −10n2033-15 104 77 49 0.745 0.743 0.12 0.3530 1.51 1948.8 25.5 0.12145 0.87 1977.6 15.4 0n2033-16-1 165 66 112 0.378 0.398 {0.00} 0.5179 1.51 2690.5 33.3 0.20007 0.66 2826.7 10.7 −2n2033-16-2 268 166 182 0.609 0.620 {0.01} 0.5006 1.51 2616.6 32.5 0.18542 0.48 2702.0 8.0 0n2033-17-1 134 14 88 0.104 0.101 {0.01} 0.5345 1.51 2760.6 34.1 0.20192 0.51 2841.8 8.2 0n2033-17-2 205 65 94 0.326 0.317 {0.00} 0.3786 1.51 2069.9 26.8 0.12376 0.62 2011.2 10.9 0n2033-21a 306 204 145 0.669 0.666 {0.01} 0.3580 1.51 1972.6 25.7 0.12432 0.49 2019.1 8.7 0n2033-21b 254 331 140 1.334 1.301 {0.02} 0.3636 1.51 1999.3 26.0 0.12543 0.52 2034.8 9.3 0n2033-22-1 850 153 286 0.174 0.180 0.22 0.2863 1.53 1623.1 21.9 0.11536 0.42 1885.5 7.6 −13n2033-22-2 95 53 41 0.568 0.559 {0.00} 0.3376 1.51 1875.1 24.6 0.11634 1.14 1900.7 20.4 0

Lazdijai 8 (Lz8)n2032-1b 381 93 113 0.273 0.244 1.54 0.2432 1.51 1403.3 19.0 0.11394 0.82 1863.2 14.8 −23n2032-4 425 157 173 0.362 0.370 0.08 0.3318 1.51 1846.9 24.3 0.11698 0.48 1910.6 8.7 0n2032-5 920 388 380 0.412 0.421 0.07 0.3340 1.51 1857.6 24.5 0.11545 0.45 1887.0 8.0 0n2032-6a 420 132 159 0.297 0.314 0.37 0.3145 1.51 1763.0 23.3 0.11526 0.51 1884.0 9.2 −4n2032-6b 2490 889 1016 0.346 0.357 0.18 0.3357 1.53 1866.1 24.9 0.11521 0.30 1883.2 5.4 0n2032-6-2 243 83 84 0.322 0.341 0.77 0.2826 1.52 1604.6 21.7 0.11225 0.79 1836.2 14.3 −10n2032-9 347 146 147 0.423 0.421 {0.01} 0.3433 1.51 1902.6 24.9 0.11505 0.55 1880.7 10.0 0n2032-10a 348 117 102 0.214 0.336 0.55 0.2434 1.55 1404.4 19.6 0.11361 0.95 1858.0 17.0 −22n2032-10b 446 140 117 0.218 0.315 0.10 0.2167 1.51 1264.7 17.4 0.11542 0.54 1886.4 9.7 −33n2032-12a 378 149 157 0.378 0.396 {0.01} 0.3406 1.51 1889.8 24.8 0.11534 0.46 1885.3 8.3 0n2032-12b 272 82 112 0.295 0.302 {0.01} 0.3446 1.51 1908.8 25.0 0.11501 0.71 1880.0 12.7 0n2032-13 558 192 162 0.304 0.345 0.39 0.2346 1.51 1358.8 18.6 0.11304 0.61 1848.8 11.0 −26n2032-14-1 498 281 164 0.387 0.563 0.21 0.2627 1.53 1503.6 20.5 0.11200 0.49 1832.2 8.9 −17n2032-14-2 786 210 311 0.266 0.267 0.06 0.3316 1.51 1846.2 24.2 0.11436 0.48 1869.9 8.6 0n2032-16-1 178 48 72 0.261 0.271 0.05 0.3382 1.51 1878.0 24.7 0.11543 0.88 1886.6 15.8 0n2032-16-2 991 366 333 0.311 0.369 0.21 0.2746 1.53 1564.2 21.3 0.11449 0.33 1871.9 6.0 −16n2032-17 485 187 205 0.407 0.386 0.02 0.3413 1.51 1893.0 24.8 0.11658 0.28 1904.4 5.0 0n2032-20 325 101 119 0.298 0.312 0.10 0.3007 1.64 1695.0 24.4 0.11586 0.40 1893.3 7.1 −9n2032-21 262 85 94 0.309 0.324 0.24 0.2949 1.52 1666.2 22.4 0.11540 0.53 1886.2 9.5 −10n2032-24a 304 73 111 0.225 0.239 0.03 0.3065 1.51 1723.3 22.9 0.11551 0.42 1887.9 7.6 −7n2032-24b 424 129 140 0.341 0.304 1.37 0.2665 1.58 1523.3 21.4 0.11561 2.31 1889.5 40.9 −11n2032-25 417 62 167 0.155 0.148 {0.00} 0.3452 1.54 1911.6 25.5 0.11562 0.31 1889.5 5.6 0

a % of common 206Pb.

16 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Fig. 3. Microphotographs of dated rocks. (a) High-grade banded, fine-grained mylonite from the metasedimentary migmatite of sample Valgu 99, NW Estonian domain. Thedark bands consist of fine-grained intergrowths of elongated garnet grains, cordierite and sillimanite with or without biotite flakes, while the white quartz-feldspar leucosomeis rich in biotite fine laminae along mylonitic foliation. (b) Coarse-grained, garnet-biotite-cordierite-orthopyroxene-bearing metagranodiorite from sample Kõnnu 300, SouthEstonian domain. (c) Meta-igneous mafic granulite from sample Gl99, Mid-Lithuanian domain. (d) Low grade granite mylonite from sample Gr7 with porphyroclastic garnet(grey) and winged feldspar porphyroclasts (white and light grey) set in a fine-grained biotite-quartz-feldspar matrix with characteristic C′-type shear bands, Grodno-Bialystok deformation zone, Eastern Mid-Lithuanian domain. (e) Lens-spotted cordierite-biotite-garnet-sillimanite-bearing granulite from sample Lk2, West Lithuaniangranulite domain. The dark restitic layers and lenses are made up of opaques (ilmenite + magnetite + spinel ± rutile), biotite, garnet, sillimanite and cordierite, while the whiteleucocratic lensoid spots mostly represent granitic melt with garnet and, in places, cordierite. (f) High grade mylonite from cordierite-biotite-garnet-sillimanite-bearinggranulite of sample Pc3. The porphyroclasts of garnet are set in a mylonite matrix of alternating garnet-sillimanite-biotite-opaques with cordierite and quartz-cordierite-feldspars plus biotite bands, West Lithuanian Granulite domain. (g) Porphyritic garnet-biotite-amphibole metadiorite from sample Lz8, West Lithuanian granulite domain.Note the brown biotite and pink skeleton garnet growing at the expense of hornblende (green), opaques (black) and plagioclase (white).

S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33 17

F nu 30m magesa

io(aig(syt

eBapbmt3

ei(wammb

ig. 4. Cathodoluminescence images of the dated zircon from (a) Valgu 99, (b) Kõnicroprobe spot analyses are indicated by ellipses traced from secondary electron i

t 1�.

rregular shapes of the overgrowths and recrystallization of thescillatory zoning suggest metamorphic origins of these zirconsFig. 5c). The upper intercept at 1809 ± 9 Ma (MSWD = 1.3) and

similar weighted mean age of 1798 ± 11 Ma (MSWD = 3.6) arenterpreted to indicate the age of metamorphism. The over-rowths are richer in uranium (up to 3042 ppm) than the cores124–271 ppm), and have lower Th/U ratios (Table 1). One analy-is of a highly metamictic grain with preserved oscillatory zoningielded 1738 ± 18 Ma, however this age may be ambiguous due tohe high degree of recrystallization.

Sample Gr7 (Grodno 7 core, depth 455 m) is from the south-rn tip of the Mid-Lithuanian domain at its boundary with thePG (site 34, Fig. 2a). This is a low grade mylonite formed from

coarse-grained biotite-garnet-Kfeldspar rich granite with theorphyroclastic garnet and the K-feldspar set in a fine-grainediotite-quartz-feldspar matrix (Fig. 3d). Peak conditions of meta-orphism were 725–730 ◦C at 7–8 kbar, while mylonitization

ook place during a near-isobaric retrogression to 575–625 ◦C at–5 kbar (Skridlaite et al., 2014).

The zircon grains are coarse (150–300 �m), euhedral, highlylongated (L/W = 3/1) and commonly display oscillatory zon-ng (Fig. 4d). Fifteen analyses were performed on eleven grainsTable 1). A concordia age of 1844 ± 8 Ma (MSWD = 5.5; Fig. 6d)as calculated from 9 points. It is very similar to the weighted

verage 207Pb/206Pb age of 1841 ± 4 Ma (MSWD = 5.6). The obtainedagmatic age likely represents the age of anatectic melting ofetasedimentary rocks. Most of the zircon grains are surrounded

y extremely thin (<5 m) CL-light rims, which at present lack a

0, (c) Gl99, (d) 7Gr, (e) Lk2, (f) Pc3, and (g) Lz8 samples. The locations of the ion- taken after the analyses. The obtained 207Pb/206Pb ages are given in Ma. Errors are

straightforward explanation. Possibly, they grew during the lastphase of crystallization or during following metamorphism. TwoCL-dark recrystallized areas (Fig. 4d) 1816 ± 8 and 1828 ± 8 Main age (Table 1) with homogenization of the magmatic oscilla-tory zoning and decreasing U and Th contents originated fromlater metamorphic reworking, which most probably accompaniedthe mylonitization. Also the monazite in this rock has domains1.85–1.82 Ga in age (Skridlaite et al., 2014).

3.2.4. Metasedimentary rocks from the West Lithuanian granulitedomain (samples Lk2 and Pc3)

The two samples of metasedimentary rock, Lk2 and Pc3, wereobtained from the cores of the Lauksargiai and the Pociai drillingsin western Lithuania (sites 44 and 46 in Fig. 2a).

Sample Lk2 is from a deformed migmatitic metapelitic gran-ulite at a core depth of 1883 m, where interbeds of fine-grainedmetagreywacke-like biotite-plagioclase gneisses are also present.The metapelite features leucocratic and melanocratic lens-shapedslivers and layers (Fig. 3e), its melanosome mostly consisting ofcordierite, biotite, sillimanite, garnet and complex opaques (spinel,magnetite, ilmenite, rutile), while the leucocratic domains containgarnet porphyroblasts set in a matrix of K-feldspar-rich graniticmaterials. The rock is rich in zircon and monazite. The peak meta-morphic conditions were 780 ◦C at ca. 8 kbar, as recorded by

garnet-biotite-plagioclase assemblages in both the leucocratic andthe melanocratic rock lenses.

Sample Pc3 from a depth of 2154 m has a similar compositionbut is strongly sheared. In the foliated matrix, new skeleton garnet

18 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Fig. 5. Concordia and distribution probability diagrams for the dated zircon.

S. Bogdanova et al. / Precambrian

0

4

8

12

16

20

1800 2000 2200 2400 2600 2800 3000 3200 3400 3600

Re

lative

pro

ba

bili

ty

Nu

mb

er

WLGdomainn=58

1935

19881894

1853

2075

2038

0

4

8

12

16

20

24

28

32

36

40

1800 2000 2200 2400 2600 2800 3000 3200 3400 3600

Re

lative

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ba

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ty

Nu

mb

er

207Pb/206Pb age, Ma

MDdomainn=99

1984

2034

18602690

0

2

4

6

8

10

1800 2000 2200 2400 2600 2800 3000 3200 3400 3600

Re

lative

pro

ba

bili

ty

Nu

mb

er

WE and SE Gdomains

n=36

1898

1948

1976

Fig. 6. Relative probability plots of concordant and nearly concordant 207Pb/206Pbdetrital zircon ages for the sedimentary rocks from the West Estonian (WE) domain,ta

a(t

gor(sada

ctg

4.1.1. Palaeoproterozoic magmatism

he West Lithuanian granulite (WLG) domain (Krzeminska et al., 2009; this study)nd the Mazowsze (MD) domain (Williams et al., 2009).

nd K-feldspar were formed at the expense of biotite and cordieriteFig. 3f). The intensely retrogressed rock does not record tempera-ures higher than 630 ◦C at 5 kbar (Skridlaite et al., 2014).

The zircon from the Lk2 metapelite features a wide range ofrain sizes (60–200 �m) and variable morphologies in the shapef sub-euhedral prismatic grains with rounded terminations, well-ounded grains, and grain fragments (Fig. 4e). The U contents47–1191 ppm, Table 1) and the Th/U ratios (0.24–1.38) vary con-iderably. Two of the recorded ages are Archaean, the remaining 11re Palaeoproterozoic (Table 1). The Archaean analyses are slightlyiscordant (Fig. 5e), with individual 207Pb/206Pb ages of ca. 2.67 Gand 2.55 Ga.

Most of the Palaeoproterozoic zircons yield concordant or near

oncordant analyses (Table 1, Fig. 5e). The 207Pb/206Pb ages scat-er between ca. 2.03 and 1.90 Ga, the two OZ cores in subhedralrains yielding similar, ca. 2.03 Ga 207Pb/206Pb ages. A thick CL-light

Research 259 (2015) 5–33 19

grey rim around a complex OZ grain fragment, a grey pyramidaltermination of a ca. 2.03 Ga core and a broken, recrystallized OZgrain all recorded ages of 1989 ± 7 and 1968 ± 6 Ma (unmix mul-ticomponent test, Fig. 5e). A group of four subhedral OZ and onerounded grain with an irregular internal zoning has 1939 ± 3 Maages (Fig. 5e), similar to the weighted mean 207Pb/206Pb age of1940 ± 12 Ma (MSWD = 8.4). One euhedral CL-light OZ grain yieldeda slightly younger, ca. 1.9 Ga age. Many grains are overgrown by thin(<10 �m) CL-dark rims.

A rounded CL-light sector-zoned zircon with a thin grey rimgave an 1801 ± 21 Ma age (Fig. 4e). Such zircons are characteris-tic of high-grade metamorphic rocks, particularly granulites (Corfuet al., 2003), thus the grain may have grown during peak metamor-phism. However, it has an extremely high Th/U ratio (1.4) causedby a very low uranium content (47 ppm).

The similar but more strongly deformed metapelites of thePociai 3 drill core (Fig. 4f) yielded a 2.94–2.70 Ga group of nearconcordant Archaean ages and another of 2.03–1.85 Ga concor-dant and near concordant Palaeproterozoic ones (Table 1). Fourmagmatic cores recorded 2021 ± 5 Ma, two 1980 ± 1 Ma, and three1918 ± 6 Ma ages (unmix multicomponent data diagram, Fig. 5f).The largest group comprises five magmatic cores with a mean ageof 1872 ± 3 (Fig. 5f). The youngest analyzed zircon was dated at1850 ± 7 Ma (207Pb/206Pb age, Table 1), indicating the maximumdeposition age of the sediments. No post-depositional meta-morphic zircon has been identified in this sample. One 1.89 Gametamorphic overgrowth on a 2.94 Ga detrital zircon (Fig. 4f) indi-cates recycling of metamorphic rocks.

3.2.5. A Latvian-East Lithuanian garnet-bearing metadiorite fromthe West Lithuanian granulite domain (WLG, sample Lz8)

The metadiorite sample Lz8 was collected from a depth of 682 min the Lazdijai 8 drilling close to the SE margin of the WLG (site 29 inFig. 2a). This drilling penetrated orthogneisses (metaigneous rocksof granodioritic to dioritic compositions) made up of plagioclase,quartz, amphibole, biotite, garnet, magnetite, apatite and zircon(Fig. 3g). The sample records a peak temperature of 650 ◦C at 9 kbarand retrogression to 500 ◦C at 4 kbar (Skridlaite et al., 2003a). Thezircon grains are coarse (100–200 �m), commonly euhedral, oscil-latorily zoned, and mostly elongated with length/width ratios upto 3/1 even though a few of them are only short prismatic. Severalmagmatic cores are overgrown by thin CL-light featureless meta-morphic rims (Fig. 4g). The prismatic oscillatory zoning appearsslightly distorted by later recrystallization.

Twenty three analyses were performed on eighteen grains(Table 1, Fig. 5g). The upper concordia intercept age is 1887 ± 9 Ma(MSWD = 1.3), which is nearly the same as the weighted average ageof 1887 ± 2 Ma (MSWD = 0.87) recorded by twelve magmatic cores.It is interpreted as the age of igneous emplacement. A 1836 ± 14 Ma207Pb/206Pb age obtained from a prismatic, featureless grain termi-nation (Fig. 4g) possibly indicates metamorphic reworking.

4. Discussion

4.1. Tectonic subdivisions of the Palaeoproterozoic crust in theSouth Baltic region revisited

The newly obtained ages of magmatism, metamorphism anddeformation, and the records of the detrital zircons in the metased-imentary rocks constrain the configurations and boundaries of thevarious tectonic domains in the South Baltic region (Figs. 1 and 2a).

The existence of two major Palaeoproterozoic magmatic eventsin the South Baltic region was confirmed by the age determinationsof the present study. The older event, with an approximate age

2 brian Research 259 (2015) 5–33

oEicLo

g(Tsmsfoo(ahp29atma

mamcfstwb

cnaadctt12imihcNcmS

mztewtcm1

Fig. 7. Tectonic setting discrimination diagram for basins of sedimentation (modi-fied after Cawood et al. (2012)) using the time span between zircon crystallizationand deposition. The basins in the Baltic Shield area and in the South Baltic region(this study) indicate their affinity with convergent tectonic settings. Basins: 1 –

0 S. Bogdanova et al. / Precam

f 1.89 Ga is known from the South Estonian (SEG) and Latvian-ast Lithuanian (LEL) domains, where it played an important rolen the formation of the crust, while our two datings of the younger,a. 1.85–1.84 Ga event were carried out on rocks from the Westithuanian domain (WLG). These have added to numerous datingsf this magmatism by various authors (Appendix A).

In the SEG, S-type garnet-orthopyroxene-cordierite-bearingranodiorites with a magmatic crystallization age of 1891 ± 3 Maour sample Kõnnu 300) originated from melts formed by high-

dry partial melting of mixed sources, probably comprisingupracrustal rocks of various compositions. A large number of mag-atic zircon grains (Fig. 4b) remained well preserved despite the

ubsequent metamorphism at ca. 1.86 Ga. In this sample, we alsoound two inherited grains with ages of 2.1 and 2.0 Ga. A crustalrigin and the involvement of mantle-related mafic source rocksr melts are reflected by positive �Nd values between +0.4 and +2for 1.9 Ga) in these and also other SEG meta-igneous granulitesnd by their 2.1 Ga TDM Nd ages (Puura and Huhma, 1993). Theigh-T mineral composition of the Kõnnu 300 granodiorites, andarticularly the presence of orthopyroxene with 7% Al2O3 (Kikas,001) suggest their initial crystallization from partial melts at ca.00–950 ◦C and 6–7 kbar (Vielzeuf and Montel, 1994; Aranovichnd Berman, 1997). Since there are no indications of a collisionalhickening of the crust, this requires a thermal anomaly in the

iddle-lower crust and extension of a crust of normal thicknesst ca. 1.9 Ga (cf. Harley, 1989).

In the neighbouring LEL domain, some metavolcanic felsic andafic rocks have juvenile �Nd values up to +4 and TDM Nd crust

ges of 2.1–1.9 Ga (Mansfeld, 2001). A fairly high contribution fromantle-related sources was probably facilitated by extension of the

rust. The newly dated garnet-bearing metadiorite (sample Lz8)rom southern Lithuania, which yielded an age of 1887 ± 2 Ma, istill another LEL magmatic rock, but it occurs in a setting of theectono-stratigraphically overlying WLG (cf. Fig. 2b). Since this rockas formed in the lower crust (Skridlaite et al., 2003a), it may have

een uplifted along a shear zone (Fig. 2a).In contrast, the large volumes and wide areas of 1.86–1.84 Ga

alc-alkaline rocks in the MLD suggest settings of an active conti-ental margin (Skridlaite and Motuza, 2001; Motuza, 2005). Thus,n amount of oceanic crust must have been subducted nearlylong the present south-western boundary of the MLD. Our newlyated 1847 ± 9 Ga metagabbro sample from the Geluva 99 drillore resembles the rocks of a number of other intrusions alonghis boundary in the southern and central parts of the MLD. Inhe north-western MLD, however, charnockites with ages between862 and 1815 Ma form large multiphase plutons (Claesson et al.,001; Motuza et al., 2008; Vejelyte et al., 2012). Their melts orig-

nated in the lower crust by high-temperature melting of bothetasedimentary and meta-igneous sources, the rocks being affil-

ated geochemically to peraluminous and metaluminous ferroanighly potassic granitoid series, which commonly associate withontinental rifting and back-arc settings (Frost and Frost, 2008).otably, the charnockitic rocks in the MLD correspond in age andomposition to the 1.87–1.84 Ga Askersund suite and its Loftaham-ar extension in the Västervik domain (VV in Fig. 1) of the Baltic

hield (Motuza and Motuza, 2011).In the southernmost MLD, the 1841 ± 4 Ma S-type granite

ylonite in our Gr7 sample occurs along the Grodno-Bialystokone of shearing and faulting that bounds the MLD againsthe Belarus-Podlasie granulite belt (BPG) in the east. Skridlaitet al. (2014) described this rock as “felsic granulites in theestern BPG”. Its least retrograded parts preserve a high-

emperature mineralogy presumably formed during the igneousrystallization of this granite at 725–730 ◦C and 6–8 kbar. Largeagmatic monazite grains contain domains with ages close to

.85 Ga.

West and South Estonian; 2 – Poceai in the West Lithuanian domain; 3 – Monki inthe Mazowsze domain (Williams et al., 2009); 4 – Västervik (Sultan et al., 2005);5–7 according to Bergman et al. (2008): 5 – Tiirismaa, 6 – Luukola, 7 – Pyhäntaka.

Altogether, the available age determinations and the chem-istry of the magmatism in the MLD characterize this domain asan active continental margin at 1.86–1.84 Ga. As different fromthat, the West Lithuanian and Polish domains feature a volumi-nous younger, ca. 1.83–1.80 Ga, plutonic and volcanic magmatismmarked by a subduction-related geochemistry (Krzeminska et al.,2005; Wiszniewska et al., 2007). Broadly, most of rocks match thedominant ca. 1.81–1.76 Ga (TIB-1) granites in the East Smålanddomain of the Baltic Shield (e.g. Wikman, 1993, 1997; Johanssonet al., 2006).

4.1.2. Palaeoproterozoic detrital zircon records and possibletectonic settings of the sedimentary rocks in the South Balticregion

In the Valgu 99 metapelitic rocks of the West Estonian (WE)domain there are several groups of detrital zircons with ages of1.90, 1.93, 1.95 and 1.97 Ga (Table 1, Figs. 5 and 6), which defines1.90 Ga as the approximate maximum age of deposition of theirdetritus. A lower age limit for the sedimentation in the WE basin isprovided by the 1.84 Ga age of the earliest metamorphism in theserocks. In the adjacent South Estonian domain (SEG), the Kõnnu 300garnet-orthopyroxene granodiorite crystallized at 1.89 Ga and wasdeformed and migmatized at ca. 1.86 Ga. An extensional back-arctectonic scenario during the accretion between 1.89 and 1.87 Gamay explain the deformation and granulite-facies metamorphism,and possibly also the initial generation of the Middle Estonianfault zone (MEFZ), which separates the WE and SEG domains. Thetight ca. 2.0–1.9 Ga grouping of the ages of the detrital zircons inthe Valgu 99 metasedimentary rocks suggests source areas fairlyclose to the West Estonian basin of sedimentation. One possiblearea is the SEG and another a magmatic arc in the north such ase.g. the 1.90–1.88 Ga Tampere belt in Finland (Kähkönen, 2005;Lahtinen et al., 2009a). Final deposition could have taken place ina convergent short lived intra-arc, back-arc or fore-arc basin (cf.Figs. 6 and 7). A previous TIMS dating of a multigrain zircon sep-arate from a metasedimentary granulite in the Varbla F-502 drillcore in the SEG (Petersell and Levchenkov, 1994) yielded a similar

range of detrital zircon ages between 2.0 and 1.96 Ga. Inherited zir-con grains with ages between 2.1 and 2.0 Ga are also present in theKõnnu 300 metagranodiorite. However, from the drill core mate-rials available at present, it appears impossible to assess whether

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here existed a system of numerous small depositional basins inter-persed between magmatic arcs throughout the entire SEG.

Farther south, in the West Lithuanian domain (WLG), the twoamples of metasedimentary rocks in our study have added sub-tantially to the knowledge of the basins of sedimentation inouthern Fennoscandia. They derive from an area of predomi-antly metasedimentary, mostly metapelitic, rocks, alternating inhe drill core sections with dominant metabasaltic and sporadic

eta-andesitic intercalations (the Pociai basin, Pc in Figs. 1 and 2a).ecurrent metamorphism reached conditions of the high amphi-olite and granulite facies (Skridlaite et al., 2014). The chemicalharacteristics of the rocks in this sequence refer them to oceanicettings such as oceanic arcs or plateaus where clays associate withrimitive mafic volcanics (Skridlaite and Motuza, 2001). TDM Ndrustal ages of 2.4–2.3 Ga and negative �Nd values (Appendix A) asell as the wide ranges of detrital zircon ages (Figs. 6 and 7; Table 1)

uggest, however, that a continental block had existed in the vicin-ty. Among the detrital zircon populations, a 2.1–1.94 Ga group wasbtained from the Lk2 sample, while the Pc3 metapelite sampleontained a 1872 ± 3 Ma cluster, their youngest 1850 ± 7 Ma grainsefining the maximum age of deposition of the WLG sediments.he zircon and monazite ages indicate that the earliest metamor-hism took place between 1.85 and 1.81 Ga (Krzeminska et al.,009; Skridlaite et al., 2014).

Similar populations of detrital zircon were found in sam-les of the metasedimentary rock in the Bliudziai 150 core fromhe WLG, and in the Mazowsze domain (MD) farther south inoland (Krzeminska et al., 2009; Williams et al., 2009). Theetected two youngest zircon grains have ages of 1870 ± 20 Ma and830 ± 35 Ma, respectively, and constrain the oldest possible depo-ition ages of the Polish sediments to ca. 1.86–1.83 Ga, which withinrror limits is nearly the age of sedimentation in the WLG. Meta-orphism at 1824 ± 14 Ma occurred in the MD close to the ages of

orresponding events in the WLG. Despite the similar ages of sedi-entation, its tectonic settings appear different in the WLG and theD (Fig. 2a). While the Monki basin (Mk) sediments were deposited

long an active continental margin or in a continental volcanic arcn the MD (Williams et al., 2009), the WLG Poceai basin (Pc) had anceanic affinity.

The deposition of the WLG and the MD sediments in convergent-ype settings appears to be mirrored by the time differencesetween the crystallization ages of the zircon and the ages ofediment deposition (Fig. 7), which according to Cawood ando-authors (2012) can be used to discriminate between variousectonic settings of sedimentary basins. In this regard, some differ-nces exist between the sediments in the WLG and the MD, whichs also true of the relative abundances of zircon populations of var-ous ages (Fig. 6). A distinct frequency peak at 1.9 Ga is seen in the

LG but not in the MD, which may indicate detritus supply fromortherly sources comparable to those for Estonia in the formerase but not in the latter (Figs. 6–8). However, the WLG and MDediments were deposited at 1.86–1.83 Ga, which is substantiallyater than the 1.89–1.86 Ga sedimentation in the WE and SEG.

Another important signature of the WLG and MD sedimen-ary rocks is the presence of Archaean zircons of various agesanging from 3.5 to 2.6 Ga, while similar deposits in the Estonianomains almost wholly lack Archaean zircon (Fig. 6). This suggestsn Archaean input into the Polish and Lithuanian basins from a con-inental block somewhere in the present south. Sarmatia, whichontains large domains of Archaean crust (Claesson et al., 2006),ould appear to be an obvious choice and was favoured as such byilliams et al. (2009), who suggested sedimentation in a system of

arginal basins even extending as far north as Finland and Sweden.

problem herewith, however, is that the Sarmatian plate includinghe Sarmatian proto-continent, the Osnitsk-Mikashevichi igneouselt along its active margin, the Okolovo oceanic terrane, and an

Research 259 (2015) 5–33 21

exotic mature arc in the shape of the former Belarus-Podlasie gran-ulite belt, appears to have been separated from the Svecofennianterranes by an ocean throughout the Palaeoproterozoic until theFenno-Sarmatian collision at 1.82–1.80 Ga.

Still, it can be argued that some protruding part of (Volgo-)Sarmatia could have approached Fennoscandia prior to theultimate oblique docking, but unfortunately there exist no palaeo-magnetic data for the period between 1.86 and 1.83 Ga forVolgo-Sarmatia, which is due to the absence of known rocksof this age. After ca. 1.90 Ga, Sarmatia had almost completelyescaped tectonothermal reworking until becoming involved inthe Fenno-Sarmatian collision. Thereafter, it underwent post-collisional extension of its lower crust and was affected by majorAMCG magmatism between 1.80 and 1.75 Ga (Bogdanova et al.,2004, 2013).

And of course there existed also other potentially involved cra-tons in the Early Palaeoproterozoic world with similar Archaeancrustal ages, such as India, Kola-Karelia, North China, Sao-Francisco,Laurentia, Amazonia, West Africa, etc. (Bleeker, 2003; Pehrssonet al., 2013).

4.1.3. Metamorphic belts and domains, ductile deformation andboundaries of tectonic domains in the South Baltic region

Some of the most important questions in the field of Palaeopro-terozoic Trans-Baltic correlations concern the times at which thepresent tectonic domains in the South Baltic region were formedand how they relate to the patterns of tectonostratigraphical ter-ranes.

Granulite- and amphibolite-facies metamorphic belts anddomains were long considered to be the first order tectonic sub-divisions of the crust in the South Baltic region (Fig. 1, inset).While the distribution of rocks of different facies and lower- orhigher-pressure types of metamorphism in each tectonic domainis still not known exhaustively due to the scarcity of drill coredata, it has by now become obvious that granulite-facies beltsand domains extend across the boundaries between the tectonicdomains/terranes. This in part, is a consequence of the varyingages of granulitic metamorphism throughout the region and vary-ing history of uplift. In general, the granulites associate with areasof high strain and deformation expressed by major swarms of linearmagnetic anomalies (Koistinen, 1994; cf. Figs. 1b and 2a).

In the present study, the ages of the granulite-facies metamor-phism in the West Estonian metasedimentary rocks (sample Valgu99) and the granulites of the South Estonian domain (sample Kõnnu300) were of particular interest because of the previously reported1.78–1.75 Ga ages of their metamorphism (Puura et al., 2004;Soesoo et al., 2006), which are much younger than anything knownfrom comparable Svecofennian parts of the Baltic/FennoscandianShield. Summing up the new and previous datings of the granulitesin Estonia, we conclude that they were formed by several events,comprising:

a) high-temperature melting of the crust and crystallization of S-type garnet-orthopyroxene granodiorites (sample Kõnnu 300)in the lower-middle crust at 1891 ± 3 Ma followed by slow cool-ing and retrogression in the granulite to high amphibolite faciesalong with deformation at 1858 ± 4 Ma;

b) burial and lower granulite-facies metamorphism of the WestEstonian sedimentary rocks (sample Valgu 99) in the middlecrust at 1842 ± 1 Ma followed by slow uplift and retrogressionalong a near-isobaric evolution path similar to the path of the

Kõnnu 300 granulite (Puura et al., 2004);

(c) shearing of the crust on both sides of the Middle Estonian faultzone at 1798 ± 2 Ma (sample Valgu 99), semi-simultaneouslywith mafic dyking at 1774 ± 20 Ma (Soesoo et al., 2006). The final

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uplift of the Estonian granulites into the upper crust occurredat 1728 ± 24 Ga (garnet Sm–Nd age, Puura et al., 2004).

Shoshonitic magmatism punctuated this metamorphic evolu-ion at 1.83 Ga (Petersell and Levchenkov, 1994), which is closeo the 1.84 Ga age of the granulite-facies metamorphism of thealgu metasedimentary rocks. In the Latvian part of the South Esto-ian domain, 1810 ± 2 Ma metavolcanic granulites were formedearly simultaneously with the Mazsalaca granodioritic intrusiont 1816 ± 5 Ma. Altogether, the HT/LP near-isobaric retrograde evo-ution in the West and South Estonian domains was apparentlyustained by magmatism and attendant heating of the crust.

Farther south in the South Baltic region, the metamorphic evo-ution in the various tectonic domains differed from that in Estoniay the substantially lower ages of peak metamorphism and dif-erent paths of the pressure–temperature evolutions (Skridlaitet al., 2014). In the MLD and the WLG, the principal peak ofranulite-facies metamorphism was attained between 1.82 and.80 Ma (Appendix A). In the MD domain in Poland rocks at the

evel of observation never reached granulite-facies conditions. Theyere metamorphosed in the amphibolite facies at ca. 1.83–1.82 Ga

Williams et al., 2009).In summary, the zircon ages indicate that the granulite-facies

etamorphism in the MLD and WLG domains and the amphibolite-acies metamorphism in the MD culminated at slightly differentimes between 1.83 and 1.80 Ga. The TIMS and EPMA analysesielded slightly lower ages of monazite at 1.81–1.79 Ga (Skridlaitet al., 2014). Altogether, these age determinations and other stud-es show that the high grade metamorphism was nearly coeval ata. 1.80 Ga throughout the various southern tectonic domains andommonly associated with zones of ductile deformation. As differ-nt from the granulites in Estonia, the 1.82–1.80 Ga MLD and WLGranulites were formed at much higher peak temperatures reach-ng 900 ◦C at 10 kbar, while their retrograde P-T evolutions followedear-isothermal paths indicating post-compressional/collisionalettings. The retrograde processes were complicated substantiallyy repeated events of magmatism, which created stepwise P-Taths with periods of enhanced heating and subsequent near-

sobaric cooling at 1.73–1.68, 1.62–1.58 and 1.52–1.50 Ga. Theemperatures were high enough to allow the formation of newranulite-facies mineral assemblages during almost every stage ofeating, sometimes superposed upon the older ones. These pro-esses have been interpreted as far-field effects of accretionaryontinental back-arc tectonics along what is now the southwesternargin of Baltica (Skridlaite et al., 2014).The BPG and the OKL belts, which are situated adjacent to the

ennoscandia-Sarmatia suture and its metamorphic core complexBogdanova et al., 2006; Fig. 2b), exemplify yet another mode ofranulite-facies metamorphism at ca. 1.8 Ga. At the peak of the.80–1.75 Ga post-collisional prograde P-T evolution, which wasssociated with and generated by widespread AMCG magmatism,rustal doming, and mantle underplating, new mineral associationsere superposed upon ca. 1.87 Ga amphibolite-facies assemblages

Taran and Bogdanova, 2003).Our new age data and the results of recent geophysical and

rill core studies are thus in line with the view of Skridlaite et al.2014) that the granulite belts in the South Baltic region are com-lex, multistage formations, comprising rocks of various ages andiverse modes of metamorphic evolution. They were formed inarious geotectonic settings and generated by a range of differenteological processes:

In the case of the West and South Estonian domains therewas extension of a crust of normal thickness and magmatic

Research 259 (2015) 5–33

underplating and intraplating in continental back-arc settingsbetween ca. 1.89 and 1.84 Ga.

- Owing to the collision between Fennoscandia and Sarmatia(the “Fenno-Sarmatian collision”) between ca. 1.82 and 1.80 Ga,the crust was thickened, which triggered the formation of1.82–1.79 Ga granulites in the MLD and WLG domains. Peakconditions of metamorphism were attained ca. 50 Ma after theonset of subduction at ca. 1.86 Ga, which was accompanied bycontinental-margin magmatism in the MLD.

- Within and along the Fennoscandia-Sarmatia suture zone,post-collisional extension of the crust, its doming, and volumi-nous AMCG magmatism caused formation of ca. 1.80–1.79 Gagranulites in the BPG and OKL belts. The following fast uplift,retrogression of the granulites, and mylonitization affected theserocks already at 1.77 Ga.

4.2. Tectonic subdivisions of the Palaeoproterozoic crust in theSouth Baltic region revised

New age determinations, geophysical modelling, and consider-ations of the nature and origins of the deformation, magmatismand metamorphism in the South Baltic region necessitate a num-ber of important changes in our previous syntheses of the character,configurations and tectonic positions of its crustal domains. Thesechanges are as follows:

- In our new model we no longer regard the Okolovo (OKL)and Belarus-Podlasie (BPG) belts as parts of or closely relatedto Fennoscandia. The reason is that at 2.0–1.95 Ga these beltsunderwent an evolution similar to that of the north-westernmargin of Sarmatia. Structurally, they are related to large faultzones penetrating the entire crust and envelop a major domeof lower crust within the Fennoscandia-Sarmatia suture zone(Figs. 1 and 2a, b). Furthermore, there is a notable discordancebetween the tectonic domains of Fennoscandia and the BPGand OKL. Very consistently, the NE-trending Grodno-Bialystokdeformation zone, which defines the north-western limit of theBPG, truncates obliquely the essentially NW-SE-trending arrayof various Fennoscandian belts and domains, which are bentand displaced in the vicinity of this boundary. It must thereforebe interpreted as one of the major shear zones related to the1.82–1.80 Ga Fenno-Sarmatian collision. The Grodno-Bialystokgranitic mylonites along this boundary (sample GR7 of thepresent study) were formed at ca. 1.82 Ga like the adjacent gran-ulites (Skridlaite et al., 2014).

Inherent in this re-interpretation is that the South Estonian(SEG) with its granulites can no longer be seen as a northernextension of the BPG or a part of the single Baltic-Belarussian beltin early publications (e.g. Gorbatschev and Bogdanova, 1993).

- Despite the differences of their metamorphic grades, the WestEstonian (WE), South Estonian (SEG) and Latvian-East Lithua-nian (LEL) domains are similar by the 1.89–1.87 Ga ages of theirprincipal magmatic events and the subsequent tectonothermalevolution. This suggests that these three domains were once partof a single crustal unit representing different crustal levels andstyles of deformation. Bimodal juvenile volcanism with associ-ated iron and sulfide (±manganese) ore deposits, quartzites andskarns, and high-T plutonic intrusions indicate extensional tec-tonic settings similar to those of continental back-arcs. Also therecorded 2.2–2.1 Ga Nd TDM model ages are equivalent through-out the three domains.

- The results of the present study and other forthcoming informa-

tion require very substantial revision also of the configurationand outlines of the Mid-Lithuanian domain (MLD) and particu-larly its boundary towards the West Lithuanian domain (WLG, cf.Figs. 1 and 2a). The new configuration proposed here agrees better

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with the seismic-gravity modelling along the EUROBRIDGE’95profile, which indicates distinct differences of crustal structureacross the former WLG (Kozlovskaya et al., 2001). Its previouslynorthern part (now the NW half of the MLD) has a thicker-than-normal crust due to a thick lower crust and the presence of alow-density low-velocity layer in the upper crust, which contin-ues across the entire MLD to the southeast. Also, large intrusionsof 1.86–1.84 Ga charnockites in northern Lithuania (Motuza et al.,2008; Vejelyte et al., 2012) and the 1.84 Ga calc-alkaline meta-diorite (sample Gl99 of this study) as well as other similar rocksin central and southern Lithuania (Appendix A) are very charac-teristic of the MLD domain as now defined. Some of its boundaries,however, are difficult to trace due to overlapping fault zonesof various ages and orientations (Figs. 1 and 2a). The N-S trenddepicted in previous maps for a Mid- (or Central) LithuanianSuture zone was mostly based on geological and geophysical indi-cations in southern and south-central Lithuania (Skridlaite andMotuza, 2001; Motuza, 2005; Motuza et al., 2008), which havelately proved irrelevant farther to the north. Thus there exist noindications at present of a suture extending all the way to thecoast north of Riga, while the suture zone in the south has nowbeen incorporated in the MLD as its sheared eastern delimitation.

New age determinations of metasedimentary rocks in theWLG (this study) and the comprehensive new geochronologi-cal database now available for Poland (Krzeminska et al., 2005,2011, 2014; Wiszniewska et al., 2007; Williams et al., 2009) allowthe outlining of Palaeoproterozoic sedimentary basins and a verylarge area of 1.81–1.76 Ga granitoid rocks and coeval volcanicbelts.

The ages of the most widespread events of metamorphism andylonitization in the area indicate that almost all of the tectonic

tructure of the crust in the South-Baltic region was shaped duringhe 1.82–1.80 Ga Fenno-Sarmatian collision but complicated sub-equently by extension that lasted until ca. 1.7 Ga and locally evenater (Bogdanova et al., 2001, 2006; Skridlaite et al., 2014).

In summary, our revision of the previous concepts on the tec-onic structure and subdivisions of the crust provides new cluesor the reconstruction of the tectonic settings and domain pat-erns in the South Baltic part of Fennoscandia. We suggest thathe West and South Estonian domains as well as the Latvian-Eastithuanian domain were once parts of one single large continen-al block with Nd TDM model ages of 2.2–2.0 Ga, for which weropose the name “Livonia” (Fig. 9). The ca. 1.89 Ga Svecofennianagmatism and the high-T metamorphism, the associated Fe-S-

i ore deposits and the terrigeneous and calcareous sedimentationn continental back-arc basins are suitable salient points for fur-her correlation. The enigmatic Latgalian domain, however, is mostommonly seen either as a virtually unreworked part of Livonianrust or a detached fragment of the Novgorod domain/terrane far-her NE. In northern Estonia, the Tallinn metavolcanic domain andhe Jõhvi belt may have shared a crustal evolution similar to thatf Livonia.

Owing to the presence in the Mid-Lithuanian domain (MLD) ofocks characteristic of the Latvian-East Lithuanian (LEL) domaint has long been considered transitional between these twoectonic units. However, a chain of 1.86–1.84 Ga calc-alkalineabbroic to granitic intrusions along its SW boundary ratheruggest an active proto-continental margin of Livonia facinghe younger, 1.83–1.82 Ga Polish-Lithuanian domains. In the NEart of the latter, the Monki and Poceai basins of sedimen-

ation (Mk and Pc in Figs. 1 and 2a) were formed between.86 and 1.83 Ga, similar in time but differing in their tectonicettings, continental and near-oceanic, respectively (cf. Section.1.2).

Research 259 (2015) 5–33 23

4.3. Trans-Baltic Palaeoproterozoic correlations and tectonicmodels for the Svecofennian orogen

The continuity of magnetic and gravity anomalies across theBaltic Sea (Koistinen, 1994; Wybraniec, 1999; Korhonen et al.,2002) and the close similarity between the 1.9–1.8 Ga crustal devel-opments in the South Baltic region and the Svecofennian orogen ofthe Baltic/Fennoscandian Shield (Fig. 8) suggest that the entire Sve-cofennian orogen and particularly its central and southern partsare almost entirely built up of regularly arranged belts of NW-trending tectonic domains and mega-domains (Figs. 1 and 9),which are complicated and separated from each other by mostlydextral transpressive shear zones, apparently connected to majorgeophysically detectable boundaries in the crust and the uppermantle (Korja and Heikkinen, 2005; Lahtinen et al., 2009b). Of sub-stantial interest is that the generation and development of thesezones occurred dominantly between 1.82 and 1.80 Ga, even thoughsome of them like, for instance, the Hassela shear zone (HSZ inFigs. 1 and 9), were initiated already at 1.87–1.85 Ga (Beunk andPage, 2001; Högdahl et al., 2009; Stephens et al., 2009; Beunk andKuipers, 2012).

The most conspicuous of these large tectonic units is the NW toSE trending assemblage of lithologically and chronologically relatedtectonic domains on both sides of the central Baltic Sea (Fig. 8).Its components are the Bergslagen, Uusimaa/North-Estonian andLivonia megadomains (the latter comprising the SEG and the LELtectonic domains) on either side of the Middle Estonian faultzone (MEFZ) and its possible continuation across the Baltic Seato Sweden, where it may pass into the Vingåker-Nyköping shearzone (Beunk and Page, 2001) that bounds the Bergslagen tectonic“domain 3” of Hermansson et al. (2008) and Stephens et al. (2009)in the south. Farther north, the Hagsta-Gävle-Rättvik system ofshear zones (HGZ-GRZ) along the northern edge of Bergslagenmeets the western end of the South Finland shear zone (Ehlerset al., 1993), which ultimately joins the Paldiski-Pskov DeformationZone (PPDZ) in Estonia (Högdahl et al., 2009). Thus it appears thatthe PPDZ and MEFZ may be conjugate dextral shear zones defin-ing a wedge-shaped structural relationship, where the Alutaguse(AL) domain of metaturbidites is a component (Figs. 1, 2 and 9).As argued above (text Sections 2 and 4.2), the ca. 1.92 Ga Tallinnfelsic metavolcanic rocks and particularly the Jõhvi belt that har-bours a large number of iron-sulfide-silica (FeSSi in Figs. 8 and 9)ore deposits, have their counterparts in the Uusimaa belt (Uu) insouthern Finland. Like the rocks in the South Finnish migmatitebelt along the southern coast of Finland (Kurhila et al., 2011), therocks in northern Estonia have undergone migmatization and arepenetrated by 1.79 Ga granitoids (Appendix A). Together, the Uusi-maa belt and the North Estonian domains form a large crustal unit,which in many ways resembles the Bergslagen region in south cen-tral Sweden (Kähkönen, 2005). The presence of Fe-S-Si ore depositsassociated with 1.91–1.89 Ga felsic volcanic and sub-volcanic rocks,and the semi-coeval sequence of turbiditic and carbonate sedimen-tary rocks and skarns, are a hallmark of the Bergslagen region (cf.Stephens et al., 2009) but are also prominent in the Uusimaa-NorthEstonian area (Fig. 8).

Whereas the close similarity and possible continuity of theBergslagen microcontinent and the Uusimaa belt have been dis-cussed for some time now, and Lahtinen with co-workers (2005,2009b) even proposed that the Bergslagen microcontinent com-prises both the Bergslagen area in Sweden and the Uusimaabelt along the coast of southernmost Finland, the three WE, SEGand LEL domains (i.e. our megadomain Livonia) emerge as yet

another conceivable Bergslagen counterpart (Fig. 9). Seismically,Bergslagen has also been traced in the middle and lower crustbeneath the Ljusdal domain farther north (Korja and Heikkinen,2005).

24 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Fig. 8. Time-space chart presenting the Trans-Baltic Palaeoproterozoic correlations. Source data from: Andersson (1997), Andersson et al. (2004), Bergman et al. (2008),Beunk and Page (2001), Hermansson et al. (2007, 2008), Högdahl and Sjöström (2001), Högdahl et al. (2004, 2009), Johansson et al. (2006), Krzeminska et al. (2009), Lahtinenand Nironen (2010), Lahtinen et al. (2002, 2009a), Mansfeld et al. (2005), Rutland et al. (2001, 2004), Stephens et al. (2009), Sultan et al. (2005), Wikström and Andersson(2004), Williams et al. (2008, 2009). Fe-S-Si indicates the presence of various iron- and sulfide ore deposits.

Fig. 9. Crustal structure in the central and southern parts of the Svecofennian orogen as integrated across the Baltic Sea. Bergslagen, Livonia and Amberland, like Keiteleand Bothnia in the north, are or may be microcontinents. All, except the Ljusdal domain in northern Bergslagen, appear marked by active-continental-margin igneous beltsalong their south-western margins Bergslagen. The Ljusdal margin could have been in a different position in a separated Bergslagen during its accretion and rotations at1.83–1.82 Ga (Beunk and Kuipers, 2012). Most abbreviations are the same as in Fig. 1. The ages in the frames are those of the principal accretionary events in the indicateddomains. The dashed-triangle lines indicate upper surfaces of dipping mantle reflectors according to Lahtinen et al. (2009b). Fe-S-Si indicates the presence of various iron-and sulfide-ore deposits.

brian

mEtp1sLT(asmasoewtgmaapcBaw

(swt(ec2gThta(FacvsBpS

toDclhStaWobcr

S. Bogdanova et al. / Precam

Although, as shown in Fig. 9, the Bergslagen and Livonianegadomains are separated from each other by the 1.8 Ga Middle

stonian Fault Zone (MEFZ) and its western continuation beneathhe Baltic Sea, their configurations allow continuity and the Palaeo-roterozoic evolutions appear similar (Fig. 8). Like in Bergslagen, ca..9 Ga felsic volcanic rocks, characteristic Fe-S-Si ore deposits, andedimentary rocks are present in the South Estonian (SEG) and theatvian-East Lithuanian (LEL) domains of the Livonia megadomain.he metavolcanic rocks in Livonia and Bergslagen have similar �Nd1.9) values varying between −0.7 and +3.8 in Livonia (Appendix A),nd between −2.6 and +4 in Bergslagen (Stephens et al., 2009), andimilar Nd TDM values of 2.2–2.0 Ga. Also, sedimentation in the twoegadomains occurred simultaneously between >1.9 and 1.89 Ga,

nd there is agreement that magmatism, Fe-S-Si ore formation andedimentation in the Bergslagen region suggest back-arc riftingf a microcontinent at 1.91–1.89 Ga (Allen et al., 1996; Stephenst al., 2009; Beunk and Kuipers, 2012). Subsequently, Bergslagenas involved in the compressional and extensional accretionary

ectonics, which was accompanied by widespread calc-alkalineranitoid magmatism at 1.90–1.87 and 1.86–1.85 Ga, and also inter-ittent sedimentation (Bergman et al., 2008), mafic underplating

nd calc-alkaline to alkali-calcic granitoid magmatism (Wikströmnd Andersson, 2004). A similar development appears to have takenlace also in Livonia. Thus there exist at present three feasibleorrelation modes across the central part of the Baltic Sea: eitherergslagen–Uusimaa/northern Estonia, or Bergslagen–Livonia, orll these three tectonic megadomains together within the frame-ork of a single block of continental crust (Fig. 8).

To the south of Bergslagen – Livonia is the Mid-Baltic beltFig. 9), which comprises the Mid-Lithuanian domain (MLD) in theoutheast and the southernmost part of the Bergslagen region asell as granites of the 1.87–1.84 Ga Askersund-Loftahammar plu-

ons in Sweden followed by intrusions of 1.81–1.76 Ga TIB-1 suitesHögdahl et al., 2004; Stephens et al., 2009). Similarly, the north-rn MLD part of the Mid-Baltic belt features large 1.86–1.82 Gaharnockitic intrusions (Motuza et al., 2008; Motuza and Motuza,011; Vejelyte et al., 2012) while 1.86–1.84 Ga mostly calc-alkalineranitic to gabbroid rocks occur along its southwestern margin.his tendency of southward younging of accretionary magmatismas been observed also in the Swedish part of the Mid-Baltic Belto the west of the Baltic Sea with its 1.86–1.85 Ga calc-alkalinective continental margin magmatism in what Hermansson et al.2008) define as their “domain 4” of the Bergslagen region (VV inigs. 1 and 9). In that domain, however, there was still a substantialmount of 1.90–1.87 Ga calc-alkaline Bergslagen magmatism. Ata. 1.86 Ga, quartz-rich sandstones were deposited in the Väster-ik area and its continuation towards the west along a one-timeouth-facing continental slope (Sultan et al., 2005; Sultan and Plink-jörklund, 2006). The south-eastern extreme part of the MLD inart coincides with the former “Mid-Lithuanian Suture Zone” ofkridlaite and Motuza (2001).

The ca. 1.83–1.75 Ga Polish-Lithuanian tectonic megadomain ishe youngest major crustal unit in the platform area to the eastf the Baltic Sea. It comprises the WLG in Lithuania and the MD,O and PM domains in Poland, which are dominantly granitic inomposition. Because of the close similarity of these rocks to theikewise mainly granitic 1.83–1.75 Ga rocks in SE Sweden, theyave long been reckoned as a separate terrane akin to the Eastmåland (ESL) tectonic domain (Bogdanova et al., 2006). In bothhese cases, granodioritic and monzonitic intrusions with associ-ted volcanic belts occur widely (Krzeminska et al., 2005, 2014;iszniewska et al., 2007). In Sweden, the northernmost margin

f the ESL is marked by the 1.83–1.82 Ga Oskarshamn-Jönköpingelt (O-J in Figs. 1 and 9) of juvenile mafic and felsic volcanics,alc-alkaline intrusions and coarse clastic metasedimentary rockselated to an assumed northward subduction zone (Mansfeld et al.,

Research 259 (2015) 5–33 25

2005; Andersson et al., 2007). This agrees well with interpretationsof the BABEL seismic reflection line B (e.g. Abramovitz et al., 1997;Korja and Heikkinen, 2005), which identify a NE dipping reflectortransecting the entire crust beneath the O-J belt as the trace of asubduction wedge (Fig. 9). Southwards, the O-J belt is succeededby major WNW-trending massifs of 1.81–1.76 Ga TIB-1 granitoidsalternating with belts of volcanic rocks (Gorbatschev, 1980, 2004).A continuation of these highly magnetized rocks southeastwardsacross the sea towards Poland and Lithuania is easily recognizedfrom large positive magnetic anomalies (e.g. Korhonen et al., 2002).In recognition of its role as one of the principal sources of amber inEurope ever since the bronze age, we propose the name “Amber-land megadomain” for this southernmost Fennoscandian tectonicunit (Fig. 9).

In summary, the current revision of the structure and the tec-tonic subdivisions of the crust in the South Baltic region and theTrans-Baltic correlation of the major tectono-stratigraphical com-plexes have created a largely new image of PalaeoproterozoicFennoscandia and its tectonic relationships with Sarmatia (Fig. 9).

- The central and southern parts of the Svecofennian orogen thuscomprise a sequence of NW-trending, 100–300 km wide tec-tonic megadomains/belts bounded by a setup of transpressional,mostly dextral shear zones, which were developed wholly byca. 1.82–1.79 Ga. These major tectonic units differ by the agesof their principal orogenic events but feature pronounced con-tinuity in the formation of their crust and its general youngingtowards the present south-southwest. Some of their parts havebeen identified as microcontinents or tectonically and magmati-cally reworked margins of microcontinents, others as probablejuvenile island arcs. As a result of the Fenno-Sarmatian colli-sion at 1.82–1.80 Ga, their southeasternmost parts were indentedand displaced along the Sarmatia-related Belarus-Podlasie andOkolovo belts.

- The Bergslagen and Livonia megadomains appear to havebeen acting as a single continental unit (microcontinent) after1.91–1.89 Ga (Fig. 9). Between 1.89 and 1.84 Ga Bergslagen-Livonia was involved in accretionary orogeny with several stagesof extension and compression (Hermansson et al., 2008; Stephenset al., 2009; Beunk and Kuipers, 2012).

- Between 1.86 and 1.84 Ga, accretion of new crust occurredparticularly in the Mid-Baltic belt along the SW edge ofBergslagen-Livonia, where there are many traits of an active con-tinental margin and magmatism ranging from calc-alkaline at itsoutermost rim to alkali-calcic and calc-alkalic in the interiors.Whether the Mid-Baltic belt was of an Andean-type, facing anocean and developed in continental back-arc settings is still anopen question.

- The subsequent accretion of Amberland occurred after a timebreak between ca.1.84 and 1.83 Ga, when the O-J juvenile vol-canic belt was formed. Possible time-correlatives of that belt arethe presumably ocean-related Poceai (Pc) sedimentary basin inLithuania with a maximum deposition age of ca. 1.85 Ga, and theMonki 1.83–1.82 volcanic belt as well as the 1.86–1.83 Ga Monki(Mk) continental sedimentary basin in Poland (Williams et al.,2009). The Pc basin, which is situated close to the SW marginof Bergslagen-Livonia, suggests an intervening marine environ-ment either of an oceanic back-arc type or the relic of an ocean.The assumedly microcontinental Amberland contains TIB-1 gran-itoids and coeval volcanic belts formed in Andean-type tectonicsettings indicating its continental affinity.

- Dominantly S to SSW directions of accretionary crustal growth

throughout the central and southern parts of the Svecofennianorogen on both sides of the Baltic Sea suggest the existence ofa master subduction zone operating consistently in the samedirection during several extension-compression cycles between

2 brian

-

-

5r

hSitnbcoGit

lo

6 S. Bogdanova et al. / Precam

1.9 and 1.8 Ga as proposed by Hermansson et al. (2008) forthe tectonic evolution of the Bergslagen region. The relativelyshort periods of compression and characteristic continental-arctype calc-alkaline magmatism alternated with stages of exten-sion when back-arc type basins of sedimentation were developedand bimodal magmatism as well as high-grade metamorphismoccurred. All these features are typical of accretionary orogensthroughout the times (Collins, 2002; Cawood et al., 2009). Fur-thermore, this tectonic model may satisfy the “marginal basinaccretion” hypothesis of Rutland et al. (2001, 2004), which, how-ever, is inconsistent with the prominence of microcontinentaccretion that characterizes much of the tectonic evolution ofthe Svecofennian orogen. Virtually all the already well-definedand still hypothetical microcontinents in that orogen have indis-tinguishable 2.2–2.0 Ga Nd TDM crustal ages, which agrees withthe model of suspect terranes or ribbon continents (Beunk andKuipers, 2012; Lahtinen et al., 2014) elaborated for the MesozoicCanadian Cordillera along the East Pacific (Coney et al., 1980;Johnston, 2008). While the northern Svecofennian microconti-nents adjacent to the Karelian and/or Norrbotten cratons (Bothniaand Keitele of Lahtinen et al., 2005, 2009b) may be autochtonous,the possible affinity of Bergslagen, Livonia and Amberland withSarmatia or some other continents requires additional study. OurTrans-Baltic Palaeoproterozoic correlations thus indicate that theSvecofennian orogen is not a haphazard collage of microconti-nents and island arcs but rather a regularly structured realm ofsouth-southwestwards younging, stepwise accreted continentalcrust.

At 1.82–1.80 Ga, the continent–continent oblique collision ofVolgo-Sarmatia with Fennoscandia interfered with the accre-tionary growth of the crust in the Svecofennian orogen. TheFenno-Sarmatian collision took place after the accretion-relatedformation of the Oskarshamn-Jönköping belt and its easternequivalents at 1.83–1.82 Ga, but before the stitching Andean-typeTIB magmatism encompassed all of the western and southwest-ern Svecofennides at 1.81–1.76 Ga.

The Fenno-Sarmatian collision during overall N-S compressioncompleted the formation of the present tectonic patterns inthe central and southern parts of the Svecofennian orogen.Subsequently, along with continuing accretion a system of trans-pressional shear zones complicated southwestern Fennoscandiaat 1.80–1.75 Ga.

. Some inferences with regard to Columbia/Nunaeconstructions

The Palaeoproterozoic Trans-Baltic correlation reported hereas nearly doubled the size of the hitherto surveyed part of thevecofennian orogen, while important changes in our understand-ng of its character and configuration prior to the initiation ofhe Columbia/Nuna assembly at ca. 1.8–1.7 Ga are reported in theew model presented graphically in Fig. 9. Thus it appears cleary now that the formation of the Svecofennian orogen betweena. 2.0 and 1.75 Ga essentially preceded most of the assemblyf Colombia/Nuna and the creation of its coherent Laurentia-reenland-Baltica margin. The Scandinavian part of that margin

s rather the Gothian orogen extending from north central Norway

o SW-most Sweden, not the Svecofennian orogen.

In Fennoscandian contexts, the existence of microcontinentsike the Keitele and Bothnia along the extended passive marginsf the Archaean Karelia and Norrbotten cratons in the north and

Research 259 (2015) 5–33

northeast has inspired recent hypotheses of orocline tectonicsof the Svecofennian terranes at ca. 1.87 Ga (e.g. Lahtinen et al.,2014). Similar accretionary tectonism at ca. 1.83–1.82 Ga has alsobeen adopted for Bergslagen in south-central Sweden (Beunk andKuipers, 2012).

As demonstrated by the new Trans-Baltic tectonic model inFig. 9, the developments in the southern half of the Svecofennianorogen at 1.87–1.75 Ga involved continuous and single-polaritynorth-eastward accretion of the microcontinent Bergslagen-Livonia and presumably also large parts or even the wholeof Amberland. This growth of the crust was interrupted at1.82–1.80 Ga by the oblique collision of Fennoscandia and Volgo-Sarmatia, which created the markedly transpressional structureof southern Fennoscandia. In the present southeast, on the otherhand, this collision caused the tectonic rotation of Sarmatia byca. 45◦ anticlockwise between 1.80 and 1.75 Ga and its exten-sion, which was associated with powerful AMCG magmatism(Bogdanova et al., 2013). The final adjustment of Fennoscandiaand entire Volgo-Sarmatia to complete the formation of Palaeo-proterozoic Baltica took place between ca. 1.75 and 1.7 Ga whenthe Laurentia-Greenland-Baltica margin of Columbia/Nuna wasestablished. Characteristically, that margin truncates the older, pre-1.8 Ga, orogenic belts at high angles, in nearly the same mannerin which the NW-SE trending Svecofennian orogen was truncatedby the N-S Gothian orogenic belt and the Trans-Hudson belt wascut by the Central Plains orogen. Still other examples are the NewQuebec and Makkovik orogenic belts cut by the Labrador orogen(Gower et al., 1990; Gorbatschev and Bogdanova, 1993; Karlstromet al., 2001). A conclusion arising from the above is that onlyca. 1.75–1.70 Ga and younger orogenic belts along the Laurentia-Baltica margin can be related to the formation of the margins andterminal outlines of assembled supercontinent Columbia/Nuna.

A further intriguing possibility is that a continuation of theSvecofennian tectonic domains might be found in some otherPrecambrian craton rifted away from Fennoscandia. If so, thatwould be an excellent tool of transcontinental correlation in caseswhich are still very controversial from both the geological and thepalaeomagnetical points of view (Zhao et al., 2011; Evans, 2013;Pisarevsky et al., 2014) and another important feedback of ourTrans-Baltic modelling of the Svecofennian orogen. Again, wheresuch continuity does not exist, that contradicts possible correla-tion.

Acknowledgements

The help provided by Martin Whitehouse, Lev Ilyinsky and Ker-stin Lindén at the NORDSIM laboratory is gratefully acknowledged.The ion microprobe NORDSIM facility in Stockholm is operatedunder an agreement between the joint Nordic research councils(NOS-N), the Geological Survey of Finland and the Swedish Museumof Natural History. This paper is NORDSIM contribution number388. The project was in part funded by the EU-supported IntegratedInfrastructure Initiative “SYNTHESYS” (SYNthesis of SYStematicresources). It is a contribution to the project of the Swedish Insti-tute’s Visby Programme “Precambrian rock provinces and activetectonic boundaries across the Baltic Sea and in adjacent areas”. Thisstudy was also supported by the Estonian Ministry of Education and

Research target research project no. SF0140016s09 and by grant no.8963 (ESF) to AS. Dr. Frank Beunk and an anonymous reviewer arethanked for their comments, suggestions and corrections of the firstversion of the manuscript.

S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33 27

Appendix A. Summary of age determinations and Sm–Nd isotopic data for Paleoproterozoic rocks of the South Baltic region

Domain,belt

No. onFig. 3A

Well, drillcore

Coordinates Rock Sm–Ndwhole rockTDM Ga,�Nd(T)

Age, Madetritalzirconpeaks in Ga

Methoda/mineralzircon (zr),monazite(mz)

Origin ofmineral

Reference

North EstonianTallinn

domain(TN)

1 JägalaF-110

59 26 N/2513 E

Metapeliticgneisses

�Nd −2.30 2.13–1.85 U–Pb zrTIMS

Magmatic,detrital

Petersell andLevchenkov(1994),Petersell(unpubl.)

TN 2 PärispeaF-113

59 39 N/2542 E

Metapeliticgneisses

�Nd −2.30 2.13–1.85 U–Pb zrTIMS

Magmatic,detrital

Petersell andLevchenkov(1994),Petersell(unpubl.)

Tapa belt(T)

3 KarkuseF-142

59 12 N/2558 E

Tonalite �Nd 0 1824 ± 26 U–Pb zr LA Magmatic Soesoo et al.(2006)

T 4 Moe F-164 59 15 N/2601 E

Metavolcanics 1918 ± 10 U–Pb zrTIMS

Magmatic Petersell andLevchenkov(1994)

T Moe F-164 59 15 N/2601 E

Charnockite �Nd −0.3 1761 ± 11 U–Pb zr LA Magmatic Soesoo et al.(2006)

Johvi belt 5 Jõhvi 1 59 23 N/2730 E

Magnetite quarzite 1.89, 1.93,1.98, 2.0,2.2

U–Pb zrSIMS

Magmatic,detrital

Soesoo(unpublished)

WestEstoniandomain(WE)

6 Muhu 590 58 38 N/2313 E

Amphibole-biotiteorthogneiss

1827 ± 7 U–Pb zrTIMS

Magmatic Petersell andLevchenkov(1994)

WE 7 Valgu 99 58 50 N/2435 E

Metasedimentarygneiss

1.94, 1.90 U–Pb zrSIMS

Magmatic,detrital

This paper

WE Valgu 99 58 50 N/2435 E

Metasedimentarygneiss

1842 ± 2 U–Pb SIMS Metamorphic This paper

WE Valgu 99 58 50 N/2435 E

Metasedimentarygneiss

1798 ± 2 U–Pb SIMS Metamorphic This paper

WE 8 F-300Voore

59 16 N/2435 E

Amphibole-biotitegneiss

2.08,�Nd(1.9)+1.5

Puura andHuhma(1993)

South Estonian Granulite belt (SEG)SEG 9 Varbla

F-50258 22 N/2350 E

Metasedimentarygneiss

�Nd −1.9to −1.6

1.93–1.84 U–Pb zrTIMS

Magmatic?metamorphic?

Petersell andLevchenkov(1994),Petersell(unpubl.)

SEG VarblaF-502

58 22 N/2350 E

Biotite gneiss 1870 ± 22 U–Pb zr LA Magmatic?metamorphic?

Soesoo(unpublisheddata)

SEG VarblaF-502

58 22 N/2350 E

Tonalite �Nd −0.4 1788 ± 16 U–Pb zr LA Magmatic Soesoo et al.(2006)

SEG 10 Kõnnu 300 58 43 N/2450 E

Garnet-orthopyroxenegranulite

2.14, �Nd(1.9) +0.4

2.1, 2.0 U–Pb zrSIMS

Magmatic,detrital

Puura andHuhma(1993), thispaper

SEG Kõnnu 300 58 43 N/2450 E

Garnet-orthopyroxenegranulite

1891 ± 3 U–Pb zrSIMS

Magmatic This paper

SEG Kõnnu 300 58 43 N/2450 E

Garnet-orthopyroxenegranulite

1858 ± 4 U–Pb zrSIMS

Metamorphic This paper

SEG Kõnnu 300 58 43 N/2450 E

Garnet-orthopyroxenegranulite

1778 ± 2 U–Pb mzTIMS

Metamorphic Puura et al.(2004)

SEG Kõnnu 300 58 43 N/2450 E

Garnet-orthopyroxenegranulite

1728 ± 24 Sm–Nd grt Metamorphic Puura et al.(2004)

SEG Kõnnu 300 58 43 N/2450 E

Mafic dyke 1774 ± 20 U–Pb zr LA Magmatic Soesoo et al.(2004)

SEG 11 HäädemeesteF-172

58 04 N/2434 E

Qz-feldspaticorthogneiss

1832 ± 22 U–Pb zrTIMS

Magmatic Petersell andLevchenkov(1994)

SEG 12 Laeva 18L 58 29 N/2623 E

Intermediategranulite

2.18,�Nd(1.9)+0.2

Puura andHuhma(1993)

28 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Appendix A. (Continued)

Domain,belt

No. onFig. 3A

Well, drillcore

Coordinates Rock Sm–Ndwhole rockTDM Ga,�Nd(T)

Age, Madetritalzirconpeaks in Ga

Methoda/mineralzircon (zr),monazite(mz)

Origin ofmineral

Reference

SEG 13 Elva 555 58 13 N/2625 E

Mafic granulite 2.08,�Nd(1.9) +2

Puura andHuhma(1993)

SEG 14 Vaki 66 58 41 N/2528 E

Qz-feldspaticorthogneiss

1828 ± 8 U–Pb zrTIMS

Magmatic Petersell andLevchenkov(1994)

Latvian-EastLithua-nian belt(LEL)

15 Strenchi 3 57 23 N/2539 E

Mafic granulite 2.17,�Nd(1.9)+1.2

Mansfeld(2001)

LEL 16 StrenchiGK-12PR

57 33 N/2539 E

Mafic granulite 2.21,�Nd(1.9)−0.7

Mansfeld(2001)

LEL StrenchiGK-12PR

57 33 N/2539 E

Metasedimentarygneiss

2.48,�Nd(1.9)−0.7

Mansfeld(2001)

LEL 17 ValmieraGK-11 K

57 4 N/2530 E

Mafic granulite 1.99,�Nd(1.9)+2.9

Mansfeld(2001)

LEL 18 Staicele 1 57 49 N/2445 E

Metarhyodacite 2.09,�Nd(1.9)+1.5

1870 ± 13 U–Pb zrTIMS

Magmatic Mansfeld(2001)

LEL Staicele 1 57 49 N/2445 E

Metarhyodacite 1810 ± 2 U–Pb zrTIMS

Metamorphic Mansfeld(2001)

LEL 19 StaiceleGK1-KC

57 50 N/2448 E

Metarhyodacite 1.97,�Nd(1.9)+2.9

Mansfeld(2001)

LEL 20 MazsalacaGK-9p

57 52 N/2458 E

Metarhyodacite 2.04,�Nd(1.9)+1.8

Mansfeld(2001)

LEL 21 MazsalacaGK-4P

57 50 N/2507 E

Granodiorite 2.02 Ga,�Nd(1.9)+1.3

1816 ± 5 U–Pb zrTIMS

Magmatic Mansfeld(2001)

LEL 22 Pulkule 57 44 N/2451 E

Metarhyodacite 2.0 Ga,�Nd(1.9)+2.3

Mansfeld(2001)

LEL 23 Kraslava104

55 53 N/2720 E

Felsicmetavolcanic

2.06,�Nd(1.9)+1.7

Mansfeld(2001)

LEL 24 Subate 2A 56 04 N/2548 E

Metarhyodacite 2.04,�Nd(1.9)+1.9

Mansfeld(2001)

LEL 25 Inchukalns2P

57 07 N/2441 E

Metasedimentaryrock

1.97,�Nd(1.9)+2.7

Mansfeld(2001)

LEL Inchukalns2P

57 07 N/2441 E

Granitoid 1.88,�Nd(1.9)+4.2

Mansfeld(2001)

LEL 26 Salcia 403 54 20 N/2455 E

Metagreywacke 2.29–2.41,�Nd(1.9)−1,−1.8

2.2, 2.1 U–Pb zrTIMS

Magmaticdetrital

Mansfeld(2001)

LEL 27 Valkininkai404

54 19 N/2449 E

Metadolerite 1.90,�Nd(1.9)+3.8

Mansfeld(2001)

LEL 28 Varena1058

54 07 N/2413 E

Gabbro 1791 ± 9 U–Pb zrSIMS

Magmatic Skridlaiteet al. (2011)

LEL 29 Lazdijai 8(Lz8)

54 14 N/2334 E

Metadiorite 1887 ± 9 U–Pb zrSIMS

Magmatic This paper

LEL Lazdijai 8(Lz8)

54 14 N/2334 E

Metadiorite 1836 ± 14 U–Pb zrSIMS

Metamorphic This paper

Mid-Lithuanian domain (MLD)MLD 30 Varlinis

26854 01 N/2403 E

Diorite 1848 ± 6 U–Pb zrSIMS

Magmatic Skridlaiteet al. (2011)

MLD 31 Zeimiai 347 54 05 N/2407 E

Metagabbro 1836 ± 17 U–Pb zrTIMS

Magmatic Rimsa et al.(2001)

MLD Zeimiai 347 54 05 N/2407 E

Diorite 1859 ± 5 U–Pb zrSIMS

Magmatic Skridlaiteet al. (2011)

MLD 32 Geluva 99 55 14 N/2330 E

Mafic granulite 1839 ± 15 U–Pb zrSIMS

Magmatic This paper

MLD Geluva 99 55 14 N/2330 E

Mafic granulite 1809 ± 9 U–Pb zrSIMS

Metamorphic Skridlaiteet al. (2014)

S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33 29

Appendix A. (Continued)

Domain,belt

No. onFig. 3A

Well, drillcore

Coordinates Rock Sm–Ndwhole rockTDM Ga,�Nd(T)

Age, Madetritalzirconpeaks in Ga

Methoda/mineralzircon (zr),monazite(mz)

Origin ofmineral

Reference

MLD 33 Grauzai105

55 13 N/2352 E

Granite 1832 ± 5 U–Pb zrTIMS

Magmatic Motuza et al.(2008)

MLD 34 Grodno 7Gr 53 28 N/2346 E

S-type granitemylonitized

1844 ± 8 U–Pb zrSIMS

Magmatic This paper

MLD Grodno 7Gr 53 28 N/2346 E

S-type granitemylonitized

1822 ± 8 U–Pb zrSIMS

Metamorphic This paper

MLD Grodno 7Gr 53 28 N/2346 E

S-type granitemylonitized

1727 ± 30 mz EPMA Metamorphic Skridlaiteet al. (2014)

MLD 35 Lazdijai 13 54 11 N/2352 E

Metavolcanics 2.08,�Nd(1.9)+1.8

1.83, 1.79 U–Pb zrSIMS

Magmatic Skridlaiteet al. (2012)and Claessonet al. (2001)

MLD 36 Lazdijai 32 54 00 N/2333 E

Anatectic granite 2.20,�Nd(1.9)+0.2

2.04, 2.0,1.94

U–Pb zrSIMS

Magmaticdetrital

Krzeminskaet al. (2009)and Mansfeld(2001)

MLD 37 Macuiciai 1 55 43 N/2125 E

Opdalite 1846 ± 12 U–Pb zrTIMS

Magmatic Motuza et al.(2008)

MLD 38 Kuziai-65 56 03 N/2315 E

S-type granite 1844 ± 5 U–Pb zrTIMS

Magmatic Motuza et al.(2008)

MLD 39 Genciai 6 55 52 N/2111 E

Charnockite 1846 ± 12 U–Pb, SIMS Magmatism Vejelyte(2011)

MLD 40 Mikoliskes1

55 48 N/2128 E

Charnockite 1814 ± 20 U–Pb, SIMS Magmatism Vejelyte(2011)

MLD 41 Virbaliskis434

54 54 N/2344 E

Metavolcanic 1842 ± 6 U–Pb zrTIMS

Magmatic Motuza(2005)

MLD 42 Vidmantai1

55 53 N/2108 E

Enderbite 2.4 Ga, �Nd(1.9) −2.4

1815 ± 20 Ma U–Pb, SIMS Magmatism Claessonet al. (2001)and Mansfeld(2001)

West Lithuanian Granulite domain (WLG)G 43 Rusne-1 55 16 N/21

18 EGranite 1810 U–Pb zr

TIMSMagmatic S. Claesson

(pers.commun.)WLG 44 Lauksargiai

255 09 N/2208 E

Metapeliticgranulite

2.0, 1.97,1.94

U–Pb zrSIMS

Magmatic,detrital

This paper

WLG Lauksargiai2

55 09 N/2208 E

Metapeliticgranulite

1801 ± 21 U–Pb zrSIMS

Metamorphic This paper

WLG Lauksargiai2

55 09 N/2208 E

Metapeliticgranulite

1833 ± 31 mz EPMA Metamorphic Skridlaiteet al. (2014)

WLG Lauksargiai2

55 09 N/2208 E

Metapeliticgranulite

1796 ± 1 U–Pb mzSIMS

Metamorphic Claessonet al. (2001)

WLG 45 Lauksargiai5

55 09 N/2209 E

Metapeliticgranulite

2.37,eNd(1.9)−3.1

Mansfeld(2001)

WLG 46 Poceai Pc3 55 30 N/2130 E

Metapeliticgranulite

2.31,eNd(1.9)−2.2

2.02, 1.98,1.92, 1.87

U–Pb zrSIMS

Magmatic,detrital

This paper;Mansfeld(2001)

WLG 47 Bliudziai150 55 15 N/2258 E

Metasedimentaryrock

2.0, 1.94,1.9, 1.89,1.85

U–Pb zrSIMS

Magmaticdetrital

Krzeminskaet al. (2009)

WLG Bliudziai150 55 15 N/2258 E

Metasedimentaryrock

1853 ± 24 mz EPMA Metamorphic Skridlaiteet al. (2014)

Mazowsze domain (MD)MD 48 Burglow 53 46 N/22

51 EMetavolcanic 1835 ± 28 Ma U–Pb zr

SIMSMagmatic Wiszniewska

et al. (2007)MD 49 Tajno IG5 53 42 N/22

52 EMetavolcanic 1831 ± 29 Ma U–Pb zr

SIMSMagmatic Wiszniewska

et al. (2007)MD 50 Monki 1 53 26 N/22

41 EMafic volcanic 1829 ± 8 Ma U–Pb zr

SIMSMagmatic Krzeminska

et al. (2005)MD 51 Monki 2

(819)53 26 N/2241 E

Metasedimentaryrock

2.0, 1.99,1.97, 1.9,1.87, 1.83

U–Pb zrSIMS

Magmatic,detrital

Williamset al. (2009)

MD Monki 2(819)

53 26 N/2241 E

Metasedimentaryrock

1824 ± 14 U–Pb zrSIMS

Metamorphic Williamset al. (2009)

MD Monki 2 53 26 N/2241 E

Granitoid 2.06 Claesson andRyka (1999)

MD 52 Jastrzebna1(598)

53 45 N/2314 E

Metasedimentaryrock

2.0, 1.99,1.93, 1.86

U–Pb zrSIMS

Magmatic,detrital

Williamset al. (2009)

MD Jastrzebna1(598)

53 45 N/2314 E

Metasedimentaryrock

1827 ± 20 U–Pb mzSIMS

Metamorphic Williamset al. (2009)

MD Jastrzebna1 (600)

53 45 N/2314 E

Pegmatite 1826 ± 12 U–Pb zrSIMS

Magmatic Krzeminskaet al. (2007)

30 S. Bogdanova et al. / Precambrian Research 259 (2015) 5–33

Appendix A. (Continued)

Domain,belt

No. onFig. 3A

Well, drillcore

Coordinates Rock Sm–Ndwhole rockTDM Ga,�Nd(T)

Age, Madetritalzirconpeaks in Ga

Methoda/mineralzircon (zr),monazite(mz)

Origin ofmineral

Reference

MD 53 Tluszcz IG 1 52 29 N/2137 E

Granitoid 2.09 Claesson andRyka (1999)

MD 54 Rajsk 52 51 N/2313 E

Granite 1826 ± 6 U–Pb zrSIMS

Magmatic Krzeminskaet al. (2007)

MD 55 Pietkowo 52 54 N/2253 E

Granite 1818 ± 15 U–Pb zrSIMS

Magmatic Krzeminskaet al. (2005)

MD 56 Lomza IG1 53 11 N/2211 E

Orthoamphibolite 2.09 1802 ± 9 U–Pb zrSIMS

Magmatic Wiszniewskaet al. (2007),Claesson andRyka (1999)

MD 57 Okuniew 52 16 N/2118 E

Granodioriticgneiss

1800 ± 6 U–Pb zrSIMS

Magmatic Valverde-Vaquero et al.(2000)

MD 58 OstrowMazowieska

52 53 N/2144 E

Granitoid 2.0 Claesson andRyka (1999)

Dobrzyn domain (DD)DD 59 Prabuty 1 53 47 N/19

11 EGranitoid 2.09 Claesson and

Ryka (1999)

Pomorze domain (PM)PM 60 Slupsk IG1 54 22 N/16

54 EGranitoid 1.99 Claesson and

Ryka (1999)

Kaszuby belt (KS)KS 61 Hel IG 1 54 43 N/18

38 E2.04 Claesson and

Ryka (1999)

Sarmatia-related beltsBelarus-Podlasie Granulite Belt (BPG)BPG 62 Mosty 8 53 31 N/24

41 EEnderbite 2.3,

�Nd(1.9)−0.5

1884 ± 7 U–Pb zrTIMS

Magmatic Claessonet al. (2001)

BPG Mosty 8 54 31 N/2441 E

Enderbite 1786 ± 15 U–Pb mzTIMS

Metamorphic Claessonet al. (2001)

BPG 63 Ivje106 53 53 N/2534 E

S-type anatecticgranite

2.3,�Nd(1.9)−1.3

1860–1880 U–Pb, SIMS Magmatic/metamorphic

Claessonet al. (2001)

BPG Ivje 106 53 53 N/2534 E

S-type anatecticgranite

1770 U–Pb zrSIMS

Metamorphic Claessonet al. (2001)

BPG 64 Ivje 192 53 54 N/2535 E

Metagabbro 2.25,�Nd(1.8)+0.8

1802 ± 2 U–Pb zrTIMS

Magmatic Bogdanovaet al. (1994)

BPG Ivje 192 53 54 N/2535 E

Metagabbro 1800 ± 7 U–Pb zrTIMS

Metamorphic Bogdanovaet al. (1994)

BPG 65 Slonim 18 52 58 N/2507 E

Mangerite 2.08,�Nd(1.85)+0.10

1800 ± 10 U–Pb zrTIMS

Magmatic Shcherbaket al. (1990)

BPG 66 Radzyn IG 1 51 47 N/2237 E

Granitoid 2.01 Claesson andRyka (1999)

BPG 67 Wigry IG 1 54 03 N/2306 E

Granitoid 2.25 Claesson andRyka (1999)

Okolovo (-Rudma) belt (OKL)OKL 68 St-10 53 39 N/26

50 EMetadacite 2.03,

�Nd(2.0)+3.3

1998 ± 10 U–Pb zrTIMS

Magmatic Bibikovaet al. (1995)and Claessonet al. (2001)

OKL 69 Zh 24 53 45 N/2643 E

Enderbite 1950 U–Pb, SIMS Magmatic Unpublished

OKL Zh 24 53 45 N/2643 E

Enderbite 1795 U–Pb, SIMS Metamorphic Unpublished

OKL 70 S.Bor 3 54 11 N/2710 E

Metadiorite 2.08,�Nd(2.0)+2.1

1903 ± 2 U–Pb zrTIMS

Magmatic Bibikovaet al. (1995)and Claessonet al. (2001)

OKL 71 St 204 53 36 N/2642 E

Granite 1900 ± 37/45 U–Pb zrTIMS

Magmatic Shcherbaket al. (1990)

OKL 72 St 31 53 07 N/2612 E

Metapelite 1777 ± 2 U–Pb mzTIMS

Metamorphic Claessonet al. (2001)

a SIMS, secondary ion microprobe mass spectrometry; TIMS, thermal ionization mass spectrometry; LA, laser ablation inductively coupled plasma mass spectrometry;EPMA, electron probe micro-analyzer.

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