EUROBRIDGE: new insight into the geodynamic evolutionof the East European Craton

SVETLANA BOGDANOVA1, R. GORBATSCHEV1, M. GRAD2, T. JANIK3,

A. GUTERCH3, E. KOZLOVSKAYA4, G. MOTUZA5, G. SKRIDLAITE6, V. STAROSTENKO7,

L. TARAN8 & EUROBRIDGE AND POLONAISE WORKING GROUPS�

1Department of Geology, Lund University, Solvegatan 12, SE-223 62 Lund, Sweden

(e-mail: Svetlana.Bogdanova@geol.lu.se)2Institute of Geophysics, University of Warsaw, Pasteura 7, 02-093, Warsaw, Poland

3Institute of Geophysics, Polish Academy of Sciences, Ks. Janusza 64, 01-452 Warsaw, Poland4Department of Geophysics, University of Oulu, FI-90014 Oulu, Finland

5Department of Geology and Mineralogy, Vilnius University, Ciurlionio 21/27, LT-2009 Vilnius, Lithuania6Institute of Geology and Geography, T. Sevcenkos 13, LT-2600 Vilnius, Lithuania

7Institute of Geophysics, NAS Ukraine, Palladin Ave., 32, 03680 Kiev, Ukraine8Institute of Geochemistry and Geophysics, Kuprievich 7, 220141 Minsk, Belarus

Abstract: The Palaeoproterozoic crust and upper mantle in the region between the Ukrainian and Baltic shields of the East European

Craton were built up finally during collision of the previously independent Fennoscandian and Sarmatian crustal segments at

c. 1.8–1.7 Ga. EUROBRIDGE seismic profiling and geophysical modelling across the southwestern part of the Craton suggest that

the Central Belarus Suture Zone is the junction between the two colliding segments. This junction is marked by strong deformation

of the crust and the presence of a metamorphic core complex. At 1.80–1.74 Ga, major late to post-collisional extension and magmatism

affected the part of Sarmatia adjoining the Central Belarus Zone and generated a high-velocity layer at the base of the crust. Other

sutures separating terranes of different ages are found within Sarmatia and in the Polish–Lithuanian part of Fennoscandia. While

Fennoscandia and Sarmatia were still a long distance apart, orogeny was dominantly accretionary. The accreted Palaeoproterozoic

terranes in the Baltic–Belarus region of Fennoscandia are all younger than 2.0 Ga (2.0–1.9, 1.90–1.85 and 1.84–1.82 Ga), whereas

those in Sarmatia have ages of c. 2.2–2.1 and 2.0–1.95 Ga. Lithospheric deformation and magmatism at c. 1.50–1.45 Ga, and Devonian

rifting, are also defined by the EUROBRIDGE seismic and gravity models.

The East European Craton (EEC) is the coherent Precambrian(mainly Archaean and Palaeoproterozoic) part of Europe thatoccupies the northeastern half of the continent. Geodynamic researchin this region, however, is hampered by the presence of an extensiveyounger sedimentary cover. Geological study of the EEC thereforecommenced in the two shields of exposed Precambrian crust, theBaltic (also Fennoscandian) Shield in the NW and the UkrainianShield in the SW. From the mid-1930s onwards, study of theseshields was complemented by geophysical surveys of the RussianPlatform, and subsequently in the 1940s and thereafter by deepdrilling into the basement.

By the 1970s, large parts of the Baltic Shield had been studied insufficient detail to allow the first attempts at plate-tectonicsinterpretation (Hietanen 1975). Despite the success of thisapproach, it was evident that the shields alone were not largeenough to fit the scales of plate-tectonic processes. Attentiontherefore shifted increasingly to the covered parts of the EEC, par-ticularly the region between the Baltic and Ukrainian shields. Inthat region, geophysical reconnaissance had indicated the presenceof large arcuate, mainly NE-trending rock units and structures in

the crystalline basement beneath the Phanerozoic and Meso-to Neoproterozoic sedimentary cover. The latter is commonly1–2 km thick, but locally reaches c. 10 km or more; forexample, in the Pripyat Trough and along the margin of theEEC towards the Trans-European Suture Zone (TESZ).

Major integrated geophysical and geological projects in thesouthwestern part of the EEC, discussed since the late 1980s, weregiven new impetus by discoveries concerning the structure and evol-ution of the entire EEC that were first presented at the EUROP-ROBE symposium in Jablonna, in 1991 (Bogdanova 1993). Thenew work demonstrated that the system of Neo- to Mesoproterozoicrifts, which subdivides the craton into three parts, was superimposedon previously unknown late Palaeoproterozoic sutures locatedwhere three independent crustal segments (Fennoscandia, Sarmatiaand Volgo-Uralia) were inferred to have collided to form theEEC (Bogdanova 1993; Gorbatschev & Bogdanova 1993b;Bogdanova et al. 1996, 2005; Khain & Leonov 1996). Thus, theEUROBRIDGE project was designed to test fundamental hypoth-eses regarding the formation of the EEC with a focus on the junctionzone between Fennoscandia and Sarmatia.

Since the EUROBRIDGE project mostly focused on a regioncovered by sedimentary deposits and the Baltic Sea, geophysicalstudies along a transect extending from the Baltic Shield in south-eastern Scandinavia to the vicinity of the Black Sea came to serveas its backbone. The work comprised seismic refraction profilingassociated with wide-angle reflection and other, mainly potential-field, geophysical studies. These were integrated with extensivegeochronological, geochemical and geological investigations,the last focusing particularly on the conditions of metamorphismand their variation in time.

In this review, emphasis is placed mainly on the integration ofgeology and geophysics. Although reports on the various EURO-BRIDGE seismic profiles, other geophysical studies, and most of

�EUROBRIDGE & POLONAISE Working Groups: V. N. Astapenko, A. A.

Belinsky, R. G. Garetsky, G. I. Karatayev, V. V. Terletsky, G. Zlotski (Belarus);

S. L. Jensen, M. E. Knudsen, H. Thybo, R. Sand (Denmark); K. Komminaho,

U. Luosto, T. Tiira, J. Yliniemi (Finland); R. Giese, J. Makris (Germany);

A. Cecys, J. Jacyna, L. Korabliova, V. Nasedkin, G. Motuza, A. Rimsa, R. Seckus

(Lithuania); W. Czuba, E. Gaczynski, M. Grad, A. Guterch, T. Janik, P. Sroda,

M. Wilde-Piorko (Poland); E. Bibikova (Russia), S. Bogdanova, R. Gorbatschev,

S. Claesson, S.-A. Elming, C.-E. Lund, J. Mansfeld, L. Page, K. Sundblad

(Sweden); J. J. Doody, H. Downes (UK); V. B. Buryanov, T. P. Egorova, T. V.

Il’chenko, O. M.Kharitonov, D. V. Lysynchuk, O. V. Legostayeva, I. B. Makarenko,

V. D. Omel’chenko, M. I. Orlyuk, I. K. Pashkevich, V. M. Skobelev, L. M. Stepa-

nyuk (Ukraine); G. R. Keller, K. C. Miller (USA).

From: GEE, D. G. & STEPHENSON, R. A. (eds) 2006. European Lithosphere Dynamics.

Geological Society, London, Memoirs, 32, 599–625. 0435-4052/06/$15.00 # The Geological Society of London 2006. 599

the isotope-geological data have been published independently(Bogdanova et al. 2001a), this synthesis also identifies crustalstructures and upper mantle irregularities that can be related topost-collisional processes and the subsequent thinning of thecrust as a result of post-collisional extension and magmatism.The sites and roles of suture zones and other boundary faults are assessedand discussed, and isotope geochronology, geochemistry and the con-ditions of metamorphism are employed to define major accretionaryand collisional events.

Geological background

The assembly of several rock belts in the region between the Balticand Ukrainian shields (Figs 1 and 2) has been interpreted to be due

to tectonic stacking of the Palaeoproterozoic crust (Gorbatschev& Bogdanova 1993a). Throughout this region, lower crustalgranulites have been juxtaposed with upper and mid-crustalamphibolite-facies rocks.

The Palaeoproterozoic ages and juvenile nature of this crusthave been assessed by geochronological reconnaissance studiesacross the various tectonic units. Several isotopic methods wereemployed, including U–Pb on zircons and monazites, Ar/Ar onamphiboles, and Sm–Nd model ages (Puura & Huhma 1993;Bogdanova et al. 1994, 2001b; Bibikova et al. 1995, 2001;Claesson & Ryka 1999; Valverde-Vaquero et al. 2000; Claessonet al. 2001; Mansfeld 2001; Dorr et al. 2002; Puura et al. 2004;Soesoo et al. 2004; Krzeminska et al. 2005). The resultsdemonstrate that virtually no Archaean crust was involved inthe Palaeoproterozoic processes in the Baltic–Belarus region,

Fig. 1. Major tectonic subdivisions of the

crust in the western part of the East European

Craton: CBSZ, Central Belarus Suture Zone;

KP, Korosten Pluton; LLDZ, Loftahammar–

Linkoping Deformation Zone; MLSZ,

Mid-Lithuanian Suture Zone; O-J,

Oskarshamn–Jonkoping Belt; PDDA,

Pripyat–Dniepr–Donets Aulacogen; PKZ,

Polotsk–Kurzeme fault zone. The dashed

light yellow line delimits the

Volyn–Orsha Aulacogen. Red lines show the

positions of the EUROBRIDGE (EB’94,

EB’95, EB’96 and EB’97), Coast and

POLONAISE (P4 and P5) seismic

profiles. The inset shows the three-segment

structure of the East European Craton

(Bogdanova 1993; Khain & Leonov 1996).

S. BOGDANOVA ET AL.600

except in the form of Archaean detrital zircons in Palaeoprotero-zoic metasediments. Thus, the Palaeoproterozoic crust of theBaltic Shield appears to continue southwards to the vicinity ofthe Meso- to Neoproterozoic Volyn–Central Russian Aulacogen,where Fennoscandia meets the Palaeoproterozoic margin ofSarmatia (Fig. 1).

The recently established age patterns suggest multistagedeformation and metamorphism during several accretionaryand stacking events, the latter affecting and reactivating alreadyexisting Palaeoproterozoic crust. It consists of several Palaeopro-terozoic terranes (Figs 1 and 2), belonging both to Fennoscandia(the Okolovo, Lithuanian–Belarus and Polish–Lithuanian

Fig. 2. Major lithotectonic units of the crust

in the EUROBRIDGE study area. The

location of the refraction and wide-angle

reflection DSS profiles are also

indicated. (a) Magnetic anomaly patterns in

the region (modified after a map provided by

S. Wybraniec, Polish Geological Institute).

(b) Tectonic domains and belts. B-I,

Borisov–Ivanovo Belt; BPG, Belarus–

Podlasie Granulite Belt; BZ, Berdichev

Zone; CBSZ, Central Belarus Suture Zone;

CnZ, Ciechanow Belt; DD, –Dobrzyn

Domain; EL, East Lithuanian Belt; FSS,

Fennoscandia–Sarmatia suture; KP,

Korosten Pluton; MD, Mazowsze Domain;

MLSZ, Mid-Lithuanian Suture Zone; Mz,

Mazury plutonic rocks; Ok, Okolovo terrane;

OMB, Osnitsk–Mikashevichi Igneous Belt;

PD, Podolian Domain; PDD, Pripyat–

Dniepr–Donets Aulacogen; Tt, Teteriv Belt;

VD, Volyn Domain; VG, Vitebsk Granulite

Domain; WLG, West Lithuanian Granulite

Domain. The dashed light yellow line delimits

the Volyn–Orsha Aulacogen. Black lines

show the position of the EUROBRIDGE

(EB’94, EB’95, EB’96, EB’97) and

POLONAISE (P4 and P5) seismic profiles.

The inset shows the three-segment subdivision

of the East European Craton (Bogdanova

1993; Khain & Leonov 1996),

and the EUROBRIDGE study area.

EUROBRIDGE 601

terranes) and to Sarmatia (the Osnisk–Mikashevichi Igneous Beltand Teterev–Belaya Tserkov belt). The Fennoscandia- andSarmatia-related terranes participating in the wide CentralBelarus Suture Zone are separated by the Minsk Fault, which isinferred to be the major lithospheric discontinuity in the EEC(Bogdanova et al. 1996; Taran & Bogdanova 2001).

The precise ages of the latest stages of formation of the belt-shaped tectonic pattern in the region were assessed by Ar/Arwork on newly grown amphiboles in mylonites along shearzones that separate the various rock belts. The results suggest asurprisingly late stage of metamorphism at 1.71–1.67 Ga through-out the study region (Bogdanova et al. 2001b). Similar ages haverecently also been obtained by the Sm–Nd method for garnetsfrom the granulites in southern Estonia (Puura et al. 2004).

Numerous faults transect and complicate the collisional struc-tures (Figs 1 and 2). The dominant NE to NNE trends of the meta-morphic belts and their bounding faults were mostly formed bycollisional tectonics before 1.80 Ga, but were reactivated later,between 1.8 and 1.7 Ga. Another set of NW- to WNW-strikingfaults may be related to similarly oriented, c. 1.8 Ga shear zonesin southern Finland and central Sweden (Beunk & Page 2001;Bergman et al. 2004).

A third important group are west–east-oriented faults and defor-mation zones, some of which are accompanied by Mesoproterozoicmafic and granitoid intrusions (Beunk & Page 2001; Bogdanovaet al. 2001b; Cecys et al. 2002; Skridlaite et al. 2003b; Cecys2004). These zones clearly cut the various rock belts and offsetthe pre-existing tectonic patterns. Bogdanova (2001) proposed torefer this deformation and the attendant magmatism to the ‘Dano-polonian’ orogeny, which affected the southwestern margin of theEEC at c. 1.50–1.40 Ga, roughly coinciding in time with the‘Hallandian’ event of Hubbard (Hubbard 1975; Soderlund et al.2002) in southwestern Sweden.

Lithotectonic units

In this section, the lithotectonic units (domains and belts) andbounding deformation zones along the EUROBRIDGE transectare considered in order from NW to SE. The metamorphicrecords (P–T– t paths) of the various terranes along this transectare important keys to the histories of their formation andinteraction (Fig. 3). The latter aspect is particularly relevant inthe tectonically complex Central Belarus Suture Zone betweenFennoscandia and Sarmatia.

Fennoscandian terranes of the Baltic–Belarus region

Potential gravity and magnetic fields show that the Precambrianbedrock of southeastern Sweden continues far to the east acrossthe Baltic Sea (Wybraniec et al. 1998; Wybraniec 1999;Bogdanova et al. 2005). Related lithologies have also beenfound in drill-holes into the basement of the island of Gotland(Sundblad et al. 1998; Sundblad & Claesson 2000).

Still farther east and SE, the crust is subdivided into several beltsand domains characterized by varying grades of metamorphism.These units form the Polish–Lithuanian, Lithuanian–Belarusand Okolovo terranes, which differ in age and tectonic position(Figs 1 and 2).

The Polish–Lithuanian terrane comprises the West LithuanianGranulite Domain, which continues southwards into theMazowsze Domain, the Ciechanow Belt and probably also theDobrzyn Domain in the crystalline basement of central Poland(WLG, MD, CnB and DD in Fig. 2). With regard to their evolutionand crustal ages, between 1.85 and 1.80 Ga, these units are inmany ways similar to each other; they also resemble the rockcomplexes in southeastern Sweden.

To the east of the Polish–Lithuanian terrane, the somewhatolder (c. 1.90–1.87 Ga) Lithuanian–Belarus terrane (Fig. 2), iscomposed of amphibolite-facies East Lithuanian Belt (EL) andthe granulitic Belarus–Podlasie Belt (BPG). They developednearly simultaneously with the crust in the classical SvecofennianDomain of the Baltic Shield.

Between the Polish–Lithuanian and Lithuanian–Belarus terranesis the c. 50 km wide Mid-Lithuanian Suture Zone (MLSZ in Figs 1and 2), across which the gravity and magnetic patterns differgreatly, the tectonic grain trending more or less west–east in thePolish–Lithuanian terrane and NNE–SSW in the Lithuanian–Belarus terrane (Skridlaite & Motuza 2001). Substantial differencesare also found with regard to the thickness of the crust and its P-wavevelocity and density images. Some rocks within the MLSZ arestrongly deformed equivalents of those in the adjoining terranes,and there is a rather sharp metamorphic break between moderate-pressure western granulites and high-pressure amphibolite-faciesrocks in the east. Thus, the MLSZ represents a major deformationzone along which the West Lithuanian Domain and the entirePolish–Lithuanian terrane appear to have been thrust eastwardsover the Lithuanian–Belarus terrane (Skridlaite et al. 2003a).

The Polish–Lithuanian terrane: the West Lithuanian Granulite Domain(WLG). According to evidence from numerous drill-cores(Skridlaite & Motuza 2001; Motuza 2005), the WLG is made upof granulite-facies para- and orthogneisses, metamorphosed inthe lower crust at depths of 30–40 km between c. 1.85 and1.80 Ga. The gravity and magnetic fields as well as the geometriesof rock distribution suggest the dominance of WNW–ESE struc-tures, but locally these trends have been rotated to align with largeNE–SW- and west–east-striking lineaments (Skridlaite & Motuza2001; Motuza 2005). In general, the structural patterns of the crustin the WLG resemble those in southeastern Sweden, but itsnorthernmost part, and farther north in Latvia, is occupied by themajor AMCG-type, c. 1.6 Ga, Riga pluton (Ramo et al. 1996).(Here and elsewhere, AMCG is the abbreviation for anorthosite–mangerite–charnockite–(rapakivi)-granite magmatic suites (afterEmslie et al. 1994).)

To judge from chemistry, the metasedimentary granulites havebeen mostly formed from marine pelites, whereas the protolithsof the orthogneisses were mostly intermediate and felsiccalc-alkaline, island-arc type magmatic rocks. The isotopic agesof the detrital zircons and the Sm–Nd isotopic characteristics ofthe metasedimentary rocks (Claesson et al. 2001) suggest prove-nance from Palaeoproterozoic sources with ages between c. 2.4and 2.0 Ga. The deposition of the sediments can have takenplace at any time between 2.0 and 1.84 Ga, the latter age beingthat of the oldest charnockitic intrusions (Motuza 2005).Whereas charnockitic and somewhat younger granitoid rocks arefairly common in the WLG, they totally dominate its continuationin the Mazowsze Domain (MD) of Poland, where a granodioriteand some metavolcanic rocks have been dated at c. 1.8 Ga(Valverde-Vaquero et al. 2000; Krzeminska et al. 2005). TheSm–Nd model ages of the MD rocks (Claesson & Ryka 1999)allow correlation with similar rocks in the TransscandinavianIgneous Belt (TIB) of Sweden (Hogdahl et al. 2004). The major,roughly north-south-trending, positive magnetic anomaly, whichaccompanies this belt in Scandinavia, appears to turn towardsthe SE beneath the Baltic Sea (Wybraniec 1999; Bogdanovaet al. 2005). This supports its continuation into Poland.

In the western part of the WLG, charnockitic and enderbiticmagmatism, and related peak metamorphism of metapelites andfelsic granulites, took place at c. 1.85 to 1.80 Ga (U–Pb zirconion probe (NORDSIM) ages. Temperatures and pressures of850–900 8C and 0.8–1.0 GPa, respectively, suggest burialdepths of 35–40 km (Fig. 3a, path 1). In the same area, asecond stage of high-grade metamorphism at c. 1.79 Ga (U–Pb

S. BOGDANOVA ET AL.602

monazite age) was related to partial melting of the metasedi-mentary and metavolcanic granulites during subsequent uplift.The temperatures were 730–850 8C at pressures of c. 0.8 GPa(Fig. 3a, path 1). A later reheating–cooling step at c. 1.64–1.61 Ga is prominent in all the P–T records from the westernWLG. During that stage, temperatures varied between 550 and700 8C at pressures around 0.6 GPa (Skridlaite et al. 2004).

In the northern part of the WLG, peak metamorphism at 760 8Cand 0.7 GPa (Fig. 3a, path 2) occurred at c. 1.62 Ga, this age esti-mate being derived from metamorphic overgrowths on igneouszircon cores in charnockites (Claesson et al. 2001). Major tecto-nothermal activity at roughly the same time is indicated also bythe c. 1.6 Ga ages of the Riga pluton.

In the central part of the WLG, P–T conditions of c. 6008C(Fig. 3a, path 2) at 0.5 GPa were reached during the metamorph-ism of volcanic and sedimentary rocks (Fig. 3a, path 3), but noisotope age data are as yet available for that event.

The tectonic setting of the WLG at 1.85–1.80 Ga resembles thatof the Svecofennian crust close to the Loftahammar–LinkopingDeformation Zone (LLDZ in Fig. 1), where back-arcrifting occurred coevally with the formation of the juvenileOskarshamn–Jonkoping Belt (Beunk & Page 2001; Mansfeldet al. 2005). However, the WLG differs from the rocks in theBaltic Shield in that all the WLG sediments appear to have beensubjected to granulite-facies metamorphism and charnockiticmagmatism.

The Mid-Lithuanian Suture Zone. In addition to reworked rocksfrom the adjoining WLG and EL terranes, the MLSZ containsnumerous c. 1.84 Ga porphyritic, predominantly andesitic anddacitic volcanic rocks; gabbroic, dioritic and tonalitic intrusionsare prominent in this zone. Such rocks also compose the south-eastern part of the WLG and probably continue into the MazowszeDomain. Island-arc type geochemical characteristics have been

Fig. 3. P–T metamorphic records of the tectonothermal evolution of the Palaeoproterozoic terranes in the EUROBRIDGE region (after Taran & Bogdanova 2001,

2003; Skridlaite et al. 2003a and new P–T work). The pressure–temperature–time (P–T– t) paths of the terrane evolution represent the following. (a) The

WLG metasedimentary and meta-igneous granulites, which demonstrate post-accretionary uplift with several steps of isobaric reheating and cooling, indicating

magmatic inter- and underplating. (b) Various records from the MLSZ. The metasediments in the east have recorded near-isothermal burial, near-isobaric

heating related to granite intrusions and uplift. In contrast, meta-igneous rocks in the west display two types of paths with different peak temperatures at the same

depths. This may be a result of their tectonic juxtaposition. Prominent steps of near-isobaric cooling and reheating were related to magmatic emplacements.

(c) Collision and subsequent uplift of the EL metasediments adjacent to the BPG, and uplift overprinted by a step of isobaric cooling. The latter was related to

granite intrusions. (d) Subduction-triggered accretion, collision and uplift of the interior and marginal part of the BPG as meta-igneous and metasedimentary rocks

show. (Note a minor step of near-isobaric cooling in the marginal parts of the BPG, which was related to Mesoproterozoic granitic magmatism.) (e) Accretion followed by

collision, near-isothermal decompression as recorded by the Okolovo metavolcanic rocks. (f ) Subduction-triggered accretion followed by magmatic emplacements, uplift

and the near-isobaric cooling as traced by metasediments of the VG Domain.

EUROBRIDGE 603

reported (Rimsa et al. 2001; Motuza 2005), and these rocks couldtherefore represent a c. 1.84 Ga rim of crustal growth in thePolish–Lithuanian terrane, similar to the 1.83–1.82 GaOskarshamn–Jonkoping Belt in southeastern Sweden (Mansfeld1996; Mansfeld et al. 2005). The former presence of an oceanicbasin between the Polish–lithuanian and Lithuanian–Belarusterranes is thus very likely.

A probable northern continuation of the MLSZ in Latvia is theIncukalns Zone, where mantle-related supracrustal rocks havebeen intruded by similarly juvenile granitoids (Mansfeld 2001).

The complex collisional structure of the MLSZ is characterizedby the presence of a variety of P–T trends in its different parts(Fig. 3b). In its eastern part (Fig. 3b, path 1), garnet-bearingfelsic gneisses have undergone amphibolite-facies metamorphismat c. 500–580 8C and 0.6 GPa (Skridlaite & Motuza 2001), andsubsequent heating to 750 8C is evidenced by migmatite-relatedanatectic granites yielding ages as young as c. 1.50 Ga (Skridlaite& Whitehouse, pers. comm.). These ages are similar to the empla-cement ages of the late- to post-collisional AMCG rocks in theMazury plutonic complex in northern Poland (Dorr et al. 2002;Skridlaite et al. 2003b).

In the western part of the MLSZ, inclusions of felsic granulitesin enderbitic and charnockitic intrusions record peak metamor-phism at c. 850–900 8C and 1.0 GPa (Fig. 3b, path 2). Togetherwith their host rocks, these inclusions underwent a secondgranulite-facies event at c. 800–850 8C and 0.8 GPa, followedby decompression and near-isobaric cooling from c. 800 to600 8C at about 0.6 GPa. Adjacent coherent units of garnet-,biotite- and hornblende-bearing metavolcanic gneisses wereinitially probably also buried to similar depths (Fig. 3b, path 3);however, their metamorphic hornblendes, yielding ages of c.1.71–1.66 Ga (Bogdanova et al. 2001b), were formed at only550–600 8C and 0.4–0.5 GPa, which suggests sites in themiddle crust. In consequence, these gneisses appear to havebeen uplifted 10–15 km after their peak of metamorphism. Byanalogy with conditions in the eastern part of the MLSZ, theisobaric stage of their metamorphic evolution was probably dueto the intrusion of the 1.5 Ga AMCG igneous suite. A subsequentevent of shearing, uplift and cooling at 1.48–1.42 Ga is well docu-mented by the 40Ar/39Ar ages of hornblendes throughout theentire southern part of the MLSZ (Bogdanova et al. 2001b).

The Lithuanian–Belarus terrane. The East Lithuanian Belt (EL) andBelarus–Podlasie Belt (BLG) (Fig. 2) making up the Lithuanian–Belarus terrane probably represent different depth levels of thesame age unit.

The EL comprises mafic metavolcanic rocks as well as sedimen-tary rocks, particularly metagreywackes and marbles (Motuza2005). Although the metagreywackes contain some Archaean det-rital zircons, Palaeoproterozoic zircons are dominant. Sm–Ndisotopic modelling indicates Palaeoproterozoic ages of deposition(Mansfeld 2001). The supracrustal rocks of the EL are marked bybanded, NNE-striking, magnetic anomalies that are clearly discor-dant in relation to the more or less west–east-trending anomalypatterns in the WLG (Skridlaite & Motuza 2001).

From the NW to the SE across the EL there are steep meta-morphic gradients from high-P rocks, presumably buriedbeneath SE-vergent thrusts, to lower-pressure amphibolite-faciesunits and, ultimately, to granulites similar to those in the neigh-bouring BPG (see below). In the easternmost EL, greywackesand pelites were buried at depths as great as 30 km and metamor-phosed at c. 680 8C and 0.8 GPa at about 1.9 Ga (Fig. 3c, path 1).From about 1.8 Ga onwards, they underwent decompression andretrogression similarly to the rocks of the neighbouring BPG(see Fig. 3d, paths 3 and 4). The P–T evolution of the easternmostEL therefore appears to have occurred roughly simultaneouslywith that of the BPG. None of these rocks seem to have beenaffected by later thermal events.

In the southern part of the central EL, interbedded metasedimen-tary and metavolcanic rocks reached their metamorphic peak atc. 700 8C and 0.6 GPa, thereafter undergoing near-isobariccooling from 580 to 480 8C at 0.2 GPa (Fig. 3c, path 2).Because these rocks are in contact with the 1.50 Ga Kabeliaigranite (Sundblad et al. 1994), their final metamorphism wasprobably under the influence of this intrusion.

The Belarus–Podlasie Granulite Belt (BPG) occupies most ofthe territory of what was previously described as the ‘Belarus–Baltic Granulite Belt’, a concept originally proposed by Aksamen-tova et al. (1982). Subsequent studies of the granulites in Latviaand Estonia (Holtta & Klein 1991; Mansfeld 2001; Puura et al.2004; Soesoo et al. 2004) have shown, however, that the Latvianand Estonian granulites differ substantially in age and provenancefrom the granulites in western Belarus and the Podlasie region ofPoland. Previously, it had been suggested that the east–west-trending Polotsk–Kurzeme belt of faulting (PKZ in Fig. 1) separ-ates the two groups of granulites from each other (Garetsky et al.2004). In the light of current insight, however, an en echelonconfiguration of two different rock belts appears more probable.

The BPG is mostly between 100 and 200 km wide and extendsfor more than 600 km in a SW–NE direction from southeasternPoland across the western part of Belarus. It is made up ofseveral large lensoid bodies of granulites, separated from eachother by fault zones and mylonites (Aksamentova & Naydenkov1990). The BPG rocks are mostly Palaeoproterozoic granuliticorthogneisses of mafic, enderbitic and charnockitic compositions.They belong to two igneous suites. The older one has an age ofc. 1.89 Ga (Claesson et al. 2001) and is calc-alkaline in compo-sition, whereas the younger (c. 1.80 Ga), is chemically more vari-able, alkali-calcic and bimodal (Bogdanova et al. 1994; Bibikovaet al. 1995; Taran & Bogdanova 2003). Metasedimentary gneissesand migmatites are relatively subordinate, but still occur in manydrill-cores. Their Sm–Nd isotopic characteristics, as well as thoseof the intrusive rocks, indicate a minor contribution of oldermaterials (Claesson et al. 2001).

The P–T history of the BPG, adjacent to the northwestern partof the Central Belarus Suture Zone (CBSZ), suggests a sequenceof tectonothermal events resembling that found in the Okolovorocks of the latter (see below). Early prograde metamorphism,recognized in the metasedimentary rocks, occurred in amphibolitefacies at conditions varying from 530–550 8C and 0.3–0.4 GPa to650–670 8C and 0.6–0.7 GPa. It was associated with deformationand tonalite–trondhjemite–granodiorite (TTG)-type magmatism,apparently related to subduction to depths of c. 30 km between1.89 and 1.87 Ga (Fig. 3d, paths 1 and 2).

Granulite-facies metamorphism at 1.8 Ga (750 8C andc. 0.8 GPa) was superimposed onto these amphibolite-faciesrocks, being caused by the input of numerous mafic and monzoni-tic to charnockitic intrusions, following late to post-collisionalextension of the crust (Bogdanova et al. 1994; Taran &Bogdanova 2003). Rapid tectonic uplift to upper crustal levelsand cooling took place between 1.78 and 1.74 Ga. Between 1.71and 1.66 Ga, there was continued extension associated with retro-grade metamorphism, strong deformation, and transtensionalrearrangement of the tectonic grain of the BPG into lensoid andanastomosing patterns (Bogdanova et al. 1994). The last majortectonothermal event in the BPG involved reactivation alongpreviously formed zones of faulting and granitic intrusions at1.6–1.5 Ga (Bogdanova et al. 2001b). This evolution is indicatedby the stepwise configuration of some of the P–T paths (e.g. path 4in Fig. 3d).

Comparison of the P–T paths from the different parts of theBPG suggests that the rock units in its interior (Fig. 3d, paths 1and 2) were uplifted and cooled more rapidly than those situatedalong the margins. Two instances of the latter are the Rudmarocks along the boundary towards the Okolovo terrane (Fig. 3e,path 2) and the rock units in the westernmost part of the BPG,along its boundary with the EL (Fig. 3d, paths 3 and 4).

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Altogether, the available data suggest that the BPG and EL wereboth formed at c. 1.9 Ga, but in widely different tectonic settings.Whereas the BPG represents a mature island arc, the formation ofthe EL most probably occurred in a back-arc environment.

The Central Belarus Suture Zone (CBSZ)

The CBSZ extends nearly 600 km along the northwestern marginof Sarmatia. Although it coincides geographically with the tra-ditional ‘Central Belarus Belt (CB)’, EUROBRIDGE studieshave shown that it is heterogeneous in structure and comprisesseveral tectonically different rock complexes separated, in particu-lar, by the major, suture-like, Minsk Fault (see below). To the NWof this fault is the Okolovo terrane; to the SE are located theVitebsk Domain and the southeasternmost part of the formerCB, referred to in the following as the Borisov–Ivanovo Belt(VG and B–I, respectively, in Figs 1 and 2).

The Okolovo terrane. The c. 2.0 Ga Okolovo terrane (Bibikovaet al. 1995; Claesson et al. 2001) forms a c. 10 km thick,WNW-dipping, tectonically delimited complex (Aksamentovaet al. 1994). Along its borders towards the overlying BPG in theNW, the rocks have been metamorphosed in the granulite facies.At the base of the Okolovo terrane in the area immediately adja-cent to the Minsk Fault, geophysical evidence suggests thatthere may exist a separate lens of granulites, isolated structurallyfrom the rest of the Okolovo terrane.

The Okolovo terrane is built up of metamorphosed komatiitic andtholeiitic, basalts, andesites, dacites and rhyolites of oceanic-arc affi-nities. Intercalated metasedimentary rocks include characteristicblack shales and ferruginous as well as siliceous volcanogenic depos-its. Although these igneous rocks are present throughout the Okolovoterrane, they are particularly abundant in its northernmost part.

The metavolcanic rocks of the Okolovo terrane (Fig. 3e, path 1)were exposed to prograde metamorphism with peak temperaturesranging from 640 to 725 8C at depths of 30–35 km. These con-ditions were attained at c. 1.9 Ga, concomitantly with the empla-cement of TTG-type juvenile melts (Claesson et al. 2001), mostprobably related to subduction-triggered metamorphism (Taran& Bogdanova 2001). The subsequent rapid, tectonic uplift ofsome crustal blocks to depths of only 12–15 km was followedby extension and high-T metamorphism associated with theinferred intrusion of mafic and monzonitic–charnockitic melts atc. 1.8 Ga. This is reflected by the deviation of some P–T pathsinto the field of higher temperatures and by their nearly isobarictrends during metamorphism in amphibolite facies. Retrogressionlasted from c. 1.77 to 1.67 Ga (Fig. 3e, paths 2 and 3).

The close relationships of the Okolovo and BPG (i.e.Lithuanian–Belarus) terranes are of great importance, militatingagainst a previous view of the Okolovo terrane as part of theSarmatian plate (Bogdanova et al. 2001b). All these terranesshared similar geodynamic evolutions after 1.9 Ga. This period,in contrast, is not characteristic of the Palaeoproterozoic historyof Sarmatia (see below).

The Vitebsk Domain (VG) and the Borisov–Ivanovo (B–I) belt. TheVG and the B–I belts, lying to the SE of the Minsk Fault(Fig. 2), are very different in age and evolution from theOkolovo terrane. These two units are characterized by tectonic set-tings and development histories very similar to those of theOsnitsk–Mikashevichi Igneous Belt (OMB), farther within Sar-matia (see below); they are interpreted as probable equivalentsof the OMB, which developed at deeper crustal levels. In bothunits, juvenile metasediments and andesitic–dacitic metavolcanicrocks were formed at c. 1.98 Ga (Bibikova et al. 1995).

Regarding metamorphism, pelitic xenoliths enclosed in 1.96 Gahigh-temperature granites in the B–I and the VG document anearly episode of regional metamorphism at P–T conditions of

c. 610 8C and 0.6–0.7 GPa. Subsequently, but still before theintrusion of the granites, some uplift occurred, which caused retro-gression (Fig. 3f). This records the burial of metapelites to depthsclose to 25 km and their subsequent uplift. A later event of contactmetamorphism, which was associated with the 1.96 Ga granites,involved heating to c. 670 8C. Substantial cooling and decompres-sion followed the initial stages of this magmatism, possibly stillcoinciding with the upward migration of the granitic melts.For the final stage of the metamorphic evolution, the recordedP–T– t path indicates cooling from c. 530 to 420 8C at pressuresdropping from 0.4 to 0.2 GPa.

The metamorphic and igneous histories of the rocks in the VGagree in many respects with those recorded from the adjacentOMB. Within the latter, some supracrustal rocks were metamor-phosed in the high-T range of high amphibolite facies at tempera-tures up to 700 8C and pressures of 0.3–0.4 GPa (Khvorova et al.1982). This event may have been caused by elevated heat flow atthe active margin of Sarmatia and recurrent igneous activity in theOMB.

West Sarmatian terranes

The Sarmatian crustal segment, which is exposed in the UkrainianShield and the Voronezh Massif, is built up of several Archaeanproto-cratonic terranes and intervening belts of Palaeoproterozoicrocks. The continental crust of the Archaean units was formedbetween c. 3.7 and 2.7 Ga, whereas that in the Palaeoproterozoicbelts was accreted to the Archaean cores mainly between 2.2and 2.1 Ga, and again between 2.0 and 1.9 Ga.

The EUROBRIDGE’96 and ’97 profiles traverse three majorlithotectonic units: the Palaeoproterozoic OMB, the similarlyPalaeoproterozoic, but somewhat older Teterev Belt in theVolyn Domain (also known as the North-Western Domain in theUkraine), and the northern part of the mostly Archaean PodolianDomain (Figs 1 and 2). The Palaeoproterozoic belts of Sarmatiadiffer in age from those of Fennoscandia by featuring 2.2–2.1 Ga crust, which is seemingly absent in northeastern Europe(see Claesson et al. 2006).

The Osnitsk–Mikashevichi Igneous Belt (OMB). This 150–200 kmwide belt occupies the northwestern margin of the Sarmatiancrustal segment. From the Trans-European Suture Zone alongthe southwestern limit of the EEC, it extends northeastwards toMoscow (Bogdanova et al. 2004a); that is, for a distance ofmore than 1000 km. It is buried almost entirely beneath Phanero-zoic sedimentary rocks (Aksamentova 2002), cropping out only inan area in the northwestern part of the Ukrainian Shield and in afew horst-type elevations within the Pripyat Trough (the westerncontinuation of the Devonian Dniepr–Donets Aulacogen). Never-theless, the OMB can be traced fairly confidently in the magneticfields, where it is marked by numerous rounded, mostly positiveanomalies associated with large batholiths of granodiorites andgranites, and intrusions of diorites and gabbros. The most reliableages of these emplacements range from 1.98 to 1.95 Ga. Betweenand inside the plutonic massifs are ‘septa’ and inclusions of maficand felsic metavolcanic and hypabyssal rocks, for which an ageof 2.02 Ga has been obtained (Skobelev 1987; Shcherbak &Ponomarenko 2000). In addition, there also exists a later,c. 1.80–1.75 Ga, generation of sub-alkaline plutonic and volcanicrocks, associated with some sedimentary complexes deposited inminor basins. These are nearly coeval with the AMCG KorostenPluton in the Volyn Domain and other similar intrusions in theUkrainian Shield.

In general, the lithologies and structures of the OMB suggestformation in an Andino-type active continental-margin environ-ment, created by c. 2.0–1.95 Ga subduction of oceanic crustbeneath the edge of Sarmatia. A period of apparent quiescencethen followed, for which age and other information are almost

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totally lacking. This period lasted until the beginning of collisionbetween Sarmatia and Fennoscandia at c. 1.84–1.82 Ga.

Kinematic analysis of faults within and in the neighbourhood ofthe OMB (Gintov 2004) indicates their formation or activationsimultaneously with the emplacement of the Korosten plutonbetween c. 1.80 and 1.74 Ga (Amelin et al. 1994). During thisperiod, the present structural patterns of the crust in northwesternSarmatia were generally established (Bogdanova et al. 2004b).

The Volyn Domain (VD). The VD consists mainly of Palaeoproter-ozoic rocks (Figs 1 and 2). Dominant granitoids were emplaced at2.06–2.02 Ga, concomitantly with high-T amphibolite faciesmetamorphism (Khvorova et al. 1982) and migmatization ofolder (2.2 Ga), strongly deformed, sedimentary and volcanicrocks of the Teterev Belt and its Belaya Tserkov continuationinto the adjacent Ros’–Tikich Domain (Stepanyuk et al. 1998;Claesson et al. 2000; Claesson et al. 2006). Palaeogeographicalreconstructions of the stages of formation of these supracrustalunits (Lazko et al. 1975; Ryabenko 1993; Shcherbakov 2005) indi-cate that at c. 2.2 Ga the entire northwestern part of the UkrainianShield was characterized by intense igneous activity and sedimen-tation in coastal and marine settings. The sedimentary successionsare terrigenous and include tuffitic and graphitic units and varioustypes of turbidites. Basaltic, andesitic, dacitic and rhyolitic vol-canism also occurred. Altogether, this suggests mature island-arcconditions and possibly derivation of detritus from Archaeansources. However, strong deformation, involving SE-vergentthrusting and associated transcurrent faulting during the for-mation of the OMB at 1.98–1.95 Ga, prevents more detailedpalaeotectonic interpretation.

Numerous intrusions of mafic to monzonitic sub-alkaline rockswere emplaced into zones of extension at 2.02–1.98 Ga; that is,almost simultaneously with the early stages of igneous activityin the OMB.

A most remarkable feature of the Volyn Domain is the giantAMCG-type Korosten pluton (Figs 1 and 2) that was formedbetween 1.80 and 1.74 Ga by the successive emplacement ofmany pulses of basic and acidic melts (Zinchenko et al. 1990;Bukharev 1992; Amelin et al. 1994; Verkhogliad 1995). Tra-ditionally, the Korosten pluton has been considered as anorogenic,formed by mantle underplating. However, it has recently beenrelated to the late and post-collisional tectonic regimes prevailingfarther NW, in the CBSZ (Bogdanova et al. 2004b). It is alsonotable that, according to recent Sm–Nd isotopic data, none ofthe Korosten igneous rocks, with 1Nd(T) values of 20.8 to21.8, were derived from a depleted mantle (Dovbush et al.2000). This supports the idea that the Korosten magmas originatedby remelting of the OMB lower crust (Bogdanova et al. 2006).

The nearly 100 km wide Berdichev Boundary Zone (BZ inFig. 2) separates the Teterev Belt from the Archaean PodolianDomain farther south. Originally, this zone may have belongedto a Palaeoproterozoic collisional boundary belt between thePodolian Domain and crustal units to the north. It is mostlymade up of high-T, upper crustal, garnet- and cordierite-bearingS-type granites, which, towards the north, pass into the 2.06–2.02 Ga granitoids of the Volyn Domain (Shcherbakov 2005).In the southern BZ, Neoarchaean charnockites intrude older,Archaean granulites (Stepanyuk et al. 1998). P–T conditionswithin the Berdichev Zone, in its deepest crustal sections in thesouth, reached c. 850 8C at c. 8 GPa (Kurepin 2003).

The Podolian Domain (PD). The PD is one of the principalArchaean units of the Sarmatian crustal segment (Claesson et al.2006). Archaean as well as Palaeoproterozoic granulites prevail.Major zones of faulting subdivide it into the Vinnitsa region inthe north and the Gayvoron region in the south. The boundary inthe east, towards the Palaeoproterozoic Kirovograd Domain, isdefined by the Golovanevsk Suture Zone, which features

numerous, nearly 3.0 Ga ultramafic and mafic rocks (Gornostayevet al. 2004).

EUROBRIDGE’97 seismic profiling particularly concerned thenorthern part of the PD, where granulites of sedimentary and maficvolcanogenic derivation are the dominant rocks. These werecomplicated structurally by doming associated with the formationof c. 2.1–1.9 Ga Palaeoproterozoic charnockites, granites andmigmatites. Archaean charnockitic intrusions are widespread inthe SE.

The supracrustal granulitic rocks of the Podolian Domainbelong to two associations. The oldest granulites are mafic, ultra-mafic and intermediate, and have been intruded by enderbites(Shcherbak et al. 2005). Sm–Nd data and ion-probe zirconstudies suggest ages of the crust around 3.7 Ga and an event ofgranulite-facies metamorphism at c. 2.8 Ga (Claesson et al.2006). Kalyaev et al. (1984) have proposed that the oldest maficrocks could represent early Archaean oceanic crust.

The apparently younger metasedimentary rocks include ferrugi-nous quartzites as well as highly aluminous, partly graphiticschists, and carbonates, all metamorphosed in granulite facies,presumably during the Palaeoproterozoic. This suggests that theearly Proterozoic tectonothermal evolution was fairly uniformthroughout the entire western part of the Ukrainian Shield.The now-exposed rocks in western Sarmatia may represent a10–15 km thick slice of Archaean- to Palaeoproterozoic crust.The granulites in the southern part of the PD were exhumedfrom depths of more than 35 km (Kurepin 2003; Shcherbakov2005).

Major stages of the Proterozoic crustal evolution

Accretionary stages

The pre-metamorphic composition of the lithotectonic units andtheir P–T evolution, as described above, suggest that, during thetime before the final assembly of the EEC, the terranes alongthe northwestern margin of Sarmatia evolved differently fromthose related to Fennoscandia. The Palaeoproterozoic metasedi-ments in the southeastern, that is Sarmatia-related part of themain suture zone, the CBSZ, were metamorphosed during subduc-tion at c. 1.96 Ga; neither this part of the suture zone, nor north-western Sarmatia in general, has yielded consistent evidence oftectonothermal activity during the period between c. 1.95 and1.80 Ga (Taran & Bogdanova 2001; Shcherbak et al. 2003).Thus, the metasedimentary rocks in the Vitebsk Domain (VG)were deformed and metamorphosed at c. 1.96 Ga in the middlecrust beneath the Osnitsk-Mikashevichi Belt (OMB) along theSarmatian continental margin. The prograde part of their P–Tpath (Fig. 3f) can be attributed to southeastwards (in present-daycoordinates) subduction, which took place simultaneously with theOMB magmatism at some time after 1.98 Ga, the earliest agedefined by the youngest detrital zircons in the VG (Claessonet al. 2001). In the supracrustal rocks of the Teteriv Belt fartherinside Sarmatia, however, an early stage of high-T metamorphism,intense migmatization and anatectic remelting occurred already atc. 2.1–2.0 Ga. This major tectonothermal event in western Sarma-tia locally also influenced the evolution of the Archaean PodolianDomain, causing repeated granulite metamorphism and associatedhigh-T magmatism.

In contrast, the Fennoscandia-related terranes, farther NW,mostly developed between 1.9 and 1.8 Ga, simultaneously withthe formation of the Svecofennian Domain proper and theoldest, 1.83–1.82 Ga rocks in southeastern Sweden. An exceptionis provided by the Okolovo terrane within the CBSZ, which con-tains c. 2.0 Ga oceanic-arc supracrustal rocks, which have not beenidentified elsewhere in the Svecofennian region. As seen fromtheir P–T– t path (Fig. 3e), the Okolovo supracrustal rocks weresubducted northwestwards (present-day coordinates) beneath the

S. BOGDANOVA ET AL.606

Lithuanian–Belarus terrane and, after 1.9 Ga, evolved togetherwith the Belarus–Podlasie Belt (BPG) and East Lithuanian Belt(EL) of this tectonic unit. The prograde parts of the metamorphicpaths of the Okolovo terrane and the BPG, and probably also someof the supracrustal rocks in the EL, imply a period of burial anddeformation between 1.90 and 1.87 Ga associated with TTG mag-matism. These data contradict a previous tectonic interpretation ofthe EL supracrustal rocks as an accretionary prism (Motuza 2005).If the latter was correct, subduction should have occurred towardthe present SE; that is, in a direction opposite to that suggested bythe evidence presented above.

In the West Lithuanian Granulite Domain (WLG) and the part ofthe Mid-Lithuanian Suture Zone (MLSZ) closest to the Polish–Lithuanian terrane, the earliest orogenic events are difficult todiscern. These lithotectonic complexes are characterized particu-larly by a stage of accretionary tectonics in the Baltic–Belarusregion between 1.85 and 1.82 Ga. During that subduction and theformation of volcanic island arcs and back-arcs, TTG magmatismand burial of supracrustal rocks at great depths occurred. Their evol-ution is described by the preserved prograde parts of the P–T– tpaths of some rocks of the MLSZ rocks (Skridlaite et al. 2003a).

Comparison of the tectonothermal events during the periodbetween c. 2.2 and 1.84 Ga (Fig. 4) thus suggests the followingevolution.

(1) Terranes related to Sarmatia and Fennoscandia, along theEUROBRIDGE transect, evolved separately until these twocrustal segments docked with each other between 1.84 and1.80 Ga.

(2) At c. 1.90–1.87 Ga, the Lithuanian–Belarus terrane and the2.0 Ga Okolovo oceanic arc were assembled to form a single unit,possibly a microcontinent. This process took place simultaneouslywith the formation of the classical Svecofennian Domain in theBaltic Shield.

(3) The development of the Polish–Lithuanian terrane wasnearly coeval and possibly even closely connected spatially withthe accretion of new crust in southeastern Sweden. In and closeto the Oskarshamn–Jonkoping Belt, subduction-related magma-tism, the formation of an island arc and a back-arc (Mansfeldet al. 2005), and nearly coeval metamorphism and partialmelting of meta-supracrustal gneisses (Beunk & Page 2001)occurred at c. 1.83–1.78 Ga.

Collision and post-collisional stages of extension

Metamorphism in the Baltic–Belarus region largely took place atc. 1.8 Ga and was broadly coeval with widespread mafic andmonzonitic magmatism, represented particularly well in theBPG, the EL and the Okolovo terranes. From geophysical data,the structure of the upper lithosphere in all these tectonic units

is ‘thick-skinned’ and thus characteristically collisional, but wascomplicated by subsequent extension (see below).

The giant Korosten Pluton in the Volyn Domain (VD) andthe presence of a thick high-velocity layer in the lowercrust beneath the CBSZ (see below) both suggest that high-Tgranulite-facies metamorphism, anatectic melting and near-isobaricretrograde cooling of most of the studied rocks were related to lateand post-collisional tectonism between 1.80 and 1.75 Ga.

The newly grown c. 1.80 Ga metamorphic zircons and monazitesin the granulites of the WLG and MLSZ indicate that a coeval tec-tonothermal event also affected the Fennoscandia-related terranesfarther NW. This was associated with 1.82–1.81 Ga charnockiticintrusions (Skridlaite & Motuza 2001). However, it is still conjec-tural whether all the terranes in the Baltic–Belarus region had beenassembled by 1.82–1.80 Ga. Possibly an oceanic ‘gap’ may stillhave existed at that time between the Polish–Lithuanian and theLithuanian–Belarus terranes (Skridlaite et al. 2003a).

Major NW-trending transpressive shearing at 1.83–1.78 Gaalso occurred in the crust of central and southeastern Sweden(Beunk & Page 2001; Hogdahl et al. 2004), coinciding with anearly event of magmatism in the Transscandinavian IgneousBelt (TIB). In that region, subduction-related magmatismand attendant metamorphism along the present southwestern andwestern margins of the Svecofennian orogen may have occurred(Andersson et al. 2004).

The subsequent, 1.71–1.67 Ga tectonothermal event (Fig. 4)may have been due either to the terminal collision and amalgama-tion of the Fennoscandian, Volgo-Uralian and Sarmatian crustalsegments to form the East European Craton, or to within-cratondeformation and reactivation of the pre-existing fault systemscaused by continuing convergence (Bogdanova et al. 2001b). Atroughly the same time (c. 1.73–1.67 Ga), a large north–south-striking belt of juvenile crust was formed as a result of eastwardsubduction in southwestern Sweden and central Norway (Anders-son et al. 2004; Gorbatschev 2004). Notably, this event did notsignificantly affect the Svecofennian Domain in Finland, wheremajor deformation and metamorphism ceased at 1.79–1.77 Ga(Lahtinen et al. 2005).

Mesoproterozoic intra-continental

deformation and magmatism

In all the ‘Fennoscandian’ terranes in the EUROBRIDGE areaand in southeastern Sweden there are numerous steep, east–west-trending zones of shearing, with associated c. 1.6–1.45 GaAMCG and A-type granitoid intrusions (Bogdanova et al.2001b; Skridlaite et al. 2003b; Cecys 2004). These may havebeen a far-field effect of continuing or renewed accretionary

Fig. 4. Timing of Proterozoic crust-forming processes in the southwestern part of the East European Craton. The principal orogenic events in each of the terranes

are shown in black. The cross-hatched lines (1, 2, 3 and 4?) indicate ‘stitching’, simultaneous events during the assemblies of the terranes and their shared

evolution as discussed in the text. Line 2 shows the approximate time of the major collision of Fennoscandia and Sarmatia. The AMCG- and A-type granitoid

igneous events (graded white–grey fill) are interpreted as indicators of late or post-collisional tectonic regimes. (See the text for the geochronological references.)

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orogeny in the westernmost Baltic Shield (Ahall et al. 2000), butthey may also have been related to collision with other plates, forexample, Amazonia (Bogdanova 2001). Somewhat unexpectedly,the P–T– t histories of metamorphism (Fig. 3) demonstrate thestrong effects of this Mesoproterozoic tectonothermal activity.Some of the ‘steps’ and indications of near-isobaric cooling,which complicate the P–T paths, suggest a relationship betweenthis younger metamorphism and the emplacement of AMCG-typemagmas, such as the c. 1.6 Ga Riga pluton and the 1.54–1.50 GaMazury complex (see Figs 3a–c and 4). The outer rims of zirconsin the metamorphic lithologies have been dated by ion probe toaround 1.50–1.45 Ga (Skridlaite et al. 2004). The AMCG magma-tism at this time also triggered growth of amphibole in mylonitesalong the major shear zones and resetting of the Ar isotopesystems of older amphiboles in various rocks (Bogdanova et al.2001b).

Seismic and density images of the crust and

upper mantle in the southwestern EEC

EUROBRIDGE and POLONAISE seismic profiling

The seismic models of the southwestern part of the EEC presentedhere are based on EUROPROBE, EUROBRIDGE and POLO-NAISE deep seismic sounding (DSS) profiles that were acquiredin the 1990s (Giese 1998, Guterch et al. 1998, 1999; EURO-BRIDGE Seismic Working Group 1999; Czuba et al. 2001, 2002;Yliniemi et al. 2001; Grad et al. 2003b; Thybo et al. 2003). Datafrom FENNOLORA and the Coast Profile (Lund et al. 2001)have also been employed (Figs 1, 2 and 5). The study area has

never previously been investigated seismically using modern tech-niques. The high-quality seismic data now obtained reveal theP-wave and, in some cases, the S-wave structures of the crust andthe uppermost mantle.

Seismic profiles: acquisition of data and the observed wave field. TheEUROBRIDGE and POLONAISE’97 profiles, altogetherc. 2300 km in length, were carried out using modern digitalseismic recorders spaced 1.2–4.0 km apart along the profiles.Shot points with charges of 300–1000 kg of TNT were locatedat intervals of 30–40 km. In the Coast Profile project, a shipborneairgun array was used to generate the seismic waves.

The wave field recorded in the southwestern part of the EEC is,in general, of very high quality (Figs 6 and 7). Because of the com-monly thin sedimentary cover (,2000 m), the refracted wavesdiving into the crust (Pg) produced clear first arrivals with apparentvelocities of c. 6–7 km s21. Strong reflected waves from the Mohoboundary (PmP), starting from the offsets at 80–120 km, andsubstantial differentiation of the arrival times, exceeding 2 s forPg, PmP and Pn phases, were observed. This reflects differentiationof the crust- and upper mantle structure.

Models of the crust and uppermost mantle. The seismic data for allthe POLONAISE’97 and EUROBRIDGE profiles were modelledusing 2D tomographic and ray-tracing techniques (Cerveny &Psencık 1983). P-wave velocity models of the crust and uppermostmantle along the profiles in the southwestern part of the EECare shown in Figures 8 and 9. The S-wave velocity model andthe Vp/Vs ratio distribution for the EUROBRIDGE’97 profileare shown in Figure 9. For the tectonic units related to Fennoscandiaand Sarmatia, respectively, the particulars of the structures based onthe P- and S-wave velocities are summarized in Tables 1 and 2. Ingeneral, the P- and S-wave velocity models both show similarity tothe results previously obtained from Scandinavia (Grad & Luosto1987, 1994; BABEL 1993; Guggisberg et al. 1991).

The crust and upper mantle of the southwestern part of the EECcan be characterized as follows.

(1) The thickest Phanerozoic cover deposits in the various partsof the Pripyat Trough and along the Trans-European SutureZone correspond to seismically defined layers with P-wave vel-ocities between 2 and 4 km s21.

(2) The crystalline crust in the surveyed area can generally bedivided into three parts, with P-wave velocities of 6.1–6.4, 6.5–6.8 and 6.9–7.2 km s21 for the upper, middle and lower crust,respectively. Relatively low velocities of c. 5.7 km s21 in theuppermost crystalline basement were found locally in the MDand WLG, and in the BPG (Figs 1 and 2). The upper crystallinecrust is commonly inhomogeneous, with low-velocity zones andhigh-velocity bodies alternating along some parts of the profiles.Normally, the rather weak, low-velocity zones reach c. 5 km inthickness and have velocity contrasts of 0.1–0.2 km s21; mostly,they occur at depths between 4 and 15 km. Low-velocity layersin the upper crust have been found in the TIB, WLG, and in theVD and PD of Sarmatia, whereas the region of the Mazury intrusions(Mz), the CBSZ, and the VD with the Korosten Pluton (KP) featurehigh-velocity bodies in the upper crust. The middle crust has P-wavevelocities of 6.5–6.8 km s21, which increase to between 6.9 and7.2 km s21 in the lower crust. The lowermost crust is marked byhigh P-wave velocities, reaching a maximum of 7.5 km s21 in thepart of the Volyn Domain underlying the KP in Sarmatia. Character-istically, that region lacks the high reflectivity in the lower crust thatis otherwise common in the VD.

The crystalline crust in the EUROBRIDGE region mostly has lowvelocity gradients and small velocity contrasts at the seismic bound-aries. Only in some places, such as, for instance, in the CBSZ, theOMB, and parts of the VD and the PD, has high reflectivity beenobserved.

(3) The average values of the Vp/Vs ratios in the crystallinecrust are 1.69, 1.70 and 1.76 in its upper, middle and lowerparts, respectively. It follows that the S-wave velocities in the

Fig. 5. Location of the refraction and wide-angle reflection DSS profiles in the

area of southwestern margin of the East European Craton. Open stars indicate

the shot points of the EUROBRIDGE (EB’94, EB’95, EB’96 and EB’97) and

POLONAISE’97 (P5 and northern part of P4) profiles. The numbered stars refer

to the locations of shot points for which examples of the record sections are

shown in Figures 6, 7 and 9. The bold dashed line shows the southwestern

edge of the craton; the fine dashed line shows the Fennoscandia and Sarmatia

suture. TESZ, Trans-European Suture Zone.

S. BOGDANOVA ET AL.608

upper and middle crust are relatively high in comparison with theP-wave velocities, whereas in the lower crust Vs is relatively low.This may explain the strong SmS reflections from the Moho, seen,for instance, in the central part of the EUROBRIDGE’97 profile(Fig. 9).

(4) High-velocity plutonic bodies in the upper crust along theEUROBRIDGE transect coincide with geologically well-knownintrusions such as the Mazury and Korosten plutons of rapakivi-granitic and gabbro–anorthositic rocks (Figs 1 and 2). In theMazury igneous complex, a high-velocity body with P-wavevelocities between 6.4 and 6.7 km s21 coincides with a gabbro–anorthosite massif (Fig. 8). The Vp/Vs ratio in this body is esti-mated to be 1.75. In the VD, the Korosten Pluton is imaged ascoinciding with a high-velocity anomaly of 6.35–6.7 km s21,extending to depths of at least 11 km. At still greater depths, thisanomaly appears to link up with another high-velocity anomalyin the lower crust where the Vp/Vs ratios are as high as 1.77–1.79.

(5) Most of the crust in the EUROBRIDGE study area has athickness between 40 and 50 km. Moho depths of c. 55 km havebeen found in the PD, whereas the shallowest Moho (c. 30 km)is that in the VD, beneath the Korosten Pluton. Mantle P-wave vel-ocities immediately beneath the Moho are generally 8.2–8.35 km s21; lower velocities (8.0–8.15 km s21) have beenfound only in the marginal zones of the EEC such as, for instance,in the Dobrzyn Domain in Poland (Figs 1 and 2). The average Vp/Vs

ratio for the uppermost mantle, determined from the Pn and Snwaves, is 1.75, with Fennoscandia having a lower average ofc. 1.72 and Sarmatia a somewhat higher value of c. 1.80 (Fig. 9).

(6) The uppermost mantle features numerous subhorizontalreflectors beneath both the Baltic Shield and the East

European Platform (Grad 1992; Sroda & POLONAISE ProfileP3 Working Group 1999; Czuba et al. 2001; Lund et al. 2001).These reflectors often are c. 10–15 km below the Moho. Amajor, SSW-dipping reflector has been recognized in the upper-most mantle beneath the EUROBRIDGE’97 profile (Figs 8 and9). It extends from the Moho to depths of c. 75 km (Thybo et al.2003). This reflector coincides with a subhorizontal reflector onthe EUROBRIDGE’96 profile, close to its crossing point withthe EUROBRIDGE’97 profile, in Sarmatia. We note, however,that the quality of data on the EUROBRIDGE’97 profile ishigher than that on the EUROBRIDGE’96 profile, and the errorof the Moho reflection positions at the crossing point isc. 10 km. The formation and age of this reflector are discussedbelow.

Gravity–seismic modelling of rock compositions

along the EUROBRIDGE profiles

The lateral and vertical variations of P-wave velocity in the crustand upper mantle along the EUROBRIDGE transect agree, ingeneral, with the global compilation of data on seismic velocitiesand composition of continental crust presented by Christensen &Mooney (1995). Seismic velocities and densities along all theEUROBRIDGE profiles, which increase gradually with depth,can be explained in terms of rock compositions that changefrom felsic to mafic, and also with increasing grades of meta-morphism. The P-wave velocities compiled in Table 1 indicatethat the silica content in the rocks of the EUROBRIDGE areadecreases from over 70% SiO2 in the upper crust to about 47%

Fig. 6. P-wave record sections for the

seismic profiles in the area along the

southwestern margin of the East

European Craton (for the locations of shot

points see Fig. 5). Pg, PmP and Pn are

crustal and Moho phases.

EUROBRIDGE 609

at the Moho boundary (Kern et al. 1993; Christensen & Mooney1995). However, the interpretation of P-wave velocities interms of rock compositions is non-unique, as a given velocityvalue can usually characterize a number of different litho-logies (Christensen & Mooney 1995). Additional constraints onrock composition can be obtained from Vp/Vs ratios (or Poisson’sratios); these are generally sensitive to the chemical compositionof the rocks (Kern et al. 1993; Sobolev & Babeyko 1994).However, correlation of Poisson’s ratios with SiO2 contents isvalid only for rocks with 55–75% SiO2, that is for felsic andintermediate compositions (Christensen 1996). For mafic rocks(SiO2 , 55%), the Poisson’s ratios vary greatly and there is nocorrelation with SiO2 content. In consequence, estimation ofthe compositions of the lower crust from the Vp/Vs ratio is notreliable.

Because of the strong dependence of rock density on compo-sition (Birch 1961), the composition of the crust can be inferredfrom the former. Thus gravity-field modelling and inversion canprovide additional information for the interpretation of seismicresults in terms of rock compositions. Another important advan-tage of gravity data is that they are more sensitive to verticaland subvertical crustal boundaries than DSS data. In conventionaltechniques of gravity data interpretation, the density models arecalculated from seismic velocity models using density–velocityrelationships estimated under laboratory conditions (Woollard1959; Birch 1961; Ludvig et al. 1970; Christensen & Mooney1995). For the EUROBRIDGE profiles, however, the integratedvelocity–density models were calculated using a technique ofjoint interpretation of seismic and gravity data that is based on a

stochastic, quasi-linear relationship between density and seismicvelocity, which is not assumed a priori, but rather obtained fromthe inversion of gravity data (Kozlovskaya et al. 2002).This approach makes it possible to obtain density models of thecrust and also analyse the relationships between density andseismic velocities in the major crustal units traversed by theEUROBRIDGE profiles.

The advantage of analysing the density–velocity relationship indetail is illustrated by Figure 10, which shows the relationshipsbetween density, Vp, Vs and Vp/Vs estimated by Monte-Carlosimulation for several types of rocks from the Fennoscandianand Ukrainian Shields. These represent major crustal lithologiesand contain various proportions of the main rock-forming min-erals. The rocks considered and their modal mineral compositionshave been described in a previous paper (Kozlovskaya et al. 2004).As can be seen from these data, both the igneous and themetamorphic lower crustal rocks have similar values of Vp/Vs,whereas their densities are significantly different. This indicatesthat the density–velocity relationships (and, in particular,the density–Vp relationship) depend on the varying mineralcompositions of the rocks (Fig. 10c).

In general, therefore, the density–velocity relationships ofactual rocks with various chemical and mineral compositionswill differ from each other and deviate also from those of anhy-drous igneous rocks with averaged chemical compositions(Sobolev & Babeyko 1994) that are used as the basis for standardcurves of reference and shown by star signs in Figure 10c. Thus, forexample, rocks with high contents of plagioclase and low contentsof amphibole have higher Vp values for any given density value

Fig. 7. P-wave record sections (for

location see Fig. 5). The large differentiation

of the arrival times (exceeding 2 s) of the

Pg, PmP and Pn phases, which reflects

differentiation of the structures in the crust

and uppermost mantle, should be noted.

S. BOGDANOVA ET AL.610

than the rocks of the standard reference curves. The opposite, ofcourse, applies to rocks rich in amphibole, but poor in plagioclase.

In another instance, mafic lower crustal rocks metamorphosed ingranulite facies have densities around or below 3.0 g cm23 andtheir density–velocity relationships follow closely the standardcurves of reference. With the appearance of substantial amounts

of garnet in high-pressure mafic granulites and eclogites,however, the densities increase to above 3.0 g cm23 and the vel-ocities to over 7.0 km s21. As can be seen from Figure 9c, thedensity–velocity relationships for these rocks differ significantlyfrom those of other mafic lower crustal rocks, whereas the Vp/Vs

ratios are all similar.

Fig. 8. Crustal and uppermost mantle models for the EUROBRIDGE transect (EB’94, EB’95, EB’96 and EB’97 profiles), the Coast profile and the

POLONAISE’97 profiles P4 (northern part) and P5. P-wave velocities are given in km s21. 1, sedimentary cover; 2, upper crust; 3, high-velocity zone in the upper

crust; 4, low-velocity zone in the upper crust; 5, middle crust; 6, lower crust; 7, high-velocity lower crust; 8, uppermost mantle; 9, elements of seismic boundaries

obtained from reflected and refracted waves; 10, zones of high reflectivity in the uppermost mantle; 11, Moho boundary; 12, mantle reflector or zone of

high-velocity gradient; 13, zone of anomalously high velocity, probably associated with the Korosten Pluton; 14, zones of rapid lateral change of the seismic structure,

probably indicating contact zones of crustal blocks. V.E., vertical exaggeration.

EUROBRIDGE 611

Figure 11a and b shows the density–Vp relationships for theEUROBRIDGE EB’95 and EB’96 profiles. Obviously, substantialscatter is present, but deviations from the reference curveare small. This indicates that the upper crust consists mainly oflow-grade felsic rocks, whereas the middle crust is made up ofcompositionally intermediate rocks, their metamorphic gradeincreasing with depth from amphibolite to granulite facies. Thescatter in the density–Vp plots is explained best by compositionaldifferences of the rocks, which affect the relationships betweendensity and P-wave velocity in the various units crossed by theEB’95 and EB’96 profiles, as follows.

(1) The upper crust of the WLG domain is composed of alumi-nous metasedimentary granulites, some pyroxene-bearing

gneisses and charnockites, whereas the MLSZ contains partlymylonitized granulites, charnockites, metavolcanic rocks and pla-gioclase–microcline granites. The EL domain, in turn, is made upof rather dense biotite–plagioclase gneisses and amphibolites(Skridlaite & Motuza 2001). This explains the scattering on theVp–density plot for EB’95 (Fig. 11).

(2) In the region of the CBSZ, the P-wave velocities anddensities in the upper crust of the Okolovo terrane (6.1–6.2 km s21 and 2.7 g cm23, respectively) are higher than thosebeneath the southeastern half of this zone (5.8–6.0 km s21 and2.6–2.67 g cm23). This reflects a distinct difference betweenamphibolite- and granulite-facies oceanic-arc rocks in the NWand a terrane of large plutons of quartz syenites and sub-alkaline

Fig. 9. Two-dimensional seismic models

along the EUROBRIDGE’97 profile

developed by forward ray-tracing. Top:

example of seismic record section for SP06

with P and S waves (for location see Fig. 5).

Middle: S-wave velocity model. The parts of

the discontinuities that have been constrained

by reflected and/or refracted S waves are

marked by bold lines. Arrows show the

positions of the shot points. Bottom: Vp/Vs

ratio distribution. The parts of the

discontinuities that have been constrained by

reflected and/or refracted P or S waves are

marked by bold lines. V.E., vertical

exaggeration.

S. BOGDANOVA ET AL.612

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2.8

2.5

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71

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55

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(km

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6.7

56

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6.5

56

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(km

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57

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(km

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58

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58

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EUROBRIDGE 613

Table 2. Modal mineralogy for selected rock types used for gravity–seismic modelling (values in per cent)

Qtz Kfs Pl(An20) Pl(An60) Cpx Opx Ol Amph

Igneous rocks

Granite 30–40 30–40 20–40 0 0 0 0 0

Rapakivi granite 17–27 46–53 14–24 0 0–1 0 0–1.5 2–6

Anorthosite 0 0–2 0 90–95 0–2 0–2 0–2 0–2

Gabbro–norite 0–2 0–10 0 43–65 10–24 0–20 0–23 0–5

Gabbro 0 2–3 0 50–60 7–10 6–14 10–30 3–6

Metamorphic rocks

Amphibolite 0–4 0 30–40 5–10 0 0 0 55–60

Mafic granulite 0–4 0 0 40–50 4–20 8–12 0 20–30

Mafic garnet granulite 0 0 0 17–41 5–10 4–10 0 19–30

Eclogite 0 0 0 0 5–41 6–16 3–15 13–57

Mineral abbreviations (after Kretz 1983): Qtz, quartz; Kfs, K-feldspar; Pl, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; Ol, olivine; Amph, amphibolite; Grt,

garnet; An, anorthite. An in Pl is shown in parentheses.

Fig. 10. Simplified density models along

EUROBRIDGE’95–96 profiles (a) and

EUROBRIDGE’97 profile (b) (after

Kozlovskaya et al. 2001, 2002, 2004). The

positions of the major tectonic units are

indicated as in Figure 2. (c) Relationships

between density and seismic velocities

obtained by Monte-Carlo simulation for

selected types of igneous and metamorphic

rocks from the Ukrainian Shield and other

Precambrian areas. The modal mineralogy of

rocks is given in Table 2.

Left: density–Vp relationship; centre:

density–Vs relationship (for Vs, the axis is

scaled by a factor of 1.73); right:

density–Vp/Vs relationship. The reference

density–velocity relationships are shown by

open stars.

S. BOGDANOVA ET AL.614

granites with subordinate metasedimentary and volcanogenicgneisses and migmatites in the SE.

The principal difference between the EB’95 and EB’96 profilesis in the composition of the lower crust. Along profile EB’95, thelower crust has P-wave velocities of 6.8–6.9 km s21 and densitiesof 2.9–3.0 g cm3, which are values typical of mafic granulites.Profile EB’96, however, in addition, shows a major high-velocitylayer at the base of the crust, which has P-wave velocities of7.1–7.4 km s21 and densities of 3.0–3.1 g cm23, and extendsbeneath the southern parts of the BPG, the CBSZ, and the OMB.Beneath the WLG and EL areas, such a high-velocity layer ismissing.

The high-velocity layer crossed by the EB’96 profile is alsovisible in the density–velocity diagram of Figure 11. Its S-wavevelocity and density values are typical of garnet-bearing maficrocks formed under high-pressure conditions. As seen in thefigure, some deviation from the reference curve towards lowervelocity values may be due to elevated plagioclase contents.

As already considered, joint interpretation of seismic andgravity data based on P-wave velocities and densities meets theproblem that rocks with high contents of calcium-rich plagioclasehave higher P-wave velocities and thus high Vp/density ratios,which do not plot on the standard reference curves. For thisreason, interpretations based on the relationship between densityand P-wave velocity have not been applied to the EB’97 profilethat intersects the Korosten gabbro–anorthosite–rapakivi granitepluton (KP) with its commonly extremely plagioclase-rich rocks.Because isotropic S-wave velocities appear to be generallybetter correlated with density than P-wave velocities, and areless affected by high contents of plagioclase, the interpretationof the gravity data along EB’97 was performed using the relation-ship connecting density to both the P- and the S-wave velocities(Thybo et al. 2003; Kozlovskaya et al. 2004).

The combined velocity and density model of EB’97 demon-strates pronounced lateral variations of these properties, whichcan be related spatially to the extents of the geological unitscrossed by this profile, that is the OMB, the VD and PD, and theKP (Table 1). In addition, the relationships between density andthe seismic velocities Vp and Vs in the geological units crossedby the EB’97 profile have been obtained. As can be seen fromFigure 12, these differ from each other and deviate also from thecorresponding reference curve relationships. Generally, thedensity–Vp and density–Vs relationships for the OMB are closeto those of the reference curves for the upper and middle crust,but are shifted upwards from the reference curve for the lowercrust (Fig. 12), indicating that the entire crust of the OMB prob-ably consists of igneous rocks (i.e. granites, granodiorites and

gabbros in different proportions). Beneath the OMB,the additional lower-crustal high-velocity layer with Vp, Vs anddensity values of 6.9–7.2 km s21, 3.95–4.0 km s21 and 2.95–3.1 g cm23, respectively, is seen in both the EB’96 and EB’97 pro-files. This layer is probably composed of mafic garnet granulitesand eclogites.

The most complicated structure of the crust along the EB’97profile is found in the Palaeoproterozoic VD, which is composedof rocks of very different character and origin. In this case, thedensity–Vp plots for the VD and KP scatter around the referencecurve (Figs 10 and 12). Furthermore, the points appear to shiftupwards from that curve for densities less than 3.0 g cm23,which characterize most of the igneous rocks of the KP (i.e. therapakivi granites, gabbro–norite–anorthosites and gabbro–norites). The values of Vp, Vs and density determined for theupper crust of the KP at depths to 5 km suggest that this body con-sists of both rapakivi granites and anorthosites, but the resolutionof the data for the EB’97 profile is not good enough to model thestructure in detail. The points closest to the reference curve in thedensity–Vp plot (Fig. 12) have been obtained for a high-velocitybody in the northern part of the KP that has high values of boththe P-wave and S-wave velocities (6.4–6.46 km s21 and 3.66–3.72 km s21, respectively), and densities of 2.75–2.80 g cm23 atdepths of 5–12 km. This combination of Vp, Vs and densityvalues can be attributed to metamorphic rocks (Lebedev &Korchin 1982; Lebedev et al. 1983), which probably belong tothe host rocks of the KP.

The lowermost crust along the EB’97 profile (Figs 9 and 10) hasP-wave velocities of 7.0–7.6 km s21 and densities over3.0 g cm23; that is, values similar to those of the high-velocitylayer at the base of the crust beneath the OMB and the CBSZ.This suggests that it may be partly composed of mafic garnet gran-ulites. However, the density–Vp relationship in this crustal unit(Fig. 12) indicates high contents of plagioclase, as a result of thepresence of igneous gabbro–anorthositic rocks in the lower crust.

For the upper and middle crust, the density–velocity ratios inthe PD plot close to the reference curves. For the lowercrust, however, they shift slightly upwards (i.e. towards higher vel-ocities). The latter indicates granulite-facies conditions of meta-morphism despite the relatively great crustal thicknesses ofmore than 50 km. These would necessitate pressure–temperatureconditions in the lower crust that correspond to eclogite facies.However, the density values of the lower crust in the area arebelow 3.0 g cm23, suggesting only minor contents of garnet butrelatively high contents of plagioclase, and implying that thelower crust had not been eclogitized extensively. The reasonmay have been its anhydrous condition (Austrheim et al. 1997).

Fig. 11. Comparison of the relationships

between density and Vp for the

EUROBRIDGE’95 and ’96 profiles.

EUROBRIDGE 615

Integrated geological–geophysical interpretation

of the structure of the upper lithosphere along the

EUROBRIDGE profiles

General structural characteristics of the crust

The interpretative models of the Earth’s crust along the EURO-BRIDGE EB’95, EB’96 and EB’97 seismic profiles presented inFigures 13 and 14 have been compiled from the overall geophysi-cal and geological information, referred to in the previoussections.

Collision between Fennoscandia and Sarmatia was decisive indetermining the seismic characteristics of the lower crust andupper mantle in the study region, and the distribution of themagnetic and gravity anomalies (e.g. Garetsky et al. 2002).Pre-collisional terrane tectonics, in contrast, is reflected best by

structures in the upper and middle crust. Tectonically, the rockbelts and domains in the Fennoscandian terranes make up anumber of ‘thick-skinned’ nappe packages, thrust towards theSSE and SE in the southern part of the Baltic–Belarus region,but towards the NE in Estonia and the area of Lake Ladoga(Fig. 1). Subsequently, these nappes were transected by sets ofNNW–NW- and NNE–NE-trending post-collisional faults andthe markedly east–west-striking Mesoproterozoic faults.

The EUROBRIDGE profiles suggest that the formation of high-velocity layers in the crust was commonly associated with detach-ment, whereas lateral undulations may have been shaped bycomplementary deformation of the whole lithosphere (Figs 8,10, 13 and 14). Almost all of the high-velocity layers areaccompanied by distinct subhorizontal seismic reflectors andmark sharp compositional discontinuities in the crust. In manycases, mafic sheet intrusions were responsible or contributed.

Fig. 12. Relationships between density, Vp

and Vs for the major geological units crossed

by the EUROBRIDGE’97 profile (after

Kozlovskaya et al. 2004). The

left panel shows the density–Vp relationship,

the right the density–Vs relationship (for Vs,

the axis is scaled by a factor of 1.73). The

reference density–velocity relationships are

shown by open stars. (a) Osnitsk–

Mikashevichi Igneous Belt; (b) Podolian

Domain; (c) Volyn Domain and Korosten

Pluton.

S. BOGDANOVA ET AL.616

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EUROBRIDGE 617

Complex crustal and upper mantle structures characterize theFennoscandia–Sarmatia junction area beneath the CBSZ andpart of the BPG, where the more ancient and more rigid crust ofSarmatia has been particularly strongly deformed, laminated andalso altered compositionally. As modelling of the gravity and mag-netic data shows, a substantial proportion of the post-collisionaland later post-orogenic faults have listric configurations (Garetskyet al. 2002). In a number of cases, their flat-lying deeper partscoincide with nearly horizontal reflectors (detachment zones) atthe boundary between the upper and middle crust. However,some large faults (e.g. the Minsk Fault in the CBSZ) appear toextend to the Moho (Figs 2 and 13).

Geophysical images of accretionary and

collisional tectonics

Terranes related to the Fennoscandian crustal segment. The Polish–Lithuanian terrane with the West Lithuanian Granulite Domain(WLG) and the Mid-Lithuanian Suture Zone (MLSZ) resembles inmany respects the crustal province of southeastern Sweden, situatedto the south of the classical Svecofennian orogen. In that province,the rocks were formed during several orogenic events at c. 1.84–1.82, 1.81–1.78 and c. 1.75 Ga, tending to young towards thesouth. The structural trends are dominantly WNW–ESE to west–east, and there exist active-margin volcanic-arc and back-arc

Fig. 14. The integrated tectonic interpretation of EUROBRIDGE’97 profile. Top: seismic model as in Figures 8 and 9. Middle: gravity–seismic model as in

Figure 10b. Bottom: the tectonic model. V.E., vertical exaggeration.

S. BOGDANOVA ET AL.618

type supracrustal belts (Sundblad et al. 1998; Beunk & Page 2001;Mansfeld et al. 2005). In southeastern Sweden, the primary,pre-collisional or pre-accretionary relationships of the different litho-tectonic complexes are recognized fairly well in the seismic images.These indicate subduction towards the NNE (BABEL WorkingGroup 1993; Abramovitz et al. 1997; Balling 2000).

In Figure 13, steep dips towards the NW are indicated for theMLSZ and the adjoining area, whereas other studies, particularlythose of metamorphism (see p. 603), have suggested that theWLG had been thrust towards the east, overriding the East Lithua-nian Belt (EL). Skridlaite et al. (2003a) inferred that the wide-spread occurrence of high- to moderate-pressure granulites inthe WLG indicated a subduction–collision tectonic regimebetween 1.84 and 1.80 Ga, whereas island-arc settings have beenidentified both in the MLSZ and the adjoining parts of the WLG(Rimsa et al. 2001; Skridlaite & Motuza 2001; Motuza 2005).The crustal thickness in the WLG ranges between 45 and 50 km,whereas the crust atop the uplifted Moho in the MLSZ has athickness of only c. 40 km. Similar differences also characterizethe lower crust, which measures 10–12 km in the WLG, butonly 5–10 km in the WLSZ (see Kozlovskaya et al. 2001;Yliniemi et al. 2001; Grad et al. 2003b). The thicker lower crustbeneath the WLG appears largely to be due to the presence of abasal crustal layer with densities as high as 3.0 g cm23 and thusprobably composed of mafic granulite (see Christensen &Mooney 1995). This layer has no equivalent in the MLSZ.

The upper and middle levels of the crust have P-wave velocitiesof 6.1–6.4 and 6.5–6.8 km s21, respectively, both in the WLGand the MLSZ. However, the upper crust in the WLG is largelygranulitic, whereas that in the MLSZ is made up of variousrocks in the granulite and amphibolite facies, such as blastomylo-nites, calc-alkaline metavolcanic rocks and gabbroic to tonaliticplutonic rocks. At depths of 12–15 km (Kozlovskaya et al.2001; Yliniemi et al. 2001), the crust of the WLG and, in part,that of the MLSZ have a low-velocity layer, which most probablyconsists of Mesoproterozoic granitic rocks created by the remelt-ing of the upper crust. A large body of a c. 1.46 Ga monzogranitoidrock is present close to the northern end of the EB’95 profile(Cecys 2004; Motuza 2005).

A prominent feature of the crust particularly in the northwesternpart of the WLG is its multi-layered structure, built up of distinctlydelimited, conformable, persistent individual layers. The seismicvelocities at the base of the crust are high (Vp 8.25–8.35 km s21). Although this might suggest a ‘platformal’ type ofcrust in the sense of Christensen & Mooney (1995), lithologicaland geophysical variations are substantial within the WLG.Thus, its northern part is largely made up of orthogneisses,whereas a mixture of metasedimentary granulites, charnockites,metavolcanic rocks and various granitic rocks dominates inthe south. Major differences of rock composition and deep struc-ture also exist between the eastern and western parts of theWLG (Kozlovskaya et al. 2001). Particularly worth noting isthe recurrent granitoid magmatism both in the Palaeo- andMesoproterozoic.

The Lithuanian–Belarus terrane, including the EL and BPG,together with the Okolovo terrane forms a composite terranewhere the crust is substantially thicker (up to c. 55 km) and, as awhole, also denser than that in the Polish–Lithuanian terrane(Fig. 13). The principal mechanisms responsible for the develop-ment of this thick crust appear to have been collisional orogenicprocesses involving compression and folding, and the stackingof large piles of nappes in the junction zone between Fennoscandiaand Sarmatia. Indications of tectonic thickening, thinning, foldingand wedging-out of the rock units are common in the seismic pro-files. With regard to the thickness of the upper and middle partsof the crust, the EL and BPG are not very different from theWLG, but no low-velocity layers appear to be present. Here, theseismic velocities vary substantially in accordance with lithologi-cal variation, but as most of the rocks are either mafic to

intermediate granulites or igneous rocks of similar compositions,the upper crustal P-wave velocities are mostly relatively high.They measure c. 6.25 km s21 in the EL, 5.8–6.0 km s21 in theBPG, and 6.1–6.2 km s21 in the largely metavolcanic Okolovoterrane. Substantially lower velocities are, naturally, found in thelarge, anastomosing systems of shear zones marked by blastomy-lonites and retrograde recrystallization of the granulites, and alsothe presence of metasediments. These occur particularly in theEL. Major west-dipping listric faults that could be traced todepths of 15–20 km (see Aksamentova et al. 1994) havepreviously been found along the Grodno–Starobin seismicprofile transecting the BPG and the Okolovo terrane in a west–east direction (see also Bogdanova et al. 2001b).

In the middle crust, which has densities around 2.8 g cm23 andP-wave velocities of 6.3–6.5 km s21, granulites and TTG-typeplutonic rocks appear to predominate. This part of the crustforms a ‘trough’ beneath the BPG.

With regard to the lower crust, the BPG and EL are similar(Table 1, Fig. 13). Both have a c. 20 km thick lower-crustallayer made up of mafic granulites with P-wave velocities of6.8–7.1 km s21 and densities of 2.9–3.1 g cm23 (Fig. 10,Table 1). A remarkable feature is the southeasterly dips of thislower crustal layer, which appears to protrude into the uppermantle; its density of 3.1–3.2 g cm23 is substantially less thanthe 3.3–3.4 g cm23 of the normal upper mantle in the region.Thus, this mantle offset–lower crustal protrusion may representa ‘fossilized’ slab of subducted Palaeoproterozoic oceanic crust.

The Fennoscandia–Sarmatia junction. Within the CBSZ, theseismic and gravity characteristics of the crust change drasticallyacross the major, west-dipping Minsk Fault (Figs 10a and 13). Thelatter extends to the Moho and, at the Earth’s surface, separatestwo very different groups of tectonic units (Figs 1 and 2), theBPG and Okolovo terrane in the NW and the Vitebsk Domain(VG) and Borisov–Ivanovo Belt (B–I) in the SE. The P-wave vel-ocities and rock densities in the upper crust are different on the twosides of the Minsk Fault, being 6.1–6.2 km s21 and 2.7 g cm23 inthe NW and 5.8–6.0 km s21 and 2.60–2.67 g cm23 in the SE.This reflects the difference between the amphibolite- to granulite-facies mafic rocks of the Okolovo terrane and the granite-intrudedmetasedimentary gneisses and migmatites in the B–I. In themiddle crust, the P-wave velocities are 6.3–6.5 km s21 in theNW and 6.4–6.9 km s21 in the SE, but there is apparently no cor-responding difference in density values (Kozlovskaya et al. 2002).The best explanation appears to be that the higher P-wave vel-ocities below the southeastern part of the CBSZ are due to a mark-edly laminated structure in a part of the crust where stronglydeformed amphibolite- and granulite-facies rocks have beenemplaced tectonically.

In the CBSZ region, these diverse structural patterns in theupper to middle crust can be followed down to depths of 25–30 km, at which level the P-wave velocities reach 7.0 km s21

and densities of 2.9–3.0 g cm23 have been modelled. Fartherdown is a rather more uniform lower crustal high-velocity layer(HVLC in Figs 8, 10a and 13) with P-wave velocities of 7.2–7.4 km s21 and densities of 3.0–3.1 kg m23. These values corre-spond best to eclogitic granulites or garnet granulites. Similarhigh-velocity layers in the lower crust, with P-wave velocitiesbetween 7.0 and 7.7 km s21, appear to be common in Precambrianregions adjoining major tectonic sutures (Guggisberg & Berthelsen1987; Korja et al. 1993; Korja & Heikkinen 1995; Funck et al. 2001;Hall et al. 2002). In some cases, they are associated with sizeableMoho offsets and are mostly explained as resulting from magmaticmantle underplating (Korsman et al. 1999; Funck et al. 2001; Hallet al. 2002).

In the case of the CBSZ, the lower crustal high-velocity layer isa relatively young feature, as it appears to underlie adjacent ter-ranes as well. These include the OMB, which adjoins the CBSZin the SE, and, some 250 km farther SE, the part of the Volyn

EUROBRIDGE 619

Domain that encloses the AMCG Korosten Pluton. In view of thisspatial association, this high-velocity, high-density layer is mostprobably coeval with the c. 1.80–1.74 Ga Korosten body. Xeno-liths from the lower crust (Markwick et al. 2001) and new dataon the isotopic compositions of both OMB and Korosten rockssuggest that this high-velocity layer may be restitic, having beenformed between 1.80 and 1.74 Ga by the successive removal ofthe AMCG melts from a 2.0 Ga lower crust similar to thatbeneath the OMB (Bogdanova et al. 2006).

As indicated by the gravity–seismic modelling in Figures 10, 13and 14, many crustal units in the CBSZ are wedge-shaped andimbricated, and are separated from each other by numerous dis-tinct reflectors dipping in opposite directions. Structural patternsof this kind are characteristic of collisional-type crust (Meissner1989; Cook et al. 1999) and therefore fit well with the locationof the CBSZ region in the zone of collision between Fennoscandiaand Sarmatia.

A conspicuous major feature of the crust beneath the CBSZ andadjoining parts of the OMB and BPG is a large antiformal (domal)structure defined by convex high-velocity middle to lower crustalrock layers apparently related to the OMB (Fig. 13). As this struc-ture coincides with the collisional zone between Sarmatia andFennoscandia, it appears reasonable to assume that stacking andthickening of the crust with attendant metamorphism and gravita-tional instability must have been part of the early stages of itsdevelopment (see Coney & Harms 1984).

An important key to deciphering the formation of this antiformbeneath the CBSZ and its evolution into a metamorphic corecomplex is the presence of a bulge of high-velocity material inthe lower crustal core of the antiform (Fig. 13). In conjunctionwith other features, such as the association of the KorostenPluton with the high-velocity lower crust, this suggests thatc. 1.8 Ga magmatism could have been a major agent causing thedoming and attendant metamorphism during the post-collisionalstage. This evolution was associated with post-collisional exten-sional tectonics leading to c. 1.8 Ga AMCG magmatism in theOMB and the Lithuanian–Belarus terrane, local granulite meta-morphism, intense mylonitization along the Minsk Fault,and, eventually, listric faulting and fast final uplift (Taran &Bogdanova 2003). Also, downfaulting of the edge of the OMBto expose the Borisov–Ivanovo and the Vitebsk tectonic unitsmust have been part of this extension. The latter are connectedwith southwards dipping reflectors in the middle and lower crustof the OMB (Juhlin et al. 1996; Stephenson et al. 1996).

The Archaean and Palaeoproterozoic crust of Sarmatia. The OMBwas formed by voluminous magmatism between c. 2.0 and1.95 Ga. In accordance with the OMB igneous mode of origin,the character of its c. 50 km thick crust is determined largely bythe presence of numerous batholiths of granitic, granodioritic,dioritic or gabbroic composition. These obviously correspond tothe seismic-velocity and density properties (Figs 12 and 13),which also suggest that the more felsic of the plutons dominatethe upper crust, whereas the mafic ones prevail in its lower parts(Kozlovskaya et al. 2002, 2004; Yegorova et al. 2004).

Markwick et al.’s (2001) study of deep crustal and mantlexenoliths indicates that the mafic plutonic rocks of the OMBhave been partly transformed into eclogite-like, garnet-bearinggranulites with P-wave velocities of 6.8–7.0 km s21 and corre-spondingly high densities.

In addition, the OMB contains younger, c. 1.8 Ga, mostly syeni-tic to quartz syenitic intrusions, which are associated with thecoeval AMCG-type Korosten Pluton farther south and define abelt of marked, more or less isometric, magnetic anomalies (seeFig. 2). The distribution of these intrusions appears to have beencontrolled by major NE-trending, NW-dipping zones of faulting,which also follow some of the OMB boundaries.

The upper and middle parts of the crust in the OMB in particu-lar feature numerous major reflectors (Figs 8 and 13) that create

an overall multi-layered structure, presumably mostly caused byrecurrent magmatism and tectonic deformation, especially in thevicinity of the Fennoscandia–Sarmatia junction. Some ofthe layering, however, must rather be due to the formationof the Devonian Pripyat–Dniepr-Donets Aulacogen (Fig. 1). Asdiscussed by Stephenson et al. (1996), the system of listricfaults in that structure coincides closely with Palaeoproterozoicwedge fabrics within the OMB, which indicates significant reac-tivation of Precambrian faults during Phanerozoic rifting. Awell-preserved fine lamination of the crust also characterizesthe OMB, presumably caused by deformation of rocks with con-trasting elastic properties (Meissner & Rabbel 1999) and prob-ably related to the latest major deformation event in theDevonian. In the lower crust, a high-velocity lower crustallayer with Vp of 7.2–7.4 km s21 exists also in the OMB,similar to that beneath the CBSZ, but substantially thinner anddenser than in the latter.

In the Volyn Domain (VD) with the large AMCG KorostenPluton (KP), the crust is only 45 km thick; that is, substantiallythinner than the 50–52 km crust in the neighbouring OMB andPD (Figs 8–10 and 14). The available seismic and gravity datasuggest, however, that beyond the limits of the Korosten Pluton,the VD is similar to the OMB. All the crustal layers of the OMBappear to continue into and across that domain, extending south-wards as far as the Berdichev Boundary Zone, which dips northand separates the Palaeoproterozoic VD from the Archaeaninterior parts of the Podolian Domain (PD).

In the Berdichev Zone (BZ), the lower part of the Palaeoproter-ozoic crust appears to wedge out at depth, and the distinctlylayered upper and middle parts are replaced southwards by seismi-cally more uniform and less reflective Archaean crust. In thisancient crust there is a rather gradual increase of the P-waveseismic velocities with depth, from 6.1–6.2 to nearly 6.9 km s21.The distribution of these velocities and the Vp/Vs ratios suggest atwo-layered crust, but the S-wave data and the gravity–seismicmodelling by Kozlovskaya et al. (2004) indicate the presence ofthree crustal layers. The relatively low Vp/Vs ratios of 1.69–1.74indicate that, the Archaean crust of the Podolian Domain is muchricher in quartz than the neighbouring Palaeoproterozoic crust(see also Yegorova et al. 2004).

The crustal structures and boundaries in the Archaean part of thePodolian Domain dip gently towards the north, conforming to theinferred northerly dips of the Archaean–Proterozoic crustalboundary in the BZ. The grades of Palaeoproterozoic metamorph-ism increase southwards from amphibolite- and granulite-faciesrocks to high-grade granulite-facies rocks. This suggests that thePalaeoproterozoic crust of the VD overlies the Archaean crust ofthe PD in a manner that may be a result of collisional tectonicsat c. 2.1–2.05 Ga, followed by extension and the formation of ametamorphic core complex (Fig. 14).

The greatest lateral and vertical variations of crustal compo-sition and structure in western Sarmatia are associated with the1.80–1.74 Ga multiphase Korosten Pluton. The influence of thisintrusion is not restricted to its area of exposure, but extends fortens of kilometres in the surrounding region. All the crustalunits have been updomed in a wide region (see Figs 8, 9 and14). The gravity and magnetic modelling of the KorostenPluton, employing also the data of the east–west-trending Geotra-verse II seismic profile (Ilchenko & Bukharev 2001), indicates anextremely complex structure in the underlying crust (Bogdanovaet al. 2004b). Whereas layered gabbro–anorthosite intrusionscan be followed only to depths of less than 10 km, granitoidrocks of various kinds, including rapakivi granites and monzo-nites, form flat-lying sheeted bodies at various levels of theupper and middle crust. The interlayering of these igneoussheets with earlier Palaeoproterozoic supracrustal and plutonicrocks is inferred to be responsible for the presence of low-velocitycrustal layers with P-wave velocities of 6.1–6.5 km s21 anddensities varying between 2.6 and 2.8 g cm23 (Fig. 14).

S. BOGDANOVA ET AL.620

At depths below 16 km, mafic rocks form a single semi-cylindrical, lensoid body beneath the eastern part of the KorostenPluton (see Fig. 14; note that the EB’97 profile crossed only thewestern half of the intrusion). This body measures c. 90 kmacross and extends to the high-velocity, high-density layer at thebase of the crust, which underlies the VD, the OMB and evenpart of the PD (see above). This giant mafic body is consideredto represent the feeding magma chamber of the Korosten Pluton(Bogdanova et al. 2004b). The underlying crust, with P-wavevelocities of 7.4–7.8 km s21, Vp/Vs of 1.77–1.79 and densitiesof 3.0–3.15 g cm23, is therefore assumed to be mostly composedof mafic, ultramafic or eclogitic rocks, presumably representing amixture between cumulates of the Korosten magma and restiticmaterial (Bogdanova et al. 2006).

Aspects of the mantle

With regard to Moho topography, the EUROBRIDGE seismicprofiling and gravity modelling indicate depths varying between40 and 55 km as well as a number of irregularities and offsets.Some of these were related to accretionary and/or collisional tec-tonics, or to superimposed late to post-collisional magmatism. Themost obvious case of the former is the mantle irregularity beneaththe Belarus–Podlasie Belt (VPG), which appears to connect withthe Fennoscandia–Sarmatia collisional junction as defined by theMinsk Fault. Another, but so far less evident instance, maybe the relationship between the Mid-Lithuanian Suture Zone(MLSZ) and a Moho offset beneath the West LithuanianDomain (Fig. 13). An offset beneath the Volyn Domain (VD)may continue the Archaean–Proterozoic boundary in the Berdi-chev Zone, but coincides with and may have been masked bythe root of the Korosten Pluton (Fig. 14). In addition, it is temptingto relate the Korosten magmatism to the extensive Moho uplift to45 km beneath the Volyn Domain and the adjacent parts of theOsnitsk–Mikashevichi Belt (OMB) and Podolian Domain (PD).

The upper mantle is rather inhomogeneous with regard toseismic velocities, the P-wave velocities ranging between 8.1and 8.35 km s21. This is probably due to lateral compositionalvariation from peridotite to eclogite. The latter composition is par-ticularly characteristic of sites of collisional thickening in zones ofdeformation and fluidization of the lower crust and upper mantle(e.g. Austrheim et al. 1997). However, as mentioned above, a res-titic origin of eclogites in the lowermost crust and upper mantlebeneath the Korosten Pluton is also possible. Xenoliths in kimber-lites of various ages and near-source alluvial placers have demon-strated that the upper mantle beneath the VD is made up of a20 km thick layer of eclogites, underlain by garnet pyroxenitesand peridotites (Tsymbal & Tsymbal 2003). According toTsymbal & Tsymbal, the age of the mantle is Proterozoicbeneath both the Volyn and the Podolian domains. A more eclogi-tic composition of the mantle beneath Sarmatia may be the reasonwhy it has substantially higher Vp/Vs ratios than the mantle ofFennoscandia (1.83 and 1.72, respectively).

Of particular interest in the tectonic interpretation of the litho-sphere are reflectors in the upper mantle. Apart from some subhor-izontal reflectors referrable to rheological and mineralogicalchanges with depth, the EUROBRIDGE transect also shows onedistinct inclined reflector and more circumstantial evidence ofseveral others. These may represent ‘fossil’ zones of subductionof oceanic as well as continental crust (see Balling 2000).

The less distinct reflectors can to some extent be extrapolated onthe basis of the Moho topography and compositional variation inthe upper mantle, and also from lower crustal lenses of meltingapparently related to post-collisional processes. The presence ofa lens of lower-velocity and lower-density mantle beneath theedge of the Lithuanian–Belarus terrane where it faces Sarmatiathus suggests subduction into the mantle of a slab of Fennoscan-dian lower crust. Similar relationships, albeit less clearly

expressed, are also found in the upper mantle beneath the MLSZand the adjoining parts of the WLG. Another zone of elevatedmantle reflectivity is associated with the Berdichev Zone outliningthe junction of the PD and VD.

The distinct SSW-inclined mantle reflector beneath the OMBalong the EB’97 profile (Figs 8 and 9) has previously been inter-preted for purely geometrical reasons as the trace of a collisionalboundary between Sarmatia and Volgo-Uralia (Thybo et al. 2003).An alternative interpretation (Aisberg & Starchik 2005) suggeststhat this reflector represents a detachment surface in the crustand upper mantle that was related to the formation of thePripyat–Dniepr–Donets Aulacogen (PDDA) in the late Devonian.The latter interpretation accounts well for the part of the mantlewith lower density and lower seismic velocity that overlies thisreflector and is best explained as consisting of mafic and ultramaficigneous rocks. Support for this interpretation is provided by aDevonian (c. 380 Ma) age of lower crustal hornblendite xenolithsfound in lamprophyric tuffites in southeastern Belarus (Markwicket al. 2001). Geochemically, these may represent remelting pro-ducts of a garnetiferous mantle. Thus, it appears possible thatthe distinct flat-lying reflectors found at depths of c. 10 kmbelow the undulating Moho can also be related to the Devonianevent and mark the presence of mantle-derived melts (see Figs 8,9 and 14). Similar conditions have also been observed alongother seismic profiles running across the PDDA (Grad et al.2003a; Maystrenko et al. 2003).

Conclusions

The EUROBRIDGE traverse project has provided a new under-standing of the structure and formation of the crust and uppermantle in the western part of the East European Craton. Althoughthe results mostly concern the key region between the Baltic andUkrainian shields, and the late Palaeoproterozoic collision ofFennoscandia and Sarmatia, they have relevance also for under-standing the upper lithosphere in the entire EEC. The majorconclusions are as follows.

(1) The crust in the region between the Baltic and Ukrainianshields is Palaeoproterozoic and juvenile. It was formed betweenc. 2.0 and 1.8 Ga by accretionary plate-tectonic processes alongthe margins of the Archaean–earliest Palaeoproterozoic nucleiof Fennoscandia and Sarmatia.

(2) Several Palaeoproterozoic terranes, related either to Fen-noscandia or to Sarmatia, are recognized on the basis of theirdifferent ages, lithologies and tectonothermal evolution. Theyinclude various tectonic settings: juvenile island arcs, back-arcsand active continental margins. The Sarmatian terranes wereformed between c. 2.2 and 1.95 Ga, whereas the Fennoscandianones are, in general, younger, ranging between c. 2.0 and1.8 Ga. Palaeomagnetic data indicate that the Fennoscandian andSarmatian terranes belonged to different plates.

(3) The complex, belt-shaped crustal structure and the faultzones that bound the various belts and domains mostly originatedduring the collision between the Sarmatian and Fennoscandianplates at some time between 1.85 and 1.80 Ga. However, thelistric character of many faults and associated late to post-collisional magmatism, retrograde metamorphism and strongmylonitization along the inter-terrane boundaries as well aswithin their interiors all suggest that post-collisional extensionaltectonics was of crucial importance for the following crustaldevelopment between c. 1.80 and 1.74 Ga, and even later at c.1.71–1.67 Ga. Emplacement of the large AMCG plutons at1.80–1.74 Ga in Sarmatia, and between c. 1.6 Ga and 1.50 Ga inFennoscandia, substantially influenced the composition, petrophy-sical properties and geophysical structure of the crust and uppermantle.

(4) The present major characteristics of the seismic profiles andpotential fields in the Baltic–Belarus region were predetermined

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by late Palaeo- and Mesoproterozoic accretionary, collisional andpost-collisional geodynamics. The last, in particular, causedrearrangement of the lithospheric structure and shaped itspresent geophysical images.

(5) Occasionally, the fault and suture zones between the Fen-noscandian belts and domains can be traced throughout theentire crust (e.g. the Mid-Lithuanian Suture Zone). Displacementsand offsets of the Moho boundary and various crustal layers alongthese zones, as well as the ‘imbricate’ character of the Palaeopro-terozoic crust in southern Fennoscandia, allow comparison with‘thick-skinned’ orogens. The offsets and irregularities of theMoho boundary and lateral changes of petrophysical propertiesand compositions in the upper mantle may be interpreted as ‘fos-silized’ Palaeoproterozoic zones of subduction and collision. Thisis particularly the case in the Central Belarus Suture Zone,between the Fennoscandian and Sarmatian terranes, where thecrust is characterized by a pronounced tectonic layering andnumerous reflectors.

(6) The boundary between Fennoscandia and Sarmatia isdefined by the major Minsk Fault, an extensional feature superim-posed on the suture zone. Beneath the Minsk Fault, the crust wasaffected by doming of the collisionally stacked crustal layers,voluminous magmatism at the base and the formation of a meta-morphic core complex.

(7) Subsequent rifting of the crust and the development of theLate Mesoproterozoic Volyn–Orsha Aulacogen was shallow anddispersed, roughly coinciding with the Central Belarus SutureZone. Also, the Devonian rifting and the formation of thePripyat–Dniepr–Donets Aulacogen did not cause substantial thin-ning of the c. 50 km Palaeoproterozoic crust or its significantreworking. However, the underlying Palaeoproterozoic faultswere reactivated and controlled the position of major listricfaults (e.g. those outlining the Pripyat Trough). The SSW-dippingreflector beneath the northwestern margin of Sarmatia, thus, mostprobably represents a detachment surface bounding this aulacogenin the NE. The low-velocity upper mantle above this reflector is prob-ably a Devonian mantle underplate.

The EUROBRIDGE project (1994–2002) has been a highly successful West–East co-operation enterprise. Despite the many economic difficulties in theEast European countries involved, it produced a wealth of results and offered aunique experience to numerous junior researchers. During its lifetime, geologicaland geophysical institutions, research councils and academies in 17 countriescontributed financially. Particular thanks go to the Swedish Institute’sVisby-Programme, the Royal Academy of Sciences in Stockholm and theINTAS organization (project 94-1664). Research and the workshops werealways conducted in a warm and cordial atmosphere, with most participantsfeeling like members of a single family. In the above text, the section reportingthe seismic results was compiled by M. Grad, A. Guterch and T. Janik, andE. Kozlovskaya authored the section on gravity–seismic modelling. L. Taranand G. Skridlaite produced most of the P–T– t data for the metamorphic rocks.G. Motuza, one of the founding fathers of the project, contributed invaluablematerial from Lithuania, and V. Starostenko did the same for the Ukraine.S. Bogdanova was the scientific leader of EUROBRIDGE, and R. Gorbatschevcoordinated the INTAS effort. The authors thank D. Kurlovich from the Belarus-sian State University in Minsk for help with the preparation of the GIS-formatedmaps presented in this paper.

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