European Lithosphere Dynamics
The Geological Society of London Books Editorial Committee Chief Editor B. PANKHURST ( U K )
Society Books Editors J. GREGORY (UK) J. GRIFFITHS ( U K ) J. HOWE ( U K ) P. LEAT ( U K ) N. ROBINS ( U K ) J. TURNER ( U K )
Society Books Advisors M . BROWN ( U S A ) E. BUFFETAUT ( F r a n c e ) R. GIER13 (Germany) J. GLUYAS ( U K ) D. STEAD (Canada) R. STEPHENSON (Netherlands)
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It is recommended that reference to all or part of this book should be made in one of the following ways: GEE, D. G. & STEPHENSON, R. •. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32.
ARTEMIEVA,I.M. THYBO, H. & KABAN,M. K. Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga. In: GEE, D. G. & STEPHENSON, R. A. (eds) European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 11-41.
GEOLOGICAL SOCIETY MEMOIR NO. 32
European Lithosphere Dynamics EDITED
BY
D. G. GEE University of Uppsala, Sweden and R. A. STEPHENSON Vrije Universiteit, Amsterdam, Netherlands
2006 Published by The Geological Society London
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Contents Preface
vii
Introduction GEE, D. G. & STEPHENSON,R. A. The European lithosphere: an introduction ARTEMIEVA, I. M., THYBO, H. 8z KABAN, M. K. Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga
ll
ZIEGLER, P. A. & DEZES, P. Crustal evolution of Western and Central Europe
43
STAMPFLI, G. M. 8z KOZUR, H. W. Europe from the Variscan to the Alpine cycles
57
COCKS, L. R. M. & TORSVIK, T. H. European geography in a global context from the Vendian to the end of the Palaeozoic
83
Europe: Alpine to Present ZIEGLER, P. A., SCHUMACHER, M. E., DEZES, P., VAN WEES, J.-D. & CLOETINGH, S. Post-Variscan evolution of the lithosphere in the area of the European Cenozoic Rift System
97
CLOETINGH, S., ZIEGLER, P. A., BEEKMAN, F., ANDRIESSEN,P. A. M., HARDEBOL, N., VAN WIJK, J. & DI~ZES, P. Thermo-mechanical controls on Alpine deformation of NW Europe
113
KISSLING, E., SCHMID, S. M., LIPPITSCH, R., ANSORGE, J. • F~3GENSCHUH,B. Lithosphere structure and tectonic evolution of the Alpine arc: new evidence from high-resolution teleseismic tomography
129
WILSON, M. & DOWNES, H. Tertiary-Quaternary intra-plate magmatism in Europe and its relationship to mantle dynamics
147
HARANGI, S. DOWNES, H. & SEGHEDI, I. Tertiary-Quaternary subduction processes and related magmatism in the Alpine-Mediterranean region
167
HORVATH, F., BADA, G. SZAFIAN,P., TARI, G., /~D~,M,A. & CLOETINGH, S. Formation and deformation of the Pannonian Basin: constraints from observational data
191
CLOETINGH, S., BADA, G., MATENCO, L., LANKREIJER, A., HORV,A,TH, F. & DINU, C. Modes of basin (de)formation, lithospheric strength and vertical motions in the Pannonian-Carpathian system: inferences from thermo-mechanical modelling
207
VERGES, J. & FERNANDEZ, M. Ranges and basins in the Iberian Peninsula: their contribution to the present topography
223
ROBERTSON, A. H. F. Contrasting modes of ophiolite emplacement in the Eastern Mediterranean region
235
BEN-AVRAHAM, Z., WOODSIDE, J., LODOLO, E., GARDOSH, M., GRASSO, M., CAMERLENGHI,A.r VAI, G. B. Eastern Mediterranean basin systems
263
SAINTOT, A., BRUNET, M.-F., YAKOVLEV, F., St~BRIER, M., STEPHENSON, R., ERSHOV, A., CHALOT-PRAT, F. & MCCANN, T. The Mesozoic-Cenozoic tectonic evolution of the Greater Caucasus
277
Mesozoic-Palaeozoic Europe PHARAOH, T. C., WINCHESTER, J. A., VERNIERS, J., LASSEN, A. & SEGHEDI, A. The Western Accretionary Margin of the East European Craton: an overview
291
GREGERSEN, S., Voss, P., SHOMALI,Z. H., GRAD, M., ROBERTS, R. G. & TOR WORKING GROUP. Physical differences in the deep lithosphere of Northern and Central Europe
313
WINCHESTER, J. A., PHARAOH, T. C., VERNIERS, J., IOANE, D. & SEGHEDI, A. Palaeozoic accretion of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories
323
FRANKE, W. The Variscan orogen in Central Europe: construction and collapse
333
vi
CONTENTS
SIMANCAS,J. F., CARBONELL,R., GONZALEZ LODEIRO, F., PgREZ ESTAON, A., JUHLIN, C., AYARZA, P., KASHUBIN,A., AZOR, A., MART[NEZ POYATOS,D., SAEZ, R., ALMODOVAR,G. R., PASCUAL, E., FLECHA,I. & MART1,D. Transpressional collision tectonics and mantle plume dynamics: the Variscides of southwestern Iberia
345
MCCANN, T., PASCAL, C., TIMMERMAN,M. J., KRZYWIEC, P., L6PEZ-GOMEZ, J., WETZEL, A., KRAWCZYK, C. M., RIEKE, H. & LAMARCHE, J. Post-Variscan (end Carboniferous-Early Permian) basin evolution in Western and Central Europe
355
OKAY, A. I., SATIR, M. & SIEBEL, W. Pre-Alpide Palaeozoic and Mesozoic orogenic events in the Eastern Mediterranean region
389
BROWN, D., PUCHKOV, V., ALVAREZ-MARRON,J., BEA, F. & PEREZ-ESTAI)N,A. Tectonic processes in the Southern and Middle Urals: an overview
407
MATTE, P. The Southern Urals: deep subduction, soft collision and weak erosion
421
KASHUBIN, S., JUHLIN, C., FRIBERG, M., RYBALKA, A., PETROV, G., KASHUBIN, A., BLIZNETSOV, M. & STEER, D. Crustal structure of the Middle Urals based on seismic reflection data
427
BOSCH, D., BRUGUIER,O., EFIMOV,A. A. & KRASNOBAYEV,A. A. U - P b Silurian age for a gabbro of the Platinum-bearing Belt of the Middle Urals (Russia): evidence for beginning of closure of the Uralian Ocean
443
SLIAUPA, S., FOKIN, P., LAZAUSKIENE,J. & STEPHENSON,R. A. The Vendian-Early Palaeozoic sedimentary basins of the East European Craton
449
STEPHENSON, R. A., YEGOROVA,T., BRUNET, M.-F., STOVBA, S., WILSON, M., STAROSTENKO,V., SAINTOT, A. & KUSZNIR, N. Late Palaeozoic intra- and pericratonic basins on the East European Craton and its margins
463
SAINTOT, A., STEPI-IENSON,R. A., STOVBA,S., BRUNET, M.-F., YEGOROVA,T. & STAROSTENKO,V. The evolution of the southern margin of Eastern Europe (Eastern European and Scythian platforms) from the latest Precambrian-Early Palaeozoic to the Early Cretaceous
481
GEE, D. G., BOGOLEPOVA,O. K. & LORENZ, H. The Timanide, Caledonide and Uralide orogens in the Eurasian high Arctic, and relationships to the palaeo-continents Laurentia, Baltica and Siberia
507
Precambrian Europe
KOSTYUCHENKO,S., SAPOZHNIKOV,R., EGORKIN,A., GEE, D. G., BERZIN, R. & SOLODILOV,L. Crustal structure and tectonic model of northeastern Baltica, based on deep seismic and potential field data
521
HJELT, S.-E., KORJA, T., KOZLOVSKAYA,E., LAHTI, I., YLINIEMI,J. & BEAR AND SVEKALAPKO SEISMIC TOMOGRAPHYWORKING GROUPS. Electrical conductivity and seismic velocity structures of the lithosphere beneath the Fennoscandian Shield
541
KORJA, A., LAHTINEN,R. & NIRONEN, M. The Svecofennian orogen: a collage of microcontinents and island arcs
561
DALY, J. S., BALAGANSKY,V. V., TIMMERMAN,M. J. & WHITEHOUSE, M. J. The Lapland-Kola orogen: Palaeoproterozoic collision and accretion of the northern Fennoscandian lithosphere
579
BOGDANOVA, S., GORBATSCHEV,R., GRAD, M., JANIK, T., GUTERCH, A., KOZLOVSKAYA,E., MOTUZA, G., SKRIDLAITE, G., STAROSTENKO, V., TARAN, L. & EUROBRIDGE AND POLONAISE WORKING GROUPS. EUROBRIDGE: new insight into the geodynamic evolution of the East European Craton
599
SLABUNOV,A. I., LOBACH-ZHUCHENKO,S. B., BIBIKOVA, E. V., SORJONEN-WARD, P., BALAGANSKY,V. V., VOLODICHEV, O. I., SttCHIPANSKY,A. A., SVETOV, S. A., CHEKULAEV,g. P., ARESTOVA, N. A. & STEPANOV, V. S. The Archaean nucleus of the Fennoscandian (Baltic) Shield
627
CLAESSON, S., BIBIKOVA, E., BOGDANOVA,S. & SKOBELEV, g. Archaean terranes, Palaeoproterozoic reworking and accretion in the Ukrainian Shield, East European Craton
645
Index
655
Preface This Memoir, 'European Lithosphere Dynamics', has a history that goes back more than 20 years. At the International Geological Congress (IGC) in 1984 in Moscow, leading Earth Scientists from Western Europe and the Soviet Union agreed to start a new pan-European project, 'EUROPROBE', focused on the European Lithosphere and modelled on the European Geotraverse (EGT, Blundell, Freeman and Muller; see Introduction for references). The latter had started a few years before and was dedicated to integrated geological and geophysical studies of a north-south transect across Europe from the Barents Sea coast of northwestern Norway to Italy and the central Mediterranean. At the 1984 IGC, EUROPROBE was conceived as a comparable, east-west profile, from the border zone between Asia and Europe in the middle Urals to the Iberian Peninsula and the Atlantic margin. EUROPROBE took a few years to climb out of the cradle, but as 'glasnost' and 'perestroika' took over the latter part of the 1980s, the opportunity for wider East-West collaboration was recognized by the President of the International Lithosphere Programme (ILP), Karl Fuchs. At that time we were involved in ILP's Global Geoscience Transects programme and the EUROPROBE initiative was seen in this context. The European Science Foundation (ESF) funded early preparatory meetings in Russia, Poland, Denmark and Germany, and advice was provided from many sources, not least EGT. It was clear by 1990 that support for EUROPROBE was to be found in many European countries and a suitable programme, directly benefiting the many nations, was designed. The simple long-range transect model was abandoned and EUROPROBE emerged 'dedicated to carrying out a new generation of major projects that will improve our understanding of the tectonic evolution of the Earth's crust and mantle, and the dynamic processes which controlled this evolution through time' (Gee & Zeyen 1996). Nine target areas (see Fig. 1) were selected for the main research activities, each run by a research team with a high level of autonomy, and all dedicated to applying integrated geological, geochemical and geophysical methods to understand surfacedepth relationships and to interpret the processes leading to the formation of major features of the European lithosphere. Most of the latter were orogens, ranging in age from Archean to the Present; intra-cratonic rifting was also prominent. For ten years, from 1992 to 2001, EUROPROBE received support from the European Science Foundation at a level that financed
Fig. 1. EUROPROBEprojects.
meetings for Science and Management committees and allowed each of the ten projects to run a workshop. These annual activities were essential for the health of the individual projects and the cohesion of the general programme. There were no central research funds other than a small budget devoted to programme travel and exchange of scientists. The individual projects defined their own goals and leadership and obtained funding for research. Those involving EastWest cooperation were often successful in obtaining support from INTAS (The International Association for the Promotion of Co-operation with Scientists from the New Independent States, NIS, of the former Soviet Union). By 2002, the time had come to review a decade of ESF sponsored research (Gee & Artemieva EUROPROBE 1992-2001). At a meeting held in conjunction with the award ceremony for the Crafoord Prize (Prize-winner Dan McKenzie) in Stockholm, a large group of EUROPROBE geoscientists agreed on the ambitions of a final Memoir. The book would not be confined to the achievements of the ESF programme, but target European Lithosphere Dynamics in general, probing the main features of the European subcontinent to give the reader an overview of the whole development through time. About 30 countries and many hundreds of geoscientists were involved in ESF's decade of EUROPROBE research. About 80 workshops were held all over Europe, from Ekaterinburg to Lisbon and Ankara to Lammi; these were great years of multinational collaboration and the publications that resulted were innumerable. Support for EUROPROBE in eastern Europe was widespread and we particularly remember the encouragement and guidance of the Academy Vice-President Alexander Yanshin in Russia, Academician Vitaly Starostenko in Ukraine, and Academician Radim Garetsky in Belarus. Dr Andrey Morozov at the Russian Ministry of Natural Resources and now with Rosnedra has been the 'foundation' for much of EUROPROBE's research in Russia; he cannot be thanked enough. And for all the workshops and research logistics in the former USSR, Elena Gornaya and, later, Nadezhda Timofeeva were vital communicators and advisers, organizers and entertainers. EUROPROBE operated from a secretariat in Uppsala where Chris Juhlin, Herman Zeyen, Monica Beckholmen and Irina Artemieva kept the programme rolling through the 1990s and Olga Bogolepova since then. We are hugely indebted to our EUROPROBE colleagues, not only for the success of the programme, but also that a foundation for international collaboration in lithosphere science has been established that recognizes no end. Many of the major projects that were planned towards the end of the EUROPROBE's ESF Programme have been realized during the last five years, e.g. IBERSEIS, CAUCASUS, POLAR URALS, KOLBAKAR, FIRE, to name but a few. This Memoir has been monumental job for the Geological Society Publishing House in Bath, UK. We thank Angharad Hills and her colleagues warmly for taking it on and seeing it through, despite many hic-ups and a hip-out en route. Our 200 authors have built a comprehensive overview of the European Lithosphere and helped us with some of the reviewing, for most of which we thank the following: U. Achauer, A. Adam, J. Ansorge, I. Artemieva, G. Bertotti, S. Bogdanova, O. Bogolepova, T. Brewer, F. Beunk, C. Biermann, D. Brown, M.F. Brunet, J. Carney, F. Chalot-Prat, R. Cocks, M. Comas, S. Daly, J. Davidson, C. Doglioni, M. Friberg, K. Fuchs, M. Gaetani, J. Golonka, R. Gorbatschev, A. Gubanov, D. Harper, A. Hegedus, R. Huismans, L. Jolivet, A. Jones, J. Knapp, Y. Mart, P. Matte, A. Mauffret, J. Mosar, F. Neubauer, S. Nielsen, A. Okay, T. Pharaoh, A. Robertson, R. Rutland, A. Saintot, J. Tait, T. Torsvik, D. White, J. Winchester, G. W6rner, T. Yegorova and P. Ziegler. D. G. GEE and R. A. STEPHENSON
The European lithosphere: an introduction DAVID G. GEE 1 & RANDELL A. STEPHENSON 2
1Department of Earth Sciences, Uppsala University, Villavagen 16, SE-75236 Uppsala, Sweden (e-mail: david.gee @geo. uu.se) 2Netherlands Centre for Integrated Solid Earth Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands (e-mail: randell.stephenson @falw. vu.nl)
Europe provides on outstanding field laboratory for studying lithospheric processes through time: for tracing the results of plate movements from the present back into the early Precambrian. This book has been designed to focus on tectonic processes in the European lithosphere through these three billion years and how they may have changed during this time. Two things are particularly striking: the importance of plate tectonics far back through the Proterozoic into the Archaean, and the significance of tectonic inheritance, older structures and rheologies guiding, even defining, the younger evolution. Basement structure has a profound influence on subsequent basin evolution and the distribution of geo-resources. The economic importance of understanding these processes cannot be overestimated. Understanding the dynamics responsible for the construction of continental lithosphere requires integrated interpretation of geological, geophysical and geochemical observations. Hypotheses often benefit from testing by numerical and analogue modelling. In practice, one technology--multi-channel, near-vertical reflection profiling--has played a leading role in connecting surface observations to the deep crust and mantle structure. Combined with other geophysical methods, deep reflection profiling has guided the interpretation of the processes that created the architecture of the lithosphere. The European part of the Eurasian continent, reaching from the Ural Mountains in the east to the Iberian Peninsula in the west and from the Mediterranean into the high Arctic, has a lithosphere that
can readily be treated in two parts, east and west (Fig. 1). Most of eastern Europe is dominated by the old, cold East European Craton (EEC), partly covered by little deformed Phanerozoic and MesoNeoproterozoic rift and platform successions. Flanking the EEC to the east are the late Neoproterozoic Timanides and both this orogen and the craton are abruptly truncated by late Palaeozoic Uralian sutures, marking the border to Asia. Northernmost Europe is dominated by the Caledonides and Timanides. The southeastern edge of the EEC, from the northern parts of the Black and Caspian seas to the southern Urals, is less easily defined, an older history of Neoproterozoic accretion and Palaeozoic tectonics being overprinted by Alpine deformation and uplift, the latter being displayed most prominently in the Caucasus, a mountain belt crowned by Europe's highest peak, Mount Elbrus (5 642 m). Western Europe, with minor exceptions, is composed of thinner, warmer, dominantly Phanerozoic lithosphere, accreted to the EEC during Palaeozoic and younger orogenesis. A broad zone of suturing, reaching from the North Sea to the Black Sea and Anatolia, separates the Craton from Phanerozoic accreted terranes, Caledonide in the north, Variscide in central regions, and Alpine in the south. The term 'Trans-European Suture Zone' (TESZ) was coined by EUROPROBE (see Preface) for this wide zone of deformation, with faulting and tectonic reactivation, many strands of which involve large displacements of the craton margin, where the latter tapers westwards beneath a Palaeozoic platform cover.
Fig. 1. Tectonic map of Europe, showing the distributionof the East European Craton and main orogens.
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 1-9. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
2
D.G. GEE & R. A. STEPHENSON
Fig. 2. Magnetic map of Europe (by courtesy of S. Wybraniecand H. Thybo).
Defining the Palaeozoic sutures, as they occur beneath the Mesozoic and younger cover of the TESZ, and tracing them through the crust into the upper mantle is more difficult along this western edge of the EEC than in the Urals, where post-Palaeozoic deformation does not obscure first-order Uralian structures, including sutures. The contrast in structure and composition of the crusts of eastern and western Europe is clearly seen in the magnetic (Fig. 2) and gravity (Fig. 3) maps of the region. In the Early Palaeozoic, the EEC formed the core of a continent, Baltica, largely covered by a wide shelf sea. Faunal evidence, supported by paleomagnetic data, has provided the foundation for Baltica's inferred Palaeozoic independence, apparently surrounded by oceanic crust, at least during the Cambrian and Ordovician. The shape of Baltica is usually considered to have been roughly circular, comparable to Australia in size. However, the shelf-edge and sutures surrounding Baltica are not all well defined and uncertainties remain, particularly in the far north and in the south, the latter affected by the opening and
closing of Palaeo-Tethys and subsequent convergence and collisional events that have left the Palaeozoic history of this margin obscure. The oldest rocks yet found in Europe occur in Ukraine and are mid-Archean in age (c. 3700 Ma). They comprise part of the Sarmatian segment of the EEC. Two other major segments of the EEC have been recognized, Fennoscandia and Volga-Uralia, and these three segments were assembled in the late Palaeoproterozoic, prior to uplift and erosion and a long subsequent period of mild intraplate deformation. The configuration of this proto-EEC in the Proterozoic and its relation to other continent-bearing plates and to larger assemblies of continents (such as Rodinia), remain to be clearly defined. Fragments of the proto-EEC were certainly rifted off the craton during the Palaeozoic and Mesozoic and possibly also during the younger Proterozoic. Grenville-age orogenesis, so prominent in other parts of the Precambrian world, is clearly defined in Europe only along the western (Sveconorwegian) margin of the EEC in southern
INTRODUCTION
Fig. 3. Gravity map of Europe (from Wybranic et
al.
3
1998), Bouguer anomaliesover land areas and Free Air over the seas, shaded relief image NE illuminated.
Norway and south-western Sweden. However, c. 1000 Ma signatures have been detected on several parts of the craton margin, from Baltica-derived allochthons of the Scandinavian Caledonides to the Timanides of the high Arctic and the Urals. Either Grenville-age terranes were accreted to the EEC during c. 1000 Ma orogeny (e.g. in the Scandes) or, subsequently, during Neoproterozoic orogeny. While there is evidence of intracratonic rifting within the EEC, at least locally during the Neoproterozoic, the eastern flank of what was then proto-Baltica, from Novaya Zemlya in the north to westernmost Kazakhstan in the south, was influenced by Timanian orogenesis. From the southern Urals westward to the Scythian Platform of southern Russia and Ukraine, there is also indirect evidence indicating that Timanian-age orogenesis may have preceded the development of an Early Palaeozoic platform. And further west within the Trans-European Suture Zone, there is local evidence of Neoproterozoic accretion of outboard terrains to the EEC prior to deposition of Cambrian strata. Only along the Baltoscandian margin of the craton is there a well-defined depositional
history that apparently denies the influence of proximal orogeny in the Neoproterozoic. Outboard of Baltica, in the Palaeozoic and younger terranes of western Europe, there is abundant evidence of Neoproterozoic, tectono-thermal activity, referred to as the Cadomian Orogeny. From the Avalonian terranes of the British Isles and northern France to the internal 'basement' fragments of the Variscide and Alpine foldbelts, there is evidence of Cadomian convergent-margin tectonics; this tectono-thermal activity apparently occurred along the subducting margin of Gondwana prior to being transferred across to the Baltica margin during closure of intervening oceans in the Early-Mid-Palaeozoic. These Cadomian terranes carry with them evidence of earlier Proterozoic lithosphere, with age signatures that are not characteristic of the western (today's co-ordinates) EEC. Closure of Tethyan ocean systems and collision of Africa with Eurasia resulted in the complexities of southern Europe's Mediterranean world, with development of Alpine fold belts, from the Pyrenees to the Caucasus. Subduction is continuing today, with
4
D.G. GEE & R. A. STEPHENSON
volcanic arcs and back-arc basins, major thrusting and transcurrent and normal faulting. The signatures of these processes in the deeper crust and mantle are well seen in seismic tomography, potential field and thermal anomalies. All this convergence was going on while the North Atlantic was opening, while the passive margins, volcanic and non-volcanic, of Europe were developing, and the Iceland plume migrating to beneath the midAtlantic ridge. 'European Lithosphere Dynamics' is not a comprehensive account of the European lithosphere, but provides an overview of many of the more important aspects of the European crust and mantle. It is arranged to lead the reader, via introductory chapters treating large parts of the subcontinent, into the Alpine world and then backwards in time through the Variscides, Uralides, Caledonides, and Timanides into the early Proterozoic and Archaean of the East European Craton. The Alps--and associated orogenic belts such as the Pyrenees, Carpathians, and Caucasus--represent the results of the most recent phase of mountain-building in Europe, the one related to the collision of the African and Arabian plates with the European (Eurasian) plate in Cretaceous and Cenozoic times. Associated with Alpine collision is the European Cenozoic Rift System, one of the main components of which is the Rhine Graben. Ziegler et al. utilizing in part inferences about the crustal structure of western and central Europe derived from the Moho compilation presented in the Introductory section of this Memoir (Ziegler & D~zes), argue that the manifestation of Alpine tectonics is to some extent an 'accident' of what was left of Variscan crustal/ lithospheric structure. This was, in turn, inherited from Variscan tectonic processes and how they had interacted with older (Cadomian and Caledonian) structures. Thus, what remains of the results of Early Palaeozoic and Neoproterozoic collisional processes in Europe is the consequence of their convolution with Variscan processes and what remains of the results of Variscan processes is, in turn, related to the extent of their interaction with Alpine processes. The record of this multiphase evolution of the lithosphere, north of the Alps, is discussed further by Cloetingh et al. [a] who show how the thermo-mechanical structure of the lithosphere, in part defined by its tectonic memory, may also control Late Neogene and neotectonic anomalies in crustal subsidence and uplift, linking these with surface processes and topography evolution. While these insights are derived primarily from modelling studies, Kissling et al. have come to similar conclusions by looking at the present-day lithospheric structure of the Alps from tomographic data constrained with high resolution seismic models of crustal structure. They show that substantial differences in the structure of the deep crust appear between the western, central, and eastern Alps. Again, it seems likely that this is a manifestation of the inheritance of particular lithospheric architectures from earlier accretionary (Variscan and older) events. Tertiary-Recent anorogenic intraplate magmatism was widespread in Europe and is spatially and temporally associated with Alpine-Pyrenean collisional tectonics, the development of the Cenozoic rift system in the northern foreland of the Alps (e.g. Ziegler et al.), and, locally, with uplift of Variscan basement of the Massif Central, Rhenish Massif and Bohemian Massif. These are much the same regional structures identified by Cloetingh et al. [a] that are related to broad thermo-mechanical lithospheric heterogeneities, at least in part due to inheritance. Wilson & Downes conclude that the partial melting of the mantle leading to volcanism was induced by adiabatic decompression of the asthenosphere, locally in small-scale, plume-like diapirs that welled up from c. 400 km depth. Tertiary-Recent volcanism in Europe may therefore be the surface expression of a 'warm' European upper mantle interacting with a compositionally heterogeneous overlying lithosphere, the latter 'filtering' the former into its diffuse pattern.
The Alpine-Mediterranean area is characterized by a system of arcuate Cenozoic orogenic belts and extensional basins, both of which can be explained by the roll-back of subducted slabs and retreating subduction zones, in the convergence zone between Eurasia and Africa. Harangi et al. summarize the main characteristics of Tertiary-Quaternary 'subduction-related' magmatism in the Mediterranean region and argue that its compositional variability can be explained by having a strongly inhomogeneous mantle--related to 'accidents' in the nearby upper mantle such as the central Atlantic plume--accompanied by crustal contamination. Ben-Avraham et al. describe the extensional basin systems developed in the central and eastern Mediterranean Sea. They show that there is a fundamental change of style at the end of the Miocene, when major adjustments in the Africa-Europe convergent plate boundary occurred, related to the collision of the Arabian plate with Eurasia and the development of the Anatolian and Aegean terranes as independent microplates. The remnants of Neotethys-related (mainly Late Cretaceous and Tertiary) ocean basins found in the deformed zones of the eastern Mediterranean are addressed by Robertson who concludes that not all ophiolites were emplaced as a result of large-scale horizontal tectonic transport, but that a strike-slip/transpressional tectonic setting may have dominated in some cases. The evidence suggests that the mode of ophiolite emplacement was strongly influenced by first-order inheritance such as the relative orientations of ophiolite emplacement vis-~-vis the adjacent continental margin. Horvath et al. discuss continental collision and back-arc basin evolution as one single, complex dynamic process, with the minimization of potential and deformational energy as the driving principle, as exemplified by the Pannonian Basin, which formed from a collapsing, over-thickened Alpine lithosphere in the Neogene. A key requirement was the presence of a 'free boundary' offered by the rollback of the subducting Carpathian slab, thus allowing orogen parallel crustal extrusion towards the east. Modelling results suggest that, as a whole, the Pannonian Basin area has displayed pronounced lithospheric weakness since Cretaceous times (Cloetingh et al. [b]) and, therefore, has been prone to repeated tectonic reactivation. Pronounced lateral variations in lithospheric strength, at least in part related to geological inheritance, have strongly influenced the thrust load kinematics and post-collisional tectonic history along the adjacent Carpathians Mountains and their foreland. Mountain-building at the European extremes of the Alpine belt--the Pyrenees (and associated topographic features of the Iberian Peninsula) and the Greater Caucasus--are discussed by Verg6s & Fern~ndez and Saintot et al. [a], respectively. Once again tectonic inheritance plays a key role. The distribution of modern topography on the Iberian Peninsula (basins and mountain ranges) is argued to be the consequence of crustal and lithospheric thickening during Tertiary compression and upper mantle thinning during the Neogene-Quaternary, superimposed upon variations in crustal (and possibly mantle) densities. These are understood to be a legacy of Late Palaeozoic orogenesis and lithospheric accretion. Similarly, the Greater Caucasus (GC) fold-and-thrust belt, developed in response to Tertiary ArabiaEurasia collision, represents the structural inversion of a deep marine Mesozoic basin. Changes in tectonic style along the GC--in which varying degrees of thick- and thin-skinned deformation are displayed--may be related to heterogeneity developed in lithosphere that was accreted or modified much earlier, in Late Palaeozoic and Mesozoic times. A key process controlling how Alpine tectonics became convolved with the older Variscan framework of Europe was the closure of the Palaeotethys and opening of Neotethys ocean systems and the development of an array of south Eurasian back-arc basins, followed or accompanied by the break-up of Pangaea and the early development of the Central Atlantic. This stage of European tectonic history can be broadly referred to as 'Cimmerian'. While Stampfli & Kozur review the plate kinematic
INTRODUCTION
record during this time, focusing on accretionary events along the south Eurasian margin, McCann et al. consider the instability and re-equilibration of western and central European intraplate lithosphere that occurred after the Variscan Orogeny. An extensive phase of Permo-Carboniferous magmatism was accompanied by transtensional activity that led to the formation of more than seventy rift basins across the region (see Ziegler et al.). These basins can be characterized according to their position relative to the Variscan Orogen and its structural trends. The geological record of Variscan and Cimmerian orogenesis in the Eastern Mediterranean-Balkan region, with a focus on Anatolia, is described by Okay et al. The pre-Alpide evolution of this region is one of episodic growth of Europe by the accretion of oceanic terranes and Gondwana-derived micro-continents, as outlined by Stampfli & Kozur. While the Palaeozoic history of the Balkans and Pontides resembles that of Central Europe, its Mesozoic (Cimmerian) evolution--because of the consolidation of Pangaea to the west--diverges. Saintot et al. [b] investigate the history of the zone between Anatolia and the EEC--from the TESZ and Carpathians in the west to the southern Urals in the east--perhaps the least known part of the European lithosphere. These authors conclude that, although this area was dominantly an oblique convergent plate margin from the Late Palaeozoic through the Mesozoic, there is no compelling evidence for accretionary orogenesis of Variscan (Carboniferous-Permian) age. Rather, the available data are more consistent with an interpretation in which the crust of the Scythian Platform, from the Pre-Dobrogean Depression across the Crimean Peninsula to the North Caucasus area, represents the thinned margin of the Precambrian continent, reworked by mainly extensional Late Palaeozoic-Early Mesozoic tectonic events (in the hanging wall of a transform to obliquely convergent plate margin; cf. Fig. 4). They suggest that the Precambrian crust may have been accreted to the EEC during the Neoproterozoic, roughly contemporaneously with the Timanian Orogeny along its northeastern and eastern margin, (cf. Gee et al.). Thus, Baltica in the Early Palaeozoic was probably not as circular as previously thought, but had prolongations to the north- and southeast comprising Neoproterozoic accreted terranes. Palaeozoic orogenesis dominated the tectonic evolution both of western Europe and the eastern margin of Baltica along the
Fig. 4. Palaeozoic peri-Atlantic orogens in the Permian (from Matte 1991).
5
border-zone to Asia (Fig. 4). The relationship of these orogens to plate movement is treated by Cocks & Torsvik, who use palaeontological, palaeomagnetic, and other lines of evidence to reconstruct the relationship between Baltica and other continents, in particular the dominating supercontinent Gondwana. For about 100 million years, from the end of the Vendian through the Cambrian and Ordovician, Baltica existed as an independent continent surrounded by oceanic domains and largely covered by platform successions. By the Early Ordovician, subduction systems were closing the oceans and collision, with accretion of microcontinents (e.g. Avalonia) started in some areas in the Late Ordovician. Caledonian, Variscan and Uralian orogenesis followed though the mid and late Palaeozoic and much of the basic architecture of today's Europe was assembled, first as part of Laurussia and thereafter Pangaea. Only along the southern margin, as mentioned above, did subsequent Tethyan tectonics and Alpine collision of Africa with Eurasia substantially change the geology of the European plate. The Caledonides of northwestern Europe (Dewey & Strachan 2005; Gee 2005) define the collision zone of Laurentia with Baltica and the suturing of Laurussia. Subduction of the Baltoscandian margin of Baltica started in Late Cambrian-earliest Ordovician times and the Iapetus Ocean, separating these two continents, closed during the Ordovician, with collisional orogeny (Scandian) lasting from the Early Silurian into and through the Early Devonian. The sedimentary basin record of these events on the northwestern EEC is described by Sliaupa et al. Contractional orogeny had ceased by the Mid-Devonian and most of the EEC was undergoing extension, accompanied, in the later Devonian by widespread mafic volcanism (cf. Stephenson et al.). A SE-trending branch off the main Caledonide Orogen has been defined along the southwestern margin of the EEC in Denmark, northern Germany, and Poland (Katzung et al. 1993) with thrusting to the NE, well defined by seismic surveys and drilling. This deformation zone occurs in the footwall to overthrust Avalonian terranes and forms part of the Trans-European Suture Zone (TESZ), the latter defining the broad boundary between Phanerozoic western Europe and the EEC (Pharaoh et al.). Suturing of Gondwana-derived terranes to Baltica had begun by the end of the Ordovician and continued through the Palaeozoic (Fig. 4). Much of the TESZ structure is obscured by younger cover (cf. McCann et al.). It follows that deciphering the Palaeozoic history has depended on comprehensive integration of geophysical techniques for defining the structure of the deeper crust and mantle. Winchester et al. draw attention to the problems involved in defining the accretion of Neoproterozoic terranes along the margin of the EEC. Some of the former, in southern Poland, the Czech Republic, and probably Romania (Moesia), were apparently accreted before the Ordovician (perhaps early in the Cambrian) and then transgressed by Baltica's platform successions. These Neoproterozoic terranes are inferred to have existed as a promontory later in the Palaeozoic when Gondwana-derived Avalonian continental fragments docked along the Baltica margin. The central European Variscides have been treated extensively in Franke et al. (2000) and Winchester et al. (2002). In this volume, Franke looks at Palaeozoic plate kinematics and divides the Variscan tectono-stratigraphic evolution into major Early and Late Palaeozoic episodes, focusing on the latter. Relationships between Baltica and Gondwana during the development of the Variscides, from oceanic separation to the close proximity and the establishment of Pangea, are discussed. Franke also draws attention to evidence of widespread Mid-Devonian extension and mafic magmatism in the central Variscides; he discusses possible genetic affinities with the intraplate extension occurring at about the same time within the adjacent EEC, specifically in the Dniepr-Donets Basin, described by Stephenson et al. The Variscide Orogen also dominates the geology of the Iberian Peninsula. Simancas et al. focus on a southern transect where a
6
D.G. GEE & R. A. STEPHENSON
deep reflection seismic profile was recently acquired. The profile runs from the southernmost terrane in Iberia--the South Portuguese Zone--northwards into central part of the peninsula, crossing a major zone of transcurrent faulting that dominates this part of the Variscides. The paper relates surface geology to deep crust and mantle structure in one of Europe's outstanding examples of transpressional orogenesis. The major differences in the character of the lithosphere across the TESZ, described by Pharaoh et al., so conspicuous in compilations of the gravity and magnetic fields, are equally apparent at sub-Moho levels (Fig. 5) and deep into the mantle (e.g. Zielhuis & Nolet 1994). Major long-range, wide-angle reflection/refraction seismic surveys such as 'Polonaise' of the 1990s and Celebration-2000 (Janik et al. 2005), that ran from the EEC westward to the Variscan terranes of central Europe, were complemented by regional tomography studies, such as TOR (Gregersen et al.). High angle, non-symmetrical features extend deep into the mantle in the vicinity of the Tornquist Zone, displacing the SW-tapeting margin of the Craton (Grad et al. 2002). The Palaeozoic orogen of easternmost Europe, exposed in the Ural Mountains from the Aral to the Kara seas, marks the edge of the EEC from 48 ~ to 60 ~ N. Further north, the Urals truncate the grain of the NW-trending Timanides to 68 ~ N and then swing northwestwards into the Pai-Khoi-Vaigach-Novaya Zemlya fold and thrust belt. The Uralide Orogen, treated in four papers in this volume, is renowned for its preservation of ophiolites (Saveleva & Nesbitt 1996), arc and back-arc volcanic rocks (Brown et al. 2000), and associated mineralization. Footwall blueschists and eclogites are well exposed from the southern to the polar Urals, crystallizing in subducted Baltica-margin protoliths. For much of the Palaeozoic, from the Late Cambrian to the Carboniferous, the Uralian edge of Baltica developed as a passive margin (cf. Saintot et al. [b]; Sliaupa et al.); outboard, subduction-related complexes that formed in the Uralian ocean did not influence the off-shelf, slope-rise facies of the eastern edge of the EEC until the Late Devonian to Early Carboniferous. Not until the Late Carboniferous did the shelf collapse and flysch followed by Permian molasse herald Uralian orogenesis. This apparently occurred somewhat earlier in the south than in the far north, where folding and thrusting did not influence Novaya Zemlya until the middle Triassic. Comprehensive geophysical investigations of the Urals during the 1990s focused on profiles through its southern and central parts; a polar profile is now in progress. These investigations
Fig. 5. Shear wave velocity variations (in %) beneath Europe at a depth of 80 km with low velocities beneath Phanerozoic Europe in the west and high velocitiesbeneath the East European Craton (based on Zielhuis & Nolet, 1994).
provided the foundation for the integrated geological-geophysical studies reported in this volume. Brown et al. focus on the southern part of the mountain belt, where a 500 km long transect across the orogen was investigated in the mid-1990s by a combination of nearvertical seismic profiling, wide-angle reflection and refraction, and potential field methods (Berzin et al. 1996; Knapp et al. 1996). Brown et al. draw attention to the bivergent character of the orogen, with evidence for volcanic-arc collision with the Kazakhstan continent in the east, prior to closure of the ocean. They also discuss the evolution of intra-oceanic arcs, which occurred prior to Baltica-Kazakhstan collision, and the western foreland fold and thrust belt with its evidence of only limited shortening. Matte also concentrates on the southern Urals, emphasizing the contrast between the Uralides and Variscides, from the existence of the deep Moho beneath the former to the remarkable preservation of the middle Palaeozoic volcanic complexes; their dense root is inferred to account for the thicker crust. Accretion of oceanic and microcontinental terranes was not achieved by orthogonal collision and extreme overthrusting, as in the Scandinavian Caledonides; instead, transpression appears to have dominated the remarkably linear Uralide Orogen. The middle Urals transect is presented by Kashubin et al.; this is based on more than a dozen years of integrated geophysical and geological investigations. These authors summarize the evidence for relating the surface geology to the deep structure defined by CDP profiling and present a synthesis of the orogenic evolution. The crustal roots beneath the middle Urals reach 60 km depth and the truncation of the craton margin is abrupt. Interestingly, this part of the Baltica margin is marked by Timanian blueschists (Beckholmen & Glodny 2004) indicating that the late Neoproterozoic suturing coincides closely with that in the Late Palaeozoic. In the middle Urals, Palaeozoic terranes dominate the accretionary complex and microcontinents are apparently absent. The Ural mountain belt is famous for its mineralization, particularly volcanic-hosted massive sulphides (Koroteev et al. 1997; Allen et al. 2002). One paper in this volume, Bosch et al. directly concerns mineralization. Mafic-ultramafic bodies of the middle Urals, occurring within ocean-derived allochthons, are notable for their locally high platinum contents. Bosch et al. provide new isotope data indicating that these ophiolite-related allochthons are middle to Late Silurian in age, apparently substantially older than other similar Pt-bearing massifs of the Late Devonian. This range of ages coincides with those of related adjacent volcanic-arc complexes and testifies to the longevity of intra-oceanic Pt magmatism. The Uralide orogenic assemblage of ocean-derived allochthons (ophiolites and arc-volcanics), footwall high-pressure blueschists and eclogites and extensively developed late-orogenic granites, reaches from the far south of the mountain belt to the Kara Sea. Further north in Pai Khoi, Vaigach, and Novaya Zemlya only a fold and thrust belt is exposed, similar to that of the Uralian western foreland. Evidence for the continuation of the classical Uralide Orogen northwards into the Barents-Kara shelf and eastwards to Taimyr, as proposed by many authors (e.g. Bogdanov et al. 1996), is reconsidered by Gee et al.; the evidence is far from compelling. Likewise, the proposal (Cocks & Torsvik) that the Timanide margin of Baltica terminates immediately north of Novaya Zemlya is also in doubt. These interpretations are important for any attempt to reconstruct the evolution of the Arctic Basin; the northernmost shelf areas of Eurasia, with their vast hydrocarbon resources, will be the focus of many new studies in the coming years. Palaeozoic orogenesis, Caledonian and Hercynian along the margins of the EEC, was interrupted in the Mid-Late Devonian to Early Carboniferous by intracratonic tiffing and extensive basaltic magmatism. In a review of the Late Palaeozoic rift basins of the EEC, Stephenson et al. point out that the DnieprDonets Basin is a true intracratonic rift basin, cutting across the
INTRODUCTION Archaean-Palaeoproterozoic structural grain of its (Sarmatian) basement, whereas the East Barents-Pechora Basin (Gee et al. Kostyuchenko et al.) and the Peri-Caspian Basin (more speculatively) are pericratonic features, developed on reworked and juvenile crystalline basement accreted to the EEC during the Neo-(?Meso)proterozoic. It is speculated that some of this late Precambrian lithosphere may have rifted away from Lanrussia in the Late Devonian. Sliaupa et al. also note that Peri-Uralian basins developed as passive continental margin basins throughout the Early Palaeozoic. Sedimentation on the EEC was confined to the cratonic margins at this time with only limited intracratonic subsidence, in two distinct geodynamic settings, one where basins formed in response to continental break-up processes (break-up of Rodinia) and the other, where basins formed in response to the reassembly of continental lithosphere fragments and associated continental accretionary processes (Neoproterozoic and Caledonian). Fundamental differences in the thermo-compositional make-up of the Precambrian lithosphere of Europe played an important role in the development of the overlying Palaeozoic sedimentary basins described and discussed by Sliaupa et al., Stephenson et al. and Saintot et al. The influence of basement character on the distribution of Barents shelf hydrocarbon resources is noted by Gee et al. Artemieva et al. present a variety of lithospheric scale geophysical data indicating that differences in structure have both a compositional and a thermal origin and are a legacy of Precambrian terrane accretion and subduction as well as Phanerozoic rifting, volcanism, subduction, and continent-continent collision. With regard to the East European Craton, dominated by Archaean and Palaeoproterozoic terranes (Bogdanova et al.
7
2005), amalgamation of its three different segments (Fig. 6) is inferred to have occurred at c. 1800 Ma (Bogdanova 1993). The craton is largely covered by younger sedimentary rocks and is best exposed in the Fennoscandian Shield; it also crops out in the Ukrainian Shield and Voronezh Massif. Deep drilling and the analysis of invaluable drillcores, together with potential field and seismic surveys, has provided the foundation for current knowledge of the crustal rocks of the EEC--their igneous, sedimentary, and metamorphic histories and their evolution through the Archaean and Proterozoic. In this volume, aspects of the EEC are treated in six papers. One of these addresses the lithosphere as a whole; the seven others concern the surface geology and crustal structure of the Fennoscandian (Baltic) Shield, and the largely unexposed, but extensively drilled, regions of southernmost Fennoscandia and their relationships to Sarmatia. The Fennoscandian Shield is readily divisible into northern and southern regions separated, in eastern Finland and western Russia, by the Karelian block--an Archaean complex, overlain by Paleoproterozoic metasediments, but little influenced by tectonothermal activity of this age. Both to the north and to the south, evidence for Palaeoproterozoic orogenesis is widespread, in the north involving substantial Late Archaean terranes. Daly et al. present a tectonic synthesis of the northern part of the Shield, dominated by the Lapland-Kola Orogen. Here, extensive studies of structure, metamorphism, and geochemistry (including isotope age) together with deep reflection and wide-angle seismic profiling (Kostyuchenko et al.) have provided new insight into a Palaeoproterozoic belt of collision and accretion involving both Archaean terranes and younger juvenile crust. Slabunov et al. focus on the Archaean terranes of the northern part of the Shield, summarizing evidence of meso-Archaean protoliths
Fig. 6. Three-segmentconfiguration of the East European Craton (by courtesy of Svedana Bogdanova and Roland Gorbatschev).
8
D.G. GEE & R.A. STEPHENSON
and late Archaean orogeny. Both meta-ophiolites and high pressure eclogite-bearing assemblages are described, indicating the existence both of oceanic crust and deep underthrusting of continental crust during Archean orogenesis. Subsequent Palaeoproterozoic tectonothermal reworking and south-vergent thrusting emplaced these older complexes onto the northern margin of the Archaean Karelian Craton. To the south of the old Karelian core of Fennoscandia, a wide variety of Paleoproterozoic complexes occur that were thrust northwards and accreted to the Archean margin towards the end of the Palaeoproterozoic. Korja et al. integrate surface geology with a substantial geophysical database in Finland and northern Scandinavia. They present a model of the Svecofennian Orogeny involving several late Palaeoproterozoic pulses of accretion, terminating with gravitational collapse at about the same time (1.78 Ga) as orogenesis ceased along the northern side of the Karelian block. Further south in the EEC, the southern parts of the Fennoscandian segment and their relationships to Sarmatia have been investigated by a large multinational group of geologists and geophysicists. Bogdanova et al. describe a wide range of largely juvenile Palaeoproterozoic terranes, beneath the Neo-(? Meso)proterozoic cover of the craton. Relationships to Sarmatia are recognized and correlated with outcrops of igneous and metamorphic complexes in the Ukrainian Shield. Potential field data and wide-angle refraction/reflection seismic profiling allows definition of the extent and geometry of the terranes and interpretation of the crustal and upper mantle structure. These provide a comprehensive foundation for definition of the suture zone between Fennoscandia and Sarmatia and tectonic modelling of accretional, followed by collisional, orogenesis. Sarmatia is the focus for Claesson et al. who present new isotope age data including evidence of the oldest protoliths in the EEC (c. 3.65 Ga; perhaps 3.75 Ga). Most, but not all, of the ancient terranes are late Archaean in age but Palaeoproterozoic reworking was widespread, particularly during collision with Fennoscandia. The deep lithosphere of specific regions of the EEC was the subject of two major studies described in this memoir. One focused on the Fennoscandian Shield and the other on the craton margin where it wedges out southeastwards beneath the Trans-European Suture Zone. In the former, Hjelt et al. summarize the results of experiments involving electromagnetic measurements and seismic tomography. The tomography indicated that the lithospheric mantle under Fennoscandian Shield part of the craton extends down to at least 300 km; no boundary zone to the asthenosphere was detected. With regard to crust-mantle relationships, no expression of the Archaean-Palaeoproterozoic suture zone along the southern margin of the Karelian Craton was inferred. Both the electrical and seismic investigations indicated considerable lateral heterogeneity in the upper mantle. In contrast to this evidence from the internal parts of the EEC, seismic tomography (Gregersen et al.) across the Trans-European Suture Zone margin clearly indicates that the thin lithosphere of westem Europe's Phanerozoic terranes thickens rapidly into the craton. The westward thinning EEC margin is seen to be displaced vertically by at least two major zones of offset, one of which coincides with the previously well-known Tornquist Zone (Blundell et al. 1992). The shallow (c. 120 km depth) asthenosphere of western Europe is seen to deepen gradually northwards beneath the craton to more that 300 krn. Artemieva et al. have inte~ated these results with many others to present a series of lithospheric thickness maps of Europe as a whole, including the EEC, and discuss the origins of the regional scale lithosphere heterogeneities evident in Europe that have played and continue to play such as important role in its tectonic and geodynamic evolution. We thank Olga K. Bogolepova and Nina Lebedeva-Ivanovafor help with the text and diagrams and Johathan Turner for comments on rearranging them.
References ALLEN, R. I., WEIHED, P. & THE GLOBAL VHMS RESEARCH PROJECT TEAM 2002. Global comparisons of volcanic-hosted massive sulphide districts. In: BLUNDELL, D. J., NEUBAUER, F. & YON QUADT, A. (eds) The Timing and Location of Major Ore Deposits in an Evolving Orogen. Geological Society, London, Special Publications, 204, 13-39. BECKHOLMEN,M. & GLODNY,J. 2004. Timanian blueschist-facies metamorphism in the Kvarkush metamorphic basement, Northern Urals, Russia. In: GEE, D. & PEASE,V. (eds.) The Neoproterozoic Timanide Orogen of Eastern Baltica. Geological Society, London, Memoirs, 30, 125-134. BERZIN, R., ONCKEN,O., KNAPP,J. H., PEREZ-ESTAUN,A., HISMATULIN, T., YUNUSOV, N. & LlelLIN, A. 1996. Orogenic evolution of the Urals Mountains: Results from an integrated seismic experiment. Science, 274, 220- 221. BLUNDELL, D., FREEMAN, R. & MUELLER, S. (eds) 1992. A continent revealed: the European Geotraverse. Cambridge University Press, European Science Foundation. BOGDANOV, N. A., KHAIN, W. E. BOGATSKY, V. I., KOSTYUCHENKO, S. L., SENIN,B. V., SHIPmOV,E. V. & SO13OLEV,S. F. 1996. Tectonic map of the Barents Sea region and the northern part of the European Russia. Institute of the Lithosphere, Russian Academy of Sciences, Moscow [in Russian]. BOGDANOVA,S. V. 1993. Segments of the East European Craton. In: GEE, D. G. & BECKHOLMEN, M. (eds) EUROPROBE Symposium in Jablonna 1991. Polish Academy of Sciences and European Science Foundation, A-20(255), 33-38. BOGDANOVA, S. V., GORBATSCHEV,R. & GARETSKY, R. G. 2005. The East European Craton. In: SELLEY, R. C., COCKS, L. R. M. & Pt.IMER, I. R. (eds) Encyclopedia of Geology. Elsevier, 2, 34-49. BROWN, D., JUHLIN,C., & PUCHKOV,V. (eds) 2000. Mountain building in the Uralides: Pangea to the present. AGU Geophysical Monographs, 132. DEWEY, J. F. & STRACHAN, R. A. 2005. Caledonides of Britain and Ireland. In: SELLEY, R. C., CocKs, L. R. M. & PLIMER, I. R. (eds) Encyclopedia of Geology. Elsevier, 2, 56-63. GEE, D. G. 2005. Scandinavian Caledonides (with Greenland). In: SELLEr R. C., COCKS R. L. M. & PEnMEn, J. R. (eds) Encyclopedia of Geology. Elsevier, 2, 64-74. GEE, D. G. & ARTEMIEVA, I. (eds) 2001. EUROPROBE 1992-2001. Uppsala University, Sweden. GEE, D. G. & ZEYEN, H. (eds) 1996. EUROPROBE 1996--Lithosphere Dynamics. Origin and Evolution of Continents. EUROPROBE Secretariat, Uppsala University. GRAD, M., GUTERCH, A., & MAZUR, S. 2002. Seismic refraction evidence of crustal structure in the central part of the TransEuropean Suture Zone in Poland. In: WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) Palaeozoic Amalgamation of Central Europe. Geological Society, London, Special Publications, 201, 295-309. FRANKE, W, HAAK, V, ONCKEN, O. & TANNER, D. (eds) 2000. Orogenic Processes: Quantification and Modelling in the Variscan Belt. Geological Society, London, Special Publications, 179. JANIK, T., GRAD, M., GUTERCH, A., RADLEZ, D., YLINIEMIJ., TIIRA, T., KELLER, G. R., GACZYNSKIE. & CELEBRATION 2000 WORKING GROUP 2005. Lithospheric structure of the Trans-European Suture Zone along the TTZ CEL03 seismic transect (from NW to SE Poland), Tectonophysics, 411, 129-156. KATZUNG,G., GIESE, U, WALTER, R. & YON WINTERFELD,C. 1993. The Rugen Caledonides, northeast Germany. Geological Magazine, 130, 725 -730. KNAPP,J., STEER, D., BROWN,L., BERZIN,R., SULEIMANOV,A., STILLER, M., SHEN, E. L., BROWN, D., BULGAKOV, R., KASHUBIN, S. & RYBALKA,A. 1996. Lithosphere-scale seismic image of the southern Urals from explosion-source reflection profiling. Science, 274, 226228. KOROTEEV, V. A., DE BOORDER,H., NETCHEUKHIN,V. M. & SAZONOV, V. N. 1997. Geodynamic setting of the mineral deposits of the Urals. Tectonophysics, 276, 291-300.
INTRODUCTION
MATTE, P. 1991. Accretionary history and crustal evolution of the Vailscan belt in Western Europe. Tectonophysics, 196, 309-337. SAVELIEVA, G. N. & NESBITT, R. W. 1996. A synthesis of the stratigraphic and tectonic setting of the Uralian ophiolites. Journal of the Geological Society, London, 153, 525-537. ZIELHIUS, A. & NOLET, G., 1994. Deep Seismic Expression of an Ancient Plate Boundary in Europe. Science, 265, 79-81.
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WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) 2002. Palaeozoic Amalgamation of Central Europe. Geological Society, London, Special Publications, 201. WYBRANIC, S, ZHOU, S, THYBO, H., FORSBERG, R., PERCHUC, E., LEE, M., DEMIANOV, G. D. & STRAKHOV, V. N., 1998. New map compiled of Europe's gravity field. EOS, 79(37), 437-442.
Deep Europe today: geophysical synthesis of the upper mantle structure and lithospheric processes over 3.5 Ga IRINA M. A R T E M I E V A t'2, HANS THYBO 2 & M I K H A I L K. K A B A N 3
1US Geological Survey, Menlo Park, CA 94025, USA (e-mail:
[email protected]) 2Geological Institute, University of Copenhagen, Copenhagen, Denmark DK-1350 3GFZ, Potsdam, Germany D-14473
Abstract" We present a summaryof geophysical models of the subcrustal lithosphere of Europe. This includes the results from seismic (reflection and refraction profiles, P- and S-wave tomography, mantle anisotropy), gravity, thermal, electromagnetic, elastic and petrological studies of the lithospheric mantle. We discuss major tectonic processes as reflected in the lithospheric structure of Europe, from Precambrian terrane accretion and subduction to Phanerozoic rifting, volcanism, subduction and continent-continent collision. The differences in the lithospheric structure of Precambrian and Phanerozoic Europe, as illustrated by a comparative analysis of different geophysical data, are shown to have both a compositional and a thermal origin. We propose an integrated model of physical properties of the European subcrustal lithosphere, with emphasis on the depth intervals around 150 and 250 km. At these depths, seismic velocity models, constrained by body- and surface-wavecontinent-scale tomography, are compared with mantle temperatures and mantle gravity anomalies. This comparison provides a frameworkfor discussion of the physical or chemical origin of the major lithospheric anomalies and their relation to large-scale tectonic processes, which have formed the present lithosphere of Europe.
'Evidence obtained under different experimental conditions cannot be comprehended within a single picture, but must be regarded as complementary in the sense that only the totality of the phenomena exhausts the possible information about the objects. ' Niels Bohr 'One cannot embrace the non-embraceable.' Kozma Prutkov The European continent comprises tectonic structures ranging in age from Archaean to Cenozoic. A great variety of past and present tectonic regimes within the European continent provides a unique opportunity to analyse the effects of processes related to plate tectonics (e.g. continent-continent or continent-ocean collisions, leading to formation of continental orogens and subduction zones) and mantle dynamics (manifesting itself in magmatism, continental tiffing and formation of large sedimentary basins) on lithospheric structure. The Precambrian part of the continent is formed by the East European craton (EEC) that crops out in the Baltic and Ukrainian shields and underlies the Archaean-early Proterozoic East European Platform (EEP) (Fig. 1). The EEP is crossed by a cratonscale system of mid-late Proterozoic rifts in its central part (Gorbatschev & Bogdanova 1993) and Palaeozoic rifts in its southern parts, perhaps of plume origin (Lobkovsky et al. 1996). A unique feature of the EEP is the existence of a thick (typically c. 2 - 4 km, although locally 20 km thick) sedimentary cover over most of the platform (e.g. Nalivkin 1976; Khain 1985). Rapid subsidence of the EEP in the Palaeozoic was associated with subduction during the formation of the Uralides orogen (Mitrovica et al. 1996). The fundamental lithospheric boundary in Europe, the Trans-European Suture Zone (TESZ), which was first discovered from geological, palaeontological and magnetic data by W. K. de Teisseyre and A. J. H. Tornquist (Teisseyre 1903; Tornquist 1908), separates the Precambrian lithosphere of the EEC from the Phanerozoic lithosphere of Western Europe. Recent seismic reflection/refraction and tomography studies show a dramatic change in all lithospheric properties across the TESZ (e.g. Zielhuis & Nolet 1994; Arlitt 1999; Sroda et al. 1999; Villasefior et al. 2001). The Phanerozoic part of Europe includes a mosaic of tectonic structures, such as Caledonian, Hercynian (Variscan) and Uralides Palaeozoic orogens, Mesozoic rifts, areas of Cenozoic rifting and tectonomagmatic activity (the Central European Rift System), and Cenozoic collisional
orogens often associated with subducting lithospheric slabs (e.g. the Alps, the Pyrenees, the Carpathians). The goal of this paper is to present a comparative overview of lithospheric structure of the major tectonic provinces of Europe, in an attempt to distinguish the effects of the tectonic evolution of the continent from the Archaean to the present. The results of numerous recent multi-disciplinary international projects in European Earth sciences, the largest of which are the European Geotraverse (EGT) (Blundell et al. 1992) and the EUROPROBE programme (Gee & Zeyen 1996; Gee & Artemieva 2001), form the basis of this paper. The extensive set of geophysical information available for Europe does not permit even simple listing of the key publications. With the goal of summarizing the present knowledge on the European lithosphere on a continent scale, we have deliberately omitted local details. The comprehensive analysis of various geophysical data accumulated by the EUROPROBE research during the past decade is presented in the subsequent papers in this book. With rare exceptions, the lithospheric mantle is inaccessible for direct studies. Images of the upper mantle structure provided by remote geophysical sampling are non-unique, and different techniques measure variations in different properties of the mantle (e.g. density, elastic moduli and conductivity, which are related to variations in composition, structure, mineral alignment, and fluid and thermal regime). Geophysical data obtained by different methods are, to some degree, complementary, such that integrated interpretations of different data types may provide a comprehensive picture of the physical properties of the lithospheric mantle. We combine the highlights of recent achievements in different disciplines of geosciences to provide the reader with comparative and diverse information on the upper mantle structure of the major tectonic structures of the continent. Numerous recent seismological surveys of the deep European lithosphere include a set of continent-scale seismic tomography models. Comparison of these models with thermal and gravity models for Europe permits us to constrain an integrated model of the European lithospheric mantle, which reflects diversity in both its structure and composition.
Precambrian lithosphere of Europe The oldest crust within the European continent (in the Ukrainian Shield, Stepanyuk et al. 1998) is c. 3.6 Ga old and thus is one of
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 11-41. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Baltic Shield
Fig. 1. Simplified tectonic map of Europe. TESZ, Trans-European Suture Zone.
the oldest known on the planet. The oldest crust of the Baltic Shield and the EEP is younger, 3.0-3.1 Ga and 1.8-2.1 Ga, respectively (Fig. 1). The basement of the EEP is buried under a thick cover of Proterozoic and Phanerozoic sediments, which complicates dating of the basement rocks. Petrological studies of mantle xenoliths from Precambrian cratons of the world suggest that the crust and the entire lithospheric mantle of the cratons were formed simultaneously and remained attached ever since (Carlson et al. 1994; Pearson et al. 1999). Therefore, one may expect that the lithospheric mantle of a large part of the continent, from the Urals in the east to the TESZ in the west, also has Archaean-Proterozoic ages. Knowledge of the ages of the subcrustal lithosphere is important for interpretations of seismic and gravity data, as petrological studies of mantle xenoliths indicate that cratonic lithosphere has a unique composition, depleted in basaltic components. The highest depletion is found globally in the Archaean roots and it decreases in Proterozoic and Phanerozoic lithosphere (Griffin et al. 1998). Low iron content in the Archaean lithospheric mantle has important geophysical consequences: it implies higher (by 3-5%) seismic velocities and lower (by c. 1.5%) density than in the Phanerozoic mantle (Jordan 1988; Poudjom. Djomani et al. 1999, 2001; Deschamps et al. 2002). On the other hand, Archaean cratons have the lowest average values of surface heat flow measured on the continents (Nyblade & Pollack 1993). Low temperatures in Archaean lithospheric roots (Pollack & Chapman 1977; Artemieva & Mooney 2001) essentially compensate for the effect of the depleted composition on densities (Jordan 1988) and thus mask gravity anomalies produced by compositional variations in the mantle. However, low temperatures in cratonic lithosphere enhance the effect of depletion on seismic velocities. High mantle velocities, as observed in the EEC, are often interpreted in terms of 'hot' or 'cold' regions, but their origin can be both compositional and thermal. For example, a 1% velocity increase can be caused either by 4% Fe depletion or by 100-150~ temperature decrease in the mantle (Nolet & Zielhuis 1994; Deschamps et al. 2002). We present seismic and gravity models for Precambrian Europe and compare them with thermal models to distinguish structural and compositional variations in the lithospheric mantle.
S e i s m i c data. Most of the data on the lithospheric structure of the EEC come from the Baltic Shield, for which interpretations of seismic reflection/refraction profiles, regional upper mantle seismic tomography, electromagnetic, xenolith, thermal and elastic data became available over recent decades. This extensive dataset provides important information on the lithospheric evolution of the Baltic Shield since the Archaean and reveals the presence of a thick lithospheric keel beneath it. A 180-230 km thick lithosphere has been interpreted from explosion P-wave data along the long-range refraction FENNOLORA profile in the northern part of the Baltic Shield (Guggisberg & Berthelsen 1987). The existence of a high-velocity upper mantle down to 200-250 km beneath most of the EEC, including the Baltic Shield, is supported by regional dispersion analysis of long-period Rayleigh waves and by large-scale P- and S-wave seismic tomography models (Calcagnile 1982, 1991" Bijwaard & Spakman 2000; Shapiro & Ritzwoller 2002; Boschi et al. 2004) (Fig. 2). However, most surface-wave models lose resolution at depths below c. 200250 km and cannot provide reliable constraints on mantle structure below this depth (e.g. Panza et al. 1986). Some regional high-resolution P-wave tomography models have been interpreted as indicators of the existence of high seismic velocities (+2% anomaly compared with the global continental model iasp91, Kennett & Engdahl 1991) down to 250 __+50 km under the Baltic Shield of Finland (Bock et al. 2001; Sandoval et al. 2004). The region with the thickest lithospheric keel is located at the suture between the Archaean and early Proterozoic provinces, and spatially coincides with the anomalously thick crust that has formed during Palaeoproterozoic accretion of Svecofennian terranes to the Archaean Karelian block (Korja et al. 1993). The small size of the region (c. 200km x 300km), where both the crust and the lithosphere have anomalous thicknesses, suggests that both crustal and lithospheric roots could have been formed during the same tectonic event and may represent a unique preserved remnant of an ancient subduction zone. This hypothesis is supported by xenolith data that indicate a compositionally stratified mantle in the region (Peltonen et al. 1999), and by an eastward-dipping high-velocity anomaly in the mantle beneath the Archaean-Proterozoic suture (Sandoval et al. 2004). The geographical distribution of mid-Proterozoic rapakivi granite intrusions at the western and southern sides of the anomalous region of thick lithosphere suggests a deflection of ascending magmas by the pre-existing lithospheric keel. This deflection of mantle heat and magma could have assisted the survival of this thick keel during the mid-Proterozoic tectonothermal activity in the region, which 'embraces' the anomalous region of thick lithosphere and led to the formation of the Baltic-Bothnian Sea basin. A layer with reduced seismic velocities (c. 8.1 km s-1 for the mean model) has been identified at the depth range of 100-160 km within the high-velocity (8.6 km s -~ at 100 km depth) lithospheric mantle of the Baltic Shield (Perchuc & Thybo 1996). Similar seismic velocity structure has been revealed for the Archaean part of the Karelian province in a recent surfacewave based seismic tomography survey (Bruneton et al. 2004), similar to recent results from the Canadian Shield and Greenland (Darbyshire 2005). Tomographic inversion for velocities in the upper mantle in the Baltic Shield, based on the FENNOLORA data, suggests that the 100-160 km depth interval is also characterized by very small S-wave velocities, corresponding to a much more pronounced reduction in velocity for S waves than for P waves (Abramovitz et al. 2002). The nature of the reducedvelocity zone is still debated. Alternative interpretations include (1) regional metasomatism (Bruneton et al. 2004); (2) the presence of pockets of small-percentage melting or fluids (Perchuc & Thybo 1996), probably associated with ancient subduction zones
DEEP EUROPE TODAY (although the layer may be at supersolidus temperatures; Abramovitz et al. 2002); (3) petrological heterogeneities in the lithosphere (e.g. a compositional boundary from a highly depleted upper lithosphere to a less depleted lower lithosphere can produce a seismic pattern similar to the top of a low-velocity zone; Artemieva 2003). However, neither the existing seismic models nor petrographic data on mantle xenoliths (Kukkonen & Peltonen 1999) require the presence of asthenospheric material in the upper 250300 km beneath the Archaean-early Proterozoic part of the Baltic Shield. This conclusion is supported by electromagnetic studies in the region (Korja 1990), in which no highly conductive asthenospheric layer has been identified beneath the Finnish part of the Baltic Shield. Earlier interpretations of a high-conductivity layer below 100-130km depth (e.g. Jones 1982, 1984)should be considered with caution, as they did not account for high-latitude ( > 6 0 ~ distortions of the magnetic field (Osipova et al. 1989). Seismic evidence f o r Precambrian plate tectonics. At present, Precambrian plate tectonic processes are reliably identified only from deep mantle reflectors and associated structures in active seismic reflection surveys. Teleseismic tomography cannot resolve small velocity contrasts (e.g. < 1%) in the lithospheric mantle beneath Archaean and Proterozoic terranes (e.g. Poupinet et al. 1997; Sandoval et al. 2004). With the exception of the Archaean-Proterozoic suture in the Baltic Shield (as discussed in the previous section) and the Southern Baltic Sea (Abramovitz et al. 1997), neither the anomalous crustal structure typical for modern collisional orogens, nor a linear high-velocity seismic anomaly in the mantle (which might indicate the presence of a subducting slab) is documented for Proterozoic collisional structures. The only robust dipping high-velocity 'slab' anomaly in a cratonic root has been distinguished recently in P- and S-seismic tomography studies along the Western Superior Transect (Canada) down to c. 660km depth (Sol et al. 2002). Otherwise, the oldest slab of subducted lithosphere individually recognized in the mantle from teleseismic tomographic data is Jurassic in age (van der Voo et al. 1999). Welldocumented evidence for Precambrian plate tectonic processes was first presented by the BABEL Working Group (1989) for the Baltic Shield. Older relict (2.7-2.8 Ga) subduction has been imaged in seismic reflection studies by the Canadian LITHOPROBE programme in the Superior province (e.g. Calvert et al. 1995; Clowes et al. 1996) and in the Slave craton (Bostok 1998; Cook et al. 1998, 1999; Aulbach et al. 2001). Analogy between the observed reflection geometries and modern subduction zones allows interpretations of seismic images as ancient subduction of former oceanic crust (van der Velden & Cook 1999). Dipping mantle reflectors are of a particular importance, as they are interpreted as relict subduction zones. Two large-scale high-resolution marine seismic reflection experiments in the Baltic Shield (BABEL in the Bothnian Gulf and 'Mobil Search' in the Skagerrak between Norway and Denmark) have found evidence for sets of dipping mantle reflectors, which provide new insights into Precambrian tectonic processes. Distinct, dipping sub-Moho reflections have been identified at 40-110 km depths (BABEL Working Group 1990, 1993; Lie et al. 1990). Dipping at a 15-35 ~ angle, these reflections can be traced laterally over distances of up to 100 km, and in two out of three occurrences they are accompanied by a sharp 5 - 7 km offset of Moho. By analogy between the reflectivity patterns in the Baltic Shield and Cenozoic (e.g. the Alps and the Pyrenees) and Palaeozoic (the Caledonides and the Appalachians) orogens, these mantle reflectors are interpreted as relics of Proterozoic (0.9-1.2 Ga and 1.8-1.9 Ga) tectonic processes related to Svecofennian and Sveconorwegian plate convergence, subduction and accretion of terranes onto the Archaean nucleus of the Baltic Shield (BABEL Working Group 1990, 1993b).
13
This tectonic interpretation is supported by S m - N d isotopic data from the exposed volcanic arc complex in the Baltic Shield (Ohlander et al. 1993). Recent analysis of lithospheric-scale seismic data from 1.90-1.85 Ga subduction zones at the Slave and Baltic cratonic margins (Snyder 2002) reveals strong similarity between them and modern tectonic analogues. Thermal and xenolith data. Surface heat-flow values within the Baltic Shield are close to the global average for Precambrian cratons, 30-50 m W m -2 (Nyblade & Pollack 1993), although extremely low values (20-30 mW m -2) have been reported for the southern part of the Finnish-Karelian province (Bailing 1995; Kukkonen & Joeleht 1996) (Fig. 3). Several thermal models for the upper mantle of the Baltic Shield indicate that variations in the surface heat flow largely result from heterogeneous heat production in the crust (Pinet & Jaupart 1987; Kukkonen 1998). Estimates of Moho temperatures vary from 350 ~ to 600~ (Bailing 1995; Kukkonen & Joeleht 1996; Pasquale et al. 2001; Artemieva 2003); large scatter comes not only from different model constraints but also from a highly heterogeneous crustal structure, varying in thickness from c. 30 km in the Caledonides to c. 60 km at the Archaean-Proterozoic suture in southern Finland. Thermal models suggest that in the Archaean-early Proterozoic part of the Baltic Shield the thickness of the thermal boundary layer with predominantly conductive heat transfer (thermal lithosphere) is in the range from 200 to 280 km (Pasquale et al. 2001; Artemieva 2003). These values are in agreement with regional seismic tomography models, in which no lowvelocity layer has been found down to a 250-300 km depth (Fig. 4). However, a direct quantitative comparison of lithospheric thickness constrained by diverse techniques is inadequate, as they measure different physical properties of the upper mantle (Artemieva & Mooney 2002). For example, the difference between 'seismic' lithosphere (defined as the seismic highvelocity region on the top of the mantle) and 'thermal' lithosphere (defined as the depth at which the geotherm intersects the mantle adiabat or becomes supersolidus) can be up to several tens of kilometres (Jaupart & Mareschal 1999); this difference approximately corresponds to the thickness of the transition zone between purely conductive and purely convective heat transfer. In tomography studies, where seismic lithosphere is considered as the layer above the convecting mantle, its base is defined either as a zone of high velocity gradient or the bottom of a layer with positive velocity anomalies. However, seismic tomography and seismic refraction models would not necessarily indicate the same depth to the base of the lithosphere. In seismic reflection surveys, strong mantle reflectors are often interpreted as the base of the seismic lithosphere, as it is assumed that they originate at the transition from the lithosphere to a zone of partial melt (Lie et al. 1990). Furthermore, the base of the seismic lithosphere should be a diffuse boundary if the decrease of the seismic velocities associated with the lithospheric base is caused by high-temperature relaxation or by partial melting (Anderson 1989). Xenolith geotherms for mantle-derived peridotites from kimberlite pipes of the Finnish part of the Baltic Shield and the Arkhangelsk region confirm low mantle temperatures (Kukkonen & Peltonen 1999; Kukkonen et al. 2003; Malkovets et al. 2003) (see Fig. 6). Peridotites from Finnish xenoliths suggest that lithospheric mantle extends down to at least 240 km depth (the depth from which the deepest xenoliths originated) (Kukkonen & Peltonen 1999) as the peridotites show no variations in texture or composition that could be interpreted as indicators of the transition zone from conductive to convective heat transfer. For example, high-temperature sheared peridotites are absent even in the deepest sampled part of the lithospheric column.
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I.M. ARTEMIEVA ETAL.
Fig. 2. Cross-section of the European lithosphere at depths of 150 km and 250 km. Most of the Precambrian part of the continent has high seismic velocities and low attenuation, at least partly caused by low mantle temperatures. In contrast, Phanerozoic Europe is characterized by low seismic velocities, high attenuation and high temperatures. (a) P-wave velocity perturbations with respect to the ak135 model (based on the tomography model of Bijwaard & Spakman (2000), smoothed by Gaussian filtering). The lateral resolution of the model is very uneven. High resolution (c. 100 km) is achieved for regions with a good coverage of events and stations (Southern and Western Europe). For the EEP the lateral resolution is very low (500-1000 km) and this region is shown white. The vertical resolution of P-wave tomography models is poor, as body waves sample the entire mantle with almost vertical propagation. Most of the anomalies seen in the map propagate to deeper levels (see. (c)). (b) Rayleigh-wave phase velocities (based on the global model of Shapiro & Ritzwoller 2002). The vertical resolution is 50-100 km for the upper 250 km and coverage disappears at deeper levels; the lateral resolution does not exceed 500-1000 km. (c) As (a) for 265 km depth (based on the model of Bijwaard & Spakman 2000). The low lateral resolution for the eastern Baltic Shield and EEP, should be noted. (d) As (b) for 250 km depth (based on the global model of Shapiro & Ritzwoller 2002). The surface wave inversion loses resolution below depths of c. 250 km.
DEEP EUROPE TODAY
15
Fig. 2. Continued. (e) P-wave velocity perturbations with respect to the sp6 reference model (based on the tomography model of Piromallo & Morelli (2003), defined over the equi-spaced nodes with 0.5 ~ spacing). The model has been smoothed by Gaussian filtering. Vertical resolution is low compared with surface-wave tomography. The model resolves similar features in the upper mantle as the model of Bijwaard & Spakman (2000). (f) Mantle temperatures (in ~ at 150 km depth (Artemieva 2003, complemented by new data for Western Europe). Temperatures for the EEC are constrained by surface heat flow for steady-state conductive heat transfer; geotherms for Western Europe are constrained by lithospheric thickness data derived from different seismic models and assuming that 1300 ~ is reached at the lithospheric base. The uncertainty in temperatures is c. 10-15%, but for western Europe can be locally larger. Lateral resolution is c. 50-500 km. (g) Rayleigh-wave tomography for velocity model at 150 km depth (based on the model of Billien et al. 2000). The model is constrained effectively to 12th-degree spherical harmonics with a vertical resolution of c. 5 0 - 8 0 km at 150 km depth. (h) Rayleigh-wave tomography for inverse attenuation at 150 km depth (based on the model of Billien et al. 2000). The model is constrained effectively to 12th-degree spherical harmonics with a vertical resolution of c. 5 0 - 8 0 km at 150 km depth.
16
I.M. ARTEMIEVAET AL.
Fig. 3. Surface heat flow in Europe (after Pollack et al. 1993, updated for new heat-flow data); a low-pass filter has been applied to remove short-wavelength anomaliescaused by shallow effects (e.g. heterogeneitiesin crustal heat production and conductivity). Stars show locations of mantle xenoliths discussed in the text.
East European Platform Seismic data. The lithospheric mantle of the EEP is not studied as extensively as the upper mantle of the Baltic Shield. Continent-scale seismic tomography models (Fig. 2), especially for body waves, have insufficient resolution for the northeastern parts of the EEP as there are few seismic events and the distribution of stations is sparse. Regional electromagnetic models are limited to models of crustal conductivity. With rare exceptions, seismic reflection or refraction profiles do not image the lithosphere deeper than 5 0 - 6 0 k m (Vinnik & Ryaboy 1981; Garetskii et al. 1990; Grad & Tripolsky 1995; Kostyuchenko et al. 1999; EUROBRIDGE Working Group & EUROBRIDGE'95 2001; Grad et al. 2002; Thybo et al. 2003). Weak mantle reflectivity along the profiles, which image the lithosphere of the EEP to a significant depth, suggests either that the entire cratonic root was formed in a fast thermal event in the Precambrian, or that pre-existing reflectivity has been erased by later tectonic processes. However, the lack of significant tectonic activity in most of the EEC since the Precambrian rules out the latter hypothesis. Recent P- and S-wave tomography of the upper mantle of the entire EEP has demonstrated that it is characterized by constant shear velocities (4.65 km s-1) in the depth range 1 0 0 - 2 5 0 k m and radial anisotropy (c. 5%) down to a depth of 200-250 km, where the anisotropy decreases sharply to c. 2% (Matzel & Grand 2004). The depth of 250 km is interpreted as a transition from dislocation deformation to diffusion creep and thus may be considered as a rheological base of the EEP lithosphere. Seismic refraction data indicate that the lithosphere of the northem EEP (along the Peaceful Nuclear Explosion (PNE) profile Quartz) is c. 200 km thick (Mechie et al. 1993; Ryberg et al. 1996); the base of the lithosphere is likely to have a transitional character as no sharp velocity contrast was found at the inferred lithospheric base. Waveform inversion for the upper mantle
structure in the western part of the EEP along the 30~ meridian revealed similar values of lithospheric thickness, c. 200 km (Paulssen et al. 1999). These estimates of the seismic base of the lithosphere are, on the whole, in agreement with thermal estimates of the lithospheric thickness of the EEP, c. 170-200 km with small regional variations within the accuracy of the model (Artemieva 2003; Fig. 4c). Similar to the Baltic Shield, a pronounced reduced-velocity channel at a depth of 105-130 km has been identified within the lithospheric mantle of the northeastern EEP along the PNE profile Quartz (Ryberg et al. 1996). According to travel-time inversion of seismic data along the PNE profiles Quartz and Kraton, this feature extends eastwards as a continuous layer for at least 3000 km into the West Siberian Basin and the Siberian Shield (Nielsen et al. 1999). Similar reduced-velocity layers have been reported earlier for other cratonic regions of the world (Grand & Helmberger 1984; LeFevre & Helmberger 1989; Pavlenkova et al. 1996; Darbyshire 2005) and suggest that it may be a global characteristic of Precambrian lithosphere (Thybo & Perchuc 1997; Thybo 2006). The proposed models for such a layer, with a relatively low seismic velocity within high-velocity cratonic root, include the presence of fluids, partial melts (or temperature close to the solidus), metasomatism, or compositional variations. For example, in North America, a low-velocity zone was found in an S-wave model but was not observed in a P-wave model, which suggests that it is an indicator of a partially molten zone (Rodgers & Bhattacharyya 2001). Thermal data. The EEP is characterized by relatively homogeneous values of the surface heat flow (35-45 m W m -z, Fig. 3), that are within the range of the global average for the Archaean-early Proterozoic cratons of the world (Nyblade & Pollack 1993). Slightly higher values (40-55 m W m -2) have been measured in the southern parts of the platform although, locally, thermal anomalies can reach values as high as 7 0 - 9 0 m W m - 2 (i.e. in the Pripyat Trough). The transition to the Phanerozoic lithosphere of Western Europe is marked by a sharp step-like increase in surface heat flow by c. 20 mW m - 2 (Fig. 3). The thickness of the thermal lithosphere within the EEP has been estimated to be 1 7 0 - 2 0 0 k m (Cermak 1982; Artemieva 2003; Majorowicz et al. 2003) (Fig. 4c). Surprisingly, the Ukrainian Shield, which is the oldest part of the European continent, has similar lithospheric thickness, 1 8 0 - 2 2 0 k m (Kutas 1979). Such values have also been reported for the Archaean lithosphere of South Africa and Australia (Jaupart & Mareschal 1999; Artemieva & Mooney 2001). These cratons are among the oldest on the Earth: the major crust-forming events in the Kaapvaal, Zimbabwe, Indian and Pilbara cratons and the Greenland Shield occurred at c. 3.0-3.5 Ga, whereas in the East European, Siberian and North American cratons the major crust-forming events occurred significantly later, c. 1.8-2.5 Ga (Goodwin 1996). The large difference in lithospheric thickness of Precambrian regions, which were assembled into cratons at different times (Artemieva 2006), poses the question of whether different tectonic and/or mantle processes have operated in the early and late Archaean and led to the formation of cratons with significantly different lithospheric structures (Artemieva & Mooney 2002; Artemieva et al. 2002). As R e - O s isotope studies indicate similar geological ages (i.e. approximately the ages of crustal differentiation; Richardson et al. 1993) for all of the Archaean cratons, it is likely that anomalously thick lithospheric roots could have formed by different intensities of tectonic modification of pre-existing terranes during the cratonization stage and not as a result of different differentiation processes within the deep mantle. Precambrian rifts within the EEP. Mantle processes have played an
important role in the evolution of the continental lithosphere since
DEEP EUROPE TODAY
17
Fig. 4. Five models of lithospheric thickness in Europe. (For (a)-(c) see caption to Fig. 2 for more details.) (a) Lithospheric base defined by a 1% P-wave velocity perturbation (based on the model of Bijwaard & Spakman 2000, interpolated with a low-pass filter) with respect to the ak135 model. (b) Lithospheric base defined by a 2% S-wave velocity perturbation (based on the model of Shapiro & Ritzwoller 2002, interpolated with a low-pass filter) with respect to the global continental model iaspei91 (Kennett & Engdahl 1991). (c) Thermal lithosphere defined by the intersection of the geotherm with a 1300 ~ mantle adiabat (the model of Artemieva 2003). (d) Lithospheric thickness in Europe based on electromagnetic surveys (compilation of Hjelt & Korja 1993, interpolated with a low-pass filter). Dark blue corresponds to regions where depth to the highly conductive layer exceeds 200 km, or where electrical asthenosphere was not detected. (e) Lithospheric thickness calculated from P-residuals (Babu~ka et al. 1988) under the following assumptions: (1) variations in lithospheric thickness are proportional to P-residuals; (2) lateral variations in average lithospheric velocities (as a result of temperature or compositional variations) are ignored; (3) a homogeneous crustal thickness of 33 km is assumed for the entire Western European region; (4) the results are scaled by data from surface-wave dispersion analysis (Panza et al. 1986) on lithospheric thickness in the Western/kips (220 kin) and the Belgo-Dutch platform (50 km).
18
I.M. ARTEMIEVA ET AL.
its formation. Giant mafic dyke swarms (the oldest known, in SW Greenland, is c. 3.25 Ga old), continental rifting (the oldest known, in the Kaapvaal and Slave cratons, is c. 3.0-3.3 Ga old), and break-up of supercontinents (the oldest known is c. 2.5-2.7 Ga old) are believed to be surface manifestations of ancient plume-lithosphere interactions (Nelson 1992). The ages of the known large-scale mantle-lithosphere interaction events within the EEC are much younger than in other cratons (Khain 1985). In the Baltic Shield, the Riphean (1.35-1.05 Ga) tiring affected the Baltic Sea region with the emplacement of rapakivi granites and a subsequent subsidence of the basin (Ga~il & Gorbatschev 1987). Within the EEP, the fundamental trans-cratonic Central Russia Rift System (CRRS) formed at c. 1.3-1.0 Ga either by a large-scale tiring event or by amalgamation of three large terranes into the EEC (Gorbatschev & Bogdanova 1993) (Fig. 1). This process was followed by intensive intraplate volcanism at c. 1.0 G a - 6 5 0 Ma (Nikishin et al. 1996). However, there is otherwise little evidence for Precambrian rifting in the present-day structure of the deep lithosphere of the EEC, although this may be due to the sparse high-resolution geophysical data coverage of the upper mantle in this region (Figs 2 - 4 ) ; much of the knowledge comes from geological data. Nevertheless, joint interpretations of different geophysical datasets indicate significant compositional variations in the lithospheric mantle of the EEP, which may be related to Precambrian (as well as Phanerozoic) tectonomagmatic activity (see discussion below). Gravity data. Density inhomogeneities in the upper mantle, related
to variations in temperature and mineral composition, can provide significant driving forces of both vertical and horizontal motions of lithospheric blocks. As the gravity field contains effects of all density heterogeneities of the Earth, it is necessary to subtract all signals that do not originate from the mantle to extract the mantle component of the gravity field. These signals include the gravity effect of the crust, which is the largest, but can be approximated from independent a p r i o r i data. The resulting residual gravity anomalies reflect density anomalies in the mantle within the accuracy of the crustal model. Although attempts to calculate mantle gravity anomalies were made since the first seismic sections became available, a reliable 3D gravity model of the lithosphere of most of Europe (Artemjev et al. 1993, 1994) could not be constructed until sufficient data on the crustal structure had been accumulated. The new model of mantle residual Bouguer gravity anomalies, based on updated data on the crustal structure of Europe (Fig. 5), shows a sharp change in the sign of anomalies across the TESZ, from positive values over the EEC to negative values over Western Europe. A strong positive anomaly over the Caucasus implies the presence of a subducting slab, which, so far, has not been resolved in tomographic models (Fig. 2). Near-zero values of mantle gravity anomalies over the Baltic Shield are in agreement with the isopycnic hypothesis (Jordan 1988) and suggest that low lithospheric densities caused by Fe depletion of the cratonic keel are well compensated by low mantle temperatures. The positive anomalies of the EEP suggest that compositional density anomalies in the lithospheric mantle of the EEP are not compensated by temperatures as a result of either a more fertile composition or very low mantle temperatures (Fig. 6). However, a strong positive anomaly in the southern part of the EEP, which has been affected by Palaeozoic tiring, rules out a temperature origin of the gravity anomaly. Spatial correlation of the strongest positive residual gravity anomaly with the position of the Central Russia Rift System (Fig. 5) also suggests a compositional rather than a thermal origin of the anomaly. Furthermore, this conclusion is supported by high average crustal velocities in the CRRS (Fig. 7), which may be caused by magmatic underplating; it implies that infiltration of basaltic magmas into the lithosphere played an important role in the tectonic evolution of the CRRS.
Fig. 5. Mantle residual gravity anomalies, which are a part of a 3D global model (Kaban et al. 1999, 2003; Kaban & Schwintzer 2001), supplemented by higher-resolution regional data (Kaban 2001). The anomalies reflect density variations produced by compositional or temperature variations, presumably in the upper 40-60 km of the subcrustal lithosphere. The model is calculated by subtracting: (1) the anomalous gravity field of the sedimentary cover and water; (2) the anomalies related to the Moho depth variations; (3) density variations within the crystalline crust from the observed gravity field (Bouguer anomalies on land and free-airanomalies offshore). The results depend critically on seismic data on the crustal structure, because during calculations seismic velocities are converted to densities. The predictions of the present model are higher by c. 50 mGal than residual gravity anomalies for the European continent based on older data on the crustal structure (Yegorova & Starostenko 2002), although the general pattern of the anomalies remains similar. Density excess in the mantle is typical for Precambrian terranes and regions of Phanerozoic subduction. Density deficit in the Phanerozoic mantle may be caused by high temperatures and partial melt.
Contrast in lithospheric properties across the Trans-European Suture Zone (TESZ) The TESZ is a fundamental tectonic boundary within the European continent. It is formed by a broad complex zone of Palaeozoic terranes accreted to the southwestern margin of the East European Craton and marks the transition from the Precambrian cratonic lithosphere to the Neoproterozoic-Palaeozoic lithosphere of Western and Central Europe. The deep structure of the TESZ is characterized by a sharp change in lithospheric properties, well established by different geophysical methods (Thybo et al. 1999, 2002). The transition from the cratonic to the Phanerozoic lithosphere is characterized by the following features. (1) Crustal thickness changes sharply from 3 5 - 4 5 km in the EEP, to 4 0 - 5 5 km in the Teisseyre-Tornquist Zone, and to 2 8 - 3 2 km with a surprisingly flat Moho beneath the mosaics of Variscan and Caledonian terranes of Westem and Central Europe (Guterch et al. 1986; Abramovitz et al. 1998; Grad et al. 2002) (Fig. 7). Furthermore, the magnetization of the crust of Central Europe is extremely weak compared with the upper and middle crust of the EEC (Banka et al. 2002). Thin crust with a flat Moho and a lack of seismic signature in the lithospheric mantle of the European Caledonides and Variscides suggests that a large portion of the lower crust and the lithospheric
DEEP EUROPE TODAY
Fig. 6. Typical geotherms in different tectonic structures of Europe. For stable parts of the EEC the geotherms are constrainedby surface heat-flow data assuming steady-state conductive regime (Artemieva2003). Models of heat production distributionin the crust were constrained taking into account: (1) wavelength of surface heat-flow variations; (2) regional seismic models for the crustal velocity structure; (3) regional and global petrological models on the bounds on bulk crustal heat production (see details in Artemieva & Mooney 2001). For tectonicallyactive regions of Western Europe, mantle temperatures are based on a nonsteady-state conductive model constrained by data on Cenozoic magmatism(Artemieva 1993) and on the conversion of regional seismic tomography models into temperatures (Sobolev et al. 1996). For comparison, P-T data on mantle xenoliths are shown (Coisy & Nicolas 1978; Seck & Wedepohl 1983; Nicolas et al. 1987; Werling & Altherr 1997; Kukkonen & Peltonen 1999; Malkovets et al. 2003). Ar-ePt, Archaean-Early Proterozoic.
mantle could have been delaminated as a result of the Palaeozoic orogenies (Ziegler et al. 2004). (2) A pronounced and sharp decrease in seismic velocities (by 2 - 3 % ) down to the depth of 100-200 km is observed at the transition from fast cratonic lithosphere to Palaeozoic upper mantle (Zielhuis & Nolet 1994; Poupinet et al. 1997; Masson et al. 1999; Villasefior et al. 2001; Cotte et al. 2002) (Fig. 2). This velocity contrast is caused by differences in lithospheric composition and mantle temperatures. Part of the velocity anomaly may possibly be attributed to palaeosubduction along the cratonic margin, which increased the fluid content in the upper mantle (Nolet & Zielhuis 1994). (3) The transition zone between the lithospheric terranes of Precambrian and Palaeozoic ages dips at a steep angle to the vertical (c. 13-20 ~ in the Irish Caledonides and the Uralides, based on teleseismic studies (Masson et al. 1999; Poupinet et al. 1997). In comparison, the dip of the transition boundary across the Caledonian Deformation Front in the southern part of the Baltic Shield is shallow (c. 15-20 ~ to the horizontal with a SW dip based on a seismic normal-incidence reflection profile) (MONA LISA Working Group 1997). A subhorizontal boundary between the cratonic and Phanerozoic lithospheres implies that high-velocity lower crust, or a part of the subcrustal lithosphere of
19
Fennoscandia, may extend far to the south (i.e. to the E l b e Oder line), underlying Phanerozoic structures of Northern Europe (Thybo 1990; Cotte et al. 2002). This conclusion is supported by the results of a joint interpretation of seismic, gravity and magnetic data (Thybo 2001; Bayer et aL 2002) and by a likely compositional origin of the velocity anomalies observed in the TOR tomography experiment (see discussion below). A similar pattern of a non-vertical transition from Archaean to Proterozoic lithosphere has been documented by LITHOPROBE data at the margins of the Canadian Shield (Bostok 1999; Ludden & Hynes 2000). (4) A strong subhorizontal upper mantle reflectivity has been documented beneath the Variscides and Caledonides at the depth range of 5 0 - 1 0 0 km (Masson et al. 1999; Abramovitz & Thybo 2000; Grad et al. 2002), as compared with a weak mantle reflectivity in the cratonic lithosphere of the EEC, where only one significant mantle reflector was found at c. 10 km below Moho (BABEL Working Group 1993; Grad et al. 2002). (5) Surface heat flow changes abruptly by 2 0 - 3 0 m W m -2 from cratonic to younger Europe (Fig. 3), and is accompanied by a significant rise in lithospheric temperatures (Cermak 1993; Artemieva 2003, 2006). (6) Lithospheric thickness sharply changes from 150-200 km in the EEC to 8 0 - 1 2 0 km in Phanerozoic Europe (Figs 2, 4, 7 and 8, and Table 2) (e.g. Panza et al. 1986; Babugka et al. 1988; Zielhuis & Nolet 1994; D u e t al. 1998; Artemieva & Mooney 2001). (7) An abrupt change in the upper mantle density structure is reflected in a transition from near-zero or weakly positive isostatic gravity anomalies in the cratonic part to strongly negative anomalies in Western Europe (Fig. 5). Strong negative residual mantle anomalies suggest the presence of low-density masses within the upper mantle and provide indirect evidence for high mantle temperatures. Near-zero isostatic gravity anomalies in the cratonic part of the continent imply that the expected density increase caused by depleted composition of the cratonic lithosphere is entirely compensated by the density increase caused by low mantle temperatures, in agreement with the isopycnic hypothesis (Jordan 1988).
Palaeozoic structures of Europe Palaeozoic orogens of Europe include the Uralides at the eastern margin of the EEP and the Caledonian and Variscan (Hercynian) structures in the western part of the continent (Fig. 1). The crustal structure of European Palaeozoic orogens has been studied in detail by numerous seismic profiles (including normal incidence and wide-angle reflection seismic profiles) in the North Sea (BIRPS, MONA LISA), Germany (DEKORP BASIN 96), France (ECORS), Poland (POLONAISE), Ireland (VARNET- 96), Spain (IBERSEIS, ILIHA, NARS), and in the Urals (ESRU, URSEIS). However, data on the properties of the mantle lithosphere of European Palaeozoic orogens still remain limited (Blundell et al. 1992) and, in the case of the Caledonides, are restricted mainly to the transitional regions from the cratonic to post-cratonic lithosphere (i.e. across the Caledonian Deformation Front) (Masson et al. 1999; Roberts 2003). The Caledonides (named after Caledonia, the Latin name for Britain) and Variscides were formed during orogenic events involving a triple plate collision (Baltica, Laurentia and Avalonia) associated with the closure of the Iapetus Ocean and Tornquist Sea, and subsequent amalgamation of a series of terranes (Dewey 1969; McKerrow & Cocks 1976). Radiometric data on abundant granitoids and metamorphic rocks provide the ages of these Palaeozoic tectonic events, which included deformation, magmatism and metamorphism, as 500-400 Ma in the Caledonides and 4 3 0 - 3 0 0 Ma (possibly as late as 280 Ma) in the Variscan belt (e.g. Stille 1951; Emmermann 1977;
20
I.M. ARTEMIEVAET AL.
Fig. 7. Ranges of (a) average Vp seismic velocities in the crust, (b) crustal thickness, and (c) lithospheric thickness in different tectonic structures of Europe (based on Table 1). CRRS, Central Russia Rift System; CERS, Central European Rift System; PDDR, Pripyat-Dnieper-Donets rift; EEP, East European Platform. Ar-ePt, Archaean-Early proterozoic; Pt, Proterozoic; Pz, Phanerozoic; Mz-Cz, Mesozoic- Cenozoic.
Matte 1986). Opening of the North Atlantic Ocean disrupted the Caledonian orogenic belt into the European (Svalbard, Norwegian, Irish-British and Danish-Polish Caledonides) and the North American (the Appalachians and East Greenland) parts (Dewey 1969). The Uralides orogen, a well-preserved arc-continent collision zone composed of a series of late Proterozoic-Palaeozoic fold belts formed at c. 4 0 0 - 2 5 0 Ma, following the closure of the Uralian palaeo-ocean at c. 4 7 0 - 4 0 0 Ma and the accretion of the Kazakh terrane at the eastern passive margin of the EEC at c. 4 0 0 - 3 2 0 Ma (Edwards & Wasserburg 1985; Savelieva 1987; Sengfr et al. 1993; Bea et al. 1997; Puchkov 1997; Brown et al. 1998). This orogen is partly exposed in the Urals mountains, Severnaya Zemlya and the Taymyr Peninsula, whereas its
eastern part is buried under the West Siberian Basin. Further collisions of the EEC with the Siberian craton resulted in the formation of the Timan Ridge in Triassic-early Jurassic time. Compared with other Palaeozoic orogens, which have been essentially reworked during the late Palaeozoic and Meso-Cenozoic tectonomagmatic processes, the Uralides have remained intact since the Palaeozoic.
E u r o p e a n CaIedonides
A thin crust (Fig. 7), in places with a seismically laminated lower crust and a sharp subhorizontal Moho, that crosses pre-existing terrane boundaries, is typical of the Caledonides, Variscides and
DEEP EUROPE T O D A Y
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6 wt%) provide important constraints on the nature of the mantle source and the conditions of partial melting. These are predominantly sodic (melilitites, nephelinites, basanites and alkali olivine basalts); however, locally, potassic magma types (olivine leucitites, leucite nephelinites) also occur. In several localities (e.g. Sicily; Vogelsberg and the Rhine Graben, Germany; Calatrava, central Spain) olivine and quartz tholeiites form a significant component of the magmatism. The sodic magmas were derived by variable degrees of partial melting (c. 0.5-5%) within a transitional zone between garnet-peridotite and spinel-peridotite mantle facies, close to the base of the lithosphere; the potassic magma types are interpreted as partial melts of enriched domains within the lithospheric mantle. Mantle partial melting was induced by adiabatic decompression of the asthenosphere, locally in small-scale, plume-like, diapirs, which appear to upwell from c. 400 km depth.
Tertiary and Quaternary volcanic activity within Europe occurs in two principal geotectonic settings, referred to as orogenic and anorogenic by Wilson & Bianchini (1999). Occurrences of calc-alkaline volcanism (orogenic) in the Alpine chain, the Carpathians and the Mediterranean region (Harangi et al. 2006) can be explained geodynamically in terms of contemporaneous subduction, and will not be considered further in this review. Here emphasis is placed on the extensive anorogenic, dominantly alkaline, volcanic province to the north of the Alpine collision zone, including the Massif Central of France, the Rhenish Massif of central Germany, the Rhine Graben, and the Eger Graben within the northern part of the Bohemian Massif in the Czech Republic (Figs 1 and 2). Further to the east mafic alkaline volcanism (anorogenic) post-dates a major phase of subduction-related volcanism in the Pannonian Basin, the Dinarides (Serbia, Slovenia, Croatia, northern Bosnia), Bulgaria and western Turkey. Further south, within the Mediterranean region, anorogenic volcanism occurs in Sicily, Sardinia, Monte Vulture and the Veneto area of Italy, in the Alboran Sea and along the northern coast of Africa, locally post-dating earlier phases of subduction-related magmatism. Anorogenic magmatism also occurs in the Iberian peninsula, mainly in the Calatrava province of south central Spain, and within the southeastern Pyrenees near Olot. Although magmatism initiated locally in the latest CretaceousPalaeocene, the major phase of activity in Western and Central Europe occurred in the Neogene ( 2 0 - 5 Ma), with a subsidiary peak in the Pliocene ( 4 - 2 Ma) (Figs 3 and 4); eruptions continued locally to a few thousand years ~p. Magmatic activity within the European foreland of the Alpine orogen (Fig. 2) is typically mafic and occurs as small-volume monogenetic centres (e.g. Eifel, Urach and Hegau provinces of central Germany), scattered necks and plugs (e.g. North Hessian Depression, Germany) and fissure-controlled plateau basalts (e.g. C~zallier, Aubrac and Coirons in the French Massif Central). Rarer central volcanic complexes (e.g. Cantal and Mont Dote in the Massif Central; Vogelsberg in central Germany) include significant volumes of more differentiated magmas that can be related to processes of magmatic differentiation in sub-volcanic m a g m a chambers (e.g. Wilson et al. 1995a).
The volcanic fields are generally concentrated in lithospheric basement terranes that have experienced tectonothermal events within the last 3 0 0 - 4 0 0 Ma (e.g. the Variscan belt of Europe); these typically have higher heat flow and thinner lithosphere than the surrounding cratons (e.g. Baltic Shield) (Prodehl et al. 1992). A number are located on uplifted basement massifs (e.g. Massif Central; Rhenish Massif; Bohemian Massif) that appear to be dynamically supported by upwelling asthenospheric mantle diapirs (e.g. Granet et al. 1995; Ritter et al. 2001; Wilson & Patterson 2001). Magmatic activity was broadly synchronous with the evolution of an extensive intra-continental rift system in Western and Central Europe, subsequently referred to as the ECRIS (East and Central European Rift System), the origins of which are intimately linked to the collision of Africa with Eurasia (Wilson & Downes 1992; Ziegler 1992; D~zes et al. 2004). The most common primitive mafic m a g m a types are sodic basanites and alkali basalts; highly silica-undersaturated, small melt fraction, nephelinites and melilitites occur much less frequently, although they are the dominant m a g m a type in some areas (e.g. Urach province of Germany). Potassic alkaline mafic magmatism (e.g. leucitites, leucite nephelinites) occurs at scattered localities throughout the province (e.g. Calatrava, Spain; Cantal and the Sillon Houiller in the Massif Central; the West Eifel, Germany; Doupovsk6 Hory and Cesk6 Stredohori in the Bohemian Massif); however, only in the Quaternary East West Eifel do potassic magmas predominate over sodic m a g m a types. More exotic m a g m a types, such as carbonatite, occur very rarely (e.g. Kaiserstuhl in the Rhine Graben, Germany; Monte Vulture, Italy). Felsic magmatic rocks occur in most of the volcanic fields, sometimes in a bimodal association with the basalts; complete differentiation series are, however, rare, and occur only in the central volcanic complexes (e.g. Cantal).
The geodynamic setting of the magmatism The Late Cretaceous-Cenozoic convergence of A f r i c a - A r a b i a with Eurasia resulted in the progressive closure of oceanic
From: GEE, D. G. & STEPHZNSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 147-166. 0435-4052/06/$15.00
9 The Geological Society of London 2006.
147
148
M. WILSON & H. DOWNES
Fig. 1. False coloured topographic map of Europe and the Mediterranean region indicating the locations of the main Tertiary-Quaternary volcanic fields discussed in this review. Also shown is the location of a zone of high-velocitymantle within the mantle Transition Zone (500-600 kin) beneath Europe (from Piromallo et al. 2001), which may be a region of subducted slabs at the base of the upper mantle. The ages of such slabs cannot be constrained; however, they most probably represent remnants of subducted Tethyan oceanic lithosphere. The variable size of the red asterisks, marking the location of individual volcanic fields, indicates, schematically,the relative volume of magmatism. TTZ, Tomquist-Teisseyre Zone.
basins in the Mediterranean region and the collision of the Alpine orogen with the southern passive margin of Europe. Compressional deformation of the lithosphere within Western and Central Europe occurred as a response to the collisional coupling of the Alpine and Pyrenean orogens with their forelands (Ziegler et al. 1995; Dbzes et al. 2004). Throughout the Tertiary there was a gradual shift of compressional tectonic activity away from the foreland of the Carpathians and Eastern Alps to the foreland of the Central and Western Alps, partly as a consequence of dextral translation between the converging blocks during the late Eocene to Pliocene. Stresses related to the collision of Iberia and Europe interfered with stresses transmitted from the Alpine collision front (D~zes et al. 2004); these stresses played an important role in the Eocene reactivation of Permo-Carboniferous fracture systems and the localization of the Cenozoic rifts (e.g. Rhine Graben). Convergence rates between Africa and Europe decreased rapidly during the late Cretaceous and Palaeocene (67-55 Ma) as the African and European plates became mechanically coupled (Rosenbaum et al. 2002). During the late Palaeocene (61-55 Ma) a pulse of intense intra-plate compression affected Western and Central Europe, the East European Craton and North Africa (D~zes et al. 2004; Fig. 3). Compressional stresses
exerted by the evolving Alpine and Pyrenean orogenic belts caused lithospheric buckling and basin inversion up to 1700 km north of the orogenic fronts. This deformation was accompanied by local intrusion of small-degree partial melts (e.g. melilitites and nephelinites) in the Massif Central, Vosges, Black Forest, Rhenish Massif and Bohemian Massif. During the early Eocene (c. 52 Ma) the convergence rate between Africa and Europe gradually increased, followed by a decrease in the early Miocene (c. 19 Ma) (Rosenbaum et al. 2002). Scattered volcanic activity occurred during the early and mid-Eocene in the Massif Central (Michon & Merle 2001), the Rhenish Massif (Lippolt 1982) and the Bohemian Massif (Ulrych & Pivec 1997). During the late Eocene extension initiated along the Massif Central, Bresse and Rhine grabens by transtensional reactivation of older Permo-Carboniferous fracture systems in a northerly directed compressional stress field (Dbzes et al. 2004). At the Eocene-Oligocene boundary, convergence of the West Alpine orogenic wedge with the European foreland changed to a NW direction (Ceriani et al. 2001), coincident with the detachment of a southerly subducted lithospheric slab beneath the Central and Eastern Alps and associated isostatic rebound of the European foreland lithosphere (von Blanckenburg & Davies 1995).
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Fig. 2. Relationship of the main Tertiary-Quaternaryvolcanic fields in Western and Central Europe to zones of uplifted basement, major rift systems and the Variscan basement terranes of Central Europe. AfterWilson & Dowries (1991) and Wilson & Patterson (2001). PBF, Pas de Bray Fault; SHF, Sillon Houiller Fault; V/BF Dome, Vosges-Black Forest dome; RG, Rhine Graben; RH, Rheno-Hercynian terrane; S, Saxo-Thuringian terrane; M, Moldanubian terrane; NHD, North Hessian Depression.
North-directed compressional stresses from the Pyrenees, combined with the forces exerted by collisional tectonics in the Central Alps, induced the main Oligocene extensional stage of the ECRIS. The Pyrenean component of compressive stress relaxed during the late Oligocene (D~zes et al. 2004). During the Oligocene the Rhine Graben propagated northwards, bifurcating into the Ruhr and Leine grabens (Fig. 2); rift propagation was associated with an intensification of volcanic activity in the Rhenish Massif. In the northern parts of the Massif Central rifting was accompanied from the late Oligocene by scattered volcanic activity (Michon & Merle 2001). In contrast, subsidence of the Eger Graben within the northern part of the Bohemian Massif commenced only towards the end of its Oligocene magmatic phase. During the Miocene extension continued in the Rhine, Ruhr and Leine grabens (Schumacher 2002); their triple junction (Fig. 2) was gradually uplifted and became the focus of increased volcanic activity (Sissingh 2003). Uplift of the Vosges-Black Forest dome commenced between 19 and 20 Ma; this has been attributed to lithospheric flexuring in the Alpine foreland (Schureacher 2002; D~zes et al. 2004). Minor volcanic activity within the Upper Rhine Graben, including its rifted flanks, was associated with this phase of uplift from 18 to 7 Ma (Jung 1999). Uplift and northward tilting of the Massif Central also commenced during the early Miocene, followed by a rapid increase in volcanic activity during the mid- and late Miocene (Fig. 3; Michon & Merle 2001). Minor compressional deformation of the European lithosphere also occurred during the late Miocene-early Pliocene and in Pliocene-Quaternary times (Fig. 3; D~zes et al. 2004). Extension continues to the present day along the Rhine and Ruhr grabens, whereas subsidence of the Massif Central grabens ceased during the Miocene. Within the Rhenish Massif, volcanic activity
shifted to the Eifel region during the Pliocene and Quaternary, coinciding with an acceleration of uplift (Garcia-Castellanos et al. 2000). Uplift of the northern Bohemian Massif, which initiated during the early Miocene, continued throughout PlioQuaternary times, accompanied by renewed volcanic activity (Ulrych & Pivec 1997; Michon & Merle 2001). This uplift has been attributed to lithospheric flexuring (Ziegler & D~zes 2006). In the northern Massif Central volcanic activity resumed at the beginning of the Pliocene (peaking between 4 and 1 Ma), whereas in the south a second peak of activity occurred between 3.5 and 0.5Ma (Michon & Merle 2001); volcanism was accompanied by renewed uplift. On a regional scale there appears to be a broad correlation between the timing of magmatic activity within the northern foreland of the Alps and changes in the regional stress field (Fig. 3). A detailed compilation of the available geochronological data for the Massif Central suggests that the main volcanic phases may be associated with periods of compressional stress relaxation in the foreland of the Alpine orogenic belt (Wilson & Patterson 2001). Magmas must rise through the crust and upper part of the lithospheric mantle through fracture systems; consequently, it is possible that the distribution of Cenozoic magmatism within Europe could be related to reactivation of pre-existing lithospheric discontinuities (e.g. Permo-Carboniferous sutures and fault systems) in response to changes in the regional stress field. The maximum horizontal stress direction within Western and Central Europe rotated from NNE-SSW to NNW-SSE during the Late Eocene-Early Oligocene to N W - S E in the Late Oligocene (Schreiber & Rotsch 1998). The orientations of linear chains of volcanic necks and scoria cones commonly reflect the orientation of the contemporary stress field. In most areas Cenozoic rifting initiated earlier than the main phase of magmatic activity and is frequently offset spatially from both
150
M. WILSON & H. DOWNES at the boundary between the Saxo-Thuringian and Moldanubian Variscan basement terranes (Fig. 2).
Age and characteristics of the volcanic fields Rhenish Massif
Fig. 3. Chronologyof Tertiary-Quaternary volcanism within Europe and the Mediterranean region in relation to major periods of lithospheric extension, basement uplift and phases of Alpine compression. (For data sources see text.) It should be noted that the time scale from 65 to 25 Ma is compressed. CS, Cesk6 Stredohori; DH, Doupovsk6 Hory; TS, Tyrrhenian Sea.
magmatism and areas of basement uplift. In a number of areas (e.g. Rhine Graben; Massif Central), however, there is evidence for minor early Tertiary magmatic activity, which pre-dates the onset of rifting (Fig. 3). The largest Cenozoic rift within Central Europe, the Rhine Graben, is about 300 km long and 35-40 km wide. It trends N N E - S S W oblique to the NE-trending structural grain of the Variscan crystalline basement of Europe (Moldanubian and Saxo-Thuringian terranes; Fig. 2). The northern end of the rift is located to the SE of the Rhenish Massif at the boundary between the Saxo-Thuringian and Rheno-Hercynian Variscan basement terranes. Extension and subsidence occurred mainly between Oligocene (c. 35 Ma) and Miocene times. Subsidence in the southern part of the rift was interrupted by basement uplift and the magmatic activity of the Kaiserstuhl volcano (Keller et al. 1990). It is notable that although the Rhine Graben is the most highly extended part of the European rift system, it is largely non-magmatic for much of its length, suggesting that lithospheric extension and decompression-induced partial melting of the upper mantle is not necessarily the main cause of magma generation within the European volcanic province. The Miocene volcanic complex of the Vogelsberg is located at the northern end of the Rhine Graben where it splits into two branches
Cenozoic volcanism in central Germany is concentrated in a 350km long, east-west-trending zone extending from the Eifel in the west to the Rh6n-Heldberg area in the east (Fig. 2; Wedepohl & Baumann 1999). Volcanic activity started during the Eocene and Oligocene in both the eastern and western extremities of the belt (e.g. Hocheifel (45-24 Ma); Fekiacova et al. 2006). The climax of volcanic activity occurred between 16 and 18 Ma in the Vogelsberg volcano in the central part of the belt. Volcanism ceased about 5 Ma ago followed by a Quaternary resurgence of activity in the East and West Eifel. Most of the volcanic rocks are relatively primitive alkali olivine basalts, nepheline basanites and olivine nephelinites; quartz tholeiites, however, occur in the Vogelsberg and North Hessian Depression. Locally (e.g. Eifel, Siebengebirge, Westerwald and Rh6n) extreme differentiation of the parental mafic magmas produced phonolites and trachytes. In the Eifel district predominantly potassic magmas, including leucitites, were erupted during the Quaternary. The Miocene Vogelsberg is a shield volcano (Bogaard & W6rner 2003), which erupted basanites, alkali basalts, quartz tholeiites and limited volumes of highly evolved magmas ranging from hawaiite to trachyte. It has an eruptive volume of c. 600 km 3, probably making it the largest volcanic centre within the European volcanic province (Jung & Masberg 1998). The volcano is located to the east of the Rhenish Massif, close to the triple junction of the Rhine, Ruhr and Leine grabens (Fig. 2). Volcanism commenced in the Early Miocene (c. 22-23 Ma); however, the main phase of activity began at c. 18 Ma and peaked between 16 and 17 Ma. The Rh6n and Northern Hessian Depression volcanic fields are closely related, both spatially and temporally, to the Vogelsberg and erupted a similar range of magma types (Wedepohl et al. 1994; Jung & Hoernes 2000). In the East and West Eifel volcanic fields about 300, typically small-volume, eruptions occurred from monogenetic centres between 700 and 10.8 ka By (Schmincke et al. 1983; W6rner et al. 1986; Schmincke 2006), associated with about 250 m of uplift. The volume of magma erupted is small (about 15 kin3), but the actual volume of magma generated at mantle depths must have been significantly greater (70-100 kin3; G. W6rner, pers. comm.). Two geochemically, spatially and temporally distinct groups of sodic-potassic alkaline volcanic rocks were erupted in the East Eifel; in the NW these include nephelinites, leucitites and their differentiates (erupted >400 ka), whereas in the SE basanites and their differentiates predominate (erupted between 400 and 10 ka; Lippolt et al. 1990). The West Eifel volcanic field covers an area ofc. 600 km 2 and comprises about 240 volcanic centres; these erupted predominantly leucitites and nephelinites with subordinate basanites. At c. 12.9 ka there was a major Plinian eruption of the Laacher See volcano, which produced c. 6.3 km 3 of phonolitic tephra, causing a major environmental impact (Litt et al. 2003). Volcanic activity in the Westerwald started in the Oligocene with the eruption of basalts and trachytes (Schreiber & Rotsch 1998); the main phase of activity had ended by 20 Ma, although there were short periods of reactivation, with the eruption of basalts, in the Miocene and Pleistocene (Fuhrmann & Lippolt 1990). Volcanic activity appears to have been synchronous with minor uplift of the Rhenish Massif, which commenced at the end of the Oligocene, strengthened during the Quaternary and continues to the present day (Meyer et al. 1983). On the basis of palaeomagnetic data, Schreiber & Rotsch (1998) proposed that the northeastern part of the Rhenish Massif has rotated clockwise by 10-16 ~ since the late Oligocene, associated
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151
Fig. 4. Summary of the changing distribution and intensity of Tertiary-Quaternary magmatism throughout Europe and the Mediterranean region.
(See text for details). with a system of dextral strike-slip faults. The Quaternary West and East Eifel volcanic fields are located in the non-rotated western Rhenish Massif block. Block rotation is considered to have initiated in the Late Oligocene as a result of small changes in the direction of the maximum horizontal compressive stress from NNW to NW. Southern Germany
Volcanism in southern Germany is confined to a few small regions including the Urach and Hegau provinces to the east of the Rhine Graben, the rift flanks of the Rhine Graben and the Kaiserstuhl volcano, which is axially located within the graben where it bisects the Vosges-Black Forest dome (Keller et al. 1990; Glahn et al. 1992; Fig. 2). On the basis of K - A r dating and stratigraphic constraints it is likely that the main phase of volcanic activity ranges from about 45 to 15 Ma (Keller et al. 2002). However, Keller et al. (2002) have recently dated amphibole phenocrysts from an olivine melilitite dyke (Trois Epis) in the Vosges at 6 0 . 9 _ 0.6 Ma; this suggests that magmatism began some 15 Ma before the onset of graben formation, contemporaneous with the onset of major horizontal crustal shortening in the Western Alps (Gebauer 1999) and a major phase of foreland compression (Ziegler et al. 1995).
Magmatism occurs as a series of dykes, plugs or necks and diatreme pipes, concentrated in two sectors: (1) the Vosges-Black Forest Dome, which is the location of the maximum updoming of the Rhine Graben rift flanks, and the only axially located volcano in the graben (Kaiserstuhl); (2) in the north between Heidelberg and Frankfurt, mostly in the crystalline basement of the Odenwald. Scattered volcanic centres also occur along the flanks of the rift (e.g. Mahlberg). The primary magmas are highly undersaturated mafic alkaline types, predominantly olivine nephelinites and olivine melilitites with high Mg-numbers, Ni and Cr contents. The Miocene Kaiserstuhl complex (15-18 Ma) is an alkaline carbonatite complex that also includes potassic magmas (Schleicher et al. 1990). The Urach province is an olivine melilitite diatreme field with more than 350 individual volcanic necks for which K - A t ages range from 11 to 17 Ma (Lippolt et al. 1973). Most of the diatremes are composed of tufts of olivine melilitite and olivine melilite nephelinite. The main period of activity is in the mid-Miocene, from 16 to 17 Ma. There does not appear to be any correlation between fault tectonics and the location of the diatremes, although the majority are located in a synclinal structure, the 'Urach Trough', in which subsidence has occurred since mid-Triassic times. The Hegau volcanic field, some 100 km further south, has a greater variety of magmatic rock types including olivine
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melilitite, olivine-nepheline melilitite and phonolite, diatremefacies pyroclastic deposits and carbonatites. K - A r ages range from 7 to 15 Ma, with emplacement of olivine melilitites between 8.5 and 12 Ma.
M a s s i f Central
Most of the alkaline magmatic activity in France has occurred within the uplifted Variscan basement of the Massif Central, with subordinate amounts further south in the Languedoc (Wilson & Patterson 2001). The Massif is divided by a NNE-SSW-trending late Variscan strike-slip fault, the Sillon Houiller, which represents a major discontinuity between two distinct lithospheric domains (Fig. 2; Alard et al. 1996). The eastern Vosges-Auvergne domain is distinguished from the western Limousin domain by thinner crust ( 1.5, whereas anorogenic magmas always have a KzO/Na20 ratio < 1. Caution must be exercised, however, as there are also examples of low-K subduction-related volcanic suites. Much
Bulgaria Italy, Sicily and Sardinia
Spain
Sources of Data Schmincke et al. (1983); Mertes & Schmincke (1985); Wrrner et al. (1986); Lippolt et al. (1990); Wedepohl et al. (1994); Jung & Masberg (1998); Wedepohl & Baumann (1999); Jung & Hoernes (2000); Wedepohl (2000); Bogaard & W~3rner(2003); Wilson & Rosenbaum (unpubl. data) Lippolt et al. (1973); Keller et al. (1990, 2002); Schleicher et al. (1990, 1991); Hegner et al. (1995); Wilson et al. (1995b) Patterson (1996); Wilson & Rosenbaum(unpubl. data) Downes (1983, 1987); Wrrner et al. (1986); Briot et al. (1991); Wilson et al. (1995a); Patterson (1996); Wilson & Patterson (2001); Wilson & Rosenbaum (unpubl. data) Kopecky (1966); Blusztajn & Hart (1989); Shrbeny (1995); Ulrych & Pivec (1997); Ulrych et al. (2001); Wilson & Rosenbaum (unpubl. data) Salters et al. (1988); Szabo et al. (1992); Embey-lsztinet al. (1993); Downes et al. (1995a,b); Dobosi et al. (1995); Embey-Isztin& Dobosi (1995); Harangi et al. (1995); Pecskay et al. (1995); Wilson& Rosenbaum(unpubl. data) Pamic et al. (1995); Tari & Pamic (1998); Jovanovic et al. (2001) Vaselli et al. (1997); Marchev et al. (1998) Beccaluva et al. (1977, 1981, 1987, 1998, 2002); Carter & Civetta (1977); Cioni et al. (1982); Condomines et al. (1982); Dostal et al. (1982); Rutter (1987); Clochiatti et al. (1988); Calanchi et al. (1989); Montanini& Villa (1993); Gillot et al. (1994); De Vecchi & Sedea (1995); Milani (1996); D'Orazio et al. (1997); Tanguy et al. (1997); Gasperiniet al. (2000); Armienti et al. (2004); Corsaro & Pompilio (2004) L6pez-Ruiz et al. (1993); Cebri~ & L6pez-Ruiz (1995); Cebrifiet al. (2000)
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M. WILSON & H. DOWNES
Fig. 5. Total alkali-silica variation diagrams for the most primitive mafic magmatic rocks. The classification boundaries are from Le Bas et al. (1986). The dashed grey line in (c) marks the subdivision between alkaline and sub-alkaline magma series. Sources of data are given in Table I. BTA, basaltic trachyandesite; TB, trachy-basalt.
continental intra-plate volcanism (e.g. Wilson & Downes 1991). Locally where the alkaline magmatism post-dates a recent episode of subduction-related magmatism (e.g. Sardinia, Pannonian Basin, Dinarides, Bulgaria, western Turkey), the earliest magmatic rocks may preserve the fingerprint of earlier fluxing of the upper mantle by subduction-zone fluids. Localized oceanic spreading centres within the Mediterranean domain preferentially sample a depleted mantle source component, similar to the source of mid-ocean ridge basalts (MORB), although in the Tyrrhenian Sea this is clearly modified by a subduction-related fluid flux (Wilson & Bianchini 1999).
S r - N d - P b isotope and trace element chemistry of the most primitive mafic magmas
In Sr-Nd isotope space (Fig. 6) the European basalts define a linear array trending f r o m 143Nd/a44Nd ratios of around 0.5130 towards Bulk Earth values (0.51264); this array may reflect mixing of partial melts derived from different mantle source components (e.g. Wilson & Downes 1991; Wilson & Bianchini 1999; Wilson & Patterson 2001). An isotopically distinct group of quartz tholeiites from central Germany fall below the array; this may be attributable to crustal contamination of the magmas. A series of broadly linear arrays are also evident in Nd-Pb and P b - P b isotope space (Fig. 7), fanning away from a common focal point, although there is considerable dispersion of the data. These S r - N d - P b isotope signatures have been attributed to the mixing of partial melts derived from a common, possibly plume-related, asthenospheric mantle source component known
as the Low Velocity Component (LVC; Hoernle et al. 1995) or European Asthenospheric Reservoir (EAR; Cebrifi & Wilson 1995) and a number of regionally heterogeneous sub-continental lithospheric mantle components (Cebrifi & Wilson 1995; Granet et al. 1995; Hoernle et al. 1995). It is generally accepted that the increasing degree of undersaturation in S i O 2 within the sequence olivine tholeiite-alkali olivine basalt-basanite-nephelinite-melilitite principally results from decreasing degrees of partial melting at increasing depth in the mantle (e.g. Wilson 1989). Constraints on both the depth and degree of partial melting can be provided by the concentration of highly incompatible trace elements in the most primitive mafic magmas, the enrichment of light rare earth elements (LREE) over heavy rare earth elements (HREE) and the CaO/AI203 ratio. Trace element ratios (e.g. La/Yb or La/Sm) where the numerator has a greater degree of incompatibility than the denominator exhibit a regular decrease from nephelinite and basanite to olivine tholeiite, consistent with increasing degrees of partial melting. High CaO/AI203, La/Yb and Nb/Y ratios suggest that the alkali basaltic (sensu lato) magmas were derived from a garnet-bearing mantle source. Enrichment of highly incompatible trace elements and LREE, and strong fractionation of LREE over HREE, can, however, be explained either by moderate degrees of partial melting of an enriched source or smaller degrees of melting of a more depleted source. Figure 8 illustrates the variation of Nb/Y v. Zr/Nb compared with model melting curves for 0.5-5% partial melting of spineland garnet-peridotite facies mantle from Harangi (2001). This clearly indicates that, in most of the volcanic provinces, the mantle partial melting column spans the transition from garnetto spinel-peridotite facies mantle, probably close to the base of
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Fig. 6. Variation of 143Nd/144Ndv. S7Sr/86Srfor the most primitive mafic volcanic rocks. EAR, isotopic composition of the European Asthenospheric Reservoir from Cebri~i& Wilson (1995). BE, Bulk Earth. Sources of data are given in Table 1.
Fig. 7. (a) Variation of 143Nd/144Ndv. 2~176 for the most primitive mafic volcanic rocks. (b) Variation of 2~176 v. 2~176 for the most primitive mafic volcanic rocks. EAR, isotopic composition of the European Asthenospheric Reservoir from Cebrifi & Wilson (1995). Sources of data are given in Table 1.
the continental lithosphere (see Bogaard & Wrrner 2003). Degrees of partial melting are typically less than 1%. Only in those regions in which sub-alkaline tholeiitic basalts occur (e.g. central Germany, Pannonian Basin and the Iblean Plateau, Sicily) does the degree of partial melting approach 5%; these tholeiitic basalts clearly equilibrated at somewhat shallower depths in spinel-peridotite facies mantle and may include a significant lithospheric mantle source component in their petrogenesis. The insets in Figure 8 show the variation of Nb/Y v. K20/Na20. There is generally a poor correlation between these two parameters. However, it is clear that the most sodic basalts (lowest K 2 0 / Na20) are the sub-alkaline tholeiites. The distinctive melilitites from the Urach and Hegau volcanic fields to the east of the Rhine Graben have the highest Nb/Y ratios, consistent with an extremely low degree of partial melting (50kin; Vissers et al. 1995). During the Early Miocene, rapid post-collisional extension took place, forming the Alboran basin underlain by thin continental
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM lithosphere. However, the mechanism of this process is highly debated. The competing geodynamic models involve: (1) back-arc extension behind the westward retreat of an east-directed subduction zone along the present Gibraltar arc or the Horseshoe seamounts in the eastern Atlantic (Royden 1993; Lonergan & White 1997; Gelabert et al. 2002; Gill et al. 2004); (2) detachment of the subducted slab (Spakman 1990; Blanco & Spakman 1993; Carminati et al. 1998; Zeck 1999; Calvert et al. 2000); (3) delamination of the lithospheric mantle (Garcia-Duefias et al. 1992; Docherty & Banda 1995); (4) convective removal of the lower part of the overthickened lithosphere (Platt & Vissers 1989; Turner et al. 1999). To the east, the presence of a Tethyan oceanic basin allowed the retreat of the Apennines-Maghrebides subduction zone (Carminati et al. 1998; Gueguen et al. 1998). Gelabert et al. (2002) argued that longitudinal shortening controlled the development of this arcuate subduction belt. They suggested that the subducting slab was split into two main fragments (Apennines and Kabylian slabs) retreating east and SE, respectively. Slab roll-back is explained by the sinking of dense Mesozoic oceanic lithosphere as a result of gravitational instability (Elsasser 1971; Malinverno & Ryan 1986; Royden 1993), by global eastward asthenospheric mantle flow (Doglioni 1992) or by lateral expulsion of asthenospheric material that was shortened and squeezed by plate convergence (Gelabert et al. 2002). Lithospheric extension behind the retreating subduction zone resulted in the formation of the Liguro-Provenqal, Algerian and Tyrrhenian basins underlain by newly formed oceanic crust, whereas the Valencia trough is underlain by thin continental lithosphere. Continental collision is thought to be associated with slab detachment along the north African margin during the Mid-Miocene (c. 16 Ma; Carminati et al. 1998; Coulon et al. 2002). Removal of the southward component of roll-back induced the eastward migration of the Apenninic arc accompanied by eastward migration of extension behind the subduction zone. Wortel & Spakman (1992) and van der Muelen et al. (1998) suggested Late Miocene-Pleistocene slab detachment beneath the Apennines based on seismic tomographic models and the lateral shifts of Apenninic depocentres. In contrast, Doglioni et al. (1994) emphasized the different roll-back rates along the arc, splitting it at least two 'sub-arc' portions. The subducting slab of the Ionian ocean still continues beneath Calabria. Behind it, new oceanic crust has been formed beneath the Vavilov and Marsili basins. Seismic tomographic models show positive seismic anomalies above the 670 km discontinuity beneath the entire Mediterranean region including the Pannonian and Aegean areas (Wortel & Spakman 2000; Piromallo et al. 2001; Piromallo & Morelli 2003). This is interpreted as accumulation of subducted residual material. In contrast to the widely accepted subduction-related models, a sharply different geodynamic scenario (i.e. a relationship with continental extension and/or upwelling mantle plume) has also been suggested to explain the evolution of the Central Mediterranean (Vollmer 1989; Lavecchia & Stoppa 1996; Ayuso et al. 1998; Lavecchia et al. 2003; Bell et al. 2004). Lavecchia et al. (2003) argued that the deformation style of the central Apennine fold-and-thrust belt, the absence of an accretionary wedge above the assumed subduction plane and the occurrence of ultra-alkaline and carbonatitic magmas within the Apennine mountain chain are evidence against the classic subduction-related models. They proposed that plume-induced lithospheric stretching and local-scale rift push-induced crustal shortening form a viable alternative model for the evolution of the Central Mediterranean region (Lavecchia et al. 2003; Bell et al. 2004). Closure of the Tethyan (Vardar) oceanic branches occurred during the Late Cretaceous-Palaeocene in the Dinaride region followed by collisional (Eocene) and post-collisional (Oligocene-Early Miocene) stages (Cvetkovi6 et al. 2004). In the Aegean-Anatolian region, north-dipping subduction terminated
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by the Eocene, when continental collision occurred along the Vardar-Izmir-Ankara suture zone (~eng6r & Yilmaz 1981). The subduction zone has migrated southwestward and its present position is found along the Hellenic-Cyprus arc (Doglioni et al. 2002). A continuous descending lithospheric slab was detected beneath the Hellenic arc by seismic tomography (Wortel & Spakman 1992, 2000). The subducted slab appears to cross the 670 km discontinuity and penetrates the lower mantle. Further east, collision between the Arabian and Anatolian plates took place during the Eocene ($engrr & Kidd 1979; Pearce et al. 1990). This collision also led to the tectonic escape of the Anatolian plate by right-lateral strike-slip movement along the North Anatolian Fault and left-lateral strike-slip along the East Anatolian Fault during the Mid-Miocene (McKenzie 1972; Dewey & Sengrr 1979). It was accompanied by widespread crustal extension and lithospheric thinning in Western Anatolia and the Aegean. The reason for the extension is, however, highly controversial. Dewey (1988) suggested gravitational collapse of the overthickened lithosphere. The orogenic collapse model was also accepted by Seyito~lu & Scott (1996) and Gautier et al. (1999), but they suggested that it occurred earlier (i.e. during the latest Oligocene to Early Miocene (24-20 Ma) time) and therefore it could not be associated with the tectonic escape of the Anatolian block. Other workers have emphasized the back-arc type extension of the Aegean region behind the retreating subduction along the Hellenic arc (McKenzie 1978; Le Pichon & Angelier 1979; Meulenkamp et al. 1988; Pe-Piper & Piper 1989). In Western Anatolia, Aldanmaz et al. (2000) invoked delamination of the lower lithosphere to explain the extension and related magmatism. Doglioni et al. (2002) ruled out the influence of both the westward extrusion of the Anatolian block and the collapse of overthickened lithosphere, and suggested an alternative geodynamic scenario for the Eastern Mediterranean. They interpreted the extension in the Aegean-Western Anatolian region as a result of the differential convergence rates between the northeastward-dipping subduction of Africa relative to the disrupted Eurasian lithospheres. Extension could be attributed to the faster southeastward motion of Greece relative to Anatolia. Thus, Doglioni et al. (2002) argued that the Aegean region cannot be considered as a classic back-arc basin. The next sections will outline the main characteristics of the Tertiary to Quaternary subduction systems in Southern Europe, the geochemical features of the magmatism, and the possible geodynamic relationships between magmatism, subduction and post-collisional processes. Figure 5 summarizes the age distribution of magmatism in this region, separated into orogenic and anorogenic types. Alpine subduction system
Subduction of Tethyan oceanic slabs occurred beneath the Alps during the Late Cretaceous to Palaeogene; however, no prominent subduction-related volcanism appears to have taken place during this period. The only evidence for subduction-related volcanic eruptions comes from Early Eocene andesitic clasts found in flysch sediments (Waibel 1993; Rahn et al. 1995). A characteristic feature of the Alpine collisional orogen is the occurrence of a chain of Oligocene to Early Miocene intrusions and dykes along the Periadriatic and Insubric lines. They continue eastward in the Pannonian Basin along the Balaton line (Downes et al. 1995; Benedek 2002) and southeastwards in the Dinarides (Pamid et al. 2002). These igneous rocks have a bimodal character (granodioritic-tonalitic intrusions and basaltic dykes; Exner 1976; Cortecci et al. 1979; Bellieni et al. 1981; Dupuy et al. 1982; Beccaluva et al. 1983; Ulmer et al. 1983; Kagami et al. 1991; Mtiller et al. 1992; von Blanckenburg & Davies 1995;
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S. HARANGIE T A L .
Paleocene
Eocene
O/igocene Miocene
Pliocene Quat Betics Alboran Rif ValenciaTrough Provence Sardinia-Corsica Tyrrhenianbasin Tuscan Region Roman-CampanianProvinces Aeol:ianIslandsand seamounts Sicily Alps (Periadriatic-Insubricline & Veneto) Western Carpathians& PannonianBasin Eastern Carpathians Apuseni Dinarides Rhodope-Thrace NE-Aegean - WesternAnatolia Central Anatolia South Aegean arc
I
I 70
60
50
40
30
20
lO
0
Age (Ma) t
Calc-alkalineto shoshonitic magmatism ('orogenic type')
j
Alkaline(sodic) magmatlsm ('anorogenic type')
Berger e t al. 1996). In addition, calc-alkaline andesites, shoshonites and ultrapotassic rocks (lamproites) also occur in subvolcanic facies (Deutsch 1984; Venturelli e t al. 1984b; Altherr e t al. 1995). All of these igneous rocks are characterized by 'subduction-related' geochemical features. Suggested models for the origin of these igneous rocks include subduction (e.g. Tollmann 1987; Kagami e t al. 1991; Waibel 1993), extension (e.g. Laubscher 1983), and gradual slab detachment (von Blanckenburg & Davies 1995; von Blanckenburg e t al. 1998). The source regions of the primary magmas of the Periadriatic line are inferred to be in the lithospheric mantle (Venturelli e t al. 1984b; Kagami e t al. 1991 ; yon Blanckenburg 1992). Mafic melts could have subsequently mixed with silicic magmas generated in the lower crust. Alkaline mafic rocks ('anorogenic' type) crop out only south of the Eastern Alps, in the Veneto region (De Vecchi & Sedea 1995; M i l a n e t al. 1999; Macera e t al. 2003; Fig. 1). The volcanism occurred in two stages, from the Late Palaeocene to Early Oligocene (30-35 Ma) and during the Early Miocene. It resulted in alkaline and tholeiitic basalts and basanites with subordinate trachytes and rhyolites (De Vecchi & Sedea 1995). The mafic volcanic rocks show an ocean island basalt (OIB)-like composition without any sign of subduction-related component. De Vecchi & Sedea (1995) and Milani e t al. (1999) interpreted this volcanism as related to lithospheric extension in the Southern Alps (Zampieri 1995). In contrast, Macera e t al. (2003) invoked slab detachment and the ensuing rise of a deep mantle plume into the lithospheric gap.
Betic-Alboran-Rif
province
(Western
Mediterranean)
Tertiary to Quaternary volcanic rocks in the Western Mediterranean are found in central Spain (Calatrava province), the Olot
qr Ultrapotassic magmat~sm
W M
C M
A L C A P A
D E
M
Fig. 5. Age distributionof the Tertiary to Quaternary magmatismin the AlpineMediterranean region. Data are from Bellon et al. (1983), Fytikas et al. (1984), Beccaluva et al. (1985, 1987, 1991), Di Battistini et al. (1987), Peccerilloet al. (1987), Aparico et al. (1991), Conticelli & Peccerillo (1992), Martf et al. (1992), Seyito~lu& Scott (1992), Serri et al. (1993), Louni-Haciniet al. (1995), Pamid et al. (1995, 2002), P~cskay et aI. (1995), Christofideset al. (1998), E1 Bakkali et al. (1998), Harkovska et al. (1998), Marchev et al. (1998), von Blanckenburget al. (1998), Wilson & Bianchini(1999), Aldanmaz et al. (2000), Ro~u et al. (2001), Coulon et al. (2002), Cvetkovid et al. (2004), Duggen et al. (2004), and further references therein. WM, Western Mediterranean; CM, Central Mediterranean;ALCAPA, Alps- Carpathians-Pannonian region; DEM, Dinafides and Eastern Mediterranean.
region, the Valencia trough, SE Spain (Betics), the Alboran basin and along the coast of Northern Africa (Morocco to Algeria; Fig. 1). The Calatrava and Olot regions are characterized by Late Miocene to Quaternary alkaline basaltic and leucititic rocks (Cebrifi & Lopez-Ruiz 1995; Cebri~ e t al. 2000), similar to those occurring in the European Rift Zone (Wilson & Downes 1991). In the other areas calc-alkaline, high-K calc-alkaline, shoshonite and lamproites can be found in addition to late-stage alkali basalts. Duggen e t al. (2003) pointed out that the transition of calc-alkaline ('orogenic') to alkaline ('anorogenic') magmatism (6.3-4.8 Ma) was coeval with the Messinian salinity crisis (5.96-5.33 Ma; i.e. the desiccation of the Mediterranean sea as a result of closure of the marine gateway). The 'orogenic' volcanism started around Malaga with intrusion of tholeiitic (basalts to andesite; Fig. 6) dykes into the Alboran block during the early Oligocene. In addition, high-K dacites also occur in this area. It was followed by volcanism in the Valencia trough during the Late Oligocene that continued during the Miocene and to the present. Calc-alkaline volcanism forming mostly dacitic to rhyolitic pyroclastic deposits characterized the first stage of volcanic activity, whereas alkaline basaltic magmas erupted during the later volcanic stage (Mart/et al. 1992). The alkaline basalts have intraplate (OIB) geochemical affinity and often contain ultramafic xenoliths. This scenario could be explained by progressive extension of the continental lithosphere and a change of the source region from lithospheric to asthenospheric. The Alboran basin is also underlain by thin continental crust similar to the Valencia trough. Volcanic rocks on Alboran island and the sea floor have been dated between 11 and 7 Ma (Duggen e t al. 2004). They are mostly low-K tholeiitic rocks with clear 'subduction-related' geochemical features (Duggen e t al. 2004; Gill e t al. 2004; Fig. 7). More widespread calc-alkaline
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM
173
Fig. 6. SiO2 v. K20 diagrams (Gill 1981; Sh, shoshonite series; HKCA, High-K calc-alkaline series; CA, calc-alkaline series; Th, low-K tholeiitic series; B, basalt; BA, basaltic andesite; A, andesite; D, dacite; R, rhyolite) for the 'subduction-related' volcanic rocks from the various segments of the Alpine-Mediterranean region. The wide compositional variations should be noted. Western Mediterranean sources: Nixon et al. (1984), Venturelli et al. (1984a), Zeck et al. (1998), Benito et al. (1999), Turner et al. (1999), Duggen et al. (2004), Gill et al. (2004). Central Mediterranean sources (A1-Fi-Sa-Pa: Alicudi, Filicudi, Salina and Panarea): Rogers et al. (1985), Ellam et aL (1988), Crisci et al. (1991), Conticelli & Peccerillo (1992), Francalanci et al. (1993), Peccerillo et al. (1993), Ayuso et al. (1998), Del Moro et aL (1998), De Astis et al. (2000), Gertisser & Keller (2000), Downes et al. (2001). Carpathian-Pannonian region sources: Downes et al. (1995a), Mason et al. (1996), Harangi et al. (2001, 2005), Seghedi et al. (2001, 2004). Dinarides-Eastern Mediterranean sources (WA, Western Anatolia; CA, Central Anatolia): Mitropoulos et al. (1987), Huijsmans et al. (1988), Pe-Piper & Piper (1989), Seyito~hi & Scott (1992), Wilson et al. (1997), Francalanci et al. (1998), Kiirkqtioglu et al. (1998), Tankut et al. (1998), Temel et al. (1998a,b), Aldanmaz et al. (2000), Cvetkovi6 et al. (2004).
Fig. 7. Normal-MORB (N-MORB; Pearce & Parkinson, 1993) normalized multi-element diagrams for representative samples of the various segments of the Alpine-Mediterranean region. (For data sources see Fig. 6.) Carp., Carpathians; UP, ultrapotassic.
174
S. HARANGIET AL.
to shoshonitic and ultrapotassic volcanism occurred on the SE coast of Spain and in Northern Africa from the Early Miocene to Pliocene (Zeck 1970, 1992, 1998; Nixon et al. 1984; Venturelli et al. 1984a, 1988; Hertogen et al. 1985; Di Battistini et al. 1987; Louni-Hacini et al. 1995; E1 Bakkali et al. 1998; Benito et al. 1999; Turner et al. 1999; Coulon et al. 2002; Duggen et al. 2004; Gill et al. 2004; Fig. 6). Calc-alkaline volcanism was associated with intrusion of granitoid magmas in Northern Africa (Fourcade et al. 2001) and southern Spain (Zeck et al. 1989; Duggen et al. 2004). The calc-alkaline volcanism resulted in andesites and dacites with subordinate rhyolites and shoshonites (Fig. 6). Sporadic cordierite- and garnet-bearing dacites were interpreted as anatectic magmas (Zeck 1970, 1992). Late Miocene ultrapotassic lamproites are found in the central and northern part of the calc-alkaline volcanic belt of SE Spain (Nixon et al. 1984; Venturelli et al. 1984a, 1988; Hertogen et al. 1985). Throughout the region, sporadic eruptions of alkaline mafic magmas followed the calc-alkaline magmatism (El Bakkali et al. 1998; Coulon et al. 2002; Duggen et al. 2004). The geodynamic setting of the Western Mediterranean calc-alkaline volcanic activity is ambiguous. The models can be divided into the following groups: (1) subduction-related; (2) subduction break-off; (3) delamination of lithospheric mantle as a result of gravitational collapse; (4) convective removal of the lower lithosphere. Torres-Roldfin et al. (1986), Royden (1993), Lonergan & White (1997), Duggen et al. (2003, 2004) and Gill et al. (2004) assumed that contemporaneous subduction occurred with the calc-alkaline volcanism. Geophysical data indicate an east-dipping subducted slab (Gutscher et al. 2002) beneath the Alboran region. Duggen et al. (2004) and Gill et al. (2004) emphasized that the 'subduction-related' nature and particularly the strong depletion in the light REE (LREE) and HFSE of the Alboran tholeiites (Fig. 7) could only be explained by formation in a metasomatized mantle wedge above a subducted slab. Furthermore, they assumed that all the other calc-alkaline volcanic rocks in the Betic-Rif province could be generated in the same geodynamic setting. Blanco & Spakman (1993) and Calvert et al. (2000) argued, however, that the seismic tomography models show a detached near-vertical lithospheric slab from about 180200 km down to the 670 km discontinuity beneath the Alboran region. Zeck (1996) considered that slab break-off could have had a major role in melt generation. Influx of hot asthenospheric mantle into the widening gap above the sinking slab induced partial melting in the overlying lithosphere (particularly in the lower crust). The close relationship between the distribution of volcanism in the Alboran volcanic province and the surface projection of the sinking slab was used by Zeck (1996) to support this model. Fourcade et al. (2001) and Coulon et al. (2002) also invoked slab break-off to explain the calc-alkaline to alkaline magmatism in northern Algeria. Other workers (Venturelli et al. 1984a; Platt & Vissers 1989; Zeck 1996; Benito et al. 1999; Turner et al. 1999) argued that the Betic-Alboran volcanism was post-collisional, following Late Cretaceous to Oligocene subduction and Late Oligocene to Early Miocene continental collision. Benito et al. (1999), Turner et al. (1999) and Coulon et al. (2002) suggested that the primary melts were generated in the lithospheric mantle, which had been metasomatized previously by fluids derived from subducted pelagic sediments. These mantle-derived magmas subsequently mixed with crustal melts. Zeck (1970, 1992, 1998) argued for a crustal anatectic origin for the calc-alkaline magmas of southern Spain. Platt & Vissers (1989), Benito et al. (1999) and Turner et al. (1999) emphasized that melt generation occurred by decompression melting caused by extensional collapse of the overthickened orogenic wedge or convective removal of the lithospheric root. In North Africa, E1 Bakkali et al. (1998) also suggested an extension-related origin for the calc-alkaline to potassic magmas of the Eastern Rif (Morocco).
Central Mediterranean
(Italy)
Tertiary-Quaternary volcanism in the Central Mediterranean resulted in extremely variable magmatic rocks including tholeiites (Vavilov basin, Tyrrhenian basin), calc-alkaline to shoshonitic (Sardinia, Aeolian Islands, Roman Province), ultrapotassic (Corsica, Central Italy) and anatectic rhyolites (Tuscany; Serri 1990; Peccerillo 1999, 2003; Figs 1 and 6). Alkali basaltic rocks with OIB chemistry also occur sporadically in Sardinia (Rutter 1987; Lustrino et al. 2000), the southern Tyrrhenian basin (Serri 1990; Trua et al. 2003), eastern Sicily (Etna and Hyblean plateau; Carter & Civetta 1977; Tonarini et al. 1995; D'Orazio et al. 1997; Tanguy et al. 1997; Trua et al. 1998) and in the Pantelleria rift (Esperanca & Crisci 1995; Civetta et al. 1998). In addition, minor occurrences of carbonatites and melilitites have been described in the central Apennines east of the Roman Province (Stoppa & Lavecchia 1992; Stoppa & Cundari 1995; Stoppa & Woolley 1996). The carbonatitic nature of these rocks has been questioned, however, by Peccerillo (1998) who suggested that they could represent a mixture of silicate magmas and carbonate material, and could be classified as ultrapotassic rocks of kamafugitic affinity. The strongly undersaturated haiiyne-bearing alkaline volcanic rocks of Mt. Vulture (De Fino et al. 1986; Serri 1990; Melluso et al. 1996) also have an exotic position (Fig. 1) and distinct magma source region compared with the Roman Province rocks. Volcanism started in the Early to Mid-Miocene in Sardinia with the eruption of tholeiitic and calc-alkaline magmas (Dostal et al. 1982; Morra et al. 1997; Downes et al. 2001; Fig. 5). The 14 Ma Sisco lamproite in northern Corsica represents the oldest ultrapotassic rock in the Central Mediterranean. After a few million years quiescence, the volcanism rejuvenated in the Tyrrhenian basin, the Tuscan region and in southeastern Sicily (Hyblean Mts) at about 7 - 8 Ma. On the west coast of Italy a gradual younging of the volcanism can be observed towards the SE, with still active volcanoes in Campania (Campi Flegrei, Vesuvius; Santacroce et al. 2003). The distinct alkaline volcanic rocks of Mt. Vulture were formed at 0.7-0.l Ma (Melluso et al. 1996). Volcanic activity in Sardinia rejuvenated with eruption of alkaline mafic magmas from 5.5 to 0.1 Ma (Di Battistini et al. 1990; Lustrino et al. 2000). A southeastward shift of volcanism has been pointed out by Argnani & Savelli (1999). Opening of the southern Tyrrhenian basin was accompanied by formation of seamounts consisting of enriched (E)-MORB to OIB type mafic rocks (Vavilov, Marsili; Serri 1990; Trua et al. 2003). Active volcanism is taking place in the Aeolian Islands (Stromboli, Vulcano), in Etna and the Sicily Channel (Pantelleria). Although most workers suggest that subduction has played an important role in the evolution of the Central Mediterranean (e.g. Keller 1982; Doglioni 1991; Serri et al. 1993), calc-alkaline volcanic rocks are volumetrically subordinate within the magmatic suites. Instead, the characteristic rocks are potassic to ultrapotassic (Fig. 6). Furthermore, the Tertiary to Quaternary volcanic rocks of this region show an extremely variable trace element and isotope chemistry (Peccerillo 2003). The close temporal and spatial relationship of this wide range of magmas indicates a heterogeneous mantle source metasomatized during several distinct events (Peccerillo 1985, 1999; Serri et al. 1993). The strongly potassic character of many of the magmas has been explained either by source contamination by subducted continental crustal material (Peccerillo 1985; Ellam et al. 1989; Conticelli & Peccerillo 1992; Serri et al. 1993; De Astis et al. 2000) or by metasomatism of deep mantle-derived melts (Vollmer 1989; Stoppa & Lavecchia 1992; Ayuso et al. 1998). West-dipping subduction of oceanic lithosphere and possibly thinned continental lithosphere is considered to have terminated in the Late Miocene (c. 13 Ma), thus most of the volcanism in the Central Mediterranean can be regarded as post-collisional.
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM
Evidence for subduction includes the introduction of continental crustal material into the mantle sources of the magmas (Peccerillo 1985, 1999) and the detection of high-velocity material either continuously extending from the surface (beneath Calabria) and accumulating in the transition zone (e.g. Spakman e t al. 1993; Piromallo e t al. 2001; Piromallo & Morelli 2003). Serri e t al. (2003) proposed that delamination and subduction of the Adriatic continental lithosphere related to the continuing collision in the northern Apennines could be a viable mechanism to explain the incorporation of crustal material in the mantle source and the eastward migration of magmatism in central Italy. Active subduction in the region occurs in Calabria. Keller (1982) directly related the recent volcanism in the Aeolian Islands to active subduction process. However, the Aeolian Islands lie 2 0 0 - 3 0 0 km above the Benioff zone (Anderson & Jackson 1987) on thinned continental basement (Schutte 1978). Furthermore, they form rather a ringshaped structure considering also the submarine seamounts at the southern margin of the Marsili basin, which is characterized by oceanic crust. The active volcanoes are located along strike-slip tectonic lines (Beccaluva e t al. 1982; Gasparini e t al. 1982). Thus, an alternative hypothesis for the Aeolian volcanism is a relationship with a back-arc environment, where magma generation is attributed to asthenospheric domal uplift developing along a NW-SE-trending extensional tectonic zone (Crisci e t al. 1991; Mazzuoli e t al. 1995). Nevertheless, subduction could have had a major, probably indirect, influence on the genesis of the magmas (release of aqueous fluids from the downgoing slab and metasomatism of the upper mantle; injection of subducted sedimentary component into the upper mantle; Ellam e t al. 1989; Francalanci e t al. 1993). Compositional features (ratios of incompatible trace elements and radiogenic isotope ratios; Fig. 8) of the potassic rocks of Stromboli are similar to those of
175
the alkaline volcanic rocks of Campania (Vesuvius and Phlegrean Fields), indicating common mantle source regions consisting of a mixture of intraplate and subducted slab-derived (continental sediment) components (De Astis e t al. 2000; Peccerillo 2001). The post-collisional volcanism in Italy has been interpreted by Wortel & Spakman (1992, 2000) as being due to gradual slab detachment, based on the absence of high-velocity structure considered to represent subducted slab beneath the Apennines, whereas a continuous slab was identified beneath southern Italy. However, Piromallo & Morelli (2003) argued that their better resolved model showed more vertical continuity of the fast structure in the top 200 km beneath the northern part of the Apennines. In contrast to the most popular subduction-related models, the presence of a mantle plume and related continental rifting was put forward by Lavecchia e t al. (2003) and Bell e t al. (2004). Gasperini e t al. (2002) also invoked upwelling of deep mantle material beneath southern Italy, but they combined it with the subduction scenario, suggesting a broad window in the Adria plate where deep mantle layers are channelled toward the surface. Lavecchia e t al. (2003) and Bell e t al. (2004) proposed that a plume arising from the core-mantle boundary could be trapped within the transition zone beneath the Ligurian-Tyrrhenian region. Asymmetric growth of the plume head within the transition zone as modelled by Brunet & Yuen (2000) could lead to a volume excess within the asthenosphere and an eastward mantle flow. This is thought to result in an eastward-migrating thinning of the overlying lithosphere. The rift-push forces generated on the eastern side of the extending system could be responsible for the fold-and-thrust belt structure beneath the Apennines. In this model, the high-velocity body above the 670 km depth (Piromallo e t al. 2001) was interpreted as reflection of compositional difference rather then abrupt change in the mantle temperature.
Fig. 8. 87sr/arsr v. 143Nd/144Nd diagrams for the Tertiary-Quaternary volcanic and plutonic rocks of the various segments of the Alpine-Mediterranean region. Data sources are as for Figure 6. Additional data sources are as follows. Western Mediterranean: Cebriit et al. (2000), Central Mediterranean: Carter & Civetta (1977), Hawkesworth & Vollmer (1979), Ellam & Harmon (1990), Esperanca & Crisci (1995), Tonarini et al. (1995, 2001), D'Orazio et al. (1997), Trua et al. (1998), Castorina et al. (2000), Lustrino et al. (2000), Alps-Carpathian-Pannonian region: Juteau et al. (1986), Kagami et al. (1991), von Blanckenburg (1992), yon Blanckenburg et al. (1992, 1998), Embey-Isztin et al. (1993), Harangi et al. (1995), Macera et al. (2003), Dinarides-Eastern Mediterranean: Briqueu et al. (1986), Gtilec (1991), Pamid et al. (1995). SV-Po-CF-Cu, melilitite-carbonatite association in San Venanzo, Polino, Colle Fabri, Cupaello; CAV, calc-alkaline volcanic rocks; DMM, depleted MORB mantle; EMI and EMIl, enriched mantle I and II; HIMU, high Ix (238U/2~ mantle components (Zindler & Hart 1986).
176
S. HARANGIETAL.
Lavecchia et al. (2003) suggested that the fast zones in the transition zone could be a highly depleted and dehydrated plume head, whereas the overlying asthenosphere was enriched by HzO-COz-rich fluids. However, Goes et al. (2000) proposed that the velocity variation in the mantle could be attributed mostly to changes in temperature, whereas the effect of mantle composition could be negligible (< 1%). Further integrated geophysical, structural and geophysical studies are needed to test the contrasting models for the evolution of the Central Mediterranean.
Carpathian-Pannonian
region
The Carpathian-Pannonian region (Fig. 1) shows many features that are similar to those of the Mediterranean subduction systems, such as arcuate and retreating subduction zones, formation of back-arc extensional basins and a wide range of magma types (e.g. Horv~ith & Berckhemer 1982; Csontos et al. 1992; Szab6 et al. 1992; Seghedi et al. 1998; Fodor et al. 1999; Tari et al. 1999; Bada & Horv~ith 2001; Harangi 2001a). Volcanic activity in this region started with eruption of Early Miocene rhyolitic magmas followed by contemporaneous calc-alkaline, silicic and sporadic ultrapotassic volcanism in the Mid- and Late Miocene (Fig. 5). Coeval calc-alkaline and alkaline mafic magmas were erupted during the Late Miocene to Quaternary (Szab6 et al. 1992; Prcskay et al. 1995; Seghedi et al. 1998, 2004; Harangi 2001b). The Miocene (21 - 13 Ma) rhyolitic volcanism resulted in extensive ignimbrite sheets. The rhyodacitic to rhyolitic pumices have 'subduction-related' geochemical features consistent with both mantle and crustal origin. The pyroclastic deposits also contain basaltic andesite and andesite lithic clasts, which are considered as cogenetic with the rhyolites. Harangi et al. (2005) interpreted their petrogenesis inferring mantle-derived mafic magmas mixed with variable amount of crustal melts. The silicic volcanism could represent the initiation of back-arc lithosphere extension (Lexa & Konern~ 1998; Harangi 2001a) or delamination of the lowermost lithosphere beneath the Pannonian Basin (Downes 1996; Seghedi et al. 1998). The decreasing age-corrected 87Sr/86Sr ratios of the pumices indicate a gradually decreasing crustal component in their genesis. A major feature of the region is the Carpathian arc, an arcuate belt of calc-alkaline volcanic complexes composed mostly of andesites and dacites (Fig. 6) along the northern and eastern margin of the Pannonian Basin. They were formed from the MidMiocene to the Quaternary and the last volcanic eruption occurred in the southernmost part of the East Carpathians only 10-40 ka ago (Fig. 5; Prcskay et al. 1995). The major and trace element compositions of these rocks show fairly similar character compared with the large variability of the Western and Central Mediterranean volcanic suites (Figs 6 and 7). However, there are major differences in spatial and temporal evolution and underlying lithospheric structure between the western and eastern segments of the Carpathian volcanic arc. These differences also appear in the geochemistry of the volcanic products, leading Harangi & Downes (2000) and Harangi (2001a) to suggest contrasting origins for the calc-alkaline magmas in the different segments. Calc-alkaline volcanism in the western Carpathian arc could be related directly to the main extensional phase of the Pannonian Basin (Lexa & Konern.~ 1998; Harangi 2001a; Harangi et al. 2001), whereas calc-alkaline magmas in the eastern Carpathian arc could have a closer relationship with subduction, particularly with gradual slab break-off (Mason et al. 1998; Seghedi et al. 1998, 2004). Gradual slab detachment was also proposed in the evolution of the Carpathian arc by other workers (Nem~ok et al. 1998; von Blanckenburg et al. 1998; Wortel & Spakman 2000; Sperner et al. 2002). Coexisting eruptions of alkaline basaltic and shoshonitic magmas in the southernmost part of
the east Carpathians led G~rbacea & Frisch (1998) and Chalot-Prat & G~rbacea (2000) to suggest partial delamination of the lower lithosphere beneath this area. In contrast to these models, Balla (1981), Szab6 et al. (1992) and Downes et al. (1995a) considered that melt generation in the whole calc-alkaline suite was a direct consequence of subduction of the European plate and occurred in the metasomatized mantle wedge above the downgoing slab. Calc-alkaline volcanic rocks also occur far from the Carpathian arc, in the inner part of the Pannonian Basin. Mid-Miocene andesites of the Apuseni Mountains are found about 200 km behind the volcanic front. Ro~u et al. (2001) and Seghedi et al. (2004) suggested that the location and compositions of these rocks is inconsistent with a typical subduction model and can be explained rather by decompression melting of the lower crust and/or of the enriched lithospheric mantle in an extensional regime. Seismic tomographic images indicate a low-velocity anomaly beneath the Carpathian-Pannonian region (except at the southeastern margin of the Carpathians) at shallow depth (Spakman 1990; Wortel & Spakman 2000; Piromallo & Morelli 2003). A fast anomaly was detected beneath the Eastern Carpathians in a depth range of 100-300 km, but it cannot be followed beneath the western Carpathian chain (Piromallo & Morelli 2003). Thus, no evidence is present for a detached subducted slab under the latter area, although Tomek & Hall (1993) interpreted the deep seismic reflection data as evidence for subducted European continental crust beneath the western Carpathians. In the southeastern part of the Carpathians, beneath the Vrancea zone, a weak fast anomaly is present, becoming more pronounced with increasing depth. The localized Vrancea slab is considered to represent the final stage of slab break-off (Wenzel et aL 1998; Sperner et al. 2001) beneath the east Carpathians. The oceanic slab is considered as either already detached from the surface (Wortel & Spakman 2000) or still attached to the continental lithosphere (Fan et al. 1998; Sperner et al. 2001). An approximately 150-200 km thick positive anomaly occurs between 400 and 600 km beneath the entire Carpathian-Pannonian region that is interpreted as accumulation of Mesozoic subducted slab material (Wortel & Spakman 2000).
Dinarides and Hellenides
A continuous belt of Tertiary igneous activity is present from the eastern Alps to the north Aegean crossing the southern Pannonian Basin (Slovenia and Croatia), the Dinarides (Serbia, Macedonia) and the Rhodope-Thrace region (Bulgaria and Greece; Fig. 1; Pami6 et al. 1995, 2002; Christofides et al. 1998, 2001; Harkovska et al. 1998; Marchev et al. 1998; Yilmaz & Polat 1998; Jovanovi6 et al. 2001; Prelevi6 et al. 2001; Cvetkovi6 et al. 2004). Subduction of part of the Tethyan Vardar Ocean occurred during the Late Mesozoic to Early Palaeogene, followed by Eocene collision and Oligocene to Pliocene post-collisional collapse (Karamata & Krsti6 1996; Karamata et al. 1999). This igneous belt comprises Eocene to Oligocene granitoid bodies and basanites, Oligocene to Miocene shoshonites, high-K calc-alkaline volcanic rocks and ultrapotassic rocks (lamproites and leucitites). Most of the Palaeogene granitoids have been interpreted as syncollisional magmas that underwent various degrees of crustal contamination (Christofides et al. 1998; Marchev et al. 1998; Pami6 et al. 2002). The Palaeocene-Eocene basanites in eastern Serbia often contain ultramafic xenoliths (Cvetkovi6 et al. 2001) and have major and trace element composition akin to OIB (Jovanovi6 et al. 2001). Similar xenolith-bearing alkaline mafic rocks also occur in the Rhodopes (Marchev et al. 1998). The primary magmas are inferred to originate in an enriched asthenospheric mantle source. Melt generation could have been triggered either by detachment of the subducted slab resulting in a slab window or by a short extensional event during the collisional phase (Jovanovi6 et al. 2001; Cvetkovi6 et al. 2004). The Oligocene to Early Miocene high-K calc-alkaline and shoshonitic series
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM comprises basalts, basaltic andesites and trachyandesites (Fig. 6). They have relatively high mg-values and high concentration of Ni and Cr, indicative of near-primary magmas. The presence of phlogopite-bearing ultramafic xenoliths clearly indicates a mantle origin for these melts, which have typical 'subductionrelated' compositions (Fig. 7). This signature could be inherited from the lithospheric mantle source metasomatized possibly during the post-collisional or collapse stage (Cvetkovid et al. 2004). The scattered ultrapotassic rocks (minettes, lamproites, leucitites, analcimites) show the most extreme enrichment of incompatible elements of all the Tertiary volcanic rocks in this region (Prelevid et al. 2001). They have fairly similar trace element patterns in the mantle-normalized diagrams to the high-K volcanic rocks, but with more pronounced anomalies. Thus, they could also represent magmas derived from metasomatized lithospheric mantle. Melt generation could be related either to slab break-off (Pamid et al. 2002) or to delamination of the lithospheric root (Cvetkovid et al. 2004). Seismic tomographic images indicate a high-velocity anomaly from about 100 km to 600 km beneath the southern Dinarides and Hellenides region, whereas a low-velocity anomaly was detected beneath the northern Dinarides (Spakman et al. 1993; Goes et al. 1999; Wortel & Spakman 2000). This feature has been interpreted as detachment of a subducted slab in the north, whereas it is still unbroken in the south and continues towards the south Aegean area (Wortel & Spakman 2000). Beneath the Dinarides a north- to NE-dipping subduction was proposed with the opposite polarity to that inferred beneath the Alps (Pamid et al. 2002). Stampfli et al. (2001) suggested that the Vardar Ocean (the Tethyan oceanic branch in the Dinaride-Hellenide region) and the Piedmont-Penninic Ocean (Alpine Tethyan oceanic branch) were not connected during the Mesozoic. Therefore, the two linear Palaeogene igneous belts along the Periadriatic line and along the Dinarides could not belong to the same subduction system.
Eastern Mediterranean
(Greece and Turkey)
Tertiary-Quaternary volcanic activity in this region was characterized by eruption of various magmas (alkaline mafic, calc-alkaline and high-K intermediate to silicic volcanic rocks and sporadic ultrapotassic rocks) in the Aegean and Western to Central Anatolia (Fig. 1; Fytikas et al. 1984; Doglioni et al. 2002). Volcanism occurred in two main phases: eruption of Oligocene to Mid-Miocene calc-alkaline to shoshonitic magmas followed by eruption of alkaline and calc-alkaline magmas during Pliocene to Recent times (Pe-Piper & Piper 1989; Seyito~lu & Scott 1992; Pe-Piper et al. 1995; Aldanmaz et al. 2000; Doglioni et al. 2002). In the Aegean-Western Anatolian region, volcanism started in the Oligocene following the Tethyan collision (Yilmaz et al. 2001). The calc-alkaline volcanism resulted mostly in andesitic to dacitic rocks, and was associated with emplacement of granitic plutons in NW Anatolia. Most of the volcanism occurred, however, during the Early to Mid-Miocene (20-14 Ma), when high-K andesitic to rhyolitic effusive and explosive volcanism with cogenetic plutons characterized the northern Aegean and the Western Anatolia (Figs 5 and 6; Fytikas et al. 1984; Pe-Piper & Piper 1989; Seyito~lu & Scott 1992; Wilson et al. 1997; Altunkayak & Yilmaz 1998; Aldanmaz et al. 2000; Yilmaz et al. 2001). In Western Anatolia, these igneous rocks are distributed mostly along the Izmir-Ankara suture zone. A Mid-Miocene (14-15 Ma) lamproite was reported by Savas~in et al. (2000). During the Mid- to Late Miocene, granitic plutonism took place in the Cycladic and Menderes massifs (Altherr et al. 1982; Innocenti et al. 1982; Delaloye & Bing61 2000). Following about 4 Ma quiescence, the volcanism resumed in the Late Miocene (10 Ma), when alkaline mafic magmas with OIB-like composition erupted mostly in the eastern Aegean and the western margin of Anatolia (Gtile~ 1991; Seyito~lu et al.
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1997; Aldanmaz et al. 2000; Alici et al. 2002). The last eruption of basanitic to phonotephritic magmas occurred in the Kula region at 1.1-0.02Ma (Gtileq 1991). In the Aegean region Pe-Piper et al. (1995) found a southward migration of volcanism. Calc-alkaline volcanism has characterized the Aegean volcanic arc from Pliocene to Recent times (Mann 1983; Barton & Huijsmans 1986; Briqueu et al. 1986; Mitropoulos et al. 1987; Huijsmans et al. 1988; Francalanci et al. 1998). Central Anatolia shows roughly the same volcanic history. MidMiocene to Pliocene high-K calc-alkaline andesitic to rhyolitic magmas erupted along major fault systems at Afyon, Konya and Cappadocia (Figs 5 and 6; Innocenti et al. 1975; Aydar et al. 1995; Alici et al. 1998; Temel et al. 1998a,b). In Cappadocia extensive dacitic to rhyolitic ignimbrite sheets were deposited from the Late Miocene to Quaternary, associated with large andesitic stratovolcanoes and alkali basaltic scoria cones and maars (Pasquare et al. 1988; Le Pennec et al. 1994; Aydar & Gourgaud 1998; Ktirkqtioglu et al. 1998; Temel et al. 1998b). Tertiary to Quaternary volcanism in the Aegean and Western to Central Anatolian region occurred mostly in a post-collisional setting and partly behind active subduction zones (Hellenic and Cyprean). The origin of the Miocene plutonic igneous rocks was interpreted as crustal anatexis related to high-T-medium-P metamorphism (Altherr et al. 1982; Innocenti et al. 1982; Br6cker et al. 1993; Delaloye & Bing61 2000). In general, the 'orogenic' volcanic rocks are potassic (high-K calc-alkaline to shoshonitic), whereas those that occur along the Aegean island arc are calc-alkaline (Fig. 6). Their trace element and isotopic compositions are consistent with involvement of a subduction component (Figs 7 and 8; Keller 1982; Briqueu et al. 1986; Mitropoulos et al. 1987; Huijmans et al. 1988; Gtileq 1991; Robert et al. 1992; Pe-Piper et al. 1995; Seyito~lu et al. 1997; Aldanmaz et al. 2000). On the other hand, the younger alkaline mafic volcanic rocks show an intraplate OIB nature (Seyito~lu & Scott 1992; Seyito~lu et al. 1997; Alici et al. 2002). Nevertheless, most workers consider that melt generation of both volcanic suites was mostly due to decompression melting because of extension of continental lithosphere. Initially, magmas with 'subductionrelated' geochemical signature were generated in the lithospheric mantle regions metasomatized by fluids and melts during an earlier subduction event. Perturbation of these metasomatic portions of the lower lithosphere could take place as a result of either lithospheric thinning to delamination of the lower thermal boundary layer allowing direct contact with upwelling hot asthenosphere (Seyito~lu & Scott 1996; Aldanmaz et al. 2000). In contrast, a closer relationship with subduction was invoked by Innocenti et al. (1975), Temel et al. (1998a,b) and Doglioni et al. (2002) to explain the calc-alkaline volcanism especially in Central Anatolia. They considered that the primary magmas originated in the mantle wedge above the subducting Africa plate and subsequently underwent assimilation and fractional crystallization to produce the intermediate to rhyolitic magmas. The Pliocene to Recent volcanic rocks in the Aegean arc seem to be a clearer candidate to have formed as a consequence of active subduction. Geophysical data clearly indicate a northeastward to northward dipping slab beneath the Aegean region (Wortel & Spakman 1992; Spakman et al. 1993). The present volcanic arc is located 130-140 km above the seismic Benioff zone (Makropoulos & Burton 1984), 200-250 km behind the subduction front. Lithospheric extension, however, may have played an important role in melt generation along the present arc, indicated by the underlying thin lithosphere and predominance of asthenosphere-derived uncontaminated mafic volcanic rocks in Santorini (Mitropoulos et al. 1987). In contrast, Briqueu et al. (1986) considered that there is a close relationship between the active volcanic arc and subduction, and assumed a contribution of a small amount of subducted sedimentary component in the genesis of the arc magmas. The late-stage alkaline basaltic rocks have a composition akin to OIB; therefore, they are interpreted as representing
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S. H A R A N G I ETAL.
asthenosphere-derived magmas. These could originate either in places where lithospheric extension progressed further allowing partial melting of the uprising asthenosphere (Seyito~lu et al. 1997) or along strike-slip zones, where localized stretching could result in production of alkaline magmas (Aldanmaz et al. 2000). Doglioni et al. (2002) suggested that the shallow subducted slab beneath Anatolia could be folded by the isostatic rebound of the mantle beneath the extensional area. The stretching between Greece and Anatolia and the differential velocity of convergence with the underlying slab could have generated a sort of window, allowing upward rise and partial melting of the asthenosphere.
Discussion Petrogenetic features
The previous sections showed that a wide variety of magmatic rocks can be found in all of the different subprovinces of the Alpine-Mediterranean region. Most of them show typical 'subduction-related' composition as reflected by the elevated potassium content (Fig. 6) along with the enrichment of LILE and depletion of HFSE (Fig. 7). However, the large geochemical variability as shown in the SiO2 v. K20 diagrams (Fig. 6) implies that complex petrogenetic processes have operated, involving different mantle sources, contamination by different crustal material and different degrees of fractional crystallization. Another important feature of the orogenic magmatic suites is the common occurrences of potassic-ultrapotassic rocks in each subprovince. They are the most characteristic of the Central Mediterranean area. OIB-type alkaline sodic mafic magmas akin to those erupted in the European foreland (Wilson & Downes 1991, 2006) overlap spatially and temporally with the 'orogenic' volcanism, although they are most characteristic of the later magmatic phases. The majority of these alkaline mafic rocks clearly indicate a distinct mantle source regions unaffected or only slightly affected by subduction-related fluids. On the other hand, interpretation of the origin of the 'subduction-related' magmatic rocks in the Mediterranean region is more difficult, because of the lack of mafic undifferentiated rocks in many areas. In multi-element diagrams that are normalized to mid-ocean ridge basalts (N-MORB; Fig. 7), the Mediterranean orogenic rocks all show fairly similar features such as enrichment in LILE and depletion in HFSE (e.g. negative Nb anomaly). As discussed previously, these characters are signs of a subduction component in the genesis of the magmas. Subduction and subduction-related metasomatism of the mantle wedge can be contemporaneous with the magmatism, but could also precede the volcanic activity. However, this geochemical feature can also be interpreted as contamination by crustal material at shallow level. Radiogenic isotope ratios (e.g. 87 Sr/ 86 Sr, 143N d / 144Nd and PbPb isotope ratios) are not changed by closed-system petrogenetic processes such as partial melting and crystal fractionation; therefore they can be used to characterize the source region of the magmas and to detect possible crustal contamination. Recognition of the nature of the pre-metasomatized mantle source region could be important to constrain the geodynamic evolution of the volcanic areas. In the 1980s four main mantle end-member reservoirs were distinguished based on the radiogenic isotope variation of oceanic basalts (White 1985; Zindler & Hart 1986): depleted MORB-mantle (DMM), high tx (238U/2~ mantle (HIMU) and enriched mantle end-members (EMI and EMII). In addition to these mantle components a primitive mantle reservoir, called as PREMA (Zindler & Hart 1986) or FOZO (Hart et al. 1992) was also suggested to be present in the mantle. These mantle components could represent distinct parts of the mantle, although they could also be spatially related (Hart 1988). Among them, the HIMU and FOZO are often interpreted to relate to upwelling mantle plumes coming from the core-mantle boundary
(Hofmann & White 1982; Weaver 1991; Chauvel et al. 1992; Hart et al. 1992; Hofmann 1997). Alternatively, this geochemical feature can reflect derivation of magmas from metasomatized lithospheric mantle (Hart 1988" Sun & McDonough 1989; Halliday et al. 1995; Niu & O'Hara 2003) and in this case no mantle plume is needed. In continental and convergent margin magmas, these isotope ratios are masked by the signature of continental crust and therefore it is difficult to discriminate between mantle and crustal sources. Subcontinental lithosphere can preserve long-lived geochemical heterogeneity (e.g. high Rb/Sr, low Sm/Nd) and therefore can develop high 87sr/g6sr and low 143Nd/144Nd values with time. Certain lithospheric mantle-derived magmas (e.g. lamproites, kimberlites) show radiogenic isotope ratios akin to those of crustal-derived silicic melts (Nelson et al. 1986). As shown in Figure 8, Tertiary to Quaternary orogenic volcanic rock suites from the Alpine-Mediterranean region show similar curvilinear trends in the 87Sr/86Sr v. 143Nd/a~Nd diagram. An exception is the calc-alkaline volcanic rocks from the Betics, which have large scatter in the isotopic ratios. Most of the volcanic series define a continuous trend, suggesting a common origin in terms of two-component mixing between a mantle component and an enriched component. Assimilation of crustal material by mantle-derived magma has been suggested for the Periadriatic magmas (Dupuy et al. 1982; Juteau et al. 1986; Kagami et al. 1991; von Blanckenburg et al. 1998) and for the East Carpathians calc-alkaline magmatism (Mason et al. 1996). However, such trends could imply also derivation of magmas from a strongly heterogeneous upper mantle without significant upper crustal assimilation, as has been proposed for the Central Italian magmas (e.g. Peccerillo 1985, 1999) and Sardinia (Downes et al. 2001). The high 87Sr/86Sr isotope ratios can be explained only by involvement of upper crustal material in the genesis of these magmas, and the most plausible explanation is that upper crustal continental material was subducted into the upper mantle and injected into their mantle source (Peccerillo 1985, 1999). To detect the processes of crustal involvement in magmagenesis, 87 Sr/ 86 Sr isotope ratios are often combined with oxygen isotope data (180/160 or 6180, w h e r e 180/160 is expressed relative to Standard Mean Ocean Water (SMOW); Fig. 9). The upper mantle is inferred to have relatively homogeneous 8180 values (+5.5 + 0.8; Mattey et al. 1994), whereas continental and oceanic crust generally display higher ~180 values as a result of weathering processes and interaction with marine or meteoritic water (Taylor 1968; Cerling et al. 1985). Interaction with crustal material could occur in two end-member processes (Fig. 9). At convergent plate margins, crustal material could be added to the upper mantle either via subduction of continental or pelagic sediments with the oceanic lithosphere or by the entry of crustal lithosphere into the mantle during mature subduction. Dehydration and melting of the subducted crustal material result in metasomatism of the upper mantle, the source region of 'subduction-related' magmas. This process is termed 'source contamination' (James 1981" Tera et al. 1986; Wilson 1989; Ellam & Harmon 1990) and is characterized by elevated 87Sr/86Sr, but relatively low ~180 values in the resulting magmas. Involvement of a crustal component could also occur within the continental crust when the ascending mantle-derived magmas assimilate fusible continental material (Taylor 1980; DePaolo 1981" Hildreth & Moorbath 1988; Davidson & Harmon 1989). In magmas formed by this process ('crustal contamination'), variable 87Sr/86Sr ratios are accompanied by high ~180 values (Fig. 9). The mechanism by which the continental crust is involved in orogenic magmatism can be constrained by combining 87Sr/S6Sr isotope ratios with the ~180 values either of phenocrysts from the volcanic rocks or of bulk rocks. The ~180 values of bulk rocks are usually higher than those of phenocrysts, because post-eruptive alteration and low-temperature weathering can increase the 180 contents (Taylor 1968; Davidson & Harmon
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM
179
Fig. 9. 87Sr/S6Sr v. 8180 diagrams for the Tertiary-Quatemary volcanic rocks of the Carpathian-Pannonian region and the Central Mediterranean. Variation of these data indicates different types of contamination ('source contamination' and 'crustal contamination'). Data sources are as for Figure 6. Additional data are from Taylor et al. (1979) and Holm & Munksgaard (1982).
1989; Ellam & Harmon 1990; Dobosi et al. 1998; Downes et al. 1995, 2001). Therefore, 8180 values of phenocrysts reflect better the isotope composition of the host magma. Unfortunately, only sporadic oxygen isotope data are available for mineral separates from the volcanic rocks of the region. In t h e 87Sr/86Sr v. 8 1 8 0 diagram (Fig. 9), the orogenic volcanic rocks of the AlpineMediterranean region show large variations. Calc-alkaline volcanic rocks from Sardinia (Downes et al. 2001) and most from the Pannonian Basin (Mason et al. 1996; Harangi et al. 2001; Seghedi et al. 2001) show only minor elevation of 8180 with increasing 87Sr/86Sr. T h e s e trends can be explained either by source contamination (Dowries et al. 2001; Seghedi et al. 2001) or by mixing of mantle-derived magmas with lower crustal metasedimentary material (Harangi et al. 2001). Mixing between lithospheric mantle-derived magmas and lower crustal melts has also been proposed by von Blanckenburg et al. (1998), for the genesis of the Alpine Periadriatic igneous rocks. For the remaining volcanic fields the higher 6180 values at a given SVSr/S6Sr could indicate upper crustal contamination (Mason et al. 1996). Contamination of the mantle source by subducted crustal material has been proposed also for Stromboli, Roccamonfina and for the potassic rocks of Vulsini (Taylor et al. 1979; Holm & Munksgaard 1982; Ellam & Harmon 1990). In contrast, contamination by upper crustal material combined with crystal fractionation in mantlederived magmas is envisaged for the calc-alkaline volcanic rocks of SE Spain (Benito et al. 1999), for most of the volcanic rocks of the Aeolian arc (Ellam & Harmon 1990) and the Aegean arc (Briqueu et al. 1986). In summary, continental crustal material has played an important role in the genesis of the 'orogenic' magmas of the Mediterranean region. Variation of 87St/ 86~ Sr and ~is O values sugge sts th at large amounts of crustal material of various types were recycled into the upper mantle during subduction and the following collision and post-collisional events. In the following, we attempt to characterize the pre-metasomatized mantle sources. One of the characteristic features of the Tertiary to Quaternary 'subduction-related' volcanic rocks of the Alpine-Mediterranean region is their close spatial and often temporal association with
alkaline sodic mafic volcanic rocks (Fig. 1). Coeval eruption of alkali basalts and calc-alkaline or shoshonitic magmas occurs at present in the Aeolian archipelago and Sicily. A similar process took place at the southeastern part of the Carpathian chain at c. 0.5-1.5 Ma (Mason et al. 1996, 1998; Seghedi et al. 2004). In the Central Mediterranean, the 'orogenic' volcanic rocks define a curvilinear trend in the 87 Sr/ 86Sr v. 206Pb/ 204Pb diagram (Fig. 10; Peccerillo 2003; Bell et al. 2004). One of the endmembers of this trend has high S7Sr/S6Sr and medium 2~176 values and is related to continental crust. The other end-member has low 87Sr/S6Sr and high 2~176 r a t i o s and could be an enriched mantle component that evolved with high U/Pb ratio over a long period of time. This mantle component shows similarities to the HIMU mantle end-member or to FOZO and is characteristic of OIB magmas. The alkaline sodic mafic rocks from Sicily and the Sicily Channel also show isotopic variation trending towards this mantle component, having a mixing trend between DMM and FOZO or HIMU. Mixing of OIB-like intraplate and subducted slab-derived components was also suggested by Peccerillo (2001) for the genesis of potassic rocks from Campania and Stromboli. Pleistocene basalts from Sardinia deviate from all of these volcanic rocks. They show a transitional geochemical character between the 'anorogenic' alkaline sodic mafic rocks and the 'orogenic' volcanic suites. However, the most peculiar feature of these rocks is the very l o w 2~176 isotope values (Gasperini et al. 2000; Lustrino et al. 2000), which are similar to the EMI-type OIB (Fig. 10). Gasperini et al. (2000) interpreted their origin as derivation from recycled oceanic plateaux material. In contrast, Lustrino (2000) and Lustrino & Dallai (2003) argued that the EMI-type nature of the Sardinian basalts can be explained by post-collisional delamination of the lithospheric mantle and the lower crust during the Hercynian orogeny, melting of the lower crust and the contamination of the lithospheric mantle by this silicic melt. In the Carpathian-Pannonian region, the Sr-Pb isotope plot suggests a more complex petrogenetic scenario (Fig. 10; Harangi 2001a). Calc-alkaline volcanic rocks from the western Carpathians (northem Pannonian Basin) fall in a curvilinear trend between a strongly
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s. HARANGIETAL.
eastern Carpathians trend toward a mantle component with lower 2~176 ratios, which is more characteristic of depleted MORB-type mantle (DMM). This mantle source was modified by addition of subducted flysch sediments and the primary magmas underwent high-level crustal Contamination (Mason et al. 1996). The youngest South Harghita shoshonites deviate from this trend, having significantly lower 2~176 ratios, possibly implying involvement of an EMI mantle component in their genesis. Alkali basalts of the Pannonian Basin (Embey-Isztin et al. 1993) form a continuous trend between the DMM and F O Z O - H I M U mantle end-members. Thus, in the CarpathianPannonian region, a multi-component mixing model can be envisaged (Salters et al. 1988; Rosenbaum et al. 1997; Harangi 2001a), where different mantle sources (DMM and F O Z O - H I M U and possibly also EMI) and lower and upper crustal components could have been involved in the genesis of the volcanic rocks. Petrogenesis of the Tertiary to Quaternary magmas in the Mediterranean region is as complex and controversial as the geodynamic evolution of the area. The mantle source regions are extremely heterogeneous, comprising all the identified mantle end-members found in oceanic basalts. This can be observed both in the undifferentiated alkaline mafic rocks and in the compositional variation of the 'orogenic' volcanic suites. In some places (e.g. Aeolian Islands-Sicily and southernmost eastern Carpathians) these alkaline sodic and 'orogenic' magmas erupted contemporaneously and spatially close to one another, implying that heterogeneity in the upper mantle could exist both horizontally and vertically on at least a 10 km scale. Crustal rocks were subducted into the upper mantle and the fluids and melts released from them thoroughly metasomatized the subcontinental mantle, the source of the 'orogenic' magmas. In addition, crustal material was also incorporated into the ascending magmas at higher crustal levels.
Geodynamic
Fig. 10. 87Sr/86Srv. 2~176 diagrams for the Tertiary-Quaternary volcanic rocks of the Carpathian-Pannonian region and the Western and Central Mediterranean. Data sources are as for Figure 6. Additionaldata are from Vollmer (1976).
radiogenic Sr isotopic component (lower crust?) and an enriched mantle component with low 87Sr//86Sr and high 2~176 ratios similar to that inferred for the Central Mediterranean magmas. In contrast, calc-alkaline volcanic rocks from the
implications
The low-K to high-K calc-alkaline volcanic rocks, the shoshonites and ultrapotassic formations, as well as alkaline volcanic rocks of the Alpine-Mediterranean region were formed in a convergent plate margin setting. The geochemistry of these magmas indicates a strongly heterogeneous mantle beneath this area. Most workers suggest that this can be explained by a lengthy period of subduction and subsequent post-collisional processes. Subduction of remnant Tethyan oceanic plates appears to have played an important role in the evolution of this region and has also had a major influence on the regional upper mantle structure. Seismic tomographic models show the presence of high-velocity anomalies beneath the Gibraltar arc, Calabria and the Hellenic arc, interpreted as recently subducted slabs (Spakman et al. 1988; Wortel & Spakman 1992, 2000; Blanco & Spakman 1993; Faccena et al. 2003; Piromallo & Morelli 2003). In addition, these models revealed an extensive coherent mass between 450 and 650 km depth that is interpreted as remnants of subducted Mesozoic oceanic slabs (Spakman et al. 1993; Piromallo et al. 2001; Piromallo & Morelli 2003). Indeed, most of the Tertiary to Quaternary volcanic rocks in the Mediterranean region show 'subduction-related' geochemical features. Some of them are found along volcanic arcs (Aeolian arc, Aegean arc) associated with the active subduction zones (Keller 1982). However, as shown in previous sections, there are debates on whether they could be considered as classic volcanic arcs. Some features would seem to indicate a back-arc tectonic setting (Mitropoulos et al. 1987; Mazzuoli et al. 1995). In this case, subduction could have only an indirect influence on the magmagenesis (Ellam et al. 1989). The principal reason for magma generation could be passive extension of continental lithosphere resulting in decompression melting of the lithospheric and asthenospheric mantle variably metasomatized by previous subduction processes.
ALPINE-MEDITERRANEAN SUBDUCTION & MAGMATISM Another candidate for volcanism directly related to active subduction is the Late Miocene low-K tholeiitic to calc-alkaline volcanic products of the Alboran basin (Duggen et al. 2004; Gill et al. 2004). However, other workers (e.g. Zeck 1996; Benito et al. 1999; Turner et al. 1999) have emphasized the post-collisional origin of these rocks. Indeed, formation of most of the 'orogenic' volcanic rocks in the Mediterranean regions post-dates the active subduction process and appear to be related to slab break-off, lithospheric mantle delamination or lithospheric extension (e.g. Mason et al. 1998; Seyito~lu et al. 1999; Turner et al. 1999; Aldanmaz et al. 2000; Chalot-Prat & G~rbacea 2000; Wortel & Spakman 2000; Harangi 2001a; Coulon et al. 2002; Seghedi et al. 2004). The 'subduction-related' geochemical character of the volcanic rocks is inherited from the mantle source regions modified previously (a few million to several tens or even hundreds of million years before) by fluids released from subducted slabs. Post-collisional or back-arc extension of the lithosphere could result in the decompression melting of the hydrous portion of the lithospheric mantle first (Gallagher & Hawkesworth 1992), followed by the melting of the deeper asthenosphere. This could explain the initial 'orogenic' magmatism and the subsequent alkali basaltic volcanism in many areas of the Mediterranean region (e.g. Wilson et al. 1997; Seyito~lu et al. 1999; Harangi 2001 a). Alkaline mafic rocks akin to those occurring in Central Europe occur sporadically in this region, often very close to the 'orogenic' volcanic formations. Furthermore, rare volcanic rocks types such as carbonatites, melilitites and/or kamafugites in the Apennines are more characteristic of intra-plate rift settings (Lavecchia & Stoppa 1996; Lavecchia et al. 2003). This may imply also another mechanism for magmatism of the Mediterranean region; that is, upwelling of hot mantle plume. The role of a mantle plume in the genesis of the volcanic rocks of Central Italy was first suggested by Vollmer (1976, 1989). Recognition of an enriched component (FOZO-HIMU) in both the 'anorogenic' and 'orogenic' volcanic rocks in many subprovinces (Hoernle et al. 1995; Ayuso et al. 1998; Wilson & Bianchini 1999; Harangi 2001a; Gasperini et al. 2002; Peccerillo 2003; Bell et al. 2004) also led some researchers to propose mantle plume activity. This could be supported by the extensive low-velocity anomaly beneath most of this area (Hoernle et al. 1995; Wortel & Spakman 2000; Piromallo & Morelli 2003) that could be interpreted as presence of anomalously hot mantle material. The estimated temperature from the P and S velocity anomalies approaches the dry solidus under the Pannonian Basin, Western Mediterranean, Tyrrhenian Basin and Aegean Basin, and the discrepancy between temperature inferred from P and S waves also indicates the presence of partial melt (Goes et al. 2000). Harangi (2001b) supposed also a relatively hot mantle beneath the Pannonian Basin, based on the composition of the Late MiocenePliocene alkali basalts. As an extreme case, Lavecchia et al. (2003) and Bell et al. (2004) argued that the geodynamic evolution of the Central Mediterranean has been controlled by an upwelling plume and subduction had no role whatsoever. Other workers combined the subduction-related models with the existence of OIB-like mantle (Fig. 11). Gasperini et al. (2002) assumed that the HIMU signature of many of the volcanic rocks in the Central Mediterranean could be related typically to an upwelling plume. They invoked a plate window beneath the central Apennines, where deep mantle plume material could be channelled toward the shallow mantle zones (Fig. l la). Decompression melting of this hot plume material could result in the enriched mantle component of many of the volcanic rocks in Central Italy. In the southern Tyrrhenian area, the coexistence of OIB and 'orogenic' magmas was interpreted as the result of lateral flow of African enriched mantle along a tear at the edge of the Ionian plate (Fig. l lb; Trua et al. 2003). Mantle anisotropy studies in this area also indicate a toroidal mantle flow around the Calabrian slab (CiveUo & Margheriti 2004), which could
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Fig. 11. Proposed models for the involvementof an enriched mantle component (EAR) of the 'orogenic magmas' in the Carpathian-Pannonian and Mediterraneanregion. (a) Deep mantle upwelling could occur via a slab window beneath Central Italy as proposed by Gasperini et al. (2002). (b) Toroidal mantle flow around the Calabrianslab from the African mantle was proposed by Trua et al. (2003) and Civello & Margheriti (2004). (e) Carpathian-Pannonian region: slab detachment could result in the suction of a hot, enriched asthenosphericmantle material possibly from the Bohemian mantle plume finger.
supply enriched mantle material beneath Campania and the southern Tyrrhenian area from behind the Calabrian subduction zone. In the Carpathian-Pannonian region the compositional variation of Miocene to Quaternary calc-alkaline volcanic rocks implies contrasting genesis. Isotopic values of andesitic to dacitic rocks in the western segment of the Carpathian arc show a mixing trend between an enriched (FOZO-HIMU-type) mantle and a crustal component (Fig. 10; Harangi 2001a) similar to other volcanic suites in the Mediterranean region, whereas the
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s. HARANGI ETAL.
calc-alkaline volcanic rocks in the eastern segment of the Carpathian arc show a mixing trend between a depleted mantle and a crustal component. A possible explanation for the contrasting mantle source regions is that enriched mantle could flow from the assumed plume finger beneath the Bohemian Massif to the thinned Pannonian Basin through the gap left behind the detached slab under the western Carpathians (Fig. l lc). In the east, no enriched mantle material could penetrate beneath the thick Ukrainian Shield, therefore slab break-off beneath the eastern Carpathians could initiate upwelling of depleted MORB-type mantle material. Deflection of the assumed mantle plume finger beneath the Massif Central towards the SE was also detected by seismic anisotropy pattern (Barroul & Granet 2002). The southeastward asthenospheric flow was explained by a combined effect of Apenninic slab roll-back and the opening of the extensional basins behind it (Barroul & Granet 2002; Barroul et al. 2004). In summary, the F O Z O - H I M U mantle component recognized in the compositional variation of many Mediterranean volcanic suites led many researchers to propose the influence of localized mantle plume(s) in the genesis of the magmas. Whether upwelling of hot mantle material was the ultimate cause of the magmagenesis and also influenced the tectonic evolution of the areas (Lavecchia et al. 2003) or subduction and post-collision processes (slab rollback, slab break-off, delamination of the lower lithosphere) initiated deflection of nearby mantle plumes, requires further combined geochemical, geophysical and tectonic research. Nevertheless, the HIMU signature of the mantle source could also be interpreted as due to metasomatic processes without assuming mantle plumes (Sun & McDonough 1989; Anderson 1994; Halliday et al. 1995; Niu et al. 1999). It is remarkable that the enriched, HIMU-like mantle component was detected also in earlier, pre-Neogene volcanic rocks of this region (Veneto region, Macera et al. 2003; Dinarides, Cvetkovid et al. 2004; C a r p a t h i a n - P a n n o n i a n region, Harangi 1994; Harangi et al. 2003), which may imply its long-lasting (at least from the Early Cretaceous) presence beneath Europe. Oyarzun et al. (1997) and Wilson (1997) suggested that this enriched mantle component could be derived from the Mesozoic Central Atlantic plume being deflected as a result of either the suction of the European thin-spots or the northeastward motion of the African plate. In any case, portions of the deflected plume material could have polluted the shallow upper mantle beneath Europe since the Early Cretaceous. In addition, subduction of crustal material could also contribute to the inhomogeneity of the shallow mantle. Statistical sampling of this heterogeneous mantle (SUMA model, Meibom & Anderson 2004) could be an alternative model for the wide variation of the Tertiary to Quaternary volcanic rocks of the Mediterranean region. Discussions with the PANCARDI Igneous Team over the last decade have helped us to develop our ideas on the Tertiary to Quaternary magmatism and its geodynamic relationships in the Alpine, Pannonian, Carpathian and Dinarides regions. Part of this study belongs to the research project supported by the Hungarian Science Foundation (OTKA # T 037974; to Sz.H.). The critical reviews of C. Doglioni and F. Chalot-Prat helped to improve the paper and are gratefully acknowledged. We also thank R. Stephenson and D. Gee. for patient editorial help, and R. Lukfics for the careful work in producing the reference list.
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ZECK, H. P., ALBAT, F., HANSEN, B. T., TORRES-ROLDAN, R. L. & GARCIA-CASCO,A. 1989. Alpine tourmaline-bearing muscovite leucogranites, intrusion age and petrogenesis, Betic Cordilleras, SE Spain. Neues Jahrbuch fiir Mineralogie, 11, 513-520. ZECK, H. P., KRISTENSEN,A. B. & WILLIAMS, I. S. 1998. Post-collisional volcanism in a sinking slab setting--crustal anatectic origin of pyroxene-andesite magma, Caldear Volcanic Group, Neogene Alboran volcanic province, southeastern Spain. Lithos, 45, 499-522. ZINDLER, A. & HART, S. R. 1986. Chemical geodynamics. Annual Review of Earth and Planetary Sciences, 14, 493-571.
Formation and deformation of the Pannonian Basin: constraints from observational data F R A N K HORV./~TH 1, GJi.BOR B A D A ~'2, PI~TER SZAFIAN ~, G A B O R TARI 3, A N T A L / ~ D A M 4 & SIERD C L O E T I N G H 2
XDepartment of Geophysics, E6tvOs University, Pdzmdny P. s. I /C, 1117 Budapest, Hungary (e-mail:
[email protected]) 2Netherlands Research Centre for Integrated Solid Earth Sciences, Department of Tectonics and Structural Geology, Vrije Universiteit, 1081 HV Amsterdam, The Netherlands 3Vanco Energy Company, Houston, TX 77046, USA 4Geodetic and Geophysical Research Institute, Hungarian Academy of Sciences, Sopron, Hungary
The past decade has witnessed spectacular progress in the collection of observational data and their interpretation in the Pannonian Basin and the surrounding Alpine, Carpathian and Dinaric mountain belts. A major driving force behind this progress was the PANCARDIproject of the EUROPROBE programme. The paper reviews tectonic processes, structural styles, stratigraphic records and geochemical data for volcanic rocks. Structural and seismic sections of different scales, seismic tomography and magnetotelluric, gravity and geothermal data are also used to determine the deformational styles, and to compile new crustal and lithospheric thickness maps of the Pannonian Basin and the surrounding fold-and-thrust belts. The Pannonian Basin is superimposed on former Alpine terranes. Its formation is a result of extensional collapse of the overthickened Alpine orogenic wedge during orogen-parallel extrusion towards a 'free boundary' offered by the roll-back of the subducting Carpathian slab. As a conclusion, continental collision and back-arc basin evolution is discussed as a single, complex dynamic process, with minimization of the potential and deformational energy as the driving principle. Abstract:
The Pannonian Basin is located in eastern Central Europe, as part of the Alpine orogenic system. The Alpine, Carpathian and Dinaric mountain belts surround this extensional basin of NeogeneQuaternary age. Its broader geological environs, the Mediterranean region, is a wide zone of convergence between the Eurasian and African plates. Because of the complex kinematics of the two major plates, the region has had a polyphase deformation history since the opening of the Atlantic Ocean, which has primarily controlled the formation of the entire Alpine orogenic belt (Biju-Duval et al. 1977; Dercourt et al. 1986; Dewey et al. 1989; Seng6r 1993; Yilmaz et al. 1996). A remarkable feature in this overall compressional setting is the abundance of extensional basins superimposed on former orogenic terranes and associated with orogen-parallel displacement of internal blocks and oroclinal bending (Horvfith & Berckhemer 1982). From west to east, these basins are the Alboran, Ligurian, Tyrrhenian, Pannonian and Aegean basins (Fig. la). Although their age, deep structure and tectonics show significant differences, there are several common features in their formation and evolution. For instance, their similar position in proximity to a once or still active subduction zone strongly suggests a causal relationship of primary importance (Royden 1993; Giunchi et al. 1996; Meijer & Wortel 1997). Other important geodynamic processes include the extensional collapse of a gravitationally unstable orogenic wedge (Dewey 1988), and the lateral escape of internal terranes (Seng6r et al. 1985; Royden 1993). Following the suggestion of Horvfith (1988), Ratschbacher et al. (1991) argued that extensional collapse occurs during tectonic escape and suggested the term extrusion to describe these interrelated processes. It is to be emphasized that this was not simply a new term but a novel concept, which postulated that orogenparallel extension and orogen-normal compression in the Mediterranean were taking place simultaneously, as in the Asian segment of the Alpine-Himalayan mountain belt (Tapponnier et al. 1986). This implies that extrusion is not a lateral displacement (translation) of one coherent block but, instead, strong internal deformations occur, which involve brittle faulting and ductile flow depending on the rheological layering of the extruding terrane (Ratschbacher et al. 1991). Based on a summary of 3D seismic tomography data, Wortel & Spakman (2000) have offered new constraints for the Cenozoic evolution of the Mediterranean region and a contribution to the
explanation of back-arc basin formation and deformation. Tomography images clearly depict the elevated asthenospheric material below back-arc basins and the lithospheric slabs underthrusting these lithospheric depressions. The geometry of the individual subducted slabs is particularly important as it shows considerable differences in the Mediterranean region (Faccenna et al. 2003). Along the Hellenic arc a continuous slab penetrates down more than 1000 km, well into the lower mantle below the Aegean Basin, and only the upper 200 km portion of this long slab exhibits seismic activity. Elsewhere, the slab dips from the subduction zone towards the interior of the back-arc basin, and then soles out between the 410 km and 660 km seismic discontinuities. This geometry is compatible with the progressive rollback of a subducted slab and arc retreat. An additional important feature is the observation that in active subduction zones the slab appears to be continuous towards the surface, whereas in inactive subduction zones a detached relict of the subducted slab can be imaged in the upper mantle. Slab detachment occurs in the form of a self-perpetuating break-off under the weight of the slab itself (Dvorkin et al. 1993; Davies & von Blanckenburg 1995). This can happen when thick and buoyant continental lithosphere enters the trench zone. The process of subduction is blocked when the buoyancy of the arriving continental lithosphere becomes equal in magnitude to the downward force exerted by the already subducted oceanic lithosphere. Lateral migration of slab detachment leads to the concentration of tensional forces on a continuously decreasing portion of the arc, which can result in a gradual, often accelerating retreat of the arc-trench system (Dvorkin et al. 1993; Royden 1993). This process is responsible for the migration and bending of the arcs and back-arc basin formation from Oligocene to recent times in the Mediterranean and the build-up of the arcuate orogenic chains of the Betics-Rif, Calabrides-Apennines and Hellenides (Faccenna et al. 2004). Formation of the Carpathian arc and the Pannonian Basin took place in a similar setting from the latest Oligocene. This basin has rapidly reached a mature stage of evolution, as extension has come to an end because of the complete consumption of the subductable lithosphere of the European foreland (Horv~th 1993). As a result, unlike other Mediterranean back-arc basins, the change of stress field from extension to compression and positive structural inversion of the Pannonian Basin system has been in progress since Late
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 191-206. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. (a) Sketch map showing the present-day configuration and the late Cenozoic structural features of the Mediterranean and adjacent areas. The overriding plates above subduction zones have undergone rigid body rotation, translation and extension, which has resulted in the formation of a set of back-arc basins. PB, Pannonian Basin; AdP, Adriatic Promontory; AeS, Aegean Sea; A1S, Alboran Sea; B1S, Black Sea; loS, Ionian Sea; LiS, Ligurian Sea; TyS, Tyrrhenian Sea. Darker shading indicates oceanic crust. (b) Dynamic model for the evolution of the same area during the Cenozoic proposed by Wortel & Spakman (2000). Continuous line indicates area of active subduction; thick grey line indicates area where slab detachment has already occurred at the given time interval. E, Eocene; O1, Oligocene; Mt, Early Miocene; M2, Mid-Miocene; M3, Late Miocene; Pr, present. Pliocene times (Horv~ith & Cloetingh 1996; Bada et al. 1998, 2001). The goal of this paper is to review and interpret recent data regarding the formation and deformation of the Pannonian Basin system. This can assist the understanding of other Mediterranean back-arc basins and lithospheric dynamics in general. We attempt to show that although the level of understanding of the formation and deformation of the Pannonian Basin was very high already in the mid-1980s (e.g. Royden & Horv~th 1988), the past decade has witnessed spectacular progress in the collection of observational data and their interpretation. The EUROPROBE programme, particularly the PANCARDI project, has clearly been a most important driving force of this progress.
Tectonic framework The Cenozoic evolution of the Alpine-Pannonian region is primarily controlled by the northward drift and indentation of the Adriatic promontory, which has produced a net convergence of at least 5 0 0 k m in the Alps (Roeder & Bachmann 1996;
Schmid et al. 1996). Adria has been pushed towards the north by the African plate, although it was not necessarily tightly attached to Africa, at least from the Early Tertiary (M~irton et al. 2003). Another important feature of Adria is that it represents a non-rigid indenter, as is shown by its dramatic deformation history throughout the Cenozoic (Fig. lb). It can be reconstructed fairly well (D'Argenio & Horv~ith 1984) that an originally broad bump at the northern boundary of the African plate gradually became a sharp nose by necking as a result of the development of the Tyrrhenian and Aegean back-arc basins (Fig. lb). The northern front of Adria also experienced strong deformations but of different nature. The formation of the West Alpine arc during the Palaeogene (Giglia et aL 1996; Schmid & Kissling 2000) can be best explained by westward extrusion of a fragment of the Adriatic indenter. On the basis of differential global positioning system (GPS) velocities and seismicity, Oldow et al. (2002) have concluded that contemporary Adria is fragmented. A northwestern block has velocities indicative of little or no motion relative to Europe, whereas a southeastern part is moving together with Africa and exhibits a spatially heterogeneous velocity field with northward displacement rates up to 1 0 m m a - t . In the Eastern Alps, the
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Austroalpine basement and cover nappes overriding the deformed European margin are obviously of Adriatic provenance. Clearly, there is not a single Adriatic-European boundary in the Alps either at upper crustal level or at a depth, as we show later in this paper, when discussing the new results of deep seismic profiling and improved mantle tomography. The most pronounced surface expression of strain partitioning (Ziegler & Roure 1996) has been the Late Oligocene to Early Miocene eastward extrusion of an Alpine orogenic wedge, primarily driven by the northward push of the 'soft' Adriatic indenter. This Alpine wedge is called the ALCAPA terrane (Fig. 2). Its boundaries, again, can only be loosely defined, which reflects the nature of the extrusion process rather than the limits of our knowledge (Frisch et al. 1998). There is a second unit in the substrata of the Pannonian Basin system, called the Tisza-Dacia terrane. Detailed description of the contrasting Mesozoic to Early Tertiary stratigraphy of these two terranes can be found in a set of recent papers (e.g. Kovfics et al. 2000; Csontos & V6rrs 2004). It is generally accepted that the Tisza-Dacia terrane rifted apart from the European margin of the Mesozoic Tethys during the Late Jurassic, and it led to the formation of the marine basins in which the AlpineCarpathian flysch complexes were deposited (Yilmaz et al. 1996; Csontos & V r r r s 2004). Although a number of papers have attempted to reconstruct the Tertiary kinematic history of the Tisza-Dacia terrane (Balla 1986; Csontos et al. 1992; Csontos 1995; Csontos & V r r r s 2004) this issue has remained obscure. This is mainly because of an acute space problem between the Bohemian massif and the assumed fixed Moesian platform (Fig. 2), and the lack of adequate kinematic indicators. Palaeomagnetic data tend to indicate counter-clockwise rotations for the ALCAPA domain and clockwise rotations for the T i s z a -
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Dacia unit (Fig. 2), but with the increase in quantity and quality of palaeo-declinations serious rotation anomalies have been revealed within the two terranes (see a review by Mfirton 2001). This is reasonable, as both units experienced strong internal deformations during the Tertiary. It is generally accepted that the two terranes became juxtaposed at the beginning of the Miocene when formation of the Pannonian Basin started (Csontos et al. 1992). Geodynamically, this formation was a stretching of the two terranes towards the eastern Carpathians facilitated by the subduction roll-back of the lithosphere of the Carpathian flysch basin (Horvfith 1993; Bada & Horvfith 2001).
Tectonic evolution The Pannonian Basin and its surroundings are characterized by a polyphase deformation history with a sequence of distinct structural episodes. There is a good knowledge of the principal kinematic features, that is the location of major fault zones, the timing and the amount of deformation (Figs 3 and 4). A rapid and dramatic change in tectonic style started in the Early Miocene (Eggenburgian to Karpatian; see Fig. 4 for local time scale), which initiated the formation of the Pannonian Basin. This process culminated in the Mid-Miocene (Badenian) and was coeval with a large-scale tectonic transport of the external flysch nappes towards the foreland of the Carpathian arc (Royden et al. 1982). It is widely recognized that in regions of continental collision intense shortening and crustal thickening can eventually lead to gravitational instability of the axial zone of the orogen
Fig. 2. Simplified Late Cenozoic tectonic map of the Alpine-Carpathian-Pannonian-Dinaric system (after Bada & Horv~ith 2001). The Adriatic promontory or microplate has been indenting and pushing the Alpine-Dinaric belt since the Cretaceous. The northern domain beneath the Pannonian Basin (ALCAPA unit) underwent a net counter-clockwise rotation of about 50-70 ~ and was translated to the ENE, whereas the southern unit (Tisza-Dacia unit) rotated c. 100~ in an opposite manner and was moving to the ESE. Green and black arrows indicate the translation and rotation, respectively, of various tectonic units. 1, Foreland (molasse) basins; 2, flysch nappes; 3, Neogene volcanic rocks; 4, Southern Alps, Dinarides and Northern Calcareous Alps; 5, pre-Tertiary units of the East Alpine-Carpathian domain and the Jura Mts; 6, Variscan basement of the European plate, and Dinaric, Vardar and Mures ophiolites; 7, Pieniny Klippen Belt; 8, Oligocene tonalites; 9, Penninic basement; 10, Penninic cover; 11, Helvetic basement; 12, Helvetic cover. T, Tauern window; R, Rechnitz window; PAL, Periadriatic lineament (PAL); G, Giudicarie fault; Za, Zagrab fault; B, Brenner fault; Tr, Trotu~ fault; IM, Intramoesian fault.
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Fig. 3. Depth of pre-Neogene basement in the Pannonian Basin system (Horvfith & Royden 1981) and related late Neogene structural pattern (Horvfith 1993). Thick red lines mark the location of structural profiles shown in Figure 6 (A-A'), Figure 7a (B-B'), Figure 7b (C-C') and Figure 8 (D-D'). B6, B6k6s Basin; D, Danube Basin; HM, H6d-Mak6 Basin; M, Mihfilyi high; PB, Pusztaf61dvfir-Battonyablock; TB, Transylvanian Basin; TR, Transdanubian Range; Z, Zala trough. 1, Foreland (molasse) basins; 2, flysch nappes; 3, Neogene volcanic rocks; 4, pre-Tertiary units on the surface; 5, Variscan basement of the European plate; 6, Dinaric and Vardar ophiolites; 7, tectonic windows in the Eastern Alps; 8, normal and low-angle normal fault; 9, thrust, anticline; 10, strike-slip fault.
(Tapponnier et al. 1986; Molnar & Lyon-Caen 1988; Bird 1991). Moreover, the central belts of orogens are often thermally weakened and, hence, are prone to strain localization. Once the gravitational forces exceed the compression exerted by plate convergence, the process of late orogenic collapse of the weakened crust can be initiated. Orogen-parallel extension in the Eastern Alps is well documented (e.g. Selverstone 1988; Ratschbacher et al. 1989) and the idea of a weakened Alpine lithosphere has been confirmed (e.g. Cloetingh & Banda 1992; Genser et al. 1996; Okaya et al. 1996). The large-scale lateral extrusion of the ALCAPA terrane took place from the Late Oligocene towards an eastern, unconstrained margin of the Carpathian flysch basin (Fig. 5). In a strict sense, lateral extrusion is a ductile flow of the lower crust confined between the brittle upper crust and mantle lithosphere (Ranalli 1995) that leads to the relaxation of the topography of the surface as well as the Moho (Bird 1991). A large amount of material was expelled towards the east from between the soft indenter (Adria) and the rigid foreland buttress (Bohemian Massif). Crustal wedges bounded by conjugate sets of strike-slip faults, that is sinistral and dextral in the north and south, respectively, were extruded and stretched in an orogen-parallel direction to the ENE (Figs 2, 3 and 5). Often, these strike-slips are still
active seismically, and magnetotelluric soundings show that they are associated with highly conducting zones containing lowviscosity material (Ad~im 2001). The extrusion process was coeval with continuing north-south compression in the central zone of the Eastern Alps, and rapid exhumation of metamorphic rocks in the Alps (Dunkl & Frisch 2002), the Pannonian Basin (Tari & B a l l y 1990; Tari et al. 1992, 1999), and the East Slovak Basin (Sotak et al. 1993). Kinematic data (Fodor et al. 1999) and numerical modelling (Bada 1999) suggest the predominant role of Carpathian subduction facilitating large-scale lithospheric extension in the Pannonian Basin from latest Early Miocene to Pliocene times. Continuous roll-back of the subducting plate along the contemporaneous Carpathian arc exerted trench-pull forces on the upper plate. The overriding plate in a subduction zone tends to passively follow the retreating hinge of the downgoing lithosphere. This induced tensional stresses and eastward stretching of the A L C A P A and Tisza-Dacia terranes. The upper plate is extending in the direction of maximum gravitational potential energy difference. Trench suction forces acting normal to the curvature of the Carpathian arc, in combination with collisional forces exerted along the Alpine-Dinaric belt, can reproduce well the reconstructed Mio-Pliocene palaeostress pattern (Bada 1999). Because
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Fig. 4. Evolutionaryscheme of the PannonianBasin system with major volcanic horizons and events, generalized stratigraphy, and main tectonic phases. The primary depositional environmentsand facies of the PannonianBasin are described in the bottom panel (after Horv~ith& Tari 1999). Major unconformitiesare labelled U1 (oldest) to U4 (youngest).
of the finite strength of the Pannonian lithosphere, tensional stresses were transmitted far behind the arc region and, as a consequence, nearly the whole Pannonian Basin system extended significantly. It is important to emphasize that there are no remarkable differences in either the style or the amount of extension between the formerly distinct ALCAPA and Tisza-Dacia terranes (compare Figs 6 and 7). There is, however, one obvious exception: the presence of the coeval non-extensional Transylvanian Basin in the eastern part of the Tisza-Dacia unit (Ciulavu et al. 2002). The formation of this basin is poorly understood and the original concept of Royden et al. (1982) still seems to be the most plausible. They explained the Transylvanian Basin as a continental sag caused by the suction force exerted to the upper plate by the downbending Carpathian slab. Tension in the Pannonian region caused about 50-120% crustal lithosphere extension and nearly an order of magnitude higher mantle lithosphere extension (Horvfith et al. 1988; Lenkey 1999). Occasionally, extension was concentrated in discrete zones where pull-apart basins developed (Horvfith & Royden 1981; Horvfith 1993; Csontos 1995; Fodor et aL 1999). Heterogeneous extension is reflected by the variation of pre-Neogene basement depth (Fig. 3). Elevated basement blocks separate deep sub-basins
where the thickness of the Neogene-Quaternary sedimentary rocks can reach 6 - 7 km. Such irregular basement morphology is mainly the result of strain localization along pre-existing crustal weakness zones inherited from Late Cretaceous thrust and nappe tectonics. Quaternary differential vertical movements, related erosion and sedimentation have significantly influenced the observed thickness of the basin fill. These processes are related to the late stage of basin evolution. Contemporary stress data, seismicity pattern, seismic profiles and Quaternary subsidence history indicate that the Pannonian Basin is in the phase of structural inversion (Horvfith 1995; Horvfith & Cloetingh 1996; Bada et al. 1999; Gerner et al. 1999). An increase of horizontal compression as a result of the changes of boundary conditions around the basin system causes buckling of the Pannonian lithosphere (Horvfith & Cloetingh 1996; Cloetingh et al. 2006) manifested in the uplift and subsidence of the basin flanks and centre, respectively. Present-day boundary conditions include active collision along the Alps-Dinarides belt (Adria-push), terminated subduction and continuing continental collision in the SE Carpathians and eastward translation of crustal wedges currently squeezed out from the region of the Alpine orogen (Bada et al. 1998, 2001).
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Fig. 5. Schematicillustrationof the Early Miocene extrusion of the ALCAPAterrane. Continentalconvergencein the Alpine collisionalbelt was able to progress because the buoyant crust decoupled from the dense mantle lithosphere. This crustal flake, the ALCAPA terrane, extruded towards a 'free' eastern boundary. This was free in the sense that the lithosphereof the Carpathian flyschBasin was subductableand its progressiveroll-back offeredthe space to be invaded. It should be noted that the extruding crustal flake was directly superimposedon the hot asthenosphere, which led to crustal melting and a dramatic decrease of its integrated strength. 1, European foreland areas, undifferentiated('stable Europe'); 2, foreland (molasse) foredeep; 3, Alpine-Carpathian nappe system; 4, flyschBasin in the Carpathian embayment;5, normal fault, strike-slipfault and thrust plane, thrust front; 6, orogen-paralleldisplacementof the extruding (collapsing)Pannonianterranes. PA indicates the Periadriaticline. The basin system has become completely landlocked and constrained from all directions, which has led to a gradual inversion in the form of multi-scale folding and fault reactivation (Fig. 8).
Basin-fill stratigraphy and magmatism Stratigraphic data provide important information on the basin evolution in terms of timing and characterization of major tectonic events. Furthermore, the petrological and geochemical analyses of magmatic rocks give further constraints on the composition, thermal state and rheological behaviour of the deforming lithosphere-asthenosphere system. The evolutionary scheme of the basin system, with major volcanic events, stratigraphic pattern, tectonic phases and local chronostratigraphic units, is shown in Figure 4. The onset of rifting in the Pannonian Basin is marked by the regional unconformity between the pre-rift strata and the overlying Lower Miocene (Eggenburgian-Ottnangian) deposits. A rhyolite tuff horizon dated at c. 20 Ma (H~mor et al. 1980) is interbedded in the lowermost part of the synrift sedimentary beds. However, in the main depocentres this tuff horizon is missing, and the thick basal conglomerates and other continental beds cannot be precisely dated. Therefore 20 Ma as a start of rifting is a somewhat uncertain, but reasonably good age estimate. The evolution of the Pannonian Basin is traditionally subdivided into a synrift (Early to Mid-Miocene) and a post-rift (Late Miocene to Quaternary) phase (Royden et al. 1983; Horv~th & Rumpler 1984), which is reflected in the sedimentary architecture. An increase of available stratigraphic data and the re-evaluation of numerous seismic profiles has allowed a slight modification of this subdivision. According to Taft (1994), Horvzith (1995) and Tari et al. (1999), a Mid-Badenian unconformity indicates the termination of the synrift period (Fig. 4). The Sarmatian-Pannonian regional unconformity is due to uplift and erosion related to an early
inversion event (Horv~th 1995; Fodor et al. 1999), which may indicate transient changes in the boundary conditions along the Carpathian subduction belt. This compressional event took place shortly after the termination of the synrift phase, that is at about 11-8 Ma. There is a second, more regional compressional event, which started during the Late Pliocene and has continued until recent time (c. 3 - 0 Ma). Widespread upwarping of basement units from below the Neogene succession has resulted in the characteristic 'inselberg' pattern of present-day Mesozoic and Palaeozoic outcrops inside the Pannonian Basin (Fig. 3). In other places, where the basement was uplifted but did not reach the surface, erosion led to a stratigraphic gap and unconformity between Miocene and Quaternary strata (Fig. 4). At the same time, rapid subsidence has taken place in the deepest sub-basins during the Quaternary. Modelling results (Cloetingh et al. 2006) suggest an increase of compressional intraplate stress magnitudes and related lithospheric folding that can fairly well explain the observed pattern of Quaternary subsidence and uplift. Depositional environments in space and time were strongly influenced by tectonics (Fig. 4). The synrift phase was characterized by continental to marine sedimentation, whereas during the post-rift phase the Pannonian Basin became an isolated brackishwater lake, which has been progressively filled up, and lacustrine sedimentary rocks were replaced by terrestrial deposits (e.g. Nagymarosy & Mtiller 1988; Jfimbor 1989; Magyar et al. 1999; H~mor et al. 2001). The late-stage terrestrial depositional environments have been particularly sensitive to climatic fluctuations driven by Milankovid-type cyclicity (Nfidor et al. 2000). However, the effect of global eustatic sea-level changes on the stratigraphy of an isolated large lake is still a matter of debate (Tari et al. 1992; Pogfics~s et al. 1994; Vakarcs et al. 1994; Juh~sz et al. 1999; Sacchi et al. 1999). Volcanic rocks identified in the Pannonian-Carpathian system show large variation in lithology, geochemical composition, and spatial and temporal distribution (P~cskay et al. 1995; Harangi 2001; Kone6n3) et al. 2002). Silicic volcanism started about
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Fig. 6. Seismic reflection section A-A' through the Great Hungarian Plain in the centre of the Pannonian Basin and a line drawing interpretation (after Tari et al. 1999). (For location of profile see Fig. 3.)
20 Ma ago with the deposition of ignimbritic flow deposits and rhyolitic tufts. Repeated activity resulted in the formation of two other rhyolitic tuff horizons (Fig. 4). Based on geochemical and isotope data, these silicic volcanic rocks can be best explained as the product of partial melting of the continental lower crust in the Pannonian lithosphere (Salters et al. 1988; Kone~n2~ et al. 2002). Harangi (2001) have argued for a mantle origin of these magmas, but with strong crustal contamination. The considerable involvement of crustal materials in early synrift magmatic processes strongly suggests that an anomalously high temperature gradient was established, leading to the low strength of the deforming lithosphere from the earliest stages of basin evolution. The second main type of magmatic activity during Miocene to Quaternary times (20-0.15 Ma) produced large bodies of calc-alkaline volcanic rocks in the northern Pannonian Basin and at the inner side of the Eastern Carpathians (Fig. 2). In the first plate tectonic models this volcanism was directly or indirectly related to the subduction of the flysch basin lithosphere (Stegena et al. 1975; Lexa & Konetn2~ 1974). Results of detailed geochemical analysis (Salters et al. 1988; Szab6 et al. 1992; Downes et al. 1995; Mason e t al. 1996; Harangi 2001) confirmed the subductionrelated origin and highlighted important spatial and temporal variations in magma generation and evolution. Volatile-rich fluids from the subducted slab migrated into the mantle wedge above it, which created large volumes of magma of mostly andesitic composition. While rising to the surface, as suggested by Downes et al. (1995), the magma was contaminated by crustal
components, resulting in both spatial and vertical variations of the isotopic composition. The widespread occurrence of garnet xenoliths suggests very rapid ascent of the magmas in an overall tensional stress regime (Harangi 2001). Kone~n3~ et al. (2002) have also argued that in the northern parts of the Pannonian Basin system the generation of calc-alkaline magmas was related to the decompressional melting of the mantle lithosphere as a result of the overall extension in the Pannonian Basin. The youngest region of calc-alkaline magmatic activity is a remarkably linear chain along the inner side of the east Carpathians (Fig. 2). A general progression from 12 Ma to 0.2 Ma in the onset and cessation of magmatism from the NW towards the SE has been documented by K - A r geochronology (Ptcskay et al. 1995). Subduction of the flysch basin lithosphere and progressive break-off of the subducted slab has been suggested as the most probable cause of the east Carpathian magmatism and its migration (e.g. Linzer 1996; Mason et al. 1998). As thick continental crust began to enter the trench zone, detachment of the subducted oceanic slab occurred, and the rupture propagated along the slab. This led to the cessation of volcanism as the slab sank out of the magma generation zone. In recent times, this process is in its final stage at the bend of the eastern and southern Carpathians. Analyses of the magmatism, structural evolution, geometry of the seismically active Vrancea slab, and mantle tomography suggest that in this area detachment is not complete and slab continuity is probably still partly maintained (Girbacea & Frisch 1998; Chalot-Prat & Girbacea 2000; Wortel & Spakman 2000; Gvirtzman 2002).
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Fig. 7. (a) Gravity model of a crustal-scale transect B-B' (after Szafifinet al. 1999) across the ALCAPAterrane. (b) Hydrocarbon exploration seismic profile C-C' across the central part of the Danube Basin (after Tari 1994). (For location of profiles see Fig. 3.)
The third main type of magmatic activity in the PannonianCarpathian region took place during Late Miocene-Pleistocene times (12-0.5 Ma) with a climax at 3 - 5 Ma (Balogh et al. 1986; P6cskay et al. 1995), producing mainly alkali basalts and some other mafic rocks. Volcanic products are located throughout the Pannonian and Transylvanian basins and are of much lower volume than the calc-alkaline rocks. Trace elements and isotope ratios indicate a predominantly asthenospheric source for the related magmas (Downes e t al. 1995; Embey-Isztin & Dobosi 1995). Moreover, the volcanic products exhibit strong similarity to other Late Neogene basalts in Western Europe,
and their origin from a common mantle source, the so-called European Asthenospheric Reservoir, has been suggested (Cebri~ & Wilson 1995).
Style of deformation To illustrate the style of deformation related to basin evolution, we present four representative structural profiles (for location see Fig. 3). The first three sections ( A - A ' , B - B ' and C - C ' ) image the basin architecture and structural pattern of the two deepest
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Fig. 8. Interpreted seismic reflectionprofile D-D ~through the Zala Basin in the southwestern part of the PannonianBasin. (For location of profile see Fig. 3.) TWT, two-way travel time.
sub-basins, the Great Hungarian Plain and Danube Basin in the central and western part of the Pannonian Basin system, respectively (Figs 6 and 7). The fourth profile ( D - D f) is located at the boundary zone of the two terranes in SW Hungary, and is a good example of the anatomy of the late-stage basin inversion (Fig. 8). Deep reflection seismic profile A - A ' extends S W - N E across the deepest depressions in the Pannonian basin, the Hdd-Mak6 and B~k~s basins (Fig. 6). This area of the Tisza terrane has undergone considerable crustal and lithospheric extension, which is clearly reflected in the large thickness of the Neogene Basin infill, high heatflow values (D6v~nyi & Horvfith 1988) and gravity anomalies (Szafi~n et al. 1997). The profile images a largely deformed and extended upper crust and the presence of a shallow Moho surface. In this part of the Pannonian Basin the average thickness of the crust and lithosphere is around 25 and 60 km, respectively. Below the B6k6s basin, however, the Moho is at a depth of about 22 km and the thickness of the lithosphere is about 50 km (/~dfim et al. 1996). The line drawing interpretation of the section (Fig. 6) shows the presence of a set of low-angle normal faults that flatten out at mid-crustal depth. This regional detachment level is located at the transition between the bottom of the brittle upper crust and the top of the ductile lower crust. Low-angle normal faults dip to the NE in the Hdd-Mak6 basin, whereas on the NE side of the B~k~s Basin a switch of fault polarity can be observed. The two deep ( > 7 kin) depressions are separated by a basement high, the Pusztaf61dvgtr-Battonya block (PB in Fig. 3). Fission-track data indicate Miocene cooling ages for the Algy6 block, suggesting a genetic relationship between the onset of extension and exhumation of the footwall block (Taft et al. 1999). The style of deformation was controlled by the collapse of an overthickened, warm and thus weak lithosphere inherited from the Alpine orogeny. In this tectonic setting, Cretaceous nappe boundaries and thrust planes served as weakness zones and were reactivated as the low-angle sole of listric normal faults in an extensional (transtensional) regime. A similar structural pattern is revealed in the density model in section B - B ' (Fig. 7a) and the seismic section C - C ~ (Fig. 7b). Section B - B ' (Szafifin et al. 1999) traverses the AlpineCarpathian-Pannonian region, starting in the molasse zone of
the Eastern Alps, running through the Vienna Basin and terminating at the southern boundary of the ALCAPA terrane. Section C - C ' (Fig. 7b; after Tari 1994) shows the interpretation of an industrial seismic section in the southern part of the Danube Basin. The abundance of low-angle normal faults in the central and SE parts of sections B - B I and C - C ~ suggests a primarily extensional origin of the Danube Basin. These SE-dipping normal faults are again reactivated Cretaceous thrusts of the Austroalpine nappe stack. The faults penetrate down to mid-crustal levels where they sole out in regional detachment levels. They separate sub-basins from basement highs of different original structural positions: from NW to SE these are the Lower, Middle and Upper Austroalpine nappes. The stratigraphically highest units of the former nappe pile are on the right sides of the sections and are exposed further to the SE in the Transdanubian Range (Fig. 2). The stratigraphically lowest units are on the left sides of the sections and are represented by Lower Austroalpine nappes with blueschist-facies metamorphism (~rkai 2001). Crustal thickness values are similar to those in section A - A ' . In section B - B t the contact zone between the underplating Bohemian Massif (European plate) and the overriding Austroalpine unit (Adriatic plate) is interpreted as a gently dipping surface. The Bohemian massif can be traced at depth as far as the Mihfilyi high in the centre of the Danube Basin (marked by well M-27 in Fig. 7a), that is some 150 km behind the front of the Alpine orogen (Taft 1996). Seismic section D - D ' (Fig. 8) highlights the main structural features of late-stage basin inversion in the SW part of the Pannonian Basin. The profile traverses the broad boundary zone between the ALCAPA and Tisza-Dacia terranes (Mid-Hungarian shear zone) and images young, even active folding. The pre-rift basement is composed of a series of Palaeozoic and Mesozoic thrust sheets. The interpretation of the profile suggests a multi-phase tectonic evolution with distinct structural episodes and deformation styles. Similarly to the previous two examples, basin formation was initiated and controlled by the reactivation of suitably oriented Cretaceous thrust planes. The orientation of these faults is relatively constant (NE-SW) although their dip may switch (Budafa anticline). During Miocene times several half-grabens formed with considerable thickness of synrift deposits. The fault pattern
200
F. HORVATHETAL.
at Hahrt high indicates an important strike-slip component of faulting. Such a transtensional regime is consistent with current tectonic reconstructions, in which a mid-Hungarian shear zone played a key role in juxtaposing the two main terranes, that is the ALCAPA and Tisza-Dacia units (Csontos & Nagymarosy 1998; Fodor et al. 1999; Csontos & V6rrs 2004). During MioPliocene times, thermal subsidence resulted in the deposition of a thick sequence of post-rift sediments with an almost complete lack of faulting (e.g. Bajcsa syncline). Some grabens (Budafa) are completely inverted, suggesting differential vertical movements of the order of 500-1000 m during the Late Pliocene to Quaternary. Uplifted anticline cores have been considerably eroded. Surface topography follows the domal architecture of basin infill only in the southern part of the section. It is to be noted that basin inversion in this area is in the most advanced stage compared with other parts of the Pannonian Basin. This is due to the close proximity of the Adriatic indenter, which is considered to be the principal source of recent compressional stresses in the Pannonian lithosphere (Bada et al. 1998, 2001).
Lithospheric structure Formation of back-arc basins is a lithosphere-scale process. Knowledge of the crustal and lithospheric thickness maps of the Pannonian Basin system and the surrounding orogens provides further constraints on the mechanism of their formation and subsequent deformations. We use the two maps presented by Horvfith (1993) as starting models, and upgrade them by taking into account new data and interpretations presented in the past decade. The new map of crustal thickness (i.e. depth to the Moho discontinuity) is shown in Figure 9. There are only minor changes in the Pannonian Basin; the 3 0 k m isoline bounding the attenuated crustal domain remains the same. Inside this isoline a few new deep seismic profiles (Posgay et al. 1995, 1996) corroborated the earlier pattern or led to a better definition of a crustal thickness minimum in the southeastern part (compare with Fig. 6).
Furthermore, the improvement by Lenkey (1999) was also taken into account, as he has made the seismically derived pattern compatible with the gravity anomalies of the basin. Similarly, gravity modelling studies (Szafifin 1999), particularly 3D modelling (Szafifin & Horvfith 2006), suggest smaller thickness of crustal root associated with the southern Carpathians and its bend towards the eastern Carpathians (Vrancea area). This map is a significant improvement in the area of the Eastern Alps, and at the transition zone between the Southern Alps and the Dinarides (45-48~ 13-15~ This is because, in addition to a new synthetic map (Waldhauser et al. 1998), all available refraction and reflection seismic data have been reprocessed and reinterpreted, most recently by Cassinis & Scarascia (2003) as a contribution to an ambitious crustal profiling campaign across the Eastern Alps (TRANSALP Working Group 2002). Although this campaign has not yet finished, the first published results (e.g. Lueschen et al. 2003) are already thought provoking because of the debate on the interpretation of the observed reflectors and refractors. The two preliminary crustal profile alternatives, the 'crocodile model' and the 'lateral extrusion model', presented by the TRANSALP Working Group (2002) differ significantly below the Tauern window and the Periadriatic line to a depth of about 30 km. However, the two models agree perfectly in the geometry of the Moho discontinuity below the colliding European and Adriatic plates, and there is agreement in interpreting the results as showing the decoupling of the light continental crust from the denser mantle lithosphere (Fig. 5). In opposition to this consensus, Schmid et al. (2003) put forward a new interpretation for the complex geometry of the subducted slabs. Their results rely fundamentally on an improved tomography image of the mantle below the Eastern Alps, the Southern Alps and their transition towards the Dinarides (Lippitsch 2002). This improvement was possible because mantle tomography studies carried out in the Alpine-Mediterranean area have generally used a simplifying assumption: they supposed that in the calculation of seismic delay-time anomalies the crustal contribution had been negligible because of its fairly constant value all over the imaged territory.
Fig. 9. Depth to the Moho in the PannonianBasin and the surroundingmountains. Values are given in kilometres. 1, foreland (molasse)foredeep; 2, flyschnappes; 3, pre-Tertiary units on the surface; 4, Penninic windows; 5, Pieniny Klippen Belt; 6, trend of abrupt change in crustal thickness. Modifiedafter Horvfith(1993), Lenkey (1999), Cassinis & Scarascia (2003), and Szafifin& Horv~ith(2006).
PANNONIAN BASIN Obviously, this is not the case in the peri-Adriatic region, which is characterized by strongly contrasting crustal domains. Accordingly, the reliability and resolution in teleseismic tomography critically depends on the precise knowledge of the 3D crustal velocity structure. With an a p r i o r i crustal correction and selection of high-quality teleseismic data, Lippitsch (2002) has been able to decrease the cell size by a factor of 10 relative to previous studies and thus obtain a high-resolution image of the upper mantle structure beneath the Alps and surrounding areas. This confirms the earlier results in the Western and Central Alps depicting a high-velocity slab from the European foreland towards the SE and south to a depth of about 300 km beneath the Po plain (Kissling 1993; Solarino et al. 1996; Wortel & Spakman 2000). This European lithospheric slab underthrusting the Adriatic region appears to be continuous, but detached slabs of earlier subduction are also present in the deeper upper mantle (Davies & v o n Blanckenburg 1995). The real novelty of the high-resolution tomography image is the demonstration that there is a change in the polarity of subduction in the Eastern Alps. To the east of the western edge of the Tauern window, the Adriatic lower lithosphere is subducting beneath the Alpine wedge, probably reaching a depth of about 230 km. This northward dipping slab is made up of the continental part of Adria and probably an oceanic segment in the frontal part related to the former Vardar ocean (Lippitsch 2002). It can be considered as an active part of the subducting Adriatic lithosphere, which is apparently inactive and already fully detached further to the SE along the Dinarides (Wortel & Spakman 2000). Schmid et al. (2003) pointed out that two major orogenperpendicular post-collisional features at the surface coincide with the change of subduction polarity at depth: the Guidicarie strike-slip and the Brenner normal faults. They suggested that the northward dipping Adriatic subduction is a young feature (c. 20 Ma), coeval with the formation of these important orogenperpendicular features. It should be noted that the Brenner fault is the western boundary of the extruding ALCAPA wedge at the
201
Tauern window. Furthermore, they interpreted the TRANSALP profile in terms of northward descending Adriatic Moho under the European crust in analogy to the lithospheric configuration revealed by high-resolution tomography. We favour this interpretation, which is reflected in our crustal and lithospheric thickness maps (Figs 9 and 10). Unfortunately, the high-resolution tomography imaging (Lippitsch 2002) did not extend well into the region of the Dinarides and Pannonian Basin. Therefore the lateral extent of the subducted Adriatic slab towards the Dinarides remains unanswered, and hence, open to speculation. We assume that the surface expression of the boundary of the active Adriatic slab should be again an orogen-perpendicular structural feature, such as the Zagreb line (Fig. 10). The presence of a subducted lithospheric slab in the southern east Carpathians has been known for a longer time because of the remarkable deep seismic activity (down to 200 kin) in the Vrancea zone (Oncescu 1984) and early results of seismic tomography (Spakman et al. 1993). More recent studies (Fan et al. 1998; Wortel & Spakman 2000) have imaged the structure with a higher resolution and have shown a high-velocity slab at the same place but to a depth of 300-350 km. A specially devised regional campaign using about 150 broadband seismometers (Wenzel et al. 1998, 2002) revealed a nearly vertical high-velocity slab to a depth of 230-240 km. These results imply that the deeper part of the slab is aseismic. In addition, a horizontal high-velocity zone has been imaged between the 410 and 660 km seismic discontinuities, which is present beneath the entire Pannonian Basin (Wortel & Spakman 2000). The 410 and 660 km seismic discontinuities are also detected as an electrical conductivity increase by magnetotelluric and magnetovariational soundings (Adfim 1993; Semenov et al. 1997). Between them there is a transitional zone in electrical conductivity which may correspond to the high-velocity zone detected by the tomography. The horizontal high-velocity body may also represent a subducted lithospheric slab. If this is the case, it could have derived from subduction related to the consumption of the Mesozoic Vardar ocean, as
Fig. 10. Lithosphericthicknessmap of the PannonianBasin and the surroundingmountains. Values are given in kilometres. 1, Foreland (molasse)foredeep; 2, flysch nappes; 3, pre-Tertiary units on the surface; 4, Penninic windows; 5, Pieniny Klippen Belt; 6, trend of abrupt change in lithosphericthickness. Modified after Horv~ith(1993) and/~dfim & Wesztergom (2001).
202
F. HORVATHETAL.
well as from detachment of subducted slabs around the rest of the Carpathian arc. Detailed studies of the tectonic and magmatic evolution of the Apuseni Mts., Transylvanian Basin and east Carpathians (Linzer 1996; Girbacea & Frisch 1998; Gvirtzman 2002), together with the results of seismic tomography, suggest that the Vrancea slab is the only existing relict of the subduction roll-back that facilitated the extension of the ALCAPA and Tisza-Dacia terranes. Slab break-off is in progress, but the vertical continuity of the slab is still partly established between two orogen-perpendicular boundaries: the Trotu~ fault and the Intramoesian fault (Fig. 10). The lithospheric thickness map has also been improved remarkably in the Pannonian Basin as a consequence of new magnetotelluric soundings and sophisticated inversion of all available apparent resistivity and phase shift sounding curves (Adfim & Wesztergom 2001). This shows that the depth to the low-resistivity asthenosphere is on average 60-65 km in the basin, and locally, as the B6k& depression, rises to 50 km depth (Adfim et al. 1996). Tomography studies (Weber 2000) revealed the presence of relatively high seismic velocities beneath the B~k~s depression. Weber (2000) explained this phenomenon by local compositional differences in the mantle as a result of a mafic intrusion corresponding to the continental rift model as indicated by interpretation of gravity and magnetic measurements (e.g. Adfim & Bielik 1998).
Extrusion tectonics Extrusion of the ALCAPA terrane into the Carpathian embayment is a well-established and widely accepted concept (Fig. 5). However, the other constituent of the substrata of the Pannonian Basin, that is the Tisza-Dacia terrane, is characterized by a more complex and, hence, less understood structural and kinematic history. Paleotectonic reconstructions, albeit different in detail, agree that this terrane underwent Eo- and Meso-Alpine deformation in the Vardar zone, where continental collision took
place in the Early Tertiary (Csontos & V6r6s 2004). It is plausible to assume that the Tisza-Dacia block was extruded from this collision zone under similar geodynamic conditions to those of the ALCAPA terrane. This primarily implies that strain partitioning also occurred there and a Tisza-Dacia crustal flake was detached from its mantle lithospheric root. This means that, from the Late Oligocene, two compressionally deformed crustal wedges were invading the embayment of the Carpathian flysch basin. Evidence to support this view is as follows. Both terranes are characterized by extensive Early to Mid-Miocene silicic volcanism of identical geochemical composition (Harangi 2001). This implies that lower crustal melting and low crustal strength were common features of the two terranes. The anomalously high temperature gradient from the Early Miocene is well constrained by subsidence, thermal and maturation history analysis carried out for Neogene basins on both terranes (Horvfith et al. 1988). It was Sclater et al. (1980) who first arrived at the conclusion that the subsidence and sedimentation history of the Pannonian Basin system could not be explained in terms of homogeneous stretching of a normal lithosphere. Instead, the data were compatible with an inhomogeneous stretching model, whereby a modest crustal extension (stretching factor of 1.2-2.5) accompanied a dramatic attenuation of the mantle lithosphere (stretching factor of 5-50). More sophisticated modelling techniques, including evolution of the thermal field and organic matter maturation data (Horwith et al. 1986, 1988), corroborated this first conclusion about the apparent disappearance of mantle lithosphere at the beginning of the synrift phase in the Pannonian Basin system. Sclater et al. (1980) disagreed on the geodynamic interpretation of this conclusion, which was indeed debatable at the time. Now, we believe that extrusion of crustal flakes rather than entire lithospheric blocks offers the most plausible explanation (Fig. 5). In other words, extrusion tectonics in the Pannonian Basin is more compatible with the concept of orogenic floating (Oldow et al. 1989), rather than extrusion of lithospheric blocks as in Central Asia (Lav6 et al. 1996).
Fig. 11. Geodynamicmodel of the Quaternaryto recent PannonianBasin and surrounding orogens. 1, West European Platform and East European Craton separated by the Tornquist-Teisseyre (TT) suture zone; 2, foreland (molasse) foredeep; 3, Alpine-Carpathian nappe system; 4, exposed metamorphic core complexes (Penninicunits); 5, upliftingpart of the Pannonianbasin; 6, subsidingpart of the PannonianBasin; 7, extinct and active (duringthe Quaternary) calc-alkalinevolcanism; 8, thrust plane, inactive and active thrust front; 9, orogen-paralleldisplacementof the Pannonianterranes. M and L denote base of the crust (Moho) and the mantle lithosphere, respectively.
PANNONIAN BASIN
203
Conclusions
References
It can be seen from this review that a mass of new data has been collected in the Pannonian Basin and the surrounding orogens in the past decade, largely as a result of co-ordinated research efforts in the framework of the EUROPROBE programme. New data and their interpretations have led to significant progress in understanding the lithospheric dynamics associated with the formation and deformation of the Pannonian Basin and surrounding orogens (Fig. 11). It can be recognized that back-arc basin evolution is a single, although very complex, dynamic process of continental collision zones. This is because the various processes involved are strongly interrelated and their fundamental physical principle is the same: to minimize the potential and deformational energy during collision. The most important processes and their mechanisms are as follows: (1) Plate convergence leads to the complete consumption by subduction of the less buoyant oceanic lithosphere between the two continents, and then continent-continent collision occurs. (2) Continuing convergence results in splitting of the two continental lithospheres, preferably along major density a n d / o r rheological boundaries, and strain partitioning occurs. (3) When the light crust is detached from the continental lithosphere, the remaining mantle lid is heavier than the asthenosphere; hence, subduction can continue. Subduction polarity may remain the same as before; however, double subduction or a change of polarity can also take place. (4) At the upper levels, intricate structural processes produce the orogenic wedge, which consists of an interfingering stack of obducted parts of the consumed oceanic lithosphere, the detached continental crustal blocks and their sedimentary cover. (5) The orogenic wedge grows vertically and also in orogennormal direction, provided that convergence continues. Vertical growth implies thickening of the wedge and an increase in its elevation as a result of isostatic forces. Orogen-normal growth leads to the propagation of thrust faults towards both the undeformed foreland and the hinterland, and their incorporation into the orogenic wedge. (6) Parts of the orogenic wedge may extrude in orogen-parallel direction, provided that it decreases the potential energy and deformation energy dissipation of the system. This requires that a 'free boundary' is available sideways, which can be related to the presence of an oceanic basin. (7) Landlocked oceanic basins are particularly important, as their subductable oceanic lithosphere can give way to lateral extrusion of the orogenic wedge. Extrusion is a sideways extension of the orogenic material accompanied by gravitational collapse of the overthickened wedge. (8) Subduction of the oceanic lithosphere of the landlocked oceanic basin occurs by roll-back of the hinge zone (trench). This process continues until the downward force of the subducted slab is balanced by buoyant forces acting on the transitional to continental lithosphere attached to the subducting slab. (9) Roll-back terminates and the subducted slab breaks off progressively along the arc produced by the retreat of the landlocked oceanic lithosphere. This results in the ending of orogen-parallel extension of the extruding wedge, and the stress field in the back-arc basin changes to compressive. The fate of the orogen critically depends on the interaction of forces driving plate convergence and resisting plate collision. This is, however, a global energetic problem of the plate tectonic system.
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This paper is an outcome of the traditional co-operation between the Ertvrs University, Budapest and the Vrije Universiteit, Amsterdam. Financial support was provided by the Hungarian National Science Fund (OTKA) projects T034928, D34598 and F043715, and the Netherlands Research Centre for Integrated Solid Earth Science (ISES). The reviewers of the paper are thanked for their useful comments and suggestions.
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Modes of basin (de)formation, lithospheric strength and vertical motions in the Pannonian-Carpathian system: inferences from thermo-mechanical modelling S. C L O E T I N G H 1, G. B A D A l, L. M A T E N C O 1, A. L A N K R E I J E R 1, F. HORV/~TH 2 & C. DINU 3 1Netherlands Research Centre f o r Integrated Solid Earth Science, Faculty o f Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 H V Amsterdam, The Netherlands (e-mail:
[email protected]) 2Department o f Geophysics, EOtvts L. University, Pdzmdny P. s. l / C , 1117 Budapest, Hungary 3Faculty o f Geology and Geophysics, University o f Bucharest, str. Traian Vuia 6, sect. 1, 70139 Bucuresti, Romania
Abstract: After a rapid multiphase evolution and a transition from passive to active rifting during late Early Miocene to Pliocene times, the Pannonian Basin has been subjected to compressional stresses leading to gradual basin inversion during Quaternary times. Stress modelling demonstrates the significance of the interaction of external plate-boundary forces and the effect of gravitational stresses caused by continental topography and crustal thickness variation. Flexural modelling and fission-track studies have elucidated the complex interplay of flexural downloading during collision, followed by rapid unroofing by unflexing and isostatic rebound of the lithosphere. The stretching and subsidence history of the Pannonian Basin, the temporal and spatial evolution of the flexure of the Carpathian lithosphere, and the lithospheric strength of the region reflect a complex history of this segment of the Eurasia-Africa collision zone. The polyphase evolution of the Pannonian-Carpathian system has resulted in strong lateral and temporal variation in thermomechanical properties in the area. Modelling results suggest that, as a whole, the Pannonian Basin has been an area of pronounced lithospheric weakness since Cretaceous time, shedding light on the high degree of strain localization in this region. This basin, the hottest in continental Europe, has a lithosphere of extremely low rigidity, making it prone to multiple tectonic reactivations. Another feature is the noticeable absence of lithospheric strength in the mantle lithosphere of the Pannonian Basin. Modelling studies suggest pronounced lateral variations in lithospheric strength along the Carpathians and their foreland, which have influenced the thrust load kinematics and post-collisional tectonic history. The inferences and models discussed in this paper are constrained by a large geophysical database, including seismic profiles, gravity and heat-flow data.
The Pannonian-Carpathian system in Central and Eastern Europe has been the focus of considerable research efforts concerning the integration of geophysical and geological data, making it a key area for quantitative basin studies. A vast geophysical and geological database has been established in recent decades as a result of a major international research collaboration in this area, largely carried out in the framework of European programmes such as the EU Integrated Basin Studies project (Cloetingh et al. 1995; Durand et al. 1999), the ILP-ALCAPA (Cloetingh et al. 1993; Neubauer et al. 1997) and EUROPROBEPANCARDI (Decker et al. 1998) programmes, and the Peri-Tethys programme (Ziegler & Horwith 1996; Brunet & Cloetingh 2003), partly funded in the context of petroleum exploration. These studies, building on previous compilations (Royden & Horvfith 1988), marked a major advance in applying basin analysis concepts to the Pannonian-Carpathian system. An important asset of this natural laboratory is the existence of high-quality constraints on basin evolution obtained through the systematic acquisition of seismic, gravity, heat-flow and magnetotelluric data by various research groups (see Radulescu et al. 1976; Royden & Horv~ith 1988; D6vtnyi 1994; Demetrescu & Andreescu 1994; Ionescu 1994; Posgay et al. 1995; Szafi~n et al. 1997; Adfim & Bielik 1998; Tari et al. 1999; Wenzel et al. 1999; Hauser et al. 2001; Lenkey et al. 2002). Extensive well coverage in the context of petroleum exploration and surface studies have allowed the construction of a high-resolution stratigraphic framework for the area (e.g. Vakarcs et al. 1994; Sacchi et al. 1999; Vasiliev et al. 2004). At the same time, the fold-and-thrust belt has been the focus of a concentrated effort, highlighting the connection between lateral variations in structural style, basement characteristics and foreland flexure development in various segments of the Carpathian orogen (Sfindulescu 1988; Roure et al. 1993; Matenco et al. 1997a,b, 2003; Taft et al. 1997; Sanders et al. 1999; Zoetemeijer et al. 1999) and its hinterland, the Transylvanian Basin (e.g. Ciulavu et al. 2002). The Pannonian-Carpathian system, therefore, allows us to test models for basin formation and subsequent deformation, for continuing orogeny and continental collision. This system
comprises some of the best documented sedimentary basins in the world, located within the Alpine orogenic belt, at the transition between the Western European lithosphere and the East European Craton. The Pannonian Basin evolved from its synrift to post-tiff phase during Early to Late Miocene times (c. 2 0 5 Ma), when back-arc extension was coupled with subduction and collision dynamics in the Carpathian orogenic arc system (Royden & Horvfith 1988). The lithosphere of the Pannonian Basin is a particularly sensitive recorder of changes in lithospheric stress induced by near-field intra-plate and far-field plate boundary processes (Bada et al. 2001). High-quality constraints exist on the regional (palaeo)stress (Fodor et al. 1999; Gerner et al. 1999) fields in the lithosphere as a result of earthquake focal mechanism studies, analyses of borehole break-outs and studies of kinematic field indicator data. A close relationship has been demonstrated between the timing and nature of stress changes in the extensional basin and structural episodes in the surrounding thrust belts, pointing to a mechanical coupling between the orogen and its back-arc basin. The Pannonian Basin, the hottest in continental Europe, is thought to have gone through a rapid transition from passive to active tiffing during Late Miocene times, simultaneously with the climax of compression in the Carpathian arc (Huismans et al. 2001). In parallel, significant efforts have been made to reconstruct the spatial and temporal variations in thrusting along the Carpathian orogen (Roure et al. 1993; Schmid et al. 1998; Zweigel et al. 1998; Matenco & Bertotti 2000) and its relationship to foredeep depocentres (Meulenkamp et al. 1996; Matenco et al. 2003; Tfirfipoancfi et al. 2003), changes in foreland basin geometry and lateral variations of flexural rigidity. A general feature of the flexural modelling studies carried out for the Carpathian system (e.g. Matenco et al. 1997b; Zoetemeijer et al. 1999) is the inferred low rigidity of the platform lithosphere downbending under the SE Carpathians. Flexural studies constrained by gravity (e.g. Szafifin et al. 1997) also point to an important role of flexural unroofing of the Carpathian mountain chain and its foredeep. A crucial element of the dynamics of lithospheric deformation is the mechanics of coupling back-arc deformation in the
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 207-221. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Pannonian Basin with continental collision and foreland basin evolution along the Carpathian arc. This has been addressed through a combination of dynamic and kinematic modelling studies constrained by integrated basin analysis in the Pannonian sector and by thermochronology, basin modelling and structural field studies in the Carpathian belt. In particular, the inversion events (i.e. Late Miocene and Late Pliocene-Quaternary) recorded in the Pannonian Basin (Horvfith 1995; Fodor et al. 1999) are coeval with the climax of thrusting in the Carpathians (e.g. S~ndulescu 1988; Hippolite et al. 1999) related to continental collision and late-stage out-of-sequence contraction. Horv~ith & Cloetingh (1996) established the importance of Late Pliocene to Quaternary compression in the Pannonian-Carpathian system, explaining its anomalous Quaternary uplift and subsidence pattern as well as its intraplate seismicity, and thus establishing a novel conceptual model for structural reactivation of back-arc basins in orogenic settings. The basin system has reached an advanced stage of evolution with respect to other Mediterranean back-arc basins and its structural inversion has been taking place for the last few millions years. Basin inversion is related to temporal changes in the regional stress field (for a general discussion, see Ziegler et al. 1995, 2002), from one of tension that controlled basin formation and subsidence, to one of compression resulting in contraction and flexure of the lithosphere associated with differential vertical movements. This paper presents an overview of the tectonic evolution of the Pannonian-Carpathian system in terms of various thermomechanical models. First, the focus is placed on the formation of the system as inferred from subsidence history and stretching models of the Pannonian Basin, and the flexural behaviour of the Carpathian lithosphere. Then a lithospheric strength map of the system is presented, which highlights rheological constraints for basin analysis studies and for the reconstruction of the deformation history of the area. The reactivation or the neotectonics of the region is described in terms of anomalous late-stage vertical movements, that is accelerated subsidence in the centre of the Pannonian Basin and fast uplift of the Carpathians orogen as a result of the rebound of the crust in the aftermath of continental convergence. The paper is concluded with a discussion on the thermo-mechanical aspects of basin inversion, lithospheric folding, and related temporal and spatial variations of continental topography in the Pannonian-Carpathian system.
Formation of the Pannonian-Carpathian system Formation and N e o g e n e evolution o f the Pannonian Basin Dynamic models of basin formation. After a long period of continen-
tal convergence in the Alpine belt during Cretaceous to Palaeogene times, a rapid change in tectonic style in the late Early Miocene led to the formation of the Pannonian Basin in the area of the Carpathian embayment. Consequently, the relatively stable Palaeogene to Early Miocene assembly of continental blocks at the axial zone of Adria-Europe convergence was completely disintegrated and these units experienced a significant amount of stretching, rigid body rotation and translation. This process was coeval with the formation and early evolution of the Pannonian Basin and the large-scale tectonic transport of the flysch nappes in the Carpathian arc (e.g. Balla 1984; S~ndulescu 1988; Csontos et al. 1992; Roure et al. 1993; Kov~i~ et al. 1994; Fodor et al. 1999). Several models have been proposed to explain the dynamics of Neogene tiffing in the Pannonian Basin. An active v. passive mode of rifting has been a matter of continuous debate in the scientific literature (for a review, see Bada & Horwith 2001), resulting in the advent of various dynamic models during the last half-century (Fig. 1). These different theoretical scenarios have been provoked by the most evident features of this basin system, that is thinned and hot v. thickened and colder lithosphere in the central and peripheral sectors, respectively. For instance, Sz~deczky-Kardoss (1967) and, at least in his early works, Stegena (1967) argued for the presence of a mantle diapir beneath the Intra-Carpathian area (Fig. 1a). In this model the ascending mantle flow resulted in thinning and subsidence (active rifting) in the central areas, whereas the nappe structure and the root of the surrounding orogens were formed above the descending branch of a local convection cell. Active tiffing was also proposed by, for example, Horv~th et al. (1975) and Stegena et al. (1975) as an ultimate origin of the Pannonian Basin. However, their model already employed the concept of plate tectonics for the Pannonian region, where basin formation and subsidence, intense Neogene-Quaternary volcanic activity, extremely high heat flow, and the presence of an anomalous upper mantle and thinned crust were considered as closely related phenomena. These features were probably controlled by subcrustal erosion of the underlying lithosphere by an active mantle diapir
Fig. 1. Dynamic models proposed for the evolution of the PannonianBasin system (after Bada & Horvfith2001). (a) Asthenospheric doming results in active rifting of the lithosphere above the central axis of the dome, whereas shorteningis taking place in the peripheral areas. (b) Active rifting may also be caused by a subduction-generated mantle diapir. (c) Hinge retreat of the subductingEuropean margin driven by the negativebuoyancyof the slab induces passive rifting in the overriding plate. (d) The same hinge retreat may be sustainedby an eastward mantle flow pushing against the downgoing slab. (For references see text.)
MODELLING THE PANNONIAN-CARPATHIANS SYSTEM above the European and Apulian (Adriatic) plates subducting under the Pannonian plate fragment (Fig. lb). Other dynamic models also suggest the key importance of subduction as a driving force, and the back-arc position of the Pannonian Basin (e.g. Royden & Horwith 1988; Csontos et al. 1992; Horv~th 1993; Csontos 1995; Linzer 1996). Extension and lithosphetic stretching are caused by the hinge retreat of the subducting European plate along the Carpathians that stretches the overriding Pannonian plate (Fig. lc). The process of slab retreat is caused by the negative buoyancy of a subducting slab (Royden & Karner 1984). Another model (Fig. l d) suggests that an eastward mantle flow may have been pushing the subducting slab and, hence, can be responsible for the same hinge retreat (Doglioni 1993). Nevertheless, both models account for the passive tiffing of the Pannonian Basin, where tension is facilitated by trench suction forces exerted at the contact zone of the overriding and subducting plates. More recently, Huismans et al. (2001) used numerical simulation to explain the temporal changes in rifting style of the Pannonian Basin. The results of their thermo-mechanical finiteelement modelling confirm the two-phase evolution scheme of the system, that is a synrift episode followed by a post-rift period. Initiation of basin formation was mainly driven by passive rifting as a result of the effect of Carpathian subduction and the gravitational collapse of the thickened pre-rift Pannonian lithosphere. This triggered a small-scale convective upwelling of the mantle lithosphere that could have led to the second, active tiffing phase of the Pannonian Basin during Late MiocenePliocene times, coeval with the climax of compression in the Carpathian arc. Stretching models, subsidence analysis. Efforts on the quantification of basin evolution started in the early 1980s by means of classical basin analysis techniques. The Pannonian Basin has been a key area for testing stretching models, because of the availability of excellent geological and geophysical constraints. At the same time, the main characteristics of the basin system, such as the extremely high heat flow, the presence of an anomalous thinned lithosphere and its tectonic position in the Alpine regime of overall convergence, made it particularly suitable and challenging for basin analysis. Interest in this research in the Pannonian Basin was mainly generated through activities in the field of deep crustal and mantle processes, local tectonic and regional correlation studies, and hydrocarbon prospecting. The stretching model of McKenzie (1978) was first applied to the intra-Carpathian basins by Sclater et al. (1980). They found that the formation of the peripheral basins could be fairly well simulated by the concept of uniform extension, using a stretching factor of about two (/3 = 2). In the more central basins, however, the considerable thermal subsidence and high heat flow suggested unrealistically high stretching factors (/3 up to five). Thus, they postulated a differential extension within the Pannonian lithosphere, with moderate crustal thinning being accompanied by a higher level of stretching at subcrustal depth. Building on this and using a wealth of well data, Royden et al. (1983) introduced the concept of modified or non-uniform stretching, in which the amount of extension is depth-dependent in the lithosphere. The subsidence and thermal history of major parts of the Pannonian Basin suggests a large attenuation of the mantle lithosphere with respect to a finite crustal extension. Horv~th et al. (1988) further improved this concept by considering radioactive heat generation in the crust, and the thermal effect of basin-scale sedimentation. By reconstructing the subsidence and thermal history, the thermal maturation of organic matter in the central region of the Pannonian Basin (Great Hungarian Plain) was calculated, resulting in a major step forward in the field of hydrocarbon prospecting by means of basin analysis techniques. These studies highlighted the considerable difficulties in explaining basin subsidence and crustal thinning in terms of uniform extension, pointing to the presence of anomalous subcrustal
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thinning. The issue has been a central theme for subsequent investigations by means of quantitative subsidence analyses (backstripping) of a more extended set of sections and wells in the Pannonian Basin, and by means of forward modelling techniques (Lankreijer et al. 1995; Sachsenhofer et al. 1997; Juhfisz et al. 1999; Lenkey 1999). Kinematic modelling incorporating the concept of necking depth and finite strength of the lithosphere during and after rifting (van Balen et al. 1999), as well as dynamic modelling studies (Huismans et al. 2001) suggested a transition from passive to active tiffing as a mechanism for subcrustal flow and small-scale convection in the lithosphere-asthenosphere system. To quantify lithospheric deformation on a whole-basin scale, Lenkey (1999) carried out forward modelling using the concept of non-uniform stretching and taking into account the effects of lateral heat flow, flexure and necking of the lithosphere. Calculated crustal thinning factors (6) indicate large lateral variation of crustal extension in the Pannonian Basin (Fig. 2). This is consistent with the areal pattern of pre-Neogene basement depth (Horvfith et al. 2006). The range of crustal thinning factors indicates 10100% crustal extension, which is in good agreement with the pre-rift palinspastic reconstruction of the Pannonian Basin, and the amount of cumulative shortening in the Carpathian orogen (e.g. Roure et al. 1993; Fodor et al. 1999). As a major outcome of basin analysis studies, Royden et al. (1983) provided a two-stage subdivision for the evolution of the Pannonian Basin, with a synrift (tectonic) phase during Early to Mid-Miocene times and a post-rift (thermal) phase during the Late Miocene-Pliocene period. Further development of the stratigraphic database, however, demonstrated the need to refine this scenario. According to Tari et al. (1999), a regional Mid-Badenian unconformity indicates the termination of the synrift period, which is followed by a post-rift phase with only minor tectonic activity. In either case, the subsidence history of the Pannonian Basin can be subdivided into three main phases as reflected in the subsidence curves of selected sub-basins (Fig. 3). First, the synrift phase characterized by a rapid tectonic subsidence, started synchronously at about 20 Ma in the entire Pannonian Basin. This phase of pronounced crustal extension is recorded everywhere in the basin system and was mostly limited to relatively small, faultbounded grabens or sub-basins. Second, the post-rift phase of extension affected much broader areas, causing general downwarping of the lithosphere manifested in a thermal phase of subsidence. This is particularly well defined in the central parts, suggesting a gradual transition from thin-skinned to wholelithosphere extension towards the centre of the basin system (e.g. Sclater et al. 1980; Royden & Drvrnyi 1988). The third (final) phase of basin evolution is characterized by the gradual structural inversion of the Pannonian Basin system during Late Pliocene-Quaternary times. As a result, the recent build-up of intraplate compressional stresses has caused basin-scale buckling of the Pannonian lithosphere associated with late-stage subsidence anomalies and differential vertical motions (Horwith & Cloetingh 1996). As seen in the subsidence curves (Fig. 3), an accelerated subsidence has been taking place in the central depressions (Little and Great Hungarian Plain; Fig. 3b and c), whereas most peripheral sub-basins have been uplifted by a few hundred metres after mid-Miocene times (Styrian and East Slovakian basins; Fig. 3d and e) or during the Pliocene-Quaternary period (Zala basin; Fig. 3f). The late-stage tectonic reactivation, as well as other episodic inversion events in the Pannonian Basin (Horvfith 1995; Fodor et al. 1999), highlight the importance of tectonic stresses not only in the stretching phase (tension) but also during subsequent inversion phases (compression). Modelling curves for the Carpathian foreland (Fig. 3h-j) indicate an important phase of basin subsidence related to the collision in the east and south Carpathians. Moreover, this period is coeval with the cessation of synrift subsidence in the Pannonian Basin (see also Horv~ith & Cloetingh 1996). Subsidence curves for the intermediate Transylvanian Basin (Fig. 3g) indicate a dual behaviour;
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Fig. 2. Crustal thinning factors (6) calculated by means of forward modelling for the PannonianBasin employingthe concept of non-uniformstretching complemented by the effects of lateral heat flow and the flexure of the lithosphere (after Lenkey 1999). The pronounced lateral variation of crustal extension, leading to the formation of deep sub-basins connected by areas of lower level of deformation, should be noted. AM, Apuseni Mts; DIN, Dinarides;EA, Eastern Alps; TR, TransdanubianRange; SC, WC, southern and western Carpathians, respectively. Local depressions of the Pannonian Basin system: B~, Brk~s; Da, Danube; De, Derecske; D, Drava; ES, East Slovakian; Jfi, Jfiszsfig;Ma, Mak6; Sa, Sava; St, Styrian; Vi, Vienna; Za, Zala. whereas the Badenian-Pannonian stage reflects a subsidence pattern similar to that in the Carpathian foreland, the major Pliocene-Quaternary inversion in the Pannonian Basin is also recorded in the Transylvanian domain.
F o r m a t i o n a n d N e o g e n e evolution o f the Carpathian system
Over the last few years, research has focused on spatial and temporal variations in thrusting along the Carpathian arc and their relationship to unusual foredeep geometry and lateral variations in flexural behaviour. The reconstruction of orogenic uplift and erosion (e.g. Sanders et al. 1999) coupled with foreland subsidence modelling (Matenco et al. 2003) have elucidated the complex interplay of flexural downloading during collision and its lateral variability, followed by isostatic rebound in the orogenic wedge and increased subsidence in the SE Carpathian corner (Matenco et al. 1997b; Zoetemeijer et al. 1999; Bertotti et al. 2003; TS.rfipoancfi et al. 2003; Cloetingh et al. 2004). The Carpathians represent a highly arcuate orogenic belt formed in response to subduction and continental collision between the European and Apulian plates and related microplates during the Alpine orogeny (Sfindulescu 1988; Csontos & Voros 2004). It consists of a nappe pile of crystalline rocks with Late Palaeozoic to Mesozoic sedimentary cover and, in an external position, an Early Cretaceous to Tertiary thin-skinned belt. The Alpine tectonic evolution of the Carpathians is traditionally subdivided into Triassic to Early Cretaceous extension followed by MidCretaceous to Pliocene shortening (e.g. Sandulescu 1988). Three main Tertiary deformation stages are recognized (e.g. Matenco & Bertotti 2000, and references therein). During PalaeogeneEarly Miocene times, the clockwise rotation of the Rhodopian fragment (Balla 1986; Schmid et al. 1998), part of the
Pannonian-Carpathian system, caused N N E - S S W to E N E WSW shortening in the internal Moldavides nappes (Sandulescu 1988). The ENE-WSW-directed shortening acting in Mid-to Late Miocene (Late Burdigalian to Sarmatian) times was responsible for the major nappe emplacement in the east Carpathians. The oroclinal shape of the thrust belt must have been initiated in the later phases of deformation, as a result of the irregular plate boundaries and the lateral variations in thickness of the sedimentary wedge involved in shortening (Matenco & Bertotti 2000). Widespread left-lateral shearing occurred along numerous eastwest-oriented faults in the north and NW-SE-trending dextral transpressive structures in the south, accommodating movement to the ESE of the intervening central sectors. Late Miocene to Pliocene N W - S E (to north-south) shortening in the east Carpathians led to further deformation mainly concentrated in the external parts of the junction zone between the east and south Carpathians (Bend Zone), and the foreland continued to subside throughout the post-collisional period, accumulating up to 6 km of Pliocene-Quaternary sediments in the Focsani Basin area (Tfir~poancfi et al. 2003). The Carpathians represent the birthplace for popular theories such as slab-retreat and slab roll-back based on flexural modelling results (e.g. Royden & Karner 1984; Royden 1993), which account for a gradual, eastward retreat of the plate contact during the Miocene contractional events. Detailed flexural modelling studies indicate that the European lithosphere primarily controls the deflection in the Carpathian foredeep (Zoetemeijer et al. 1999; Fig. 4b). It appears that small remnants of oceanic slab cause slab-pull forces, increasing in magnitude in an eastward direction. Modelling results indicate the regional importance of post-collisional uplift, which is most pronounced in the vicinity of the Bohemian massif, related to flow of mantle material away from the multiple rifting Pannonian Basin (Huismans
MODELLING THE PANNONIAN-CARPATHIANS SYSTEM
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Fig. 3. Subsidence curves for selected sub-basins of the Pannonian Basin and Carpathian foreland (compiled after Horvfith et al. 1988; D6v6nyi 1994; Lankreijer et al. 1995; Sachsenhofer et al. 1997; Ma~enco et al. 2003). It should be noted that after a rapid phase of general subsidence throughout the Pannonian Basin, the sub-basins show a distinct subsidence history from Mid-Miocene time. Arrows indicate generalized vertical movements. Timing of the synand post-rift phases is after Royden et al. (1983). Timing of Carpathian collision after Matenco et al. (2003). For the time scale the central Paratethys stages of R6gl (1996) are used. O, K, Ottnangian and Karpatian, respectively (Early Miocene); BAD, SA, Badenian and Sarmatian, respectively (Mid-Miocene); PAN, PO: Pannonian and Pontian, respectively (Late Miocene); PL, Pliocene; Q, Quaternary. AM, Apuseni Mts; DIN, Dinarides; EA, Eastern Alps; PanBas, Pannonian Basin; SC, WC, southern and western Carpathians, respectively.
e t al. 2001). The European foredeep configuration is clearly controlled by lateral changes in the European lithosphere strength. These changes are dominated by the amount of internal deformation caused by the curvature of the lower plate during the emplacement of the Carpathians, weakening of the elastic plate
coinciding with the areas of m a x i m u m flexural bending stresses (Zoetemeijer e t al. 1999). Flexural modelling studies also predict a general back-stepping of the lower plate system in the transition area between European and Moesian lithosphere (Matenco e t al. 1997b) along the crustal-scale Trotu~ fault
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(Fig. 4c), accounting for crustal lower plate tearing and disruption of the mechanical continuity between the two types of lithosphere involved in collision. The kinematics of this fault records a large-scale sinistral movement during collision, and continuing subsidence in the post-collisional period has induced the downward throw of this fault to the south. Modelling the deflection of the Moesian lithosphere below the south Carpathian nappe pile indicates a plate contact placed in an advanced, foreland position, demonstrating the small amount of general n o r t h - s o u t h oriented shortening in an overall transpressional regime (Matenco e t al. 1997b). Recent 3D flexural studies (Tfir~poancfi e t al. 2004) indicate that the subsidence of the unusual deep SE Carpathian foredeep is affected by pre-orogenic extension followed by a 3D emplacement of the thrust load on top of large lateral changes in lithospheric strength in the downgoing plate. Under these conditions the m a x i m u m deflection moves out of the orogenic load and accounts for large-scale subsidence in the Bend Zone (Fig. 4d).
The subsidence pattern recorded in the Carpathian foreland units during the Tertiary is characterized by significant vertical motions (e.g. Matenco e t al. 2003; Bertotti e t al. 2003; Fig. 3h-j). Opening of the Early Miocene transtensional basin in the Getic depressionwestern Moesian platform led to the deposition of up to 5 km of sediments, whereas the other platform areas were characterized by non-deposition a n d / o r erosion. Starting from the Mid- to Late Miocene, the entire Carpathian foreland started to subside (e.g. at a rate of 1 5 0 0 - 3 0 0 0 m M a -1 in and around the Focsani depression in the Late Miocene) as a direct response to coeval thrust loading, the major depocentre of the foreland basin being located in the Bend Zone. Following the collisional event at the beginning of the Late Miocene, the subsidence continued particularly in the Focsani depression, where a rate of 2 0 0 - 3 0 0 m Ma -1 is recorded for the entire Pliocene-Pleistocene). Subsidence analysis in the Carpathian foreland demonstrated comparable kinematically related vertical motion episodes simultaneously occurring in the frontal part of both the east and south Carpathians
Fig. 4. Flexural studies across the Carpathians foreland. (a) Location of the flexural studies in the Carpathians orogen. (b) Simplified geological map of the western Carpathians, indicating the predicted flexural bending stresses (magnitudes are in GPa) in the 2D modelled profiles (after Zoetemeijer et al. 1999). (c) Computed contour map of the basement shape along the Romanian Outer Carpathian foreland system, juxtaposed over the structural map of the autochthonous platforms (for details see Fig. 9). BF, Bistri~aFault; TF, Trotu~ Fault; LR, lateral ramp. Inset A shows a model of the late Miocene strike-slip escape above a late Cretaceous-Palaeogene lateral ramp. Inset B shows a model of Mid- to late Miocene deformations in the vicinity of the Trotu~ Fault (after Matenco et al. 1997). (d) A 3D flexural model for the east and south Carpathians foreland. Underlined numbers represent the predicted EET for each flexural block. Shaded zones represents simplified topography of the area. CSmax represents the predicted maximum deflection (after T~poanc~ et al. 2004).
MODELLING THE PANNONIAN-CARPATHIANS SYSTEM (Bertotti et al. 2003; Matenco et al. 2003), with limited to no depocentre migration along the east Carpathian foreland, in contrast to previous inferences (e.g. Meulenkamp et al. 1996). Despite the apparent continuity of sedimentary facies and tectonic units at the surface in the thin-skinned units of the east Carpathians, the analysis of kinematic thrust emplacement has revealed a non-cylindrical Miocene shortening, as a result of lateral variations in the strength of d~collement horizons and mechanical properties of the foreland lower plate (Matenco & Bertotti 2000). This non-cylindrical character refers to the amount of internal shortening (higher to the north) and to the variable character of the associated uplift and erosion. This is particularly important in the SE Carpathians, where a reduced amount of internal shortening during collision has failed to induce significant uplift (in any case below fission-track resolution; see Sanders et al. 1999) to induce significant erosion. Thermo-mechanical modelling (Cloetingh et al. 2004) suggests that collision with the Moesian platform in this sector of the belt locked the system in the early phases and led to thermal re-equilibration in the post-collisional phase, when the subducted slab started to stretch vertically. In contrast, continuation of the oceanic-type subduction in the central-northern domain means that a large part of the former passive margin, comprising thinned to stable continental crust of the lower plate, was underplated in the collision process. This induced significant internal shortening in the thin-skinned nappe pile at the contact and 4 - 5 km of uplift followed by erosion (Sanders et al. 1999). During the post-collisional phase in this central-northern part, the system became welded, the minor uplift observed being related to orogenic rebound. In contrast, the post-collisional phase of the Bend Zone is characterized by large subsidence in the foredeep, whereas a large amount of uplift is recorded towards the neighbouring western nappe pile with no orogenic-related shortening, probably in response to crustal folding (Bertotti et al. 2003). The drag of the vertical stretching slab accounts for a reduction in the orogenic load needed for the observed deflection in the foreland, as inferred also by 3D flexural modelling (Tfir~poancfi et al. 2004).
Lithospherie strength in the Pannonian- Carpathian system The mechanical behaviour of the lithosphere can be formulated in different theological models providing depth-dependent stressstrain relationships (see Ranalli 1995). The vertical strength distribution of the lithosphere is often represented by means of rheological profiles, or yield envelopes, in which the yield limit is plotted against depth. These profiles rely on the results of laboratory experiments extrapolated to different levels in the crust and mantle, and a number of assumptions regarding the thermal structure, strain rate, presence of fluids and several other factors. The depth dependence of these parameters results in a rheological stratification of the lithosphere. Accordingly, strain is mostly concentrated in the weak, mostly ductile levels of the lithosphere. These levels are good candidates for necking during extension or the formation of large-scale shear zones, detachment horizons, etc., and can lead to decoupled deformation between layers of higher strength. The Pannonian-Carpathian system shows a remarkable variation of thermo-mechanical properties of the lithosphere. Lithospheric rigidity varies in space and time, giving rise to important differences in the tectonic behaviour of different parts of the system. As theology controls the response of the lithosphere to stresses, and thus the formation and deformation of basins and orogens, the characterization of rheological properties and their temporal changes has been a major challenge to constrain and quantify tectonic models and scenarios. This is particularly valid for the Pannonian-Carpathian region, where tectonic units of different history and rheological properties are in close contact.
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Figure 5a displays three strength envelopes through the western, central and eastern part of the Pannonian lithosphere, constructed on the basis of extrapolation of rock mechanics data and incorporating constraints on crustal and lithospheric structure, and presentday heat flow along the modelled rheological section. These strength profiles show that the average strength of the Pannonian lithosphere is very low (see also Lankreijer 1998); this is mainly due to high heat flow, which is the consequence of an elevated asthenosphere dome below the basin system. The hottest basin in continental Europe has a lithosphere of extremely low rigidity, making it prone to repeated tectonic reactivation. This is the result of earlier phases of tectogenesis, namely nappe emplacement, crustal accretion and related orogenic thickening and destabilization during Cretaceous-Palaeogene times. At the same time, the strength of the Pannonian lithosphere segments gradually decreased, leading to the collapse of these orogenic terranes, a high level of strain localization and, eventually, the formation of the Pannonian Basin, that is rifting and extension of former compressional domains. Another essential feature is the noticeable absence of present-day lithospheric strength in the mantle lithosphere of the Pannonian Basin. Strength appears to be concentrated in the upper 7 - 1 2 km of the lithosphere. This finding is in very good agreement with the distribution of seismicity at depth. Earthquake hypocentres are restricted to uppermost crustal levels, suggesting that brittle deformation in the lithosphere is taking place not deeper than 5-15 km (T6th et al. 2002). Figure 5b shows estimates of the total integrated strength (TIS) of the Pannonian-Carpathian lithosphere along section A - A ' . Rheology calculations suggest major differences in the mechanical properties of different tectonic units within the system (Lankreijer et al. 1997, 1999). In general, there is a gradual increase of TIS from the centre of the basin towards the basin flanks in the peripheral areas (see also Fig. 5c). The centre of the Pannonian Basin and Carpathian foreland are the weakest and strongest parts of the system, respectively. The presence of a relatively strong lithosphere in the Transylvanian Basin is due to the differences in tectonic history with respect to the Pannonian Basin. The absence of large-scale Tertiary extension left this lithosphere segment relatively strong. The Carpathian arc, particularly its western parts, shows a high level of rigidity, with the exception of the southeastern Bend Zone, where a striking decrease of lithospheric strength is noticeable. Calculations for the seismically active Vrancea area indicate the presence of a very weak crust and mantle lithosphere, indicating mechanical decoupling between the Transylvanian Basin and the Carpathian orogen. The pronounced contrast in TIS between the Pannonian Basin (characterized by TIS < 2.0 x 1012 N m -1) and the Carpathian orogen and its foreland (characterized by TIS > 3.0 x 1012 N m -1) indicates that recent lithospheric deformation is more likely to be concentrated in the hot and hence weak Pannonian lithosphere than in the surrounding Carpathians. The results of rheology calculations are often expressed in terms of the effective elastic thickness (EET) of the continental lithosphere by integrating the thickness of the mechanically strong layers of the lithosphere (Burov & Diament 1995). EET values can also be obtained from forward modelling of extensional basin evolution and from flexural models of the lithosphere, which in turn can be compared with the inferences from rheological studies. This provides an important means to test the validity of different modelling techniques. By the conversion of strength predictions to EET values on a regional scale, Lankreijer (1998) constructed a map of EET distribution for the entire PannonianCarpathian system (Fig. 6). Calculated EET values are largely consistent with the spatial variation of lithospheric strength in the system. Lower values are characteristic for the weak central part of the Pannonian Basin (5-10 km), whereas EET increases towards the neighbouring orogens (15-30 kin) and, particularly, in the foreland areas in the Bohemian Massif and Moesian Platform (25-40 km). This trend is in good agreement with EET
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S. CLOETINGH E T A L .
Fig. 5. (a) Typical strength envelopes from the western (A), central (B) and eastern (C) part of the Pannonian Basin (for location see Fig. 6) constrained by data on lithospheric structure, heat flow, strain rate and stress regime (for details see Lankreijer et al. 1997; Sachsenhofer et al. 1997; Lenkey 2002). The nearly complete absence of mantle strength predicted by the model should be noted. (b) Total integrated lithospheric strength (TIS) along a regional profile through the Pannonian-Carpathian system (Lankreijer 1998). (c) Schematic cross-section showing the non-uniform stretching of the Pannonian lithosphere and its effect on depth-dependent rheology. The thickness of the crust (c) and mantle lithosphere (m) is reduced by the stretching factors 6 and/3, respectively, in the basin centre. Because ascending asthenosphere is heating the system, the isotherms become significantly elevated. As a result, the thinned and hot Pannonian lithosphere becomes extremely weak and, thus, prone to subsequent tectonic reactivation.
estimates obtained from flexural studies and forward modelling of extensional basin formation. Systematic differences, however, can exist. This may be the consequence of significant horizontal intraplate stresses (e.g. Cloetingh & Burov 1996) or the mechanical decoupling of the upper crust and uppermost mantle, which can lead to a considerable reduction of EET values. As lithospheric rheology is primarily controlled by temperature, a general relation is expected between the thermal age (time elapsed since the last thermal event) and the rheology of the lithosphere. This relation is straightforward for oceanic crust but is much more complex for continental areas. In Figure 7 estimated EET values and thermal age for selected locations in the Pannon i a n - C a r p a t h i a n system are plotted. In general, EET increases with thermal age, that is parallel to cooling of the lithosphere. Differences may exist, however, between situations when the crust and the mantle lithosphere behave in a coupled and a
decoupled manner. The presence of a weak lower crust can lead to the mechanical decoupling of the lithosphere, which may substantially change the response of the deforming plate. According to Burov & Diament (1995), EET follows the 3 0 0 - 4 0 0 ~ and 6 0 0 - 7 0 0 ~ isotherm in the case of a decoupled and coupled lithospheric configuration, respectively. In spite of significant variations in EET as a function of thermal age, shown in Figure 7, these estimates do not exceed the depth of the 600 ~ isotherm as predicted for the model of a cooling lithosphere (Cloetingh & Burov 1996). The range of calculated EET values reflects the distinct mechanical habitat and behaviour of different domains in the P a n n o n i a n - C a r p a t h i a n system, which is mainly controlled by the memory of the deforming lithosphere. The considerable differences in the tectonic and thermal history of these domains from Cretaceous-Palaeogene times (Alpine orogeny) and through the Neogene (extension in the Pannonian domain) resulted in a wide
MODELLING THE PANNONIAN-CARPATHIANSSYSTEM
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Fig. 6. Effective elastic thickness (EET, in km) of the lithosphere in and around the Pannonian Basin predicted from rheological calculations (Lankreijer 1998). A, B and C indicate the location of strength envelopes shown in Figure 5a. Regional strength profile A-A' is shown in Figure 5b. BM, Bohemian Massif; MP, Moesian Platform; PB, Pannonian Basin; EC, SC, WC, eastern, southern and western Carpathians, respectively.
Fig. 7. Thermal age of the lithosphere v. calculated effective elastic thickness (EET) for selected sites in the Pannonian-Carpathian system (Lankreijer 1998). DDR-EET and Flex-EET represent EET values based on depth-dependent rheology estimates and flexural models, respectively. Upper limits of DDR-EET boxes represent EET calculated for wet rheology; lower limits represent dry rheology. Isotherms (in ~ are calculated for a cooling half-space model taking radiogenic heat production into account (Cloetingh & Burov 1996).
spectrum of lithospheric strength within the system. This in turn strongly influences its present-day behaviour, leading to a complex pattern of continuing tectonic activity.
Deformation of the Pannonian-Carpathian system The present-day deformation pattern and related topography development in the Pannonian-Carpathian system is characterized by pronounced spatial and temporal variation of the stress and strain fields (Fig. 8). Horv~ith & Cloetingh (1996) established the
importance of Late Pliocene to Quaternary compression in the Pannonian Basin, explaining its anomalous uplift and subsidence as well as intraplate seismicity. Those workers established a novel conceptual model for structural reactivation in back-arc basins within orogens through the case study of the PannonianCarpathian system. At present, the basin has reached an advanced stage of its evolution with respect to other Mediterranean back-arc basins, and its structural inversion has been taking place during the last few millions of years. Basin inversion is related to temporal changes in the regional stress field, from one of tension that controlled Miocene basin formation, extension and subsidence, to one of Pliocene-Quaternary compression resulting in basin deformation, contraction and flexure of the lithosphere associated with differential vertical motions. The spatial distribution of uplifting and subsiding areas inside the Pannonian Basin can be therefore interpreted as a result of an increasing level of intraplate compressional stresses. The general inversion of the basin can be explained in terms of stress-induced lithospheric deflection, that is increasing intraplate stresses cause large-scale bending of the lithosphere at various scales. This includes basin-scale positive reactivation of Miocene normal faults, and large-scale folding of the system leading to differential uplift and subsidence at the anticlinal and synclinal segments of the Pannonian crust and lithosphere. Model calculations are in good agreement with the overall topography of the system (Fig. 8). Several flat-lying, low-altitude areas (e.g. the Great Hungarian Plain, and the Sava and Drava troughs) have been continuously subsiding since the onset of basin formation in the Early Miocene, and are filled with a sequence of Quaternary alluvial deposits of 300-1000 m in thickness. In contrast, the periphery of the basin system, as well as the Transdanubian Range, the Transylvanian Basin and the neighbouring Carpathian orogen, have been uplifting and considerably eroded since Miocene-Pliocene times (see Figs 3 and 9). Quantitative subsidence analysis by van Balen et al. (1999) confirmed that compressive stresses can cause accelerated subsidence in the central parts of the Pannonian Basin. Similar studies at the rim of the Pannonian Basin, including the Styrian Basin (Sachsenhofer et al. 1997), the Vienna and East Slovak basins (Lankreijer et al. 1995) and the Transylvanian Basin (Ciulavu et al. 2002), have demonstrated a major uplift of several hundred metres starting from Mio-Pliocene times. The mode and degree of coupling of the Carpathians with their foreland controls the Pliocene-Quaternary deformation
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Fig. 8. Topographyof the PannonianCarpathian system and present-day maximumhorizontal stress (Srt.....) trajectories (after Bada et al. 2001). Plus and minus signs mark areas of Quaternary uplift and subsidence,respectively.BA, Balkanides;BM, BohemianMassif; D, Drava trough; D, Dinarides;EA, Eastern Alps; EC, eastern Carpathians;F, Focsani depression; MP, Moesian Platform;PB, Pannonian Basin; S, Sava trough; SC, southern Carpathians; TB, Transylvanian Basin; TR, TransdanubianRange; WC, western Carpathians.
patterns in the hinterland, and particularly interesting, in the Transylvanian Basin (Ciulavu et al. 2002). The western-eastern Carpathians is coupled with the strong lid of the European lithosphere, showing comparable coeval deformations in both the upper and the lower plate (Matenco & Bertotti 2000; Krzywiec 2001). The Bend Zone is decoupled in terms of deformation from its Moesian lower plate. It records contraction in the upper plate (e.g. Hippolite et al. 1999) and large-scale extensional collapse in the Moesian unit, as demonstrated by structural studies and crustal earthquake focal mechanisms (e.g. Bala et al. 2003). These findings have been recently corroborated by results from fission-track studies in the Romanian Carpathians, demonstrating up to 5 km of erosion with a systematic migration from the northwestern and southwestern part of the Romanian Carpathians, uplifted and eroding since 12 Ma, towards the bend area where uplift and erosion was initiated around 4 Ma ago (Sanders et al. 1999; Fig. 9). This region coincides with the actively deforming Vrancea zone in Romania, where considerable seismic activity is observed both at crustal levels and in the mantle. These findings can now be connected with seismic tomography results highlighting the development of hot mantle material under the Pannonian Basin and the presence of late-stage detachment of the lithosphere in the Vrancea area (Wortel & Spakman 2000). At the same time, these rapid differential motions at the rim of the Pannonian Basin and the Carpathians have important implications for the sediment supply to the depocentres as well as for the hydrocarbon habitat (Dicea 1996; Taft et al. 1997; Horv~ith & Tari 1999). In summary, the results of forward basin modelling show that an increase in the level of compressive tectonic stress during Pliocene-Quaternary times can explain the first-order features of the observed pattern of accelerated subsidence in the centre of the Pannonian Basin and uplift of the basin flanks in the peripheral areas. Therefore both observations (see Horvfith e t al. 2006) and modelling results lead to the conclusion that compressive stresses can cause considerable differential vertical motions across the back-arc basin-orogen system in the Pannonian-Carpathian area.
The sources of compression in the context of basin inversion were investigated by means of finite-element modelling (Bada et al. 1998, 2001). The results suggest that the state of recent stress in the Pannonian-Carpathian system (Fig. 8), particularly in its western part, is controlled by the interplay of plate boundary and intra-plate forces. The former includes the counter-clockwise rotation and northward indentation of the Adriatic microplate against the Alpine-Dinaric orogen, whereas intra-plate buoyancy forces are associated with the elevated topography and related crustal thickness variation of the Alpine-Carpathian-Dinaric belt. Model predictions indicate that uplifted regions surrounding the basin system can exert compression on the thinned Pannonian lithosphere of about 4 0 60 MPa, which is comparable with values calculated for far-field tectonic stresses (Bada et al. 2001). The combined analysis of stress sources of tectonic and gravitational origin has provided estimates for the magnitude of maximum horizontal compression. Significant compressional stresses (up to 100 MPa) are concentrated in the elastic core of the lithosphere, consistent with the continuing structural inversion of the Pannonian Basin. Such high-level stresses are close to the integrated strength of the system, which may lead to its whole-lithosphere failure in the form of large-scale folding and related differential vertical motions, and intense brittle faulting in the form of seismoactive faulting.
Discussion and conclusions Previous studies of the Pannonian Basin-Carpathian system underlined the importance of crustal stretching and lithospheric flexure as controls on the main features of basin formation and present-day lithospheric structure. These studies have revealed major discrepancies between subcrustal and crustal stretching factors inferred from subsidence analyses and seismic reflection data, and predictions from uniform stretching models. The same is true for the late-stage (Pliocene-Quaternary) anomalous acceleration in subsidence and uplift in the Pannonian Basin and
MODELLING THE PANNONIAN- CARPATHIANS SYSTEM
Fig. 9. Contours of amount of erosion (km) inferred from fission-track analyses of samples from the Romanian Carpathians and Apuseni Mts (Sanders et al. 1999) and thickness of foredeep sediments in the foreland. Numbers in elliptical boxes indicate tinting of onset of erosion (Ma); square boxes indicate the main moment of subsidence. The pronounced lateral differences in uplift ages along the arc should be noted; in contrast, the main subsidence period is coeval along the studied area.
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surrounding Carpathian arc. This pattern of vertical motions deviates from scenarios of decaying thermal subsidence and downflexure of the lithosphere inferred from stretching models and foreland flexural models, respectively. Lower crustal flow and small-scale lithosphere-asthenosphere convection appears to be a viable mechanism to explain the occurrence of a second, Late Miocene rift phase in the Pannonian Basin coeval with the climax of compression in the surrounding Carpathian arc. The Pannonian B a s i n - C a r p a t h i a n system displays pronounced lateral variation in lithospheric structure and sedimentary basin configuration. The tectonic evolution of the system is characterized by a polyphase history with Early to Late Miocene extension in the Pannonian Basin and simultaneous contraction in the Carpathian arc. Pre-existing structures and pre-rift rheology of the lithosphere play a key role in basin formation and subsequent reactivation, explaining anomalous features in subsidence characteristics and inferred thinning factors. The P a n n o n i a n - C a r p a t h i a n system has been subjected to repeated tectonic reactivation and, thus, a high level of strain concentration resulting in an overall weakening of the deforming lithosphere. Cretaceous to Palaeogene convergence in the Alpine belt led to nappe stacking and crustal accretion, and related thickening of the orogen. The overthickened and unstable Pannonian lithosphere, with a laterally unconstrained plate boundary towards the Carpathian embayment in the east, underwent gravitational collapse and extrusion and, eventually, the Miocene formation of the Pannonian Basin by the rifting and extension of former contractional domains. Models for the Pannonian Basin therefore underscored the importance of the pre-rift lithospheric structure for the mode of extension in the basin system. Lithospheric memory is important on different spatial scales. On a lithospheric scale, the presence of a thickened continental Alpine crust was manifested in shallow necking levels during subsequent extension (van Balen e t al. 1999). On an upper crustal scale, reactivation of Alpine thrust-faults in basin extension has been widely documented at
Fig. 10. Example of crustal-scale folding in south Transdanubia, Hungary, based on the interpretation and sequence stratigraphic analysis of regional seismic reflection profiles (Sacchi & Horvfith 2002). During the Quaternary the whole area has been uplifting. The basement units of the Bakony and Mecsek Mts, however, represent regions of higher amounts of uplift and, thus, large-scale anticlines, whereas the folded Miocene strata in the Somogy can be regarded as a syncline between the S.S., sensu stricto; S.L., sensu lato.
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Fig. 11. Example of lithosphere-scale folding in the Pannonian Basin system documented by stratigraphic, geophysical and geodetic data. (a) Crustal profile A-B transects the whole Pannonian Basin from the Eastern Alps to the eastern Carpathians, connecting uplifting and subsiding areas. 1, Areas of Quaternary subsidence; 2, areas without Quaternary subsidence or uplift; 3, areas of Quaternary uplift; 4, NeogeneQuaternary volcanic rocks; 5, internal basement units; 6, flysch nappes; 7, AlpineCarpathian foredeep; 8, Tauern-window; 9, direction of present-day maximumhorizontal stress (SHmax).Adr, Adria; AM, Apuseni Mts; D, Drava trough; Da Danube Basin; DIN, Dinarides; EA, Eastern Alps; EC, eastern Carpathians; EEP, East European Platform; GHP, Great Hungarian Plain; S, Sava trough; SA, Southern Alps; SC, southern Carpathians; TB, Transylvanian Basin; TR, Transdanubian Range. (b) Stratigraphic cross-section along transect AD. (e) Basement deflection profile based on forward modelling of stress-induced subsidence-uplift pattern. Compiled after Jo6 (1992), Horvfith (1993), Horvfith & Cloetingh (1996) and Gerner et al. (1999).
both seismic and outcrop scale in tectonic windows at the margins of the Pannonian Basin (e.g. Horvfith 1993). Recently, independent constraints on the bulk rheology of the pre-rift Alpine lithosphere have been obtained from rheological modelling (e.g. Sachsenhofer et al. 1997; Lankreijer 1998; Willingshofer et al. 1999). A marked contrast in recent rheology between the Pannonian Basin area, the surrounding Carpathian orogen and the foreland lithosphere is directly related to crustal configuration and thermal properties. In general, the Bohemian Massif, the East European Platform and, to a lesser extent, the Moesian Platform form strong and rigid buttresses in the foreland areas, whereas the intra-Carpathian regions are characterized by a much weaker rheology. Lateral and temporal variations in lithospheric strength exert a major control on the tectonic evolution of the study area. The extremely low rigidity of the Pannonian Basin and its tectonic setting locked in the interior of the Carpathians arc has made it a sensitive recorder of stress changes induced by various tectonic processes. Stress studies have underlined the close
relationship between the timing and nature of the stress changes in the extensional basins and the timing inferred from kinematic studies of the surrounding thrust belts. This suggests a mechanical coupling between the orogen and the back-arc domain. Quantitative modelling of recent and palaeostress fields constrained by a vast database of stress indicators has demonstrated that the stress fields in the area are primarily controlled by near-field intraplate and far-field plate boundary forces (Bada et al. 1998) in the context of continuing continental collision in the Alpine-Dinaric belt. The strain and stress pattern are also strongly influenced by the effects of topography and crustal thickness variation (Bada et al. 2001), and by asthenospheric ascent (Huismans et al. 2001). The recent tectonic activity of the region is largely controlled by the counter-clockwise rotation of the Adriatic microplate relative to Europe around an Euler pole in the Alps. As a result of the indentation of this crustal block against the Southern Alpine-Dinaric fold-and-thrust belt, intense oblique shortening (dextral transpression) is taking place in these orogens, as shown
MODELLING THE PANNONIAN-CARPATHIANS SYSTEM
by the general seismicity and crustal deformation pattern. The present-day kinematics of the Pannonian Basin shows that the area is pushed from the SSW by the Adriatic microplate and compressional stresses propagate well into the interior of the basin system. As a result, strike-slip to reverse faulting is observed inside the Pannonian Basin. Furthermore, the nearly complete absence of normal faulting throughout the study area suggests that extension in the Pannonian Basin has been terminated and its positive structural inversion is in progress. The intra-plate seismicity and neotectonic deformation indicate that the Pannonian Basin is affected by strong tectonic coupling with the surrounding parts of the African-European collision system. It appears that the polyphase evolution of the Pannonian Carpathian system has resulted in strong lateral variation in thermo-mechanical properties in the area, with weak lithosphere in the Pannonian Basin making it prone to late-stage basin reactivation. A particular feature of the Pannonian Basin is the build-up of a compressional stress regime resulting in a gradual inversion of the basin system. Tectonic inversion is manifested in a relatively high level of seismicity, abundant fault reactivation, late-stage folding and erosion documented by high-resolution reflection seismic data collected in Hungarian rivers and lakes (e.g. T6th & Horvfith 1998; Sacchi et al. 1999). Crustal deformation in the form of folding can take place at several scales from the inversion of a half-graben (e.g. in the Sava folds in Slovenia and SW Hungary; see Horvfith et al. 2006) to crustal-scale buckling. The area of south Transdanubia in Hungary is a key example for upper crustal folding (Fig. 10). Sacchi & Horvfith (2002) found that although the area has been uplifted as a whole during the Quaternary, differential vertical motions indicate continuing flexure, with maximum uplift taking place in the Bakony and Mecsek Mrs. The wavelength and amplitude of crustal folding can be estimated at about 1 0 0 - 1 5 0 k m and 500-1000 m, respectively (Fig. 10). On the other hand, large-scale bending of the Pannonian lithosphere, manifested in the Quaternary subsidence and uplift history (Horvfith & Cloetingh 1996), is also characteristic. As such, the Pannonian Basin has been interpreted in terms of irregular lithosphere folding (Cloetingh et al. 1999; Fig. 11) with a spectrum of wavelengths from several hundreds of kilometres (lithosphere-scale folding) through a few tens of kilometres (crustal-scale folding) to a few kilometres (basin-scale inversion). Folding and related structural development of the Pannonian lithosphere and related differential vertical movements have several important consequences for the hydrocarbon habitat (e.g. sealing of fault systems, reservoir integrity, maturation history and hydrodynamic regime, overpressure zones) and environmental issues (including landscape and slope stability). Recently, much attention has been paid to the spatial and temporal variations in thrusting along the Carpathian arc and its relationship to migrating depocentres, foreland basin geometry and lateral variations in fiexural rigidity (Matenco et al. 1997b; Zoetemeijer et al. 1999). At the same time, it has become increasingly evident that the second rift phase of the Pannonian Basin occurred simultaneously with the climax of compression in the Carpathian arc (Huismans et al. 1999), suggesting a mechanical link in terms of lower crustal flow induced by the rifting process, directed towards the Carpathian orogen. A general feature of all the flexural modelling studies carried out for the Carpathian system is the relatively low rigidity of the platform lithosphere downbending under the Carpathian belt, with effective elastic thickness (EET) estimates consistently below predictions inferred from rheological models of corresponding thermotectonic ages (Cloetingh & Burov 1996). These low EET values may be partly the result of stress-induced weakening associated with steep bending of the platform lithosphere under the arc, and partly the result of inherited weaknesses related to pre-orogenic Mesozoic extensional faults (Zoetemeijer et al. 1999), and the reactivation of inherited deep-seated
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weakness zones may have also played a key role (Matenco et al. 1997b; Lankreijer 1997). This paper reflects concepts developed during a decade-long scientific cooperation between the Vrije Universiteit Amsterdam, the Netherlands, the E6tv6s University, Budapest, Hungary and the University of Bucharest, Romania. Financial support, received from the Netherlands Research Centre for Integrated Solid Earth Science (ISES), the Hungarian National Science Fund (OTKA) projects T034928, D34598 and F043715, and the Romanian Ministry for Education and Research are acknowledged. The International Lithosphere Programme is thanked for providing the framework for scientific collaboration and discussions.
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Ranges and basins in the Iberian Peninsula: their contribution to the present topography JAUME VERGt~S & MANEL FERNANDEZ
Group of Dynamics of the Lithosphere (GDL), Institute of Earth Sciences 'Jaume Almera', CSIC, 08028 Barcelona, Spain (e-mail: jverges @ ija. csic. es)
Abstract: The Iberian Peninsula,at the western end of the Alpine-Himalayan Belt, displaysa complexstructure with mountainranges of diverse structural trends and sedimentarybasins between them. The Iberian Peninsulaalso shows an elevated mean topography, the highest in Europe. In this short paper, we investigatethe Alpine evolution of the Iberian Peninsula since Mesozoic times, when Iberia was isolated as an independentplate. This occurred from Albian (formationof the northern plate boundary) to Oligocenetimes (end of the PyreneanOrogeny). Iberia was squeezedbetween Africa and Europe during Tertiary times and all previouslyestablishedMesozoic extensional basins were inverted, as were some of the Hercynian structures. The opening of the Valencia Trough, cutting the eastern margin of the Iberian Peninsula, began in Oligocene times. Concomitant crustal and lithospheric stretching during the Neogene along the eastern margin of Iberia produced limited uplifts, some of which are still active. The modern topography of the Iberian Peninsula was developedmainly as the result of three main tectonothermalmechanismssince late Palaeozoictimes: variationsin crustal densities, and possibly mantle depletion, inherited from the Hercynian Orogeny; crustal and lithospheric thickening during Tertiary compression; and upper mantle thinning during the Neogene-Quaternary.
The Iberian Peninsula constitutes the westernmost segment of the 12 000 km long Alpine-Himalayan Belt formed as a result of the Tertiary closure of the Tethys Ocean during the collision of India, Arabia and Africa with Asia and Europe (e.g. Dercourt et al. 1986). Rifting initiated during Triassic times (c. 250 Ma) and culminated in crustal break-up along the Atlantic margin (e.g. Ziegler, 1988, 1992). Continental break-up of the African Plate occurred during the Late Jurassic (c. 156 Ma). The Atlantic Ocean propagated northwards through the proto-Azores-Gibraltar plate boundary, producing the continental rupture of the Iberian Plate in Early Cretaceous times (c. 118 Ma; e.g. Srivastava et al. 1990). The mid-Cretaceous northern boundary of the Iberian Plate formed along the oceanic lithosphere of the Atlantic Ocean and its eastern continuation along the continental lithosphere of the Pyrenees (Fig. 1). Towards the end of Late Cretaceous (chron 33, 80 Ma) Africa shifted its motion northwards, initiating convergence with Eurasia with the consumption of the Tethys Ocean (e.g. Dercourt et al. 1986). At the westernmost termination of the AlpineHimalayan Belt, the Iberian Plate underwent a protracted deformation phase, resulting in orogenic belts along the plate boundaries (Bay of Biscay-Pyrenees and Azores-Gibraltar) and severe intraplate deformation. Several large Tertiary sedimentary basins developed on the Iberian Plate close to the bounding mountain chains (Friend & Dabrio 1996). Most of these basins began as flexural basins and continued as intermontane basins during the growth of the complex Iberian mountain system (Fig. 1). Iberia initially moved together with the African Plate, from latest Cretaceous to mid-Eocene times (chron 19, 42 Ma), deforming mainly the Bay of Biscay-Pyrenees plate boundary. From midEocene to the end of Oligocene times (chron 6c, 24 Ma), it moved independently and both plate boundaries were active. Subsequently, during the last 24 Ma, most of the deformation was accommodated along the complex and poorly understood plate boundary between Iberia and Africa, leading to the formation of the Betics, the Gibraltar Arc, and the Rif. The end of the Oligocene also coincided with extension along the proto-Western Mediterranean Sea, which affected the entire eastern margin of the Iberian Plate. This extension formed the oceanic lithosphere below the Liguro-Proven~al Basin north of the Paul Fallot Fault. To the south of this fault, thinned lithosphere below the Valencia Trough and Alboran Sea and oceanic lithosphere below the Algeria Basin formed (Fig. 1). The present contact
between Africa and Iberia changes progressively from pure fightlateral strike-slip along the Gloria Fault to a diffuse transpressive boundary from the Gorringe Bank to the Gulf of Cadiz region (e.g. Argus et al. 1989). The present structure of the Iberian Peninsula developed through the interplay of several geodynamic processes related to the Atlantic opening, the formation of two plate boundaries limiting the Iberian Plate, the north- south Africa-Europe convergence, and the concomitant rapid retreat and consumption of the oceanic Tethyan realms. Different geodynamic processes related to these large-scale tectonic events were to some extent coeval over particular morphotectonic regions. Both the diversity of geodynamic processes and their potential conjunction complicate the unravelling of the evolution of the Iberian Plate in general and, in particular, the southern plate boundary between Iberia and Africa (Betic Cordillera, Rif, Alboran Sea and Gulf of Cadiz tectonic units). This paper documents in brief the Alpine evolution of the onshore Iberian Peninsula mountain ranges and sedimentary basins, emphasizing the geodynamic processes that created positive topographic relief. This evolution took involved the following major tectonic events: (1) formation of extensional Mesozoic basins at the intersection of the proto-Atlantic and the Tethys oceans; (2) generation of Late Cretaceous-Tertiary fold-and-thrust belts and basins by the northwards motion of Africa; (3) formation of basins by Neogene extension along the eastern margin of the proto-Western Mediterranean. The paper concludes with the present topographic configuration of the Iberian Peninsula and its heritage from Hercynian times including the relatively recent lithosphefic thinning along the Mediterranean province of Spain. Two recently published books on the geology of Spain give a detailed description of the mountains and basins documented in this brief paper (Gibbons & Moreno 2002; Vera 2004). Andeweg (2002) has also illustrated the evolution of the palaeostress field in the Iberian Peninsula through the Cenozoic.
Mesozoic extensional basins
Preceding the opening of the central Atlantic during mid-Jurassic times (chron BSMA at c. 170 Ma), the Iberian Peninsula (Iberian Plate) was deformed by large-scale stretching that resulted in numerous extensional basins with different orientations. Rift systems developed along the western margin of the Iberian Plate
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 223-234. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. Map of Western Europe with location of principal orogenic chains related to the Africa-Europe collision (based on Verg6s & Sfibat 1999). The Pyrenean Range corresponds to the westernmost limit of the about 12 000 km long Alpine-Himalayan Belt. CCR, Catalan Coastal Ranges; CM, Cantabrian Mountains; P.F.F., Paul Fallot Fault.
The separation between Europe and Africa for this period is about 240 km along this transect (e.g. Boccaletti et aL 1977; Olivet 1996; see position B for Africa in Fig. 2). The proposed position of Africa provides very little room for the restored Betic domain, thus creating a significant space problem, which has already been recognized (e.g. Andrieux et al. 1971; Mauffret et aL 1989; Frizon de Lamotte et al. 5995; Lonergan & White 5997; Spakman & Worte12000). Andrieux et al. (1971) proposed a model, still used with modifications by a number of workers, in which the Alboran Block was displaced towards the west by lateral extrusion during the north- south convergence of Africa and Iberia. Crustal and lithospheric thinning was the common process that formed the extensional basins. The thinned regions constituted weaker zones at the end of the Cretaceous just before the onset of Tertiary compression. Most of these extensional basins were tectonically inverted, preserving their original basin orientation. An extensive distribution of Triassic evaporites controls the geometry of the thrust system in both previous tectonic basins and structural highs, as in the case of the southern end of the Iberian Chain along the Altomira thrust system (the western boundary of the Iberian Range; Fig. 3).
Alpine Orogeny: Tertiary compressive belts and sedimentary basins (offshore Galicia and Portugal; e.g. Malod & Mauffret 1990), within the continental plate (Iberian and Catalan rifts; e.g. Salas et al. 2005), and along the two plate boundaries (Pyrenean rift in the north and the rifted south Iberian margin; Fig. 2). The opening of the central Atlantic produced an eastwards motion of Africa along the former Azores-Gibraltar transform fault. The Alpine-Tethys Ocean opened along the eastern margin of the Iberian Plate (e.g. Stampfli et al. 2002). This opening led to the cessation of rifting processes occurring within the Iberian Plate at the end of the Late Jurassic. A renewed phase of extension took place before the opening of the North Atlantic during the early Aptian (chron M0 at about 118 Ma). Pyrenean rift events occurred before the onset of ocean formation in the Bay of Biscay during Aptian-Albian times. The northern Iberian Plate transform boundary propagated eastwards and the northern segment of the Alpine-Tethys formed along the southeastern margin of Western Europe (Stampfli et al. 2002). These processes ended by the end of Coniacian at about 86 Ma. The Iberian rift system formed a linked configuration of extensional basins with different orientations as observed in the NE corner of Spain at the eastern end of the east-west-trending Pyrenees, the NW-SE-trending Figueres-Montgrf branch (F-M in Fig. 2), and the NE-SW-directed Catalan Basin (Fig. 2). The Iberian Peninsula at the end of Cretaceous times (Fig. 2), before the onset of Africa-Europe convergence, shows extended regions within the plate as well as along its margins. It is interesting to note that, during this period, at the end of Mesozoic extension, there are about 125-150km of separation between the central and eastern sides of France and Spain, and at least 35 km of extension in the Iberian Basin (Salas & Casas 1993; Salas et al. 2001). The restoration of the Prebetic and Subbetic units shows that their former, common, southeastern boundary was at least 90 km to the SSE of its present position (restoration according to GarcfaHernfindez et al. 1980). If we add the Internal Betics to the reconstruction by unfolding them to a minimum of double their present width (shortening of 50%), the SE boundary of the Internal Betics restores to about 250 km to the SE of its present position. Adding a counter-clockwise rotation of about 25 ~ to fit reasonably well with the palaeomagnetic rotations observed in the Betics (Platt et al. 2003) before the end of the late Tortonian (Krijgsman & Garc6s 2004), then the SE border of the Internal Betics restores to about 300 km to the SE (Fig. 2). This restoration is approximately in agreement with the proposed restoration by Platt et al. (2003).
In Late Cretaceous times, the northern motion of Africa against Europe significantly deformed the Iberian Plate along the previously extended Mesozoic basins. These basins, of diverse orientations and sizes, developed along both the plate boundaries and within the interior of the plate. Along the northern Iberian Plate margin, the Pyrenees represent a continental collisional orogeny with limited northwards subduction, whereas a more complex region deformed in the Southern Iberian Plate, including the Betics and Rif, the Alboran Sea, and the Gulf of Cadiz. The western Iberian margin along the Portuguese coast represents a slightly inverted margin, especially in its southern segment (e.g. Alves et al. 2003; Zitellini et al. 2004). During the roughly north-south Alpine convergence the interior of the Iberian Peninsula deformed while mostly preserving the original trends of the previously extended basins: N E - S W in the Catalan Coastal Ranges, N W - S E in the Iberian Range, N E - S W in the Central System, and north-south in the Altomira Range (Mufioz Martfn & De Vicente 1998; Fig. 3). All these fold-and-thrust systems are connected and at the intersections of any two of them there is always a linking zone in which the two different trends coexist as well as intermediary trends. The linkages between all the compressive fold-and-thrust belts indicate the synchronicity, at least partially, of several of these deformational events (Fig. 3). These connections between the different thrust systems that shape the present Iberian Peninsula have been described since the early 5980s in several transects crossing the Iberian Peninsula. The linkage of different thrust systems and the partial synchronicity of the deformation are agreed upon by most workers (e.g. Guimerfi 1984; Banks & Warburton 5995; Anad6n & Roca 1996; Casas S ainz & Faccenna 2005). The remarkable repetition of N E - S W trending chains in the central part of the Iberian Peninsula was interpreted as being produced by lithospheric folding during the Neogene (Cloetingh et al. 2002), in contrast to crustal and lithospheric thickening. In the following sections, we describe the principal compressive mountain systems of the Iberian Peninsula and their associated sedimentary basins starting in the Pyrenees (see regional transect of Roca et al. 2004) and ending in the Betics (see regional transect of Frizon de Lamotte et al. 2004). Between these, all the smaller mountain ranges developed in an interior tectonic setting, within the Iberian Plate. Casas Sainz & Faccenna (2001) have published an overview of these compressive mountain chains.
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225
Fig. 2. Reconstructed map at the end of Late Cretaceous times to show the distribution of principal sedimentary basins in and around the Iberian Peninsula. Ca, Cameros Basin; Ma, Maestrat Basin; Col, Columbretes Basin; F-M, Figueres-Montgrf Basin; NPFZ, North Pyrenean Fault Zone. Position and extent of Columbretes Basin after Roca (1996). Sardinia is shown in its restored position prior to Neogene opening of the Gulf of Lyons (Olivet 1996). The Balearic Islands are shown in their restored position prior to Neogene opening of the Valencia Trough (Vergrs & S~bat, 1999). Position A of Africa corresponds to the present position and position B to the restored position (e.g. Boccaletti et al. 1977; Olivet 1996). Pyrenees
The partial subduction to the north of the Iberian lithosphere underneath the European one along the Iberia-Europe plate boundary shaped the large-scale Pyrenean fold-and-thrust belt (Choukroune & Team, 1989; Roure et al. 1989; Mufioz 1992; Beaumont et al. 2000). The Pyrenean orogen is asymmetrical and double-sided, with the most significant thrust system developed towards the south, on top of the subducted zone as in most orogenic belts (e.g. Capote et al. 2002; Verg6s et al. 2002; Fig. 3). Although Africa compressed the entire Iberian Plate during its northwards shift, the Pyrenees recorded the initial stages of generalized compression around Santonian times (e.g. Puigdef~bregas & Souquet 1986), about 35 million years after the end of major rifting events along the Pyrenean branch of the North Atlantic in Albian times. The geometry at different scales of the northern boundary of the Iberian Plate clearly controlled the structural evolution of
the compressive belt. The northwards subduction was initiated along the extremely thinned lithosphere of the North Pyrenean Fault Zone (Mufioz 1992). The northern and southern thrust systems tectonically inverted the earlier Mesozoic basins together with their extensional fault systems. The location and extent of the early Mesozoic (Late Triassic) evaporites also exerted an important influence on the geometry and extent of the fold-and-thrust belts in the Iberian Peninsula. Most of the important Pyrenean shortening processes lasted for about 40 million years and were partitioned along several thrusts describing an overall foreland-directed propagation of deformation. Maximum shortening occurred across the Central Pyrenees and decreased towards the west (e.g. Vergrs et al. 2002). The northwards subduction of Iberia below Europe continued towards the western Pyrenees (Teixell 1998) and the Cantabrian Mountains (Fern~indez-Viejo et al. 2000; Pedreira et al. 2003). However, the degree of shortening decreased towards the west (e.g. Teixell 1998) and the age of initial deformation was younger in the same direction (e.g. Vergrs et al. 2002).
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Fig. 3. Tectonic map of the Iberian Peninsulabased on Rodrfguez-Fern~ndez (2004) and the detailed tectonic map of the Pyrenees of Vergrs et al. (1995). Tertiary mountain ranges: Pyrenees, Cantabrian Mountains (CM), Catalan Coastal Ranges (CCR), Iberian Range, Central System (CS), and Betics. Foreland basins: Ebro, Duero, Tajo, and Guadalquivir (GB). Iberian Massif: South Portuguese Zone, Ossa Morena Zone, Central Iberian Zone, AsturianLeonese Zone (ALZ), and CantabrianZone (CZ). AP, As Pontes Basin, located in NW Spain; CR, Ciudad Rodrigo Basin, located in the SW side of the Duero Basin. A, Altomira fold system, forming the western margin of the IberianRange.
The South Pyrenean flexural foreland basin was underfilled and marine from 55 to 37 Ma and then became overfilled and continental, starting with the deposition of mid-late Eocene Cardona evaporites until the end of shortening during the Oligocene (Puigdeffibregas & Souquet 1986; Puigdeffibregas et al. 1992; Vergrs et al. 1995, 1998). At m i d - l a t e Eocene times (c. 37 Ma), uplift of the western Pyrenees triggered the end of the foreland basin stage and originated an intermontane basin bounded by the Pyrenees, the Catalan Coastal Ranges and the Iberian Range (e.g. Burbank et al. 1992; Garcfa-Castellanos et al. 2003). A long period of lacustrine deposition that lasted through the Oligocene and most of the Miocene characterized this endorheic period (Riba et al. 1983). An internal fluvial network delivered sediments to the Ebro Basin, which was characterized by a large central lake (e.g. Anad6n et al. 1979; Arenas & Pardo 1999). The end of deformation occurred during late Oligocene times (c. 24.7 Ma; Meigs et al. 1996) although major basement uplift, based on fission-track cooling ages, ended at about 30 Ma (Fitzgerald et al. 1999). The late Oligocene-early Miocene age of younger compression in the western Pyrenees was synchronous with extensional processes affecting the eastern Pyrenees related to the formation of the Western Mediterranean basins. During late Miocene times, the endorheic Ebro fluvial system opened towards the Mediterranean Sea (e.g. Coney et al. 1996; Garcfa-Castellanos et al. 2003).
Roca & Guimerfi 1992). This thick-skinned style of tectonics affected the entire crust (Sfibat et al. 1997; Roca et al. 2004) producing a frontal monocline with relatively high topography (L6pez-Blanco et al. 2001). The oblique position of these basins with respect to the direction of compression also produced a limited sinistral strike-slip component (e.g. Anad6n et al. 1985; Guimerfi & Alvaro 1990).
Iberian R a n g e
The Iberian Range, with a N W - S E trend (nearly orthogonal to the Catalan coastal system), shows a complex structure involving cover and basement units (Alvaro et al. 1979; Casas Sainz & Faccenna, 2001; Fig. 3). The thrust system along the chain shows double vergence corresponding to the inversion of previous Mesozoic extensional faults. The NW termination of the range plunges beneath the Duero Basin and continues at depth below the frontal thrust of the Cantabrian Mountains. The northern tectonic displacement of the Iberian Range in its northwestern terminus is about 30kin (Casas Sainz 1993; Guimerfi et al. 1995). The Sierra de Altomira on the southwestern side of the Iberian Range flanks the Tajo Basin to the east (Fig. 3). This almost north-south-trending fold-and-thrust system is detached above Triassic evaporites (Mufioz Marffn & De Vicente 1998).
Catalan Coastal R a n g e s Central S y s t e m
The Catalan Coastal Ranges have a N E - S W trend (AlpineTethys Ocean trend) and display the effects of multiple tectonic events, including Eocene compression and Oligocene-Miocene extension, the latter related to the opening of the Valencia Trough (e.g. Anad6n et al. 1985; Fig. 3). The Catalan Coastal Ranges represent the inversion of earlier extensional structures, with the inversion affecting both cover and basement units (e.g.
The Central System, with a N E - S W structural trend, comprises an uplifted block, like a large pop-up structure (Vegas et al. 1990), with relatively high topography and associated with crustal thickening (Mezcua et al. 1996; Fig. 3). This block bounds two large sedimentary basins in the central part of the Iberian Peninsula: the Duero Basin to the NW and the Tajo Basin to the SE. The
BASINS & RANGES, IBERIAN PENINSULA Central System thrusts the Duero and Tajo basins to the NW and SE, respectively (e.g. Querol 1989; De Vicente et al. 1996). The Duero Basin is filled by a maximum of 2.5 km of Oligocene and Miocene continental deposits, mostly from the Cantabrian Mountains (e.g. Alonso Gavilfin et al. 2004). The age of the younger sediments in the basin is late Miocene at about 9.6 Ma (Krijgsman et al. 1996). The Ciudad Rodrigo Basin, at the SW termination of the Duero Basin, opened to the Atlantic at the Oligocene-Miocene boundary (Santisteban et al. 1996; its location is shown in Fig. 3). Subsequently, after 9.6 Ma, the Duero River captured the closed Duero Basin and opened it to the Atlantic basin. The Tajo Basin filled with 2 - 3 km of continental deposits ranging in age from the latest Oligocene to the latest Miocene (e.g. Alonso Zarza et al. 2004). Other compressive areas
Towards the NW corner of the Iberian Peninsula, small transpressional basins (such as the As Pontes basin) were filled with alluvial to lacustrine deposits during late Oligocene and earliest Miocene times (e.g. Cabrera et al. 1996; Fig. 3). The western boundary of the Duero Basin also formed during this period, closing it. Betics
The Betic Cordillera, trending generally ENE-WSW, corresponds to part of the former northern Africa-Iberia plate boundary, developed on top of Iberian crust and cropping out now in the southern part of the Iberian Peninsula (Figs 1 and 3). Although some aspects of the evolution of the Betic Cordillera are well known, there is still no agreement about the mechanisms that created it (see discussion of models by Calvert et al. 2000). The Betic Cordillera is divided into the Internal Betics, comprising metamorphic basement, the External Betics, consisting of cover rocks, and the Guadalquivir Foreland Basin (e.g. Azafion et al. 2002). The Alboran Sea, with a complex tectonic history, formed as a result of Neogene extension. Neogene to Quaternary depositional sequences fill curved, elongated and deep basins (e.g. Comas et al. 1999). The Internal Betics comprise a tectonic pile of three different tectonic units separated by thrusts (Nevado-Filfibride at the base, Alpujfirride in the middle, and Malfiguide at the top). Each unit displays a different degree of Alpine metamorphism, decreasing from the bottom to the top. The Nevado-Filfibride unit was affected by H P - L T metamorphism indicating that the rocks were located at depths of 50-70 km (the metamorphic evolution of the Internal Betics has been described by Comas et al. (1992)). The metamorphic units crop out as large antiforms that exhibit east-west trends in contrast to the E N E - W S W regional direction of the External Betics. Towards the western end of the Internal Betics, relatively large massifs of peridotites have been incorporated into the Alpujfirride thrust system (e.g. Ronda Peridotites; Tubfa et aL 1997). The External Betics constitute a system of thrust sheets carrying different Mesozoic SE Iberian passive margin palaeogeographical units towards the foreland. The Subbetic Zone is located to the SE and the Prebetic Zone to the NW (Garcfa-Hernfindez et al. 1980). The stratigraphy of the External Betic units includes Triassic to Miocene successions. The Prebetic Zone is mostly composed of shallow-water deposits whereas the Subbetic Zone units are deeper and pelagic. The Campo de Gibraltar Unit consists of terrigenous deposits forming an Oligocene and early Miocene accretionary prism developed above the Subbetic Zone towards the WSW (e.g. Crespo-Blanc & Campos, 2001; Bonardi et al. 2003) that was actively deformed until late Tortonian times (e.g. Grficia et al. 2003). The external part of the Prebetic Zone is composed mainly of Triassic evaporites, unconformably overlain by an incomplete and discontinuous succession of Mesozoic to Neogene deposits. The most external part of the Prebetic system forms an
227
intricate system of tectonic thrust imbricates and chaotic units, which possibly correspond to an imbricate thrust system emplaced in a marine depocentre filled with numerous olistoliths and olistostromes (e.g. Azafion et al. 2002). Contacts between units in the External Betics are principally foreland-directed thrusts involving different cover palaeogeographical domains mostly detached above Triassic evaporites. The cover-basement contact is a major hinterland-directed thrust in the east and centre of the Betics (Banks & Warburton 1991), which changes to a foreland-directed one to the west, where it over-thrusts the Subbetic Unit as well as the Campo de Gibraltar Unit. This hinterland-directed thrust was a response to tectonic wedging produced by the antiformal stack of basement units flattening on top of the Triassic detachment level (Banks & Warburton 1991). As for the rest of the mountain ranges of the Iberian Peninsula, the Triassic evaporites constitute an excellent detachment level between basement and cover rocks. Several relatively small intermontane basins filled with Neogene deposits are located along the contact between the Internal Betics and the External Betics (e.g. Iribarren et al. 2003). The Neogene to Quaternary Guadalquivir Basin, to the NW and WSW of the External Betics, corresponds to a foreland basin in front of the Betics thrust system (e.g. Berfistegui et al. 1998; Garcfa-Castellanos et al. 2002). Hercynian rocks of the Iberian Massif constitute the NW boundary of the basin.
Neogene formation of the Western Mediterranean basins Beginning in mid-Oligocene times the opening of the Valencia Trough created an extensional fault system paralleling most of the eastern coast of northeastern Spain (e.g. Roca et al. 2004; Fig. 3). This system cut obliquely across the Early Tertiary compressive Catalan Coastal Ranges and cut the SE termination of the Iberian Range almost perpendicularly, forming an extensional arrangement of basins parallel to the Mediterranean coast (e.g. Roca et al. 1999). Concomitant uplift of segments of the Catalan Coastal Ranges as well as of the SE margin of the Ebro Basin initiated the development of the present landscape configuration (e.g. Morgan & Fernfindez 1992; Lewis et al. 2000; GasparEscribano et al. 2004). Fission-track studies indicate that more than 1.5 km of uplift occurred, responsible for the significant dissection of this margin (Juez-Larr6 & Andriessen 2002). To the SE of the Iberian Plate, an early Miocene large-scale system of normal faults connected to the Alboran Sea extensional system cuts the rear flank of the Internal Betics antiform. Most of these large normal faults are subparallel to the Internal Betics antiformal thrusts, reactivating some of them (e.g. Platt & Vissers 1989; Garcfa-Duefias et al. 1992; Crespo-Blanc et al. 1994; Comas et al. 1999). The Alboran Basin is filled by up to 8 km of early MioceneQuaternary sedimentary sequences (e.g. Comas et al. 1999).
Present-day topographic configuration of the Iberian Peninsula The present-day mean elevation of the Iberian Peninsula, slightly over 600 m, is almost certainly the highest in Western Europe. Smith (1996) pointed this out but did not propose a good solution for what is sustaining it. However, the integration of the tectonic structure of the Iberian Peninsula with its topography (Fig. 4), Bouguer anomalies (Fig. 5), and existing 2D and 3D lithospheric models can explain most of this high topography. The greatest negative Bouguer anomalies of the Iberian Peninsula show a very good match with its principal Tertiary mountain chains such as the Pyrenees, the Iberian Range, the Central System and the Betics (e.g. Casas Sainz & Faccenna 2001; Cloetingh et al. 2002). Available geophysical modelling
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J. VERGES & M. FERN,~NDEZ
indicates that these negative anomalies, corresponding to compressive systems, are primarily explained by crustal thickening, which in most cases is combined with lithospheric thickening such as has been inferred for the Pyrenees (Zeyen & Fern~mdez 1994) and the Central Betics (Tome et al. 2000). Cloetingh et al. (2002) proposed an alternative mechanism, lithospheric folding, as the main cause of alternating mountain ranges and basins of the central part of Iberia, but combined crustal and lithospheric modelling, as for the Pyrenees by Zeyen & Fernhndez (1994), does not support this interpretation.
To the west and SW of the Iberian Range there is a very large segment of the Iberian Peninsula with an elevated topography between 600 and 1000 m (Fig. 4). This region corresponds to the Central Iberian Zone of the Variscan Iberian Massif and includes the Tajo and Duero basins, and also shows a significant negative Bouguer anomaly (Fig. 5), suggesting that the crust of this region is either thickened or less dense than the surrounding crust, or a combination of both (Fernandez et al. 1998). The former interpretation can be applied to the NE domain (Duero and Tajo basins) where thick Tertiary sedimentary successions
Fig. 4. Combined tectonic and topographic map of the Iberian Peninsula.
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229
Fig. 5. Combinedtectonic and Bouguer anomaly map of the Iberian Peninsula (from Mezcua et al. 1996).
are present. However, to account for the regional distribution of the negative Bouguer anomaly and elevated topography the hypothesis of a less dense crust for the whole of the Central Iberian Zone is more realistic. Finally, in the southwestern Iberian Peninsula, the Ossa Morena and South Portuguese zones are dominated by a Bouguer anomaly maximum (Fig. 5), coinciding with an average topographic height
of about 200 m. Field studies demonstrate an increase in crustal density for these two Hercynian domains. A 2D lithospheric model, which integrates elevation, gravity, geoid and heat-flow data, indicates that the present lithospheric structure in these two domains, with relatively high crustal density, must be underlain by a thinned lithosphere or by a depleted lithospheric mantle (FernSndez e t al. 2004). According to the model, the mass deficit
230
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Fig. 6. Lithosphericthicknessfor the Valencia Trough (Ayala et al. 2003) and Alboran Sea (Tome et al. 2000) superimposed on the tectonic map of the Iberian Peninsula.The volcaniccentres along the Mediterranean coast coincide with the main entrances of the 70 km thick lithosphere line (note that all these fields contain alkaline volcanic rocks). The Campo de Calatrava volcanicprovince crops out in the centre of Iberia. (Abbreviations as in Fig. 2; v.p., volcanic province).
related to mantle thinning or depletion implies a lithospheric buoyancy-driven uplift of about 120-150 m to fit the observed elevation and geoid data (Fernandez et al. 2004). On the Mediterranean margin of Iberia, crustal and lithospheric mantle thinning show dissimilar patterns. Thin lithospheric mantle underlies onshore Iberia at three main localities (marked by the 70 km thick lithosphere contour; Fig. 6): in NE Spain, on the southern flank of the Valencia Trough, and in SE Spain. Positive Bouguer anomalies, crustal and lithospheric thinning, high topography, and asthenospheric volcanism characterize these three regions. In NE Iberia, several studies show that thinning was more intense in the lithospheric mantle than in the crust, producing additional uplift as well as basic volcanism (e.g. Cabal & Fernhndez 1995; Lewis et aL 2000; Ayala et al. 2003; Fig. 7). Radiometric ages for volcanic rocks of the NE volcanic province indicate that there was a migration of volcanism to the SW and west from 14 to 0.011 Ma (e.g. Saula et al. 1994; Lewis et aL 2000; Marti 2004). This indicates that lithospheric thinning and concomitant uplift started during late mid-Miocene times and is still active at present. At the southern end of the Valencia Trough, the SE volcanic province, in Levante (Fig. 6), shows ages ranging from 8 to 1 Ma (Ancochea & Huertas 2004). Towards the SE of Iberia, 3D crustal and lithospheric modelling shows that this area is also supported dynamically (Tome et al. 2000). The SE volcanic province in Almer~a and Murcia shows only a few alkaline volcanic edifices (L6pez-Ruiz et aL 2004; Fig. 6). The lithospheric geometry and the existence of onshore alkaline volcanic provinces along the coast as well as in the interior of Iberia (Campo de Calatrava volcanic province; Fig. 6) has been interpreted as being indicative of a long-lived deep process related to the opening of the Atlantic Ocean (Oyarzun et al. 1997). According to Oyarzun et al. (1997), the thinning along the Western Mediterranean region corresponds to a long,
sublithospheric channelling starting in the Cape Verde and ending in the northern North Sea (Ziegler 1990). Thus, the present-day topography of the Iberian Peninsula was mostly acquired by means of three main tectonothermal mechanisms effective since the Late Paleozoic: variations of crustal densities and possibly mantle depletion inherited from the Hercynian Orogeny, crustal and lithospheric thickening during Tertiary compression, and upper mantle thinning during the Neogene-Quaternary. However, the present landscape of the Iberian Peninsula has also been sculpted by the opening of numerous endorheic basins. The most spectacular of these openings occurred in the Ebro Basin during the early late Miocene (Coney et al. 1996; Lewis et al. 2000; Garcfa-Castellanos et al. 2003). The best preserved basin is the Duero Basin, in which fiver incision is still relatively minor.
Summary Between a widespread Triassic phase of extension and the opening of the central Atlantic in Late Jurassic times, Iberia was affected by extensional processes that created numerous sedimentary basins with different orientations, including the Pyrenees, Catalan and Iberian basins, and the Betic basin on the southern margin of Iberia. The Alpine-Tethys Ocean opened at this time, forming the eastern margin of the Iberian Plate. An additional extensional phase, forming the western margin of Iberia, took place before the opening of the North Atlantic during the early Aptian. This transform boundary propagated eastwards from the Bay of Biscay and constituted the northern plate boundary of the Iberian Plate, resulting in its isolation. These processes ended by the end of the Coniacian. Basins formed during these extensional phases were inverted during Late Cretaceous-Tertiary compression, which started in
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231
Fig. 7. Map of the NE corner of the Iberian Peninsula to show two areas affected by Neogene and Quaternary uplift (from Lewis et al. 2000). The SW area is limited to the footwall of the Oligoceneearly Miocene normal faults that parallel the coastline. The NE area is wider and affects part of the previous area and both the footwall and hanging wall of the late Miocene-Quaternary NW-SE-trening system of normal faults. The volcanism migrated westwards from the Empord~t Basin to the la Selva Basin and finally to La Garrotxa, where it is as young as 0.011 Ma.
Santonian times. The extent and orientation of the basins controlled the size, width and trends of the compressive belts, which normally show a double vergence. In almost all the Iberian fold-and-thrust belts, Triassic evaporites provided an excellent decoupling level between the basement and the cover units. The Iberian mountain chains produced lithospheric flexural bending and thus foreland basins, especially in front of the Pyrenees and the Betics: the Ebro and Aquitaine basins flank the Pyrenees and the Guadalquivir Basin lies to the north of the Betics. The Duero and Tajo basins developed to the south of the Cantabrian Mountains and to the south of the Central System, respectively. The Ebro, Duero and Tajo basins became intermontane basins at different stages of their evolution by the closure of their Atlantic connections. The basins were filled by radial fluvial systems feeding central lakes. Starting in the late Oligocene, the opening of the Valencia Trough initiated the development of a system of normal faults and linked basins aligned with the present Mediterranean coastline of Iberia. The deformation history of the Iberian Peninsula produced a significant mean elevation that is the highest in Europe. Tertiary compressional processes contributed strongly to an increase of mean elevation, but inherited Hercynian lithospheric structures as well as late Cenozoic upper mantle thinning related to the opening of the Western Mediterranean also contributed to the high average elevation of the Iberian Peninsula. Some of the thermo-mechanical
processes affecting the eastern margin of the Iberian Peninsula are still active at present. This is a contribution of the Group of Dynamics of the Lithosphere (GDL), Department of Geophysics and Tectonics, Institute of Earth Sciences 'Jaume Almera', CSIC. Partial support for this paper was provided by MCYT projects REN2001-3868-C03-02/MAR, REN2002-11230-E-MAR, NATO Grant EST-CLG978922. 2001 SGR 00339, and project 2001 SGR 00339 Grup d'Estmctura i Processos Litosf~rics, funded by the Comissionat per Universitats i Recerca of the Generalitat de Catalunya, Grups de Recerca Consolidats, II Pla de Recerca de Catalunya. We finally thank an anonymous reviewer for constructive remarks and suggestions.
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Contrasting modes of ophiolite emplacement in the Eastern Mediterranean region A L A S T A I R H. F. R O B E R T S O N Grant Institute of Geology and Geophysics, School of GeoSciences, University of Edinburgh, Edinburgh, EH9 3JW, UK (e-mail:
[email protected])
Abstract: The Eastern Mediterranean region is characterized by one of the largest concentrations of ophiolites anywhere in the world. Many of these ophiolites are fragmentary or highly deformed, such that their initial mode of tectonic emplacement cannot easily be inferred from the local field relations. The emplacement of many of these ophiolites can usefully be compared with the intact Oman ophiolite, one of the largest and best-studied ophiolites in the world. The Oman ophiolite is commonly believed to have been created in Late Cretaceous time (c. 95 Ma) above an oceanward-dipping, intra-oceanic subduction zone. This was followed by collision of the subduction zone with the downflexed Arabian passive margin, facilitating the emplacement of the ophiolite onto the continental margin. A less likely alternative is that the Oman ophiolite formed at a mid-ocean ridge that then collapsed, initiating the emplacement of the ophiolite. An Oman-type model is applicable to many of the Mid-Jurassic and the Late Cretaceous ophiolites of the Eastern Mediterranean region that were thrust over former passive continental margins. These ophiolites are again mainly of suprasubductionzone type. Such ophiolites include many of the Jurassic ophiolites of Greece, Albania and former Yugoslavia, and also the Late Cretaceous ophiolites of Turkey and northern Syria. These ophiolites were emplaced from both more northerly and southerly Neotethyan ocean basins. In contrast, the opposing (northerly) margins of these oceanic basins experienced a history of subduction-accretion, marginal arc volcanism and back-arc basin formation ('Cordilleran-type' ophiolites). Ophiolites that were emplaced associated with active margin settings range from large accreted thrust sheets to small slices within accretionary prisms and back-arc basins. Examples include the Late Cretaceous ophiolites that are related both to the northern margin of the southern Neotethys and to the northern margin of the northern Neotethys in Turkey. Not all ophiolites were emplaced in response to large-scale horizontal tectonic transport (e.g. Jurassic Guevgueli ophiolite, northern Greece), and several ophiolites experienced dominantly strike-slip or transpression (e.g. the Late Cretaceous Antalya ophiolites, SW Turkey). In general, the mode of ophiolite emplacement, especially the direction of emplacement relative to the orientation of the adjacent continental margin was influenced by the regional palaeogeographical setting.
The objective of this paper is to discuss and interpret the tectonic processes related to the emplacement of ophiolites exposed in the Eastern Mediterranean region (Fig. 1). Many but by no means all of these ophiolites were emplaced by processes that were similar to that for the emplacement of the Semail ophiolite in the Oman Mountains, the most complete, widely exposed and bestdocumented Tethyan ophiolite. The main focus here is on information gained from the tectonostratigraphy of the ophiolites in various settings and of different ages throughout the Eastern Mediterranean region. A glance at a tectonic map of the world shows a greater density of ophiolites per unit area in the Eastern Mediterranean region than anywhere else, except possibly Alaska or Indonesia (see Hoeck et al. 2002). With the advent of plate tectonics theory, ophiolites came to be seen as allochthonous slices of oceanic crust and mantle emplaced onto continental margins, as exemplified by the Semail ophiolite, Oman (Glennie et al. 1973, 1974) and the Bay of Islands ophiolite, Newfoundland (Williams & Smythe 1973; Jenner et al. 1991). Alternative tectonic scenarios were initially envisaged for the emplacement of various ophiolites (Dewey & Bird 1970; Dewey 1976; Casey & Dewey 1984), including several from the Eastern Mediterranean region (Woodcock & Robertson 1985). A complete understanding of ophiolite emplacement, however, has remained elusive, as few modern analogues exist. Onland and marine evidence of oceanic crust emplacement (potential future ophiolites) in different modern settings remains sparse (e.g. Dilek et al. 2000). Also, some ophiolites have reached their final locations following re-thrusting associated with regional continental collision, thus concealing the processes of initial emplacement from the ocean onto a continental margin. Different ophiolites were emplaced by several different tectonic processes and no one tectonic model is applicable to all ophiolites. Recent reviews of ophiolite emplacement processes worldwide (Dilek & Newcomb 2003; Dilek & Robinson 2003) re-emphasize the wide range of opinions concerning ophiolite genesis and emplacement, and highlight the need for them to be interpreted in the context of their regional tectonic settings. An attempt will be made here to identify several contrasting modes of ophiolite emplacement that mainly apply to Mid-
Jurassic and Late Cretaceous ophiolites of the Eastern Mediterranean region (Mukasa & Ludden 1987; Liati et al. 2004). The first setting relates to the emplacement of ophiolites onto former passive continental margins (Oman-type model). A second setting relates to ophiolite emplacement along active continental margins ('Cordilleran-type' model) by processes including subduction-accretion or the collapse of back-arc marginal basins. Because the ophiolites of the Eastern Mediterranean are commonly incomplete or dismembered, it is useful to draw comparisons with the largest and best-documented Tethyan ophiolite, in Oman (Fig. 1). The Troodos ophiolite, Cyprus, is undeformed by emplacement-related processes (Gass 1990) but cannot be used as a model for ophiolite emplacement, in view of the absence of any exposed base to the ophiolite, a complex local tectonic setting and its insular position. Several tectonic models for the genesis of the Oman ophiolite exist but there is a general consensus concerning its mode of emplacement over the Arabian passive margin. Should the Omantype emplacement be regarded as a 'one-off" or is it more widely applicable; for example, to the Eastern Mediterranean region? A summary of alternative models for the genesis of the Oman ophiolite and its emplacement is given first, followed by more detailed comparisons with Eastern Mediterranean ophiolites. Ophiolites that have formed in settings that differ from that of Oman are discussed later in the paper. The O m a n trench-collision model A widely accepted model for the tectonic setting of ophiolite genesis and emplacement in Oman (Fig. 2) is that presented by Lippard et al. (1986), with minor modifications. The future Oman ophiolite was generated by spreading above an oceanwarddipping (i.e. NE-dipping) intra-oceanic subduction zone in early Late Cretaceous time (c. 95 Ma; Pearce et al. 1981, 1984; Fig. 3). The exact location of subduction zone initiation is still unclear, whether near or some distance from the former spreading axis. A location removed from the ridge axis is probable in view of the argument that very young oceanic lithosphere may not be
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 235-261. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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A.H.F. ROBERTSON
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subductable (e. g. Cloos 1993; Shervais 2001). The well-developed metamorphic sole of the Oman ophiolite was accreted to the base of the oceanic slab beneath the depleted mantle tectonites of the over riding ophiolite. The high-temperature metamorphic sole was probably created some distance from the former spreading axis, if it is again assumed that young oceanic lithosphere is not normally subductable. This high-temperature metamorphism took place either during steady-state subduction (Searle & TERTIARY
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E. MEDITERRANEAN OPHIOLITE EMPLACEMENT
shallow-level thrust (i.e. related to a collapsed spreading axis). Continental margin-derived sediments were thrust deep beneath the oceanic mantle wedge, giving rise to granulite-facies metamorphism, implying the presence of a subduction zone (Searle & Cox 1999). In addition, granitic rocks were locally formed by melting of sediments at depth, followed by intrusion up to the level of the ophiolitic Moho in the northern Oman Mountains (Cox et al. 1999). A model of subduction trench-passive margin collision in Oman can also explain the occurrence of blueschists, south of the Semail Gap (e.g. E1-Shazly et al. 1990). The passive margin in this area underwent attempted subduction (oceanwards) until underthrusting of a thickening wedge halted subduction. This was rapidly followed by buoyancy-driven exhumation (Searle & Cox 1999, and literature cited therein). A more complex scenario is required in the MOR model, involving genesis of a separate intra-oceanic subduction zone and the transfer of the emplacing ophiolite from one convergence zone (i.e. collapsed ridge) to another (subduction zone) prior to final emplacement over the Arabian margin.
constructed as subduction proceeded (Fig. 2). The subduction zone collided with the Arabian passive margin several million years later, as indicated by combined palaeontological and radiometric dating (Tilton et al. 1981; Hacker & Gnos 1997). As the oceanic slab impinged on the passive margin, this then flexurally subsided beneath the advancing thrust load to form a foredeep (Robertson 1987a). This subsidence accommodated the subaqueous emplacement of the ophiolite over the Arabian continental margin as one, or several, vast thrust sheets (Glennie et al. 1973, 1974; Gealey 1977; Rabu et al. 1990; Robertson & Searle 1990; Le Metour et al. 1995; Searle & Cox 1999). A promontory of the Arabian margin in the south (to the south of the Semail Gap) was subducted and metamorphosed to high-pressure facies, then rapidly exhumed (see Searle et al. 2003, for literature review). In contrast, continental margin units further north (e.g. slope sediments of the Sumeini Group) remained unmetamorphosed. There is an alternative tectonic model for the genesis and emplacement of the Oman ophiolite. In this model the ophiolite formed by spreading at a mid-oceanic ridge (MOR), rather than above a subduction zone (Coleman 1971, 1981; Hopson & Pallister 1980; Coleman & Hopson 1981; Boudier & Nicolas 1988; Nicolas 1989), in a setting perhaps akin to the small East Pacific Juan de Fuca spreading axis. In the mid-ocean ridge interpretation the spreading ridge later collapsed to form an intra-oceanic thrust along which the ophiolite and its underlying metamorphic sole were emplaced towards the adjacent continental margin (Nicolas & Le Pichon 1980; Boudier et al. 1988). The view favoured here is that the Oman ophiolite formed above a subduction zone. However, in both the suprasubduction zone (SSZ) model and the MOR model young, hot oceanic crust and upper mantle were emplaced over a passive continental margin. This general mode of ophiolite emplacement contrasts with some other settings where ophiolites were emplaced along active continental margins ('Cordilleran ophiolites'), as discussed later in the paper. Several features of the emplacement of the Oman ophiolite appear to support the subduction-related hypothesis. There is strong evidence that the emplacement-related thrusting was deep-seated (i.e. involving a subduction zone), rather than a !
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237
Palaeogeography of the Eastern Mediterranean region A brief summary is given below to facilitate comparisons of the settings of ophiolite emplacement (Fig. 4). There is still considerable discussion about the palaeogeography of the Eastern Mediterranean region, especially during the MesozoicEarly Tertiary time of important ophiolite genesis and emplacement. In outline, the African and European continents were separated by generally older and younger oceans known, respectively, as Palaeotethys (Late Palaeozoic-Early Mesozoic) and Neotethys (Mesozoic-Early Tertiary). Palaeotethyan ophiolites are mentioned only briefly here as they are fragmentary, commonly metamorphosed and still poorly understood. However, pre-Late Jurassic 'Palaeotethyan' ophiolites were emplaced in the Pontides, northern Turkey (e.g. Ydmaz et al. 1997; Fig. 4). According to some workers (~engrr et al. 1980, 1984; Okay et al. 1991; Grnctio~lu et aI. 2000) this emplacement was the result of southward closure of a wide Palaeotethys to the north.
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238
A.H.F. ROBERTSON
However, others have inferred emplacement related to northward subduction of a Palaeotethyan ocean located to the south (rooted towards the present Izmir-Ankara-Erzincan suture zone: Robertson & Dixon 1984; Dercourt et al. 1986, 1992, 2000). Recently, the Palaeotethys of the above workers was reinterpreted as one, or several, marginal basins related to back-arc spreading behind a wide 'Palaeotethys' located further south (Stampfii et al. 2001). However, recent work has discounted the existence of a southerly Palaeotethys as in this tectonic model (Robertson 2006). Also, the related Palaeotethyan units of northern Turkey differ strongly from the nature of modern back-arc basins, as documented from the modern SW Pacific regions (see Robertson 1994, for comparisons). It is instead assumed that Palaeotethys was rooted in the north near the Izmir-Ankara-Erzincan suture zone. This persisted in some form or other as an oceanic area from Late Palaeozoic to Early Tertiary time, evolving into the Northern Neotethys, now represented by the Izmir-Ankara-Erzincan zone in Turkey (Robertson & Dixon 1984; Robertson et al. 2004b). The inferred northward subduction of a Palaeotethys rooted to the south of the Eurasian margin in time led to the construction of a regionally extensive accretionary wedge (Tekeli 1981; Pickett & Robertson 1996, 2004), known as the Karakaya Complex. The type area of the Karakaya Complex is located in the western Pontides (Okay et al. 1991, 1996), with counterparts in the central Pontides and eastern Pontides (Yflmaz et al. 1997; Okay & Sahintiirk 1997; Ustarmer & Robertson 1997, 1999). These inferred accretionary units are dominated by mrlange, including accreted volcanic seamounts (Ntilifer Unit) and inferred continental margin units (e.g. ~al Unit), as exposed in the NW Pontides. Most of the accreted material was derived from the upper levels of the oceanic crust and only exceptionally includes plutonic ophiolitic rocks. Small slices of ophiolitic rocks, including MOR-type basalts covered by radiolarites slices (Pickett & Robertson 1996), are found within the Karakaya Complex, presumably representing accreted fragments of Palaeotethyan ocean floor. There is also the much larger Denizgrren ophiolite, which can be correlated with the Chios ophiolite on the offshore Greek island (Pickett & Robertson 1996). The field relations are compatible with the Denizg6ren ophiolite being a Palaeotethyan ophiolite (Okay et al. 1991; Pickett & Robertson 1996). However, Ar-Ar dating of the Denizgrren metamorphic sole has suggested an Early Cretaceous age (Okay et al. 1996), pointing to an origin more related to the emplacement of the Jurassic ophiolites of the Inner Hellenide Ophiolite Belt in northern Greece (see below). Although the Karakaya Complex is dated as Triassic in NW Turkey (Okay et al. 1991, 1996), it is possible that the Eurasian margin was long lived and that similar but older Karakayatype units may exist elsewhere in the Pontides. Palaeotethyan ophiolites are, indeed, also present in other areas adjacent to the Eurasian margin, including Iran (Strcklin 1974) and the Caucasus (Adamia et al. 1981, 1995), but require additional studies to determine the processes of genesis and emplacement in detail. Northward subduction of Palaeotethys, possibly obliquely (Robertson & Dixon 1984), is considered to have given rise to rifting behind a continental margin arc (~angalda~ unit) in the central Pontides (Ustarmer & Robertson 1997), dated as being active during Late Palaeozoic-Early Tertiary time (Kozur et al. 2000). Back-arc rifting culminated in opening of a small marginal basin floored by Eurasian-derived terrigenous sediments (Ustarmer & Robertson 1994; Robertson et al. 2004; Fig. 5). This basin later closed by southward subduction, accreting small ophiolitic slices, known as the Kfire ophiolite, within mainly terfigenous turbidites, followed by covering by Early Jurassic neritic carbonates. In contrast to the Palaeotethyan ophiolites of pre-Late Jurassic age, ophiolites of Jurassic age are well exposed in the Balkan region and large ophiolites of Late Cretaceous age dominate the region further east, including Turkey, Cyprus, Syria, Iran and Oman. The emplacement of these ophiolites is the main subject of this paper. According to some researchers most, or all, of
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Southerly Mediterranean ophiolites The 'southerly' ophiolites (Figs 6 and 7) include the Jurassic Balkan ophiolites, of which the best known are the Pindos, Vourinos and Othris ophiolites in Greece, the Jurassic Mirdita ophiolites in Albania and the Dinaride ophiolites of former Yugoslavia. The Jurassic Balkan ophiolites are discussed first, below. These ophiolites are all associated with the 'Pindos suture' shown in Figure 4. For the purposes of discussion here these 'southerly' ophiolites also include the Late Cretaceous ophiolites that are exposed along the Arabian margin (Croissant peri-Arabe of Ricou 1971) from Oman, through Iran, to Syria and the Late Cretaceous ophiolites of Cyprus and Antalya (southern Turkey). These ophiolites are associated with the SE Turkish, Antalya and Mamonia sutures shown in Figure 4. All of these ophiolites are believed to have been emplaced from a southern Neotethys ocean that was located outboard of the Arabian continent (Robertson et al. 1991). The' southerly' ophiolites also include a belt of Late Cretaceous ophiolites further north in Turkey, associated with the IzmirAnkara-Erzincan suture (e.g. the Lycian and Bey~ehir-Hoyran ophiolites), which are restored to an origin within a more northerly Neotethyan oceanic basin. J u r a s s i c B a l k a n ophiolite c o m p a r e d with O m a n
The Oman-type trench-passive margin collision model is generally applicable to ophiolites located within the Pindos zone of Greece (e.g. Othris, Vourinos, Pindos, Evia, Argolis and its
E. MEDITERRANEANOPHIOLITEEMPLACEMENT
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continuation into Albania in the Mirdita zone as the 'Western-type' and 'Eastern-type' ophiolites (Shallo et al. 1990; Robertson & Shallo 2000). This model was proposed for the Pindos ophiolites in NW Greece by Jones & Robertson (1991;
239
see also Jones et al. 1991; Robertson et al. 1991) and has been widely applied to other Tethyan ophiolites since then by many workers (with varying modifications). More recently, a similar model was applied, for example, to the Albanian ophiolites (e.g. Robertson & Shallo 2000; Dilek et al. 2005; Koller et al. 2006) and the Greek ophiolites (Clift & Dixon 1998; Smith & Rassios 2003). Many of the Jurassic Balkan ophiolites (e.g. Pindos, Vourinos, Evia, Eastern-type Albanian ophiolite) show geochemical evidence of formation in a subduction-related setting (e.g. Capedri et al. 1980; Noiret et al. 1981; Jones et al. 1991; Beccaluva et al. 1994; Robertson & Karamata 1994; Cliff & Dixon 1998; B~bien et al. 2000; Rassios & Smith 2000; Pami~ et al. 2002; Bortolotti et al. 2004; Koller et al. 2006; Rassios & Moores 2006). Remnants of MOR-type oceanic crust are also rarely preserved, including some of the ophiolitic units of the Othris area (Smith et al. 1975) and the Western-type Albanian ophiolite (Shallo et al. 1990). Many of these ophiolites are underlain by metamorphic soles and volcanic-sedimentary mrlange. Also, in several cases, the sedimentary cover is preserved, and includes deep-sea sediments and debris-flow deposits containing exotic rocks derived from continental margin units. These associated units, as well as the ophiolites themselves, provide important information about the processes of ophiolite emplacement. One of the most problematic aspects, assuming a SSZ origin of the Jurassic Balkan ophiolites is accepted, is the setting of the initiation of subduction to generate these ophiolites. Most researchers have assumed that this subduction was initiated at, or near, the spreading axis for the Oman ophiolite (Lippard et al. 1986; Searle & Cox 1999) and the Jurassic Balkan ophiolites (Jones & Robertson 1991; Robertson et al. 1991). Recently, Smith & Rassios (2003) have adopted an Oman-type model in which ophiolite emplacement was driven by the collision of a subduction trench with the passive margin (Pelagonian zone). However, they suggested that subduction was initiated, not within the ocean, but instead adjacent to the western margin of the ocean basin. Specifically, the subduction zone was located along the eastern side of the marginal Parnassus carbonate platform and within marginal oceanic crust to the north (and presumably to the south). Thus, after subduction began the preexisting ocean basin was, in effect, replaced by new SSZ oceanic crust, although marginal oceanic crust remained to the north and south of the Parnassus block. This obviates the apparent difficulty of initiating spreading within relatively young oceanic lithosphere. Unfortunately, the eastern margin of the Parnassus platform is a thrust contact with the Pelagonian zone, itself part of the eastern margin of the oceanic basin, and thus the model cannot be directly tested in the field in this area. If applicable, however, the above model could apply regionally throughout the Balkans (Pindos-Mirdita ophiolites) and possibly more widely. In this context several points can be made. (1) If spreading originated near the westerly passive margin, a record of this event might be observed, as margin sediments are well preserved in the Pindos-Olonos Nappes (e.g. in the Peloponnese). If subduction began near this margin, significant tectonic disruption could be expected, combined with formation of a bathymetric trench, which could become a sediment sink with a good chance of ultimate preservation. The Pindos-Olonos thrust sheets record proximal to distal successions that prograded over Triassic oceanic crust (or transitional crust). During Triassic-Jurassic time these successions record passive margin subsidence and a switch to pelagic radiolarian deposition, with the first major siliciclastic input not being until Cretaceous time, after genesis and emplacement of the Pindos-Mirdita ophiolites (Degnan & Robertson 1998, 2006). Similar constraints apply further NW throughout Albania and former Yugoslavia (e.g. Robertson & Shallo 2000). There is thus no stratigraphical evidence for the initiation of a westward-dipping subduction zone along the westerly margin of the ocean basin. (2) ff a subduction zone developed near, or along, this passive margin, a sedimentary record of this event could be expected
240
A.H.F. ROBERTSON
within the sedimentary cover of the ophiolitic extrusive rocks, where well preserved. This cover might be expected to resemble that of the distal passive margin, as preserved in the Pindos-Olonos thr-ust sheets (i.e. mixed muddy, siliceous and calcareous sediments). However, the cover of the ophiolific extrusive rocks, as best preserved in the Albanian ophiolites (Mirdita zone), consists of well-dated ribbon radiolarites without any relatively coarse marginderived sediments lying directly on the ophiolite (e.g. Prela et al. 2000; Danelian & Robertson 2001). (3) If nearly the entire width of the Pindos-Mirdita ocean was 'replaced' by SSZ lithosphere it is surprising that the western-type ophiolites in the extension of the Greek ophiolites into Albania are of MOR-type (e.g. Shallo et al. 1990). According to Dilek et al. (2005), radiometric dating of plagiogranites within both the western- and eastern-type ophiolites indicates that both are of similar age (c. 162-165 Ma) and thus that two contrasting magma types were involved in oceanic crust genesis within a relatively short period of time. This is explicable if a MOR-type setting evolved into an SSZ-type setting relatively rapidly following the initiation of a subducfion zone at, or near, a ridge crest. It is less easy to envisage why MOR-type lavas were erupted in an entirely subduction-controlled setting. (4) The model implies that the original Late Triassic-Early Jurassic spreading ridge was subducted beneath the passive margin, but no obvious thermal effect of this is seen along the emplaced passive margin (e.g. in the Parnassus unit). The available evidence, therefore, provides little obvious support for the hypothesis of subduction initiation at or near the western margin of the oceanic basin, although this option cannot be ruled out. It is interesting to apply the above model of subduction initiation along a passive margin more widely. If applied to the Oman region, the entire ocean floor of the Gulf of Makran and related accreted material within the Makran accretionary prism should be of similar SSZ crust. Applied to the southern Neotethys in Turkey, the model would have difficulty in explaining the presence of arc-type rocks associated with the northern margin, where, by comparison, subduction would have initiated. A similar difficulty applies to the northern Neotethys in Turkey, as again arc rocks are present in a northerly part of the basin (see discussion later in the paper). Applied to the modern oceans, the question arises as to why SSZ spreading apparently ceases, followed by construction of a volcanic arc (Bloomer et al. 1995), rather than the subducting slab continuing to roll-back indefinitely. The main objection to the hypothesis of subduction zone initiation at or near a spreading ridge is the theoretical difficulty of subducting young oceanic lithosphere. This problem may, however, be diminished for the Jurassic Balkan ophiolites, as the MOR-type western ophiolite in Albania is interpreted as a rifted ridge (marked by low magma production), which might be subductable, especially if convergence was initiated along a fracture zone. In general, it is possible that subduction was initiated in different places in different oceans, basically where the crust was weakest. On balance, the writer prefers a model of subduction initiation well within the ocean, at least for the Jurassic ophiolites of the Balkan region. This leaves part of the initial ocean basin intact until much later final closure of the ocean. However, subduction initiation near a continental margin may apply in some cases (e.g. Late Cretaceous, northern Neotethys; see below). For the present discussion, we are more concerned with the mode of emplacement of the ophiolite onto a continental margin itself, rather than the initial setting of subduction, which may remain hypothetical as so little evidence has survived following regional convergence events. One problem is how to emplace a vast slab of oceanic crust and mantle, apparently upwards against gravity onto a continental margin, as depicted by Glennie et al. (1973). In Oman it was found from palaeo-environmental analysis of Late Cretaceous synemplacement sediments (Muti Formation) that the abyssal plain-slope-margin underwent thrust loading and flexural collapse into deep water ahead of the advancing ophiolite
(Robertson 1987a,b). Thus, no major vertical uplift of the emplacing ophiolite was required. Similar flexural processes related to ophiolite emplacement can be inferred for many of the Jurassic Balkan ophiolites and the Late Cretaceous ophiolites further east, as discussed below. In Oman, the adjacent carbonate platform underwent flexural uplift and erosion, locally down to near the base of the Mesozoic carbonate platform succession, creating the regional WasiaArumu break (Glennie et al. 1973; Robertson 1987a; Figs 2 and 3 - l b ) . Alkaline volcanic rocks were locally erupted during this stage (Le Metour et al. 1995). The erosion surface was karstified and in places capped by glauconite- siderite ironstones (Robertson 1987c). The unconformity was attributed to the passage of a flexural bulge ahead of the advancing ophiolite. Rather than an idealized flexural bulge, however, the uplift was probably partially related to the reactivation (i.e. inversion) of older (Triassic) rift faults. After passage of the flexural bulge, the uplifted margin gradually submerged and was covered by deepening-upward neritic, to hemi-pelagic, sediments (Robertson 1987c). The platform margin later collapsed dramatically as a result of flexural loading of the advancing thrust pile. This collapse was associated with large-scale slumping and mass wasting of carbonate debris flows derived from the upper levels of the carbonate platform succession. A switch to non-calcareous muds and quartzose sandstone turbidites ensued. The inferred source of these sediments was rift-related, or basement, rocks exposed to erosion by uplift of the continental margin, rather than transport from the Arabian landmass, then far to the SW. Very similar uplift and collapse features are seen associated with the emplacement of the Jurassic Greek ophiolites (Fig. 7a). The underlying Pelagonian carbonate platform (Fig. 4), where exposed, shows evidence of block faulting prior to the arrival of ophiolite-derived clastic sediments. The carbonate platform subsided and was overlain by non-calcareous radiolarian sediments of Bathonian-Callovian age (e.g. in the Kalidrommon Mountains, central Greece; Danelian & Robertson 1995). This was followed by the arrival of ophiolite-derived sediment in Tithonian time (Thiebault et al. 1994). Evidence of the initial emplacement of oceanic units onto the continental margin is documented from successions within the Pelagonian carbonate platform. On the large island of Evia, eastern Greece (Figs 4 & 8), the Late Triassic to Mid-Jurassic carbonate platform succession culminates in an unconformity, interpreted as a subaqueous erosion surface, overlain by tuffaceous sediments and radiolarian sediments (Robertson 1991). Comparable tuffaceous sediments occur on the adjacent mainland, in the Othris Mountains (Price 1976). The local volcanism, as in Oman, is attributed to extensional processes accompanying flexural upwarping of the carbonate platform (Robertson 1991). Elsewhere on Evia (at Achladi beach), the typical Jurassic platform carbonates pass upwards into mudstones and radiolarites of Late Jurassic (Oxfordian) age. These sediments document the collapse of the carbonate platform in response to loading by the advancing allochthon (B aumgartner & Bernoulli 1976; Robertson 1991). The appearance of chromite grains preceded the arrival of ophiolitederived material generally. There are several possible alternative origins for the chromite grains. They were possibly derived from ultramafic rocks at the base of the overriding ophiolite, from slices of ultramafic rocks within an advancing accretionary wedge, or from submarine erosion of diapiric serpentinite formed within the rifted continent-ocean transition zone (e.g. as in the modern Iberia (Atlantic) rifted margin; e.g. Boillot et al. 1980). The last mentioned possibility was not previously considered. In this case, the ultramafic rock was exposed along the rifted margin during the original continental break-up and later reworked during tectonic deformation of the passive margin, prior to the arrival of exotic ophiolitic material. On Evia (e.g. at Achladi beach), the collapsed platform succession is overlain by debris-flow deposits and then by allochthonous oceanic-derived mrlange (Baumgartner & Bernoulli 1976;
E. MEDITERRANEANOPHIOLITE EMPLACEMENT
S
241
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Fig. 7. Schematic cross-sections of Greece and Turkey to illustrate the settings of emplacement of the Jurassic ophiolites (a) and the Late Cretaceous ophiolites (b).
•
Continental crust Oceanic crust (undifferentiated)
Robertson 1991; Scherreiks 2000). Further south, in the Argolis Peninsula, comparable emplacement-related debris flows contain abundant ophiolitic material, especially serpentinite (Baumgartner 1985). These debris flows are seen as the result of gravity TERTIARY
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Fig. 8. Summaryof the tectonostratigraphy of the Early-Mid-Jurassic Evia ophiolite, eastern Greece. Modified after Robertson (1991).
redeposition of material that was shed off advancing thrust sheets of continental margin sedimentary and volcanic rocks and oceanic crust. In Oman, continental margin-derived material forms a wellorganized thrust stack, including sedimentary units (Hawasina Complex; Glennie et aI. 1973; Searle & Stephens 1984; Cooper 1990; Bernoulli & Weissert 1997; Fig. 2) and largely volcanic units (Haybi Complex; Searle & Malpas 1980; Searle et al. 1980). The Hawasina Complex is interpreted to record an original transition from proximal rifted continental margin units to oceanic sediments. This complex is assumed to have been underlain by transitional-type or oceanic crust that has been entirely subducted (Bernoulli & Weissert 1987; Bgchennec et al. 1988; Robertson & Searle 1990). The Hawasina Complex is interpreted as an accretionary prism related to northeastward subduction of the distal edge of the rifted Arabian margin and adjacent, marginal Tethyan oceanic crust of Permo-Triassic age (Glennie et al. 1973; Lippard et al. 1986). The Haybi Complex is dominated by Permian and Triassic limestone exotic units (Oman Exotics; Glennie et al. 1973; Pillevuit et al. 1997) and Late Triassic alkaline volcanic rocks that variously formed both along the rifted Oman margin and as seamounts within the oceanic crust (Searle & Graham 1982; Lippard et al. 1986). The combined Hawasina and Haybi complexes were emplaced over the Oman margin (beneath the Oman ophiolite) when the subduction trench collided with the Arabian continental margin. Similarly, in Greece the Pelagonian carbonate platform was overthrust by rifted passive margin units, as seen in the Othris
242
A . H . F . ROBERTSON
21~
21~
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urassic ophiolite, mainly peridotite, ultramafic. fic cumulates, DRAMALA OPHIOLITE - Cretaceous platform carbonates and related ~ts; blocks assoc, with Avdella Melange L. Cretaceous accretionary melanges, _A MELANGE - Cretaceous d e e p - sea sediments, Zone, Dio Dendra Group
tceous - L. Tertiary Pindos flysch , Miocene cover sediments 3ic Hellenic Trough
)ic- Early Tertiary basinal sediments, ?_one
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area (Smith et al. 1975, 1979). The well-organized thrust stack there comprises proximal (marginal), to more distally derived thrust sheets of sedimentary (e.g. turbidites; radiolarites), volcanic (e.g. Agrilia lavas) and ophiolitic lithologies (Sipetorrema lavas) (Smith et al. 1975). Elsewhere, coherent thrust sheets are commonly absent and only a large-scale 'mrlange' is exposed, dominated by broken formation (dismembered thrust sheets) and debris-flow deposits ('olistostromes' of Ferri~re et al. 1988), all exposed beneath the ophiolites (e.g. Evia, Atalanti, Pindos). An excellent example of a mrlange, mainly derived from continental margin units, is exposed in the Pindos Mountains (Figs 9 and 10). The Pindos ophiolites are underlain by a thick unit of mrlange (Avdella Mrlange), including large blocks of Triassic W i PL!OCENE u MIOCENE M U ionian
Fig. 9. Simplified geological map of to show the setting of the Jurassic Dramala (Pindos) ophiolite in the Pindos Mtns, NW Greece (from Robertson 2002).
neritic limestone and locally coherent thrust sheets of Mid-Late Triassic volcanic-sedimentary lithologies, together with debrisflow deposits (Jones & Robertson 1991; Fig. 10). Similar, but volumetrically subordinate, mrlange locally underlies the Albanian ophiolites (Robertson & Shallo 2000). Elsewhere, in northern Greece the Pelagonian carbonate platform was overthrust by ophiolite-related mrlange and ophiolite slices (Sharp & Robertson 2006) All of the above mrlange units can be explained by the accretion of a variety of oceanic to continental margin units of sedimentary and volcanic origin, as shown in Figure 11 (Jones & Robertson 1994). Pre-existing topographic highs located along the rifted margin (e.g. rift fault blocks and carbonate build-ups; Fig. 1 la) were preferentially accreted (Fig. 1 lb and c), culminating in collapse of the
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Fig. 10. Summary of the tectonostratigraphy of the Jurassic Dramala (Pindos) ophiolite, NW Greece. Modified after Jones & Robertson (1991).
Fig. 11. Reconstruction of the rift-drift-emplacement history of the Pindos ophiolites and related units, northern Greece. (a) Late Triassic rifting; (b) Early to Mid-Jurassic initial subduction and accretion (Avdella Mrlange); (c) Bathonian-Oxfordian; approach of the Pelagonian continental margin, leading eventually to emplacement of the Pindos ophiolite and underlying Avdella Mrlange. From Jones & Robertson (1994). WPB, within-plate basalt.
E. MEDITERRANEAN OPHIOLITE EMPLACEMENT
adjacent margin of the Pelagonian platform and overthrusting by the accretionary prism and the ophiolites (Fig. 1 lc). The Oman ophiolite is underlain by a metamorphic sole. The sole is typically up to tens of metres thick, with a lower, greenschist-facies thrust sheet and an upper, amphibolite-facies thrust sheet (Searle & Malpas 1980, 1982; Ghent & Stout 1981; Gnos & Peters 1993). In some areas the metamorphic sole is laterally intact (e.g. Sumeini area), whereas in others it is more dismembered (e.g. Hawasina window), or thickened by structural repetition (e.g. Dibba Zone, Northern Oman Mountains; Searle & Cox 1999). In the Dibba Zone, detailed petrological and mineralogical studies indicate that continentally derived material was entrained with hot peridotite to form granulite-facies rocks. The protoliths were potassium-rich terrigenous and calcareous rocks, which, unusually, are exposed directly beneath the hanging wall of the metamorphic sole. In addition, partial melting of these highgrade metamorphic rock produced peraluminous granitic melts that percolated upwards into the overlying ophiolite (Gnos & Nicolas 1996). Searle & Cox (1999) argued that these processes could only have taken place as a result of the attempted subduction of continental crust beneath a hot mantle wedge. This particular setting of granulite-facies rocks and associated melts was not so far documented in the Eastern Mediterranean region. Very similar metamorphic soles to those of the Oman are present beneath the Greek and Albanian Jurassic ophiolites (Spray et al. 1984; Dimo-Lahitte et al. 2001). These range from intact (as in Evia, Vourinos and the Eastern-type Albanian ophiolites), with well-developed amphibolite- and greenschist-facies units, to blocks of variable lithologies strewn through underlying mrlange, as in the Pindos Mountains. The dismembering of the metamorphic sole of the Pindos ophiolite is largely attributable to re-thrusting of the allochthon associated with continental collision in Early Tertiary time (Jones & Robertson 1991). The details of the petrogenesis of the metamorphic sole are outside the scope of this paper. However, it is notable that the protoliths of the metamorphic soles vary considerably based on chemical evidence; they vary from near-MORB (mid-ocean ridge basalt) composition in Oman (Searle & Malpas 1982), to mainly within-plate-type (alkaline) basalts in Baer-Bassit, northern Syria (A1-Riyami et al. 2002a,b) and Mersin, southern Turkey (Parlak et al. 1995), to subduction-influenced tholeiites, for example, in Evia (Simantov et al. 1990; Danelian & Robertson 2001). Also, the protoliths may include various intrusive as well as extrusive ophiolitic rocks. This evidence shows that oceanic crust from different tectonic settings was incorporated into the metamorphic sole in different areas. The nature of the overlying relatively intact ophiolite provides additional information concerning emplacement processes. The emplaced ophiolites range from a single relatively coherent thrust sheet, as in Oman (although buried thrust sheets may exist offshore), to composite units made up of several different thrust sheets originating in different oceanic settings. Specifically, the Pindos ophiolites include a thin (several hundred metres) lower dismembered thrust sheet of boninitic lavas (high-magnesian andesites) (Aspropotamous ophiolite), overlain by the much thicker main Pindos ophiolite (Dramala unit; Fig. 10). The boninites are seen as forming in a forearc-type setting and provide additional evidence for a subduction-related origin of the Pindos ophiolites as a whole (Jones & Robertson 1991; Jones et al. 1991). The Jurassic Balkan ophiolites also provide evidence for the processes of initial deformation related to tectonic emplacement. In the Oman ophiolite it was found that the preferred orientations of olivine and orthopyroxene crystals, located within hightemperature mylonites (at the base of the ultramafic tectonites) could be used to determine the emplacement direction (Boudier et al. 1985). Similarly, in the Vourinos ophiolite, the primary fabric within the depleted upper mantle tectonites is overprinted by evidence of high-temperature deformation. This can be related to initial tectonic displacement of the ophiolite while it
243
cooled through the brittle-ductile transition (Rassios et al. 1994; Rassios & Smith 2000). Similar observations were made in the Pindos, Vourinous and Othris ophiolites. This evidence incidentally indicates that these ophiolites were initially emplaced to the NE (in present co-ordinates), towards the Pelagonian continental margin. In a few cases the structures in the underlying metamorphic soles may also suggest emplacement directions (e.g. Zlatibor ophiolite; Dimitrijevic & Dimitrijevic 1979). Interpretation is, however, difficult as the early transport lineations that developed at high temperatures commonly lack vergence indicators, and the sole may be disrupted and tectonically rotated during or after ophiolite emplacement. During its emplacement, the Oman ophiolite was segmented into crustal blocks and has undergone clockwise rotation, as shown by palaeomagnetic studies (Perrin et al. 1994; Thomas et al. 1988). These rotations are attributable to oblique subduction, or to diachronous collision with the Arabian margin. Comparable, to more extreme, rotations have affected the Late Cretaceous Troodos, Hatay and Baer-Bassit ophiolites (see later discussion). The nature of the sedimentary cover of the ophiolites can provide additional clues concerning emplacement processes. In Oman, the Semail ophiolite is depositionally overlain by a thin (< 10 m thick), but well-dated cover of metalliferous umbers, radiolarites and pelagic carbonates (Fleet & Robertson 1980; Tippett et al. 1981). In comparison, most of the Jurassic Balkan ophiolites lack intact sedimentary covers (see Danelian & Robertson 2001, for a recent review). Exceptionally, both the Eastern- and the Western-type Albanian ophiolites are depositionally overlain by thin (
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Fig. 23. Setting of ophiolite emplacement in the Isparta Angle, SW Turkey. (a) The area restores as an embayment of the southern Neotethys; (b) oceanic lithosphere formed in a SSZ setting; (c) ophiolites and marginunits were initially displaced in latest Cretaceous time and finally emplaced over the continental margin in Early Tertiary time; (d) the Isparta Angle began to form, while Neotethys still remained partially open to the south. From Robertson (2003).
a review). The Antalya area formed part of the northern, active margin of the southern Neotethys in Late Cretaceous time. Within this segment, the northern margin was oriented at an oblique angle to the generally east-west trend of the southern Neotethys. As a result of this obliquity, the emplacement of the Late Cretaceous ophiolites was strongly influenced by strike-slip faulting (Woodcock & Robertson 1982, 1985; Fig. 23). In the SW area of the Isparta Angle, the Cretaceous Tekirova ophiolite (Juteau 1975), and more marginal (proximal) units further west (i.e. G6dene zone) were dismembered into small (tens of kilometres or less) tectonic 'terranes' separated by highangle shear zones, interpreted as strike-slip faults (Woodcock & Robertson 1982). Internally, the Late Cretaceous Tekirova ophiolite, exposed along the coast (Reuber 1984), is relatively intact, although the extrusive rocks are missing. Geochemical data suggest a SSZ setting of formation (Ba~cl et al. 2006). The ophiolite is depositionally covered by a chaotic unit of
251
Maastrichtian age, composed of detached ophiolitic rocks and clastic sediments (Robertson & Woodcock 1982b; Lagabrielle et al. 1986). These clastic sediments lack a terrigenous component, implying that they formed in an oceanic setting before the ophiolite was emplaced over the adjacent Bey Da~larl carbonate platform. In addition, large-scale mass-wasting of ophiolite-derived debris flow deposits, commonly rich in serpentinite, took place on the adjacent (more inboard) terrane (Kemer zone) that is underlain by Late Palaeozoic pre-rift basement, Triassic synrift sediments and Jurassic-Cretaceous post-rift, mainly carbonate, sediments (Robertson & Woodcock 1982b). This unit is interpreted as one or several small continental fragments, or large rift blocks that were isolated along the rifted continental margin of the southern Neotethys when spreading began in Late Triassic time (Robertson & Woodcock 1980b). The adjacent small 'terrane' further west (G6dene zone) includes strongly dismembered Late Cretaceous ophiolitic rocks, Triassic rift-related volcanic rocks and sediments, and very rare small exposures of metamorphic sole-type rocks (Robertson & Woodcock 1982b; Ydmaz 1984; (~elik & Delaloye 2003). The G6dene zone also includes distinctive ophiolite-derived breccias and mass-flow deposits related to tectonic amalgamation of this unit in a strikeslip setting (Cmarqlk Breccias; Robertson & Woodcock 1980a). The Late Cretaceous Tekirova ophiolite was initially deformed in latest Cretaceous (Maastrichtian) time associated with strikeslip or transpression, outboard of the margin of the Bey Da~lan carbonate platform. The ophiolite was later emplaced onto the adjacent Tauride continental margin (Bey Da~lan platform) in response to partial closure of Neotethys during Early Tertiary (Late Palaeocene-Early Eocene) time (Robertson & Woodcock 1982b; Poisson 1984; Dilek & Rowland 1993; Robertson 1993). The neighbouring Bey Da~lan carbonate platform to the west remained undeformed in latest Cretaceous time as a submerged platform undergoing pelagic deposition into Early Tertiary time (Poisson 1977; Robertson & Woodcock 1982b). In summary, this is an example of ophiolite emplacement in which strike-slip, rather than orthogonal emplacement played an important role, in contrast to most of the other settings of ophiolite emplacement discussed above. The final setting of ophiolite 'emplacement' discussed here is one in which the ophiolite was not actually tectonically emplaced onto a continental margin at all, but instead was rotated as a microplate, still within an oceanic setting, associated with deformation of its margins. An example of this is the latest CretaceousEarly Tertiary palaeorotation of the Troodos ophiolite (Fig. 24). The Troodos ophiolite, Cyprus, was generated around 90 Ma (Mukasa & Ludden 1987) in a SSZ setting within the southern Neotethys ocean (e.g. Pearce et al. 1984). The Troodos ophiolite, together with the Hatay, Baer-Bassit, Tekirova and other Late Cretaceous southern Neotethyan ophiolites formed above a northdipping subduction zone (Robertson 1998). To the east the BaerBassit, Hatay and other ophiolites underwent southward emplacement onto the Arabian margin (Oman-type model) as discussed earlier, and to the west the Antalya ophiolites (e.g. Tekirova) were strongly affected by oblique (strike-slip) emplacement. However, the Late Cretaceous tectonic setting of the Cyprus region was strongly influenced by 90 ~ anticlockwise palaeorotation of the Troodos ophiolite. The rotation was first discovered by palaeomagnetic study of the Troodos ophiolite (Moores & Vine 1971) and was accurately dated by detailed palaeomagnetic studies of the in situ deep-sea sedimentary over of the ophiolite (Clube et al. 1986; Morris et al. 1990). The rotation began in Campanian-Maastrichtian time and continued at an approximately constant rate until Early Eocene time (Morris 1996). Recent work shows that the Hatay ophiolite also underwent anticlockwise rotation, but to a smaller extent (Morris et al. 2006). Also, tectonic rotations that may be extreme are recorded in the more dismembered Baer-Bassit ophiolite further south (Morris et al. 2002).
252
A.H.F. ROBERTSON
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Fig. 24. Outline geological map of Cyprus showingthe Troodos ophiolite, which underwent palaeorotation in latest Cretaceous-Early. Tertiary time while still within a remnant of the southern Neotethys ocean. Associated with this rotation, margin units in western Cyprus (MamoniaComplex) and northern Cyprus (KyreniaRange) were deformed in latest Cretaceous time and then covered with deep-sea carbonate sedimentsin Early Tertiary time.
One explanation for the rotation of the Troodos ophiolite is that this resulted from oblique northward subduction beneath the northern active margin of the southern Neotethys (Clube et al. 1985). Another possibility is that the palaeorotation relates to
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., Z?+.
- 9
Accretionary prism Slices of MORB or SSZ ophiolites
..-.____
9
9
~
II
~
' Subducting Oceanic crust
(b)Generalised
active margin-related
model
Fig. 25. Summaryof the main models of ophiolite emplacementin the Eastern Mediterranean, Tethyan region. (a) Collision of SSZ ophiolite with a passive margin (Oman-typemodel); (b) emplacementof ophiolites in active margin settings associated with subduction and accretionary complexes.Model (a) produces very similar geological features regardless of age or location (e.g. Jurassic Greek v. Late Cretaceous Turkish ophiolites), whereas model (b) shows considerablevariation in different examples.
collision of the inferred north-dipping subduction zone with the Arabian promontory to the east, thus triggering pivoting and anticlockwise rotation of a Troodos microplate (Clube & Robertson 1986). It was also proposed that the Late Cretaceous tectonic history of the Kyrenia Range, northern Cyprus, and of the Mamonia Complex, western Cyprus, were strongly influenced by strike-slip around the periphery of the Troodos microplate (Robertson 1990). The probable explanation for the sparsity of evidence for tectonic emplacement of the Troodos ophiolite in Late Cretaceous time, compared with ophiolites further east (e.g. Hatay; Baer-Bassit), is that the Troodos remained protected from deformation within an embayment of the Arabia-North Africa continental margin, today known as the Levant Sea (Clube & Robertson 1986; Robertson 1990). The palaeorotation of the Troodos microplate was spatially and temporally associated with the deformation of adjacent units. To the north, the Troodos ophiolite is inferred to terminate abruptly against the Kyrenia Range, based on geophysical evidence (Aubert & Baroz 1978). The Kyrenia Range is restored as part of the former southern margin of Tauride-related continental units to the north (Robertson & Woodcock 1986). The Troodos is interpreted as having been thrust beneath this margin in latest Cretaceous time (Clube & Robertson 1986) in a similar manner to the eastern Tauride ophiolites discussed earlier. The same explanation for the halting of subduction may apply; that is, the arrival of young buoyant SSZ-type crust at a trench. Furthermore, to the west and SW the Troodos is tectonically juxtaposed with Mesozoic rocks of continental margin-oceanic affinities, represented by the Mamonia Complex (Lapierre 1972). These Mamonia units (Dhiarizos and Ayios Photios Groups; Swarbick & Robertson 1980) can be restored to a position to the NE of the present position of the Troodos ophiolite. This inference is partly based on lithological correlation of deep-sea sediments (e.g. Early Cretaceous quartzose sandstones) exposed in the Mamonia Complex (Akamas Sandstones) with counterparts in the southern part of the Antalya Complex (Kumluca area) on the Turkish mainland (Robertson & Woodcock 1982b). The Mamonia Complex is unlikely simply to represent a continental margin located to the south, as there
E. MEDITERRANEANOPHIOLITE EMPLACEMENT is no evidence of southward emplacement of the Troodos ophiolite onto a continent to the south of Cyprus in latest Cretaceous time (see Dilek & Flower 2003). The first evidence of collisional deformation is, instead, related to collision of the Eratosthenes Seamount with the subduction trench in Plio-Quaternary time (Robertson 1998, and references therein). Deformation of the Mamonia Complex took place in latest Cretaceous time by a combination of thrusting, strike-slip faulting and gravity tectonics (e.g. Robertson & Woodcock 1979; Swarbrick 1980; Malpas et al. 1992; Bailey et al. 2000; see Robertson & Xenophontos 1993, for review). The emplacing Mamonia terrane was sealed by debris-flow deposits (Kathikas Mrlange) shed from the Mamonia continental margin-oceanic units during latest Cretaceous time (Swarbrick & Naylor 1980) while still in an oceanic setting. This setting is confirmed by the cover of deep-sea carbonate sediments of latest CretaceousEarly Tertiary age (Urquhart & Banner 1994; Lord et al. 2000). Gravity emplacement is, additionally, exemplified by the Moni Mrlange, southern Cyprus, in which occur large blocks of, for example, Cretaceous continentally derived sedimentary rocks, including proximal quartzose sandstones (Parekklisha Sandstone) and deeper-water siliceous pelagic limestones (Monagroulli Limestones). These exotic units were emplaced in a matrix of Late Cretaceous deep-sea clays (Moni Clay) that depositionally overlie the Troodos ophiolite (Robertson 1977; Gass et al. 1994; Urquhart & Robertson 2000). The deformation of the continental margin units exposed in Cyprus (Mamonia Complex and Kyrenia Range) is, therefore, closely related to the palaeorotation of the Troodos microplate and differs markedly from that seen in adjacent areas of the southern Neotethys. The question remains as to whether the palaeorotation of the Troodos ophiolite relates more to the tectonic evolution of the active Tauride margin to the north, or to the emplacement of ophiolites along an embayment of the North African-Arabian passive margin to the south (as in the pivoting slab model). Much depends on the width of the southern Neotethys remaining by latest Cretaceous time when the palaeorotation began. According to evidence from palaeomagnetic inclinations, the Troodos ophiolite originated at c. 20~ closer to Gondwana than Eurasia (Morris 2004). The Troodos and other southern Neotethyan ophiolites (e.g. Hatay, Baer-Bassit, Kocali) presumably migrated southwards as northward subduction proceeded and Arabia-North Africa drifted northwards. Indeed, the surviving southern Neotethys was probably not much more than a few hundred kilometres wide by latest Cretaceous time, as there is little evidence of arc volcanism associated with the latest stages of Tertiary destruction of the southern Neotethys (Akta~ & Robertson 1984; Robertson et al. 2006, 2007). Remnants of Late Cretaceous oceanic crust, chemically of subduction-related affinity (Floyd et al. 1992), remained within the remnant southern Neotethys. These units were only finally accreted to the Tauride active margin to the north after Late Eocene time, based on evidence from the Misis-Andlnn Complex in coastal southern Turkey (Robertson et al. 2004). It is, therefore, probable that the present scale of Cyprus represents a large proportion of the width of the southern Neotethys ocean that remained to the east of the north-south Levant embayment by latest Cretaceous-Early Tertiary time. If so, the palaeorotation of the Troodos microplate could have been triggered by collision of the subduction trench with the irregular southerly ArabiaNorth African continental margin, while also involving units associated with the northerly, active margin of the southern Neotethys (i.e. Kyrenia Range and Mamonia Complex). In summary, the well-known palaeorotation of the Troodos microplate is an example of displacement of an ophiolite while still within the ocean, but without involving emplacement onto a continental margin. This, in turn, re-emphasizes the diversity of the processes of ophiolite emplacement and displacement exemplified within the Eastern Mediterranean region.
253
Conclusions The Oman-type trench-margin emplacement model can be applied to a wide range of Mid-Jurassic and Late Cretaceous ophiolites throughout the Eastern Mediterranean region (Fig. 25a). In this interpretation, most of these ophiolites were generated above oceanward-dipping subduction zones. These subduction zones consumed Neotethyan oceanic lithosphere, creating an accretionary prism of deep-sea sediments and volcanic rocks, until the subduction trench collided with a passive margin, either a continent or a microcontinent. This margin was then flexurally loaded and collapsed, facilitating final emplacement of the ophiolites onto submerged former passive margins. The Oman-type emplacement model can be successfully applied to many of the ophiolites of the Eastern Mediterranean region. These include those of Mid-Jurassic age that were emplaced onto collapsed continental fragments, such as the Pelagonian carbonate platform and counterparts further NW along strike (e.g. Othris, Vourinos, Pindos-Mirdita and Dinaride ophiolites of Greece and former Yugoslavia). This model applies regardless of different interpretations of the locations of the root zones of these ophiolites. In addition, the Oman-type model applies well to the emplacement of Late Cretaceous ophiolites of Turkey, Cyprus, Syria and Iran. These ophiolites were emplaced separately, from two Neotethyan basins: a southern Neotethys to the south of the Tauride carbonate platform, and a northern Neotethys to the north of this inferred microcontinental unit. The mode of emplacement was influenced by the regional palaeogeography. In some areas the collided passive margin was subducted and metamorphosed under high-pressure conditions and soon exhumed. However, the ophiolites typically remained attached to the overriding upper plate and remained unmetamorphosed. In addition, ophiolites were also formed and emplaced along active continental margins during Late Triassic-Early Jurassic, Mid-Jurassic and Late Cretaceous time. These include 'Cordillerran-type' ophiolites formed in association with accretionary prisms (Fig. 25b) and collapsed marginal basins. These ophiolites, again mainly of SSZ-type, show considerable variation in emplacement style. Throughout the Balkan region, the northern margin of the Jurassic Vardar ocean experienced active margin processes, including subduction, marginal arc volcanism and back-arc basin opening. Oceanic crust formed in a marginal basin setting by rifting of the Eurasian margin and was later emplaced as the 'Inner Hellenide' ophiolites, but without evidence of large-scale horizontal transport. In addition, the pre-existing Late Palaeozoic-Early Mesozoic Palaeotethys, exposed in the Pontides of northern Turkey, shows an active margin history of subductionaccretion (e.g. Karakaya Complex), arc volcanism and back-arc opening during Late Palaeozoic-Triassic time. When this marginal ocean basin later closed, prior to Early Jurassic time, a dismembered ophiolite (Ktire ophiolite) was emplaced northwards onto the Eurasian margin. Contrasting settings of ophiolite emplacement characterize the northerly margins of the Jurassic and Cretaceous Neotethyan ocean basins. These were characterized by active margin processes including subduction, accretion, arc magmatism and opening of back-arc basins. Jurassic ophiolites in Greece (e.g. Guevgueli ophiolite) formed by rifting of the southern margin of Eurasia (Serbo-Macedonian zone), followed by emplacement without major horizontal transport over adjacent continental units. Late Cretaceous ophiolites also originated in a back-arc basin setting associated with the Eurasian margin in Turkey (Pontides). However, these ophiolites formed in a more oceanic setting, bordered by an intra-oceanic arc and forearc basin, and were later emplaced northwards over the Eurasian margin prior to Early Tertiary time.
254
A.H.F. ROBERTSON
Ophiolites were also emplaced along the northern, active margin of the southern Neotethys by a variety of processes in different geographical areas. The Late Cretaceous ophiolites of the SE Turkish thrust belt were accreted to the hanging walls of a northdipping subduction zone, where they were then intruded by Late Cretaceous calc-alkaline plutons related to continuing northward subduction. Where subduction was oblique or at a high angle to the adjacent margin the ophiolites and related units were emplaced as small 'terranes' by dominantly strike-slip, or transpressional processes. These ophiolites retain relatively intact vertical successions (e.g. Late Cretaceous Antalya ophiolites), rather than forming the lowangle thrust sheets that are typical of ophiolites emplaced by orthogonal overthrusting. The presence of a highly irregular palaeogeography remaining from the previous r i f t - p a s s i v e margin stage influenced the mode of ophiolite emplacement in the southern Neotethys. For example, the well-known 90 ~ anticlockwise palaeorotation of the Troodos microplate was possibly triggered by collision of the subduction trench with the Arabian promontory to the east. The southern Neotethys was by then relatively narrow and it is likely that both southern and northern margin units were affected by the regional palaeorotation. Many other ophiolites, commonly of a fragmentary highly deformed nature, are present in the wider Eastern Mediterranean area, including former Yugoslavia, Romania, Bulgaria, Armenia, Georgia and Iran. The discussion has covered many of the larger ophiolites in the Eastern Mediterranean region but has necessarily excluded many others. In future, it would be useful to determine whether the emplacement of these other ophiolites can be related to the two main settings of ophiolite emplacement discussed here: (1) ophiolites emplaced by t r e n c h - m a r g i n collision onto former passive margins (Oman-type model); (2) ophiolites emplaced along active margins undergoing subduction and arc volcanism. D. Baty assisted with drafting many of the figures, and Y. Cooper helped prepare the illustrations for publication. Helpful comments on this manuscript were received from J. Winchester and L. Jolivet.
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Eastern Mediterranean basin systems ZVI B E N - A V R A H A M 1, J O H N W O O D S I D E 2, E M A N U E L E L O D O L O 3, M I C H A E L G A R D O S H 4, M A R I O G R A S S O 5, A N G E L O C A M E R L E N G H I 3 & G I A N B A T T I S T A V A I 6
1Department of Geophysics and Planetary Sciences, Tel Aviv University, Tel Aviv, Israel (e-mail:
[email protected]) 2Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1082 HV Amsterdam, Netherlands 3Istituto Nazionale di Oceanografia e di Geofisica Sperimentale (OGS), Trieste, Italy 4Geophysical Institute of Israel, P.O. Box 182, Lod 71100, Israel 5Dipartimento di Scienze Geologiche, Corso Italia 55, 95129, Catania, Italy 6Dipartimento di Scienze della Terra e Geologico-Ambientali, Via Zamboni 67, 40127, Bologna, Italy
Abstract: The basins in the Eastern Mediterranean can be divided into those that were formed mainly in post-Miocene time and those
that were formed during the rifting episodes that led to the formation of the Neotethys. The younger basins can be further divided into those that were formed mainly in post-Miocenetime and those that were formed in post-Pliocene time. The separation is not only one of convenience but also corresponds to major adjustments in the plate tectonic situation in the Eastem Mediterranean. The late Miocene deposition of thick evaporites throughout the Mediterranean region, or, where evaporites are missing, the creation of an important erosional unconformity during the extreme lowstand of the Mediterranean, makes the Miocene-Pliocene boundary relatively easy to identify, especially on seismic reflection records. At about the same time, following the collision of the Arabian plate with Eurasia, the Anatolian and Aegean microplates came into existence between the convergent African and Eurasian plates to accommodate tectonic escape between them. The general configurationof the Eastern Mediterraneanbasins reflects the tectonic and structural gradients between the collisional domain of southeastem Turkey and Iran, and the continuing but increasingly limited subductionalong the Calabrian and Hellenic arcs, with the Cyprus and Levantine zones between them. Several distinct zones can be identified in the Eastern Mediterranean. The Dead Sea Fault system marks the edge between the collisional and pre-collisional zones to the east and west, respectively. The meridian through the Anaximander Mountains (30~ forms a rough boundary between the zone of incipient collision to the east and the zone of continuing but late-stage subduction to the west. The Malta Escarpment forms the Eastern boundary of the Eastern Mediterranean basins. The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea share this progressive evolution, with those containing Messinian evaporites to the east and those without to the west. The Sicily Channel with its associated basins is an extensional zone between the Eastern and Western Mediterranean. The basins discussed in this paper are divided into two groups, the larger and older basins and the smaller and younger basins. In the first group are the Ionian Basin and the Levantine Basin, and in the second group the Cilicia Basin, Antalya Basin, Finike Basin, Rhodes Basin, Aegean basins, Sicily Channel basins, Latakia Basin and Larnaca Basin. The Eastern Mediterranean represents the last stage in the evolution of an ocean basin. Given the current motion between Africa and Eurasia, the Eastern Mediterranean will cease to exist in about 6-8 Ma from now. As a result, the larger and older basins are shrinking, whereas the younger and smaller basins are growing. Eventually the smaller basins will also disappear.
The Central and Eastern Mediterranean is dominated by the convergence of the African plate with the Eurasian plate. The relative motion between these two plates has produced, after the closure of an oceanic-type domain, a complex system of contractional structures (i.e. the Cyprus, Hellenic, Calabrian and Maghrebian arcs) along which the stress field is being dissipated. There are very few places in the world where the o c e a n - c o n t i n e n t crustal transition at a passive continental margin is approaching a subduction zone. A thick pile of sediments as old as Mesozoic is being deformed at the collisional margin (Mediterranean Ridge), although the irregular shape of the colliding continental margins leaves portions of the former basin still relatively u n d e f o r m e d (Ionian and Herodotus basins). In the Eastern Mediterranean, the transition between the remnants of a thick Mesozoic oceanic crust located in the Ionian Basin (DeVoogd et al. 1992) and in the Levantine Basin (Ben-Avraham et al. 2002) and the African continental margin (Fig. 1) is approaching the Calabrian, Aegean and Cyprus subduction zones (Dewey et al. 1973; McKenzie 1978). The relative plate motion between Africa, Eurasia and the Aegean microplate produces convergence. The African plate moves in an approximately northward direction at a rate of 1 cm a -1 (Rebai et al. 1992). The Aegean region undergoes extension at a rate of 3 . 5 4.3 c m a -1 in a SSW direction relative to Africa (Le Pichon et al. 1995). A large doubly vergent accretionary complex, the eastern Mediterranean Ridge, forms in a complicated setting of altemation of zones of continental collision (the Pelagian and Cyrenaica promontories) and zones of terminal subduction (the
Ionian Basin). The irregular shape of the converging plate boundaries causes diachronous continental collision, strain partitioning and lateral escape of both the Ionian and Levantine inner zones (Finetti 1976; Chaumillon & Mascle 1995; Le Pichon et al. 1995; Polonia et al. 2002). The southern, A f r i c a n - A r a b i a n , continental margins were considerably less affected by the closing of the basins and remained more or less stable in their position near the present-day coastal areas of northern Africa, Sinai, Israel and Lebanon (Bein & Gvirtzman 1977; Garfunkel 1998). The structure of the southern Levant margin and the adjacent deep marine basin is relatively simple, and the Mesozoic and Cenozoic rock record found on land and in the sea is more or less continuous. Both structure and stratigraphy preserve the signature of the main tectonic events that shaped the basin and its margin. These can be separated into four distinct tectonostratigraphic phases. Similar tectonostratigraphic phases are observed in the Hyblean continental margin in eastern Sicily and the adjacent deep Ionian Basin. The young basins of the Eastern Mediterranean can be divided conveniently into those that were formed mainly in post-Miocene time and those that were formed in post-Pliocene time. The separation is not only one of convenience but also corresponds to major adjustments in the plate tectonic situation in the Eastern Mediterranean. The late Miocene deposition of thick evaporites throughout the Mediterranean region (Hsu et al. 1973a,b) or, where evaporites are missing, the creation of an important erosional unconformity during the extreme lowstand of the Mediterranean, makes the M i o c e n e - P l i o c e n e boundary relatively
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 263-276. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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264
Z. BEN-AVRAHAMETAL.
Fig. 1. Map of the Eastern Mediterranean region. Bathymetriccontours are at a 500 m interval; circles with numbers mark the location of basins, and squares with letters mark other structures.
easy to identify. At about the same time, following the collision of the Arabian plate with Eurasia, the Anatolian and Aegean microplates came into existence between the convergent African and Eurasian plates to accommodate tectonic escape between them (McKenzie 1972; Angelier et al. 1982). The general configuration of the Eastern Mediterranean basins reflects the tectonic and structural gradients between the collisional domain of southeastern Turkey and Iran (McClusky et al. 2000), and the continuing but increasingly limited subduction along the Calabrian and Hellenic arcs, with the Cyprus or Levantine zone between. Just as the Dead Sea Fault system marks the edge between the collisional and pre-collisional zones to the east and west, respectively, the meridian through the Anaximander Mountains (30~ forms a rough boundary between the zone of incipient collision to the east and the zone of continuing but late-stage subduction to the west (Ben-Avraham & Nur 1976; Nur & Ben-Avraham 1978). The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea share this progressive evolution, with those containing Messinian evaporites to the east and those without to the west. In the western part of the Eastern Mediterranean within the regional convergent system, NW-SE-trending troughs have developed in the Strait of Sicily. These depressions represent relatively young (late Miocene to Present) extensional structures that cut across the undeformed Pelagian block, a continental crust that represents a promontory of the African plate margin. In this paper we briefly describe the large and small basins in the Eastern Mediterranean. We also describe the main tectonostratigraphic stages and review the evolution of the Ionian and Levantine basins as the result of the large-scale plate motions that took place in the Eastern Mediterranean region during the Mesozoic and Cenozoic.
Older basins Ionian Basin
The Ionian Sea is a deep marine basin separated by the conspicuous Malta escarpment from the shallow, Hyblean-Malta Plateau on the west. Most researchers agree on the oceanic nature of the
Ionian lithosphere, although its uppermost layer is made of up to 10 km thick sedimentary cover. Its age, however, is debatable, with suggestions ranging from Early Permian (Ben-Avraham & Ginzburg 1990; Catalano et al. 1991; Vai 1994, 2003), to Late Permian to Early Triassic (Stampfli et aI. 2001b), Triassic (Finetti 1984), mid-Jurassic or broadly Mesozoic (Robertson & Grasso 1995; Cantarella et al. 1997), Cretaceous (Dercourt et al. 1986) or even Messinian (Fabricius & Hieke 1977). Evidence of deepening and/or accelerated subsidence of the continental margin sequences around the long-lasting Sicilian Sicani Basin, before Tertiary and Quaternary deformation took place, are known during mid-Cretaceous, Late Triassic to early Jurassic, Mid-Triassic, and especially Early to Mid-Permian times (Charier et al. 1988; Catalano et al. 1991, 1992; Vai 1994, 2003; Dercourt et al. 2000; Stampfli et al. 2001a,b). This evidence is consistent with rift pulses producing WNW-ESE-trending basins. The Early Mesozoic rifting events are similar to those in other domains in the Central and Eastern Mediterranean, and mark a major continental break-up followed by rapid tectonic subsidence that initiated the opening of the Neotethyan ocean (Yellin-Dror et al. 1997). This early evolutionary stage of the Hyblean margin was followed by slow thermal subsidence (Early JurassicLate Cretaceous), northward movement and thrusting (Late Cretaceous-Palaeogene), and continued convergence, uplift and subsidence (Neogene-Quaternary; Yellin-Dror et al. 1997). The Ionian Basin, of which the crustal structure is shown in Figure 2 and a seismic section in Figure 3, is located near the orthogonal plate convergence zone. In the western Ionian Basin plate convergence is nearly orthogonal in a N E - S W direction. Because of the curved shape of the Mediterranean Ridge deformation front, however, a variable plate convergence angle has to be expected, especially in the Messina foredeep where the deformation front approaches a north-south direction. Thrusting and folding in a direction roughly parallel to the Ionian rigid crustal backstop are reflected in shortening and uplift on the Mediterranean Ridge accretionary complex, which results in the tectonic addition of post-Messinian material to the outer deformation front. The main regional detachment surface is located at the base of the Messinian evaporites at the outer wedge. The evaporites, of variable thickness and lateral extent, were deposited during
EASTERN MEDITERRANEANBASIN SYSTEMS
Fig. 2. Main structural and morphological elements of the Ionian Basin. 1, Alpheus seamount; 2, Malta escarpment; 3, Medina seamounts; 4, Messina abyssal plain (Messina Foredeep); 5, Victor Hensen seahill; 6, Medina-Victor Hensen structure (Hieke & Dehghani 1999), including the Medina-Cephalonia line; 7, Sirte abyssal plain (Sirte Foredeep); 8, Bannock buried seamount and related tectonic lineament; 9, Cyrene seamount; 10, Hellenic trench system on the continental backstop of the Mediterranean ridge accretionary complex. The doubly vergent structure of the Mediterranean ridge accretionary complex, identified by an outer deformation front (line with filled) and an inner deformation front, should be noted. Asterisk marks position of section shown in Figure 3. the Messinian (late Miocene) desiccation of the Mediterranean Sea (Hsu et al. 1973a,b). The rheology of the deforming materials is expected to have undergone drastic changes in the last 5 Ma (Kastens et al. 1992; Kukowsky et al. 2002; Reston et al. 2002a,b), permitting the identification of pre- and post-Messinian wedges. This has led to the post-Messinian development of an accretionary complex whose features are similar to those of salt-bearing fold-and-thrust belts (Davis & Engelder 1985). Below the inner portions of the wedge the detachment becomes progressively deeper, and cuts into the top of the Mesozoic carbonates belonging to the African foredeep and foreland (Cita e t al. 1981; Camerlenghi et al. 1995; Chaumillon & Mascle 1995, 1997). Thrusting to the NE of accreted sediments above the Ionian rigid backstop (Le Pichon et al. 2002) occurs along the inner deformation front. The Sirte and Messina foredeep areas remain completely undeformed. The structural style is affected by geological structures rooted in the incoming continental crust of the African plate. There are SW-NE-aligned isolated structural highs found either buried in the accretionary wedge or having morphological expression as seamounts in the western
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Fig. 3. Seismic section in the Ionian Basin (modified after Cernobori et al. 1996). The stratigraphic interpretation follows regional correlation with main seismic boundaries tied to known wells. The definition of the oceanic basement is correlated with the results of expanded spread seismic profiles (DeVoogd et al. 1992). The location of the section is indicated by a bold asterisk in Figure 2. Similar seismic images of the Ionian foredeep have been obtained by Polonia et al. (2002). foredeep and foreland (Della Vedova & Pellis 1989; Von Huene et al. 1997; Hieke & Dehghani 1999; see Fig. 3). In the northeastern Ionian basin, plate convergence is controlled by the fast rates of N W - S E extension of the Tyrrhenian Sea and associated southeastward migration of the Calabrian-Peloritan arc. The directions of motion of the Africa plate and the Calabrian-Peloritan arc therefore differ by about 90 ~ Deformation is developed offshore as a series of imbricated thrusts involving the entire sedimentary section (Finetti 1976; Cernobori et al. 1996) with detachment cutting into progressively deeper stratigraphic levels, from the Messinian salts at the outer deformation front to Mesozoic rocks and the crystalline basement in the continental blocks of the Calabrian arc in southern Italy (Tortorici 1982; Tortorici et al. 1995). Similarly to the case of the Mediterranean Ridge accretionary complex, it can be inferred that the introduction of a ductile salt layer in the deforming sequence has modified drastically the rheology of the arc. Strain partitioning induced by obliquity of subduction must exist only in the pre-Messinan deforming domain. The viscous behaviour of the detachment surface has produced thrusting trending normal to the shortening direction (thus the high curvature of the deformation front), at much higher rates of outward propagation (Costa & Vendeville 2002; Costa et al. 2004).
266
z. BEN-AVRAHAMETAL.
Levantine Basin Regional setting. The present-day southeastern Mediterranean Sea,
also known as the Levantine Basin (Fig. 4), is a remnant of a larger, Neotethyan oceanic basin that opened between several fragments of the Pangaea supercontinent in Early Mesozoic times (Dewey et al. 1973; Bein & Gvirtzman 1977; Garfunkel & Derin 1984; Robertson & Dixon 1984; Garfunkel 1998). During the M i d - L a t e Cretaceous, as the basin started to close, its northern, Eurasian part underwent a dramatic transition from a passive to an active margin, and the basin was subsequently largely subducted or accreted at the present-day areas of Cyprus, southern Turkey, northern Syria and Iran (Ben-Avraham 1989; Robertson 1998). Closing of the northern Levantine Basin continued through the Cenozoic; the active convergent front is at present located at the area of the Cyprian arc (Kempler & Ben-Avraham 1987; Ben-Avraham et al. 1995; Fig. 4).
Fig. 4. Tectonic map of the Eastern Mediterranean area (modifiedafter Garfunkel (1998) and Robertson (1998), and incorporating results from Ben-Avraham et al. (1995) and Woodside et al. (2002). The Levantine Basin is a remnant of a larger, Early Mesozoic (Neotethyan) oceanic basin that was largely consumed during Late Mesozoic-Cenozoic convergence of the African-Arabian and Eurasian plates. Subduction and wrenchingis taking place at the tectonically active, northern part of the basin (Cyprian arc), whereas only minor thrusting and contraction (Syrian arc folds) is recorded on the southeastern, Levantine margin. The Eratosthenes Seamount is a large continental block that was stranded within the basin and is now approaching its northern margin. The Carmel fault (CF) is an active tectonic lineament that separates two crustal types, and probably originated during the Early Mesozoic rifting phase.
The structure of the southern Levant margin and the adjacent deep marine basin is relatively simple and preserves the signature of the main tectonic events associated with opening and closing of the Neotethyan ocean. These can be separated into several distinct tectonostratigraphic phases: (1) Triassic to Early Jurassic continental break-up and rifting; (2) Mid-Jurassic to MidCretaceous subsidence and formation of passive continental margin; (3) M i d - L a t e Cretaceous to Early Tertiary large-scale convergence and contraction; (4) Late Tertiary to Quaternary minor convergence and subsidence. Triassic - E a r l y Jurassic rifting phase. During the Palaeozoic the area of the Levantine Basin was located south of the Palaeotethys Ocean on the northern edge of the Gondwana continental platform. Most workers agree that a break-up of the northern part of Gondwana into several microplates occurred in the latest Palaeozoic to Early Mesozoic (Bein & Gvirtzman 1977; Hirsch et al. 1995; Garfunkel 1998). Continental break-up processes are indicated by two types of observations: (1) rift-related phenomena identified in wells and multi-channel, seismic reflection profiles; (2) variation in crustal thickness and composition interpreted from deep seismic refraction profiles, gravity and magnetic data. An extensive graben and horst system is recognized in the deeper stratigraphic level of central Israel and the Levant margin. Part of this system is observed in multi-channel, seismic reflection profiles (Fig. 5; Gardosh 2002; Gardosh & Druckman 2006). Early Mesozoic structural highs and lows are further interpreted from thickness variations in the Permian to Lower Jurassic strata, identified in seismic and well data (Garfunkel & Derin 1984; Bruner 1991; Druckrnan et al. 1995; Garfunkel 1998). The system is composed of fault lines several tens of kilometres long that are generally oriented N E - S W , roughly perpendicular to the present-day coastline of Israel (Fig. 4). An early tectonic pulse of this fault system is indicated by the northward thickening trend of the Permian to Lower Triassic section, identified in wells at southern and central Israel (Garfunkel & Derin 1984). The continuation of faulting activity is evident from the occurrence of a 350 m thick, polymictic and angular, Middle Triassic carbonate breccia discovered in wells near the southern coastal plain of Israel (Druckrnan 1984). A conspicuous graben, filled with Carnian dolomite and gypsum, was identified in outcrops and wells in central and southern Israel (Druckman 1974; Garfunkel & Derin 1984). Extensive series of the Lower Liassic Asher volcanic rocks, found in the subsurface of northern Israel, were partly accumulated within a 2.5 km deep trough (Gvirtzman & Steinitz 1983; Garfunkel 1989). Abrupt thickness changes are not recognized in Bathonian and younger strata of the Levant margin, suggesting the cession of activity of the Neotethyan graben and horst system (Garfunkel 1998). The upper age limit of the faulting is further indicated in some of the structures where Triassic and older tilted beds are overlain by horizontal Lower to Middle Jurassic strata (Gardosh 2002; Gardosh & Druckman 2006; Fig. 5). The Early Mesozoic rifting and extension resulted in profound changes in crustal composition and thickness. Long-range seismic refraction profiles, gravity and magnetic data show that a major transition between two types of crust takes place several tens of kilometres NW of the present-day coastline of Israel (Ben-Avraham et al. 2002). A 30-35 km thick crust of continental character underlies the area of southern and central Israel. The SE Mediterranean Sea is underlain by a 10 km thick crust of an oceanic character (Ginzburg et al. 1979; Ginzburg & Folkman 1980; Makris et al. 1983; Ginzburg & Ben-Avraham 1987; Ben-Avraham et al. 2002). A 2 5 k i n thick crust of continental character underlies the Eratosthenes Seamount, located south of Cyprus (Fig. 4; Makris et al. 1983; Ben-Avraham et aL 2002). These observations indicate that a major part of Early Mesozoic rifting and extension took place west of the present-day coastline. This activity resulted
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Fig, 5. Multi-channel, migrated, marine seismic reflection profile across the Levant margin showing the main stratigraphic units and faults. The line was shot in 1983, offshore Israel, using seven Bolt air guns (2180 cubic inches) and was recorded with 120 channels on a 2975 m cable. The interpreted seismic horizons are regional seismic markers, most of which are correlated to offshore wells: I, top of Messinian evaporites; II, base of Messinian evaporites; III, Mid-Tertiary unconformity; IV, Upper Cretaceous; V, Mid-Cretaceous unconformity; VI, near top Lower Jurassic, VII, near top Basement. The four major tectonostratigraphic stages identified on the southeastern margin of the Levantine Basin are: Triassic to Early Jurassic continental break-up and rifting (horizons VII-VI); Mid-Jurassic to Mid-Cretaceous subsidence and formation of passive continental margin (horizons VI-V); Mid-Late Cretaceous to Early Tertiary large-scale convergence and contraction (horizons V-III); Late Tertiary to Quaternary minor convergence and subsidence (horizon III-sea bottom). Location of section is shown in Figure 4. YW1, Yam West-1 well.
in the separation of the Tauride block of southern Turkey from the African continental mass (Smith 1971; Ben-Avraham 1989; Robertson et al. 1991, 1996; Garfunkel 1998). Additional tiffing, extension, creation of oceanic crust and development of deep marine basins took place when the Eratosthenes block separated from the African continent, resulting in the creation of 2 0 0 - 3 0 0 km wide deep marine basin in the southern Levant area (Ben-Avraham 1989; Garfunkel 1998). The maximum size of the palaeo-Levantine Basin, extending south of the Tauride block and north of the African-Arabian continental mass, is hard to estimate, as much of its northern part was later consumed. Garfunkel (1998) estimated that it was at least twice its present width. Another pronounced transition of crustal properties is identified across the Carmel fault on the Levant margin (Fig. 4). The 35 km thick continental crust of central Israel thins to about 20 km in the Galilee area, NW of the fault (Ginzburg & Folkman 1980). The Carmel fault is an active tectonic lineament; its origin is associated with Palaeozoic-Mesozoic plate motion (Ben-Avraham & Ginzburg 1990; Ben-Gai & Ben-Avraham 1995). The relationship between the fault and the rifting processes that led to the formation of the Levantine Basin are not clear. Mid-Jurassic-Mid-Cretaceous continental margin phase. During the later part of the Mesozoic, subsidence and the development of passive continental margins dominated the Levant region (Bein & Gvirtzman 1977). The M i d - U p p e r Jurassic strata of the southern margin show no evidence of large-scale faulting and magmatic activity, indicating that the opening motion in the Levantine Basin was considerably reduced. A shallow marine shelf was developed near the present-day coastline, whereas a deeper marine basin prevailed throughout the SE Mediterranean Sea. The Levant margin evolved through the accumulation of various types of low- and high-order depositional sequences; their geometry and stratal pattern reflect the combined effects of eustatic sealevel cycles, local subsidence and uplift, change of environmental conditions and rate of sediment supply (Flexer et al. 1986; Gardosh 2002). The Lower to Middle Jurassic depositional cycles are characterized by rapid growth and aggradation of carbonate margin and relatively minor bypass into the basin, reflecting fast thermal subsidence coupled with long-term eustatic rise. The uppermost Jurassic to Lower Cretaceous depositional cycles are dominated by intense transport of siliciclastic strata
and accumulation of deep-water turbidites in the basin and margin. These are associated with tectonic uplift and erosion SE of the basin (Gvirtzman et al. 1998) coupled with eustatic sealevel drop. Finally, the Middle Cretaceous depositional cycles are characterized first by the progradation of carbonate slopes into the basin, followed by intense growth of carbonate platforms on the margin and restricted bypass into the basin. These were the results of tectonic quiescence, followed by marked eustatic rise during the Cenomanian-Turonian (Gardosh 2002). Late Cretaceous-Early Tertiary main convergence phase. Closing of the Levantine Basin started during the late Mid-Cretaceous and continued through the Late Cretaceous and Early Cenozoic. The relative motion of the African-Arabian plate towards the Eurasian plate resulted in the development of several subduction and collision zones in the northern part of the basin, in the present-day areas of the Cyprian arc and the Taurus Mountain belt (Fig. 4; Ben-Avraham & Nur 1986; Ben-Avraham 1989; Robertson et al. 1991). The most conspicuous tectonic element associated with this motion is a series of contractional structures found throughout the Levant region, termed the Syrian arc folds (Krenkel 1924; Fig. 4). Well and seismic data from the inland part of Israel show that reverse faults are found in the core of many of the Syrian arc folds. These structures are often superimposed on the older graben and horst system, and their formation is controlled by the reactivation, in a reverse direction, of Early Mesozoic normal faults (Freund et al. 1975; Druckman et al. 1995). Seismic data from the Israeli offshore reveal more details on the relation between the Levant margin and the contractional deformation. A belt of dense Syrian arc type folds and reverse, upthrust faults is observed at the southeastern edge of the basin (Figs 4 and 5). The folded zone terminates some 5 0 - 7 0 km west of the present-day coastline. The entire Precambrian to Mesozoic sequence in the offshore fold belt is uplifted, and the lowered area to the west contains a several kilometres thick sedimentary section of assumed Late Cretaceous to Early Tertiary age (Fig. 5). The western edge of the fold belt coincides with a zone of transition between two types of crust identified by Ben-Avraham et al. (2002). Based on the geophysical evidence it is suggested that the western edge of the fold zone marks an area of convergence and possibly thrusting of an oceanic or transitional type crust under continental crust on the east. Ben-Avraham & Nur (1986) have
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suggested a similar process along the continental margin of the Sinai Peninsula. A deep marine flexural foredeep developed west of the thrust front where large amounts of detrital carbonate strata were accumulated (Figs 4 and 5). Late Tertiary-Pleistocene minor convergence phase. Contractional
deformation in the southern Levant became rather limited during the last tectonostratigraphic phase (Fig. 5), indicating that the rate of convergence was reduced. A possible cause for this reduction is the emergence of a new plate boundary that developed east of the Levant margin contemporaneously with the opening in the Red Sea (Freund et al. 1970). Part of the northward motion of the African plate was probably taken up by this new boundary, the Dead Sea transform zone (Fig. 4). The activity along the Dead Sea plate boundary during the Miocene-Pliocene was also associated with considerable uplift and erosion of the transform shoulders (Garfunkel 1989). This was followed by the development of vast drainage systems that carried large amounts of detrital material westward and northward across the Levant margin. Three distinct depositional cycles are recognized during the Late Tertiary-Pleistocene tectonic stage (Fig. 5). The lower, Oligocene-Miocene cycle is characterized by intense transport of siliciclastic strata into the basin. The middle, uppermost Miocene cycle is associated with the accumulation of thick evaporitic section within the basin. The Messinian salinity event is well recognized throughout the Eastern Mediterranean. The deposition of the Messinian salt within Mediterranean basins is considered to be the result of a dramatic sea-level drop, in the range of 800-1300 m (Ben-Gai et al. 2005), as a result of the isolation of the Mediterranean Sea from the Atlantic Ocean (Hsu et al. 1973a,b; Ryan 1978). Quantitative basin analysis of the Levant margin suggests the existence of a deep basin, similar to the present one, in pre-Messinian time (Tibor et al. 1992). A final Plio-Pleistocene cycle is characterized by the development of conspicuous, prograding delta systems along the western margin as well as the formation of the Nile river delta further to the south (Figs 4 and 5). These systems were associated with anomalously high sedimentation and subsidence rates, influenced by flexural response of the lithosphere to the loading of the Messinian salt and Nile-derived sediments as well as to uplift of the Judea Mountains east of the margin (Tibor et al. 1992).
The present-day active northern edge of the Levantine Basin is located along the Cyprian arc in the NE Mediterranean (Fig. 5). The Cyprian arc is divided into three distinct segments (Kempler & Ben-Avraham 1978; Ben-Avraham et al. 1995). In the western segment subduction of the the African lithosphere under the Turkish plate is assumed to have led to the creation of the small Antalya Basin (Fig. 4). In the central segment subduction was interrupted, as a result of the presence of the Eratosthenes continental block in front of the arc. The eastern segment is a system of wrench faults dominated by shear motion, with no active subduction (Ben-Avraham et al. 1995). The small Latakia and Larnaca basins (see below) were formed in this segment of the arc (Fig. 4).
Younger basins The Strait o f Sicily rift s y s t e m
The morphological and structural evidence of elongated bathymetric lows in the Strait of Sicily (Pantelleria, Malta and Linosa troughs; Figs 1 and 6) were mapped during the early exploration of the Mediterranean Sea (Finetti & Morelli 1972, 1973). The interpretation of these troughs as rift-related structures was proposed by various workers (lilies 1981; Finetti 1984), and some of them have emphasized the role played by transcurrent tectonics in their development (Cello et al. 1984; Jongsma et al. 1985; Reuther & Eisbacher 1985; Boccaletti et al. 1987; Cello 1987). The depressions found within the Pelagian block are viewed in general as pull-apart transtensional basins generated in a dextral wrench system. The proposed interpretations, mainly based on structural analyses carried out in the islands within the Strait of Sicily or in restricted parts of the surrounding areas, present some differences, mostly related to the poorly constrained stretching mechanisms and deformational history. The morphostructural features present in the Strait of Sicily do not have an equivalent in the central and eastern regions of the Mediterranean Sea, where major changes in crustal nature, structural trends and tectonic styles occur within short distances. In the Strait of Sicily, a prevalent extensional regime dominated from late Messinian to early Pliocene time, as determined from the analyses of the seismic sequences found in the depocentral
Fig. 6. Bathymetricmap of the Strait of Sicily obtained by combiningunder-way soundings and depth-converted sea-floor reflectors from a seismic grid. The parts of seismic profiles MS-19 and MS-120 that are presented in Figures 7 and 8, respectively, are indicated by bold continuous lines. The main troughs associated with the rift system (Pantelleria, Malta and Linosa) are indicated by grey shading.
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269
areas of the troughs, and was responsible for the development of the three main depressions within the Pelagian block. Regional structural setting. The Strait of Sicily rift system is located within the Pelagian block (Burollet et al. 1978; Boccaletti et al. 1984), which geologically corresponds to the northern
leading edge of the African plate. The stretched continental crust thins to less than 20 km (Colombi et al. 1973) underneath the Pantelleria trough. The sedimentary cover is made up of MesoCenozoic carbonate sequences that crop out in both the northern (southeastern Sicily) and southem (Tunisia and western Libya) domains of the platform. To the east, the shallow Pelagian block is separated from the deep Ionian Basin by the Malta escarpment, a structural boundary where deep-seated faults have throws exceeding 2000 m. The escarpment extends over a length of about 300 km from the eastern coast of Sicily southwards to the Medina seamounts with a steep slope that descends to more than 3000 m below sea level. It separates the 23 km thick continental crust (Finetti 1984) of the Strait of Sicily (Hyblean-Malta Plateau) from the 13 km thick crust beneath the Ionian Sea (Ferrucci et al. 1991). On land, normal and strike-slip faults formed by left-lateral movements exist on the Ionian side of the Hyblean Plateau, where they represent second-order structures superimposed on the dominant vertical displacement along the Malta escarpment. Sinistral wrenching along NNW-SSE-oriented fault planes is accommodated by extension along the N E - S W oriented faults. The contrasting structural behaviour of the Hyblean Plateau and the Ionian crust to the east could account for the kinematics of the Plio-Pleistocene structures along the onshore extension of the Malta escarpment (Grasso 1993). Three subparallel principal troughs (Fig. 6), roughly trending Nll0~ and deeper than 1000 m, form the Strait of Sicily rift system. The northwestern basin of the rift system (Pantelleria trough) is the widest depression, and is separated from the other two troughs (Malta and Linosa troughs) by localized structural highs at around longitude 12~ However, the geometric relationships between these depressions and their relative boundaries are not yet precisely identified because the lack of detailed bathymetric information. These depressions are filled with turbiditic Plio-Pleistocene deposits (Maldonado & Stanley 1977; Biju Duval et al. 1985), reaching thicknesses of about 1000 m in the Pantelleria trough, 2000 m in the Linosa trough, and 1500 in the Malta trough (Winnock 1981). In the other parts of the Pelagian block, the Plio-Pleistocene sequence has a maximum thickness of about 500 m. Geometry of the rift system. The basic information from which we derive the morphostructural and sedimentary setting of the Strait of Sicily region is the bathymetric map and the grid of seismic reflection profiles collected in the region since the early 1970s, calibrated with boreholes and well data. Seismic profiles provide the most striking evidence for analysis of the structural configuration and geometric disposition of the fault-related structures associated with stretching. After the earlier work of Finetti & Morelli (1972, 1973), a large amount of data have been collected in the Pelagian block, mainly in the course of a series of oil exploration surveys. Here we analyse the reprocessed, migrated version of seismic line MS-19 (Fig. 7), originally presented by Finetti (1982). This profile illustrates in detail the structural architecture of a continent-type crust affected by extensional tectonics. It runs N N E - S S W from the offshore part of the Hyblean Plateau, which constitutes the northern sector of the African foreland in the Pelagian block, to the Lampedusa Plateau, a generally flat-lying carbonate platform that borders to the SE the stretched region of the Strait of Sicily. The seismic profile, located between the two horsts constituting the islands of Malta and Lampedusa, crosses the eastern part of the Malta trough and the central-eastern Linosa trough. The stratigraphy along the seismic profile has been derived from correlation with adjacent well logs and
Fig. 7. Above: part of a reprocessed seismic line MS-19, originallypresented by Finetti (1984). This part of the profile crosses the central-eastern portion of the Linosa trough (see location in Fig. 6), and shows the geometry of the rift-related structures of the Pelagian block. The subvertical faults that progressively deepen the trough symmetrically should be noted. TWT, two-way travel time. Below: line drawing interpretation in which only the principal subvertical faults are indicated. (For the seismostratigraphic control along the profile, see Finetti (1984).) boreholes, and extrapolating the information across the entire grid of data (Finetti 1984; Pedley et al. 1993). Here we utilize the same interpretation in terms of seismo-stratigraphic control and sedimentological character of the sequences. In the Linosa trough a set of mostly subvertical, equi-spaced faults, separating rotated and, in some cases, uplifted blocks, dominate the structural framework along seismic line MS-19 (Fig. 7). Block rotation is particularly evident in the central sector of the Linosa trough. The core of the trough, corresponding to the deepest part of the system, is bounded on both sides by prominent faults with opposite polarities. All the subvertical faults reach the sea floor, indicating recent tectonic activity. Towards the SW, within the flat Lampedusa Plateau, another significant graben (the Lampedusa trough), composed of at least two tilted blocks, marks the southern boundary of the rift zone within the Pelagian block. The Malta trough is imaged by the profile MS-120 presented in Figure 8. The two basin shoulders are symmetrical, and possibly are composed of single normal faults. However, the strong erosion and the possible presence of localized slide structures do not allow clear identification of subsidiary discontinuities on the hanging walls. The total throw along these flanking faults is of the order of 1000 m or more. Secondary normal faults are visible mainly on the southern flank of the trough. Significant uplift occurs on the southern shoulder of the graben. The seismic units filling the Malta trough onlap the basin margins. These units may represent post-Messinian sequences, as identified by Ryan (1978), and are covered by Plio-Quaternary strata that are
270
Z. BEN-AVRAHAMETAL. by many workers, are detectable only in local, very detailed seismic surveys, such as in the western offshore of the Maltese islands (Gardiner et al. 1995). Analogue sandbox experiments in oblique rift models (where there is an angle between the rift axis and the extension direction) show remarkable similarities to the fault architecture and geometric disposition of graben structures found in the Strait of Sicily. In particular, the along-strike offsets in depocentres and the en echelon fault pattern parallel to the zone of rifting are the most striking evidence. En echelon stepping and segmentation of the axial depocentre is interpreted to occur across accommodation zones formed by complex interfingering extensional fault systems, as seen in natural examples such as the Central Graben of the North Sea (Roberts et al. 1990). Conceptual models based upon the analogue experiments show footwall uplift on the individual faults and mantle upwelling below the rift zone. Both these features are seen in the Strait of Sicily region. In SE mainland Sicily the Plio-Pleistocene fault pattern observed along the western margin of the Hyblean Plateau implies rightlateral movements along a broad NNE-oriented fault system that traverses the southern Sicilian foreland oblique to the front of the Maghrebian arc. Grasso et al. (1990) have argued that this foreland strike-slip system played the role of a transform fault linking zones of modern rifting within the Strait of Sicily with zones of recent underthrusting in south-central Sicily.
Small basins along the Cyprian a n d Hellenic arcs
Fig. 8. Above: near-trace monitor of part of seismic profileMS-120 (see location in Fig. 6), where it crosses the central-western part of the Malta trough. The basin is flankedby prominent subverticalfaults, and the basin fill is characterizedby mostly horizontal reflectors onlapping the basin margins. Significantuplift of both of the hanging walls of the depression can be seen on the profile. Below: line, drawing interpretation. affected by structural discontinuities and possibly by block rotation underneath the sedimentary cover. Some evidence of asymmetry within the basin fill is evident from the section in association with recent contractional deformation, as noted by Argnani (1990) on the basis of sparker profiles acquired within both the Malta and Linosa troughs. Rift model. Intense block rotation and tilting generated by differen-
tial motion corresponding to the subvertical faults are the most notable tectonic features characterizing the entire rift system. The polarity of the block rotation varies along the rift system and the most pronounced deformation is found within the axial areas of the troughs, where block subsidence and tilting is particularly impressive. In general, the tectonic style that characterizes the Strait of Sicily rift system is remarkably symmetrical, a mode of extension that has been described kinematically as pure shear (McKenzie 1978), with an upper brittle layer overlying a ductile lower layer, producing a symmetrical lithospheric crosssection. The model predicts the formation of sediment-filled grabens, which causes isostatic disequilibrium and the compensatory rise of the asthenosphere, eventually accompanied by surface volcanism in a late evolutionary stage of the rift system. The origin of possible asymmetries within the rift system is generally ascribed to inherited inhomogeneities in the lithosphere or/and in the uppermost crustal layers. The structural configuration of the Strait of Sicily resembles most, if not all of the tectonic elements predicted by the model. In particular, seismic data have shown that the normal fault pattern dominates, and appears to have controlled the evolution of the trough within the Pelagian block. This suggests a mechanism in which the stress field is largely extensional, and appears to have acted mostly as dip-slip faults. Dextral strike-slip mechanisms along N W - S E - or east-west-trending faults, as proposed
The series of basins along the northern margin of the Eastern Mediterranean and the Aegean Sea can be divided into two groups, those containing Messinian evaporites to the east and those without to the west. In the following, we describe briefly the various small basins in the same sense, from east to west. A number of small basins are found both on- and offshore (Fig. 1) in the northeastern corner of the Mediterranean. They started to form at roughly the same time, probably at least from the early Miocene, but have since been tectonically separated to some degree. On the basis of the presence at depth of shallow marine and continental deposits overlying Messinian evaporates, the Adana and Iskenderun basins may have subsided at least 3 km in the Pliocene-Quaternary; however, the Mut Basin, to the north of the Cilicia-Adana Basin, contains at least 1500 m of almost horizontal Neogene sediments at a present elevation of 1500-2000 m, a vertical difference between the basins of almost 5 km. This is not unusual in the Eastern Mediterranean. The Cilicia Basin. The Neogene Cilicia Basin (Fig. 4), a roughly 3 km deep N E - S W trough (shallower to the SW) lying between Cyprus and Turkey, is a seaward continuation of the Adana Basin, to the NE in Turkey (but offset to the south by the Ecemis Fault, bounding the western part of the Adana Basin on land). It contains Messinian evaporites, which show diapirism and lateral flowage away from the northern and northeastern depocentres (Evans et al. 1978; Aksu et al. 1992a,b). Asymmetry in the sedimentation has developed as a result of the major sources being to the north and northward tilting of the northern Cyprus margin. The two most important structural boundaries to the basin are the transpressional Kyrenia-Misis Ridge to the SE and the A n a m u r Komakiti Ridge, which separates the Cilicia Basin from the deeper Antalya Basin to the west. The Latakia and Larnaca basins. These are relatively shallow Miocene basins lying to the east of Cyprus on steps formed by the development of several sinistral wrench zones along the boundary of the African and Anatolian plates (Ben-Avraham et al. 1995; Vidal & Alvarez-Marrrn 2000; Vidal et al. 2000). They are formed on northward tilting basement rocks that are continuous from Cyprus to Syria and Turkey. Post-Miocene tectonics
EASTERN MEDITERRANEAN BASIN SYSTEMS
resulting from the extrusion of the Anatolian plate to the west (with Cyprus forming the southernmost part of this, pushing more southward) caused both the tilting and uplift of the basins. Uplift has resulted in the basins forming two steps from the Levantine Basin towards the Cicilia Basin to the north: the Latakia Basin forms the southern step and the Larnaca the second step just to the north (Fig. 4). The fault-bounded basins are separated by the Larnaca Ridge, and limited to the north and south by the Kyrenia-Misis Ridge and the Latakia Ridge, respectively (Ben-Avraham et al. 1995).
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lying west of the Cilicia Basin between Turkey and the Florence Rise (Figs 4 and 9). It was shown to contain thick Messinian evaporites and display northward thickening sediments suggestive of active and long-term northward tilting (Woodside 1977; Taviani & Rossi 1989; Sage & Letouzey 1990). A Palaeogene and Mesozoic basement is proposed from the relationship of the basin to Alpine nappe structures inferred at the Florence Rise (Sage & Letouzey 1990; Woodside et al. 2002). The ophiolites that have been shown to continue from SE Turkey (Baer-Bassit) into Cyprus (Delaloye & Wagner 1984; Delaune-Mayere 1984; Robinson & Malpas 1990) and to reappear again in the Antalya Nappes Complex (e.g. Robertson & Woodcock 1982) of southwestern Turkey have been thought also to continue through the Florence Rise or the Antalya Basin (e.g. Monod 1976), but there is little evidence for this in the form of typical magnetic or gravity anomalies (Woodside 1977; Woodside et al. 2002). A major roughly north-south structural discontinuity separates Cyprus and the Cilicia Basin to the north from the Antalya Basin, which has substantially greater depth than the western Cilicia Basin. A gravity discontinuity (low to the west, higher to the east) defines this boundary (Woodside 1976). The western boundary of the Antalya Basin is the fault-bounded Anaximander Mountains (Zitter et al. 2003; Ten Veen et al. 2004), and the fault-bounded western margin of the Gulf of Antalya to the north. Within the Gulf of Antalya to the north, there appears to be a buried, abrupt N W - S E boundary (defined also by a change in the gravity field; Woodside 1976) of the deepest part of the basin, with a shallower part to the north. The shallow part is likely to be the offshore extension of the Manavgat Basin, which is located onshore along the eastern margin of the Gulf of Antalya (e.g. Necker et al. 1998; Flecker & Ellam 1999). The Finike Basin. The Finike Basin is a narrow (about 20 km) depression about 80km long lying between southwestern Turkey to the north and the Anaximander Mountains to the south. It appears at first glance to be a continuation of the Antalya Basin to the west, but this is only a morphological continuity. An absence of Messinian evaporites and a relatively thin post-Miocene sedimentary fill (no more than about 1200m, assuming a seismic velocity of 1700 m s -1) indicate that this basin is geologically relatively young, probably no older than late Pliocene to Present. Northward-tilting sediments indicate that it is still forming, with the appearance of a rift basin being created by listric faulting along its northern boundary. This narrow and very linear (WSW-ENE) section of the margin of southwestern Turkey is inferred to mark the upper part of the listric fault. The fault may have originated as a strike-slip fault connected through the Rhodes Basin with transpressional faulting through the Pliny and Strabo trenches to the west (Ten Veen et al. 2004). Parallel faulting is mapped on land in Turkey (Gutnic et al. 1979). Basement rocks in the western part of the basin are shown to be similar to the Suzug Dag and Bey Daglari section in Turkey (Woodside et al. 1997), and in the east they form a continuation of the Antalya Nappes Complex, which is traced as far south as the Anaximander Mountains at 35~ about 60 km south of the Turkish coast (Woodside et al. 1997). Thus the Finike Basin acts as a rift basin separating the Anaximander Mountains from
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Turkey (Nesteroff et al. 1977; Woodside et al. 1997). To some degree, sediments from the Finike Basin can be followed southward onto the Anaximander Mountains that form the basement below an erosional unconformity. The Rhodes Basin. The Rhodes Basin shows many similarities with the Finike Basin, having no Messinian evaporites, and a relatively thin post-Miocene section (no more than about 1000 m, assuming a seismic velocity of 1700 m s -1) overlying inferred Alpine basement (Woodside et al. 2000). It is also, at about 4500 m depth (Fig. 10), one of the deepest basins of the Mediterranean Sea, which implies a relatively recent and rapid subsidence. It is bounded to the west and NE by faults (Woodside et al. 2000). The western boundary fault shows transpression with reverse faulting (Woodside et al. 2000; Kontogianni et al. 2002) and was thought to form the plate boundary between the Aegean and African plates, in continuity with sinistral transpressive shear along the Pliny Trench. Because of the poorly constrained nature of the Fethiye-Burdur fault zone to the NE, and a number of normal fault-plane solutions for earthquakes along it (rather than strike-slip; McKenzie 1978), the idea that the plate boundary continues in this direction has been disputed by Ten Veen et al. (2004), who have suggested instead that the plate boundary crosses the Rhodes Basin and continues along the northern edge of the Finike Basin. The Rhodes Basin is separated into two sub-basins by a zone of deformation that could be seen as the link between the Finike Basin and the Strabo Trench (Ten Veen et al. 2004). The Aegean basins. Within the Aegean Sea are a number of generally shallow basins that formed by extension of Alpine (Hellenide) basement during the southwestward extrusion of the Aegean and Anatolian plates (McKenzie 1972; Mascle & Martin 1990; Jackson 1994). The two most important of these basins are the Crete Basin (over 2000 m deep in the east), just to the north of Crete and parallel to its west-east long axis (Angelier et al. 1982), and the North Aegean Trough (over 1000 m deep in two connected E N E - W S W linear depressions, the Saros Trough to the east and the Sporades Trough to the west; Papanikolaou et al. 2002) which formed probably as a pull-apart along the principal strand of the North Anatolian Fault zone at its westernmost end (Yaltlrak et al. 1998; Yaltlrak & Alpar 2002). The North Anatolian Fault zone, the key tectonic element in the development not only of the Sea of Marmara but also of the North Aegean
Fig. 10. Six-channelseismic profile (using two GI 75 guns as source; with total air chamber volume of 2.451) from the French PRISMED II expedition (line 24, modifiedafter Woodside et al. 2000) running roughly north-south in the southern Rhodes Basin. It should be noted that there are no Messinianevaporites overlying the pre late-Miocene eroded basement, and that the post-Miocene sediments are relatively thin and undeformed.
Trough, abuts the Greek mainland at the western end of the Sporades Basin; however, the motion is absorbed by a system of extension with rapid clockwise motion, which, within the past 1 2 Ma, has passed motion on to the Gulf of Corinth and the Cephalonia Transform Fault to the west (Le Pichon et al. 2002). Thus an age of about 2 Ma can be given to the tectonic regime now prevailing in the Aegean basins (Le Pichon et al. 2002), although the beginnings of extension date back to as early as 25 Ma (Jackson 1994). Between the Cretan and the North Aegean Troughs lie a number of smaller basins with similar structural trends. Included in these are, from north to south, the Edremit Trough, the North IkariaSamos Basin and the North Mikonos-Andros Basins, and the South Ikarian Basin. These are mainly grouped in the eastern half of the Aegean (e.g. Lykousis et al. 1995). To the west are small Plio-Quaternary basins such as the Saronikos Basin (following roughly the Gulf of Corinth structural trend to the east of the Peloponnesus), the Mirthes Basin, and the Argolide Basin. The differing structural trends between west and east are related to the southward migration of the Hellenic Arc and probably rollback in the back-arc, and have been modelled by Jackson (1994), among others (e.g. Kreemer et al. 2003), as well as imaged by seismic tomography (Spakman et al. 1988).
Discussion The Eastern Mediterranean basin systems were formed in several stages. During the first stage of evolution, continental fragments rifted away from Africa to form the Ionian and the Levantine basins. This stage was followed by other rifting events. In the Levantine Basin, during the first stage, continental fragments, now part of Turkey, rifted away from the Levant and Sinai, and in the second stage, the Erathostenes Seamount and possibly other microcontinental blocks rifted away from Africa and moved northward. This process has caused the formation and destruction of oceanic crusts in the basins. The large-scale process that dominates the evolution of the Eastern Mediterranean basin systems is the approach of two large lithospheric plates, the African plate and the Eurasian plate, toward each other (Ben-Avraham 1989). The sea-floor spreading process in the Ionian and Levantine basins was interrupted occasionally by the collision of the rifted fragments with the southern margin of the Eurasian plate in the north. The rifting of continental fragments away from Africa, while the African plate was moving northward relative to the Eurasian plate, means that subduction along the Calabrian, Hellenic and Cyprian arcs had to be faster than the convergence of the two plates. Le Pichon et al. (1982) suggested that because of the large slab-pull force, the subduction of land-locked deep-sea basins will, in general, occur much faster than the collision rate. As a consequence, subduction of the old deep-sea basins in the Eastern Mediterranean will be compensated by the formation of young deep-sea basins behind the subduction zones. This is, in fact, what is taking place in the Eastern Mediterranean. Large basins, such as the Tyrrhenian and Aegean, as well as small basins, such as Cilicia and Antalya, are opening behind the Calabrian, Hellenic and Cyprian arcs. The mechanism responsible for the origin of the rift system in the Strait of Sicily is a crucial point that needs to be addressed. Very few workers have attempted to unravel this question. The difficulty of understanding the genetic evolution of the area is attributed to the inherent geological complexity of the region, which is surrounded by an assemblage of relatively small lithospheric blocks with a wide variety of rheologies and thicknesses that evolved during the pre-Tertiary tectonic evolution of the Africa passive palaeo-margin. The geological nature of these crustal segments is still poorly known because of the lack of distinctive geological constraints, and, in particular, the poor seismic coverage and stratigraphic information on both the
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pre-Mesozoic substratum and the crystalline basement. Ben-Avraham & Grasso (1990, 1991) stressed that one of the most important elements triggering segmentation along collision zones is probably the crustal structure variation along leading edges of the impinging Africa plate. The most important tectonic event that occurred in the central part of the Mediterranean Sea in the latest Tertiary was the rifting and opening of the Tyrrhenian basin, which started in late Miocene time and continued until the early Pleistocene with the possible formation of oceanic-type crust (Kastens et al. 1988). Because the Strait of Sicily rift system and the Tyrrhenian Sea formed in the same time span, Argnani (1990) suggested that the two geodynamic events could be in some way correlated. Rollback of the subducted slab and lithospheric mantle delamination have been proposed as feasible mechanisms that have produced a limited amount of extension within the Pelagian block, as a consequence of slab-pull forces and secondary mantle convection. On the other hand, Reuther & Eisbacher (1985) suggested that the origin of the dramatic change in the stress pattern during the Messinian might be related to an abnormal tectonic context; for example, a northeasterly subduction of the Ionian lithosphere beneath the Aegean Arc, as argued by Le Pichon et al. (1982). According to this hypothesis, crustal extension affecting the lithosphere underlying the Pelagian block occurs where it pulls away from its African anchor, giving rise to graben development and associated basaltic volcanism. Ben-Avraham et al. (1987) have considered the activity of a 1000 km long transcurrent fault running along the north African passive margin to explain the crustal extension within the Strait of Sicily. The seismic data presented here have shown that most of the faults in the Strait of Sicily affect the sea floor, indicating recent tectonic activity. No significant evidence of reactivation, inner deformation of the fault-rotated blocks and azimuth changes within the throws has been detected. Considering that the Strait of Sicily rift zone is a relatively young tectonic feature (early Pliocene to Present), we may assume that the extensional tectonic regime did not change significantly during this time span. Further support is provided by the fact that the volcanic edifices within the rift system are all Quaternary in age, indicating a progressive evolution of the rift from an immature stage to a more developed configuration in which partial melting is taking place.
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The Mesozoic-Cenozoic tectonic evolution of the Greater Caucasus A L I N E S A I N T O T t'2, M A R I E - F R A N ~ O I S E B R U N E T 3, F E D O R Y A K O V L E V 4, M I C H E L S E B R I E R 3, R A N D E L L S T E P H E N S O N l, A N D R E I E R S H O V 5, F R A N ~ O I S E C H A L O T - P R A T 6 & T O M M Y M C C A N N 7
1Netherlands Centre for Integrated Solid Earth Sciences, Faculty of Earth and Life Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands 2present address: NGU, Leiv Eirikssons vei 39, N-7491 Trondheim, Norway (e-mail:
[email protected]) 3UMR 7072 Tectonique CNRS-UPMC, Case 129, Universitd Pierre et Marie Curie, 4 place Jussieu, 75252 Paris cedex 05, France 4Institute of Physics of the Earth, RAS, Moscow, Russia 5Moscow State University, Geological Faculty, Moscow, Russia 6CRPG, 15 rue Notre Dame des Pauvres, BP 20, 54501 Vandoeuvre-Les-Nancy Cedex, France 7Geologisches Institut, Universitdt Bonn, NuBallee 8, 53115 Bonn, Germany
Abstract: The Greater Caucasus (GC) fold-and-thrust belt lies on the southern deformed edge of the Scythian Platform (SP) and results from the Cenozoic structural inversion of a deep marine Mesozoic basin in response to the northward displacement of the Transcaucasus (lying south of the GC) subsequent to the Arabia-Eurasia collision. A review of existing and newly acquired data has allowed a reconstruction of the GC history through the Mesozoic and Cenozoic eras. In Permo(?)-Triassic times, rifting developed along at least the northern part of the belt. Structural inversion of the basin occurred during the Late Triassic corresponding to the Eo-Cimmerian orogeny, documented SE of the GC and probably linked to the accretion of what are now Iranian terranes along the continental margin. Renewed development of extensional basin formation in the area of the present-day GC began in Sinemurian-Pliensbachian times with rift activity encompassing the Mid-Jurassic. Rifting led to extreme thinning of the underlying continental crust by the Aalenian and concomitant extrusion of mid-ocean ridge basalt lavas. A Bathonian unconformity is observed on both sides of the basin and may either correspond to the end of active rifting and the onset of post-rift basin development or be the record of collision further south along the former Mesozoic active margin. The post-rift phase began with deposition of Late Jurassic platform-type sediments onto the margins and a flysch-like unit in its deeper part, which has transgressed the basin during the Cretaceous and Early Cenozoic. An initial phase of shortening occurred in the Late Eocene under a NE-SW compressional stress regime. A second shortening event that began in the Mid-Miocene (Sarmatian), accompanied by significant uplift of the belt, continues at present. It is related to the final collision of Arabia with Eurasia and led to the development of the present-day south-vergent GC fold-and-thrust belt. Some north-vergent retrothrusts are present in the western GC and a few more in the eastern GC, where a fan-shaped belt can be observed. The mechanisms responsible for the large-scale structure of the belt remain a matter of debate because the deep crustal structure of the GC is not well known. Some (mainly Russian) geoscientists have argued that the GC is an inverted basin squeezed between deep (near)-vertical faults representing the boundaries between the GC and the SP to the north and the GC and the Transcaucasus to the south. Another model, supported in part by the distribution of earthquake hypocentres, proposes the existence of south-vergent thrusts flattening at depth, along which the Transcaucasus plunges beneath the GC and the SP. In this model, a thick-skinned mode of deformation prevailed in the central part of the GC whereas the western and eastern parts display the attributes of thin-skinned fold-and-thrust belts, although, in general, the two styles of deformation coexist along the belt. The present-day high elevation observed only in the central part of the belt would have resulted from the delamination of a lithospheric root.
The Greater Caucasus (GC) belt forms a morphological barrier along the southern margin of the Scythian Platform (SP; contiguous with the southern East European Platform, EEP), running from the northern margin of the eastern Black Sea Basin to the South Caspian Basin (Fig. 1). It developed during several phases of deformation in M e s o z o i c - C e n o z o i c times (Milanovsky & Khain 1963; Adamia et al. 1977, 1981; Rastsvetaev 1977; Khain 1984; Muratov et al. 1984; Gamkrelidze 1986; Dotduyev 1989; Zonenshain et al. 1990; Nikishin et al. 1998a,b, 2001). The geology of the GC has been studied for at least 150 years and a significant volume of published literature deals with its evolution, although much of this is difficult to access for the international scientific community. The GC is located in the Black S e a - C a s p i a n Sea region, which is regarded as a mosaic of terranes of Gondwanan, Tethyan and Eurasian affinity that are sometimes controversial in origin (see discussions by ~eng6r 1984; Zonenshain et al. 1990; Dercourt et al. 1993, 2000). Accretion of these blocks along the SP occurred throughout the Phanerozoic and, accordingly, orogenic events developed in the GC as such: the Palaeozoic Variscan orogeny, the Triassic-Jurassic Cimmerian orogeny, and the Cenozoic Alpine orogeny. Structural styles of the GC belt are not yet unequivocally fixed and different proposed geometries exist in
the literature, even for the major boundary faults separating fundamental tectonic units. It follows that there is still considerable disagreement regarding tectonic mechanisms, simply because there are insufficient diagnostic data. The GC orogenic events are also not well understood in terms of the driving mechanisms. There are major discrepancies concerning the rate of shortening and the nature of the M e s o z o i c - C e n o z o i c basement of the GC. Did oceanic crust and lithosphere form during this time or not? In other words, did a complete orogenic Wilson cycle from opening of an ocean to its consumption by subduction and collision of its margins take place along the GC during the M e s o z o i c - C e n o z o i c ? There is rough agreement regarding the continuing Late Cenozoic pulses of mountain building and uplift, which have resulted from collision-accretion of the Transcaucasus continental block along the southern margin of the SP (Fig. 2). At a regional plate tectonic scale, this corresponds to the final stage of the Alpine orogenic cycle involving the collision between Eurasia and Afro-Arabia 'mega'-continental plates, with the main suture zone of the Tethyan Ocean running through Anatolia and the Lesser Caucasus (Fig. 1). The aim of this paper is first to assess and present the existing data, and then to describe and reinterpret them, as necessary, as well as to present some new data to constrain better the M e s o z o i c - C e n o z o i c orogenesis of the
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 277-289. 0435-4052/06/$15.00
9 The Geological Society of London 2006.
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Fig. 1. Digital topographic map of the Greater Caucasus, from the global topography 2 minute database; illumination from N135. GC: from the development of basins, oceanic or not, to their inversion and/or the collision of continental blocks subsequent to the consumption of an oceanic plate by subduction.
Structure of the Greater Caucasus The GC belt comprises a basement-core containing strata as old as Proterozoic (Figs 2 and 3), with Jurassic to Eocene formations lying on its flanks. The deep crustal structure of the GC and Transcaucasus are not known, as the resolution of deep seismic sounding (DSS) lines acquired more than 20 years ago and crossing the belt is insufficient to provide an accurate crustal image. Consequently, discrepancies exist regarding the published deep structure of the belt. The two main competing models of the deep structure of the GC are described below. (1) The first model argues for a subvertical geometry of all the GC main faults, including the border faults. Somin (2000), for example, argued for a subvertical disposition of the Main Caucasian Thrust (MCT; shown in Fig. 3) at great depths, given its steep near-surface dip (65-80 ~ along all of its strike and at imaged depths to 3 - 5 kin. Nevertheless, such a steep geometry could also have resulted from deformation of the fault during the final stage of collision. Both reprocessed old and newly acquired geophysical data (Shempelev et al. 2001, 2005; Grekov et al. 2004; Prutsky et al. 2004) were used to show a similarly deep inclination of the MCT at a depth of 80km. Shempelev et al. (2001) and Rastsvetaev et al. (2004) proposed the same subvertical geometry of regional faults along profiles crossing different parts of the GC. It was concluded that the boundary of the western GC with the Black Sea is a steep (60-80 ~ and deep (80 kin) major fault (Shempelev et al. 2001) linking with the Racha-Lechkhumy Fault Zone (RLFZ; see location in Figs 3 and 4) to the east (Yakovlev 2002, 2005). (2) The alternative model, proposed by Gamkrelidze (1986), Dotduyev (1987), Giorgobiani & Zakaraya (1989), Baranov et al. (1990), Zonenshain et al. (1990) and Gustchin et al. (1996), and referenced by many others, differs strongly and favours instead flat-dipping thrusts at depth. This is a thick-skinned tectonic model for the GC (Fig. 3) with the orogen interpreted as a collage
of two or three northward underthmst slabs. These authors have argued that the SP is thrust upon the Transcaucasus continental block. The MCT sensu stricto is thus considered to be a northdipping flat thrust at depth (Fig. 3) along which the pre-Jurassic basement of the Main Range zone and its overlying Mesozoic cover were presumably displaced southwards some 100 km or more during the Cenozoic. (The MCT sensu lato comprises at least two parallel branches at the surface (see Figs 2 and 4) and the northern branch is the MCT sensu stricto, along which the crystalline basement thrusts onto the sedimentary succession (see Fig. 3). Similarly, units of the GC belt have been thrust southwards over the Transcaucasus along the RLFZ (Fig. 3). Accordingly, the GC is regarded as a large south-vergent fold-and-thrust belt with its northern limb forming a gently north-dipping monocline toward the SP. Some north-vergent thrusting could be locally present along the western and central part of the GC (Milanovsky & Khain 1963). Back-thrusting is more developed onto the Terek-Caspian foreland, where the northward propagation of the Dagestan nappes contributes to the fan-shaped structure of the eastern part of the belt (Fig. 4; see Ershov et al. 2003). Thick-skinned deformation is reported along a north-south profile cutting across the internal part of the GC, with thin-skinned deformation prevailing on the southern front in the Rioni and Kura basins. A north-south profile across the western GC (Fig. 5; Robinson et al. 1996) also shows thick-skinned deformation with imbricate structures involving the basement. It can be seen that the N W - S E and W N W ESE faults parallel to the general grain of the belt are thrust faults flattening at depth whereas the NNW-SSE faults transverse to the belt are steeper (Koronovsky 1984; Giorgobiani & Zakaraya 1989; Philip et al. 1989; Giorgobiani 2004; Fig. 4). The Fore-Caucasus region, which lies on the SP, evolved in conjunction with the GC. From Latest Eocene-Oligocene times, two flexural basins developed, separated by the elevated zone called the Stavropol High (Figs 2 and 3). This comprises a north-south elongated and anomalously thick crustal block (see Kostyuchenko et al. 2004) that from early Mesozoic times never significantly subsided. The Terek-Caspian foreland basin to its east and the Kuban foreland basin to its west developed during the Cenozoic, both showing a high subsidence rate during the Oligo-Miocene ('Maykop' facies). Thus, what is peculiar about the Fore-Caucasus area is that 'foreland type-like basins' developed in front of the more topographically subdued eastern and western parts of the belt but not in front of its topographically highest central part (see Ershov et al. 2003). Several basins also developed south of the GC belt, in the Shatsky Ridge-Transcaucasus area, in Oligo-Miocene times. These are, from west to east, the Tuapse, Rioni and Kura basins (Figs 2 and 5). They are reported to be flexural in type and related to Eocene compression (Milanovsky & Khain 1963; Gamkrelidze 1986; Nikishin et al. 1998b). However, the history of the Kura Basin is more complex, as it is the western prolongation of the South Caspian Basin (Brunet et al. 2003). A review of the Early Mesozoic tectonic evolution of the Greater Caucasus The Triassic and Jurassic history of the area is not well constrained and is still a matter of considerable debate. The extent and age of Cimmerian orogenic phases, in Late Triassic or Early Jurassic, MidJurassic and Late Jurassic times, as well as the successive rifting events, are not confidently known (see Nikishin et al. 1998a,b). E a r l y Triassic basin d e v e l o p m e n t a n d Late Triassic E o - C i m m e r i a n tectonics
Permo(?)-Triassic rifting and volcanism (and, probably, magmatism-related doming) are widespread in the Fore-Cancasus
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Jurassic-Cretaceousophioliticcomplex
Volcanicplateaus (dacites,andesiticbasalts)
region and in the northern part of the GC (Nazarevich et al. 1986; Lordkipanidze et al. 1989; Tikhomirov et al. 2004). The geodynamic setting of such tectonics is still debated: was this a back-arc setting or not? Throughout the area, including the northern GC, there was also a period of Late Triassic compression (Eo-Cimmerian
Fig. 2. Simplified geological map of the Greater Caucasus and Lesser Caucasus (from Milanovsky & Khain 1963) and locations of cross-sections shown in Figures 3 and 5. PTF, Pshekish-Tyrnauz Fault; MCT, Main Caucasian Thrust; RLFZ, Racha-Lechkhumy Fault Zone. Encircled numbers: 1-6 are zones of the Scythian Platform: 1, Stavropol High; 2, Azov-Berezan High; 3, Manych Basin; 4, Kuban Basin; 5, Terek-Caspian Basin; 6, Kussar-Divitchi Basin; 7-13 are zones of the Greater Caucasus: 7, Peredovoy Zone; 8, Betcha Anticline; 9, Svanetia Anticline; 10, Laba-Malka Monocline; 11, Dagestan Folded Zone; 12, Flysch Zone of southeastern GC; 13, Flysch Zone of north-western GC; 14-24 are zones of the Lesser Caucasus: 14, Somketo-Karabakh Zone; 15, Artvin-Bolnisi Zone; 16, Adzharo-Trialet; 17, Talesh; 18, Sevan-Akera; 19, Kafan; 20, Vedin; 21, Zangezur; 22, Mishkhan-Zangezur Massif; 23, Ararat-Djulfa Massif; 24, Araks Basin; 25-29 are intramontane zones of the Transcaucasus and Black Sea: 25, Rioni Basin; 26, Kura Basin; 27, Dzirula Massif; 28, Tuapse Basin; 29, Shatsky Ridge.
tectonic phase) during which all the Permo(?)-Triassic basins were inverted (Nikishin et al. 1998a,b, 2001; Gaetani et al. 2006). The compressive event is probably related to the coll i s i o n - a c c r e t i o n of Gondwana-derived blocks (which together form the composite Iran plate; S ~ d i 1995; Besse et al. 1998) SE of the GC w h e n the Palaeotethys Ocean closed along the
Fig. 3. Section across the central part of the Greater Caucasus showing the southward vergence of the whole belt and the major thrusting of the belt over the Transcaucasus (Dotduyev 1987). (Section location is shown in Fig. 2.)
280
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Fig. 4. Tectonic map of the Greater Caucasus area from Ruppel & McNutt (1990) (other sources: Milanovsky& Khain 1963; Kotansky 1978; Dotduyev 1987; Philip et al. 1989). Talesh-Alborz-Aghdarband zones, in the area of the present-day South Caspian Basin (Sttcklin 1968; Davies et al. 1972; ~engtr 1984; Alavi 1991; Ruttner 1993; Dercourt et al. 2000, and references therein). This Eo-Cimmerian orogeny has also been clearly identified as a major event in the Turkish Pontides (Okay 2000; Okay et al. 2006). It should also be noted that (1) the Black Sea was certainly not developed by the Late Triassic, and the Pontides were therefore close to the Transcaucasus-GC; (2) the Pontides-Transcaucasus- Talesh-Alborz- Aghdarband-GC zones probably together formed a contiguous part of the widespread Eo-Cimmerian orogenic belt.
E a r l y J u r a s s i c to M i d - J u r a s s i c
A field study carried out in 2003 led to the postulation of a model of the GC in Jurassic times by Saintot et al. (2004). A new rifting phase occurred during the Early Jurassic (Zonenshain et al. 1990; Nikishin et al. 1998a,b, 2001, and references therein) under a transtensional stress regime with a nearly east-west-directed tensional stress axis (Stbrier et al. 1997; Saintot et al. 2004). Thus, this transtensional Early Jurassic rifting shares some similarities
with the model of Banks & Robinson (1997) for the Black Sea region, which surmises that the Early Jurassic GC Basin corresponded to an en echelon set of rhomb-shaped depocentres. Early Jurassic rift activity is also reported in the Eastern Pontides, which were adjacent to the GC at that time (e.g. Okay & Sahintfirk 1997) and in the South Caspian Basin (Early(?) to Mid-Jurassic times; Brunet et al. 2003). GC rifting continued through part of the Mid-Jurassic. Extrusive magmatism (mainly rhyolitic) accompanied the GC rifting phase during Sinemurian-Pliensbachian times (Lordkipanidze et aL 1989). Sediments of this age are represented by deltaic(?) coarse sandstones NW of the belt, by deep marine mudstones-sandstones in the central part, and by shallower mudstones-sandstones to the south. In Toarcian times, from north to south, shelf to deep marine mudstones-sandstones were deposited, with no record of volcanic activity. In AalenianBajocian units, sedimentary facies laterally vary from continental to deep marine. The Aalenian and Bajocian periods are also characterized by bimodal rhyolitic and basaltic extrusive rocks (from mantle and crustal sources) in a subaerial as well as a shallow-marine environment. In the model, the western part of the GC evolved during the Early and Mid-Jurassic as the western margin of the rift with shallow-water sedimentation and subaerial extrusion of lava flows. Deep-water sediments are encountered towards the present-day central part of the belt (crossing the inferred, north-south-oriented normal faults), associated with mid-ocean ridge basalt (MORB)-like tholeiitic basalt extrusion during the Aalenian. Not only partial melting of asthenosphere is implied, but also a high degree of extension, approaching that required for oceanic crust development in the present-day central part of the GC belt. The total thickness of the Lower Jurassic to Aalenian unit in some parts of the GC Basin is more than 5000 m, and it is mainly composed of black shales and deep-water sandstone turbidites (as well as the volcanic rocks and pyroclastic deposits). The Aalenian extensional phase has been well documented in the field with, for example, the presence of a large NW-SE-trending normal fault, east of the Kuban Basin. Toarcian to Aalenian units are tilted along this fault (Fig. 6) and the minimum downthrow should be of several hundreds of metres. The age of fault activity is constrained by overlying, sealing Upper Aalenian units. In Bajocian times, a huge quantity of pyroxene-bearing basalts were extruded, and formed a subaerial to shallow-water volcanic chain on the southern margin of the basin (presumably accompanied by uplift at the rift margin). Synchronously with the formation of this relief, conglomerates (reworking the lavas) were deposited toward the depocentre of the basin to the north.
Fig. 5. Section across the western part of the Greater Caucasus showing the basement involved in south-vergent, flat thrusting. Offshoreis shown an interpretation of the seismic line SU8040 (from Robinson et al. 1996). (Section location is shown in Fig. 2.)
THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS
Fig. 6. Photograph of a progressive unconformitycreated by a major synsedimentarynormal fault within Aalenian deposits dipping to the NE. The normal fault, with a NW-SE strike, is located to the left of the photograph and may be followedfor some 10 km along the upper rims of the Upper Kubanvalley, on the northern side of the central GC. This normal fault testifies to the Mid-Jurassic extensionaltectonics of the GC Basin and contributes to the along-strike variation of sedimentarythickness (photograph at 43~ 42~ by M. Srbrier).
The calc-alkaline nature of the Bajocian lavas has been formerly interpreted as indicative of a subduction-related volcanic arc marking the incipient subduction of a large oceanic plate such as the Palaeotethys along the GC (see, for example, the southward subduction of Palaeotethys along the GC as described by ~eng6r (1984)). However, this hypothesis seems very unlikely because the above-mentioned Bajocian volcanic rocks are also spread over the Transcaucasus and there is no evidence for a subduction zone along the southern edge of the GC (i.e. remnants of an accretionary prism, high-pressure metamorphism, ophiolitic fragments, etc.). The calc-alkaline nature of the Bajocian lavas can also be explained by the GC rift being in a back-arc setting relative to a subduction zone located far to the south in the Lesser Caucasus (see Adamia et al. 1981; Gamkrelidze 1986; Panov 2004). The Artvin-Bolnisi zone lying between the Transcaucasus and the suture zone of the Lesser Caucasus (location shown in Fig. 2) is a good candidate for a subduction-related volcanic arc during the Early and Mid-Jurassic, with shallow-water to continental sediments and major calc-alkaline volcanism (Adamia et al. 1981; Gamkrelidze 1986; Panov 2004). It is not uncommon for lavas extruded in a back-arc rift setting, but close to the volcanic arc, to show such calc-alkalinity. Therefore, it cannot be excluded that the Bajocian lavas were extruded during what could still be considered as a synrift stage of basin evolution, continuing from the Aalenian. However, it is noted that, whereas structural constraints (e.g. synsedimentary normal faulting) clearly exist to define the Aalenian succession as synrift, there are no such structural constraints for the Bajocian units. Indeed, the widespread occurrence of Bajocian calc-alkaline volcanic rocks that can be encountered from the Lesser Caucasus to the MCT may also simply suggest an expansion of the subduction-related volcanic arc from some 5 0 - 1 0 0 km width in Aalenian time, restricted by the Artvin-Bolnisi zone, to nearly 200 km in the Bajocian, thus merging with the southern part of the GC Basin. Shallowing of the subducted slab could explain such an encroachment of the arc into the previously back-arc setting. If this were the case, rift activity in the GC Basin would have stopped (given that a rather flat-dipping slab does not favour the opening of a back-arc basin; eg. Lallemand et al. 2005, and references therein) and, therefore, the Bajocian volcanic rocks should be considered as occurring at the onset of the post-rift stage of GC Basin development. The available observations, relating to only
281
the calc-alkaline character of the Bajocian lavas and their widespread occurrence, cannot discriminate between these two possibilities. The Bathonian unit (where not absent) is composed of a greywacke siltstone unit into the basin and regressive coal-bearing terrigeneous sediments on its southern margin. The Upper Jurassic unit lies transgressively and discordantly on the Middle Jurassic unit. It reportedly lies conformably on the Middle Jurassic unit along the present-day southern slope of the central and eastern part of the GC Basin (Gamkrelidze 1986; Zonenshain et al. 1990; Nikishin et al. 1998a,b), although field observations made in the same central area (by M. Srbrier in 2004) revealed an unconformity between the Callovian deposits and underlying units. It is also worth noting that Cenozoic deformation is so intense in the so-called Flysch Zone (see Fig. 2) that no clear conclusion can be made regarding the detailed relationships between Mesozoic units. A compressional event has been proposed to have occurred in Bathonian times, resulting in the inversion of the margins of the basin (Adamia et al. 1981), although it may be that this unconformity is simply related to the cessation of rifting and the onset of post-rift basin development, such as recorded in many rift basins (see, e.g. Coward et al. 1987; Tankard & Balkwill 1989; Frostick & Steel 1993; Williams & Dobb 1993; Busby & Ingersoll 1995; Stephenson et al. 1996; Cloetingh et al. 1997; McCann & Saintot 2003). (Brunet et al. (2003) also pointed out that the regional Bathonian unconformity around the South Caspian Basin may be a 'break-up unconformity' marking the onset of sea-floor spreading rather than the occurrence of a compressive tectonic event.) Nevertheless, in the southernmost part of the GC (in Georgia), Callovian strata overlie open folds in Middle Jurassic strata, constraining a gentle folding event to the Bathonian. In the central part of the GC, north of the MCT, highly folded Early Jurassic strata are overlain by subhorizontal layers of Upper Jurassic and Cretaceous platform-type deposits. According to Belov et al. (1990) and Somin (2000), they are in place and indicate that the Bathonian folding was significant and involved intense shortening. Other authors (e.g. Korsakov et al. 2001) have considered that the Upper Jurassic and Cretaceous strata are in an allochthonous position and, accordingly, that the thrust sheet and the folding developed together during Alpine orogenesis (implying the occurrence of a folding phase, followed by the development of an erosional surface and then thrusting of nappes along a drcollement level). On the northernmost slope of the belt, the angular discordance between transgressive Callovian and older rocks disappears. Published cross-sections (e.g. Panov 2002, 2004) show south-vergent folds and thrusts affecting strata older than and including Bajocian, and no sealing by younger sediments (which are absent). In the northern part of each of these crosssections lies a gentle monocline composed of Upper Jurassic and Cretaceous units underlain by Lower Jurassic units without any angular unconformity as might be expected to be related to a compressional phase during Bathonian times. (A Bathonian stratigraphic gap indeed exists locally, the Bathonian being a time of worldwide regression.) No important or diagnostic compressive structures (such as folds and thrusts) were observed in Middle Jurassic rocks sealed by the Callovian by the senior author during fieldwork in 2003 in the northern part of the belt (see Saintot et al. 2004). What was observed is a gently tilted unit (like the Aalenian unit) below the Callovian transgressive unit. Going southward across the belt, closely and tightly folded Lower and Middle Jurassic strata can be observed (Fig. 7). The same style of folding is observed some 10 km towards the Black Sea coast in Palaeocene rocks (Fig. 8). SE-vergent minor thrusts are also common in Lower and Middle Jurassic units, similar to the younger strata. The systematic analysis of brittle structures within the GC also strongly suggests that only one set of reverse faults developed in Jurassic and younger strata and that this set is related to the Cenozoic palaeo-stress field (Fig. 9; see discussion
282
A. SAINTOT ETAL. indeed, only isostatic readjustments at the syn- and post-rift transition), affecting units from place to place, rather than the complete inversion of the GC Basin. (In Lower Middle Jurassic rocks there is no evidence of intense folding and thrusting that can be ascribed unequivocally to a Bathonian compressional event, most of the deformation being clearly Cenozoic in age). In summary, the pre-Callovian unconformity remains a matter of debate. It could record either the transition between syn- and post-rift phases in the GC Basin or, alternatively, a weak compressive event related to the accretion of crustal blocks along the active continental margin to the south.
The Callovian-Eocene Greater Caucasus Basin
Fig. 7. Photograph of folded Aalenian-Bajocian unit (PshishFormation) of the western Greater Caucasus (photograph by A. Saintot; S. Korsakov for scale). Fold axes strike NW-SE to WNW-ESE. and analyses of structures related to Cenozoic shortening by Saintot & Angelier (2002)). The localized angular unconformity at the base of Upper Jurassic strata thus probably records not more than a phase of gentle compression of the GC Basin (or,
The GC Basin evolved dominantly as a post-rift (thermally subsiding) basin from the Callovian until the Late Eocene following its Early to Mid-Jurassic episodes of rifting. A thickness of 6 - 8 km of calcareous, mainly Cretaceous, flysch-type sediments was then deposited and most of the Greater Caucasus Mountains corresponds to the so-called Flysch Zone of the southern limb of the GC (Fig. 2; Milanovsky & Khain 1963; Lordkipanidze 1980; Koronovsky 1984; Gamkrelidze 1986; Belousov et al. 1988; Adamia & Lordkipanidze 1989; Zonenshain et al. 1990). The nature of the underlying crust has not been established, although Ershov et al. (2003) estimated a crustal thickness of 15-17 km, suggesting that it corresponds to thinned continental crust. Such an interpretation is in agreement with the absence of oceanic crustal remnants in the belt. It follows that the basin was probably not floored by significant oceanic crust (see also the important discussion by Ershov et al. 2003, p. 102). The Callovian conglomerates and calcareous sandstones clearly belong to the post-rift succession of the GC Basin. They unconformably overlie the oldest units on an erosional surface. Upward, the Callovian unit becomes marly, indicating platform subsidence. In Late Jurassic times, sandstones and clays filled in the sedimentary basin and reef limestones developed towards its margins. Kimmeridgian-Tithonian gypsumbearing and lagoonal sediments were deposited on the northern (Laba-Malka zone) and southern margins (in Georgia). A very thick Cretaceous to Eocene greywacke siltstone flysch-like unit with clastic limestones in the Lower Cretaceous interval conformably overlies the Upper Jurassic sequence. (The Lower Cretaceous succession is 750-1600 m thick, the Upper
Fig. 8. Two photographs of the Lower Palaeocene flysch-likefolded unit of the western Greater Caucasus along the Black Sea coast. Fold axes strike NW-SE to WNW-ESE (photographs by A. Saintot).
THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS
283
Shcherba 1998). In general, according to the Russian literature, there are two distinct orogenic processes recorded in the GC. The first of these is understood to have occurred in Late EoceneEarly Oligocene times, and involved folding with significant crustal shortening but without major uplift of rocks at the surface. The second is considered to have occurred in Miocene times, and is characterized by limited shortening (reported to be 5-10%) and no significant folding but rather by significant uplift, thus explaining the deposition of Sarmatian conglomerates.
Fig. 9. (a) Stereoplot of 60 close and tight fold axes collected in rocks of Early and Mid-Jurassic age of the western Greater Caucasus (Schmidt's projection, lower hemisphere). (b) Attitudes (strikes and dips) of 215 bedding planes collected in rocks of the western Greater Caucasus from Late Jurassic to Eocene in age. The attitude of folds is the same in both stratigraphic intervals: the WNW-ESE close folds in Late Jurassic to Cenozoic rocks are also observed in Early and Mid-Jurassicrocks. Most of the compressivestructures and tectonic contacts (such as thrusts) measuredin Early and Mid-Jurassicrocks are consistent with a NNE-SSW Cenozoic compression(see Saintot & Angelier 2002).
Cretaceous 5 0 0 - 9 0 0 m thick, and the Palaeocene-Eocene, 600-850 m thick). Restorations by Yakovlev (2002, 2005) along several profiles crossing the GC clearly show that the RLFZ (Figs 2 and 3) was the southernmost normal fault bordering the GC Basin, involved in controlling the northward increase of sedimentary thickness from the Mesozoic-Cenozoic Transcaucasus stable platform to its subsiding part, the Gagra-Dzhava zone (see Fig. 3). They also show that the cumulative primary normal displacement along the RLFZ was much larger than the secondary reverse one. The most important normal fault at this time, however, was the Utsera fault, limiting the Gagra-Dzhava zone and the flysch zone, in which the thickness of Mesozoic and Cenozoic deposits reaches 12-15 km. It was also during this time that the eastern Black Sea Basin developed, close to the GC Basin and south of the Shatsky Ridge (a western prolongation of the Transcaucasus; see Fig. 5), although the precise timing is still matter of debate: Late Cretaceous and Palaeocene according to Finetti et al. (1988), Eocene according to Lordkipanidze (1980), and Late PalaeoceneEocene according to Robinson et al. (1996) and Shreider et al. (1997). Analysis of kinematic data by Saintot & Angelier (2002) revealed that a transtensional stress field affected the GC Basin during the Eocene (with an east-west trend of extension), which those workers considered to be a far-field effect of rifting (or, at least, of rift reactivation) in the eastern Black Sea Basin. In any case, during Palaeocene-Eocene times, prior to the main shortening event, the GC Basin was a deep-water basin, with a sedimentary infill of 10 km on average (Borsuk & Sholpo 1983), similar to the eastern Black Sea and South Caspian basins (Zonenshain & Le Pichon 1986; Nikishin et al. 1998b) and linked with the latter but apparently separated from the former by the Shatsky Ridge (see Fig. 2).
Cenozoic to present-day shortening along the Greater Caucasus The main orogenic phase is considered to extend from Late Eocene to Early Oligocene time by Shardanov & Peklo (1959), Beliaevsky et al. (1961), Milanovsky & Khain (1963), Grigor'yants et al. (1967), Khain (1975, 1994), Milanovsky et al. (1984), Muratov et al. (1984), Giorgobiani & Zakaraya (1989), Robinson et al. (1996), Lozar & Polino (1997), Robinson (1997), Nikishin et al. (1998a,b, 2001) and Mikhailov et al. (1999), with pulses of orogeny encompassing the rest of Cenozoic to the present day. According to other workers, the orogeny did not begin prior to the Miocene, in Sarmatian time (e.g. Dotduyev 1987; Shcherba 1987, 1989, 1993; Zonenshain et al. 1990; Kopp 1991; Kopp &
Late Eocene
What follows is a summary of the main arguments used by authors to demonstrate that the inversion of the basin began in Late Eocene times. (1) An angular unconformity of Maykop (Oligo-Miocene) on deformed older units is regionally observed in the field (Milanovsky & Khain 1963; Khain 1975, 1994; Borukaev et al. 1981; Rastsvetaev & Marinin 2001; Banks, pers. comm.) and on seismic lines (Tugolesov et al. 1985; Robinson et al. 1996; Banks, pers. comm.) on both sides of the GC. On the southern slope of the central GC, there is (a) a Late Eocene olistostrome unit (10-400 m thick) coeval with southward thrusting of the GC Basin (Khain 1975, 1994) and (b) a deltaic, southward prograding sandy facies in the Oligocene unit coeval with the uplift-emergence of part of the GC (M. S6brier's field observation). Sharafutdinov (2003) dated the folding event as latest Eocene-Early Oligocene on the northern slope of the GC and in the Fore-Caucasus, and he reported Early Oligocene folds, olistostromes (confirming the existence of back-thrusts), angular unconformities and the tectonic removal of a large part of the section with everything being overlain by flat-lying strata of MidOligocene age. The youngest unit prior to the onset of deformation is Late Eocene in age (Khadum Formation). The reported features, including the folds (especially if they are related to slumping), imply synsedimentary deformation in a foreland developing at the front of a propagating back-thrust. (2) A high rate of tectonic subsidence occurred at the beginning of the Oligocene in the Indolo-Kuban and Terek-Caspian basins as shown by burial history modelling (back-stripping analyses of wells and numerical modelling of lithospheric deformation) by Nikishin et al. (1998a), Ershov et al. (1999) and Mikhailov et al. (1999). Those workers assumed that the Indolo-Kuban, TerekCaspian and Tuapse, Kura and Rioni troughs developed as flexural foreland basins in response to lithospheric compression from the south during Late Eocene times (resulting from the closure of Neotethys and collision south of the Transcaucasus area). Ershov et al. (2003) discussed mechanisms other than foreland flexure for the formation of these basins, including the role of mantle processes occurring at the cessation of shortening, related to the underthrusting of thinned continental crust beneath the basins. (3) Lozar & Polino (1997) carried out a study based on nannofossils occurring in Maykop sediments of the Kuban Basin and of Upper Cretaceous rocks on the northern slope of the western GC. The base of the Maykop group is inferred to be Late Eocene-Early Oligocene in age and its lowermost part contains a reworked assemblage (80% of the total assemblage) of Late Cretaceous and Palaeogene nannofossils. These are very well preserved (with, for example, intact spines), implying that they were not transported over long distances. The sediment source was the area of the present GC where, indeed, Late Cretaceous and Palaeogene sediments were eroded. However, although there is general agreement on the Late Eocene uplift of the central part of the GC, this area was not yet actually above sea level by this time according to Kopp & Shcherba (1985) and Ershov et al. (1999, 2003). The types of Maykop nannofossils found in situ suggest restricted environmental conditions, leading to the interpretation that environmental changes occurred during
284
A. SAINTOT ETAL.
Late Eocene-Oligocene times, either with the onset of a cooler climate or as a result of the isolation of the Paratethys domain by the uplift and emergence of an orogenic belt acting as a barrier along the Pontides-Lesser Caucasus. The Late Eocene compressional palaeostress field responsible for the inversion of the GC Basin has been determined through tectonic analysis by Saintot & Angelier (2002). It was oriented N E - S W to N N E - S S W , leading to the development of N W - S E and W N W - E S E dip-slip thrusts in the GC. The main features and chronology of this tectonic phase have been established as follows: (1) on the northern slope of the GC, where the regional structure is a monocline, the Palaeocene strata are clearly affected by the inferred compressional palaeostress field (see details of measurements and site numbers given by Saintot & Angelier (2002)), whereas no related reverse and strike-slip faults can be observed affecting the overlying Miocene rocks studied by Saintot & Angelier; (2) the palaeostress field has also been recorded in Middle Eocene rocks along the southwestern coast (see details of measurements and site numbers given by Saintot & Angelier (2002)); (3) this palaeostress field is the only one recorded during pre- (e.g. Fig. 10), syn- and post-folding phases (Saintot & Angelier 2002).
M i o c e n e to p r e s e n t day
From Sarmatian times (Mid-Miocene) until the present, pulses of compressional deformation have affected the GC (Belousov 1940; Shardanov & Peklo 1959; Beliaevsky et al. 1961; Milanovsky & Khain 1963; Shcherba 1987, 1989, 1993; Giorgobiani & Zakaraya 1989; Kopp 1989, 1991, 1996; Rastsvetaev 1989; Zonenshain et al. 1990; Milanovsky 1991; Khain 1994; Kopp & Shcherba 1998; Nikishin et al. 1998b). However, it appears as though the present-day structure of the GC is inherited mainly from the Sarmatian compressional pulse. The Sarmatian sedimentary unit surrounding the belt comprises syndeformational conglomerates reflecting the growth of topography at this time (Mikhailov et al. 1999) and, indeed, the emergence of the GC belt as a whole, the central GC having already been uplifted since the latest Eocene (Khain 1994; Lozar & Polino 1997; Ershov et al. 1999, 2003). The present-day displacement of Arabia relative to Eurasia by several centimetres per year is recorded throughout the GC. The indentation of Arabia occurs at Bitlis-Zagros and deformation propagates towards the GC. This indentation has produced large strike-slip faults along which the Anatolian block escapes westward. In the GC, both strike-slip faults and thrusts actively accommodate deformation. Earthquake focal mechanisms reveal that the
Fig. 10. Photograph of reverse faults developed prior to the tilting of beds under a NE-SW compression (Late Cretaceous flysch-likeunit of the western GC) and stereoplots of the related stress tensors (calculatedfor both attitudes of beds: present-day and restored to horizontal stress tensors obtained from inversionof the fault slip data as givenby Saintot & Angelier2002). (Photograph by A. Saintot; J. Angelier for scale.)
whole Caucasian area is under a north-south compressional stress regime (Gushtchenko et al. 1993; Gushtchenko & Rebetsky 1994; Mikhailov et al. 2002), a continuation of the inferred Sarmatian palaeostress regime (Saintot & Angelier 2002). The 'Caucasian' N W - S E and W N W - E S E faults act as oblique reverse faults. The depth distribution of earthquakes is limited to the crust and the overlying sedimentary succession; no deeper earthquakes are observed, nor has a Benioff Zone been imaged. Earthquakes at depths of 10-15 km are related to strike-slip faults, whereas deeper hypocentres are along thrust faults. Also, it is observed that along single focal zones, the depths of hypocentres increase northwards (Gamkrelidze 2005) along gently north-dipping planes. In particular, the identified focal plane of the 29 April 1991 Racha earthquake (Mw = 7) exhibits a dip angle of 2 0 40 ~ north (Triep et al. 1995). The distribution of seismicity also indicates the propagation of the GC front southwards to the offshore Shatsky Ridge (a western prolongation of the Transcaucasus; Fig. 5), and to the Rioni and Kura basins (Figs 2 and 4). On seismic lines crossing the offshore western GC (Finetti et al. 1988), it can be observed that, with the continuing compression, the Tuapse Basin as a whole overthrusts the Shatsky Ridge with a southward propagation of the GC deformation front. Active thrusting of the GC also affects the Rioni and Kura basins. The Oligocene-Early Miocene sedimentary infill of these two basins has been incorporated into the south-vergent fold-and-thrust belt during the Mid-Miocene compressional phase. The faults transverse to the GC belt have been invoked as conduits for Quaternary volcanism (Milanovsky et al. 1984; Giorgobiani & Zakaraya 1989; Lordkipanidze et al. 1989; Koronovsky et al. 1997). These faults, which were very active during the Cenozoic, have segmented the GC and Transcaucasus area as well as the W N W - E S E 1250km long south-vergent frontal thrust of the GC (Giorgobiani 2004). Similarly, a large NE-striking left-lateral fault, with a reported offset of 90 km (Philip et aL 1989), was proposed as the conduit for the Kazbek volcano (see location of volcanoes shown in Fig. 4). However, the evidence for large strike-slip displacements along such structures in the GC belt and Transcaucasus area remains very speculative. S o m e characteristics o f the inversion o f the Greater C a u c a s u s Basin
Using simple area-balancing restoration of cross-sections, Ershov et al. (2003) estimated the amount of shortening along the GC to have been 2 0 0 - 3 0 0 k m (as also reported by Khain 1982; Zonenshain & Le Pichon 1986; Shcherba 1993; Nikishin et al. 1998b). Such an estimate is in agreement with the inferred plate kinematics of the area, which suggests a 400 km displacement of Arabia northwards to (fixed) Eurasia from Oligocene times and takes into account the amount of shortening in the Lesser Caucasus area. However, field observations (M. Stbrier) of the structural relationships between Mesozoic GC formations indicate that the shortening accommodated by the MCT sensu lato is of the order of some tens of kilometres and that each of the few other major thrusts should accommodate some 2 - 5 km (e.g. along the eastern part of one of the MCT branches, the Lower Jurassic units are thrust over themselves). It follows that the shortening across the GC as a whole could be much less, as little as 100 km. The central part of the GC belt, which has the highest elevation and the highest rate of Neogene to present uplift, corresponds to the thinnest part of the Aalenian GC rifted lithosphere. Thus, the anomalously high elevation in this area could be a consequence of the subduction of highly thinned continental lithosphere (if not partly oceanic, as mentioned earlier). The lithospheric root might also be comparatively more important in the central part of the GC because collision and shortening was concentrated there, directly in front of the indenting Arabian plate. The Quaternary and still active uplift of the central part of GC could
THE MESOZOIC-CENOZOIC TECTONIC EVOLUTION OF THE GREATER CAUCASUS
285
Fig. 11. Main subsidence-driving mechanisms for the foreland stage of basin evolution and uplift of the central part of the GC as a result of delaminationof the root (from Ershov et al. 2003).
be the result of delamination of a lithospheric root, as suggested by numerical modelling (Ershov e t al. 2003; Fig. 11) and tomography (Brunet e t al. 2000) results, with the Quaternary volcanism being linked to this deep process. Lithospheric roots would not have been so well developed in the western and eastern GC because there was less shortening there and it was accommodated differently. In the eastern GC, shortening is symmetrically accommodated by the fan-shaped development of foreland structures. To the west of the central GC, no large lithospheric root is expected because the cumulative shortening there is significantly less because of the western escape of Anatolia.
The structural style of the GC belt agrees with the inversion of a deep basin developed on very thin continental crust, perhaps similar to what Gamkrelidze & Giorgobiani (1990) referred to as 'intraplate subduction' in an intraplate setting. As such, the GC can be viewed as a Pyrenees or Atlas Mountains analogue (see, e.g. the overview of the Pyrenees by Grup de Geodin~mica i Anfilisi de Conques 2005). No lateral escape during shortening and consequent development of large nappes (and rootless nappes), such as in the Alps, occurred. The absence of any remnants of an ophiolitic suture supports such a model. Furthermore, there is no obvious record of any subduction zone along the GC
Fig. 12. Summaryof the tectonic evolutionof the Greater Caucasus. Absolute ages are from Gradstein et al. (2004).
286
A. SAINTOT ETAL.
during Mesozoic and Cenozoic times. There is no volcanic arc or blueschist and high-grade metamorphic rocks, and no accretionary complexes are present. (It is noted, however, that older, Palaeozoic, ophiolites and associated high-grade metamorphic rocks do crop out in the central part of the belt.)
Conclusions The crustal structure of the Greater Caucasus remains a matter of debate, and two different models have been postulated. One model considers the GC belt as a former deep marine Mesozoic basin that was subsequently squeezed between steep crustal faults, these faults separating the GC from its adjacent tectonic units, the Transcaucasus and the SP. The alternative model considers the GC as a south-vergent, crustal-scale, imbricated fold-and-thrust belt with the SP thrust over the Transcaucasus massif along north-dipping planes, which flatten at depth. More and better geophysical data are needed to discriminate between these two models. However, the latter appears in general to satisfy better the available data, although some interpretations remain questionable (such as the geometry of the boundary fault zone between the GC and the SP and the amount of shortening in the GC belt). The tectonic evolution of the Greater Caucasus during Mesozoic and Cenozoic times can be summarized as follows (see Fig. 12): (1) Permo(?)-Triassic rifting; (2) Eo-Cimmerian shortening related to collision of the Iranian Block with Europe; (3) development of Early-Mid-Jurassic rift basins, possibly related to north-dipping subduction south of the Transcaucasus (i.e. in the Lesser Caucasus); (4) development of a Bathonian (Mid-Cimmerian) unconformity related either to the syn- to post-rift transition or to a collisional event at the active margin; (5) M i d - L a t e Jurassic to Eocene post-rift subsidence; (6) Late Eocene basin inversion related to the final closure of the Tethys oceanic domain; (7) a second shortening phase from Late Miocene time to the present accompanied by uplift and magmatism and corresponding to the final stages of A r a b i a - E u r a s i a collision. This paper has benefited from many fruitful discussions with Russian colleagues, including S. Korsakov, P. Fokin and P. Tikhomirov. Part of the research was funded by the MEBE programme and, in the past, by the Peri-Tethys Programme (A. S. would especially like to thank A. Ilyin for his help in the field during 'PeriTethys years', as well as J. Angelier for her earlier work). The Netherlands Research Organization (NWO/ALW) funded part of A.S.'s research, and F.Y.'s research was partly supported by NATO 1997 (202025D). The authors also thank M. L. Somin, L. M. Rastsvetaev and A. V. Marinin, who kindly discussed some important scientific aspects of the manuscript, as well as the two reviewers, D. Brown and A. Okay, whose comments led to improvements incorporated in the present manuscript, which is Netherlands Research School of Sedimentary Geology contribution 2005.05.02.
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SAINTOT, A. & ANGELIER, J. 2002. Tectonic paleostress fields and structural evolution of the NW-Caucasus fold-and-thrust belt from Late Cretaceous to Quaternary. Tectonophysics, 357, 1-31. SAINTOT, A., CHALOT-PRAT, F., MCCANN, T., FOKIN, P., KORSAKOV, S. & STEPHENSON, R. 2004. The Early Middle Jurassic basins of the Western Greater Caucasus (Russia). AAPG European Region Conference with GSA, Prague, 10-13 October 2004. SI~BRIER, M., POLINO, R., GALKIN, V. A. & YUNGA, S. 1997. Caucasus stress evolution from Mesozoic to present. LUG 9, Strasbourg, 23-27 March 1997. Terra Nova, 9 (Abstracts Supplement 1), 337. ~ENGOR, A. M. C. 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia. Geological Society of America, Special Papers, 195. SHARAFUTDINOV,V. F. 2003. Geological structure and patterns of development of the Maykopian deposits on the North-Eastern Caucasus in connection with oil-and-gas content. DSc thesis, Lomonosov Moscow State University, Moscow [in Russian]. SIqARDANOV, A. N. & PEKLO, V. P. 1959. Tectonics and history of formation of buried folds on southern flank of western Kuban trough and hydrocarbon potential of Mesozoic deposits. Gostoptekhizdat, Moscow, 3-27 [in Russian]. SHCHERBA, I. G. 1987. Olistostromes and problems of Cenozoic tectonics of the Caucasus. In: MILANOVSKY, E. E. & KORONOVSKY, N. V. (eds) Geology and Mineral Resources of the Great Caucasus. Nauka, Moscow, 191-200 [in Russian]. SHCHERBA, I. G. 1989. Paleogeography and tectonics of Maikop basin, Caucasus. Bulletin of MOIP, Geologia, 306(5), 1196-1200 [in Russian]. SHCHERBA,I. G. 1993. Stages and Phases of Cenozoic Evolution of Alpine Domain. Nauka, Moscow [in Russian]. SHEMPELEV, A. G., PRUTSKY,N. I., FELDMAN, I. S. & KUHMAZOV,C. U. 2001. Geological-geophysical model along cross-section TuapseArmavir. The tectonics of Neogey (Pz + Mz + Cz)---common and regional aspects. GEOS, 2, 316-320 [in Russian]. SHEMPELEV, A. G., PRUTSKY, N. 1., KUHMAZOV, C. U., P'YANKOV, V. YA., LIGIN, V. A. & MOROZOVA, A. G. 2005. The materials of geophysical investigation along Near-Elbrus cross-section (volcano Elbrus-Caucasian mineral waters). The tectonics of Earth crust and mantle. The tectonic regularity of mineral resources localization. GEOS, 2, 361-365 [in Russian]. SHREIDER, A. A., KAZMIN, V. G. & LYGIN, V. S. 1997. Magnetic anomalies and age of the Black Sea deep basins. Geotectonics, 31(1), 54-64. SOMIN, M. L. 2000. About structure of axial zones of Central Caucasus. Reports of the Russian Academy of Sciences, 375(5), 662-665 [in Russian]. STEPHENSON, R. A., WILSON, M., DE BOORDER, H. & STAROSTENKO, V. I. (eds) 1996. EUROPROBE: intraplate tectonics and basin dynamics of the Eastern European Platform. Tectonophysics, 268. STOCKLIN, J. 1968. Structural history and tectonics of Iran: a review. AAPG Bulletin, 52(7), 1229-1258. TANKARD, A. J. & BALKWILL, H. R. (eds) 1989. Extensional Tectonics and Stratigraphy of the North Atlantic Margins. AAPG, Memoirs, 46.
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ZONENSHAIN,L. P. & LE PICHON, X. 1986. Deep basins of the Black Sea and Caspian Sea as remnants of Mesozoic back-arc basins. Tectonophysics, 123, 181 - 211. ZONENSHAIN, L. P., KUZMIN, M. I. & NATAPOV, L. M. 1990. Geology of the USSR: a plate tectonics synthesis. American Geophysical Union, Geodynamics Series, 21.
The Western Accretionary Margin of the East European Craton: an overview T. C. P H A R A O H 1, J. A. W I N C H E S T E R 2, J. V E R N I E R S 3, A. L A S S E N 4 & A. S E G H E D I 5
1British Geological Survey, Kingsley Dunham Centre, Keyworth NG12 5GG, UK (e-mail:
[email protected]) 2School of Physical and Geographical Sciences, Keele University ST5 5BG, UK 3Ghent University, Palaontologie, Krijgslaan 281/$8, B 9000, Gent, Belgium 4Copenhagen University, Ostervoldgade 10, DK 1350, Copenhagen, Denmark 5Geological Institute of Romania, 1 Caransebes St, 012271 Bucharest 32, Romania
Abstract: Multidisciplinary investigations of the western margin of the East European Craton (EEC) by EUROPROBE projects since
1992 have confirmed that the Trans-European Suture Zone (TESZ) is the most fundamental lithospheric boundary in Europe, extending 2000 km from the North Sea to the Black Sea-Crimean region. The crust of the EEC is thicker and denser than that of Phanerozoic-accreted Europe, and the base of the lithospheric mantle significantly deeper. These characteristics persist throughout the length of the TESZ, despite the variation in age of the accreted crust along strike. Geological studies of key deep borehole cores and the limited outcrop data confirm that the crust of Phanerozoic-accreted Central Europe comprises a number of terranes, each thought to be derived from Gondwana during several episodes of rifting, ocean formation, ocean destruction and sequential accretion to the EEC throughout Palaeozoic time. There is still much discussion about the identity, provenance and history of these orogenic terranes. The process of accretion led to the formation of terrane-bounding orogenic sutures, which may be marked in outcrop by ophiolitic and eclogitic relics. Recognition of concealed sutures is obviously more difficult,and relies on a variety of geophysical techniques, used in an integrated way by multidisciplinaryteams; the evidence from deep seismic reflection and refraction surveys, teleseismic tomography, magnetotelluric experiments and from geophysical potential-field modelling is crucial for such studies. Since the European Geotraverse, much has been learnt about the geometry of the Thor, Iapetus, Rheic, Saxo-Thuringian and Moldanubian oceanic sutures, through the crust and sometimes into the mantle. This has led to a much better understanding of the 3D crustal structure of the Western Accretionary Margin of the EEC, and the lithospheric processes that have shaped it. From this, the influence of tectonic heterogeneities within the orogenic crust on the development of post-orogenic structures and basins can be much better constrained.
The Western Accretionary Margin of the East European Craton (EEC) is the most fundamental lithospheric boundary in Europe, separating the thick, cold, ancient crust and lithosphere of the exposed Baltic Shield and the partly concealed EEC from the younger, warmer and much thinner crust and lithosphere of Western Europe (Gee & Zeyen 1996), and extending deep into the mantle (Zielhuis & Nolet 1994; Babugka et al. 1998; Plomerova et al. 2002), perhaps as deep as 250 km. The transition from the thick crust and lithosphere of the EEC to the thinner crust of the accreted margin takes place over a broad zone some 400 km wide and 2000 km long, extending from the North Sea to the Black Sea (Fig. 1), for which the name 'Trans-European Suture Zone' was coined by the E U R O P R O B E Programme (Gee & Zeyen 1996). This region, recognized as a particularly significant lithospheric boundary, was identified as a key target for EUROPROBE research (Gee & Beckholmen 1993). For the past 10 years, it has been the focus of multidisciplinary investigations within E U R O P R O B E ' s Trans-European Suture Zone (TESZ) Project. Over most of its length this zone is concealed by deep sedimentary basins of Permian to Cenozoic age; thus geophysical experiments and multidisciplinary studies of samples from deep boreholes are crucial to understanding its history. A number of major WNW-trending crustal lineaments, in particular the Sorgenfrei-Tornquist (STZ), Teisseyre-Tornquist (TTZ) and Elbe Lineaments, are present within the TESZ. The significance of these has long been recognized, as a consequence of the influence they have persistently exhibited during the evolution of overlying late Palaeozoic and Mesozoic sedimentary basins, and as loci for Alpine inversion (Berthelsen 1992a). EUROPROBE seismological research has confirmed that the STZ and TTZ are steep features associated with displacements of up to 5 km at the Moho level. Although these are spectacular crustal features, they do not represent the original, orogenic sutures between the orogenic terranes making up the collage of 'Old' and 'Young' Europe, however. For instance, seismic reflection profiling in the Danish and north German areas indicates that the original
(Ordovician) oceanic suture between Baltica and Avalonia dips at an angle of about 15 ~ through the crust, such that the crust of Baltica extends SW some 140 km beneath the crust of Avalonia (Bayer et al. 2002); the Elbe Lineament corresponds to the SW limit of Baltica at the Moho level. In Central Europe too, seismic reflection experiments suggest that younger Palaeozoic orogenic sutures associated with the accretion of the Variscide terranes of the 'Armorican Terrane Assemblage' (Tait et al. 1997) to the EEC also dip at moderate angles through the crust. The relationship between the inclined orogenic sutures and the steep lineaments is complex, and the latter cannot have developed solely as a consequence of crustal-scale reactivation of the former. Rather more likely is reactivation at the lithospheric scale, with the fundamental differences in the lithospheric properties of 'Old' and 'Young' Europe resulting in a variety of reactivation styles along the various early formed lineaments. The crust of Europe to the SW of the EEC comprises a mosaic of orogenic terranes accreted throughout Phanerozoic time (Ziegler 1982, 1990), 'traditionally' (i.e. in the 20th century) regarded as developing during a series of distinct orogenic cycles, notably the Caledonian, Variscan and Alpine orogenic cycles. The application of modern methods of analysis, in particular high-precision radiometric techniques, has revealed that the traditional interpretation is too simplified: each 'orogeny' comprises a number of distinct deformation phases, corresponding to the terrane evolution described above; specifically to phases of ocean destruction, terrane docking (or 'soft collision'), collision (between larger crustal blocks) and dispersal. There is overlap between some cycles, particularly in late Palaeozoic time, when rifting of Gondwana proceeded virtually without interruption (Stampfli & Kozur 2006). In some places (e.g. the Rheno-Hercynian Basin), the continuous record of synorogenic clastic sedimentation from Frasnian to late Westphalian appears to indicate a continuum of deformation, rather than discrete phases (W. Franke, pers. comm.). The principal evidence for the identity (and duration of existence) of individual terranes comes from the uniqueness (endemicity) of
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 291-311. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. Revised basement tectonic sketch map of the TESZ and adjacent areas. (Compare with fig. 4.1 of Gee & Zeyen (1996) and fig. 1 of Pharaoh (1999).) Revisions incorporate information from Matte et al. (1990), Dallmeyer et al. ( 1995, 1999), Franke (1995b), Geluk (1997), Bula et al. (1997), Seghedi (1998), Franke & Zelainiewicz (2002), Verniers et al. (2002) and Winchester et al. (2002). Oceanic sutures, filled ticks; orogenic frontal zones, open ticks. Post-Palaeozoic basins and platforms: ADB, Anglo-Dutch Basin; ADF, Alpine Deformation Front; MNSH, Mid-North Sea High; NDO, North Dobrogea Orogen; NGB, North German Basin; POT, Polish Trough; RFH, Ringk0bing-Fyn High; RG, RCnne Graben; RMFZ, R0mr Fracture Zone; SP, Scythian Platform. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: BST, Bruno-Silesian (Brunovistulian) Terrane; BT, Bohemia Terrane; DSHFZ, Dowsing-South Hewett Fault Zone; EL, Elbe Lineament; EMT, East Moesian Terrane; IMF, Intra-Moesian Fault; KLZ, Krak6w-Lubliniec Zone; LRL, Lower Rhine Lineament; LT, Lysogory Terrane; MT, Matopolska Terrane; MDT, Moldanubian terranes; MST, Moravo-Silesian Terrane; NT, Normannian Terrane; PO, Palazu Overthrust; SNSLT, Southern North SeaLtineberg Terrane; SGF, Sfantu Gheorghe Fault; TT, Tulcea (North and Central Dobrogea) Terrane; WMT, West Moesian Terrane. Proterozoic-Palaeozoic tectonic elements: ABDB, Anglo-Brabant Deformation Belt; AD, Ardennes Massifs; AM, Armorican Massif; BB, Brabant Massif; BM, Bohemian Massif; CBT, Central Brittany Terrane; CDF, front of Caledonian deformation (see text for explanation); CM, Cornubian Massif; CPPR, Central Polish Palaeo-Rift; DR, Drosendorf Unit (of BM); EA, Ebbe Anticline; EEC, East European Craton; EFZ, Elbe Fault Zone; GF, Gf6hl Unit (of BM); HM, Harz Mountains; HCM, Holy Cross Mountains; LF, Loire Fault; L-W, Leszno-Wolsztyn Basement High; MC, Massif Central; MMC, Midlands Microcraton; MGCH, Mid-German Crystalline High; MH, Mazurska High; MN, Mtinchberg Nappe (of BM); MO, Moldavian Platform; NASZ, North Armorican Shear Zone; NBT, North Brittany Terrane; Pom, Pomerania; PP, Pripyat Trough; RM, Rhenish Massif; USM, Upper Silesian Massif (=MST); SH, South Hunsrtick; SNF, Sveconorwegian Front; SASZ, South Armorican Shear Zone; S-TZ, Sorgenfrei-Tornquist Zone; TB, Teplfi-Barrandian Basin (of BM); T-TZ, Teisseyre-Tornquist Zone; UM, Ukrainian Massif; VF, Variscan Front.
their palaeo-faunas and palaeo-floras (e.g. Cocks & Fortey 1982); and the principal evidence for their magnitude and direction of motion and rotation comes from palaeomagnetic constraints (e.g. Torsvik 1998). These topics are discussed in more detail elsewhere in this volume. It is relatively simple to establish faunal endemicity for the platform successions of palaeocontinents, from abundant shelly macrofauna present in little-deformed shelf sedimentary strata. It is much more difficult to do this for the accretionary margins of terranes, or indeed suspect terranes, frequently poorly exposed or known only from boreholes, comprising strongly deformed sequences of deep-water strata with sparse (if any) macrofauna. Palaeomagnetic studies of terrane margins suffer similar geological constraints. There is therefore vigorous debate about the status of many of the terranes reviewed here. Another important matter of debate is the provenance of individual terranes. The crystalline basement of most terranes in the
TESZ has historically been referred as being of 'Cadomian', 'Avalonian' or occasionally 'Pan-African' affinity, usually based on rather sparse and imprecise radiometric data. In the literature, usage of these terms has often been applied inconsistently, to lithofacies rather than to precisely dated rock suites. The increasing availability of large numbers of precise U - P b zircon ages has revolutionized the study of terrane provenance. Thus, although there is an abundance of Neoproterozoic zircon grains dated at about 600 Ma, other peaks in the grain population have been used to attribute sources in Eastern (North Africa) or Western Gondwana (Northern A m a z o n i a - G u y a n a ) and in Baltica. Such attribution depends on the presence of a robust database for the Precambrian shield areas for comparative purposes, and the quality of the palaeogeographical reconstructions used. Combined with other evidence, such as palaeomagnetic evidence for palaeolatitude, sedimentological evidence for palaeoclimate, and
WESTERN ACCRETIONARYMARGINOF THE EEC lithological association, this type of analysis can be a very powerful technique. The Palaeozoic time scale used throughout is that published by McKerrow & Van Staal (2000). Accreted terranes of the western margin of the EEC P a l a e o z o i c terrane motions, accretion and deformation phases
Most of the terranes were derived by phases of rifting along the margins of the Gondwana palaeocontinent, which lay at high southerly latitude for much of Palaeozoic time (Torsvik 1998). A large ocean, Iapetus, opened during late Neoproterozoic time, separating Gondwana from other large relics of the RodiniaPannotia supercontinent (Dalziel 1991, 1997), such as Laurentia and Baltica. That part of Iapetus separating the terrane of Avalonia from Baltica is referred to as the Tornquist Sea (Cocks & Fortey 1982). The subsequent history of the newly rifted terranes varies, but typically involves northward transport as a result of the creation of new oceanic crust between the newborn terranes and Gondwana; the destruction of older oceanic crust lying between the newborn terrane and the EEC, principally by subduction; the episodic accretion of the Gondwana-derived terranes to the EEC margin; and finally, dispersal along the margin of the EEC by strike-slip displacement. The location of the litho-tectonic elements described below is indicated in Figure 1. The Bruno-Silesian Terrane may have been one of the first to leave Gondwana, and was certainly accreted to Baltica by the time of the late Cambrian Sandomierz deformation phase (Znosko 1974; Winchester et al. 2002; Nawrocki et al. 2004). Next, Avalonia left Gondwana in the early Ordovician, migrating from high to low southerly latitudes throughout remaining Ordovician time (Trench & Torsvik 1992), driven by the opening of the Rheic Ocean (Cocks & Fortey 1982) to the south of Avalonia ('ridge-push'), as well as by rapid destruction of the Iapetus Ocean to the north ('slab-pull') in a number of subduction systems. Closure of the Tornquist Sea segment of this ocean involved a significant dextral oblique component (Trench & Torsvik 1992; Oliver et al. 1993). Soft collision ('docking') of Avalonia and Baltica, producing Balonia (Torsvik 1998), occurred during the Shelveian Phase in Ashgill time (Samuelsson et al. 2002b), and is associated with amphibolite-facies metamorphism in the Mid-North Sea region (Frost et al. 1981; Pharaoh et al. 1995). Northward drift of Baltica in late Cambrian-early Ordovician time was accompanied by 55 ~ counter-clockwise rotation (Torsvik & Rehnstr6m 2001). Terranes in central Poland with a basement of supposed 'Cadomian' affinity and Acado-Baltic faunal association in the Cambro-Ordovician (e.g. the inferred M a t o p o l s k a Terrane) are the most controversial. They may have migrated from the vicinity of the southern Urals along the Tornquist margin of Baltica (Pharaoh 1999); or crossed the Iapetus Ocean from Gondwana to Baltica prior to mid-Cambrian time, as envisaged for the Bruno-Silesian Terrane (Betka et al. 2000, 2002); or they may always have been located close to their present position (Cocks 2002). A foredeep developed along the margin of Baltica in Silurian time (Dadlez et al. 1994; Berthelsen 1998) as a result of loading caused by the newly accreted crust. Subduction continued to the NW beneath Laurentia, leading to final closure of the Iapetus Ocean in Wenlock time (Leggett et al. 1979; Kneller et al. 1993), marked by the Scandian event in NW Scotland and the amalgamation of Laurussia (Ziegler 1990). The terranes now comprising the internides of the Variscan orogen and exposed in the Bohemian, Armorican and Iberian massifs, were located along the Gondwana margin at high southerly palaeolatitudes in late Ordovician time (Krs et al. 1986; Tait et al. 1995). Persistent plume-induced magmatism (Floyd et al. 2000) resulted in several phases of tiffing from the margin; rapid dispersal northward was driven by subduction of
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the Rheic Ocean (Franke 1995), and the opening of the Saxo-Thuringian and M a s s i f Central oceans closer to Gondwana. An early accretion of at least part of Saxo-Thuringian Armorica to Laurussia is recorded by emplacement of the Lizard Peridotite and the Acadian deformation phase, in Emsian time (Soper et al. 1987; Pharaoh 1999). Some parts of the 'Armorican Archipelago' (Franke et al. 1999) collided at c. 400 Ma, with significant HP metamorphism, predating their amalgamation into the TESZ. Provenance studies indicate that Iberia remained attached to Gondwana until early Devonian time (Martinez-Catal~in et al. 2004). Palaeomagnetic evidence (Tait et al. 1997, 2000) supports independent motion of distinct M o l d a n u b i a n and Perunica terranes until at least late Devonian time (c. 370 Ma) when collision with Saxo-Thuringian Armorica occurred. In early Carboniferous time, widespread HP metamorphism in the Variscan internides records rapid crustal thickening following closure of the Massif Central Ocean and collision with Iberia (Ziegler 1990). The terminal phase of collision between Laurussia (the Old Red Continent) and Gondwana, to produce Pangaea, is recorded by late Carboniferous (Variscan) and early Permian (Alleghenian) orogenic phases in Europe and America. The crust of the TESZ continued to undergo modification as a result of post-orogenic 'reordering' (Meissner 1989), Permian to Mesozoic basin development, rifting along the pre-Alpine Tethyan margin, Cimmerian inversion and subsequent Alpine-Carpathian thrusting, particularly in Romania.
Terrane analysis
The principles of terrane analysis (e.g. Coney et al. 1980) have been successfully applied to the TESZ (Franke 1990; Po~aryski 1990; Pogaryski et al. 1992). The principal characteristics of the major Palaeozoic terranes depicted in Figure 1 are briefly reviewed here. Geographical extent is described using present geographical coordinates, locations and distances, without palinspastic reconstruction. Provenance studies, using characteristic isotopic, geochemical or biostratigraphic assemblages, aim to identify the source continent of a rifted terrane. Of particular value are studies of detrital zircon suites using the single-crystal or SHRIMP (sensitive high-resolution ion microprobe) methods of U - P b isotopic analysis. The internal structure of a terrane is deduced from outcrop studies and seismic reflection data. Particularly informative for this purpose are images generated from geophysical potential fields (e.g. aeromagnetics and gravity; Banka et al. 2002). Examples of such maps, with a structural overview template, are presented in Figures 2 and 3. Of course, not all the features visible on these images are associated with the basement; for example, Permo-Carboniferous volcanic rocks locally cause magnetic anomalies, particularly close to the mid-North Sea rifts and in the North German Basin (Fig. 2), and the expression of younger sedimentary basins (which may represent extensional reactivations of basement structures) is clear in the Bouguer gravity image (Fig. 3). For further discussion of these topics the reader is referred to Banka et al. (2002), who have also listed and acknowledged the numerous sources of these data. Rifting history is typically determined by sequence stratigraphic studies, subsidence rates or magmatic episodes. The period of terrane isolation is most easily deduced from endemic faunal assemblages (e.g. see Cocks & Torsvik 2006), but also sometimes from isotopic signatures (e.g. Thorogood 1990; Samuelsson et al. 2002a). To determine the direction, rotation and rate of drift requires an excellent palaeomagnetic record, but can be done successfully (e.g. for Baltica in Cambro-Ordovician time; Torsvik & Rehnstr6m 2001). The history of ocean closure is deduced from arc-related magmatic suites. Timing and nature of terrane collision are deduced from the breakdown of faunal endemism, the arrival of orogenic flysch sediment at the foreland (e.g. Franke 2000), structural evidence and isotopic
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T.C. PHARAOHET AL.
Fig. 2. Colour shaded relief map of aeromagnetic potential field of the Western AccretionaryMargin of the EEC, after Banka et al. (2002), who listed the data sources. Colour scale ranges from red (+2000 nT) to green (0 nT) to blue (-700 nT). The localizedcontribution of magnetic sources shallowerthan the basement, (e.g. Permian volcanic rocks near the North Sea graben intersections), should be noted. Key to basin, platform, terrane and other tectonic elements as in Figure 1.
data. The geometry and location of terrane-bounding sutures is deduced by mapping all of the above characteristics, structural evidence, and seismic reflection and other geophysical data. Finally, the post-accretion history is deduced from structural evidence, sedimentology of overstep sequences, reworking of microfossils, isotopic age of stitching plutons, etc. T e r r a n e s a c c r e t e d in e a r l y P a l a e o z o i c t i m e Baltica. This
large palaeocontinent comprised the exposed Fennoscandian and Ukrainian shields and other parts of the concealed EEC, extending eastward from the North Sea to the Urals, and northward to the Arctic Ocean to include the Timanides. Its NW boundary lies within the Scandinavian Caledonides, where it was overthrust by terranes of Laurentian (North American) affinity (Fig. 1) in Silurian time. The long and complex evolution of the Precambrian crust of this palaeocontinent has been described by Bogdanova et al. (2006). The crust of Baltica is characterized by high-frequency aeromagnetic anomalies (Fig. 2), whose zonation bears witness to a complex Proterozoic accretion history (Banka et al. 2002; Williamson et al. 2002). A phase of mafic magmatism ('Older Dykes' and Volhyn Basalts) and tiffing ('Sparagmite Basins') at c. 600 Ma (Andr~asson 1998) reflects rifting from the rest of the RodiniaPannotia supercontinent (Dalziel 1991, 1997) in the Neoproterozoic. The Cambro-Ordovician palaeogeographical history deduced from biostratigraphic and palaeomagnetic data has been described by Cocks & Torsvik (2006). The whole palaeocontinent is thought to have rotated counter-clockwise by at least 100 ~ in late
Neoproterozoic to mid-Ordovician time, with at least 55 ~ of this in late Cambrian to early Ordovician time (Torsvik & Rehnstr6m 2001), a fact of considerable importance for palaeogeographical evaluation (Cocks 2002). Thin Cambro-Ordovician strata are of platformal type with rich shelly faunas. After collision with Avalonia in late Ordovician (Ashgill) time, when faunal isolation ended (Cocks et al. 1997), a rapidly subsiding foredeep developed along the SW margin of the EEC (Dadlez et al. 1994; Poprawa et al. 1999). Boreholes in Denmark (Vejbaek 1997), northern Germany (Katzung et al. 1993) and Poland (Dadlez 1982) prove up to 7 km of basinal Silurian strata. The latter were incorporated in a northward-vergent foreland thrust belt of Scandian (late Silurian) age. The so-called 'Caledonian Deformation Front' (Fig. 1) delimits the western edge of the autochthonous, lightly deformed Baltic platformal sequence. The Teisseyre-Tornquist Zone sharply truncates (Figs 2 and 3) NE-trending belts of granulites, anorthosites and granite-gneiss in eastern Poland and the Ukraine (Bogdanova et al. 1996). In Romania, the basement of the Moldavian Platform lithologically resembles that of the Ukrainian Shield, whereas that of the Scythian Platform comprises Neoproterozoic granitic and dioritic rocks (Neaga & Moroz 1987), atypical of the EEC. Subduction-related magmatism (other than ash-fall bentonites) is absent in Baltica; thus dominantly NE-directed subduction appears unlikely (Pharaoh 1999; Balling 2000); however, an alternative view has been given by Meissner et al. (2002). NE-dipping zones of reflectivity in the subcrustal lithosphere of Baltica may represent the relics of Proterozoic subduction or a late switch in Ordovician subduction polarity, or may post-date subduction entirely (Berthelsen 1998).
WESTERN ACCRETIONARYMARGIN OF THE EEC
295
Fig. 3. Colour shaded relief map of Bouguer gravity potential field of the Western Accretionary Margin of the EEC, after Banka et al. (2002), who listed the data sources. Rainbow colour scale ranges from red (+60 mGal) to yellow (0 mGal) to blue (-235 mGal). The local enhancementof 'basement grain' by basins resulting from post-orogenic extensionalreactivation should be noted. (See Banka et al. (2002) for further discussion). Key to basin, platform, terrane and other tectonic elements as in Figure 1.
The Holy Cross Mountains (Fig. 1) in south-central Poland are the largest exposure of Palaeozoic rocks within the TESZ. Two groups of strata, the Lysogdry unit in the north, and the Kielce unit, part of the largely concealed Matopolska Massif in the south, are separated by the WNW-trending Holy Cross Fault. The relationship of these units to the EEC during Cambrian time has been the subject of much recent debate. Po~aryski (1990) and Franke (1995a) recognized distinct L y s o g d r y and M a t o p o l s k a t e r r a n e s on the basis of stratigraphic and structural contrasts between these units (and the EEC). The crystalline basement of both units is unknown, but the aeromagnetic anomaly map (Fig. 2) indicates that highly magnetic basement (of typical EEC type) is absent. Ediacaran (Vendian) silty and volcaniclastic rocks (Buta et al. 1997) resemble those of the adjacent EEC (Vidal & Moczydtowska 1995). A prominent latest CambrianTremadoc unconformity (Matopolska Massif) records folding and low greenschist-facies metamorphism during the Sandomierz deformation phase (Znosko 1974). The traditional view is that the Cambrian shelly fauna is diagnostic of Baltica (Dzik 1983; Bergstr6m 1984; Ortowski 1992). This was challenged by Belka et al. (2002), who claimed that although the Ordovician faunas are certainly Baltican, the Cambrian faunas of the Lysogdry unit are unknown in Baltica, and those of the Matopolska Massif are dominantly Avalonian in aspect. Cocks (2002) pointed out that as Avalonia existed as a separate entity only in the Ordovician, such affinity cannot be assigned in the Cambrian. From a study of detrital zircon and muscovite ages, Betka and colleagues (Betka et al. 2000, 2002; Valverde-Vaquero et al. 2000) concluded
that the Matopolska Massif (and less certainly, the Lysogdry unit) was detached from Gondwana in early Cambrian time and accreted to Baltica by the end of the Cambrian. However, the recognition of Neoproterozoic basement in the southern Uralides (Glasmacher et al. 1999) and in the Scythian Platform suggests that the presence of 'Cadomian' age detritus is not necessarily diagnostic of a Gondwanan provenance (Winchester et al. 2002). Unrug et al. (1999) have speculated that the Matopolska Massif may have formed the accretionary wedge to the BrunoSilesian Terrane. Detrital mineralogical and biostratigraphic evidence apparently do not support such a linkage, however (Belka et al. 2002; Cocks 2002). In the Holy Cross Mountains (as in the rest of Baltica) a highly condensed Arenig-Lower Silurian carbonate-clastic sequence contains many bentonites. The Upper Silurian sequence comprises up to 1500 m of greywackes deposited on the EEC foredeep (Dadlez et al. 1994; Berthelsen 1998) and strongly affected by Scandian phase deformation (Tomczyk 1980; Dadlez et al. 1994). Palaeomagnetic evidence indicates that the Matopolska Massif may have been displaced dextrally along the TTZ (Lewandowski 1993), but the faunal evidence constrains any displacement to a maximum of a few hundred kilometres (Cocks 2002). Another possibility is that the Polish Trough may have a more ancient antecedence than hitherto realized. It might have been initiated as a rift (referred to as the 'Central Polish Palaeo-Rift' in Figs 1, 2, 3 and 6) controlled by the ancestral TTZ, during the Neoproterozoic break-up of the Rodinia-Pannotia supercontinent, analogous to rift structures seen elsewhere in
296
T.C. PHARAOHETAL.
the EEC (Puchkov 1998). Cambrian strata were deposited on the passive margin of the EEC (see Fig. 6) before undergoing at least partial inversion in the late Cambrian Sandomierz phase, as recognized in the Holy Cross Mountains. Subsequently, the rate of subsidence matched that of the rest of the EEC. Only the NW part of this proposed palaeo-rift then subsided in Carboniferous, Permian and Mesozoic time to form the Polish Trough. Bruno-Silesian Terrane. Also known as the Brunovistulian Terrane (Dudek 1980; Aleksandrowski & Mazur 2002; Nawrocki et al. 2004), this terrane lies at the eastern edge of the Bohemian Massif in SW Poland and the NE Czech Republic (Fig. 1), and is concealed towards the east by the Carpathian Orogen and its foreland basin. It has traditionally been regarded as a component of the Rheno-Hercynian Zone of the Variscides, whose Devonian-Carboniferous evolution it strongly resembles (W. Franke, pers. comm.). It comprises two main sub-regions: the 'Brunovistulian' Block (Dudek 1980), extending south to Brno and the Krems-Vienna Line, and a western part, reworked into parautochthonous Variscan nappes. The concealed Upper Silesian Massif is separated from the Matopolska Massif by the Krakdw-Lubliniec tectonic zone, a narrow (about 0.5 km wide) belt of polyphase ductile deformation and magmatic intrusion, representing a possible terrane boundary (Dadlez et al. 1994; Buta et al. 1997). Deep boreholes prove metasedimentary rocks dated at 610-580Ma and granites emplaced at c. 585Ma (Finger et al. 2000; Belka et al. 2002). Amphibolites from the Rzeszotary Horst have been dated at 2.5 Ga by the U-Pb SHRIMP method (Bylina et al. 2000). These characteristics indicate that the basement of this terrane probably has a Gondwanan provenance (Friedl et al. 2000). Finger et al. (2000) correlated it with Avalonia, but a distinct gap in the range of detrital ages at about 570-590 Ma in the latter terrane (Murphy et al. 2004) suggests that such a correlation is unlikely, as does the presence of Acado-Baltican trilobites (Orlowski 1975; Buta et al. 1997) in gently deformed Cambrian strata. Nawrocki et al. (2004) argued that the presence of the endemic taxon S c h m i d t i e l l i u s p a n o w i in both the Matopolska Massif and the Bruno-Silesian Terrane supports their proximity in early Cambrian time. Acritarchs have closest affinity with the EEC (Jachowicz & Pf-ichystal 1998) and with Iberia (Moczydtowska 1995), but apparently not the rest of Gondwana. Detrital zircon and muscovite ages in Cambrian strata suggested a Gondwanan source to Friedl et al. (2000) and Belka et al. (2002). According to Cocks (2002), neither of the Holy Cross blocks were part of the same terrane as Bruno-Silesia; the faunal evidence on whether the latter terrane was derived from Baltica or Gondwana (or neither) is currently inconclusive. Middle-Upper Ordovician carbonates contain Baltican conodont faunas (Belka et al. 2002). Lower Devonian to Namurian A strata, resting with gentle discordance on lightly deformed early Palaeozoic strata, underwent strong Variscan folding and thrusting. Clear contrasts with the adjacent Matopolska Massif support the view that the Krakdw-Lubliniec tectonic zone is a terrane boundary (Franke 1990; Po2aryski 1990; Dadlez et al. 1994; Buta et al. 1997), along which the Bruno-Silesian Terrane was sutured to Baltica prior to early Devonian time (Nawrocki et al. 2004), although probably not at its current location (M. Lewandowski, pers. comm.). Another provenance could be a Neoproterozoic source in the Scythian Platform or the Uralian margin of Baltica (Puchkov 1998; Glasmacher 1999), which in Ediacaran time faced Gondwana. As a compromise, Winchester et al. (2002) suggested that it might represent a 'bridge' between Baltica and Amazonian Gondwana. Palaeomagnetic data suggest that in the early Cambrian, the Bruno-Silesian Terrane lay near the Equator, far from the Avalonian margin of Gondwana (Nawrocki et al. 2004) at 40-50~ and reached its present location on the Tornquist margin of the EEC before mid-Ordovician time. A further alternative, more 'fixist' model, proposed by Zelazfiiewicz et al. (2001), envisages development
of the terrane at the Tornquist margin of Baltica in Neoproterozoic time, although this is not supported by subsidence modelling (Poprawa et al. 1999) and other regional geological considerations (Nawrocki et al. 2004). Thus, the Bruno-Silesian Terrane is a suspect terrane in early Cambrian time and its provenance is hotly debated. Winchester et al. (2002, 2006) have suggested that, once accreted to Baltica, the Bruno-Silesian Terrane may have acted as an orogenic promontory on the EEC margin predating accretion of Avalonia in late Ordovician time. The 'Moldanubian Thrust' defining the western limit of the terrane (Fig. 1), and associated dextral transpression (see later discussion in the section ' M o r a v i a n S u t u r e ' ) , may have played an important part in the geometrical development of the postulated Variscide orocline in this region (Schulmann et al. 1991, 1995; Franke 1995; Franke & Zelainiewicz 2002). Avalonia. The eastern part of this microcontinent extends from
southern Ireland to the Mid-North Sea High, northern Germany and Poland. Avalonian faunas are recognized in Ordovician strata in Britain and Ireland, south of the Iapetus Suture, in the Rheno-Hercynian nappes north of the Lizard Thrust, in the Brabant and Ardennes massifs of Belgium, and in the Northern Phyllite Belt of Germany (Dallmeyer et al. 1995; Cocks et al. 1997). Only that part of Avalonia lying east of the Atlantic Ocean, referred to as Eastern Avalonia, is discussed here. It includes a heterogeneous Neoproterozoic basement comprising metamorphosed magmatic and sedimentary rocks generated in volcanic arcs and marginal basins (Thorpe et al. 1984; Pharaoh & Gibbons 1994) accreted initially to the Rodinia-Pannotia supercontinent, and following break-up of the latter, to protoGondwana, between 680 and 545 Ma. Xenocrystic zircons with ages of about 1.45 Ga (Tucker & Pharaoh 1991) and Nd isotopic studies (Noble et al. 1993; Nance & Murphy 1996) indicate possible involvement of 'Rondonian' type (Northern Amazonia and Guyana) crust, suggesting affinities with the South American (western) part of Gondwana (Murphy et al. 2000; Winchester et al. 2002). The Midlands Microcraton has a thin cover of lightly deformed lower Palaeozoic strata. Flanking deep-water basinal successions in Wales, northern and eastern England and Belgium obscure the Precambrian basement, and at the extremities of the microcontinent (e.g. in the Lake District, beneath the southern North Sea and northern Germany) the latter may be attenuated or absent. Here, the crust probably comprises juvenile lower Palaeozoicaccreted material. A calc-alkaline magmatic arc is traced from northern England to Belgium and is inferred to result from SW-directed subduction of Iapetus-Tornquist oceanic lithosphere beneath Avalonia (Noble et al. 1993; Pharaoh et al. 1993), possibly with a significant oblique component (Pharaoh 1999). The progressive change in age of volcanic onset from Wales (Tremadoc), to northern England (late Llanvirn), eastern England (Caradoc) and Belgium (Ashgill) and rotation of the Welsh Basin from arc to back-arc position (Kokelaar et al. 1984; Stillman 1988) through Ordovician time (Pharaoh 1999) is compatible with palaeomagnetic evidence for the counter-clockwise rotation of Avalonia (Piper 1997) with respect to this subduction zone (Pharaoh et al. 1995). Deformation is strongest in the Acadian (early Devonian) slaty cleavage arc developed in the basinal areas (Turner 1949; Soper et al. 1987; Van Grootel et al. 1997), contiguous with the Anglo-Brabant Deformation Belt (ABDB; Winchester et al. 2002) of eastern England and Belgium (Fig. 1), where, once again, a strong rotational component has been postulated (Verniers et al. 2002). Granite plutonism in northern England is of early Devonian age. An earlier, Shelveian (Ashgill) phase of deformation recognized in the Welsh Borders (Toghill 1992) is localized along major crustal lineaments (Pharaoh et al. 1995). This may correlate with the inferred late Ordovician phase of deformation affecting the Ardennes massifs in Belgium (Verniers et al. 2002). The parautochthonous Rheno-Hercynian nappes of
WESTERN ACCRETIONARYMARGINOF THE EEC Cornubia, the Ardennes in Belgium and central Germany (Fig. 1) represent the southern margin of Eastern Avalonia, extensively reworked by the Variscan Orogeny (Cocks et al. 1997; Verniers et al. 2002) and containing areally restricted, but significant, exposures of early Palaeozoic rocks (Franke 2000). 'Far Eastern Avalonia'. The existence of a distinct southern North Sea-Ltineberg Terrane has been proposed on the basis of geophysical criteria (Franke 1995a; Pharaoh et al. 1995) and more recently dubbed 'Far Eastern Avalonia' (Winchester et al. 2002). A possible terrane boundary with Avalonia proper may lie in the vicinity of the Dowsing-South Hewett and Lower Rhine Lineaments (Lee et al. 1993; Pharaoh et al. 1995). A small, perhaps marginal, oceanic basin may have been subducted here, giving rise to short-lived (Caradoc-Ashgill) volcanism in the ABDB (Verniers et al. 2002). Unfortunately, the available geophysical data provide little information on the internal structure or composition of this crust. One of the very few basement provings by deep boreholes in this region, the A/17-1 granite in the Netherlands sector, emplaced at 410 + 7 Ma (A. Gerdes, pers. comm.), contained no evidence of older crustal inheritance, compatible with the presence of only juvenile crust in this region. Boreholes on the Mid-North Sea and Ringk~bing-Fyn Highs penetrated metamorphic rocks of uncertain provenance. Whole-rock 4~ plateau ages indicate prograde greenschist-amphibolite metamorphism at 450-425 Ma, with retrogression at 415-400 Ma (Frost et al. 1981). The older age group has been interpreted as the age of docking or 'soft collision' of Avalonia with Baltica (Pharaoh et al. 1995; MONA LISA Working Group 1997b; Torsvik 1998; Pharaoh 1999), during the Shelveian deformation phase, an interpretation supported by biostratigraphic evidence for the reworking of microfossils (Samuelsson et al. 2002b). Gneisses in the Hunsrtick (SE Rhenish Massif) within the Rheno-Hercynian Zone (Fig. 1), yield U - P b ages in the range 560-574 Ma (Baumann et al. 1991) and may represent the only exposed Precambrian basement in this terrane (Winchester et al. 2002). Near Rtigen Island, in the southern Baltic (Fig. 1), deep boreholes encounter thrust and deformed anchizonal graptolitic greywackes of Ordovician age (Katzung et al. 1993) in the presumed hanging wall of the Thor Suture. They yield early Ordovician acritarchs similar to those of the English Lake District (Servais & Molyneux 1997). Many species are common to other periGondwanan areas in early Ordovician time (e.g. Spain, Bohemia and the Taurides of Turkey). Lithologically similar rocks and fossil assemblages are found in the Skibn6 Borehole in Pomerania (Cocks 2002; Samuelsson et al. 2002b). As acritarchs are planktonic, they cannot definitively indicate palaeocontinental affinity (Cocks & Verniers 2000), but they do suggest that terranes of probable Avalonian affinity extend eastward as far as northern Poland. It should be noted that Figure 1 represents a significant modification of earlier mapping (e.g. Pharaoh 1999) in this regard. Ashgill strata in the G14 Borehole lying just to north of the Thor Suture contain reworked Llanvirn microfossils with clear Gondwanan affinities (Samuelsson et al. 2002b), thus constraining the docking event to a period of about 10 Ma. Detrital muscovites from these same strata yield an 4~ plateau age of about 609 Ma, compatible with a Neoproterozoic provenance of Gondwanan affinity (Dallmeyer et al. 1999). There is also a significant detrital contribution from an unknown, immature volcanic arc (Giese et al. 1994; McCann 1998). Geophysical evidence, primarily from aeromagnetic data, supporting the presence of such a 'lost arc' concealed within 'Far Eastern Avalonia', was presented by Williamson et al. (2002). The Rtigen sequence was subsequently overthrust onto Silurian strata of the EEC foredeep during the Scandian (late Silurian) deformation phase. Dallmeyer et al. (1999) inferred that the Loissin-1 Borehole proved a culmination of the EEC (Pharaoh 1999; Fig. 1), emphasizing the low dip angle of the suture (see Fig. 5).
297
T e r r a n e s a c c r e t e d in late P a l a e o z o i c t i m e The Armorican Archipelago. The Rheic Suture represents a fundamental divide within the Variscan Orogen, separating crust reworked into a foreland thrust belt along the southern margin of Avalonia (Rheno-Hercynian or Externide Zone) from crust of the Internide Zones, which separated from Gondwana after the Ordovician (Dallmeyer et al. 1995). The Variscan internides have long been referred to as the Saxo-Thuringian and Moldanubian Zones (Kossmat 1927). These are more complex than the Rheno-Hercynian Zone (RHZ), incorporating ancient crust rifted from Gondwana at high southerly palaeolatitudes in early Silurian (Ziegler 1990) to earliest Devonian time (Paris 1998; MartfnezCatalan et al. 2004). Significant plume-related magmatism from about 500 Ma initiated tiffing at the Gondwana margin (Floyd et al. 2000; Crowley et al. 2002b) and facilitated generation of a progression of terranes referred to as the Armorican Terrane Assemblage (Tait et al. 1997) or Armorican Archipelago (Franke et al. 1999). The first convergence of the internide terranes with Laurussia is believed to have occurred in late Silurian-early Devonian time (Cocks & Fortey 1982), and was a possible cause of the Acadian deformation phase found throughout Eastern Avalonia (Soper et al. 1987). Subsequent collisions gave rise to the various phases of the Variscan Orogeny.
Saxo-Thuringian terranes. The Saxo-Thuringian Zone (STZ) as defined by Kossmat (1927) can be traced from SW Poland (Fig. 1), through central Germany and northern France, possibly extending as far as the Man of War rocks off the Lizard in SW England (Sandeman et al. 1997). The only good exposure is in the northern part of the Bohemian Massif, however, Franke (2000) recognized the following components: Franconia Terrane; Vesser Rift Basin; Saxo-Thuringia Terrane; the Saxo-Thuringian Ocean Basin. The microcontinental terranes comprise Neoproterozoic basement of 'Cadomian' affinity, largely greywackes and granitoid intrusions (Hammer et al. 1998; Linnemann et al. 1998); Cambrian-early Ordovician shallow marine clastic strata; bimodal magmatism at 500-480 Ma (Fumes et al. 1994; Sandeman et al. 1997; Krrner & Hegner 1998; Floyd et al. 2000); mid-late Ordovician hemipelagic shales, turbidites and glacigene strata (Erdtmann 1991); and pelagic shales, cherts and carbonates of Silurian to mid-Devonian age. The MidGerman Crystalline High, associated with a prominent aeromagnetic anomaly in Figure 2, represents the active magmatic margin resulting from southward subduction of ocean crust in Silurian-Devonian time (Franke 1998). Late Devonian-Visran flysch was fed NW from the developing orogenic belt. The apparent lateral continuity of the zone (Fig. 1) favours interpretation as (one or several) forearc or arc terranes accreted to a number of microcontinental terranes now forming the Moldanubian Zone (Ziegler 1990). Palaeomagnetic data (Krs et al. 1986; Tait et al. 1995) support derivation of the STZ terranes from Gondwana at high southerly palaeolatitudes after late Ordovician time. Suspect terranes ( s e n s u Coney et al. 1980) are found at the northern and southern margins of the STZ: the Lizard Peridotite (Clark et al. 1998) and Giessen Ophiolite (Franke 1995) formed at c. 397 Ma (early Devonian), either as marginal basins bordering the Rheic Ocean (Ziegler 1982, 1990) or as relics of Rheic mid-ocean ridge (Franke 2006). In the south, the Cambro-Ordovician Marifinsk~ Lfizn~ Complex (St~drfi 1999; Crowley et al. 2002a) and Silurian Sl~2a-Klodzko Ophiolite (Oliver et al. 1993; Floyd et al. 2002), are interpreted as accreted relics of the Saxo-Thuringian Ocean. A prolonged period of accretional thickening in the Sudetes, associated with HP metamorphism at c. 380-365 Ma (Maluski & Pato~ka 1997; Marheine et al. 2002) was followed by rapid uplift and greenschistfacies retrogression to 340 Ma (Kryza et al. 1990). Late and posttectonic granitoids (e.g. Karkonosze) were emplaced at c. 333 Ma
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(Kr6ner et al. 1994). The granulites of the Saxonian Dome and Erzgebirge (see Fig. 8), which also experienced high-grade metamorphism at c. 340 Ma (Kr6ner & Hegner 1998), were tectonically emplaced (or intruded?) beneath the floor of the contemporaneous early Carboniferous flysch basin (Franke et al. 1999; Franke & Stein 2000). Bohemia Terrane. This terrane (Franke & Zelainiewicz 2000) corresponds to the northern part of the Perunica Terrane recognized by Havff6ek et al. (1994). The boundary of the Moldanubian Zone with the STZ is poorly exposed, except in the western part of the Bohemian Massif (Fig. 1). The upper levels of the Teplfi-Barrandian Unit comprise slightly deformed Neoproterozoic volcanic rocks overlain by a Cambrian-middle Devonian sedimentary cover (Chlupfi6 1993). Deformation is mainly of Variscan age. The Bohemia Terrane forms a crustal block separating NW- and SE-verging parts of the Variscides (Matte et al. 1990), reflecting the opposing polarity of subduction zones closing the inferred Saxo-Thuringian and Massif Central oceans (Fig. 1). Palaeomagnetic evidence indicates 140 ~ counterclockwise rotation of the terrane before late Devonian time (Tait et al. 1997), apparently independently of the STZ terranes. The high-grade part of the Teplfi Unit (Czech Republic) and Erbendorf-Vohenstrauss Zone (Germany) form the root zone to allochthonous nappe outliers of Moldanubian Zone rocks overlying the STZ (Franke 1989), represented by the Mfinchberg, Wildenfels and Frankenberg massifs in Germany. These comprise paragneiss and orthogneiss with ophiolitic protolith metamorphosed to eclogite facies at c. 400-380Ma (Gebauer & Grtinnenfelder 1979) with Nd model ages c. 100 Ma older. Pressures >25 kbar (Klemd et al. 1994; O'Brien, pers. comm.) indicate subduction of the narrow Saxo-Thuringian Ocean to >75 km depth beneath Bohemia (Franke 2000). The G6ry Sowie of the Polish Sudetes (Cymerman et al. 1997; Kr6ner et al. 1994) lie in a similar structural position and are metamorphosed to granulite facies. Further collisions with STZ arc terranes caused amphibolite-facies retrogression at 370 Ma (Timmermann et al. 2000). Further HP granulite metamorphism at c. 340 Ma affecting both intemide zones in central Germany may be a consequence of lithospheric delamination or crustal thickening (Franke et al. 1999), probably during collision with Laurussia. Nappe emplacement occurred in latest Visgan time c. 330-325 Ma (Franke 1998). Moldanubian Terrane. The Teplfi Unit is separated from the high-
grade Gf6hl and Drosendorf units of the Bohemian Massif by steep ductile shear zones on its western and southern sides (Rajlich 1987; Zulauf 1994). The structurally higher Gf6hl Unit comprises anatectic ortho- and paragneisses and felsic granulites. HP metamorphism at c. 400Ma (Pin & Vielzeuf 1983) was followed by widespread amphibolite-facies metamorphism at c. 340 Ma (as described above) and emplacement of late tectonic granites. An accretionary complex of imbricated early Palaeozoic (c. 480 Ma) oceanic crust and passive margin components, it is comparable to the Massif Central Terrane (Matte et al. 1990). The Drosendorf Unit, overthrust by the Gf6hl unit, comprises >6 km of pelitic metasediments with a Neoproterozoic-lower Palaeozoic protolith interpreted as a passive margin sequence, similar to that of the C6vennes-Vend6e Terrane (Matte et al. 1990). The presence of ophiolitic fragments within this ductile shear zone again indicates the likely presence of an oceanic suture here. Romanian Terranes. The Trans-European Suture Zone reappears
from beneath the Carpathian Orogen and its foreland basin in Romania (Fig. 1) NW of the Black Sea. In the Dobrogea region at least three and possibly four, distinct fault-bounded terranes, described below, are recognized within the Carpathian Foreland, all suspect with respect to the Scythian Platform and Moldavian Platform of the EEC. Cimmerian inversion structures control the
disposition of the pre-Mesozoic basement blocks, but do not represent simple reactivations of the original terrane boundaries. (1) Tulcea (North a n d Central Dobrogea) Terrane. In North and Central Dobrogea, north of the (pre-Jurassic) Palazu Overthrust (Visarion et al. 1979), the crystalline basement comprises dismembered ophiolitic and metasedimentary rocks interpreted as a Neoproterozoic accretionary complex (Seghedi et al. 1999); upper Ordovician-Devonian anoxic distal turbidites and radiolarian cherts occupy a younger accretionary prism near Tulcea (Seghedi 1998); Silurian distal shelf strata pass up conformably into Lower Devonian shelf clastic and carbonate strata. The northern part (North Dobrogea Orogen) of this terrane was strongly affected by the Variscan Orogeny, with NE-directed thrusting and granite intrusion (S~ndulescu 1984; Seghedi & Oaie 1995; Liszkowski et al. 1998; Seghedi 1998) although the orogenic front is poorly located (Banks 1997). Monazite ages suggest a late Carboniferous-early Permian age for amphibolite-facies metamorphism (Seghedi et al. 2003). The basement in Central Dobrogea comprises Neoproterozoic metaturbidites affected by 'Cadomian'-age folding (Kr~iutner et al. 1988, and references therein; Seghedi & Oaie 1995), and is regarded by many researchers as a terrane separate from North Dobrogea. The evolution of the pre-Variscan basement of the Tulcea Terrane shows closest parallels to that of the SaxoThuringian (southern margin) of the Rheic Ocean, described in an earlier section. Unconformably overlying upper Palaeozoic strata (Carapelit Fro) have significant volcaniclastic input from a calc-alkaline magmatic arc. In Permian-early Mesozoic time the terrane was dislocated, by rifting along the peri-Tethyan margin (aided by strike-slip along possible correlatives of the TeisseyreTornquist Zone, e.g. the Peceneaga-Camena and Sfante Georghe Faults, crust-penetrating structures associated with offset of the Moho), from the remainder of Variscan Europe (and the TESZ) during break-up of the Pangaea supercontinent (Ziegler 1990). (2) East M o e s i a n Terrane. The Moesian Platform (MesozoicCenozoic) extends SW from the Capidava-Ovidiu Fault in South Dobrogea towards the Carpathian Foreland in Bulgaria (Visarion et al. 1988), but the pre-Mesozoic terrane boundary is the concealed Palazu Overthrust. The Intra-Moesian Fault is the boundary with the West Moesian Terrane (see below). To the north of the latter, in what is referred to here as the East Moesian Terrane, a higher-grade basement of Archaean gneisses and Palaeoproterozoic banded iron formation, similar to that of the Ukrainian shield, is overlain by a low-grade Neoproterozoic volcano-sedimentary succession (Seghedi 1998), comparable with the Volhyn volcanic units. These poorly dated units are overlain by Cambro-Ordovician siliciclastic strata (Iordan & Spassov 1989). The lithostratigraphic similarities to the EEC of Poland and the Ukraine are therefore strong. Claims that mid-Cambrian trilobites show affinities with England, Bohemia and the EEC (Iordan 1999), or support a Baltican affinity (Rushton & McKerrow 2000) cannot at present confirm the faunal provinciality of East Moesia. These rocks are unconformably overlain (following a Llandovery hiatus) by largely pelitic upper Silurian-Lower Devonian strata of North Gondwanan (i.e. Armorican Terrane Assemblage; ATA) affinity (Vaida et al. 2005). Midddle Devonian coarse clastic strata are overlain by upper Devonian-lower Carboniferous carbonate platform strata thickening into a foredeep north of the Craiova High. Westphalian-Stephanian coal measures are unconformably overlain by Permian strata (Banks 1997). Thus it is possible that the East Moesian Terrane may have been displaced from the EEC in early Ordovician time and, following transcurrent displacement and intercalation within the ATA, reincorporated to the EEC in post-Variscan time. (3) West M o e s i a n Terrane. That part of the Moesian Platform west of the Intra-Moesian Fault is referred to here as the West Moesian Terrane (Fig. 1). The Precambrian basement here is overlain by a Palaeozoic sequence up to 6.5 km thick, and is poorly
WESTERNACCRETIONARYMARGINOF THE EEC known and undated. The basement reworked into the internal massifs of the adjacent southern Carpathian Mountains may provide clues to its affinity. The structurally lowest, supposedly parautochthonous, Danubian nappe complexes contain a metaplutonic and metasedimentary basement of Neoproterozoic age and 'Cadomian' affinity, considered by some to represent the exhumed basement (Sandulescu 1994) of the West Moesian Terrane. Arguments have been presented for both Armorican (the traditional view) and Avalonian provenance (Winchester et al. 2006) for this crust. In contrast, the higher, allochthonous Getic Nappes apparently show more certain ATA affinity (Iancu et al. 2006). It is apparent that the Moesian Terrane (as defined by Haydutov & Yanev 1997) had a complex Palaeozoic history, and is probably composite. The larger part of the Carpathian Foreland in Romania (Tulcea and West Moesian terranes) has a basement showing affinities with Gondwana and the ATA, and was probably accreted to the TESZ during the Variscan Orogeny. Tectonically interleaved is a rather narrow wedge of crust about 100 km wide, the East Moesian Terrane, apparently originating in the EEC (Cambrian), but showing increasingly Gondwanide affinity through Ordovician and Silurian time, which was probably reamalgamated with the EEC during or soon after the Variscan Orogeny. The affinities of West Moesia are even more controversial. Subsequent opening of the proto-Pannonian marginal basin in late Triassicearly Jurassic time (Banks & Robinson 1997) may have caused further dispersal of the newly amalgamated Moesian composite terrane along the EEC margin, as did mid-Cretaceous opening of the Black Sea Basin. Correlation of the terranes of the Moesian Platform with the Moravo-Silesian Terrane, which is in a similar structural position with regard to the EEC margin, has been proposed by Burchfiel (1975) and Matte et al. (1990), but is here considered unlikely because of the lithostratigraphic contrasts described above.
Peri-Tethyan dispersal of TESZ terranes The Variscide basement of the Western and Central Pontides in Turkey (Okay et al. 1994), was rifted away from the Moesian Platform in mid-Cretaceous time, following oblique slip along the TTZ, to form the Western Black Sea Basin. It is therefore clear that the TTZ crustal discontinuity also affected post-Variscan tectonics very significantly (C. Tomek, pers. comm.), during Mesozoic extension and Cimmerian and Alpine inversion. Further consideration of the evolution of terranes first accreted to, and then lost from, the Western Accretionary Margin of the EEC (e.g. the Zonguldak Terrane (?Avalonian affinity) and Sakarya and Eastern Pontide Blocks (?Armorican)), has been given by Winchester et al. (2006). The Alpine and Carpathian orogens (Fig. 1) contain numerous internal massifs (e.g. the Tatra Mountains) comprising crust reworked from the TransEuropean Suture Zone following the end of the Variscan Orogeny. Some were displaced from the margin of the Palaeotethys and Neotethys oceans (Stampfli et al. 2001), prior to Alpine thrust displacement. The peri-Gondwanan and Gondwanan affinities of such massifs have been demonstrated by detailed U-Pb zircon dating studies (e.g. Schaltegger et al. 1997; Von Raumer et al. 1999), but a detailed description of them is largely beyond the scope of this review. Further details may be found in the sections of this volume dealing with the younger orogens.
Orogenic sutures within the Western Accretionary Margin of the EEC The most significant terrane boundaries (Fig. 1) are sutures associated with destruction of oceanic lithosphere: the Iapetus Suture, separating Avalonia from the Laurentian terranes (beyond the
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scope of this review); the Thor Suture (Berthelsen 1998) separating Baltica and Avalonia, marking closure of the Ordovician Tornquist Sea; the R h e i c Suture (Cocks & Fortey 1982), separating the early Palaeozoic accreted terranes of Laurussia from an 'archipelago' of Gondwana-derived terranes accreted in late Palaeozoic time. The latter terranes were separated by seaways and larger ocean basins, such as the Saxo-Thuringian and Massif Central oceans (Matte et al. 1990), the closure of each being marked by the suture zones that are now summarized.
Iapetus Suture
The Iapetus Ocean (Harland & Gayer 1972) was initiated during the break-up of the Rodinia-Pannotia Supercontinent (Bond et al. 1984; Dalziel 1997) in late Neoproterozoic time. This event was preceded by the intrusion of tholeiitic dykes throughout the Scandinavian margin of Baltica (Andrrasson 1998) and the conjugate Laurentian margin in Newfoundland, Scotland (Tayvallich) and Ireland at c. 600 Ma. The 'Sparagmite Group' in Scandinavia (Kumpulainen & Nystuen 1985) and the volcanic-rich Volhyn strata of Ukraine and Poland (Moczydlowska 1997) are manifestations of this same phase of rifting. The Tornquist Sea (Cocks & Fortey 1982) is generally regarded as a segment of Iapetus separating Avalonia from Baltica only in Ordovician time (see below). The final closure of the Iapetus Ocean occurred in Silurian time (McKerrow et al. 1990). The location of the Iapetus Suture is well constrained within Britain and Ireland by deep seismic data (Freeman et al. 1988; Soper et al. 1992) and extends into the western part of the North Sea (Fig. 1). However, the thermo-mechanical effects of extension in the North Sea graben have so modified lower crustal reflectivity that it is difficult to identify the location of the suture farther east. Fichler & Hospers (1990) inferred a position on the East Shetland Platform west of the graben, but a location closer to the Norwegian coast (Fig. 1) is equally feasible (Pharaoh 1999).
T h o r Suture
This suture marks the closure of the Tornquist Sea separating Baltica and Avalonia in Ordovician time, and was originally recognized using faunal provinciality criteria (Cocks & Fortey 1982). It closed slightly earlier than the rest of the Iapetus Ocean, probably in late Ordovician time (Picketing 1989; Pharaoh et al. 1995). The apparent absence of any subductionrelated magmatism on the Baltica margin (other than bentonites representing ash-fall deposits) argues for subduction towards the SW. The near-surface location of the suture is constrained by deep boreholes on the Ringkcbing-Fyn High of Denmark and NE Germany, as described earlier (Fig. 1). A set of rather weak, SW-dipping, mid-crustal reflections observed on seismic profiles in the southern Baltic Sea has been correlated with the suture (BABEL Working Group 1993; Tanner & Meissner 1996; DEKORP-BASIN Research Group 1998, 1999). Similar features have also been identified west of Denmark (MONA LISA Working Group 1997a,b) where they maintain a rather constant dip (c. 10-12 ~ into the lower crust. The concept of the suture as a steep 'Trans-European Fault' (Berthelsen 1992b, 1993) is contradicted by evidence from normal-incidence and wide-angle seismic profiles on the North German Plain (EUGENO-S Working Group 1988; Krawczyk et al. 2002), which indicate the presence of a wedge of high-velocity, EEC-type basement in the lower crust as far south as Hamburg (Thybo et al. 1990; Aichroth et al. 1992; Rabbel et al. 1995). The suture is presumed to continue from the vicinity of Rtigen Island into northern Poland (Winchester et al. 2002) towards the Moravian Line marking the western limit of the Matopolska and Moravo-Silesian (=Brunovistulian) Terranes, which had Baltican affinity by mid-Cambrian time at the
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T.C. PHARAOHETAL.
latest. The Sorgenfrei-Tornquist Zone (STZ) has a long history of reactivation (Berthelsen et al. 1992a; Thybo 1997) and is associated with a Moho offset (EUGENO-S Working Group 1988; Kind et al. 1997; Cotte & Pedersen 2002) but does not appear to have a simple genetic link to the Thor Suture. Instead, it may represent a younger 'orogenic back-stop' or 'boundary detachment' defining the northern limit of Variscan crustal deformation (Berthelsen 1998), subsequently reactivated in the Alpine Orogeny. The Teisseyre-Tornquist Zone (TTZ) in Poland (Guterch et al. 1986; Krdlikowki & Petecki 1997) may have originated as part of the passive margin architecture of the EEC as early as the Neoproterozoic, and acted as a transcurrent fault zone for much of Palaeozoic time. It is unlikely to represent an oceanic suture, however (Pharaoh 1999).
2000) at the northern edge of the Bohemia Terrane (=Tepl~Barrandian Unit). Gf6hl Suture
To the south of the Tepl~-Barrandian Unit, which separates NWand SE-verging parts of the Variscides (Matte et al. 1990), a suture within the Moldanubian Zone dips to the north (see Fig. 8), with an opposite sense of vergence (and, presumably, subduction polarity) to that of the Saxo-Thuringian Suture. It is associated with accreted oceanic protoliths (e.g. pods of mantle peridotite and eclogite) in the Gf6hl Unit of the Bohemian Massif. It may represent a continuation of the inferred M a s s i f Central Suture (Matte et al. 1990).
R h e i c Suture
M o r a v i a n Suture
The Rheic Suture is a complex structure, and is almost certainly compound in nature (W. Franke, pers. comm.). The early history of the ocean is poorly known, but it may have begun to open in early Ordovician time with the rifting of Avalonia and the earliest terranes of the Armorican archipelago from Gondwana (Cocks & Fortey 1982; Pharaoh 1999). The Lizard Peridotite and Giessen Ophiolite appear to represent the only likely relics of Rheic ocean floor. Plagiogranite in the former crystallized at 397 _+ 2 Ma (Clark et al. 1998) in early Devonian time. This age is identical within error to a white mica age (Ar-Ar laser microprobe on Acadian cleavage pressure fringes) of 396.1 __ 1.4 Ma obtained by Sherlock et al. (2003). The ophiolites may represent relics of original marginal-basin crust associated with northward subduction at the northern margin of the ocean (Ziegler 1990), apparently contemporaneous with Acadian deformation farther north in Avalonia. If this subduction was at a low angle, it might explain the transfer of compressional stress deep into the interior of Avalonia to produce Acadian deformation, as well as the apparent absence of subduction-related magmatism. Obduction of the Lizard complex, associated with overthrusting by the Normannian Complex, began at about 380 Ma (Clark et al. 1998). Subsequently, in late Devonian to early Carboniferous time, a smaller Rheno-Hercynian Ocean Basin may have opened not quite coincident with the original Rheic Suture (Franke 2000). The Mid-German Crystalline High (Dallmeyer et al. 1995; Franke 2000) is a magmatic arc produced by later southward subduction of this ocean, and it is this later phase of suture development that is imaged by deep seismic reflection profiling. Profiles in the English Channel (Leveridge et al. 1984) show that the suture maintains a constant southward dip of c. 20 ~ into the lower crust and suggest that the Variscan orogen is distinctly thick-skinned in aspect. The DEKORP-2 profile (see Fig. 8) provides a complete transect across the Variscides in Germany. The Rheic Suture Zone is imaged as several SE-dipping reflector zones in the mid-crust, interpreted as thrusts (Meissner & Bortfeld 1990). In North Dobrogea in Romania, terranes showing affinity with the Armorican Terrane Assemblage, including Variscan-age amphibolite-facies metamorphism (Seghedi et al. 2003) directly abut the EEC, and the Rheic Suture probably lies close to the Galati-St. George Fault. Its original geometry has, however, been severely modified by post-Variscan events in this region.
On the eastern flank of the Bohemian Massif, the Gf6hl Unit is in tectonic contact with the Bruno-Silesian Terrane, along a ductile shear zone containing ophiolitic fragments (Schulmann et al. 1991) long referred to as the 'Moldanubian Thrust'. Structures along this line show highly oblique (dextral sense of shear) overthrusting to the east in early Carboniferous times (Schulmann & Gayer 2000). This geometry is inferred to reflect the northward convergence of the Armorican Terrane Assemblage with Laurussia, obliquely against the orogenic promontory on the EEC formed by the former Bruno-Silesian Terrane (Banka et al. 2002; Winchester & PACE TMR Network 2002). However, the ophiolitic fragments are of uncertain age and might be derived from older Giessen, Mfinchberg or Gf6hl-type protoliths. An alternative view (Franke 2006) is therefore that the inclusion of the ophiolitic fragments is fortuitous, and the shear zone does not represent a suture. Furthermore, the Moldanubian Thrust truncates three different terranes along its western edge, and studies of the flysch provenance suggest that it is of late tectonic age. The extension of the 'Moravian Line' toward the NE has also been invoked as the eastern boundary of 'Far Eastern Avalonia' (Winchester & PACE TMR Network 2002). All of the sutures within the ATA, described above, reflect closure of (perhaps small) ocean basins that originally separated the Gondwana-derived elements of the Armorican Archipelago (Franke 2000). Faunal evidence (McKerrow et al. 2000) indicates that no large (i.e. > 1 0 0 0 k m wide) oceans existed in late Palaeozoic Europe.
S a x o - T h u r i n g i a n Suture
This SE-dipping suture forms the root zone of the Mfinchberg Nappe, and is well imaged by seismic reflection data, both by the DEKORP-4 profile at the western margin of the Bohemian Massif (Fig. 1) and the DEKORP-2 profile, which is incorporated in Crustal Transect 5 (see Fig. 8). It is marked by early Palaeozoic mid-ocean ridge basalts (Dallmeyer et al. 1995; Linnemann et al.
Crustal transects through the Western Accretionary Margin of the EEC The crustal structure of the Western Accretionary Margin of the EEC is illustrated by a number of transects, shown in Figures 4 - 8 . These attempt to build on the cross-sections presented in the European Geotraverse (Blundell et al. 1992) by incorporating information from geophysical experiments carried out in the past 15 years. Transect 1 (Fig. 4) extends some 1200 km from the English Midlands to the Stockholm archipelago. The crustal structure is entirely concealed in this region by Carboniferous and younger strata of the Southern North Sea Basin, thickest in the Central Graben, Horn Graben and Norwegian-Danish Basin. The core of the transect is based on interpretations of deep seismic reflection profile SNST 83-07 and refraction profile EUGENO-S2, tested by modelling of the gravity and magnetic potential fields by Williamson et al. (2002). The most significant crustal boundary in this transect is the Thor Suture between crust of Avalonian affinity (in the SW) and crust of Baltican affinity to the NE. A SW-inclined, poorly reflective zone recognized on the MONA
WESTERN ACCRETIONARYMARGIN OF THE EEC
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Fig. 4. Transect 1, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC in the central North Sea-southern Scandinaviaregion; 5 x vertical exaggeration. Based on an interpretation of deep seismic reflectionprofile SNST 83-07 (Klemperer& Hobbs 1991) and refraction profile EUGENO-S2 (EUGENO-S Working Group 1988). MR, mantle reflectors shown schematicallyfollowingBlundell et al. (1991). Slightly modifiedafter Williamsonet al. (2002). Post-Palaeozoic basins and platforms: CG, Central Graben; HG, Horn Graben; NDB, Norwegian-Danish Basin; RFH, Ringkcbing-Fyn High; RG, R0nne Graben. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: DSHFZ, Dowsing-South Hewett Fault Zone; EEC, East European Craton; ?SNSLT, Inferred Southern North Sea-Luneberg Terrane; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: ABDB, Anglo-BrabantDeformation Belt (=Eastern EnglishCaledonides); 'CDF', front of Caledoniandeformation (see text for explanation);EEC, East European Craton; MMC, MidlandsMicrocraton; SFD, Svecofennian Domain; SND, SveconorwegianDomain; SNF, SveconorwegianFront; S-TZ, Sorgenfrei-TornquistZone; TSB, Trans-ScandinaviazlBatholith;VA, concealed volcanic arc inferredfrom magnetic signaturein Southern North Sea (Williamsonet al. 2002). For mantle reflectors: key boreholes (in red): G1, Glinton;Gr, Grinsted;NC, North Creake; No, Novling; Ro, R0nne; WF, WithycombeFarm.
LISA deep seismic profiles across the RingkCbing-Fyn High has been attributed to the suture (MONA LISA Working Group 1997a). This hypothesis was supported by the modelling work of Williamson et al. (2002), which used the fundamental contrast in magnetic susceptibility of these two types of crust to map the location and geometry of the suture. This analysis indicated that the suture dipped SW at an average angle of about 14 ~ being steeper in the upper crust (up to 40 ~) than in the lower crust (average 7~ This interpretation also invoked tectonic imbrication of platform cover and crust in the footwall of the suture, with strong deformation dying out eastwards towards a 'Caledonian Deformation Front' some 100 km east of the suture. Within the crust of the EEC, strong magnetic contrasts were not found across the Sorgenfrei-Tornquist Zone, suggesting that it does not bound crust of fundamentally different type. MONA LISA Working Group (1997a) also reported inclined zones of reflectivity in the mantle beneath the crustal suture, with both SW and NE dips. Pharaoh (1999) interpreted the SW-dipping set as being related to Avalonia-Baltica collision, invoking lithospheric delamination at or close to the Moho to explain the observed offset of about 200 km (Fig. 4). The crust of Avalonia is laterally heterogeneous. In the SW, crust of the Midlands Microcraton, known from exposures and deep boreholes, virtually unaffected by Caledonian deformation, passes towards the NE into the AngloBrabant Deformation Belt (----Eastern English Caledonides), which experienced much stronger Caledonian deformation. In the west, SW-vergent thrusting in the shallow crust involves the Precambrian basement, and may be of Acadian age. In the mid-crust, SW-dipping zones of inclined reflectivity described by Reston (1990) and Blundell (1993) may have developed during the earlier, Shelveian, 'soft' collision between Avalonia and Baltica (Pharaoh et al. 1995). The Dowsing-South Hewett Fault Zone and Lower Rhine Lineament (Fig. 1) may have been initiated at this time, as a suspect suture between Eastern Avalonia s e n s u s t r i c t o and the poorly known crust of the Southern North
Sea region (Far Eastern Avalonia), for which separate terrane status has been invoked (Franke 1995; Pharaoh et al. 1995; Pharaoh 1999). Williamson et al. (2002) invoked the presence of a concealed volcanic arc within this inferred terrane to explain the linear magnetic anomaly lying along the southern flank of the Mid-North Sea High and R y n k c b i n g - F y n High, clearly seen in Figure 2, possibly the 'lost arc' generated by subduction of the Tornquist Sea. Transect 2 (Fig. 5) runs SSW from the Harz Mountains in Germany, across the North German Plain towards the Baltic Sea in the NNE. The interpretation presented is slightly modified from that presented for the deep seismic reflection profiles DEKORP-BASIN 9601 (onshore) and DSB-9 (offshore extension) by Lassen et al. (2001) and Lassen (2005). In this region, the crystalline basement is almost completely concealed by the thick Permian-Cenozoic fill of the North German Basin. The most important crustal structure is the Thor Suture between Avalonia in the SW and Baltica in the NE. The DEKORP profile clearly images the high-velocity crystalline basement of the EEC extending for 165km as a SW-tapering wedge beneath the North German Basin (Bayer et al. 2002; Lassen 2005) to about the location of Hamburg. The average dip of the suture is about 10 ~ The tip-line of this wedge (at the Moho level) is apparently juxtaposed with the postulated Elbe Lineament, inferred to separate crustal regions with distinct basement tectonic grains and geophysical attributes, referred to as the Pompeckj and Holstein blocks (Aichroth et al. 1992; Rabbel et al. 1995), and a possible suspect terrane boundary (e.g. Franke 1995b; Rabbel et al. 1995; Tanner & Meissner 1996). In the interpretation presented here, the Elbe Lineament does not represent the SW boundary of the EEC, which lies at the base of the crust here. The lineament may represent a postcollisional structure in the upper crust focused at the boundary between contrasting lithospheric types; or it may represent a boundary between contrasting types of upper crust (e.g. Variscan
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T.C. PHARAOHET AL.
Key to Figs 4-8.
and Caledonian deformed crust), in which case the geographical coincidence with the lower crustal wedge of Baltica is fortuitous. Various 'interwedging' structures in the lithosphere of this region are attributed to Caledonian compressional tectonic phases (Meissner et al. 2002). Anorthosite xenoliths of shield type entrained in Permian lavas encountered by the Schwerin-1 Borehole (K/impf et al. 1994; Breitkreuz & Kennedy 1999), and isotopic data from the Loissin Borehole (Dallmeyer et al. 1999), support this geophysical interpretation. No attempt is made to depict the crustal structure of the EEC in this profile, which is poorly constrained by seismic refraction data. Other features of the EEC shown here are the presence of inferred half-grabens infilled by Ediacaran strata (Lassen 2005), probably generated during the rifting and break-up of Rodinia-Pannotia; and Cambro-Ordovician strata of the Baltica rifted passive margin
(Scheck et al. 2002) thinning eastward. The latter are known from deep boreholes to the north of Rtigen Island (Katzung et al. 1993; Vejbaek 1997). The carbonaceous Alum Shale is a likely tectonic detachment horizon, forming, as in the Scandinavian Caledonides, an important geophysical marker: a structural boundary ('O-Horizon') on seismic reflection profiles; and a zone of high conductivity on resistivity profiles (ERCEUGT Group 1992). On Rfigen Island, several boreholes penetrate deep-water strata of Ordovician age, which have been assigned an Avalonian affinity (Verniers et al. 2002). These, together with boreholes further east in Pomerania, and to the west in Heligoland (Frost et al. 1981), represent rare provings of the crust of the suspect terrane of 'Far Eastern Avalonia'. Little detail is shown within the largely concealed Caledoniandeformed crust of Avalonia presumed to underlie much of the
WESTERN ACCRETIONARYMARGIN OF THE EEC
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Fig. 5. Transect 2, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC beneath the North German Basin; 5 x vertical exaggeration. Modified after interpretations of seismic reflectionprofiles DEKORP BASIN 9601 (onshore) and DSB-9 (offshore) by Lassen et al. (2001) and Lassen (2005). Post-Palaeozoic basins, platforms and geographical features: ADF, AlpineDeformationFront; GT, GardelegenThrust; NGB, North German Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: EL, Elbe Lineament;TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: CDF, front of Caledoniandeformation(see text for explanation);EEC, East European Craton; HM, Harz Mountains; SHT, Sub-HercynianTrough; VF, VariscanFront. Key boreholes (in red): G14, G14; Lo, Loissin; Pz, Pritzwalk; Ru, Rtigen 5, Other geographical locations mentioned in text: BS, Baltic Sea; ER, Elbe River.
North German Plain. Towards the south, several south-dipping reflector sets have been correlated with Variscan thrusts within the Rheno-Hercynian Zone. Some of these structures (e.g. the Gardelegen Thrust) were reactivated during Alpine compression. Transect 3 (Fig. 6) is an 800 km long profile extending from the Elbe Fault Zone (not to be confused with the Elbe Lineament, defined above) in the SW across the central part of the Polish Trough to the Mazurska High of the EEC in the NE. This interpretation is based on the results from the POLONAISE P4 deep seismic refraction profile published by Grad et al. (2002a,b). The crust of the EEC exhibits characteristic three-layered structure, with average P-wave velocities of about 7.1 km s-1 (lower crust), 6.5 km s -1 (middle crust) and 6.2 km s -t (upper crust). The Suwalki anorthosite complex forms a localized high-velocity (6.4 km s - i ) anomaly within the upper crust. At up to 50 km thick, the crust of the EEC is much thicker than that of young, accreted Europe. The velocity model presented by Grad et al. (2002a) suggests that at the Teisseyre-Tornquist Zone, the high-velocity upper and middle crust terminates rapidly against the Polish Trough. This abrupt termination is also clear in the magnetic potential field image (Fig. 2). The lower crust appears to extend as an attenuated wedge for perhaps 250 km to the SW of the Teisseyre-Tornquist Zone, however (Fig. 6), beyond the Polish Trough, to underlie a less heterogeneous crust with much lower average velocity assigned to Avalonia (Winchester & PAGE TMR Network 2002). It is clear that in this region, the geometry of the Thor Suture has been considerably modified by the effects of Permian-Mesozoic extension in the Polish Trough. The midcrust beneath the Polish Trough has a P-wave velocity of about 5.8 km s- 1. This is too low for crystalline basement, and therefore may represent deeply buried Neoproterozoic-Palaeozoic strata. These could be of Devonian and Carboniferous age, as in other rifts within the EEC; but more probably comprise thick Ediacaran (Vendian) and thinner lower Palaeozoic strata comparable with those of the Lysogory and Matopolska Blocks, exposed in the Holy Cross Mountains some 250 km to the SE of the transect.
They could be filling a 'Central Polish Palaeo-Rift' (see above, for discussion) antecedent to the (Permian-Cenozoic age) Polish Trough, and initiated during the rifting of Rodinia-Pannotia. The eastern bounding fault of the Polish Trough may represent a simple reactivation of an early Palaeozoic syndepositional fault, perhaps associated with the passive margin development of Baltica; or, if the interpretation of these blocks as suspect terranes is correct (see earlier discussion), a reactivated terrane boundary (between the Lysogory Terrane and Baltica). It was a locus of Alpine inversion. The western boundary of the Polish Trough also appears sharp, but not quite as steep as the eastern boundary. It coincides with the northward extension of the Moravian Line, which Winchester et al. (2002) have identified as the boundary at the eastern extremity of (Far Eastern) Avalonia. The crust here is comparatively homogeneous and unlayered, exhibiting a P-wave velocity of about 6.3 km s- 1 to 25 klTIdepth, characteristics that support an Avalonian affinity. At shallow crustal level, deep boreholes on the Leszno-Wolsztyn High sample metamorphic rocks of the Rheno-Hercynian Zone, derived from (Franke & Zelainiewicz 2002), and resting upon, the Avalonian crust. Transect 4 (Fig. 7) extends for 240 km across the Mid-North Sea High off the NE coast of Britain (Fig. 1). It crosses the various accreted terranes bordering the Iapetus Suture between Avalonia and Laurentia. The interpretation presented is adapted from those presented by Freeman et al. (1988) and Chadwick & Holliday (1991). The location of the Iapetus Suture is well constrained in the western part of the North Sea by deep seismic data (Freeman et al. 1988; Soper et al. 1992). The northward dip of reflector packages (IN, IS) in the lower crust and in the mantle (P) at and just below the Moho (Fig. 7) indicates a northward-dipping suture, with a wedge of Avalonia (probably comprising juvenile accreted material, such as the Skiddaw Group) extending some 7 0 k i n beneath the largely Silurian age Southern Upland Acccretionary Complex. The lower crust of the latter is distinguished not by sonic velocity, but by its reflection character. In the north, the Southern Upland Fault is
304
T.C. PHARAOHETAL.
Fig. 6. Transect 3, to illustrate inferred crustal structure of the Western AccretionaryMargin of the EEC in central Poland; 5 x verticalexaggeration. Interpretation based on crustal model of POLONAISE P4 deep seismic refractionprofilepublished by Grad et al. (2002a,b) and Sroda et al. (2002). Post-Palaeozoic basins and platforms: CPT, Central Polish Trough; L-W, Leszno-Wolsztyn Basement High. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: DF, Dolsk Fault; EEC, East European Craton; KLZ, Krakdw-Lubliniec Zone; LT, LysogoryTerrane; ML, MoravianLine; OF, Odra Fault; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: CPPR, Central Polish Palaeo-Rift;MH, Mazurska High; ML, MoravianLine; SU, Suwalki AnorthositeMassif; T-TZ, Teisseyre-TornquistZone; VF, Variscan Front. the terrane boundary with the Midland Valley Terrane, another component of the Laurentian terrane collage. The development of Carboniferous sedimentary basins overlying the accretionary complex (e.g. the Northumberland Trough and Tweed Basin)
Fig. 7. Transect 4, to illustrate inferred crustal structure across the 'Caledonide' accreted crust in the central North Sea region; 2.5 x vertical exaggeration. Modifiedfrom interpretationsof NEC deep seismicreflectionline by Freeman et al. (1988) and Chadwick & Holliday (1991). Post-Palaeozoic basins and platforms: MNSH, Mid-NorthSea High; NFF, Ninety Fathom Fault; NT, NorthumberlandTrough; TB, Tweed Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: IN, IS; Iapetus Suture Zone; MVT, Midland Valley Terane; SNSLT, SouthernNorth Sea Terrane. Mantle reflectors: P.
was facilitated by extensional reactivation of the early Palaeozoic structures (Chadwick & Holliday 1991). Transect 5 (Fig. 8) extends from Kiel, southward across the North German Basin (Variscan Foreland), crossing the various internal zones of the German Variscides towards the Alpine Molasse Basin near Zurich. The interpretation of the crustal structure is based on that published for the central segment of the European Geotraverse (EGT) by Aichroth & Prodehl (1990) and Prodehl & Aichroth (1992). The detail of shallow crustal structure in the central part of the transect is derived from interpretations of the DEKORP-2N and -2S deep seismic reflection profiles, which are slightly divergent from the path of the EGT (Fig. 1), published by Giese (1995) and Oncken et al. (2000). As on Transect 2, a tapering wedge of high-velocity EEC lower crust is depicted extending southward beneath (Far Eastern) Avalonian crust underlying the North German Plain, to the vicinity of the Elbe River. Crossing the Variscan Front near Celle, the DEKORP-2N profile (Meissner & Bortfeld 1990; Meissner et al. 1994) provides an excellent view of the internal structure of the Rheno-Hercynian nappe pile. The profile images a ramp-flat geometry for the basal detachment, unreflective crust of the Brabant Massif forming a south-tapering wedge in the footwall, and a Moho that brightens southward. Despite the considerable shortening and displacement implied, the early Palaeozoic faunal affinities of the RhenoHercynian Zone lie with Avalonia (Cocks et al. 1997) so that the Variscan Front and basal detachment represent an entirely intra-Avalonian boundary. The section balancing reconstruction of Oncken et al. (2000) indicates at least 200 km of horizontal shortening in the upper crust in this zone of the Variscides. The lower crust has a much higher P-wave velocity of about 6.8 km s -1. In Figure 8, following the known inherited ages of many Neoproterozoic Avalonian granitoids, this is speculatively interpreted as an extensive block of Rondonian (Mid-Proterozoic)
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Fig. 8. Transect 5, to illustrate inferred crustal structure of the 'Variscide' accreted crust in the central German region; 5 • vertical exaggeration. Interpretation of full crustal structure from the EGT deep seismic refraction experiment (Aichroth & Prodehl 1990; Prodehl & Aichroth 1992). Detail of shallow crustal structure in central part of transect is after interpretations of DEKORP deep seismic reflection profiles by Giese (1995) and Oncken et al. (2000). Post-Palaeozoic basins and plaOeorms: HD, Hessen Depression; MB, Molasse Basin; NGB, North German Basin. Postulated Palaeozoic terranes and possible terrane or sub-terrane boundaries: MGCH, Mid-German Crystalline High; NPZ, Northern Phyllite Zone; PT, Perunica (Bohemia) Terrane; MGS, Moldanubian-Gf6hl Suture; MT, Moldanubia Terrane; RS, Rheic Suture; STS, Saxo-Thuringian Suture; TS, Thor Suture. Proterozoic-Palaeozoic tectonic elements: BM, Bohemian Massif; DR, Drosendorf Unit (of BM); ET, Ebbe Thrust; GF, Gf6hl Unit (of BM); GN, Giessen Nappe; MN, Mtinchberg Nappe (of BM); SG, Saxonian Granulites; TBZ, Tepl~-Barrandian Basin (of BM); TT, Taunus Thrust. Other geographical locations mentioned in text: Ki, Kiel; ER, Elbe River; SJ, Swabian Jura; Zu, Zurich.
crust lying at the heart of Avalonia. The transect then crosses the Saxo-Thuringian Zone, passing just east of the Odenwald. The Rheic Suture, defining the southern edge of pre-Variscan Avalonia, appears to be a rather steeply dipping feature in the upper crust (perhaps 70~ but is less steeply dipping in the lower crust, where a wedge of Rheno-Hercynian (=Avalonian protolith) apparently extends beneath the Saxo-Thuringian domain (Giese 1995; Krawczyk et al. 2002). To the south of the Mid-German Crystalline High, the Saxonian Granulite Dome forms a key, but enigmatic, element of the crust (Krawczyk et al. 2000). The Mtinchberg Nappe, resting on the Saxo-Thuringian Suture, is interpreted as an outlier of the Moldanubian Zone, which forms the southern part of the transect. The Moldanubian Gfrhl Suture separates correlatives of the Drosendorf and Gfrhl units of the Bohemian Massif, amongst the latest of the TESZ terranes to leave Gondwana before the opening of the Palaeo-Tethys Ocean.
Conclusions Multidisciplinary studies by the EUROPROBE Programme, including reinterpretation of older geophysical datasets, have led to considerable improvements in knowledge of the structure and evolution of the Trans-European Suture Zone (TESZ), the most fundamental lithospheric boundary in Europe. The TESZ represents a tectonically complex zone of crustal terranes accreted, throughout Palaeozoic time, to the passive margin of the East European Craton (EEC). Various geophysical techniques have been applied to define the geometry of the sutures, representing destroyed ocean basins, which define the boundaries of these various terranes, and most are shown to be non-vertical. Other well-known structures and lineaments (e.g. the SorgenfreiTornquist Zone (STZ) and Teisseyre-Tornquist Zone (TTZ), are steep structures unrelated to suturing, but may have originated during late Neoproterozoic rifting of the Rodinia-Pannotia supercontinent. They have subsequently been reactivated many
times, most recently during Alpine inversion. Standard methods of terrane analysis have been applied to identify and characterize individual terranes within the TESZ, but the status of some remains controversial. Fortunately, the application of more sophisticated studies of isotopic composition (e.g. of detrital grain provenance), palaeontology (e.g. to recognize timing of ocean closure) and palaeomagnetism (e.g. to identify terrane rotation) is helping to resolve some of these controversies, and add detail to the history of accretion and dispersal. These studies have demonstrated that numerous crustal terranes were rifted away from various margins of Gondwana at low southerly palaeolatitude more or less continuously, for much of early Palaeozoic time. After a northward passage, driven by the opening and closure of numerous (perhaps not very large) ocean basins, at least some of these ended up being accreted to the passive margin of the EEC. As studies proceed, the evidence for the rates of rotation and closure amidst this archipelago of Gondwanaderived terranes is starting to contribute to a dynamic model for the evolution of the TESZ. Subsequent to initial accretion, dispersal of some terranes to other locations on the EEC margin was facilitated by major, crust-penetrating steep faults such as the TTZ. This process is most advanced in the SE part of the TESZ in Romania and Turkey, where terranes originally accreted during the Variscan Orogeny have been displaced and re-accreted to the EEC margin during the development of the peri-Tethyan margin, opening of the Black Sea and the subsequent AlpineCarpathian Orogeny. Some of the studies reported here were carried out in the EU-funded PACE (Palaeozoic Amalgamation of Central Europe) TMR Network, no. ERBFMRXCT97-0136. The contribution of T.P. is published with the permission of the Executive Director, British Geological Survey (NERC). Numerous fruitful discussions with EUROPROBE TESZ and PACE project participants are gratefully acknowledged. PACE colleagues P. Williamson and D. Banka contributed significantlyto the quality of the diagrams. J. Carney and W. Franke are thanked for their helpful reviews, and D. Gee for editorial comments. D. Gee is sincerely thanked for his significant contribution to the TESZ project, and for the unstoppable motivation he provided during the 10 EUROPROBE years, 1992-2002.
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Physical differences in the deep lithosphere of Northern and Central Europe SOREN G R E G E R S E N l, PETER VOSS 1'2, Z. H O S S E I N S H O M A L I 3'4, M A R E K G R A D 5, R O L A N D G. ROBERTS 3 & TOR W O R K I N G GROUP
IGEUS, Ostervoldgade 10, DK-1350 Copenhagen K, Denmark (e-mail:
[email protected]) 2Department of Geophysics, Niels Bohr Institute, University of Copenhagen, DK-2100 Copenhagen O, Denmark 3Department of Earth Sciences, Uppsala University, SE-752 36 Uppsala, Sweden 4present address: Geological Survey of Sweden, SE-751 28 Uppsala, Sweden 5Institute of Geophysics, University of Warsaw, Pasteura 7, PL-02-094 Warsaw, Poland
A number of large-scale integrated studies, including the TOR and POLONAISE'97 projects, with an emphasis on seismic methods, have been used to elucidate the southwestern boundary (suture zone) between the East European Craton and the Phanerozoic terranes of Western Europe. Results indicate that a thick slab of mantle lithosphere below the craton thins southwestwards beneath the Trans-European Suture Zone and is not seen south of the Variscan front. The thinning is not gradual, but is interrupted by at least two abrupt deep boundaries, the most significant of which corresponds to the surface position of the Tornquist Zone, a major fault. The present geometry of the lithosphere is the result of modification of the margin of the Neoproterozoic continent Baltica by Phanerozoic processes, including the development of the Tornquist Zone and the stretching of the lithosphere in a broad central block SW of this zone. Seismic results and their interpretations from the TOR tomographic project are presented and compared with results from the POLONAISE'97 controlled source project to the SE. Both investigations have shown high-angle, non-symmetrical features extending deep into the mantle. Abstract:
The lithosphere of Eastern Europe is dominated by the East European Craton (EEC), a crystalline complex largely composed of Archaean and Early Proterozoic rocks, assembled in the Late Palaeoproterozoic (Bogdanova et al. 2005). It is flanked to the SW by younger terranes of Neoproterozoic and Phanerozoic age that were accreted to the EEC in the Palaeozoic and subsequently overlain by Mesozoic and Cenozoic successions (Pharaoh et al. 2006). The boundary between the EEC and these western accretionary complexes (Fig. 1) is a wide belt of thrust-emplaced terranes, called the Trans-European Suture Zone (TESZ) (see Fig. 1 and Gee & Zeyen 1996). Here, we present the results of some more recent studies designed to provide information on the lithospheric and asthenospheric mantle across the TESZ. The area investigated covers the southern part of Scandinavia from the exposed Proterozoic crystalline basement rocks of the Baltic (Fennoscandian) Shield, in southern Sweden, to the Palaeozoic, Mesozoic and Cenozoic cover successions of Denmark and northern Germany (Fig. 1). It is an area that has been subject to a variety of very different stress regimes since its establishment as the passive continental margin of the EEC in the Late Neoproterozoic. This thinned southwestern edge of the EEC, composing, in the Early Palaeozoic, the margin of the continent Baltica (Cocks & Torsvik 2006) was subject to mid-Palaeozoic (Caledonian) N E - S W compression, Late Palaeozoic transcurrent faulting, Early Mesozoic extension with lithospheric stretching, and Late Cretaceous to Early Cenozoic Alpine inversion. Thus, the configuration of the lithosphere today is a result of a complex interplay of lithospheric processes that have influenced the edge of the craton over the last 600 Ma. Geophysical investigations of this area were an important component of the European Geotraverse Project (Blundell et al. 1992) and the crustal structure of the area has been investigated by a number of projects (e.g. EUGENO-S Working Group 1988; BABEL Working Group 1993; Rabbel et al. 1995; Abramovitz & Thybo 2000; Thybo 2000; Gregesen et al. 1992; Thybo et al. 1998). These geophysical investigations, supported locally by drilling (e.g. to the Precambrian basement of the Ringkcbing-Fyn High (RFH), across central Denmark, just north of the Thor Suture of Fig. 1), have been interpreted to show that the old crystalline complexes of the Baltic Shield can be followed southwestwards across Denmark beneath the Phanerozoic and partly
Neoproterozoic sedimentary successions at least as far as the Thor Suture (see Fig. 1 and Pharaoh et al. 2006). The crystalline basement of the EEC has been interpreted on crustal seismic data to taper southwestwards from the southern exposures of the Baltic Shield, across the NW-trending fault system of the Sorgenfrei-Tornquist Zone in southernmost Sweden and northern Denmark, and beneath the North German Basin (e.g. Abramovitz & Thybo 2000; Thybo 2000), interrupted only by the thicker crust of the RFH. It is inferred to reach to a line that trends from southernmost Denmark southeastwards across northern Germany into northern Poland. This c. 200 km wide zone of tapering cratonrelated rocks is overthrust by the Neoproterozoic and Early Palaeozoic complexes of Eastern Avalonia and Late Palaeozoic Variscan nappes (Franke 2006). Studies providing information about the deeper structure have also been published, including a number of analyses of fundamental-mode and higher-order Rayleigh waves. The highermode data of Nolet (1977) have been interpreted to distinguish Scandinavia from the Western European structure (see also Zielhuis & Nolet 1994). Concerning the deep lithospheric structure of the Baltic Shield, conflicting results have been published by Cara et al. (1980) and by Dost (1990). Dost (1990) found a low-velocity zone of up to 2% in the depth interval 150-220 km whereas Cara et al.'s (1980) interpretation of the higher modes found no need for a low-velocity zone in the mantle. Some information on very deep structure is also available from earlier P-wave studies. From explosion studies Guggisberg (1986) interpreted several deep low-velocity channels (of about 5% in velocity) in the lithosphere, with the bottom of the lithosphere at about 200 km depth marked by a velocity change of a couple of per cent. Later, Ryaboy (1990) and Thybo & Perchuc (1997) have interpreted low-velocity zones of several per cent beneath the craton at depths of 105-135 km and somewhere between 100 and 280 km, respectively. Because of the limited resolution of earlier studies, the deeper structure of Northern and Central Europe has recently been investigated by a number of projects including TOR, POLONAISE'97, CELEBRATION 2000, SVEKALAPKO and Eifel Plume. The TOR results are discussed below. To the NE of TOR, in Scandinavia, the subcrustal lithosphere of the shield area has been investigated by the SVEKALAPKO project (Sandoval 2002), and to the SW
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 313-322. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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lithosphere. High-quality record sections were obtained for the longest offsets of about 600 km from the shot points, with clear first arrivals and later phases of waves reflected or refracted in the subcrustal lithosphere. The CELEBRATION 2000 project (Guterch et al. 2003) slightly farther to the SE has also provided some information about the upper mantle. Below, results from the POLONAISE'97 project are discussed and compared with results from TOR. With new fieldwork commencing in 2006, this area in Poland will soon be further investigated in a large-scale tomographic project, similar to TOR.
TOR field experiment, data and crustal structure
Fig. 1. The location of seismographs during the field-workof the TOR project, 1996-1997. The red dots indicate short-period seismographs and the blue dots indicate broadband seismographs. The 2D interpretation profile of Figure 5 is perpendicular to the Teisseyre-Tornquist Zone (TTZ) in Poland, whereas the 2D interpretation profile of Figure 4 follows the middle of the cloud of seismographs from 59~ 16~ to 51~ 9~ and is perpendicular to the Sorgenfrei-Tornquist Zone (STZ). The geology (a) and crustal interpretation (b) are from Pharaoh et al. (2006). The Ringk0bing-Fyn High (R-F H), a basement high separating sedimentary basins to the north and south, is located just north of the interpreted Thor Suture. The Elbe Line is a geophysically recognized lineament, locatedjust south of the interpreted Thor Suture.
of TOR, in Germany, by the Eifel Plume project (Ritter et al. 2000). In Poland the POLONAISE'97 project (Guterch et al. 1999) investigated that part of the TESZ known as the Teisseyre-Tornquist Zone (the continuation southeastwards of the the Sorgenfrei-Tornquist Zone in Sweden and Denmark) using a wide-angle controlled source method, revealing seismic velocity structures and reflective interfaces in the subcrustal
In contrast to, for example, POLONAISE'97, where the seismic signals were produced using explosions, TOR was primarily a passive project, recording signals from distant earthquakes and using the characteristics of these recordings to deduce structures below the recording array. Figure 1 shows the position of the recording array, with a strike roughly N E - S W (perpendicular to the Tornquist Zone) from southern Sweden, through Denmark into northern Germany (Gregersen et al. 2002). The position of the array was chosen partly because the crustal structure here had been well investigated by a number of earlier seismic projects, referred to above. This was considered important because the size of the TOR area to be investigated, together with the number of recording instruments available, implied that the distance between recording stations was too large to resolve the crustal structure in any detail. This is further discussed below. The TOR seismic antenna was designed to be relatively long and narrow, as geological evidence and previous geophysical studies have indicated that the large-scale subcrustal lithospheric inhomogeneities in the area are predominantly 2D and oriented N W - S E . However, an array, rather than a single profile, was used to allow some control of the effects of a possibly more complex subsurface geometry. A pilot project was undertaken in 1995 (Kind et al. 1997) consisting of 26 broadband seismographs on a 120kin line from NE Denmark to SW Sweden. It confirmed a rapid change in crustal thickness and seismic velocity across the Tornquist Zone and demonstrated the feasiblity of the proposed main TOR project. The TOR seismic antenna was then deployed in a 900 km long by 100 km wide strip around Profile 1 of the EUGENO-S study and along the German DEKORP reflection seismic line (DEKORP-BASIN Research Group 1999) as depicted in Figure 1. Short-period seismographs (more than 100) were operated from October 1996 to April 1997 (close to half a year) to record distant earthquakes for teleseismic body-wave P- and S-wave travel-time tomography. By chance, they also recorded an earthquake inside the array (Schmidt 1998). The TOR antenna also contained 31 broadband seismographs operated from summer 1996 to summer 1997. These broadband seismographs were mainly used for studies of surface waves (Cotte et al. 2002), anisotropy (Wylegalla et al. 1999; Plomerova et al. 2002), and receiver functions (Gossler et al. 1999; Wilde-Pi6rko et al. 2002). The average station spacing (all stations) is 20 km in the centre of the array in Denmark and southern Sweden, and 2 5 - 3 5 km in northern Germany and central Sweden. The broadband seismographs were placed in triangles of side lengths 4 0 - 6 0 krn. The locations of the triangles were selected so that the data could also be used together with those from permanent broadband seismographs in the area and a few strategically placed temporary stations located off the TOR array, as shown in Figure 1. All the seismographs were placed in isolated areas, as far away as possible from traffic and other man-made disturbances. In the shallow low-velocity lithosphere, rays from distant earthquakes propagate with steep incidence angles. Together with the relatively large station spacing, this severely limits TOR's potential to resolve structures in the sedimentary cover and underlying crystalline crust. It was therefore decided to establish a crustal
DEEP LITHOSPHERE DIFERENCES
model for the area based on other available geophysical data. Two versions of the 3D crustal model were derived, based on different choices of interpolation method between areas with well-resolved crustal velocities, which lie along previous seismic profiles. One of these was derived by Arlitt et al. (1999) following the procedure of Waldhauser et al. (1998), where only interfaces that have confinned seismic reflections are accepted, and smooth mathematical interpolation is done between these. The sediments are included in an a d hoc manner, with little geological reference (Arlitt et al. 1999; Gregersen et al. 1999). In contrast, the model derived by Pedersen (1999) and Pedersen et al. (1999) considers also geological information and other geophysical data, such as gravity. This model was produced by interpolation between structures in published crustal models (EUGENO-S Working Group 1988; Green et al. 1988; Thybo et al. 1989; Stangl 1990; Thybo 1990; Thybo & Sch6nharting 1991; Aichroth et al. 1992; BABEL Working Group 1993; Guterch et al. 1994; Rabbel et al. 1995). Other available information was used to constrain the interpolation. The sediments are included in the modelling through realistic average velocities in the upper layer. Thus, the model of Pedersen et al. (1999) includes more data, which should improve reliability, but also includes more choices in the interpolation procedures, giving greater scope for preconception and bias. The crustal influence on the TOR travel-time anomalies to be interpreted is shown in Figure 2. The Arlitt model version implies crustal corrections of the order of half a second, whereas the Pedersen model involves crustal corrections twice as large. The effects of using these 3D crustal models on the analysis of the mantle structures based on the TOR tomography has been investigated and discussed by Shomali et al. (2002). They showed that, in the specific case of the TOR experiment, because of the character and geometry of the large-scale lithospheric structures, the dominant features of these are well constrained, largely independent of differences between the two crustal models. The largest part of the TOR project was the bodywave tomographic study. Almost 300 earthquakes were well recorded by the array, but not all of these were analysed in detail, partly because earthquakes with sources relatively close to each other do not provide independent information in the tomography, and partly because the inclusion of data with lower signal-to-noise ratios can degrade the results. After the definition of a suitable suite of events, P-wave first arrival times were independently read by workers in the several groups involved in TOR, before being combined into a single dataset for interpretation. An important aspect of this work was a series of tests and comparisons to ensure that students and scientists of the many groups would pick arrival times in a consistent manner. Travel time residuals relative to a standard 1D global model were then computed for each event and station. Examples of P-wave travel time residuals for two events are given in Figure 2 (Pedersen et al. 1999). The lowermost row of Figure 2 presents those residuals caused by the deeper lithosphere-asthenosphere structure after the effect of the crustal model has been removed. Even from each single event, it is clear that major structures exist at depth below the array, and the size and character of these structures is such that they are not masked by the overlying crustal structure.
Interpretation of the TOR data Several different analytical procedures have been applied to the TOR data, including P- and S-wave tomographic, surface-wave, anisotropy, scattering and receiver function analyses. We now discuss some of the main results. Using slightly different methods and datasets, several interpretations, including those by Arlitt (1999), Pedersen et al. (1999) and Shomali et al. (2002), have been made of the P-wave travel time residuals. All of these have found the largest
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subcrustal lateral velocity contrast directly below the Tornquist Zone (see Figs 1 and 4). Although the existence of this feature is very clear, some of the details, including the direction of dip of the boundary (northwards or southwards), are less well resolved. Further independent analyses include that of Horn (2001) using a Monte Carlo inversion and that of Busche (2001) using eigenfunction analysis. A recurring theme in these different analyses is that of the resolution of specific features. Clearly, we would like to extract as much information as possible from the available data, and a natural question concerns which features in the derived models are reliable. Some questions remain, and will be further analysed. However, all analyses confirm the important division of the lithosphere into three blocks, with thick (>200 km) high-velocity lithosphere NE of the Tomquist Zone, a thinner lithosphere (about 100 km) in a central block beneath Denmark, and an even thinner lithosphere in a southwestern block. This model is consistent with analyses of crustal and even some sub-crustal data, but the TOR data now provide resolution down to several hundred kilometres depth. In a separate investigation, Shomali et al. (2006) have extended the tomographic study to include the S waves, in addition to the P waves. Fewer data exist because of lower signal-to-noise ratio for the S phases, so the S-wave velocity model has larger uncertainty than that for the P waves. One outcome of this study has been a confirmation of the most prominent features of the P-wave models. From the P- and S-wave velocity models, the ratio between the two, Vp/V~, has been derived. A difference has been distinguished between the velocity ratios in the three blocks, northern, central and southern. The difference is a couple of per cent, with a ratio of < 1.8 in the two outer blocks, and > 1.8 in the central block. In addition to direct station-to-station analyses, the triangular broadband sub-arrays allow array processing methods to be applied to the surface-wave data, providing phase velocity dispersion information to periods as long as 90 s. The fundamental mode Rayleigh-wave dispersion characteristics of the various paths have been grouped and classified, and the results have been compared with the tomography results. The dispersion characteristics can be grouped into three classes, covering essentially the same areas as those found in the P-wave travel time anomaly study: the North German Basin area, the Danish area almost up to the Tornquist Zone, and the Baltic Shield, NE of the Tomquist Zone. The Rayleigh-wave dispersion curves of each of the three regions have been inverted to estimate the deep S-wave velocity structure. In Figure 3 the S-wave velocities in each area are shown as a function of depth, with a line for each model that is acceptable within the standard deviation of the dispersion curves (i.e. the spread of the lines indicates the resolution). The southern areas (1 and 2) are both interpreted to contain low-velocity channels for the S waves (interpreted as asthenosphere), but at different depths. No low-velocity channel is observed in the shield below Sweden. Exactly where the boundaries between these three areas occur is discussed in a separate paper on the surface-wave results by Cotte et al. (2002). Those researchers have claimed that the northern broadband instrument triangle in Denmark (Fig. 1) shows the same trend of Rayleigh wave dispersion, for periods of 3 0 - 9 0 s, as that in the shield. Therefore they have proposed that the separation between the shield and the Danish lithospheric block dips southwestwards from the northern rim of the Tornquist Zone. However, a consideration of surface-wave theory in laterally inhomogeneous media (e.g. Gregersen 1976) suggests that this dip may not be correctly resolved. Cotte et al. (2002) argued that the dispersion similarity in northeastern Denmark and in the shield to the north indicates similar vertical velocity profiles. The counter-argument, based on surface-wave propagation in laterally inhomogeneous media (e.g. Vaccari & Gregersen 1998), is that this is not a valid argument because
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Fig. 2. P-wave travel time residuals from two earthquakes, in Japan (19 October 1996; 32~ 132~ wave arrival to TOR area from NE) and in Mexico (11 January 1997; 18~ 103~ wave arrival to TOR area from NW), each in one column. For each of the earthquakes, the upper diagram shows total observed residuals (observed arrival times minus expected arrival times according to the global average tables IASP91). The middle diagram shows computed crustal residuals to a depth of 50 km (from Pedersen et al. 1999) which, when subtracted from the observed ones, give the lower lithosphere residuals below 50 km depth; these are shown in the lower diagram.
these long-wavelength surface waves are sensitive to structures at a lateral distance from the site comparable with the depth of the feature. Two studies of shear-wave splitting in the TOR data have been carried out (Wylegalla et al. 1999; Plomerova et al. 2002). Good SKS, SKKS or diffracted and refracted S signals have been collected from the many seismographs, and the polarization of the fast S waves has been determined. One of the studies on horizontal anisotropy (Wylegalla et al. 1999) found the fast S-wave azimuth to be almost e a s t - w e s t in the shield area and beneath the German Basin. Close to the Tornquist Zone, the fast S-wave azimuth is
interpreted to be along the trend of the Tornquist Zone, even where this bends to the south of Sweden (Fig. lb). The other study (Plomerova et al. 2002) retrieved dipping high-velocity structures in three dimensions. Plomerova et al. identified three geographical zones consistent with those found in the isotropic tomography studies. In the NE the fast S-wave structures dip NE, in the central zone NW, and in the southern area the dip is SW. The methods used in the two shear-wave splitting investigations are different, so the two investigations do not necessarily disagree. Using P-wave coda, H o c k et al. (2000) investigated scattering. A difference between the shield and the area to the south of the
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Discussion
Fig. 3. LithosphericS-wave velocity models from the surface-wave investigations of Cotte et al. (2002) in the three blocks of different Rayleigh-wavedispersion.
Tornquist Zone was observed, but no smaller-scale differences were resolvable. Hock et al. deduced that the scattering in the shield is mostly confined to the crust, where the correlation lengths are as short as c. 1 km, with P-wave velocity fluctuations of the order of 4%. Farther south, correlation lengths in the crust are 5 - 1 0 k m , with P-wave velocity fluctuations of c. 8%, and the subcrustal lithosphere also produces some scattering, with correlation lengths 10-20 km and P-wave velocity fluctuations of 6 - 8 % . The scattering attenuation was found to be more important than the anelastic attenuation, which beneath the craton is negligible and further to the SW is small.
All analyses of TOR data are consistent with a sharp and steep boundary that penetrates the whole lithosphere below the Tornquist Zone, and a more shallow near-vertical boundary below the Elbe Line, just south of the Thor Suture of Figure 1, which in recent crustal investigations has been interpreted to define the southwestern boundary of the EEC. There are also some indications of a significant boundary about 100 km NE of the Tornquist Zone within the shield, which could be connected to the known crustal differences across the Protogine Zone (see, e.g. Plomerova et al. 2002) between southwestern and southeastern Sweden. Given the low heat flow of the craton, temperature considerations alone would suggest that there should be significant velocity differences between the northeastern and southwestern parts of the TOR area. However, it is clear that the transition zone is not gradual between the craton in Sweden and the thin lithosphere in Germany, but contains at least two rapid lateral changes in velocity in the upper mantle. In terms of seismic velocity contrasts, the most pronounced lateral variations at sedimentary, deeper crustal and subcrustal depths occur in different geographical locations along the profile. This is illustrated in Figure 4 using the teleseismic tomography model of Shomali et al. (2002). Superimposed on the velocity image are boxes showing where the major structural transitions are deduced to be, based on all available data, not just the P-wave tomography. The teleseismic tomography has poor resolution in the uppermost 50 km or so, and this range is left uncoloured in the figure. However, as referred to above, there is information about shallower structure based on other data (Arlitt et al. 1999; Pedersen et al. 1999). The deduced crustal thickness is shown by the light dotted line. The sloping crustal transition from Baltica to Avalonia in box A, and the wedging out southwestwards of the mantle lithosphere are noteworthy features. The blue and red P-wave velocity anomalies are computed with reference to a 1D global travel-time model (Kennett & Engdahl 1991). The boxes B and C in Figure 4 emphasize two sharp and steep velocity changes. These steps are seen as red-blue steps. The third step is less obvious, between light blue and dark blue, 100 km into Sweden from box C. That the change from red to blue occurs at these points depends, of course, on our choice of presentation parameters. However, a closer examination of the actual velocities involved, or a presentation of lateral velocity gradients (as opposed to estimated velocities) clearly shows that Figure 4 is not misleading because of the choice of colours. The existing uncertainties in the exact locations of the transitions and their slopes are illustrated by broad boxes (B and C) in the transition regions. Further studies are in progress to, as far as possible, enhance and define the resolution of these features. Although the various seismological investigations, such as the tomographic inversions of Arlitt (1999) and Shomali et al. (2002), agree regarding the major features of the derived models (zones B, C, etc.), some important details in the images differ, notably the slope of the deep boundary beneath the Tornquist Zone. Box D has been introduced as the bottom of the low-velocity layer, deduced from higher-mode surface-wave studies (Nolet 1977; Dost 1990). Box D does not extend to the NE, consistent with Cotte et al.'s (2002) interpretation from the fundamental mode Rayleigh waves that there is no low-velocity layer below the shield. This conclusion is not, however, consistent with some previously published studies of higher-mode surface-wave data and some P-wave controlled source data, which deduced lowvelocity layers (see above). Both fundamental-mode and highermode surface waves are mainly sensitive to the S-wave velocity structure, but because of their shorter wavelengths, higher modes are more sensitive to rather thin low-velocity layers. However, the TOR measurements provide a considerable quantity
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Fig. 4. Tomographic image of the lithosphere-asthenosphere system in a profile along the TOR array (from 59~ 16~ to 51~ 9~ STZ, Sorgenfrei-Tornquist Zone; RFH, RingkCbing-Fyn High; EL, Elbe Line. Blue areas show high P-wave velocities; red areas show low P-wave velocities with respect to the laterally homogeneous IASP91 model (Kennett & Engdahl 1991). Box A is the crustal transition region (e.g. Abramovitz & Thybo 2000). The generalization boxes B and C are drafted taking into account the TOR results of Arlitt (1999), Pedersen et al. (1999), (2001), Horn Busche (2001), Cotte et al. (2002), Shomali et al. (2006) and Voss et al. (2006). Moho depths are from Thybo et al. (1998) and Gossler et al. (1999). Box D in the southern part delimits the material below the asthenosphere at depth 230 km, from Nolet (1977) and Dost (1990). of new very high quality data from the area, suggesting that the previously deduced low-velocity zones may be incorrect. In this paper only the most significant well-resolved features have been discussed. There is, of course, more information in the data than this. Thus, for example, several other features in Figure 4 could be geologically meaningful. However, resolution generally decreases with the spatial size and velocity contrast of a feature. It can be difficult to assess just which features in the models should be regarded as well defined and suitable to interpret in geological terms. It follows from this that issues regarding spatial resolution, accuracy and uniqueness of the models are important. Within TOR, considerable effort has therefore been dedicated to examining these issues in, for example, tomographic inversions. In such an inversion it is straightforward to calculate, for example, the variance of model parameters, which quantifies the reliability of each parameter. The resolution is limited by station spacing, ray geometry between earthquakes and seismic stations, and the frequency content of data, and it is different for different parts of the model. The calculated resolution is dependent on choices in the mathematical inversion procedure. Furthermore, the variance of a model parameter describes only a part of the problem, partly because it makes a simplifying mathematical assumption of linearity, and partly because it ignores the often
complicated interactions between model parameters (covariance). The complexity of these problems means that several different approaches can be considered to quantify reliability. Some information on this can be gained by solving the mathematical inverse problem stepwise linearly, as was done by Arlitt (1999), Shomali (2001) and Shomali e t al. (2002). Other approaches used have been relative model testing (Shomali e t al. 2002), resolution kernel and synthetic tests (Shomali 2001), sigmoid function and Fourier component computations (Busche 2001), and Monte Carlo model evaluations (Horn 2001; Voss e t a l . 2006). All the results indicate that the major transitions and blocks that have been discussed here are resolved well by the observed data. It is interesting to compare the TOR results with those from POLONAISE'97, which also crossed the TESZ, but farther to the SE. In POLONAISE'97, 2D interpretation was carried out using a ray tracing approach for the reversed system of P- and S-wave travel times (Grad e t al. 2002). The model of the seismic P-wave velocity structure beneath profile P4, perpendicular to the Tornquist Zone in Poland, includes several reflectors in the lower lithosphere (Fig. 5). The corresponding reflected waves are interpreted as originating from a sub-Moho reflector (PI), the top of a low-velocity zone (PH, usually poor) and from discontinuities at depths of c. 80 and c. 90 km (Pro and Piv, respectively; Fig. 5). The velocity
Fig. 5. Subcrustal reflectors on profile P4 crossing the Tornquist Zone from the POLONAISE'97 experiment (Grad et al. 2002). The displacement of sharp lateral transitions in various depth ranges should be noted.
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beneath the Moho is found to be rather high (c. 8.25 km s -1) in the Palaeozoic terranes in the SW and normal (c. 8.1 km s -1) in the East European Craton to the NE. The thickness of the crust is 30 km in the SW and 4 0 - 5 0 km in the EEC. The subcrustal, almost vertical inhomogeneities between the various regions (Fig. 5) are displaced into the EEC, with respect to the crustal inhomogeneities, similar to the sloping derived by Pedersen et al. (1999), Busche (2001), Shomali et al. (2002) and Voss et al. (2006). The variations in P- to S-wave velocity ratio reported by Shomali et al. (2005) from TOR P- and S-wave tomographic studies, with an anomalous central block, are consistent with receiver function results for the uppermost mantle in Poland (Wilde-Pirrko et al. 2002), where the central block is the Tornquist Zone. Thus, the TOR and POLONAISE'97 data both reveal deep structures of the TESZ, and several significant components in the derived models show similar properties, despite the considerable geographical separation of the projects. In the TOR area, the extreme southeastern boundary of the EEC crust deduced from geological and geophysical data lies below the Thor Suture (just north of EL in Fig. 4), but there is a clear subcrustal transition zone as far north as the Tornquist Zone, which can be characterized as the edge of the undisturbed (non-stretched) craton lithosphere. A different and complementary way to describe the edge of the craton is by its rheological behaviour (i.e. its reaction to regional stress). In studies on the small, local earthquakes of the region Gregersen et al. (1996, 1998) suggested that the rheological edge of the craton is located between the STZ and RFH. Furthermore, analyses of recent crustal movement data (Lykke-Andersen & Borre 2000; Gregersen & Schmidt 2001) have shown that the STZ is a separation between different geodetic movement patterns. As shown in the papers by Cotte et al. (2002) and Gregersen et al. (2002), we interpret the surface-wave results of Figure 3
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and the tomography results of Figure 4, as well as the other geophysical data, to show a lithosphere of thickness a little under 100 km in the southwestern part of the profile, a little over 100 km in the central part, and more than 200 km in the northeastern part of the profile. The Baltic (or Fennoscandian) Shield (i.e. the area of the craton where old, Precambrian, mostly Palaeoproterozoic-Archaean crystalline rocks are exposed), terminates to the SW at the Tornquist Zone. From deep drilling data and seismic crustal studies, it is deduced that material originating from the craton extends southwards from the Tornquist Zone, gradually thinning below the overlying sediments, and terminating at the Thor Suture, just south of the RFH. A seismic transition is seen in this area, consistent with craton material extending to a depth of the order of 100 km. In the latest Precambrian to early Palaeozoic (c. 600 Ma), the EEC composed the core of the palaeo-continent Baltica, and the TOR area was a passive continental margin. The tomographic images reveal the structure originating from this episode, reworked by a number of later events, as shown in Figure 6. The large-scale development of the transition zone between the Proterozoic craton to the NE (right side of Fig. 6) and the Phanerozoic lithosphere to the SW (left side of Fig. 6) can be described in the block diagrams of Figure 6. Stage 1 is the situation of two separate lithospheric plates colliding in mid-late Palaeozoic times. In stage 2, the area was deformed by lithospheric stretching and transcurrent faulting, the R i n g k c b i n g - F y n High (RFH) with its thick crust detached from the Baltic Craton, and the various blocks of the RFH rotated slightly, separately. In the ensuing stage 3, the lithospheric stretching was perpendicular to the trend of the plate transition and the RFH. The compressional stage 4 is very different, involving inversion of the SorgenfreiTornquist Zone.
Fig. 6. Generalizedsummarydiagram of the broad-scale geologicaldevelopmentof the TOR area (Fig. 1). Blue shows sediments, orange shows crystallinecrust, and red is the uppermost mantle lithosphere. The Toruquist Zone (STZ) and the Ringkoebing-Fyn High (RFH) act through time as compression, spreading, shearing and compressionzones. Very thick arrows show regional stress field. (For further explanation, see text).
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Conclusions The TOR project successfully collected an extensive high-quality seismic dataset, suitable for many seismological analysis methods and allowing the production of high resolution P- and S-wave velocity models for the transition across the Trans-European Suture Zone, which separates the East European Craton from the Phanerozoic lithosphere of Western Europe. The models produced are consistent with previous models based primarily on geological modelling of controlled source crustal seismic and potential field data. They relate surface geological observations and crustal structure with the deeper lithosphere. A major, very deep, near-vertical boundary is observed below the Tornquist Zone. The velocity contrast occurs over too short a lateral distance to be explained by increasing temperature farther from the Shield. Although anisotropic velocity structures in the mantle could contribute to the observations, the feature suggests a compositionally significant boundary. If, as indicated by crustal seismic data and drilling, material from the proto-continent Baltica exists as far south as the Thor Suture, this suggests that this deep feature below the Tornquist Zone cannot have recent origins. Presumably, it originated in some form at the time of development of the Baltica passive margin during Caledonian collision, and has been altered as a result of later phases of regional deformation. The boundary in the neighbourhood of the Thor Suture appears to be much shallower, but still extends to a depth of over 100 km. We infer this to mark the boundary of stretched material from Baltica. No 'seismic asthenosphere' in the sense of a low-velocity layer (relative to the material both above and below) appears to be present below the craton, but is apparent to the SW. Striking parallels in the models from the TOR and POLONAISE projects are seen, suggesting that significant parts of the derived models cannot be related to some local component of the geological evolution, but rather reflect a more fundamental aspect of the large-scale evolution of the area. Clearly, the geological evolution of the crust is dependent upon the evolution of the entire lithosphere in the area, and, to fully understand the crustal evolution, we must understand that of the underlying lithosphere. TOR was a large-scale project, but nevertheless (as we can see from e.g. Fig. 4) there are lithospheric-asthenospheric features on a scale comparable with that of the TOR array. It seems therefore that, to fully understand the TOR data, it will be necessary to place the data into a larger spatial perspective. A profile for the TOR area extracted from the global tomographic results of Bijwaard et al. (1998) shows some similar features to the TOR teleseismic travel-time tomographic results (W. Spakman, pers. comm.), but the locations of the transitions and the regional differences are much better delineated in Figure 4 than in the large-scale global model. This supports the reliability of both models, and it follows that a combination of the two datasets will be very valuable. Similarly, the TOR model fits very well with later, as yet unpublished, tomographic studies to the north in Sweden (R. Roberts, pers. comm.). The data from the SVEKALAPKO (Sandoval 2002) and Eifel Plume (Ritter et al. 2000) projects may also be incorporated. Clearly, one future direction will be to integrate the data from these different projects to create a geophysical and geological model including length scales of thousands of kilometres, depths to 1000 km, and resolution of a few tens of kilometres, even at great depth. The TOR work has been carried out within the framework of EUROPROBE's international TOR Working Group with the following members, besides the five main authors: L. B. Pedersen, A. Berthelsen, H. Thybo, K. Mosegaard, T. Pedersen, R. Kind, G. Bock, J. Gossler, K. Wylegala, W. Rabbel, I. Woelbern, M. Budweg, H. Busche, M. Korn, S. Hock, A. Guterch, M. Wilde-Pi6rko, M. Zuchniak, J. Plomerova, J. Ansorge, E. Kissling, R. Arlitt, F. Waldhauser, P. Ziegler, U. Achauer, H. Pedersen, N. Cotte, H. Paulssen and E. R. Engdahl. Many scientists have supported the project as members of the original TOR planning group or through the field-work with
advice. The following scientists are thanked for their participation: D. Gee, R. Gorbatschev, A. Tryggvason, N. Juhojuntti, H. Wagner, N. Balling, B. H. Jacobsen, P. H. Nielsen, W. Hanka, P. and E. Bankwitz, M. Weber, H.-P. Harjes, A. Biegling, J. Skamletz, E. Perchuc, W. Spakman, J. Zednik, T. Hyvonen, S.-E. Hjelt and L. N. Solodilov. The TOR project has been supported in Germany by the GeoForschungsZentrum Potsdam through personnel and seismographs from the German seismograph pool, and by the Deutsche Forschungsgemeinschaft; in Switzerland by the Swiss National Science Foundation under contract 21-43444.95; in Denmark by the Danish Natural Science Research Council, grant 9401105; and in Sweden by the Swedish Natural Science Research Council, contract G-AA/GU 04990-336. The field-work of the Polish groups was partly supported by the Institute of Geophysics, Polish Academy of Sciences. The research of the Czech group was partly supported by a grant from the Czech Academy of Sciences.
References ABRAMOVITZ, T. & THYBO, H. 2000. Seismic images of Caledonian.
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Palaeozoic accretion of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories J. A. WINCHESTER a, T. C. PHARAOH 2, J. VERNIERS 3, D. IOANE 4 & A. S E G H E D I 5 1School of Physical and Geographical Sciences, Keele University ST5 5BG, UK (e-mail j. a. winchester @esci. keele, ac. uk) 2British Geological Survey, Kingsley Dunham Centre, Keyworth NG12 5GG, UK 3Ghent University, Palaontologie, Krijgslaan 281/$8, B 9000, Gent, Belgium 4Faculty of Geology and Geophysics, University of Bucharest, Bucharest, Romania 5Geological Institute of Romania, 1 Caransebes St, 012271 Bucharest 32, Romania
Abstract: Recent work in Central Europe, combined with emerging information about basement massifs in SE Europe and NW Turkey, permits a new look at the relationships between crustal blocks abutting the East European Craton (EEC) along the Trans-European Suture Zone (TESZ). The simplest model indicates that the end-Cambrian establishment of the Bruno-Silesian, Lysogory and Matopolska terranes close to their present location on the SW margin of the EEC formed a major promontory on this margin of the continent. Moesia may also have formed part of this block. Both late Ordovician accretion of Avalonia and early Carboniferous accretion of the Armorican Terrane Assemblage (ATA) attached new continental material around the Bruno-Silesian Promontory (BSP). Palaeozoic faunal affinities and inherited isotopic signatures similar to those of Avalonia seen in the Istanbul block of NW Turkey, and in minor thrust slices in Moravia and Romania, suggest that easternmost Avalonia was severed, on collision with the BSP, and migrated east along the southern margin of the EEC. Likewise, the similarities to the ATA of the Balkan, Istranca, Sakarya and eastern Pontides blocks suggests that more easterly components of the ATA were detached at the BSP and migrated east. All the newly accreted blocks contain similar Neoproterozoic basement indicating a peri-Gondwanan origin; Palaeozoic plume-influenced metabasite geochemistry in the Bohemian Massif may explain their progressive separation from Gondwana before their accretion to the EEC. Inherited ages from Avalonia contain a 1.5 Ga 'Rondonian' component arguing for proximity to the Amazonian Craton at the end of the Neoproterozoic; Armorican terranes lack such a component, suggesting that they have closer affinities with the West African Craton. Models showing the former locations of these terranes and the larger continents from which they rifted, or later became attached to, must conform to both these constraints and those provided by palaeomagnetic data. In the late Neoproterozoic and Palaeozoic, these smaller terranes, some containing Neoproterozoic ophiolitic marginal basin and magmatic arc remnants, probably fringed the end-Proterozoic supercontinent as part of a 'Pacific-type' margin. When this margin fragmented, most resulting fragments accreted to the EEC.
The S W margin of the East European Craton (EEC) is marked by the Trans-European Suture Zone (TESZ), traceable from the Black Sea coast of Romania to North Germany, the Baltic Sea, Denmark and the North Sea (Fig. 1), despite being everywhere concealed beneath thick sedimentary cover (Gee & Zeyen 1996; Pharaoh 1999). A description of the nature, age and geometry of this fundamental feature of European geology has been provided by Pharaoh et al. 2006). On the SW side of this zone a collage of blocks accreted to the EEC margin during the Palaeozoic following the end of the Cambrian. SW of the TESZ and north of the A l p i n e - C a r p a t h i a n Front, the basement structure of Central Europe has long been known in outline. Evidence from geophysical compilations, geological information provided by deep boreholes, and outcrops of Palaeozoic and older rocks across central Germany and in the Bohemian Massif reveals a mosaic of microcontinental blocks, derived from different sources and shown by isotopic dating, and biostratigraphic and palaeomagnetic evidence, to have become attached to the EEC in their present locations during the Palaeozoic. An early phase of terrane emplacement was largely complete by the end of the Cambrian. These terranes, including the BrunoSilesian Terrane, with the Lysogory and Ma~opolska blocks of the Holy Cross Mountains in Poland appear to have been situated in approximately their present position since that time. They may extend southwards to the Danube, approximately as far as the K r e m s - V i e n n a Line in Austria (Dudek 1980) and also be linked to the SE beyond the Carpathians with the central and southern Dobrogea and the Moesian Platform in Romania. Whether these terranes comprise displaced portions of the EEC (Cocks 2002) or an early accreted fragment derived from Gondwana (e.g.
Belka et al. 2000, 2002) has been hotly debated, and the arguments have been set out more fully by Pharaoh et al. (2006). However, the end-Cambrian attachment of these blocks to the EEC also precludes any pre-Ordovician association with terranes accreted later, particularly Avalonia, which was still attached to Gondwana in the early Ordovician (Winchester et al. 2002). Whatever their derivation, if it is accepted that the subsequent Devonian displacement of these terranes, suggested from palaeomagnetic and structural evidence (Lewandowski 1993; Mizerski 1995), was restricted in extent (Cocks 2002), they must have formed a major promontory extending from the SW margin of the EEC during most of the Palaeozoic. This is referred to below as the Bruno-Silesian Promontory (BSP) and its geometry is crucial in explaining the mechanisms of attachment of the microcontinents that subsequently accreted to the EEC. Excluding the small portions of Laurentian crust forming Scotland and N W Ireland, the main microcontinental blocks that subsequently accreted to the SW margin of Europe during the Palaeozoic are known as Avalonia and the Armorican Terrane Assemblage (ATA; Franke 2000; Tait et al. 2000). Both of the latter were derived from Gondwana, but rifted from it at different times. They therefore possess characteristic Proterozoic basement, affected by end-Proterozoic Panafrican (locally termed Cadomian) magmatism and deformation, which therefore does not distinguish between them. Factors distinguishing these microcontinental blocks are: (1) the timing of accretion to the EEC; (2) the presence or absence of an inherited c. 1.5 Ga 'Rondonian' event, which seems, in particular, also to be a characteristic feature of the 'Ganderian' portion of Avalonia; (3) the occurrence of either distinctive
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphere Dynamics. Geological Society, London, Memoirs, 32, 323-332. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. A map showing the distribution of crustal blocks and Palaeozoic deformation belts in Central and SE Europe. ABDB, Anglo-Brabant Deformation Belt; AD, Ardennes; ADF, Alpine Deformation Front; AM, Armorican Massif; BB, Brabant Massif; BK, Balkan Terrane; BM, Bohemian Massif; BSM, Bruno-Silesian Massif; CACC, Central Anatolian Crystalline Complex; CAU, Caucasus; CD, Central Dobrogea; CDF, Caledonian Deformation Front; CDO, Central Dobrogea; CM, Cornubian Massif; COF, Capidava-Ovidiu Fault; DR, Dronsendorf Unit; EA, Ebbe Anticline; EL, Elbe Lineament; EP, Eastern Pontides; GF, Gf6hl Unit; HCM, Holy Cross Mountains; HPDB, Heligoland-Pomerania Deformation Belt; IB, Istanbul Block; IMF, Intra-Moesian Fault; Istr, Istranca Terrane; KLZ, Krakow-Lubliniec Zone; LU, Lysogory Unit; L-W, Leszno-Wolsztyn High; MC, Midlands Microcraton; MM, Matopolska Massif; MN, Mfinchberg Nappe; MNSH, Mid-North Sea High; MP, Moesian Platform; MST, Moravo-Silesian Terrane; NASZ, North Armorican Shear Zone; NBT, North Brittany Terrane; NDO, North Dobrogea; NGB, North German Basin; PCF, Peceneaga-Camena Fault; Pom, Pomerania; POT, Polish Trough; R, Rfigen Island; RFH, Rynkcbing-Fyn High; RG, RCnne Graben; Rh, Rhodope; SASZ, South Armorican Shear Zone; SBT, South Brittany Terrane; SDO, South Dobrogea; SGF, Sfantu Gheorghe Fault; SNSLT, South North Sea-Luneberg Terrane; SP, Scythian Platform; S-TZ, Sorgenfrei-Tornquist Zone; Su, Sudetes; TB, Teplfi-Barrandia; T-TZ, TeisseyreTornquist Line; VF, Variscan Front; ZZ, Zonguldak Zone. 'Celtic' (e.g. Avalonian) faunas that were unique to Avalonia during the mid- and late Ordovician, or of distinctive mixed Siluro-Devonian faunas characteristic of the ATA; (4) the presence of late Ordovician glaciogenic sediments, a Gondwana feature shared by the ATA, but not by Avalonia, which had by that time already migrated into lower latitudes (Cocks et al. 1997).
Avalonia Precambrian and early Palaeozoic basement exposed in central England, Belgium and western Germany is widely accepted as part of Avalonia, this Ordovician microcontinent extending west as far as New England, and being best exposed in the Avalon Peninsula of Newfoundland, after which it is named. Avalonian basement in central England typically consists of late Proterozoic intrusive, volcanic and sedimentary rocks (e.g. Thorpe et al. 1984; Pharaoh & Gibbons 1994; Strachan et al. 1996), affected by end-Proterozoic or pre-Early Cambrian deformation. In the English Midlands it was little affected by later movements, and is overlain by a thin Lower Palaeozoic shallow marine sedimentary sequence, succeeded conformably by Devonian terrestrial deposits: the 'Old Red Sandstone'. For this reason it has sometimes been called the 'Midlands Microcraton' (e.g. Turner
1949; Pharaoh et al. 1987). A fuller description has been given by Pharaoh et al. (2006). The Midlands Microcraton is flanked to the NW by much thicker Lower Palaeozoic successions, strongly deformed in an early Devonian 'Acadian' event (Soper et al. 1987), deposited on Avalonian basement and exposed in Wales, the English Lake District and SE Ireland. Boreholes in eastern England reveal that similarly deformed rocks (Pharaoh et al. 1987) containing Upper Ordovician calc-alkaline volcanic rocks, as an apparent continuation of the Lake District Arc, also extend from eastern England to Belgium, where they are exposed in the Brabant Massif (Andr6 et al. 1986; Pharaoh et al. 1991). These rocks have been termed (Winchester et al. 2002) the Anglo-Brabant Deformation Belt (ABDB), in which the deformation is inferred to have developed in the early Devonian (Acadian). This deformation belt is thought to mark a zone of crustal suturing inherited from the late Ordovician soft collision of Avalonia and Baltica (Verniers et al. 2002). Unusually thick lower Cambrian deposits in Brabant suggest the presence of rifting, which may have thinned and weakened the Proterozoic basement, thereby controlling the location of this deformation belt (Winchester et al. 2002). Because the ABDB contains no known ophiolitic material, it is not thought to mark a zone of microcontinent collision and seaway destruction, even though,
DETACHED TERRANE FRAGMENTS IN EEC as suggested by Pharaoh et al. (1993), it may separate crusts with somewhat differing structures. An area of stable basement, indicated by seismic traverses and termed the Southern North Sea-Luneberg Terrane (SNSLT) by Pharaoh et aL (1995), lies east of the ABDB, NE of the DowsingSouth Hewett Fault Zone-Lower Rhine Lineament, themselves younger reactivations of earlier major fault-lines (Pharaoh 1999). It has been more recently dubbed 'Far Eastern Avalonia' (Winchester et al. 2002), because of its geological similarities to Avalonia. Although crystalline basement is mostly concealed, one outcrop area, the 574 _+ 3 Ma Wartenstein Gneiss (Molzahn et al. 1998), is exposed far to the south in the South Hunsriick at the SE margin of the Rhenish Massif; lying to the south of the Variscan Front, this typically calc-alkaline granitoid gneiss of late Neoproterozoic age is broadly comparable with granitoid rocks in the Avalonian basement in central England. In addition, Samuelsson et al. (2002a) noted that the eNd(t) trends of Ordovician sedimentary rocks from the Ebbe Anticline of NW Germany (Fig. 1), situated NE of the Lower Rhine Lineament and therefore on SNSLT basement, match those from the Welsh Basin and the Brabant Massif, but are different from those in Brittany and Iberia. They also concluded that these rocks formed part of Avalonia. Hence the ABDB is interpreted as an intra-Avalonian zone of local subduction, initiated above a failed Cambrian rift, where the basement had been thinned and weakened, in response to the distortion and anticlockwise rotation of part of Avalonia as it moulded itself on to the margins of the EEC and Laurentia (Verniers et al. 2002). The late Ordovician timing of the accretion of the SNSLT to the EEC (Vecoli & Samuelsson 2001; Samuelsson et al. 2000b), which only slightly predates Avalonian convergence with Laurentia, based on evidence from Atlantic Canada (e.g. Cawood et al. 1994), and the onset of Windermere Supergroup sedimentation in the English Lake District (Cooper et al. 1993), also suggests that the SNSLT should be considered as part of Avalonia. If so, the Heligoland-Pomerania Deformation Belt (HPDB), which largely comprises a zone of overthrusting, marks the collision zone between Avalonia and the EEC. It shows little evidence of contemporary magmatism in boreholes, but geophysical evidence may indicate the presence of buried arc volcanic rocks (Williamson et al. 2002; Pharaoh et al. 2006), suggesting that convergence was accompanied by southdirected subduction beneath the Avalonian margin.
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(Fig. 2). Recognizing the Avalonian affinities of such detached extensions, where later metamorphism may have destroyed faunal evidence, is problematic; discrimination between Avalonia and the ATA must rely instead on the presence (as in Avalonia), or absence (as in the ATA) of mid-Proterozoic (1.45 Ga) inherited zircon dates, related to the previously adjacent (pre-rifting from Gondwana) Rondonian event in the Amazonian Craton.
Avalonian easternmost extremities Although totally concealed by thick Mesozoic and Cenozoic sequences in the Polish Trough, the easternmost end of Avalonia appears to abut the Lysogory and Matopolska blocks of the Holy Cross Mountains, which form part of the Bruno-Silesian Promontory (BSP). Further south, tectonic structures along the western margin of the Bruno-Silesian Block show highly oblique (dextrally transpressive), complex overthrusting to the east (Moldanubian and Drinova thrusts) in the early Carboniferous, between 350 and 330 Ma (Schulmann & Gayer 2000). This junction is traceable northwards beneath the thick sedimentary cover of the Polish Trough, using seismic profiling. Both the Polonaise PI and TeisseyreTornquist Zone (TTZ) profiles (Grad et al. 1999; Jensen et al. 1999) show a clear change of mid-crustal structure north of the Moldanubian Thrust, suggesting that it continues northward as a major feature (Moravian Line of Winchester et al. 2002). The TTZ profile shows the mid-crustal break to be displaced eastwards compared with Polonaise P 1, indicating dextral displacement of the Moravian Line by strike-slip faulting between the two profiles, perhaps along the Dolsk Line (Grad et al. 2002). On accretion, Avalonia was unlikely to fit exactly into the position against the EEC margin that it now occupies, bounded to the east by the BSP. Therefore, former eastern extensions are likely to have been detached by sheafing initiated by collision with the BSP
Fig. 2. (a) Sketch map illustratingthe supposed configurationof Avalonia on its initial impact with the Bruno-SilesianPromontory. (Note the detachment and displacementeastward of its eastern extremity). Abbreviationsas in Figure 1. (b) Sketch illustrating the likely configurationof the Armorican Terrane Assemblageon initial impact with the Bruno-SilesianPromontory in the early Carboniferous. GWSO, Giessen-Werra-Sudharz Ocean; MGCH, Mid-German CrystallineHigh. Other abbreviationsas in Figure 1. (c) Likely configuration of crustal blocks followingthe main Variscan Orogeny. (Note the eastward displacementof the eastern Variscides.)
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The M o r a v i c u m Nappe
In NE Austria and Moravia, the Dobra Gneiss (Gebauer & Friedl 1993) and Bittesch Gneiss (Friedl et aL 2000) both yield midProterozoic inherited zircon dates of c. 1.5 Ga, contemporary with the Rondonian orogeny affecting the NW side of the Amazonian Craton (Tassinari et al. 2000; Cawood et al. 2003). They also show clusters of inherited dates at 1.2 and 1.78 Ga, which have also been noted both from Ganderian (Avalonian) basement in southern New Brunswick and from the Amazonian Craton (where they are termed the Sunsas and Rio Negro provinces; Tassinari et al. 2000). Because these rocks were thought to be linked to the Bruno-Silesian massif (BSM), these dates were interpreted to assign an Avalonian affinity to the BSM (Finger et al. 2000). However, if the latter was attached to the EEC, at least since the end of the Cambrian, it cannot also have been part of Avalonia. However, both the Dobra and Bittesch gneisses are situated in the deformed Moravicum nappe, emplaced onto the western margin of the Bruno-Silesian block by movement on the Drinova Thrust (Hock et al. 1997; Melichar & Kotkova 2003). They therefore need not belong to the BSM, and could instead represent small detached slivers of Avalonian crust thrust obliquely onto the margins of the BSM. The Danubian basement o f the southern Carpathians in Romania
Exposed in nappes towards the western end of the southern Carpathians of Romania are rocks of the Lower and Upper Danubian basements (Berza et al. 1983, 1994, 2004; Iancu & Berza 2004). Late Neoproterozoic magmatic rocks (Liegeois et al. 1996) including calc-alkaline granitoids such as the Tismana Pluton (567 + 3 Ma, U - P b ) are unconformably overlain by a clastic sedimentary succession of Late Ordovician-Early Silurian age, apparently devoid of recorded glaciogenic diamictites. These rocks underwent a Devonian (?Acadian) deformation and, although clear faunal or palaeomagnetic evidence remains lacking, and some claim that the older rocks may be exhumed Moesian basement (Sandulescu 1984, 1994), these characteristics make an Avalonian affinity an alternative possibility. The Istanbul Block
Further east, in the Istanbul Block of NW Turkey, lithologically similar Ordovician rocks have yielded 'Celtic' (e.g. Avalonian) faunas (Kozur & G6nctio~lu 1998; Dean et al. 2000). Some studies have distinguished, in this area, separate Istanbul and Zonguldak terranes, based on differences of facies between Palaeozoic rocks close to Istanbul and those further east (Goncuoglu & Kozur 1998, 1999; Kozur & Gincuoglu 2000). Also, whereas sedimentation near Istanbul seems to have continued uninterruptedly from the Early Ordovician to the MidCarboniferous, further east there are increased signs of Early Devonian uplift and deformation. However, no clear boundary between these terranes can be mapped, and it seems more likely that these differences relate to facies changes within a single terrane, analogous to those seen within Eastern Avalonia, in which Central English and Welsh successions can be contrasted. According to this analogy, the continuous sequence in the Istanbul area corresponds better to the Lower Palaeozoic shelf deposits in Central England, whereas the overlying Upper Palaeozoic rocks, containing mixed shales, cherts and limestones and their benthic fauna (Tokay 1955) have more in common with the marine Devonian and Carboniferous units of the Rheno-Hercynian zone, which in both SW England and Germany overlie Avalonian basement. By contrast, the Zonguldak area, after the deposition of an initial pebbly quartzite, reveals a shale-dominated Ordovician sequence more characteristic of the Welsh Basin, and this
analogy is enhanced by the presence of a Lower Devonian unconformity. Ensuing Late Devonian and Early Carboniferous sedimentation is dominated by limestone, which is in turn succeeded by regressive flood plain deposits containing coals (Yanev et al. 2006). Both this Upper Palaeozoic sequence and the major unconformity above which Permo-Triassic continental clastic rocks occur are analogous to the sequence seen in England overlying Avalonian basement. The Avalonian link for the Istanbul Block, suggested by both the lithological sequences and the faunal evidence, is further strengthened by a mid-Proterozoic discordia date of 1 4 4 5 _ 24 Ma obtained from a late Neoproterozoic granite in the Karadere basement (Chen et al. 2002). A further, less well-constrained inherited age of 1189 _+ 110 Ma from a metatonalite in the same area (Chen et al. 2002) also resembles some of the ages obtained from the Moravian Nappe, and from the Ganderian part of Avalonia in southern New Brunswick. This information, combined with evidence of Silurian deformation in the northern part of the block near Zonguldak (although not nearer to Istanbul itself) also is consistent with Siluro-Devonian docking with Baltica of an Avalonian fragment, and post-Carboniferous (Variscan) deformation, as in the Rheno-Hercynian zone of Western Europe, may attest to the deformation of the southern Avalonian margin caused by accretion of terranes of the ATA. Moesia and Dobrogea
In contrast, there is little evidence to link the basement rocks of any part of either the Moesian Platform or the Dobrogea with Avalonia. Neoproterozoic and Lower Palaeozoic rocks of the central and southern Dobrogea, in eastern Romania, appear to have similarities to those in the Holy Cross Mountains of Poland rather than the Istanbul Block, and borehole sections indicate that this is also true of the northeastern part of the Moesian Platform, NE of the Intra-Moesian Fault (IMF). Metamorphic basement to the southern Dobrogea has also yielded Mid-Proterozoic ages (Giusca et al. 1967), and is considered to have affinities with rocks in the Ukrainian Massif of the EEC. Residual gravity and magnetic anomalies also indicate a link with the EEC (Ioane & Atanasiu 2000). However, as the Bruno-Silesian Massif itself may also have originated adjacent to the southern EEC, it is possible that this basement may also underlie the Central Dobrogea at greater depth. West of the IMF, Neoproterozoic granitoids have also been recovered from boreholes (Savu & Paraschiv 1985), and Cambrian rocks from deep boreholes (Mutiu 1991) have yielded numerous fragmentary specimens of the trilobites Paradoxides (species undetermined) and Peronopsis fallax (Linnarsson). A. Rushton (pers. comm.) considered that the latter fossil, although apparently showing some affinities with species associated with the Baltican margin (Rushton & McKerrow 2000), is a widely recorded member of an outer shelf fauna, which may have been able to cross geotectonic boundaries. Thus, although neither trilobite can be used to prove conclusively a periBaltican affinity of the Moesian Platform, similar to that proposed for the Dobrogea or Bruno-Silesian regions, this remains its most likely affinity. However, the presence of a widespread unconformity beneath Silurian rocks (Iordan 1984) may indicate a Late Ordovician uplift, which could record deformation associated with collision of Avalonian fragments immediately to the south.
Avalonian eastern extremities: mechanism of emplacement Lower Palaeozoic rocks deposited directly on the EEC margin are all consistent with a passive margin setting: the narrowing of the Tornquist Sea appears to have occurred exclusively by subduction under the Avalonian margin up to the time of collision in the late
DETACHED TERRANE FRAGMENTS IN EEC Ordovician. Further east, however, the southern margin of the BSP is now concealed beneath younger rocks associated with the much later formation of the Carpathians. Yet, for fragments of Avalonia to migrate east with dextral transpression, continued subduction was probably needed, but the Silurian rocks in the Istanbul Block do not include magmatic rocks indicative of continued subduction. It therefore seems likely that, on collision of easternmost Avalonia with the BSP, a change of polarity of subduction occurred analogous to that recorded in New Brunswick with closure of the Iapetus Ocean in that sector (van Staal et al. 1991). With continued subduction, but this time northwarddirected, the buoyant continental fragments of Avalonia could be transported eastwards with sinistral transpression, along the southern margin of the EEC, until 'trapped' in re-entrants of the continental margin (Fig. 2a). Smaller fragments, such as the Moravicum Nappe and the Danubian Terrane, may be interpreted as slivers on the continental margin, abandoned during the eastward progress of the Istanbul Block. The present position of the Istanbul Block partly arises from its southward displacement during the opening of the Black Sea basins, since the Cretaceous. There seems to be no clear westward continuation into Moesia or the Dobrogea, which suggests that almost the entire Avalonian fragment was displaced southwards as the Istanbul Block.
Accretion history of the Armorican Terrane Assemblage: mechanisms of migration and ocean closure The Armorican Terrane Assemblage (sensu Franke 2000; Tait et al. 2000), also previously referred to as 'Peri-Gondwanan Terranes' or 'Northern Gondwana terranes' (e.g. Erdtmann & Kraft 1999), is exposed in a series of massifs across much of SW to Central Europe from Iberia to Poland. In Western and Central Europe, these terranes were accreted to Laurussia during the Late Palaeozoic. The term 'Variscan Orogeny', which has been used to describe the deformation and magmatism associated with the closure of the Rheic Ocean, its successor basins, and basins separating constituent terranes within the ATA, does not fully convey the complexity of these multiple accretions: a revised overview following intensive study of the constituent terranes in Central Europe and their accretion histories has been given by Pharaoh et al. (2006). In summary, early Devonian metamorphism and magmatism (sometimes called 'Caledonian', but historically and collectively termed Eo-Variscan elsewhere in Hercynian Europe; e.g. Faure et al. 1997; Shelley & Bossi~re 2000) was confined in the northern Bohemian Massif to isolated high-grade metamorphic rocks in the Gdry Sowie Block (GSB; Brueckner et al. 1996; O'Brien et al. 1997) and the Mtinchberg klippe (395-390 Ma; Kreuzer et al. 1989; Stosch & Lugmair 1990). It may record local tectonothermal and hence collisional activity between migrating platelets of the ATA, with subsequent exhumation. Whereas high-P metamorphism was initiated somewhat earlier in the GSB than further west, as indicated by growth of metamorphic (granulitefacies) zircon at 402 __ 0.8 Ma (O'Brien et al. 1997), subsequent late Devonian H T - M P metamorphism in the GSB is well constrained by U - P b monazite ages (van Breemen et al. 1988; Br6cker et al. 1998; Timmermann et al. 2000) and appears to be contemporary with H P - L T metamorphism along the contact zone of the Saxo-Thuringian and Teplfi-Barrandian blocks between 380 and 365 Ma. Further west, recently obtained mid- to late Devonian dates for the emplacement and metamorphism of the Lizard Peridotite and associated rocks at the southern margin of the Cornubian Massif (Sandeman et al. 2000; Nutman et al. 2001) reinforce the parallel with the Giessen-Werra-Sudharz Ocean (Franke 2000), as the latter also underwent contemporary metamorphism and deformation. The latter has been interpreted as an obducted successor basin to the Rheic Ocean.
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In the Karkonosze-Izera complex (central West Sudetes) tectonic exhumation was earlier and greater in the SE. This is shown by: (1) early kinematic indicators in mylonitic ductile shear zones (Mazur 1995; Seston et al. 2000); (2) decrease in metamorphic grade from garnet zone in the SE to chlorite zone in the NW (Baranowski et al. 1990; KachlN & Patorka 1998; Collins et al. 2000); (3) northwestward decrease of 4~ cooling ages (Marheine et al. 1999); (4) progressively later flysch sedimentation onsets towards the NW. Also, late Devonian unconformities in the central West Sudetes occur between the Ktodzko metamorphic complex and the overlying Bardo Unit (Hladil et al. 1998; Kryza et al. 2000), while Late Devonian and Carboniferous coarse-grained clastic sedimentary deposits, derived from exhumed metamorphic complexes to the east, were deposited in syntectonic basins (Aleksandrowski & Mazur 2002). Deformation and metamorphism, which started in the central West Sudetes in pre-late Devonian times (e.g. Hladil et al. 1998) continued until the Tournaisian in both the northwesternmost frontal parts of the West Sudetic orogenic wedge, where m~langes formed in the Kaczawa Complex (Collins et al. 2000), and in the metamorphic core of the complex, as in the OrlicaSnieznik area, where HP metamorphism produced eclogites. This range of dates suggests that a series of small-scale collisional events occurred, consistent with a progressive aggregation of the constituent terranes of the ATA. In the West Sudetes Carboniferous metamorphism was followed by tectonic exhumation of deeply buried crustal slices (353-350 Ma) and the superimposition of a greenschist- to lower amphibolite-facies overprint dated at 345-340 Ma). 4~ dating (325-320 Ma) suggests that metamorphism was complete by the mid- to late Carboniferous (Marheine et al. 2000), a timing supported by the age of deposition in adjacent intramontane basins. These Carboniferous events are generally considered to reflect the docking of the amalgamated ATA with the Avalonian and Bruno-Silesian margin of the growing Laurussian supercontinent. The range of dates suggests that collision was not a simple process: it probably began earlier where the accreting ATA first impinged on promontories, such as that of the Bruno-Silesian Massif, and occurred later further west. Deformation of Devono-Carboniferous sedimentary sequences on the Laurussian passive margin in the Cornubian, Rhenish and Bruno-Silesian massifs, as a result of this collision, produced the only significant late Palaeozoic deformation to affect both Avalonia and Bruno-Silesia. As the ATA approached Laurussia, subduction was south-dipping beneath its leading edge, causing the formation of an arc edifice preserved as volcanic rocks of the Mid-German Crystalline High (MGCH), with its associated oceanic back-arc basin, the Giessen-Werra-Siidharz 'ocean'. Subduction of this successor back-arc basin, which developed on the south side of the Rheic Ocean, occurred in DevonoCarboniferous time, with obduction of fragments of it, originally developed on the southern side of the ocean, eventually thrust northwards across the Rheic Suture, so that they are now preserved as ophiolitic outliers assigned to the Giessen-Werra-Stidharz or Selke Nappe (e.g. Franke 2000), north of the Rheic Suture. Thus, the MGCH marks the superimposition of both late Silurian-Devonian arc magmatism on the Avalonian margin below the south-dipping Rheic Suture, and Carboniferous age volcanism above it (Oncken 1997). Small magnetic highs seem to indicate a continuation of the volcanic centres within the MGCH eastwards into Poland as far as a point just NE of the LesznoWolsztyn High, corresponding to the location of the Moravian Line.
Eastern extremities As with Avalonia, the eastern extremities of the ATA abut the Bruno-Silesian Massif (BSM), which must have still formed a
328
J.A. WINCHESTERETAL.
promontory on the Laurussian margin at the time of ATA accretion. Without a perfect fit, the ATA presumably included crustal blocks that converged with Laurussia further east, and that might be expected to be accreted to the southern margin of the BSM. However, because the latter margin is overthrust by the Carpathian-Alpine Front, the mechanism for distribution of ATA-related blocks further east is obscured. However, rocks apparently subjected to Variscan-age metamorphism, often intruded by mid-Carboniferous post-orogenic granitoids, occur as basement inliers in the Carpathians, such as the Tatra Mts. In the western Tatra Mts, metamorphic rocks containing amphibolites with similar chemistry to those in the West Sudetes (Gaw~da et al. 2000) are cut by post-metamorphic Variscan granitoid rocks, dated by both 4~ and Rb-Sr methods at 300-330 Ma (Burchart 1968; Janak 1994). In the Romanian Carpathians, the Getic-Supragetic basement (Iancu & Berza 2004) contains similar lithologies subjected to Variscan deformation and metamorphism. Further SE, the Balkan terrane exposed in western Bulgaria (Fig. 1), and also sampled north of Sofia in the Svoge borehole, contains mid-Ordovician faunas similar to those of Bohemia and North Africa (Gutteriez-Marco et al. 2003), and typical of a cold, peri-Gondwanan environment (Haydoutov & Yanev 1997). Mid-Ordovician trilobites ( C y c l o p y g e p r i s c a ) occur in shales overlain by Ashgill diamictites, indicating that the Balkan Terrane remained attached to Gondwana in high latitudes long after Avalonia had rifted off and migrated to lower latitudes. Built upon a basement of Neoproterozoic ophiolites and Cambrian calc-alkaline volcanic rocks, the thick Palaeozoic sequence also includes Silurian argillites, Devonian clastic deposits and an unconformity above the Lower Carboniferous units. All these indicators point to an 'ATA' Gondwana affinity, with collision with Moesia during the Carboniferous. However, the presence of a late Cambrian subduction-related sequence (493 Ma, Carrigan et al. 2003) also needs explanation. Although this could be interpreted as the product of intercontinental collision, it could also be the result of a Cambrian arc-continent collision, closing the intervening oceanic back-arc basin that had been formed in the late Neoproterozoic. If so, the Balkan terrane could represent yet another portion of the NeoproterozoicCambrian supercontinent-fringing series of arcs and back-arc basins. To the SE, the Balkan Terrane is structurally juxtaposed with the Rhodope (Thracian) and Strandja terranes, which nevertheless seem to share its Palaeozoic continental affinities. Still further east, in NW Turkey, the basement to the Sakarya Zone shares a similar Palaeozoic history to blocks comprising the ATA, in that it underwent Carboniferous metamorphism, followed by intrusion of late Carboniferous post-orogenic granites (Yilmaz et al. 1997). A rupture of the ATA, similar to that experienced by Avalonia on collision with the Bruno-Silesian Promontory, might explain, in the same way, the eastward migration of displaced ATA-related blocks.
Why did Avalonia and the ATA separate from Gondwana? The composition of Palaeozoic magmatic rocks provides clues to the causes of the separation of Avalonia and the ATA from the Gondwana margin. In the northern Bohemian Massif extensive bimodal magmatism occurred in the early Ordovician, with bursts of magmatism continuing until the Devonian. Early, mainly acidic magmatism of Cambro-Ordovician age (e.g. Korytowski et al. 1993; Krrner et al. 1994; Philippe et al. 1995; Hammer et al. 1997) shows calc-alkaline chemistry, which some interpreted as evidence for an arc or active continent margin tectonic setting (e.g. 0liver et al. 1993; Krrner & Hegner 1998). Others suggested that the absence of supporting geological evidence for an arc edifice at the time suggested that chemical
characteristics of the intrusions were inherited from extensive melting of the calc-alkaline Panafrican basement (Kryza & Pin 1997; Aleksandrowski et al. 2000; Floyd et al. 2000). Subsequent, dominantly basic volcanism was associated with clastic basin-fill metasedimentary rocks, typical of magmatism associated with an extensional tectonic setting. Minor associated felsic volcanic rocks were shown by Sm-Nd systematics and their REE distribution to result from continued melting of continental crust (Fumes et al. 1994; Patofika et al. 1997; Dostal et al. 2000), whereas the compositional range of the basic rocks (e.g. Floyd et al. 1996, 2000; Winchester et al. 1995, 1998) indicated magma production resulting from the interaction of an enriched plume with both asthenospheric and sediment-contaminated lithospheric mantle sources (Floyd et al. 2000). Although the preserved volume of magmatic rocks is smaller than younger plume-influenced magmatic provinces, it has widespread correlatives in many parts of Western Europe, including the Massif Central (Briand et al. 1991, 1995) and Massif des Maures (B. Briand, pers. comm.) in France and NW Spain (e.g. Peucat et al. 1990). Floyd et al. (2000) suggested that plume-induced magmatism could also explain the amount of heat needed to melt substantial volumes of lower crust to produce the major granitoid bodies, this providing a possible mechanism for the fragmentation of the Armorican Terrane Assemblage (ATA) as it separated from Gondwana, and the repeated rifting of crustal fragments from the Gondwana margin, including Avalonia and the ATA.
Palaeozoic palaeogeographical evolution and accretions to the EEC Recent reconstructions show that the main pre-Alpine, Central European and related microcontinents formed an active continental margin (ACM) to the Pannotian supercontinent, with Avalonia adjacent to the Amazonian Craton, based on the presence of inherited 1.5 Ga 'Rondonian' ages obtained from rocks in Nova Scotia (Nance & Murphy 1994) and central England (Tucker & Pharaoh 1991). To the east (present co-ordinates) the ACM extends through the ATA (shown adjacent to the North African Craton as it lacks inherited 'Rondonian' ages) and other blocks that are thought to have separated from their peri-Gondwanan positions later, notably the basements of Italy, the Pannonian blocks, and the Tauride basement of southern Turkey. The presence of late Neoproterozoic ophiolitic fragments within this ACM (e.g. Yifgitba~ et al. 1999; Scarrow et al. 2001) attests to the obduction of successor basins and suggests that the continental margin was originally of West Pacific rather than Andean type. Shared end-Proterozoic calc-alkaline magmatism and deformation affecting all the accreted blocks records their former location along an active margin to the end-Proterozoic supercontinent Pannotia. During the Cambrian, subduction along this margin appears to have ceased or been greatly reduced, whereas during the Tremadoc, renewed subduction resulted in calc-alkaline magmatism and the formation of large back-arc basins (the Gander Arc and associated ophiolites) in the western part of the margin, now preserved in Atlantic Canada. During the Llanvirn Stage, bimodal acid-basic magmatism marks the detachment of Avalonia, possibly as more than one block (Pharaoh 1999; Banka et al. 2002; Winchester et al. 2002), marking the break-up of the Gondwana margin, and renewed arc magmatism in the 'Caradoc' Stage (Exploits and Lake District arcs) marks its rapid northward migration, narrowing the Iapetus Ocean. By this stage, a widening Rheic Ocean opened between Avalonia and the Gondwana margin, from which parts of the ATA were already starting to rift as a series of linked blocks. By the early Silurian, Avalonia had moulded itself onto the TESZ margin of Baltica, with its easternmost extremity detached and displaced eastwards along the southern margin of the new
DETACHED TERRANE FRAGMENTS IN EEC
supercontinent of Laurussia, comprising Avalonia, Baltica and Laurentia. By this time also, many blocks of the ATA, already rifted into an archipelago or related microcontinents, had separated from Gondwana, narrowing the Rheic Ocean, although the widespread occurrence of late Ordovician glacial deposits (lacking in Avalonia) indicates that significant separation from Gondwana by even the earliest blocks occurred only after the end of the Ordovician. However, the contrast between Silurian microfaunas of the French Armorican terranes and those of the Brabant Massif (Verniers 1982), suggests that the Rheic Ocean remained broad. Subduction was initiated along the southern margin of Avalonia, marking the earlier stage of volcanism in the Mid-German Crystalline High. As terranes of the ATA moved away from the Gondwana margin, the new seaway being formed was the Proto-Tethys Ocean. During the Devonian (Emsian), high-P, low-T metamorphism, recording subduction and closure of intervening seaways within the ATA, suggests that amalgamation of individual ATA terranes had begun, eventually resulting in the production of a single ATA microcontinent. Southward subduction, marked by renewed volcanism in the Mid-German Crystalline High, recorded the final stage in the approach of the now-amalgamated ATA to Laurussia, also impelled by Gondwanan convergence. Contact with the BSP, still not firmly enough attached to Baltica to prevent some displacement and relative rotation, was marked by dextral strike-slip faulting along its western margin. This was followed by the docking of most ATA blocks along the southern margin of Laurussia. Easternmost parts of the ATA were detached on collision with the BSP and displaced eastwards by sinistral faulting to form the Variscide basement seen in Carpathian inliers, the Balkan and Thracian terranes of Bulgaria, and the Sakarya and Eastern Pontide crustal blocks of N W Turkey. These investigations, and the collation of information, were supported by the EU-funded PACE (Palaeozoic Amalgamation of Central Europe) TMR Network, No. ERBFMRXCT97-0136. Part of the study is sponsored by the FWO Research Project No. G.0094.01 'Tectonics of the Early Palaeozoic basin development in NW Europe: basin analysis and magnetic fabric analysis in the Belgian Caledonides'. The contribution of T.C.P. appears with permission of the Executive Director, British Geological Survey (NERC). Particular thanks are expressed to M. C. G6nctio~lu (Turkey), I. Haydoutov and S. Yanev (Bulgaria), A. Okay (Turkey), and A. Rushton (UK), who all provided valuable additional data for inclusion.
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VECOLI, M. & SAMUELSSON,J. 2001. Quantitative evaluation of microplankton palaeobiogeography in the Ordovician-Early Silurian of the northern TESZ (Trans-European Suture Zone): implications for the timing of the Avalonia-Baltica collision. Review of Palaeobotany and Palynology, 115, 43-68. VERNIERS, J. 1982. The Silurian Chitinozoa of the Mehaigne area (Brabant Massif Belgium). Professional Paper of the Belgian Geological Survey, 1982(6), 192. VERNIERS, J., PHARAOH,T. C., ANDRI~,L., ETAL. 2002. The Cambrian to Mid-Devonian basin development and deformation history of Eastern Avalonia east of the Midlands Microcraton: new data and a review. In: WINCHESTER, J. A., PHARAOH, T. C. & VERNIERS, J. (eds) Palaeozoic Amalgamation of Central Europe. Geological Society London, Special Publications, 201, 47-93. WILLIAMSON,J. P., PHARAOH,T. C., BANKA, D., THYBO, H., LAIGLE,M. & LEE, M. K. 2002. Potential field modelling of the BalticaAvalonia (Thor-Tomquist) suture beneath the southern North Sea. Tectonophysics, 360, 47-60. WINCHESTER, J. A., FLOYD, P. A., CHOCYK, M., HORBOWY, K. & KOZDROJ, W. 1995. Geochemistry and tectonic environment of Ordovician meta-igneous rocks in the Rudawy Janowickie Complex, SW Poland. Journal of the Geological Society, London, 152, 105-115. WINCHESTER, J. A., FLOYD,P. A., AWDANKIEWICZ,M., PIASECKI,M. A. J., AWDANKIEWICZ,H., GUNIA, P. & GLIWICZ, T. 1998. Geochemistry and tectonic significance of metabasic suites in the G6ry Sowie Block, SW Poland. Journal of the Geological Society, London, 155, 155-164. WINCHESTER, J. A. & PACE TMR NETWORK 2002. Palaeozoic Amalgamation of Central Europe: new results from recent geological and geophysical investigations. Tectonophysics, 360, 5-22. YANEV, S., GONCUOGLU,M. C., GEDIK, I., E T AL. 2006. Stratigraphy, correlations and palaeogeography of Palaeozoic terranes in Bulgaria and NW Turkey: a review of recent data. In: ROBERTSON, A. H. F., MOUNTRAKIS, D. & BRUN, J.-P. (eds) Tectonic Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 260, 51-67. YIdlTBA~, E., ELMAS, A. & YILMAZ, Y. 1999. Pre-Cenozoic tectono-stratigraphic components of the Western Pontides and their geological evolution. Geological Journal, 34, 55-74. YILMAZ, Y., TOYSI)Z,O., YIGITBA$,E., GEN(~,S. C. & ~ENGOR,A. M. C. 1997. Geology and tectonic evolution of the Pontides. In: ROBINSON, A. G. (ed.). Regional and Petroleum Geology of the Black Sea and Surrounding Region. American Association of Petroleum Geologists, Memoirs, 68, 183-226.
The Variscan orogen in Central Europe: construction and collapse WOLFGANG FRANKE Geologisch-Paliiontologisches Institut, Johann Wolfgang Goethe-Universitiit, Senckenberg-Anlage 32, Frankfurt, Germany (e-mail:
[email protected])
Abstract: On the basis of a brief survey of the subdivision and evolution of the Variscides, this paper addresses controversial issues relating to the plate kinematic assembly and the 'collapse' of the orogen. A widespread phase of Devonian extension and basaltic magmatism is at variance with overall convergence. This episode either reflects subduction of the Rheic mid-ocean ridge, or else relates to a set of mantle plumes that also produced the Dniepr-Donets aulacogen. Another controversy regards the position of Gondwana in Devonian and Early Carboniferous time. Contrary to recent proposals of a wide Palaeotethys ocean, biogeographical and palaeomagnetic data suggest, until the Late Carboniferous, a Pangaea B model with Gondwanajuxtaposed against Southern Europe. Contrary to the concept of Late Carboniferous-Permian 'collapse' of a central Variscan high plateau, major crustal thickening occurred only in relatively narrow belts, and parts of the central Variscides were close to sea level from the Late Devonian onwards. Collision occurred in a hightemperature regime from c. 350-340 Ma onwards. Heating by several independent mechanisms effected the reduction of orogenic roots by buoyant rise and lateral spreading of thermally softened crust. However, major flysch wedges reflect the importance of erosion and uplift. Late Carboniferous-Permian magmatism and extension associated with strike-slip zones affected a largely equilibrated crust. These events probably relate to the westward displacement of Gondwana and the opening of the Palaeotethys embayment (Pangaea B to Pangaea A).
It is generally agreed that the Variscan crust is a collage of Gondwana-derived microplates (Avalonia, Armorican Terrane Assemblage; ATA), which were sequentially accreted to Baltica and eventually caught up in the collision of Gondwana and the 'Old Red Sandstone Continent' (Laurussia plus Avalonia). Late Devonian and Carboniferous subduction and collision created a large and heterogeneous orogen, with two zones of subduction on the northern flank and one on the southern flank of the belt (Figs 1 and 2). Shortening of continental crust involved in the Variscan collisions (Fig. 3) amounts to at least 800 km. Much higher values are probable. As discussed by Franke et al. (1995) and Franke (2000), the Variscan sutures have been overprinted by important dextral strike-slip movements that are difficult to constrain. The evolution of major parts of, or the entire Variscides has repeatedly been summarized (see, e.g. Martin & Eder 1983; Matte 1986, 1991; Matte et al. 1990; Franke et al. 1995; Franke 2000). Pharaoh et al. (2006) has provided a summary that includes the broader plate-tectonic context. However, the plate kinematic evolution and its geodynamic background as well as build-up and destruction of the orogen still present many unsolved questions. The present paper attempts a brief review of the main facts and highlights major open problems encountered in the German segment of the orogen, with references to some neighbouring areas. Main issues concern Devonian extension and its geodynamic causes, the existence of a Palaeotethys ocean, and the processes that destroyed the orogen. For a more detailed assessment and a survey of earlier literature, the reader is referred to Franke (2000) and, for eastern parts of the Bohemian Massif, Franke & Zelainiewicz (2000, 2002). Correlation of exogenic and endogenic events is based upon the time scales of Gradstein et al. (2004) and, for the Carboniferous, Menning et al. (2000). Plate kinematic scenario Major oceanic basins and constraints on their closure (Figs 2 and 3) Rheic ocean. Avalonia had rifted off from Gondwana during the Ordovician, opening the Rheic ocean in its wake (e.g. Tait et al. 2000). Avalonia made contact with Baltica in Late Ordovician time, thus producing a narrow tectonic belt known only from drillholes (see the review by Pharaoh et al. 2006). In Silurian-Early
Devonian times, the Rheic ocean was closed by intra-oceanic subduction, giving rise to an island arc now preserved in the Mid-German Crystalline High and at the southern margin of the Rhenish Massif (Rheno-Hercynian belt). Felsic members of the arc have been dated to the latest Ordovician to Early Devonian (444 _+ 22 to 398 _+ 3 Ma; U/Pb, Pb/Pb zircon and a few Rb/Sr whole-rock ages, see Franke 2000). By the earliest Devonian, the ATA must have been juxtaposed against Avalonia, as both microplates share the same Lochkovian non-marine fish (around 415 Ma, Young 1990; see biogeographical summary by McKerrow et al. 2000). Silurian and Devonian sedimentary sequences of the Rheno-Hercynian belt do not show any evidence of deformation and synorogenic clastic sedimentation during the relevant time span. Instead, important Gedinnian to Siegenian subsidence and sedimentation suggest crustal extension, probably effected by subduction toward the north and resulting back-arc spreading. Rheno-Hercynian narrow ocean. Shortly after, in Emsian time, a new spreading episode started to open the Rheno-Hercynian (Lizard-Giel3en-Harz) basin. The age of the oceanic crust is constrained by the oldest pelagic sediments overlying the pillow lavas (Emsian and Eifelian near Giegen, Birkelbach et al. 1988) and by a U - P b zircon age of 397 +__2 Ma from the Lizard allochthon in SW England (Clark et al. 1998), which again falls into the Emsian or early Eifelian, thus matching the biostratigraphic evidence in Germany. During Variscan collision, nappes in SW England and Germany transported oceanic fragments over the Avalonian foreland. At their base, these thrust sheets contain fragmented sedimentary sequences of Ordovician to Early Devonian age with Armorican faunas (see discussion by Franke & Engel 1982; Franke & Oncken 1995; Plusquellec & Jahnke 1999). Hence, opening of the Rheno-Hercynian ocean must have sprit off a fragment from the ATA adjacent to the south, which was left stranded on the N W shore of the nascent Rheno-Hercynian basin. Because of this Rheno-Hercynian reworking, the Rheic suture does not correspond exactly to the present-day fault zone at the southern margin of the Rhenish Massif, but is contained within the Northern Phyllite Zone, a narrow belt of Variscan pressuredominated metamorphic rocks at the southern margin of the Rhenish Massif and Harz Mts. (Anderle et al. 1995; Franke 2000). Rheno-Hercynian extension is reflected in three volcanic episodes. Early Devonian calc-alkaline rhyolites probably relate to incipient rifting (Jones & Floyd 2000). Younger, intraplate basaltic volcanism occurs in two peaks at about the
From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 333-343. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. Structural map of the European Variscides (Permian plate configuration), with foreland basins marking the major sutures (yellow, Rheno-Hercynian-Moravo-Silesian; orange, Saxo-Thuringian; blue, retro-arc basin of the Moldanubian zone; brown, Mediterranean-Alpine, resulting from the Massif Central-Moldanubian collision). After Franke (2000).
Fig. 2. Plate kinematic model for the assembly of minor and major plates in the German segment of the Variscides. North is to the left. After Franke (2000).
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Fig. 3. Structural map and diagrammatictectonic cross-section of the German Variscides. (Note the distinctionbetween terranes (colours) and tectonic zones (black & white lettering).) Section is combined from a northwestern part (RheinischesSchiefergebirge-Spessart)and a southeastern part (SW part of the Bohemian Massif). KTB, site of deep continentaldrilling (KontinentalesTiefbohr-Programm). After Franke et al. (2004).
Givetian-Frasnian boundary and in the Tournaisian to early Vis~an (see summaries by Floyd 1995; Nesbor 2004) These volcanic episodes are also recorded in lateral equivalents of the Rheno-Hercynian in Germany, such as the Moravo-Silesian belt (Dvo}fik 1995) and SW England (see Holder & Leveridge 1986, and references therein). The Carboniferous volcanic episode is also represented in the Pyrite Belt of southern Portugal, with a predominance of felsic lavas (Oliveira & Quesada 1998; Boulter et al. 2001). Because the intra-plate basalts post-date the early Devonian onset of oceanic spreading, and the passive, northern margin of the Rheno-Hercynian ocean is clearly non-volcanic (Franke 2000), they require a separate geodynamic cause. The Rheno-Hercynian ocean is not detectable in the palaeomagnetic and biogeographical records. This is understandable, as flysch greywackes deposited on the oceanic crust suggest that subduction was already active in mid-Frasnian time (c. 380 Ma), which leaves only c. 30 Ma (Emsian to Givetian) for the drift stage. The Frasnian to Namurian flysch sediments are derived from the active, southern margin of the Rheno-Hercynian basin (Mid-German Crystalline High, MGCH; see Kopp & Bankwitz (2003) for the latest review of regional geology). The Crystalline High evolved from a north Armorican microplate (Franconia), which is documented only in Neoproterozoic detrital micas from Late Devonian greywacke turbidites (Huckriede et al. 2004). Magmatic activity in the arc of the MGCR is detectable from c. 360 Ma onwards. Retrogressed eclogites in the eastern Odenwald have been dated at 357 _+ 7 and 353 + 11 Ma ( L u - H f garnet-whole
rock, Scherer et al. 2002). Collision is documented by the crossover of greywacke turbidites onto the Avalonian foreland from the Devonian-Carboniferous boundary onwards (see Franke 2000, p. 50). During the Early Carboniferous to Namurian B, the front of synorogenic clastic sedimentation migrated across the foreland (Kulick 1960; Engel & Franke 1983; Franke & Engel 1986). From the Namurian C to the Westphalian C, sedimentation continued in a paralic molasse basin with coal seams, which can be traced from the Ruhr district of the northern Rhenish Massif westwards through northern Belgium as far as south Wales, and eastwards in Silesia, on the SE flank of the Bohemian Arc. S a x o - T h u r i n g i a n n a r r o w ocean. The Saxo-Thuringian basin originated from Cambro-Ordovician tiffing, which separated Bohemia from the north Armorican microplates of Saxo-Thuringia and Franconia. The Vesser Rift (Fig. 2; Kemnitz et al. 2002) between the latter terranes probably failed and did not evolve into a separate orogenic belt. Prolonged Saxo-Thuringian extension is documented in episodes of basaltic intra-plate volcanism in Silurian, Early Devonian and early Frasnian time (see summary by Falk et al. 1995). Southward subduction of Saxo-Thuringian crust commenced no later than c. 400 Ma (a summary of metamorphic events has been given by Franke et al. 1995; Scherer et al. 2002) and produced eclogites now preserved at the deformed northwestern margin of the Bohemia terrane (Tepl~-Barrandian unit) and in allochthons emplaced on the Saxo-Thuringian foreland (Franke 1984a,b).
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Collisional closure of the basin is recorded in widespread medium-pressure metamorphism around 380 Ma; in rocks on both sides of the suture and is also constrained by the onlap of early Famennian flysch greywackes on the foreland (c. 375 Ma; Sch~ifer et al. 1997). Massif Central-Moldanubian (MCM) narrow ocean. Oceanic separ-
ation between the ATA and mainland Gondwana is inferred from allochthons in the Moldanubian belt of the Bohemian Massif, which contain rocks that have undergone metamorphism at pressures up to 4 GPa derived from Palaeozoic ultramafic-mafic as well as continental protoliths (e.g. Becker & Altherr 1992; O'Brien 2000; VrLqa & Fryda 2003). Tectonic transport was directed generally southwards. As in the Rheno-Hercynian and Saxo-Thuringian cases, this southern ocean is not documented in the palaeomagnetic and biogeographical records (McKerrow et al. 2000; Robardet 2003). However, the presence of mantle rocks in the (ultra-)high-pressure metamorphic assemblage of the Gfrhl Moldanubian (Medaris et al. 1995) requires a zone of major crustal extension or narrow ocean, which guided continental subduction. Several lines of evidence suggest that closure of the MCM ocean occurred in latest Mid-Devonian to Late Devonian time. Late Givetian (c. 380 Ma) clastic sediments in the Barrandian syncline near Prague (Chlup~i6 1993) may be taken as a first sedimentary signal of orogeny preserved in the upper plate. Metagranites at the southern margin of the Tepl~i-Barrandian block were intruded at c. 370 Ma (Ko~ler et al. 1993; Ko~ler & Farrow 1994). In an eastern part of the Tepl~i-Barrandian unit, now concealed under Cretaceous cover (Fig. 3), limestones of late Famennian age unconformably overlie folded and cleaved Palaeozoic rocks, which indicates deformation prior to c. 360 Ma (Chlup~i~ 1994). This post-tectonic sedimentary cover extends into the Bardo basin of the West Sudetes, where limestones of Late Devonian age unconformably overlie a greenschist-grade basement (Ktodzko unit), whose protoliths include Neoproterozoic and early Ordovician felsic magmatic rocks (Mazur et al. 2003) as well as Givetian carbonates (Hladil et al. 1999); see also discussion by Franke & Zelainiewicz (2000, 2002). The suture zone at the southern margin of the Tepl~i-Barrandian block is sealed by the largely undeformed Central Bohemian batholith, which was intruded between 354 and 337 Ma (Drrr et al. 1997; Holub et al. 1997; Janousek & Gerdes 2003). In the French Massif Central, a western equivalent of the Moldanubian Zone, collision likewise occurred in Mid- to Late Devonian time (before c. 380 Ma; e.g. Lardeaux et al. 2001; Cartier & Faure 2004). High-pressure granulites from the Moldanubian allochthon further south (Gfrhl unit) have consistently yielded U-Pb zircon ages around 340 Ma (e.g. Krrner et al. 2000). Overthrusting of the Gfrhl unit over the less allochthonous Drosendorf unit led to widespread metamorphism around 335 Ma (see compilation by Franke 2000). In the SE part of the Moldanubian unit, post-metamorphic durbachites with U-Pb zircon ages between 338 and 335 Ma reflect HP melting in mantle or slab rocks (Kotkowi et al. 2003). Tectonometamorphic and magmatic events in the Moldanubian unit are difficult to interpret, as there is evidence for two subduction-collision events: the earlier one (>340 Ma) probably records collision between Bohemia and Gondwana (or some other Gondwana-derived fragment), with the suture extending westwards into the Massif Central (Figs 1 and 2). Later, the Moldanubian, Saxo-Thuringian and Rheno-Hercynian belts were dissected by NW-trending dextral shear zones (the Elbe and Intrasudetic fault zones) and rotated clockwise to form the Bohemian Arc. Shortly after, the rotated tectonic belts were truncated by dextral transpression along the Moldanubian Thrust, and juxtaposed against the Moravo-Silesian block (Figs 1 and 3; Franke & Zelalniewicz 2000, 2002; Gayer & Schulmann 2000). This latter process is documented by Vis~an flysch sediments deposited on the Moravo-Silesian foreland from the early Vis~an
(c. 340 Ma) to the early Namurian, which was followed by Late Carboniferous fluvio-lacustrine molasse with coal seams. The latest Visran Moravice Formation of the Moravo-Silesian flysch contains granulite pebbles derived from the Moldanubian allochthon and transported across the Moldanubian Thrust (Hartley & Otava 2001). This indicates that, by c. 325 Ma, the transpressional event must have been completed. The Moldanubian belt is conventionally correlated with the southern Black Forest and Vosges in SW Germany and, beyond, the French Massif Central (e.g. Matte 1986, 1991). At the southern margin of the Massif Central, in the Mouthoumet Massif and in the Pyrenees, synorogenic clastic sediments record the southwestward advance (in present-day coordinates) of the orogenic front in late Vis~an to Namurian time (Engel 1984; Franke & Engel 1986). The foreland is taken to represent Gondwana, as thick sequences of Cambro-Ordovician shelf sediments cannot be derived from a microplate, but require a large catchment area. The Variscan basement fragments dispersed over the Alps and the Mediterranean realm cannot be treated in this paper. However, it is important to note that the Carboniferous flysch of southern France is generally correlated with the southwarddriving flysch wedge exposed in the Southern Alps (see Fig. 1 and, e.g. Franke & Engel 1986; for the Carnic Alps and Karawanken Mts, see L~iufer et al. 2001). Therefore, it can be expected that evolution of the Variscan basement units now incorporated in the Alps was similar to that of France. The south Alpine Hochwipfel Flysch is unconformably overlain by shallow-marine deposits of Late Carboniferous to Permian age (Krainer 1993; Sch6nlaub & Histon 2000), which probably indicate incipient rifting at the tip of the westward-propagating Palaeotethys. O p e n questions Devonian plate divergence and basaltic magmatism. It is difficult to
understand why Devonian sea-floor spreading should occur within an assembly of plates that, during this time interval, underwent large-scale convergence. Back-arc extension (as proposed by Ziegler; e.g. Ziegler & D~zes 2006) can account for only the late Silurian-earliest Devonian part of Rheno-Hercynian extension in areas to the north of the intra-oceanic arc (Fig. 2). However, formation of Rheno-Hercynian ocean crust occurred within the northern part of the ATA and would, therefore, require southward subduction under the northern margin of the ATA (for which there is no evidence), or else subduction of the Saxo-Thuringian narrow ocean towards the north (which, in fact, was towards the south under Bohemia). Instead, it might be speculated that the narrow Rheno-Hercynian ocean was formed, when the northward moving ATA overrode the mid-ocean ridge of the Rheic ocean, much like the Bay of California is being opened, today, because the North American plate overrides the East Pacific Ridge. A back-arc model is feasible only for the early Frasnian intra-plate volcanism in the Saxo-Thuringian basin adjacent to the south, which might have been caused by southward subduction of Rheno-Hercynian ocean crust under the nascent Mid-German Crystalline High. A separate explanation is required for the Givetian to Frasnian and Early Carboniferous intra-plate basalts in the RhenoHercynian autochthon of the Rhenish Massif and Harz Mts (and equivalents from Portugal to Moravia). The geometry of the subduction zones active during this time interval (Fig. 2) precludes back-arc spreading. It is interesting to note that the majority of Devonian alkaline magmatic rocks in the Kola Peninsula were emplaced between 382 and 362 Ma (Sindern et al. 2003). Devonian to Early Carboniferous extension and magmatism are also important within the East European Craton (see Stephenson et al. 2006). It may be speculated that all these magmatic provinces represent a large-scale cluster of mantle plumes, whose activities were independent of the convection systems driving Variscan plate tectonics. Late Devonian-Early Carboniferous extension and magmatism in the Brevenne unit of the French
VARISCAN OROGEN IN CENTRALEUROPE Massif Central and Early Frasnian basaltic magmatism in the Saxo-Thuringian belt (see above) possibly record back-arc extension, but could equally well be part of the plume scenario. Position o f Gondwana and evolution of the Palaeotethys ocean. The
position of Gondwana during the Devonian and Early Carboniferous is still controversial. Most palaeomagnetic scenarios (e.g. Tait et al. 2000; Cocks & Torsvik 2002) propose an oceanic separation between Gondwana and the ATA, which corresponds to the Massif Central-Moldanubian ocean. This ocean is inferred to have opened from the Early Devonian and to have widened during the Carboniferous. A detailed assessment of the biogeographical and palaeomagnetic findings involved would go beyond the scope of this paper. It should be pointed out, however, that recent biogeographical reviews by McKerrow et al. (2000) and Robardet (2003) do not reveal indications of the Palaeotethys. Differences in the floras and faunas in post-Early Devonian times are attributed, by those workers, to climatic belts. In addition, the U - P b signatures of zircons from Ordovician to Devonian clastic sediments in Iberia suggest derivation from the West African Craton and the surrounding Pan-African belts (Martfnez-Catal~in et al. 2004). Lastly, the tectonic and sedimentary records of Variscan collision in southern France, the Southern /kips, and the Moldanubian part of the Bohemian Massif consistenfly document deposition of flysch and emplacement of thrust sheets on continental forelands. In the best-preserved section (southern France), the pre-flysch sedimentary and faunal records clearly suggest that this foreland was part of Gondwana (Robardet 2003). As noted above, the oldest marine ingression attributable to the opening of Palaeotethys occurs in the Late Carboniferous of the Carnic Alps. This is consistent with the findings of Muttoni et al. (2003), who have documented that the change from Pangaea 'B' (with Gondwana juxtaposed to Europe) to Pangaea 'A' (with the Palaeotethys to the south of Europe) occurred not before the Permian, and was accommodated by large-scale dextral shear zones (as proposed already by Arthaud & Matte 1977; Matte 1986). These findings rule out the concept of an oceanic subduction zone dipping to the NW under the ATA in Carboniferous time (as shown, e.g. by Cocks & Torsvik 2002), for which there is no evidence in any part of the southern Variscides. Likewise, the southern Variscides do not show any indications of Late Devonian-Early Carboniferous extension and magmatism. To avoid these difficulties, Tait et al. (2000) have proposed that the continental foreland found in the southern Variscides (southern France, Southern Alps) does not represent Gondwana mainland, but another Gondwana-derived microplate. This explanation, however, just shifts the problem to unknown areas further south, so that evidence of Late Devonian-Early Carboniferous rifting and Carboniferous northward subduction remains elusive.
Destruction of the orogen: where, when and why? Destruction of an orogen is brought about by reduction of the thickened crust either by plate boundary forces leading to lithospheric extension or else by buoyancy forces. In the latter case, crustal roots may be reduced by erosional or tectonic removal of orogenic topography. Alternatively, topography may be reduced by gravitational spreading of hot, low-viscosity lower crust. In reality, these processes will often work together. Their evolution in the Central European Variscides is assessed in the following paragraphs. P a l a e o - t o p o g r a p h y , vertical crustal m o v e m e n t s a n d thermal regime Late Devonian and Tournisian events ( 3 6 0 - 3 4 0 Ma). During the Late
Devonian, subduction of oceanic and continental materials
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occurred both on the northwestern margin (Saxo-Thuringian belt) and southeastern margin (Moldanubian belt) of Bohemia. Major erosion and some uplift is documented at the active, northwestern margin of Bohemia, where 400 Ma high-pressure and 380 Ma medium-pressure metamorphic rocks were already being eroded in Fammenian time and deposited in the marine Saxo-Thuringian foreland basin (Sch~ifer et al. 1997). Clues to the palaeo-topography have survived only in eastern parts of the Tepl~i-Barrandian block, where the transgression of Late Devonian to Toumaisian marine sediments on deformed Palaeozoic rocks indicates zero elevation. Also, the preservation of very low-grade sedimentary rocks and moderate tectonic shortening argue against major crustal thickening and uplift. These findings clearly rule out the concept of a Tepl~i-Barrandian high plateau proposed by Zulauf (1997, 2002). The absence of siliciclastic debris from these deposits argues against major elevation also in the neighbouring regions in the time between c. 360 and 340 Ma. Coarse-grained clastic sediments of Late Devonian to Early Carboniferous age do occur in the SwiCbodzice pull-apart basin of the West Sudetes (PorCbski 1990), but late Frasnian and Famennian intercalations of marine mudstones and limestones indicate that the sediments of the intra-orogenic basin were deposited close to sea level. These findings demonstrate that, < 20 Ma after the closure of the Saxo-Thuringian and Massif Central-Moldanubian oceans, parts of the central Variscides were inundated by the sea, and orogenic topography existed only in marginal parts of Bohemia. This clearly precludes an areally extensive high plateau in the area. Similar considerations apply to the Morvan in the northern part of the French Massif Central, where marine shales have been dated as Tournaisian (Weyer 1976), and to the Beaujolais and Brevenne units further south, which contain Devonian to Early Carboniferous marine sedimentary and volcanic rocks of very low to low metamorphic grade (Leloix et al. 1999; Lardeaux et al. 2001). A relatively cool thermal regime is indicated by the observation that most of the metamorphic rocks of the Bohemian Massif dated at c. 400-380 Ma were formed in, or else exhumed through, the amphibolite facies. H P - H T metamorphism is restricted to one locality in the Saxo-Thuringian region of NW Bavaria (Kleinschrodt & Gayk 1999) and to the Grry Sowie of the West Sudetes (Zelalniewicz 1990; O'Brien et al. 1997; Timmermann et al. 2000). The 340 Ma event. Areas surrounding the Tepl~i-Barrandian unit (i.e. the Bohemia microplate and rocks accreted to it) are characterized by high-temperature metamorphic rocks and granitoids dated at c. 340 Ma. The Saxonian Granulites of the SaxoThuringian belt were equilibrated at c. 22 kbars and 1050 ~ (Rrtzler et al. 2004), the highest metamorphic temperatures hitherto documented in the Variscan belt. Ultrahigh pressures are documented in metamorphic diamonds in continental rocks of the Erzgebirge dated at c. 340 Ma (Massonne 2001, 2003; Massonne et al. 2001). Shortly after their formation, granulites and eclogites were emplaced in the continental crust of the foreland (e.g. Reinhardt & Kleemann 1994). Emplacement was probably driven by hydraulic forces and occurred under the floor of the Saxo-Thuringian foreland basin (DEKORP & Orogenic Processes Working Groups 1999; Franke & Stein 2000; Henk 2000), a process requiring low viscosity (Zulauf et al. 2002a). In the southeastern, internal part of the Saxo-Thuringian belt, the exhumed high-pressure rocks were subsequently involved in the accretion of the foreland (Erzgebirge: Franke & Stein 2000; Konopfisek et al. 2001). Eclogites in the Orlica-SnieZnik unit of the West Sudetes (Fig. 3) probably represent an eastern continuation of the Erzgebirge (see discussion by Franke & Zelainiewicz 2000). On the southern (Moldanubian) flank, high-pressure granulites are widespread in the allochthonous Gfrhl unit. These granulites have consistently yielded U - P b zircon ages around 340 Ma
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(compilation by Franke 2000; Kr6ner et al. 2000). However, deformed granites within allochthonous Bohemian granulite gneisses show U-Pb zircon minimum ages of 354 Ma (Svojtka et al. 2002). Zircon growth possibly occurred during decompression melting in felsic granulites, so that the time of deepest burial might be older (Finger et al. 1996; Roberts & Finger 1997). This is compatible with the findings of Kr6ner et al. (2000), who have demonstrated that zircons from granulites with U-Pb ages of c. 340 Ma were repeatedly formed at various stages of decompression. Zulauf (1997, 2002b) and Zulauf et al. (2002) have shown that the Teplfi-Barrandian is surrounded, today, by rocks whose metamorphic grade implies subsidence of the Tepl~i-Barrandian (or uplift of the surrounding rocks) by c. 10-15 km. Overthrusting of the Gf6hl unit over the less allochthonous Drosendorf unit led to widespread metamorphism around 335 Ma (see compilation by Franke 2000). In the southeastern part of the Moldanubian unit, post-metamorphic durbachites with U-Pb zircon ages between 338 and 335 Ma reflect HP melting in mantle or slab rocks (Kotkovfi et al. 2003), probably belonging to the underthrust Moravo-Silesian belt. 3 4 0 - 3 2 0 Ma. Orogenic uplift and erosion from 340 to 320 Ma is
documented by flysch sediments in the foreland basins of the Rheno-Hercynian, Saxo-Thuringian and Moravo-Silesian belts. Clastic sediments were derived from the Mid-German Crystalline High, the northwestern margin of Bohmia, and from Moldanubian sources (Fig. 1). The Variscan topography, during this time span, was characterized by an alternation of foreland basins and flyschproducing collisional uplifts. Whereas sedimentation in the Rheno-Hercynian and Moravo-Silesian basins continued into Late Carboniferous coal-bearing molasse deposits, sedimentation in the Saxo-Thuringian was terminated by folding around 330 Ma. A short time later, there is again evidence of low elevations. A eustatic sea-level rise occurred in the Goniatites crenistria zone of the traditional European Culm zonation (Go oL of the classical goniatite stratigraphy, Late Asbian; see Herbig 1998), which corresponds to an isotopic age of c. 327 Ma (Menning et al. 2000). Marine sediments of this age were deposited on folded Cambrian rocks in the the northwestern part of the Saxo-Thuringian basin (Doberlug-Kirchhain, Vis~an 3b, Weyer 1965) and in the strike-slip, related Intra-Sudetic Basin (Fig. 3: SW of the Gdry Sowie; Zakowa 1963; Herbig 1998; Tumau et al. 2002). In both these basins, the marine beds represent an early phase of sedimentation, so that the orogenic topography must have been low from the beginning of basin evolution. During the time interval between 340 and 320 Ma, there is no evidence of pressure-dominated metamorphism, and P - T conditions suggest collisional stacking and heating. The Moldanubian zone contains large volumes of granites intruded between c. 335 and 325 Ma (see references in Franke et al. 2000; discussion by Gerdes et al. 2002, 2003). A narrow, NW-trending belt along the SW margin of the Bohemian Massif cuts across the tectonic zonation. It is characterized, between c. 325 and 320 Ma, by an especially high-temperature regime. Low-pressure-high-temperature metamorphism with anatexis (Tanner & Behrmann 1995; Tanner 1999) dated to a narrow interval of 327-320 Ma (see reviews by Franke 2000; Kalt et al. 2000) was immediately followed by the intrusion of post-tectonic granitoids (e.g. Siebel et al. 2003; Chen& Siebel 2004). This transverse zone extends northwestwards into the very low grade rocks of the Saxo-Thuringian foreland (Kosakowski et al. 1999). Younger granites (315-290 Ma) occur in this SW Bohemian Transverse Zone, but also in the Fichtelgebirge-Erzgebirge antiform of the southeastern Saxo-Thuringian zone and in the Moldanubian zone. Evolution in SW Germany, France and Iberia. A detailed assessment
of metamorphism and granitoid magmatism in other parts of the
Variscan intemides would go beyond the scope of this paper. It should be noted, however, that the time interval of granitoid intrusion observed in the Bohemian Massif is the same as that in Iberia (352-297 Ma, maximum at 335-305 Ma; Montero et al. 2004). Intense metamorphism and granite emplacement at about 340 Ma have also been recorded from the Black Forest, the Vosges and the Massif Central (Costa 1992; Boutin et al. 1995). In the intramontane basins of the French Massif Central, there are no marine sediments, which would suggest moderate elevations also in neighbouring regions. Becq-Giraudon & Van den Driessche (1994) and Becq-Giraudon et al. (1996) even claimed to have found petrographic, sedimentological and palaeobotanical evidence of sediments deposited >5000 m above sea level. However, their sedimentological findings are equivocal. Also, the floras of the Permo-Carboniferous basins in the Massif Central do not reveal a cold environment, but warm and humid conditions (H. Kerp, Mtinster Univ., pers. comm.).
Geodynamic model
Whereas the closure of the Rheno-Hercynian basin conforms to a classical model of subduction-collision, the tectonothermal evolution of the internal Variscides (Bohemian Massif, Saxo-Thuringian and Moldanubian belts) is problematic. In Mid- to Late D e v o n i a n time, the Saxo-Thuringian and Moldanubian narrow oceans were closed by subduction from the SE and the NW under the Bohemian microplate (now largely represented by the Teplfi-Barrandian tectonic unit). Eclogites in the Saxo-Thuringian belt were formed and obducted in a mediumtemperature regime. Obduction may be explained by buoyant rise of subducted continental material according to the model of Chemenda et al. (1995; see Franke & Stein 2000). Late Devonian subduction on the Moldanubian flank cannot be excluded, but isotopic evidence has been obliterated by later high-temperature metamorphism. Crustal thickening, resulting in uplift and erosion, is documented only for the northwestern margin of Bohemia (the active margin of the Saxo-Thuringian basin). The central part of Bohemia (Teplfi-Barrandian) was only moderately thickened and had already been inundated by the sea by Late Devonian time. B e t w e e n c. 350 a n d 340 Ma, granitoids were emplaced along the southeastern and western flanks of the Teplfi-Barrandian block and remained largely unaffected by later ductile deformation. During the same time, or shortly thereafter (c. 340 Ma), large volumes of low-viscosity, high-pressure granulites and some eclogites rose on both flanks of the Bohemian 'median massif' (Fig. 4). The buoyancy of felsic material will certainly have contributed to uplift (Reinhardt & Kleemann 1994; Gerya et al. 2002a,b; see also Lardeaux et al. (2001) for a similar situation in the French Massif Central). However, expulsion by hydraulic or compressional forces probably also played a major role in their emplacement (Franke & Stein 2000; Henk 2000; Stfpskfi et al. 2004) Ultrahigh pressures in metamorphosed continental rocks on both flanks of the Teplfi-Barrandian suggest that at least part of the heat was derived from contact with the asthenosphere. In addition, crustal shortening on the NW and SE flanks of Bohemia amounts to at least 500 km. This implies subduction of the same amount of lithospheric mantle, which cannot have been accommodated under the narrow 'median massif' of the Tepl~-Barrandian block (Franke 2000). Loss of parts of the subducted lithospheric mantle slabs and subsequent rise of hot asthenosphere probably added to the high-temperature regime. Because the density contrast between lithospheric and asthenospheric mantle is rather low (c. 0.05 g cm-3; e.g. Grow & Bowin 1975), neither subduction nor break-off of the lithospheric mantle slabs can be expected to have important isostatic consequences. This is in accord with Late Devonian to Tournaisian shallow-marine environments in the West Sudetes.
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Fig. 4. Tectonic model (not to scale) of continental subduction and subsequent exhumation at the northwestern and southeastern margins of Bohemia, around 340 Ma. Downward thinning of subducted lithospheric slabs is intended to indicate transition into oceanic lithosphere (already subducted).
B e t w e e n 340 and 325 Ma, accretion and thickening of Moldanubian crust propagated toward the SE, in a medium- to low-pressure metamorphic regime with widespread migmatization and intrusion of large volumes of granitoids (not depicted in Fig. 4). Both heating by radiogenic decay and transfer of mantle heat through a thinned mantle lithosphere may be responsible for the thermal environment and have caused extensional spreading of thickened crust. However, shedding of flysch sediments into the MoravoSilesian foreland basin indicates that crustal thinning was also effected by erosion and uplift. These considerations also apply to the active, southeastern margins of the Rheno-Hercynian and Saxo-Thuringian basins. In any case, the late Asbian (c. 328 Ma) marine sediments in intramontane Saxo-Thuringian basins preclude, also for this time interval, the existence of a central Variscan 'Tibetan' plateau. A r o u n d 325 Ma, low-pressure metamorphism and granitoid intrusion in the SW Bohemian transverse zone cut across the collisional zonation, from the Moldanubian belt in the SE to the Saxo-Thuringian foreland in the NW. This argues against causes acting 'along-strike', such as crustal thickening or delamination of mantle lithosphere. Franke et al. (1995), Franke (2000) and Kalt et al. (2000) have proposed advective heating by melts, probably triggered by processes in the asthenospheric mantle independent of the mechanics of the orogen. Younger granitoids ( 3 2 0 - 2 9 0 Ma) are widespread in Western and Central Europe. This magmatic pulse was associated with re-equilibration of the Moho (e.g. Ziegler et al. 2004). In particular, the youngest plutonic and volcanic rocks (_ 100 000 km 3, Neumann et al. 2004). A
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possible explanation for this is the presence of an underlying thermal anomaly (i.e. a mantle plume) below the European lithosphere in Permian times, which in turn could also explain the observed widespread rifting and magmatism. However, for various reasons (see discussion by Pedersen & van der Beek 1994; Pascal et al. 2004) this hypothesis is questionable. Melting modelling of the Oslo Graben was carried out by Ro & Faleide (1992) and Pedersen & van der Beck (1994). From a model in which crust and lithospheric mantle are equally stretched, Ro & Faleide (1992) argued for the mantle plume hypothesis. In contrast, Pedersen & van der Beck (1994) showed that the volumes of melts of the Oslo Rift can be accounted for by differential stretching between crust and lithospheric mantle (i.e. the lithospheric mantle is more stretched than the crust) and reduced melting temperatures for the mantle owing to the presence of volatiles (i.e. water and CO2). Based on geophysical observations, Pascal & Cloetingh (2002) proposed a rheological model that considers lithosphere thickness heterogeneities in the Oslo region (Fig. 25). Their modelling shows that such heterogeneities could have resulted in strong localization of deformation in the Oslo Rift. A similar study by Pascal et al. (2004) showed that the introduction of lithosphere thickness contrasts in the models results in pronounced differential stretching between crust and mantle lithosphere, which, in turn, leads to decompression melting of the mantle over relatively short time periods subsequent to the onset of rifting. In summary, the models of Pascal & Cloetingh (2002) and Pascal et al. (2004), in which the mechanical behaviour of the rocks and a more realistic configuration for the lithosphere are included, complement the study by Pedersen & van tier Beek (1994). Although modelling results are very often more suggestive than firmly conclusive and need to be compared with nature, whenever it is possible, they appear here to go against a plume hypothesis for the Permian rift event in Europe. Henk (1999) used rheological modelling of Permian basins of Europe to examine the post-convergence evolution of the region. The purpose of his modelling approach was to explore whether the Variscides simply collapsed following the end of the orogenesis, thus leading to Permian tiffing, or whether the region was also influenced by far-field extension. Various 2D models were presented by Henk (1999), and he concluded that far-field extension superimposed on gravity stresses are required to overcome the strength of the post-Variscan lithosphere. Along the LT-7 deep seismic refraction profile in the NW Polish Basin (Guterch et al. 1994), 1D rheological modelling using a simplified petrological model of lithospheric layering was completed. The results suggest that the lithosphere, except for the East European Craton (EEC), is mechanically decoupled, and that the upper crust is separated from the upper mantle by extremely weak and ductile middle and lower crustal layers (c. 20 km thick). Only within the Tornquist-Teisseyre Zone and the EEC can the lower crust remain strong. The lithosphere of the EEC is probably entirely coupled except for the edge of the craton, where, with the low strain rates, mechanical discontinuity may occur at the middle-lower crust or lower crust-mantle boundaries. Laterally, the cumulative strength of the lithosphere changes by more than an order of magnitude (Jarosinski et al. 2002; Grad et al. 2003).
Tectonic a n d structural m o d e l s
Based on geological and geophysical data, tectonic and structural modelling of an object usually summarizes and tests the admissibility of combined information measured and observed in the field and laboratory. Balanced sections thus provide geologically reasonable constraints (Dahlstrom 1969), a concept that has been widely used in the hydrocarbon industry (Bally et al. 1966; Rowan & Kligfield 1989), but also is used to reveal the nature
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Fig. 25. Numerical modelling of the Oslo Rift, involving rock rheology and heterogeneity in lithosphere thickness (after Pascal et al. 2004). The thickness of the lithosphere in the left and right parts of the model is initially equal to 125 km and 180 kin, respectively. The modelled line is 500 km long at t -----0 Ma. The model is stretched using a velocity of 1.6 cm a- 1. The upper panel presents the horizontal strain distributions (i.e. exx) 1 Ma and 9 Ma after rift initiation. (Note the strong strain localization at the middle of the model and the Earth surface depression simulating basin formation.) The lower panel presents the thermal evolution (i.e. isotherms) of the lithosphere. Note the rise at t ~ 9 Ma of hot mantle rocks below the area that is depressed at the surface. The finite-element grid used for the computations is also shown. U.C., upper crust; L.C., lower crust; L.M., lithospheric mantle.
of tectonic processes and kinematic evolution in the area of interest (e.g. Oncken 1989). In the Central European Variscides, extensive studies were carried out to determine the pre-Variscan and Variscan evolution (see summary by Franke et al. 2000), but only few comprise 2D and 3D geometric and tectonic modelling of late, Variscan (e.g. Plesch & Oncken 1999; Oncken et al. 2000, and references therein; Schtifer et al. 2000) or even postVariscan development (Tanner et al. 1998). In the NE German Basin, the only palinspastic reconstructions available are by Kossow & Krawczyk (2002), based on results from the BASIN96 and commercial seismic surveys (Krawczyk et al. 1999; Kossow et al. 2000). The flexural cantilever model (see Kusznir et al. 1991, for model details) was also applied for forward modelling of the initial phase of NEGB formation in combination with detailed analysis of core material (Rieke et al. 2001; Fig. 26). NE German Basin formation was initiated during the Early Permian and was largely controlled by normal faulting related to deep-seated ductile shearing, with a steep and faulted eastern and a gently dipping western basin margin. A post-rift subsidence phase of 35 Ma immediately followed this east-west extension. The cantilever model predicts a stretching factor of/3 = 1.2 in the basin centre and 1.0 at the margins, which would have only
a slight effect on the crustal structure. The resulting smooth Moho uplift would fit well with the observed seismic data (Krawczyk et al. 1999). Restoration of the subsequent postZechstein kinematic evolution of the NEGB along a 260 km long N E - S W cross-section further indicates two major uplift periods at the Jurassic-Cretaceous and the Cretaceous-Tertiary boundaries (Kossow & Krawczyk 2002). Quantification of geological processes yields a total basement subsidence of 2850 m in the basin centre from end-Zechstein to present, maximum erosion of 860 m during the Cretaceous-Tertiary event at the southern NEGB margin, and at least 9 km of basin shortening. Interestingly, there is a clear correlation between the deformation intensity and the amount of uplift and erosion associated with the Cretaceous-Tertiary deformational period in the NEGB. Deformation intensity decreases from south to north, as do uplift rates, thus suggesting compression from the south, which was probably related to Alpine-induced intraplate deformations (Kossow & Krawczyk 2002). The Permian-Mesozoic development and tectonic inversion of the Polish Basin has been modelled using a 3D structural model combining analysis of 3D depth views and thickness maps (Lamarche et al. 2003a; Lamarche & Scheck-Wenderoth 2005). The model confirms earlier ideas that the Polish Basin and the
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Fig. 26. Schematiccross-section across the northern part of the NE German Basin showingthe faulted basement comprising Permo-Carboniferousvolcanic units, which were subsequentlyoverlain by Rotliegend sediments(after Rieke et al. 2001).
Mid-Polish Swell are genetically related to the Teisseyre-Tornquist Zone, which seems to have tectonically controlled the development of the area through time (e.g. Kutek & Glazek 1972; Dadlez et al. 1995; Kutek 2001). When the Mid-Polish Trough started to form, the Teisseyre-Tornquist Zone constituted a zone of crustal weakness that was prone to extensional deformation. Crustal thinning along the Teisseyre-Tornquist Zone, rifting, and the following Mesozoic subsidence resulted in additional weakening along the zone. As a result, when the stress conditions changed from transtensional to compressional at the end of the Cretaceous, the Teisseyre-Tornquist Zone was preferentially deformed, inducing the inversion of the Mid-Polish Trough and the uplift of a central NW-SE-elongated anticlinorium along the former basin axis, as well as the formation of two bordering marginal troughs (see Krywiec 2002a; Lamarche et al. 2003a for details). This geometry is the surface expression of the tectonic squeezing of the Teisseyre-Tornquist Zone, which played the role of an intra-continental zone of crustal weakness as modelled by Nielsen & Hansen (2000), Hansen et al. (2000) and Gemmer et al. (2002). Although the stress magnitudes may have significantly decreased after the climax of the tectonic inversion, the stress pattern remained compressional, as indicated by the Cenozoic central horst and marginal troughs developed above the Mid-Polish Swell (Lamarche et al. 2003a; Lamarche & Scheck-Wenderoth 2005). The Teisseyre-Tornquist Zone can be considered as a regional weakness zone within which the deformation was localized. A strong tectonic inheritance of Palaeozoic and Precambrian basement structures influenced the deformation during the tectonic inversion (Krzywiec 2004). As a result of the mosaic nature of its Palaeozoic basement, the southwestern flank of the Mid-Polish Trough was tectonically unstable during the Mesozoic, in contrast to the stability of the Precambrian East European Craton beneath the northeastern part of the Mid-Polish Trough. The model of Lamarche & Schech-Wenderoth (2005) and tectonostratigraphic models based on seismic reflection data (Krzywiec 2004) also show the Zechstein salt-beating layer acting as a decoupling level between the pre-Zechstein basement and the Mesozoic cover in the central and northern segments of the Polish Basin, inducing disharmonic deformation during the tectonic inversion. Thus, the idea is that the TTZ was a zone of weakness allowing the Polish Trough to form. Such an idea is supported by the fact that long-lived shear zones (in the crust, but probably also in the mantle) tend to focus strain without regard to the past tectonic context of the area. This is a fact, and is totally independent of theoretical models. For example, the border faults of the Viking Graben are at present the loci of a high degree of micro-seismic activitiy (e.g. Olesen et al. 2004). This observation is in clear contradiction to the idea that crustal thinning implies (following thermal relaxation) lithospheric strengthening (with respect to nearby non-rifted areas). Furthermore, recent advances in fault zone rheology suggest that repetitive deformation of the fault zone results in the development of an in situ mylonitic foliation and concentration of weak phases, which imply a drastic decrease in the coefficient of friction in the
fault zone and potentially a local drop in crustal strength (Bos & Spiers 2002; Holdsworth 2004).
Discussion The Variscan orogen was characterized by a particularly long period of intracontinental deformation, associated with the collision of Gondwana and Laurussia. The post-collisional evolution of Europe (i.e. within the latest Carboniferous-Early Permian time frame) was characterized by the formation of a series of rift, and wrench-induced, basins across the continent, together with significant magmatic events. From the above outline it can be seen that although we have a reasonable understanding of the broad evolution of the late stages of the Variscan orogenic event and the subsequent period of wrench fault activity that was widespread across both the internal and external Variscan provinces, there are many problems relating to our understanding both of the underlying mechanisms that controlled the various observed events and of the detailed integration of the various observations. In particular, there are problems relating the internal and external zones, which have, at times, remarkably similar evolutionary histories (e.g. coeval graben formation and associated volcanism in northern Spain, Italy and northern Germany). Although certain events may be interpreted in terms of plume-related activity, how do we interpret similar successions thousands of kilometres apart? Although it is clear from the above outline that the post-Variscan period in Central, Southern and Western Europe was a period of intense tectonic, magmatic and sedimentary change, any attempt at summarizing these changes must, by necessity, try to assess the various possible driving mechanisms involved in the generation of the post-Variscan basins. It is clear that the coincidence of tectonic activity (both compressional and extensional), magmatic activity and basin formation (with subsequent sedimentation) was very different from the periods immediately before and after. The geological evolution of the region, however, is problematic given the relative lack of significant Early Permian extensional structures. The large amounts of crustal-derived and crustalcontaminated volcanic rocks are also problematic. The processes controlling the post-orogenic modification of the Variscan lithosphere have been variably attibuted to such mechanisms as slab detachment, delamination of the mantle-lithosphere, crustal extension and plume activity during the Stephanian-Early Permian phase of wrench faulting and magmatism that overprinted the Variscan orogen and its foreland (see Ziegler et al. 2006, for references). The following sections will attempt to examine the main controlling mechanisms within the basin, to try to isolate those that are of greatest importance in terms of overall basin evolution.
Rifting history
The examination of a variety of basins across Europe has allowed us to compare and contrast the various successions within the
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basins, as well as features such as basin form, controls on basin formation, and the magmatic, tectonic and infill history. The observed contrasts suggest that the underlying processes that controlled the post-Variscan evolution of Europe were very different between those areas located in the former Variscan foreland basin and those within the thrust front. The various modelling studies carried out on the post-Variscan Permian basins suggest very different mechanisms for each area. Permian rifling in the former Variscan hinterland seems to have been strongly controlled by the collapse of the mountain chain (Brunet & Le Pichon 1982; Prijac et al. 2000) with a possible far-field extension component also being plausible (Henk 1999). This process may also have been modified by the slow decay of the associated thermal anomaly (e.g. Paris Basin). In contrast, rifting in the former Variscan foreland appears to have been dominated by late Variscan wrench tectonics (van Wees et al. 2000), particularly along the boundary between Precambrian and Phanerozoic Europe (Dadlez et al. 1995). Numerical modelling highlights the role of such lithospheric discontinuities in controlling tiffing (Pascal & Cloetingh 2002; Pascal et al. 2003). Butler et al. (1997) have noted that pre-existing heterogeneities in the continental lithosphere are thought to influence its response to subsequent deformation. From the late Early Carboniferous onwards, Laurussia was transected by the Arctic-North Atlantic Rift System, which was partially superimposed on the Caledonian suture zone (Ziegler 1990). Indeed, Variscan exploitation of older Caledonian structures has been reported from other areas (e.g. offshore Ireland; for details, see Shannon 1991; McCann 1996). In Cornwall, early Variscan thrusts were reactivated as late Variscan extensional faults (Shail & Alexander 1997). Additionally, the interaction of the Variscan structures with the pre-Variscan east-west dextral (Badham 1982) transform fault system (running from the Uralides through Europe (Pitra et al. 1999) to the Appalachians) and the NNW-SSE-trending wrench fault system produced a complex series of conjugate shear zones and pull-apart structures in the Cornwall area (Willis-Richards & Jackson 1989) that remained active throughout the early Permian. It is, therefore, highly likely that, within the area under discussion, older structures, both Caledonian and Variscan, were reactivated by later Variscan tectonic activity. However, more recent work (Ebbing et al. 2006) has suggested that even older structure may be involved. In their study of the Oslo Graben they suggested that the rifting in the region is coupled to a reactivation of Precambrian fault systems, and indeed, the very location of the Oslo Graben is more strongly dependent on the pre-rift structure of the area than previously assumed. One factor of note is that Permian wrench activity was not merely limited to 'accreted' Europe, but is also evident in other parts of the craton where there is sufficient stratigraphic evidence. In particular, there is evidence of late Carboniferous-early Permian transtensional tectonic activity in the Dniepr-Donets Basin (Stovba & Stephenson 1999) and even further afield on the margins of the East European Craton (Saintot et al. 2006).
Mantle plume dynamics
Another important issue addressed by modelling of Permian basins, and in particular of the Oslo Rift (Ro & Faleide 1992; Pedersen & van der Beek 1994), is the eventual role played by a mantle plume (although this idea has recently been questioned; see Ebbing et al. 2006, for details). Despite significant differences in the tectonosedimentary setting and the type of magmatic activity within the various basins examined, the StephanianAutunian volcanic rocks in the internal Variscides comprise a high proportion of pyroclastic deposits and are generally of intermediate to felsic composition, of calc-alkaline character, and often have a significant crustal component, as shown by Sr-Nd isotope data and the presence of crustal xenoliths, magmatic garnet, and
(locally) topaz, and (for the volcanic rocks in the NE German Basin) the large amount of inherited zircons necessitating sensitive high-resolution ion microprobe (SHRIMP) dating (Breitkreuz & Kennedy 1999). The calc-alkaline character may reflect the derivation of the melts from a subduction-modified mantle source, extensive assimilation of crustal material, or perhaps inheritance resulting from the melting of older calc-alkaline, crustal sources (such as Cadomian basement). However, the relative scarcity of more primitive mafic melts precludes a more precise interpretation of the mantle source compositions. In addition, numerical studies suggest that huge volumes of magmas can be produced with small amounts of stretching and without the need for any underlying thermal anomaly (Pedersen & van der Beek 1994). Crustal p r o c e s s e s
The Stephanian-Autunian magmatic rocks in the internal Variscides comprise a high proportion of pyroclastic rocks and are generally of intermediate to felsic composition. Their generally calc-alkaline character suggests a subducfion-related origin. With the possible exception of some magmafic rocks in the Alpine basement, this contradicts their intracontinental setting and the fact that the Variscan oceans had closed by mid-Carboniferous times. However, Sm-Nd isotope data and the presence of garnet and crustal xenoliths indicate that many contain a significant crustal component. This is corroborated by the predominantly negative ENd(t) values of the 290-300 Ma volcanic and intrusive rocks of felsic to intermediate composition: - 2.1 to - 6.0 for the Krkonoge Basin (Ulrych et al. 2002), -2.7 to -6.1 for the Intra-Sudetic Basin (Ulrych et al. 2004); -0.8 to - 7 . 0 for the rhyolites of the Halle Volcanic Complex (Romer et al. 2001), -4.3 to -7.5 for the granites in Comwall (Darbyshire & Shepherd 1994), and - 0 . 6 to - 5 . 7 for the Saar-Nahe Basin (Schmidberger & Hegner 1999; von Seckendorff et al. 2004a, and references therein). The parent magmas of the granitoids, rhyolites and andesites may, therefore, have assimilated large amounts of crustal material, or alternatively, be derived from mantle sources that had been modified by earlier subduction events (e.g. Cabanis & Le Fur-Balouet 1989; Schmidberger & Hegner 1999; Innocent et al. 1994; Cortesogno et al. 1998). As in the North German Basin, the granites and rhyolites may be of crustal origin, and their calc-alkaline signature inherited through partial melting of calc-alkaline basement (Schaltegger 1997b; Romer et al. 2001). The possible mechanisms for mantle melting in the internal Variscides may have been the break-off of subducted oceanic crust (e.g. Schaltegger 1997b; Cesare et al. 2002) or even the oblique subduction of the mid-ocean ridge of Palaeotethys beneath the active Eurasian margin (Stampfli 1996). Regional extension leading to lithosphefic thinning and decompressional melting of updoming asthenosphere may have been a contributing factor in the late Carboniferous-early Permian period. Compared with the foreland, Stephanian-Autunian mafic rocks are much rarer in the internal Variscides, which suggests that the mantle-derived parent melts were unable to reach the surface, but stalled at lower to midcrustal levels. This may have been due to the large contrast between the density of the parent melt and a low average density of thinned Variscan crust. Only after fractionation and assimilation of sufficient amounts of crustal material did the melts attain a low enough buoyancy to be able to escape the magma chambers and erupt on the surface. M a g m a t i c - t e c t o n i c activity
The relatively short and widespread pulse of StephanianAutunian magmatism is likely to have taken place in response to changes in the regional stress field at the Westphalian-Stephanian boundary and subsequent thermal equilibration of the lithosphere. The change of stress may have been due to a change in VisranWestphalian crustal shortening and orogen-parallel extension,
POST-VARISCAN BASIN EVOLUTION, EUROPE
and to Stephanian-Autunian gravitational collapse of the Variscan orogen. The latter process was possibly superimposed and aided by a far-field dextral extensional stress-feld that was due to the collision of Gondwana with eastern southeastern North America and concomitant dextral translation (Torsvik & Van der Voo 2002). The invocation of far-field effects is something that has previously been noted in discussions of postVariscan tectonics (e.g. discussion on the origin of the NE German Basin; see DEKORP-BASIN Research Group 1999, for details). In terms of the magmatic history of the post-Variscan there are some indicators that far-field effects might also have played an important role. For example, the alkaline composition, style of volcanism and the presence of abundant megacrysts and mantle xenoliths in the Early Permian mafic rocks in Scotland indicate derivation by low-degree melting of local mantle sources and rapid, vertical transport. In contrast, the sub-alkaline mafic dyke and sills complexes (such as the Whin Sill Complex) indicate higher degrees of mantle melting, and do not necessarily reflect a mantle thermal anomaly of the same extent. The geometry and orientation of the dyke swarms suggest a magmatic focal region in the vicinity of the Denmark-Skagerrak region, which suggests that magma transport may have been horizontal, westwards into the North Sea and Scotland. Thus, the position, trend, number and size of the dykes may have been controlled by the far-field dextral extensional stress field.
Conclusions The end Carboniferous-early Permian history of Europe represents a period of crustal instability and re-equilibration throughout Western and Central Europe. An extensive and significant phase of Permo-Carboniferous magmatism led to the extrusion of thick volcanic successions across the region. Coeval transtensional activity led to the formation of more than 70 rift basins, which differ both in form and infill according to their position relative to the former Variscan Orogenic Front as well as to the controls that acted on basin development. Despite the fact that no unified model for the Permian event can at present be unequivocally proposed from the results of the various modelling studies, recent studies do agree on two fundamental and relevant points: (1) Permian rifting was widespread in Europe with progressively propagated development; (2) its signature strongly influenced the evolution of the European lithosphere during Mesozoic and Cenozoic times (S0rensen 1986). It may not, however, be possible to provide more detailed models for the evolution of the region. Numerical modelling of lithospheric rifting, for example, requires numerous parameters, among which the pre-rifl crust and mantle-lithosphere structure are crucial. Because the pre-Permian lithosphere structure has been obscured by repetitive tectonic phases in most parts of Europe, lithosphere-scale modelling of the Permian event remains difficult and modelling results need to be treated with a high degree of circumspection. The best approach, therefore, to elucidating the tectonosedimentary and magmatic history of the region is to adopt a broad approach, examining the various basins at a range of scales and making use of a variety of techniques. This manuscript was greatly improved by the reviews of two anonymous reviewers. T. Beilfuss is thanked for the production of the excellent diagrams.
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Pre-Alpide Palaeozoic and Mesozoic orogenic events in the Eastern Mediterranean region A. I. O K A Y 1, M. SATIR 2 & W. SIEBEL 2
lJstanbul Teknik (Yniversitesi, Avrasya Yerbilimleri Enstitiisii, Ayazafi, a 80626, Istanbul Turkey (e-mail: okay@ itu.edu.tr) 2Institut fiir Geowissenschaften, Universitiit Tiibingen, Wilhelmstrafle 56, D-72074 Tiibingen, Germany
Abstract: We review the Palaeozoic-Early Mesozoic evolution of the Eastern Mediterranean-Balkan region with special reference to Anatolia, and provide new isotopic data on the Palaeozoic magmatic and metamorphic rocks. The pre-Alpide evolution of the region involves episodic growth of Laurussia by accretion of oceanic terranes and Gondwana-derived microcontinents. Terrane accretion, associated with deformation, magmatism and regional metamorphism, took place in the Late Ordovician-Early Silurian, Carboniferous, Late Triassic-Early Jurassic and Mid-Jurassic. The Late Ordovician-Early Silurian accretion is inferred from stratigraphic and faunal records in the Pontides; other evidence for it is buried under young cover on the northern margin of the Black Sea. The Carboniferous orogeny is related to southward subduction and continental collision on the southern margin of Laurussia. It is marked in the Pontides by high-grade regional metamorphism, north-vergent deformation and post-orogenic latest CarboniferousEarly Permian plutonism. The latest Triassic-Early Jurassic Cimmeride orogeny involved the collision and amalgamation of an oceanic plateau to the southern margin of Laurasia. It is represented by voluminous accretionary complexes with Late Triassic blueschists and eclogites. Late Jurassic regional metamorphism and deformation is confined to the Balkans, and is the result of continental collision between the Rhodope-Serbo-Macedonian and Strandja blocks in the Late Jurassic. The Palaeozoic geological history of the Balkans and the Pontides resembles that of Central Europe, although the similarities end with the Mesozoic, as a consequence of the formation of Pangaea.
Orogenic belts and Mesozoic oceanic basins occupy the Eastern Mediterranean region between the stable areas of the East European Craton in the north, and NE Africa and the Arabian Platform in the south (Fig. 1). The East European Craton, as represented by the Ukrainian Shield north of the Black Sea, is an A r c h a e a n Palaeoproterozoic crystalline terrane. The consolidation of the southern part of the East European Craton was completed by 2300-2100 Ma (e.g. Bogdanova et al. 1996; Claesson et al. 2001). In the Early Palaeozoic, the East European Craton formed part of the Balfica plate, which collided in the west with Laurentia, Avalonia and Armorica, creating Laurussia in the Late Palaeozoic (e.g. Pharaoh 1999; Matte 2001; Wart 2002). In contrast, Africa and the Arabian Platform constituted part of Gondwana, which preserved its unity until the Early Mesozoic opening of the southern Atlantic. Very large areas in the northern margins of Gondwana are characterized by NeoproterozoicCambrian plutonism and metamorphism forming part of the Pan-African-Cadomian orogenic cycle (e.g. Stern 1994), and are therefore readily distinguished from the Palaeoproterozoic basement of the East European Craton. During the Late Palaeozoic and Mesozoic, Tethyan oceanic basins separated Laurussia from NE Africa-Arabia. Parts of the present Eastern Mediterranean Sea represent a Triassic to Jurassic remnant of a Tethyan oceanic crust, whereas the Black Sea is a Late Cretaceous oceanic back-arc basin that opened during the northwards subduction of a Tethyan ocean (e.g. ~eng6r & Yllmaz 1981; Garfunkel 1998). The Anatolian-Balkan region between the Eastern Mediterranean and the Black Sea consists of several small continental fragments or terranes bearing evidence of various periods of deformation, metamorphism and magmatism, the latest and strongest of which is the Alpide orogeny. The Alpide orogeny resulted in the amalgamation of the continental fragments into a single landmass in the Tertiary. Previous to this amalgamation, these continental fragments were situated on the margins of the Tethyan oceans, or formed small edifices within the ocean. The pre-Alpide orogenic history of these terranes forms the subject of this paper.
Terranes in the Eastern Mediterranean-Balkan region As orogenic events are restricted to the plate margins, identification of former plates is important for an understanding of the orogenic evolution. Only the deformed continental parts of the former microplates would be expected to be preserved, and they would be rimmed by sutures marked by linear zones of accretionary complex, blueschist, eclogite and ophiolite, and would show distinctive strafigraphic features, especially if the intervening oceans were large. The main methods used in the differentiation of the former plates include recognition of sutures, palaeomagnetism, faunal provinciality and stratigraphy. A complication in this picture is that the number and configuration of the plates change through time. For example, the Anatolian microplate, which makes up most of the present Anatolian landmass, did not exist before the Miocene (e.g. Seng6r 1979). Furthermore, late-stage strike-slip faulting may lead to the dispersal of a single palaeoplate, as happened during the Cretaceous opening of the Black Sea (Okay et al. 1994). Therefore, a better term for such a palaeo-plate would be 'terrane', and this would be distinguished by its distinctive stratigraphic, palaeomagnetic, faunal, structural, metamorphic and magmatic features. The following terranes are defined in the Anatolian region from south to north: the Arabian Platform comprising part of SE Anatolia, the Anatolide-Taufide Block, the Kir~ehir Massif, the Sakarya Zone, the Istanbul Zone and the Strandja Massif (Fig. 1; Okay & Tfiystiz 1999). Of these, the last three are grouped together as the Pontides. Pre-Alpide orogenic events are especially strong and well documented in the Ponfides (Fig. 2). With the possible exception of its northwestern margin, the Anatolide-Tauride Block was largely free of Palaeozoic and early Mesozoic deformation. Therefore, most of this review concerns the pre-Alpide geological history of the Pontides. In addition to the Anatolian terranes listed above, several other terranes have been defined in the Balkans (e.g. Burchfiel 1980; Stampfli 2000; Stampfli et al. 2001). The major ones include the Moesian Platform, the Rhodope-Serbo-Macedonian Massif, Pelagonia (which comprises the Pelagonian Zone and the
From: GEE, D. G. & STEPHENSON, R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 389--405.0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. Tectonic map of the Eastern Mediterranean region showing the major terranes and the bounding sutures. The filled triangles indicate the polarity of the subduction (modified from Okay & Ttiystiz 1999). NAF, North Anatolian Fault; EAF, East Anatolian Fault.
Cyclades) and Apulia, which includes Greece south of the Pindos suture (Fig. 1; Stampfli et al. 2001). Palaeozoic rocks are not exposed on Apulia, so it is not known whether this region was affected by the Variscan orogeny. Assuming that it was not, then Apulia probably forms a single terrane jointly with the Anatolide-Tauride Block. Isotopic data indicate Carboniferous (325-295 Ma) plutonism and regional metamorphism in the Pelagonian Zone and in the Cyclades (see the discussion by Vavassis et al. 2000), which contrasts with the Neoproterozoic magmatic and metamorphic basement ages for the Anatolide-Tauride Block. Their Mesozoic histories are also different, with Late Jurassic-Early Cretaceous ophiolite obduction on the northern margin of Pelagonia contrasting with the Late Cretaceous ophiolite obduction on the Anatolide-Tauride Block. The contact between Pelagonia and the Anatolide-Tauride Block is represented by an Eocene thrust, where the cover sequence of the Cycladic Massif, metamorphosed at blueschist-facies conditions in the Eocene, is thrust on the Menderes Massif (Fig. 3; Okay 2001). This contact, which may represent the extension of the Pindos suture, may link up with the Izmir-Ankara suture, in which case Pelagonia will be correlated with the Sakarya Zone (Fig. 1). Such a correlation is supported by the similar Variscan magmatic and metamorphic ages from Pelagonia and the Sakarya Zone (see below), although their Mesozoic histories are separate. The Alpide orogeny in the Mediterranean area started with the convergence between the Africa-Arabian and Eurasian plates during the Late Mesozoic. The relative movement between these two plates was sinistral strike-slip from Early Jurassic to midCretaceous time (e.g. Savostin et al. 1986; Dewey et al. 1989). Starting with the Cenomanian-Albian (100-90 Ma) the Africa-Arabian and Eurasian plates started to converge, presumably with the initiation of subduction. In the geological record the Albian flysch of the Central Pontides (Ttiystiz 1999), the Turonian (c. 91 Ma) high-temperature-medium-pressure metamorphism in the Klr~ehir Massif (Whitney et al. 2003), and the Campanian (c. 80 Ma) high-pressure-low-temperature metamorphism in the Anatolide Tauride Block (Sherlock et al. 1999) are the first
recognized events of the Alpide orogeny in Turkey. Therefore, the period discussed in this review extends to the Early Cretaceous. We also do not discuss the Neoproterozoic-Cambrian-aged Pan-African orogenic events in Anatolia, which, although important (e.g. Krrner & ~engrr 1990; Yi~itba~ et al. 2004), are poorly preserved and documented.
The Anatolide-Tauride Block The Anatolide-Tauride Block has a Neoproterozoic crystalline basement overlain by a sedimentary succession ranging from Mid-Cambrian to Miocene in age (e.g. Gutnic et al. 1979; Ozgiil 1984, 1997). It was strongly deformed and partly metamorphosed during the Alpide orogeny, and now consists of metamorphic regions in the north (the Anatolides) and a south-vergent Eocene nappe stack in the south (the Taurides). Stratigraphy of several nappe units in the Taurides reveals Palaeozoic to Mesozoic sedimentary sequences with no evidence of pre-Cretaceous deformation or metamorphism (e.g. Gutnic et al. 1979; Ozgtil 1984, 1997). Rare reports of Late Triassic deformation in the Anatolide-Tauride Block (e.g. Monod & Akay 1984) have been questioned and need confirmation (G6nctio~lu et al. 2003). The Anatolide-Tauride Block is here considered as not affected by significant pre-Alpide Phanerozoic contractional deformation, which contrasts with the regions to the west and north that were deformed and metamorphosed during the Variscan orogeny. The largest outcrops of the Precambrian basement in the Anatolide-Tauride Block are found in the Menderes and Bitlis massifs (Figs 3 and 4). In the Menderes Massif, the metagranitoids, which make up most of the crystalline basement, have been dated as to c. 550 Ma using a stepwise Pb evaporation method on zircons (Hetzel & Reischmann 1996; Loos & Reischmann 1999); this is a similar age to those found in the Arabian Platform and in NE Africa. The eclogite-facies metamorphic rocks in the basement of the Menderes Massif are also Neoproterozoic in age (Candan et al. 2001). The Palaeozoic stratigraphy in the Anatolide-Tauride Block is also similar to
OROGENS IN THE EASTERN MEDITTERANEAN
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Fig. 2. A chronostratigraphic chart showing generalizedgeological relationships of the Pontic terranes. The main sources of the data are: for the Istanbul Zone, Gedik (1975), Ali~an& Derman (1995), G6rtir et al. (1997) and Dean et al. (2000); for the Strandja Zone, Chatalov (1988), and Okay et al. (2001); for the Sakarya Zone, Altaneret al. (1991) and Okay & Leven (1996).
that of the northern margin of the Arabian Platform. Hence, since Smith (1971), all palaeogeographical reconstructions place the Anatolide-Tauride Block into the Eastern Mediterranean Sea between the Levant and Egyptian margins. Recent evidence for the latest Ordovician (Hirnantian) glaciation in the AnatolideTauride Block, including the presence of striated pebbles and striated basement (Monod et al. 2003), supports this pre-drift position. Stratigraphic evidence from the Levantine (e.g. Bein & Gvirtzman 1977; Garfunkel & Derin 1984) and Gondwana margins in SE Anatolia (Fontaine et al. 1989) indicates that the Anatolide-Tauride Block rifted away from Gondwana during the Triassic or Early Jurassic with the opening of the Tethyan ocean. However, it was never far away from Gondwana, and drifted with Gondwana to the north, as shown by its Jurassic palaeomagnetic record (Piper et al. 2002) and by its Jurassic-Cretaceous stratigraphy, which resembles that of the SE Anatolia.
The Istanbul Zone The Istanbul Zone consists of a Neoproterozoic crystalline basement overlain by Lower Ordovician to Eocene sedimentary rocks (Fig. 2; e.g. Haas 1968; G6rtir et al. 1997). Before the Late Cretaceous opening of the Black Sea, the Istanbul Zone was situated east of the Moesian Platform and adjacent to the Scythian Platform (Okay et al. 1994). These three tectonostratigraphic units have a similar basement and show a similar Palaeozoic-Mesozoic stratigraphic development (e.g. Tari et al. 1997; Nikishin et al. 1998), and formed a single late Proterozoic-Early Palaeozoic terrane, named here the MOIS (Moesian-Istanbul-Scythian) terrane. The crystalline basement of the Istanbul Zone is characterized by voluminous granitoids, which intrude low- to medium-grade metasediments and metavolcanic rocks (Usta6mer & Rogers 1999; Yi~itba~ et al. 1999, 2004). The granitoids have yielded U - P b and Pb/Pb
392
A.I. OKAY ETAL.
Fig. 3. Tectonic map of the western Anatolia illustratingthe geological features discussedin the text. The cross-hatched area shows the extent of the metamorphosed Palaeozoic and Mesozoic rocks of the Menderes Massif.
zircon ages of 590-560 Ma, and the surrounding metasediments have provided similar R b - S r mica ages (Chen et al. 2002). The geochemistry of the granitoids and the metavolcanic rocks indicates a subduction-zone setting in the Neoproterozoic. In terms of age, lithology and geochemistry, the basement of the Istanbul Zone is similar to the Pan-African basement of northern Gondwana, and unlike the East European Craton. Therefore, the Istanbul Zone is generally regarded as a Peri-Gondwana terrane (e.g. Stampfli 2000). The Neoproterozoic basement of the Istanbul Zone is overlain by a thick Palaeozoic sedimentary succession extending from the Ordovician to the Carboniferous (Fig. 2). There are significant stratigraphic differences between the westem and eastern parts of the Istanbul Zone, which led to a suggestion that the Istanbul Zone consists of two terranes, the Istanbul terrane in the west and the Zonguldak terrane in the east (Kozur & G6nctio~lu 1998; Stampfli et al. 2002; von Raumer et al. 2002). The most important difference is in the Carboniferous system, which in the west is represented by Vis~an radiolarian cherts overlain by siliciclastic turbidites, but in the east, in the Zonguldak region, is
represented by Vis6an neritic carbonates overlain by Namurian to Westphalian coal measures (Figs 2 and 5). However, evidence for a Phanerozoic ocean, in terms of pelagic sedimentary rock, m61ange, ophiolite or blueschist, is missing between the western (Istanbul s e n s u stricto) and eastern parts (Zonguldak) of the Istanbul Zone (see Fig. 5), and the stratigraphic differences are a result of facies changes. A similar situation has been reported in the Moesian Platform, where neritic carbonate deposition in the Tournaisian in the north is replaced by radiolarian chert sedimentation in the south in the Elovitza region (Haydoutov & Yanev 1997). During the Carboniferous, the MOIS terrane was part of the southern continental margin of Laurussia. Turbidite deposition in a continental slope setting took place in the western part of the Istanbul Zone and on the southern margin of the Moesian Platform, whereas coal deposition occurred in swamps in the north (Fig. 6). The palaeogeographical situation was similar to that of SW Britain at the same period, when coal measures were being deposited in Wales and siliciclastic turbidites (Culm facies) in Cornwall and Devon (Fig. 6; e.g. Guion et al. 2002).
OROGENS IN THE EASTERN MEDITTERANEAN
393
Fig. 4. Tectonic map of eastern Anatolia illustrating the geological features discussed in the text. (For legend see Fig. 3.) The cross-hatched area shows the extent of the metamorphosed Palaeozoic and Mesozoic rocks of the Bitlis Massif.
Fig. 5. The distribution of pre-Jurassic rocks in the Istanbul Zone (simplified from Aksay et al. 2002; Ttirkecan & Yurtsever 2002). Noteworthy features are the different Carboniferous and Triassic facies in the west and east, and the trend of the facies boundary, which is highly oblique to the Intra-Pontide suture.
394
A.I. OKAY ETAL.
Fig. 6. Carboniferous palaeogeography in the southern margin of Laurussia (a) compared with that of Britain at the same period (b). Both maps are of the same scale. In (a) the Istanbul Zone is restored to its predrift position before the Cretaceous opening of the Western Black Sea basin (Okay et al. 1994). In the Tournaisian-Vis~an, neritic carbonate deposition took place in Moesia and in the eastern Istanbul Zone; this was succeeded by the accumulation of coal during the Namurian and Westphalian. In the same period radiolarian chert sedimentation gave way to siliciclastic turbidite deposition in a continental slope setting in the western Istanbul Zone and the southern margin of Moesia. A similar picture exists in Britain, where, in addition, the Early Devonian Lizard ophiolite in Cornwall provides another indication of the Rheno-Hercynian ocean in the south. It should be noted that the Intra-Pontide suture truncates the facies belts. The Moesia data are from Dachev et aL (1988), Popova et al. (1992), Tenchov (1993), Haydoutov & Yanev (1997) and Tari et aL (1997); the data for Britain are from Guion et al. (2002) and Warr (2002).
The Intra-Pontide suture, which marks the southern boundary of the Istanbul Zone, truncates the Palaeozoic and Triassic facies boundary between the westem and eastern parts of the Istanbul Zone (Figs 5 and 6). This suggests removal of a major section of the Istanbul Zone, possibly by post-Triassic strike-slip faulting. Late C a r b o n i f e r o u s d e f o r m a t i o n a n d p l u t o n i s m
The Palaeozoic sequence in the Istanbul region ends with Visran to Namurian siliciclastic turbidites, whereas in the east, in the Zonguldak region, it extends into the Westphalian coal measures (Fig. 2; G6riir et al. 1997). The Palaeozoic rocks in the Istanbul region are deformed in a contractional mode, with the generation of recumbent folding, local cleavage and minor thrusting, whereas deformation is less intense in the Zonguldak region. The minor folds generally show an east to NE vergence (Seymen 1995; Zapcl et al. 2003), although the timing of deformation, whether Variscan or later, is difficult to constrain. Nevertheless, the observation that the lowermost Triassic red beds step down from Carboniferous to Ordovician (Ttirkecan & Yurtsever 2002) indicates significant deformation and erosion in the Late Carboniferous-Permian interval. The deformed Palaeozoic rocks are intruded by a Permian granite east of Istanbul, which has biotite K - A r and whole-rock R b - S r ages of c. 255 Ma (early Late Permian; Figs 3 and 5; Yllmaz 1977). The age of the undeformed pluton constrains the Variscan deformation in the Istanbul Zone to the Late Carboniferous-Early Permian. The Triassic restoration
The Palaeozoic rocks in the Istanbul Zone are unconformably overlain by Triassic continental clastic rocks with basaltic flows
(Fig. 2). In the Istanbul region, the Triassic sequence continues with neritic to pelagic carbonates, capped by Carnian or Norian siliciclastic turbidites, showing a typical transgressive..passive margin type of development (e.g. Gedik 1975; Yurtta~-Ozdemir 1971), whereas in the east the Triassic is represented mainly by continental clastic rocks and lacustrine limestones (Figs 2 and 5). The change in the Triassic facies closely follows that of the Palaeozoic, suggesting a long-term hinge, possibly controlled by a deep-seated fault (Fig. 5). The termination of deposition in the Carnian or Norian in the Istanbul Zone probably reflects the Cimmeride orogeny, which is particularly strong in the Sakarya Zone farther south.
P a l a e o g e o g r a p h i c a l affinity
The Infra-Cambrian to Cambrian granitoids (590-560 Ma) and Neoproterozoic metamorphism in the basement of the Istanbul Zone suggest a location on the Gondwana margin in the latest Precambrian. This is supported by the Ordovician trilobite faunas, which are similar to those from Central European and Anglo-Welsh successions, and differ from those of Baltica, as well as from those of typical Gondwana realms of the Anatolide-Tauride Block and the Arabian Platform (Dean et al. 2000). Therefore, a location of the MOIS terrane on the western margin of Baltica during the Early Ordovician, as shown in some reconstructions (von Raumer et al. 2002) is not possible. The absence of the latest Ordovician (Hirnantian) glaciation in the Istanbul Zone provides another constraint on its location on the Gondwana margin. However, from the Late Silurian onwards the Istanbul Zone became part of Laurussia, as indicated by its palaeomagnetic record from sediments of Late Silurian, Devonian, Carboniferous and Triassic age (Sarlbudak et al.
OROGENS IN THE EASTERNMEDITTERANEAN 1989; Evans et al. 1991), and by the Devonian-Carboniferous foraminiferal assemblages (Kalvoda 2003; Kalvoda et al. 2003). These data imply that the MOIS terrane separated from Gondwana during the Ordovician, and docked with Baltica in the Late Ordovician-Early Silurian; however, there is little evidence for Ordovician-Silurian collision in the geological record of the Istanbul Zone. Apparently, the zone of collision is hidden under young cover on the northern margin of the Black Sea. The Early Palaeozoic history of the MOIS terrane appears to be remarkably similar to that of Avalonia (Stampfli et al. 2002; Winchester & the PACE TMR Network Team 2002).
The Strandja Massif The Strandja Massif forms part of large metamorphic region in the Balkans, which includes the Rhodope, Serbo-Macedonian and Peri-Rhodope zones (Fig. 1). The relationship between these metamorphic units, and their ages of regional metamorphism are poorly known. The Strandja Massif crops out both in Turkey and in Bulgaria, and is bordered in the west by the Rhodope Massif. It consists of a metamorphic basement of unknown age, intruded by Permian granitoids, and overlain by continental to shallow marine sedimentary rocks of Triassic to Mid-Jurassic age (Fig. 2). During the Late Jurassic, the cover and the basement of the Strandja Massif underwent contractional deformation and regional metamorphism, and Triassic allochthons were emplaced on the Mid-Jurassic metasediments. Cenomanian and younger sediments lie unconformably over the metamorphic rocks (Chatalov 1988; Okay et al. 2001). The basement of the Strandja Massif consists of gneisses and micaschists intruded by voluminous plutonic rocks, several of which have been dated as Early Permian (c. 271 Ma) using stepwise Pb evaporation method on single zircon grains (Okay et al. 2001). The overlying Triassic sequence of the Strandja Massif shows affinities to the Central European Germanic Triassic facies, with a basal continental clastic series overlain by Middle Triassic shallow-marine carbonates (Chatalov 1988, 1991). A hiatus between the Late Triassic and Early Jurassic is probably a distant echo of the Cimmeride deformations farther south (Fig. 2). The shallow marine sedimentation continued into the Mid-Jurassic (Bathonian), and was terminated by the Late Jurassic Balkan orogeny. Late Jurassic deformation and metamorphism in the Strandja Zone
The Triassic to Jurassic sedimentary cover sequence of the Strandja Massif, together with its crystalline basement, underwent deformation and greenschist-facies metamorphism during the Late Jurassic. The age of regional metamorphism is constrained to the Late Jurassic-Early Cretaceous (Callovian-Albian) by the Bathonian age of the youngest metamorphosed strata (Chatalov 1988), and by the Cenomanian post-metamorphic cover (Fig. 2). R b - S r and K - A r biotite ages from the deformed and metamorphosed Permian granites of the Strandja Massif fall in the range of 155-149 Ma (Aydln 1988; Okay et al. 2001), indicating a Late Jurassic age for the regional metamorphism. The Late Jurassic metamorphism in the Strandja Massif was associated with north-vergent thrusting, folding, and the generation of foliation and lineation (Okay et al. 2001). Permian granitoids were penetratively deformed and thrust north over the Triassic to Jurassic mylonitic metasediments and marbles. Large allochthons, composed of Triassic deep-sea metasediments and metavolcanic rocks, were thrust northwards over the epicontinental Triassic-Jurassic rocks of the Strandja Massif (Chatalov 1985, 1988; Dabovski & Savov 1988). A foreland basin, called the Nij-Trojan trough, developed in the Oxfordian between
395
the Strandja-Rhodope massifs and the Moesian Platform. The Nij-Trojan trough migrated northward and persisted until the Early Cretaceous (Barremian; Tchoumatchenko et al. 1990; Harbury & Cohen 1997).
The Sakarya Zone The Sakarya Zone forms a continental sliver, over 1500 km long, south of the Istanbul Zone and the eastern Black Sea (Fig. 1). It consists mainly of Jurassic and younger sedimentary and volcanic rocks, which unconformably overlie a heterogeneous basement. The only sign of the Late Jurassic-Early Cretaceous deformation and metamorphism that is so intense in the Strandja Massif is a parallel unconformity at the base of CallovianOxfordian limestones (Fig. 2; Altlner et al. 1991). The pre-Jurassic basement of the Sakarya Zone includes Devonian plutonic rocks, Carboniferous plutonic and metamorphic rocks, and Triassic accretionary complexes with blueschists and eclogites (Figs 3 and 4). The pre-Jurassic relation between these basement units is strongly overprinted by Alpide deformations. The Devonian and Carboniferous units, and the Triassic accretionary complexes, are described below.
Early Devonian plutonism in the Sakarya Zone
The Devonian was a period of widespread granitoid plutonism in the Caledonides in NW Europe (e.g. Woodcock & Strachan 2002), whereas granitoids of this age were unknown in the Eastern Mediterranean region. Therefore, it was a surprise when a single sample from a granitoid in NW Turkey was dated as Early Devonian (Okay et al. 1996). The ~amllk granodiorite in the Biga peninsula (Fig. 3) forms a 20 km long and 3 - 4 km thick thrust sheet in an Alpide thrust stack. It is a leucocratic granodiorite consisting mainly of quartz, plagioclase and chloritized biotite, and is unconformably overlain by Upper Triassic arkosic sandstones. As the age of the granite is tectonically significant, zircons from a second sample from the (~amllk granodiorite were dated using the stepwise Pb-evaporation method. The details of the dating method have been given by Okay et al. (1996). Two zircon grains from the ~amllk granodiorite gave an Early Devonian age of 397.5 + 1.4 Ma (Fig. 7, Table 1), confirming the earlier less precise age of 399 +_ 13 Ma obtained by Okay et al. (1996). The relationship between the ~amlik Granodiorite and the other pre-Jurassic basement units of the Sakarya Zone are not known. However, the proximity of the highgrade Carboniferous metamorphic rocks of the Kazda~ and the essentially unmetamorphosed Devonian ~amllk granodiorite in NW Turkey (see Fig. 3) suggest major pre-Jurassic shortening between these two units.
Carboniferous deformation and metamorphism in the Sakarya Zone
The high-grade Variscan metamorphic basement of the Sakarya Zone is exposed in only a few areas throughout its 1500 km length. These include the Kazda~ and Uluda~ massifs in the west, the Devrekani Massif in the Central Pontides, and the Pulur Massif in the Eastern Pontides (Figs 3 and 4). These metamorphic regions are composed of gneiss, amphibolite and marble metamorphosed at amphibolite- to granulite-facies conditions, and in the Kazda~ and Pulur massifs there are also meta-ultramafic rocks within the sequence (Okay 1996; Okay et al. 1996; Duru et al. 2004; Topuz et al. 2004a). Isotopic age data exist only for the Pulur and Kazda~ massifs. Monazite Pb ages from a Pulur gneiss are late Early Carboniferous
396
A.I. OKAY ETAL.
Fig. 7. Histograms showing the distribution of radiogenic Pb isotope ratios derived from the evaporation of two zircon grains from the ~amhk granodiorite (a) and from a gneiss and an amphibolite of the Kazda~ Group (b) in the Sakarya Zone, NW Turkey.
(331-327 Ma, Namurian), considered as the age of high-grade metamorphism (Topuz et al. 2004a). The 315-310 Ma (Westphalian) N d - S m , R b - S r and A r - A r ages from the Pulur gneisses are regarded as cooling ages. Zircons from two gneiss samples from the Kazda~ Massif, dated by the stepwise Pb-evaporation method, gave an age of 308 • 16 Ma (Okay et al. 1996). To further refine the age of high-grade metamorphism in the Kazda~ Massif, we have dated a gneiss and an amphibolite from the Kazda~ Massif using the same method (Okay et al. 1996). Six zircon grains from the gneiss sample produced a relatively precise age of 319.2 ___ 1.5 Ma (early Late Carboniferous, latest Namurian), and one zircon grain from the amphibolite gave an age of 329 + 5 Ma (Fig. 7, Table 1). The isotopic data indicate high-grade metamorphism and associated deformation in the midCarboniferous (Namurian) in the Sakarya Zone. Permo-Carboniferous plutonism in the Sakarya Zone
Pre-Jurassic granitoids are common in the Sakarya Zone, although few are dated. The S6~tit granite in the western Sakarya Zone gave an A r - A r biotite plateau age of 290 • 5 Ma (CarboniferousPermian boundary, Okay et al. 2002) confirming earlier U - P b zircon and K - A r biotite ages (~o~ulu et al. 1965; ~o~ulu & Krummenacher 1967). K - A r biotite ages from the G6nen and Karacabey granites, east and west of Bandlrma, respectively, are
Table 1. Isotopic data from single-grain 2~ Lithology and sample number
Kazda~ gneiss K14{
Kazda~ amphibolite K4 ~amhk metagranite CL1
Grain 1 2 3 4 5 6 mean 1 t 1 / 2 t mean
/2~
in the range 286-298 Ma (Delaloye & Bing61 2000). These data indicate late orogenic acidic plutonism in the Sakarya Zone in the latest Carboniferous to early Permian period. The Variscan granites and high-grade metamorphic rocks in the Sakarya Zone were exhumed and unconformably overlain by the latest Carboniferous continental to shallow marine sedimentary rocks in the Eastern Pont• (Okay & Leven 1996; Okay & ~ahinttirk 1997; ~apklno~lu 2003) and in the Caucasus (e.g. Khain 1975).
Accretionary complexes in Anatolia: data on the spatial and temporal aspects of the Tethyan oceans There are widely differing views on the number, location, age span and name of the Tethyan oceans that existed during the Phanerozoic (e.g. Seng6r & Ydmaz 1981; Robertson & Dixon 1984; ~eng6r et al. 1984; Dercourt et al. 1986; Ricou 1994; Robertson et al. 1996; Stampfli et al. 2001, 2002). One way to approach this problem is through a biostratigraphic and isotopic study of the accretionary complexes. Because of their relatively low density the accretionary complexes have a wide preservation potential, and crop out widely in orogenic belts. The accretionary complexes may comprise three types of constituents: (1) pelagic sedimentary and basic magmatic rocks scraped at subduction
evaporation analyses of zircons from the basement of the Sakarya Zone, NW Turkey
Number of scans
2~176
Z~176
Mean value of 2~176 ratios
124 1l0 66 372 143 217
0.00210 0.00115 0.000255 0.000179 0.000115 0.000095
8.2 8.4 10.6 5.8 8.2 8.5
0.053178 + 102 0.052621 + 104 0.052407 • 130 0.052832 _+ 59 0.052792 • 90 0.052609 • 88
68 296 165
0.000071 0.000074 0.000336
7.7 12.0 12.1
0.053009 ___112 0.054640 _+ 37 0.054631 • 87
Errors are given at 95% confidence level and refer to the last digits.
Z~176 age (Ma) 336.4 • 4.4 312.5 + 4.5 303.2 • 5.7 321.6 • 2.5 319.9 • 3.9 312.0 _+ 3.8 319.2 • 1.5 329.2 • 4.8 397.6 • 1.5 397.2 +_ 3.6 397.5 • 1.4
OROGENS IN THE EASTERN MEDITTERANEAN zones from the downgoing oceanic crust; (2) greywacke and shale, which represent the trench infill; (3) blueschists and eclogites brought up along the subduction channel. The pelagic sedimentary rocks in the accretionary complexes provide an age range for the subducted ocean, whereas the greywackes and the isotopic age of the blueschists give an indication of the duration of subduction. The structural position of the accretionary complexes provides clues to the location of the associated oceans. Accretion ends by collision, or when the subduction zone is clogged by large oceanic or continental edifices, such as oceanic islands, oceanic plateaux, or isolated continental slivers (e.g. Cloos 1993). Termination of subduction in the Eastern Mediterranean south of Cyprus by the collision of the Eratosthenes Seamount provides a presentday example (e.g. Robertson 1998). The age of the accretionary complex can be defined as the age of subduction. At least four distinct accretionary complexes can be defined in the Balkan-Anatolian region. Karaburun-Chios
accretionary complex (Carboniferous)
Carboniferous accretionary complexes are found on the Karaburun peninsula and the adjacent island of Chios (Fig. 3; Stampfli et al. 1991, 2003). Both areas are situated in the Aegean on the northwestern margin of the Anatolide-Tauride Block immediately south of the Neotethyan Jzmir-Ankara suture. The complexes consist of strongly deformed siliciclastic turbidites, regarded as a Franciscan-type trench infill, which are unconformably overlain
397
by Lower Triassic basinal sedimentary and volcanic rocks (Robertson & Pickett 2000; Zanchi et al. 2003). The Lower Triassic pelagic sediments pass up into a typical Tauride carbonate platform of Triassic to Early Cretaceous age (Erdo~an et al. 1990). The intensely deformed and tectonically sliced and repeated turbidites comprise exotic limestone, radiolarian chert and volcanic blocks, up to kilometre scale, and Silurian-Carboniferous in age (Fig. 8; Kozur 1995). The age of the turbidite matrix is probably Early Carboniferous (Groves et al. 2003; Zanchi et al. 2003). Karakaya-Kiire accretionary complex ( T r i a s s i c - E a r l y Jurassic)
Triassic-Early Jurassic accretionary complexes are widely exposed in the Sakarya Zone below the Jurassic unconformity (Figs 3 and 4; Tekeli 1981; Ttiystiz 1990; Usta6mer & Robertson 1993, 1994; Pickett & Robertson 1996; Yflmaz et al. 1997; Okay 2000; Okay & G6nctio~lu 2004). In the western part of the Sakarya Zone they are attributed to the Karakaya Complex, and in the central Pontides to the Kiire Complex. Some Triassic palaeogeographical reconstructions show the Karakaya and Ktire accretionary complexes as belonging to different oceans separated by a continental sliver, attributed to the western part of the Istanbul Zone and to the northern parts of the Sakarya Zone (e.g. Usta6mer & Robertson 1993; Stampfli et al. 2001; Ziegler & Stampfli 2001). However, no coherent continental fragment can be defined
Fig. 8. A chronostratigraphic chart showing biostratigraphicand isotopic data from the Anatolian accretionary complexes. Data for the Karaburun complex are from Kozur (1995) and Groves et al. (2003); for the KarakayaKiire Complex from Kozur & Kaya (1994), Okay & Mosfler (1994), Kozur (1997), Okay & Moni6 (1997) and Okay et al. (2002); for the izmir-Ankara accretionary complexes from Bragin & Tekin (1996), Sherlock et al. (1999) and Tekin et al. (2002).
398
A.I. OKAYET AL.
between the outcrops of the Kiire and Karakaya complexes (Fig. 3). Furthermore, no Phanerozoic accretionary complex exists in the Istanbul Zone (Fig. 5). The Kiire and Karakaya complexes are similar in lithology, tectonostratigraphy and in structural position, but slightly differ in age, and will be treated together. The youngest palaeontological ages from the Karakaya Complex are latest Triassic (Leven & Okay 1996; Okay & Altmer 2004), whereas the age of the Kiire Complex extends to Early Jurassic (Kozur et aL 2000), and the complex is cut by granitoids of Mid-Jurassic age (Boztu~ et al. 1984; Yflmaz & Boztu~ 1986). Before the Cretaceous opening of the Black Sea, the Kiire Complex was contiguous with the Taurian Flysch of the Crimean Peninsula. The Karakaya-Ktire Complex consists of a lower metamorphic unit made up of a strongly deformed thrust stack of metabasitephyllite-marble with tectonic slices of ultramafic rock, broadly referred to as the Niliifer Unit. The depositional age of the Niltifer Unit, based on scarce conodonts in the marbles in NW Turkey, is Early to Mid-Triassic (Kaya & Mrstler 1992; Kozur et al. 2000). The geochemistry of the metabasites in the Niliifer Unit suggests a within-plate tectonic setting (Genq & Yalmaz 1995; Pickett & Robertson 1996, 2004; Genq 2004). The Niliifer Unit generally shows a high-pressure greenschist-facies metamorphism, although, in several localities in the Sakarya Zone, it also includes tectonic slices of blueschist and eclogite. The H P - L T metamorphic rocks in the Niliifer Unit are dated in the Bandlrma and Eski~ehir regions of NW Turkey (Fig. 3) as latest Triassic (205-203 Ma) using A r - A r method on phengites (Okay & Moni~ 1997; Okay et al. 2002). The structural setting and the lithological, metamorphic and geochemical features of the Niliifer Unit suggest an origin as an oceanic plateau or oceanic island, which was accreted to a Late Triassic active margin (Pickett & Robertson 1996, 2004; Okay 2000; Genq 2004). Recently, Topuz et al. (2004b) reported Early Permian (263-260Ma) R b - S r and A r - A r hornblende and muscovite ages from a metabasite-phyllite sequence from the Pulur region in the Eastern Pontides (Fig. 4). The metabasite-phyllite sequence, which is correlated with the Niliifer Unit, is tectonically overlain by the granulite-facies gneisses of mid-Carboniferous age (Okay 1996; Topuz et al. 2004b). If these isotopic data are confirmed then the subduction-accretion represented by the KarakayaKtire Complex will extend back to the Early Permian (Fig. 8). In the Sakarya Zone, the Niltifer Unit is overlain by Triassic to Lower Jurassic siliciclastic and volcanic sequences, which were strongly deformed, probably in a subduction zone setting, in the latest Triassic-earliest Jurassic (Okay 2000). In NW Turkey, the siliciclastic rocks comprise olistostromes with numerous Carboniferous and Permian shallow marine limestone blocks (Leven 1995; Leven & Okay 1996), and smaller numbers of Middle Carboniferous (Bashkirian), Permian and Triassic radiolarian chert and pelagic limestone exoticblocks (Fig. 8; Kozur & Kaya 1994; Okay & Mostler 1994; Kozur 1997; Kozur et aL 2000; Grnctio~lu et al. 2004). Olistostromes with the Carboniferous and Permian shallow marine limestone blocks form a belt, over 150km long and 5 - 1 0 k m wide, in NW Turkey immediately NW of the Izmir-Ankara suture (Fig. 3). The origin of the Permo-Carboniferous limestone blocks is controversial; the fauna in the blocks is interpreted either as Laurussian (Leven & Okay 1996) or as Gondwanan in origin (Altiner et al. 2000). Strandja accretionary complex (Late T r i a s s i c - E a r l y Jurassic)
The Triassic-Middle Jurassic epicontinental sediments of the Strandja Massif are tectonically overlain by a highly deformed volcano-sedimentary complex of siliciclastic turbidites, carbonates, mafic and acidic volcanic rocks of Early to Late Triassic
age (Chatalov 1980, 1988). Sengrr et al. (1984) and Ustarmer & Robertson (1993) interpreted the Strandja allochthons as an accretionary complex, although definite evidence for the oceanic origin of the Strandja allochthons (e.g. ultramafic rocks or deep-sea radiolarian cherts) is missing. Dismembered ophiolites of Late Jurassic age, including peridotite, gabbro and basalt, occur on the eastern margin of the Rhodope Massif (Fig. 3; Tsikouras & Hatzipanagiotou 1998), and are associated with a slightly metamorphosed epicontinental sequence of Triassic to Early Jurassic age (Kopp 1969). The ophiolites of this Circum-Rhodope Zone may represent the root zone of the Strandja allochthons. Although the Karakaya-Ktire and the Strandja accretionary complexes are similar in age, they differ markedly in their structural setting and lithology, and, as discussed below, are ascribed to different oceans. i z m i r - A n k a r a accretionary complex (Late Cretaceous)
Late Cretaceous accretionary complexes cover large regions in the Anatolide-Tauride Block south of the Izmir-Ankara-Erzincan suture (Figs 3 and 4; Okay 2000). They generally form tectonic imbricates sandwiched between the ophiolites above and the Anatolide-Tauride carbonate platform below. In many regions near the izmir-Ankara suture, such as north of Eski~ehir (Okay et al. 2002), east of Ankara (Koqyi~it 1991) and in the Tokat Massif (Bozkurt et al. 1997) the Izmir-Ankara accretionary complexes are imbricated with those of the Karakaya-Ktire Complex (Figs 3 and 4). The Late Cretaceous accretionary complexes consist mainly of basalt, radiolarian chert, pelagic shale and pelagic limestone, and in the old literature were often referred to as ophiolitic mrlange or coloured mrlange. In the north near the Izmir-Ankara suture, the accretionary complexes have undergone low-grade blueschistfacies metamorphism dated at c. 80 Ma (Sherlock et aL 1999). Palaeontological study of radiolarian cherts and pelagic limestones in these accretionary complexes has shown the presence of Triassic, Jurassic and Cretaceous rocks (Fig. 8; Bragin & Tekin 1996; Tekin et al. 2002). In contrast, no Palaeozoic pelagic sedimentary rocks were described in the Izmir-Ankara accretionary complexes, the oldest ones being Late Triassic (Late Carnian) radiolarian cherts (Tekin et al. 2002). Age and location o f the Tethyan oceans
Biostratigraphic and isotopic data from the accretionary complexes in Anatolia indicate the presence of several Tethyan oceans. A Tethyan ocean north of the Anatolide-Tauride Block, called the izmir-Ankara ocean, had a minimum age span from Mid-Late Triassic to the Cretaceous, suggesting an opening as early as the Early Triassic. This is in accord with the Triassic stratigraphy from the Karaburun Peninsula, which is indicative of rifting in the Early Triassic (Robertson & Pickett 2000). The widespread unconformity at the base of the Lower Triassic rocks in the northern margin of the Anatolide Tauride Block (e.g. Eren 2001; G6ncfio~lu et al. 2003) is probably also related to shoulder uplift before the rifting. The outcrop pattern of the Karakaya-Kiire Complex indicates an ocean in a similar position to the lzmir-Ankara ocean (e.g. north of the Anatolide-Tauride Block and south of the Istanbul Zone) but older (at least Mid-Carboniferous to Early Jurassic). The Karakaya-Kiire ocean must also have been located south of the Variscan basement of the Sakarya Zone, as the exotic blocks in the Karakaya-Kiire Complex were most probably derived from the south rather than the north (Okay 2000). The absence of continental fragments between the Karakaya-Kiire and Izmir-Ankara accretionary complexes (Figs 3 and 4) implies that the Karakaya-KiJre ocean corresponds to the main Palaeotethys (Fig. 9), rather than to a small back-arc basin as
OROGENS IN THE EASTERN MEDITTERANEAN
399
Fig. 9. Palaeogeographicalreconstructions of the Tethyan realm for the Late Carboniferous (a), Early Triassic (b), Late Triassic (c) and Late Jurassic (d), showingthe possible locations of the terranes and oceans discussed in the text. The general palaeogeographicalframework is taken from Stampfliet al. (2001).
shown in most models (e.g. ~engrr et aL 1984; Stampfli et aI. 2001). Biostratigraphic data from the Karaburun-Chios accretionary complexes suggest an ocean of Silurian to Carboniferous age, again situated north of the Anatolide-Tauride Block. The Karaburun-Chios accretionary complex may be related to the Variscan subduction (Zanchi et al. 2003), in which case it must have been displaced eastwards from its original position by strike-slip faulting (Fig. 9a). The Strandja allochthons indicate the presence of a Triassic to Early Jurassic ocean between the Strandja Massif in the north and the Rhodope-Serbo-Macedonian massifs in the south (Fig. 3). The Triassic stratigraphy of the Istanbul Zone indicates rifting in the Early Triassic. This ocean probably formed an eastern extension of the Hallstatt-Meliata ocean, described
farther west in the Eastem Alps and the Balkans (Kozur 1991; Channel & Kozur 1997). The relation between these Tethyan oceans and the surrounding continental terranes, which gave rise to the orogenic events, is discussed below.
Variscan orogeny in the Balkans and the Black Sea region The Late Carboniferous orogeny in the Pontides forms a link between the Variscan orogen in Central Europe and the Uralides of Eastern Europe. The Variscan orogeny comprises Carboniferous to Early Permian deformation, metamorphism and magmatism linked to the collision and amalgamation of Gondwana, Laurussia and the intervening terranes (e.g. Matte 2001; Wart 2002). The
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eastward extension of the Variscan orogen towards the Balkans and Anatolia is obscured by the strong overprint of the Alpide orogeny, or is concealed by the younger cover. The East European Craton is bordered in the south by a narrow tectonic belt, called the Scythian Platform, which is generally considered as a Late Palaeozoic (Early Carboniferous) orogen (e.g. Nikishin et al. 1998, 2001). The Palaeozoic stratigraphy in the Scythian Platform is concealed beneath the Mesozoic and younger strata, and there are only patchy data from a few boreholes, which indicate a thick Lower Devonian continental sandstone succession overlain by Middle Devonian to Lower Carboniferous shallow marine limestones. The overlying Vis6an-Namurian sequence consists of paralic and limnic deposits, and the Permian of red clastic rocks (Vaida & Seghedi 1997). During the Devonian and Carboniferous the Istanbul Zone and the Moesian Platform were adjacent to the Scythian Platform, and formed a south-facing passive continental margin (Figs 6 and 9a). The western part of the Istanbul Zone was the site of deep marine sedimentation in the Devonian and Carboniferous, in a continental slope setting, and hence was closer to the ocean compared with its eastern part and the Scythian-Moesian platforms (Fig. 6). The Late Carboniferous deformation in the Istanbul Zone is coeval with the high-grade metamorphism in the Sakarya Zone. It is plausible to relate the deformation and regional metamorphism to collision of the Laurussia margin with an ensialic arc represented by the basement rocks of the Sakarya Zone and possibly of the Strandja Zone. Absence of Palaeozoic magmatism in the MOIS terrane suggests southward subduction, which is compatible with the general north to NE vergence of Carboniferous deformation in the Istanbul Zone (G6rtir et al. 1997) and the northward migration of the coal deposition in the Moesian Platform (Tari et al. 1997). The ocean between the Laurussia margin and the Sakarya-Strandja microplate probably started to close by the Early Devonian, producing a magmatic arc represented by the ~amlkk Granite in the Sakarya Zone. The Late Carboniferous collision was followed by the latest Carboniferous-Early Permian plutonism in the core of the orogen in the Strandja and Sakarya zones, possibly linked to crustal thickening. The latest Carboniferous-Early Permian molasse deposition in the Eastern Pontides and the Caucasus marks the end of the Variscan orogeny in northern Turkey. The Intra-Pontide suture between the Istanbul and Sakarya zones probably links up with the Late Carboniferous Rheic suture in Central Europe (Fig. 9a; Ziegler & Stampfli 2001). The Variscan evolution of northern Turkey and the Balkans appears to be similar to that of NW Europe, with the Istanbul-Moesia-Scythian Block corresponding to Avalonia, and the Sakarya-Strandja zones to Armorica (Stampfli et al. 2002; Winchester & The PACE TMR Network Team 2002).
Early Triassic rifting and magmatism The Early Triassic is characterized by widespread tiffing and mafic magmatism in the Eastern Mediterranean region, possibly associated with mantle plumes (e.g. Dixon & Robertson 1999). The Istanbul Zone started to rift from the Sakarya Zone along the former Carboniferous suture, as shown by the deposition of earliest Triassic continental sandstones and conglomerates intercalated with basaltic flows (Fig. 9b). In the Mid-Triassic the Istanbul Zone became separated from the Sakarya Zone, as the rift turned into the Intra-Pontide-Meliata ocean. On the Gondwana side in the south, mafic magmatism was associated with the break-up of Permo-Carboniferous carbonate platforms, and the separation of the Anatolide-Tauride Block from Gondwana (Fig. 8b). Possibly a thin carbonate sliver, corresponding to the Cimmerian continent of ~eng6r et al. (1984), rifted away from the Anatolide-Tauride Block in the Early Triassic. Associated with this rifting, major intra-plate mafic magmatism occurred and an abnormally thick oceanic crust or oceanic plateau was created adjacent to the
passive continental margin. The northward drift of this narrow continental sliver is shown to close the Palaeozoic Tethys and open up the Mesozoic Tethys in the Triassic (Fig. 9b and c), although as discussed below there is no unequivocal evidence for the Cimmerian continent in the Pontides.
Cimmeride orogeny in the Pontides In Turkey deformation and metamorphism of latest Triassic to earliest Jurassic age is particularly marked in the Sakarya Zone. It is associated with the emplacement of large oceanic allochthons over the Variscan basement. In contrast, the Cimmeride deformation is weak, and the Cimmeride metamorphism is absent in the other Pontic zones, where this period is generally marked as an unconformity (Fig. 2). The cause of the Cimmeride orogeny in Anatolia was generally thought to be the collision and amalgamation of a Cimmerian continent with the Laurasian margin (e.g. ~eng6r 1984; ~eng6r et al. 1984). However, it has not been possible to define a Cimmerian continent in the field, which would have been readily recognized by its Gondwana-type stratigraphy, free of Variscan deformation and metamorphism. In many regions along the Izmir-Ankara suture the accretionary complexes of the Izmir-Ankara and Karakaya-Ktire oceans are tectonically intercalated with no evidence of an intervening continental fragment (Figs 3 and 4; Bozkurt et al. 1997; Okay et al. 2002). Apparently, the narrow Cimmerian continental sliver, responsible for the opening of the Izmir-Ankara ocean, was completely subducted, with only its Permo-Carboniferous limestone cover providing blocks to the accretionary complex. The Cimmeride orogeny in the Pontides was largely accretionary, caused by the collision and partial accretion of an oceanic plateau to the Laurasian margin during the latest Triassic (Fig. 9c; Okay 2000). This is compatible with the short duration of deformation and regional metamorphism observed in the Karakaya-Ktire Complex.
Late Jurassic Balkan orogeny Apart from the Strandja Massif, Late Jurassic deformation is strangely absent, or is marked by only a slight disconformity in the Pontic zones. The Late Jurassic was a period of opening of the Alpine Tethys in the west, where contractional deformation is also not reported. This leaves a relatively small space for the Balkan orogeny on the southern margin of Laurasia (Fig. 9d). The Balkan orogeny is possibly linked to the subduction of the Intra-Pontide-Meliata ocean between the Strandja and the Rhodope-Serbo-Macedonian massifs, and the ensuing collision (Fig. 9d). The north-vergent deformation in the Strandja Massif indicates a southward subduction under the Rhodope Massif, with the implication that the latest Jurassic-Early Cretaceous granitoids in the Serbo-Macedonian Massif were generated in a magmatic arc. The eastern part of the Intra-Pontide-Meliata ocean between the Sakarya and Istanbul zones did not close until the mid-Cretaceous, suggesting the existence of a transform fault between Pelagonia and the Sakarya Zone (Fig. 9d).
Conclusions The Pre-Alpide geological history of the Eastern MediterraneanBalkan region can be viewed as the episodic growth of Laurussia by the accretion of oceanic and continental terranes, interrupted by the opening of narrow back-arc basins on the southern margin of Laurussia. The continental terranes were invariably derived from Gondwana, and were accreted to Laurussia during the Late Ordovician-Early Silurian and Late Carboniferous,
OROGENS IN THE EASTERN MEDITTERANEAN
whereas a major accretion of oceanic crustal material occurred during the Late Triassic-Early Jurassic. Orogenic deformation associated with the Late OrdovicianEarly Silurian accretion of the Istanbul-Moesia-Scythian Platform (the MOIS terrane) is buried under young cover in the northern margins of the Black Sea. The Carboniferous accretion of the Strandja-Sakarya terrane to the Laurussian margin, along a south-dipping subduction zone, resulted in strong deformation, mid-Carboniferous metamorphism, and latest CarboniferousEarly Permian post-orogenic plutonism. The ensuing suture is probably an extension of the Rheic suture in Central Europe (Ziegler & Stampfli 2001). In contrast to these Palaeozoic continental collisions, a major accretion of oceanic crustal rocks occurred during the Late Triassic-Early Jurassic. The accretionary complexes in the Pontides comprise voluminous metabasic rocks with latest Triassic blueschist and eclogite ages. The Late Jurassic deformation and metamorphism, observed only in the Balkans, were the result of the closure of a narrow back-arc basin, the Meliata ocean between the Rhodope-SerboMacedonian and the Strandja massifs. In contrast to the polyorogenic history of the Pontides and the Balkans, the Anatolide-Tauride Block south of the I z m i r - A n k a r a suture was largely free of Palaeozoic-early Mesozoic deformations, except along its northwestern margin, where a Carboniferous accretionary complex has been recognized (Stampfli et al. 1991). Biostratigraphic and isotopic data from the Anatolian accretionary complexes and their structural position indicate the presence of three oceanic realms north of the Anatolide-Tauride Block during the Phanerozoic. Two of them correspond to the mid-Palaeozoic-Early Jurassic Palaeotethys, and Early Triassic-Tertiary Neotethys, respectively, both of which were subducted along the I z m i r - A n k a r a suture, which represents the main boundary between Laurussia and Gondwana. The third ocean, the Meliata-Intra-Pontide ocean, opened as a marginal back-arc basin on the Laurussian margin (e.g. Stampfli 2000). A major point from this review and that is also implicit in some recent studies (e.g. Dean et al. 2000) is the mobility of the small plates that make up Anatolia. The assumption of conjugate margins, common in the old Tethyan reconstructions (e.g. ~engrr & Ydmaz 1981; Robertson & Dixon 1984) is clearly not correct. Prior to the Tertiary, the Pontides and the Anatolide-Tauride Block never formed a single contiguous terrane, which is implicit in recent palaeogeographical reconstructions (e.g. Stampfli et al. 2001). Translation of continental terranes oblique to the rifted margin, and margin-parallel strike-slip faulting led to the juxtaposition of unrelated continental fragments. For example, in the Early Ordovician both the Anatolide-Tauride Block and the Istanbul Zone were probably located on the northern margin of Gondwana but separated by several thousand kilometres (see Dean et al. 2000). The sutures separating the terranes in the Eastern MediterraneanBalkan region were major zones of weaknesses, and were rejuvenated at various times. For example, the Intra-Pontide suture started as a Late Carboniferous suture, and later became the site of an Early Triassic rift, which developed into the MeliataIntra-Pontide ocean. This ocean closed in the Late Jurassic-Early Cretaceous, generating a second suture. In the Miocene the suture was reused by the North Anatolian Fault, which at present defines the northern margin of the Anatolia microplate. This study was partly funded by the Turkish Academy of Sciences. We thank J. Winchester, A. Saintot and L. Jolivet for constructive and helpful comments on the manuscript.
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Tectonic processes in the Southern and Middle Urals: an overview D. B R O W N 1, V. P U C H K O V 2, J. A L V A R E Z - M A R R O N 1, F. BEA 3, & A. P E R E Z - E S T A U N t
Xlnstitute of Earth Sciences 'Jaume Almera', CSIC, c/Llufs Sold i Sabarfs s/n, 08028 Barcelona, Spain (e-mail: dbrown @ija. csic. es) 2Ufimian Geoscience Center, Russian Academy of Sciences, ul. Karl Marx 16/2, Ufa 45000 Bashkiria, Russia 3Department of Mineralogy and Petrology, Fuentenueva Campus, University of Granada, 18002 Granada, Spain
Abstract: The tectonic evolution of the Uralide orogen began during the Late Palaeozoic as the continental margin of Baltica entered an east-dipping (today's coordinates) subduction zone beneath the Magnitogorsk and Tagil island arcs. The subsequent arc-continent collision resulted in the development and emplacement of an accretionarycomplex over the continental margin, the development and deformation of a foreland basin, and the extrusion of high-pressure rocks along the arc-continent suture. There is mounting evidence that, at about the same time as arc-continent collision was occurring along this margin of Baltica, eastward-directed subcontinental subduction of the Uralian oceanic crust was also taking place beneath the Kazakhstan plate. This subcontinental subduction is thought to have resulted in the formation of a continental volcanic arc. The final closure of the Uralian ocean basin and the start of collision between the Baltica and Kazakhstan plates occurred during the Late Carboniferous. This continent-continent collision resulted in development of the Late Carboniferous to Early Triassic western foreland fold and thrust belt and foreland basin of the Uralides. The foreland fold and thrust belt displays a large amount of basement involvement, extensive reactivation of pre-existing faults, and a small amount of shortening. At the same time, widespread strike-slip faulting accompanied by melt generation and granitoid emplacement took place in the interior part of the Uralides, leading to the transfer of material laterally along the strike of the orogen. The final crustal structure of the Uralides that resulted from the combination of all of these tectonic events is bivergent, with a crustal root reaching c. 53 km depth.
Extending for nearly 2500 km from near the Aral Sea in the south to the islands of Novaya Zemlya in the Arctic Ocean, the Uralide orogen of Russia marks the eastern boundary of the Early Palaeozoic continent Baltica and its collision zone with the Siberian and Kazakhstan plates during the Palaeozoic assembly of Pangaea. For descriptive purposes the Uralides have traditionally been divided into a number of longitudinal zones (Fig. la) that are largely based on the ages and palaeogeography of the dominant rocks within them (e.g. Ivanov et al. 1975; Khain 1985; Fershtater et al. 1988; Puchkov 1997). From west to east these zones are the Pre-Uralian zone, the West Uralian zone, the Central Uralian zone, the Magnitogorsk-Tagil zone, the East Uralian zone and the Trans-Uralian zone. Additionally, the Uralides have been divided geographically into the Southern, Middle, Northern, CisPolar and Polar Urals. The Pre-Uralian, West Uralian and Central Uralian zones contain syntectonic Late Carboniferous to Early Triassic sediments of the foreland basin, Palaeozoic platform and continental-slope rocks, and Archaean and Proterozoic rocks of the East European Craton (part of Baltica). These three zones were affected by Uralide deformation and make up the foreland thrust and fold belt (e.g. Kamaletdinov 1974; Brown et al. 1997b). The Magnitogorsk-Tagil zone consists of Silurian to Devonian intra-oceanic island arc volcanic rocks and overlying volcaniclastic sediments. The Magnitogorsk-Tagil zone is sutured to the former continental margin of Baltica along the Main Uralian fault. The East Uralian zone is composed predominantly of deformed and metamorphosed volcanic arc fragments with minor amounts of Precambrian and Palaeozoic rocks thought to represent continental crust (Puchkov 1997, 2000; Friberg et al. 2000b). The East Uralian zone was extensively intruded by Carboniferous and Permian granitoids (Fershtater et al. 1997; Bea et al. 1997, 2002), forming the 'main granite axis' of the Uralides. The East Uralian zone is juxtaposed against the Magnitogorsk-Tagil zone along the East M a g n i t o g o r s k - S e r o v - M a u k fault system. The Trans-Uralian zone is composed of Carboniferous volcano-plutonic complexes (Puchkov 1997, 2000). Ophiolitic material and high-pressure rocks have also been reported (Puchkov 2000). The contact between the East Uralian and Trans-Uralian zones is exposed only in the Southern Urals, where it is a serpentinite mrlange. Rocks that unequivocally belong to either the Kazakhstan or Siberia plates do not crop out in the Uralides.
It is generally accepted that the tectonic evolution of the Uralides (Hamilton 1970; Zonenshain et al. 1984, 1990; Puchkov 1997, 2000; Brown & Spadea 1999; Alvarez-Marron 2002; Bea et al. 2002) began with the development of intra-oceanic island arcs in the palaeo-Uralian ocean, which were then accreted to the margin of the East European Craton. Meanwhile, subcontinental subduction is thought to have been taking place along the margin of the Kazakhstan plate, forming Andean-type arcs. The Uralian orogeny began in the latest Carboniferous as the Uralian ocean basin closed and the Kazakhstan plate, followed by the Siberia plate, collided with Baltica. Continent-continent collision continued until the Early Triassic. With the exception of minor Triassic transtension, intra-plate volcanism, erosion and basin inversion during the development of the West Siberian Basin, the Uralide orogen has been preserved, relatively intact, since the Permian, providing an ideal place to study Palaeozoic orogenic processes. The aim of the paper is to summarize a number of the key tectonic processes that formed the Southern and Middle Urals (Fig. lb). It begins with the earliest recognizable event and progresses through time to the final crustal structure that is observable today. Emphasis is placed on two transects, which are focused around two deep seismic surveys, EUROPROBE's Seismic Reflection Profiling in the Urals (ESRU) survey in the Middle Urals and the multicomponent Urals Seismic Experiment and Integrated Studies (URSEIS) survey in the Southern Urals (Fig. lb).
Tectonic units and processes
Arc-continent collision (Mid-Devonian to Early Carboniferous) The Tagil and Magnitogorsk volcanic arcs developed during the Silurian-Devonian (Tagil) and the Early Devonian-Early Carboniferous (Magnitogorsk) in an intra-oceanic setting (Seravkin et al. 1992; Yazeva & Bochkarev 1996; Spadea et al. 1998, 2002; Brown & Spadea 1999; Herrington et al. 2002) and began to collide with the margin of Baltica in the late Mid-Devonian (Magnitogorsk) and the Early Carboniferous (Tagil) (Puchkov 1997; Brown & Spadea 1999). The Tagil arc, in the Middle Urals, is made up of Silurian andesitic basalts and Lower Devonian trachytes and volcaniclastic rocks, overlain by 2000 m of Lower
From: GEE,D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 407-419. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. (a) Map showing the zones of the Urals and its geographical divisions from north to south. The area discussed in this paper is indicated by the box. (b) Geological map of the Southern and part of the Middle Urals. The legend shows the disposition of the various tectonic units discussed in this paper. The locations of the cross-sections in Figures 2 and 5 are shown, as is the location of Figure 6 and the ESRU and URSEIS seismic profiles.
and Middle Devonian limestone that, in the east, is intercalated with calc-alkaline volcanic rocks (Antsigin et al. 1994; Yazeva & Bochkarev 1994). The Tagil arc has been deformed and folded into an open synformal structure (e.g. Bashta et al. 1990; Ayarza et al. 2000b) and has been metamorphosed to lower greenschist facies. By far the best preserved and exposed, and therefore the most studied of the Uralide arcs, is the Magnitogorsk arc in the Southern Urals. It is composed of Emsian boninite-bearing arc-tholeiites in the forearc region, followed by Emsian to Givetian arc-tholeiite to calc-alkaline volcanic rocks of the Irendyk volcanic front; all of which display a clear intra-oceanic island arc signature (Fig. 2a; Seravkin et al. 1992; Spadea et al. 1998, 2002; Brown & Spadea 1999; Herrington et al. 2002). These volcanic units form the basement on which up to 5000 m of Frasnian- to Famennian-age forearc basin volcaniclastic sediments were deposited (Fig. 2a; Maslov et al. 1993; Brown et al. 2001). Lower Carboniferous shallow-water carbonates and, locally, basalt-rhyolite volcanic rocks unconformably overlie the arc edifice. Locally, Lower Carboniferous granitoids intrude the arc. Deformation in the Magnitogorsk volcanic arc is low, with only minor open folding and thrusting (Brown et al. 2001). The metamorphic grade barely exceeds sea-floor metamorphism. In the Southern Urals, a well-preserved accretionary complex developed during the Magnitogorsk arc-continent collision (Figs lb and 2b) (e.g. Bastida et al. 1997; Brown et al. 1998;
Brown & Spadea 1999; Alvarez-Marron et al. 2000). The accretionary complex is composed of Silurian to Middle Devonian continental slope and platform sedimentary rocks (Suvanyak Complex) that were detached from the East European Craton, and were overthrust by c. 5 km of late Frasnian and Famennian syncollisional volcaniclastic turbidites (Zilair nappe) sourced predominantly from the accretionary complex to the east with minor input from the Magnitogorsk arc (e.g. Puchkov 1997; Brown et al. 1998; Brown & Spadea 1999; Alvarez-Marron et al. 2000; Willner et al. 2002) (Figs lb and 2b). These units are flanked to the east by eclogite- and blueschist-bearing gneisses of the Maksutovo Complex that record a peak metamorphic pressure and temperature of 20 _+ 4 kbar and 550 __ 50 ~ (Beane et al. 1995; Hetzel et al. 1998; Schulte & Blfimel 1999), and a peak metamorphic age of c. 380-370 Ma (Fig. 2c; Matte et al. 1993; Lennykh et al. 1995; Beane & Connelly 2000; Glodny et al. 2002). Recently, microdiamond aggregates have been described from the Maksutovo Complex, suggesting that even higher pressures were achieved than those recorded by the metamorphic mineral assemblages (Bostick et al. 2003). The highest structural level of the accretionary complex is the Sakmara Allochthon in the south and the Kraka lherzolite massif in the north. The accretionary complex is at present sutured to the Magnitogorsk arc along the east-dipping Main Uralian fault zone, a m~lange that contains several kilometre-scale ultramafic fragments, one of which records
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Fig. 2. (a) Geochemicaland isotope data for Magnitogorskextrusive rocks. Plots of Emsian age-corrected Nd and Sr isotope ratios for Baimak-Buribai, Irendyk and Karamalytashformationsshow depleted mantle sources and secondary radiogenic Sr enrichment.Th/Yb v. Ta/Yb plot shows mostly intraoceanic arc affinities.A stratigraphiccolumn for forearc basin stratigraphyis also shown. After Brown & Spadea (1999). (b) Upper crustal cross-sectionacross the Magnitogorskforearc and the accretionary complex showing the structural architectureof the arc-continent collision zone in the SouthernUrals (after Mvarez-Marron et al. 2000). The location of the section is shown in (a). (e) Radiometric age determinationsfrom the Mindyakand Maksutovo complexes and a P - T path of the lower unit of the Maksutovo Complex. Data are taken from the sources discussed in the text. The upper path is for a garnet-mica schist and the lower path for an eclogite. The open arrows indicate a generalized retrograde path.
metamorphism under mantle conditions (Savelieva & Nesbitt 1996; Savelieva et al. 1997, 2002; Scarrow et al. 1999). The geochemistry of the Magnitogorsk arc volcanic rocks (Spadea et al. 1998, 2002; Herrington et al. 2002), the structure of the accretionary complex and its forearc (Brown et al. 1998, 2001; Alvarez-Marron et al. 2000), the high-pressure rocks beneath and along the suture zone (e.g. Hetzel et al. 1998; Hetzel 1999; Beane & Connelly 2000; Brown et al. 2000), and the ophiolitic, mafic and ultramafic material (Savelieva et al. 1997, 2002; Scarrow et al. 1999) show that the Palaeozoic tectonic processes that went into its formation can be favourably compared with those in currently active settings such as the west Pacific (Fig. 3a; Puchkov 1997; Brown et al. 1998; Brown & Spadea 1999; Herrington et al. 2002; Spadea et al. 2002). For example, boninitic lavas found in the oldest arc volcanic units provide a geodynamic marker that records the initiation of intra-oceanic subduction and the early development of the arc (Spadea et al. 1998; Brown & Spadea 1999). High-pressure rocks along the backstop of the accretionary complex were in part derived from continental margin material (Hetzel 1999), and the Mid-Devonian age of the high-pressure metamorphism provides a constraint for determining the timing of the entry of the continental crust into the subduction zone (Brown et al. 1998). The pressure,
temperature and thermochronology of the Maksutovo Complex and other high-pressure rocks along the arc-continent suture provide evidence for the flux of material in the subduction zone channel during its evolution (Fig. 3b; Brown et al. 2000). The sediments overlying the volcanic arc record (near) surface processes such a growth folding (Brown et al. 1998, 2001; Alvarez-Marron et al. 2000). The widespread occurrence of debris flows within the Late Devonian Zilair formation is thought to represent seismic events (seismites), and may be related to the arrival of the full thickness of the continental crust at the subduction zone (Brown et al. 2001). The accretionary complex was subsequently reworked during the formation of the foreland fold and thrust belt (see below).
S u b d u c t i o n b e n e a t h the K a z a k h s t a n p l a t e (Late D e v o n i a n to Late Carboniferous)
To date, little is known about what happened along the margin of the Kazakhstan plate prior to or during its collision with Baltica, as no rocks that can be unequivocally assigned to its plate margin have been recognized in the Uralides. Nevertheless, some recent studies suggest the presence of a continental volcanic arc that may have developed on the active margin of the Kazakhstan
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Fig. 3. (a) The early convergent history in the Southern Urals is marked by the generation of boninite-bearing arc-tholeiites in the Magnitogorsk forearc (T1), followed by arc-tholeiite to calc-alkaline volcanism. With the entry of the East European Craton continental crust into the subduction zone, volcanism waned and stopped, and high-pressure metamorphism of its leading edge took place (T2). The arrival of the full thickness of the continental crust at the subduction zone is marked by increased sedimentation in the forearc basin and deposition of arc-derived volcaniclastic turbidites across the subducting slab (T3). These, together with offscraped continental material, the exhumed high-pressure rocks, and a lherzolite massif, formed an accretionary wedge. A broad m61ange zone containing ultramafic fragments separates the forearc basement from the accretionary wedge, and marks the damage zone that developed along the backstop region. From Brown & Spadea (1999). (b) T l: during the Early Devonian suprasubduction-zone material was subducted to upper mantle depths. T2: by the end of the Early Devonian, when the East European Craton appears at the subduction zone, steady-state intra-oceanic subduction was under way. Geotherms are from van den Beukel (1992). T3: with the entrance of the East European Craton into the subduction zone the thermal regime would have departed from steady state. Geotherms are from van den Beukel (1992) for a continental heat flow of 70 mW m -2. The dotted line indicates van den Beukel's continental crust, whereas we have chosen to show a thinned continental crust (dark grey). The lowest frame shows an enlargement of the area shown in the box in T3. When the downgoing slab had reached a depth of 50-70 km, the Proterozoic sediments with a quartz rheology were detached, interacted with the mantle wedge, and the exhumation history began.
plate. In particular, data f r o m granitoids in the East Uralian zone point in this direction (Bea et al. 2002). A n u m b e r o f Uralide granitoids formed in what is thought to be two subduction settings from the Late D e v o n i a n to Late Carboniferous (Fig. 4; B e a et al. 1997, 2002; Montero et al. 2000). The first subduction-related m a g m a t i s m occurred from about 370 M a to 350 Ma, and is found in the eastern sector o f the East Uralian zone. B e a et al. (2002) interpreted this phase of m a g m a t i s m to have b e e n related to an east-dipping subduction zone located to the east o f the accreted Magnitogorsk arc, and to have produced I-type granitoids such as those o f the C h e l y a b y n s k and the Chernorechensk batholiths. A n older continental c o m p o n e n t in these granitoids can be interpreted to be the result o f their formation on the continental margin o f the Kazakhstan continent (Bea et al. 2002). A second phase of subduction m a g m a t i s m occurred from about
335 M a to 315 Ma, and is found in the western part of the East Uralian zone, between 55~ and 58~ (Bea e t al. 2002). B e a et al. (2002) have interpreted this phase to have been related to a subduction zone located to the east of the accreted Tagil arc, and that dipped eastward underneath the older continental arc. This subduction event produced batholiths c o m p o s e d of I- and M-type granitoids with little, if any, continental component. Magmatic activity directly related to subduction e n d e d before the Permian. Friberg et al. (2000b) h a v e described mafic to felsic gneisses o f largely Silurian and D e v o n i a n age (note, h o w e v e r , that there are large errors on the age determinations) and v o l c a n o - s e d i m e n t a r y rocks in the East Uralian zone that have b e e n interpreted to represent a volcanic arc complex. It is into this arc c o m p l e x that the above-discussed granitoids intrude, suggesting that the gneisses m a y be a deep, m e t a m o r p h o s e d part of the arc. In the
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Fig. 4. (a) Schematic map of the Southern and Middle Urals, outlining the late orogenic strike-slip fault system and the location of subduction-type granitoids. (b) Continental-crust normalized REE plots of Early Carboniferous subduction-related granitoids (from Beaet al. 2002). These subduction granites are enriched in trace elements of continental affinity such as Rb, Ba, Th, U and Li, suggesting that the protolith was composed of oceanic materials plus a significant fraction of old crustal materials. (e) eNa(t) V. esr(t) of Uralide subduction granitoids (from Beaet al. 2002). Neither 87Sr/S6Sr(t) nor 143Nd/144Nd(t) values bear any relation to the age, but depend on the geographical longitude. The Early Carboniferous batholiths in the east, at Chelyabinsk and Chernorechensk, are composed of granitoids with significantly higher 87Sr/86Sr(t) but lower 143Nd/144Nd(t)than similar rocks of the Late Carboniferous batholiths in the west, which have identical (in some cases more primitive) 87Sr/86Sr(t) and only slightly lower 143Nd/144Nd(t) compared with oceanic plagiogranites.
Southern Urals, however, the East Uralian zone is primarily composed of amphibolite-facies metapelites that are thought to represent continental crust (e.g. Puchkov 1997, 2000). The presence of Early Carboniferous subduction-related granitoids in this zone may indicate that the continental crust was part of the Kazakhstan plate at some stage. Finally, the eastern parts of the URSEIS and ESRU seismic reflection profiles image west-dipping reflectivity throughout the crust of the Trans-Uralian zone which has been interpreted to possibly represent east-vergent structures related to imbrication along the margin of the Kazakhstan plate (Tryggvason et al. 2001; Brown et al. 2002). The f o r e l a n d f o l d a n d thrust belt (Late C a r b o n i f e r o u s to E a r l y Triassic)
The foreland fold and thrust belt of the Middle and Southern Urals (which includes the Pre-Uralian, West Uralian and Central Uralian zones) contains syntectonic Late Carboniferous to Early Triassic sediments of the foreland basin, Palaeozoic platform and slope sediments, the Archaean and Proterozoic basement of Baltica, and the a r c - c o n t i n e n t collision accretionary complex. The foreland fold and thrust belt developed from the Late Carboniferous to the Late P e r m i a n - E a r l y Triassic (Kamaletdinov 1974; Brown et al. 1997b; Puchkov 1997). The foreland fold and thrust belt
between c. 56~ and 59~ is a narrow, north-south-trending, west-verging basement-involved thrust stack measuring c. 5 0 75 km in width from the Main Uralian fault (the a r c - c o n t i n e n t suture) to the frontal folds (Fig. 1). In this area it is flanked to the east by the Precambrian-cored Kvarkush Anticline, and to the west by the foreland basin (Fig. 1). Balanced cross-sections and the amount of shortening have not been determined for this part of the orogen, and farther discussion of it is beyond the scope of this paper. By far the best studied area of the foreland fold and thrust belt is in the Southern Urals (from c. 56~ to 51~ where its architecture has often been compared with that of other thrust belts from around the world, especially that of the Appalachians (e.g. Kamaletdinov 1974; Kruse & McNutt 1988; Rodgers 1990). However, recent structural mapping and seismic reflection data have shown the southern Uralides to be different (see below) (Brown et al. 1997b, 1998, 1999; Perez-Estaun et al. 1997; Giese et al. 1999; Alvarez-Marron 2000; Alvarez-Marron et al. 2002). Between c. 56~ and 53~ the Southern Urals foreland fold and thrust belt is a c. 150 km wide, west-vergent thrust wedge made up of Precambrian basement in the Bashkirian Anticline, the accretionary complex, Palaeozoic platform and foreland basin sediments (Figs 1 and 5). Palaeozoic shortening in this part of the thrust belt is c. 20 km or less (Fig. 5a and b; Brown et al. 1996, 1997b, 1999; Perez-Estaun et al. 1997; Giese et al. 1999). South of 53~ the foreland fold and
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Fig. 5. (a) Balanced and restored cross-section across the central Bashkirian Anticline (after Perez-Estaun et al. 1997). The calculated shortening is c. 20 km. The location is shown in Figure 1. (b) Balanced and restored cross-section across the southern Bashkirian Anticline (after Brown et al. 1997b). The calculated shortening is c. 17 kin. The location is shown in Figure lb. (c) Surface slope v. basal dip angle relationships for critical wedges. Calculation of the surface slope of section 3 using the equation a = arctan(tan f i / H ~ ) - fi (where a is the surface slope, fl is the basal slope and H is shortening) yields an c~ value of 1.1 ~ for a taper (r -----ct + t ) of 4.1 ~ This value of r requires only a small amount of material to have been eroded from the frontal part of the belt, and is in agreement with fission-track data (Seward et al. 1997; 2002). It also yields realistic values for average strain (AR = (tan r/tan/3)) of 1.3 : 1, and would place the section within the subcritical field. (See Brown et al. (1997a) for further explanation).
TECTONICS IN SOUTHERN AND MIDDLE URALS
thrust belt is dominated by the Southern Urals accretionary complex (Brown et al. 1998, 2004; Alvarez-Marron et al. 2000; Fig. lb). Cross-sections by Brown et al. (2004) indicate a very different structural style from that of the thrust belt to the north, in the Bashkirian Anticline, although the amount of shortening in this part of the thrust belt also appears to be small. The Uralides foreland fold and thrust belt exhibits a number of features that differentiate it from other Palaeozoic thrust belts. For example, the amount of shortening is very small, with vertical displacement along faults nearly equal to horizontal displacement (Brown et al. 1997b; Perez-Estaun et al. 1997). Mechanically, the thrust belt may never have reached a critical taper, and developed as a subcritical wedge (Fig. 5c; Brown et al. 1997a). The amount of basement involvement is high, and in many cases thrusting appears to have been localized by reactivation of two sets of pre-existing structures in the basement (Brown et al. 1997b, 1999; Perez-Estaun et al. 1997; Giese et al. 1999). Reactivation of structures parallel to the developing Uralide structural grain resulted in the incorporation of crystalline thrust sheets into the thrust belt at an early stage in its development, whereas those at a high angle to the Uralide structural grain influenced the location and development of lateral structures that can explain along-strike structural changes (Fig. 6; Perez-Estaun et al. 1997; Brown et al. 1999). The small amount of shortening, together with the localization of thrusts along pre-existing structures, suggests that the basal detachment may also be controlled by a Precambrian feature within the basement, or is absent completely.
413
Late o r o g e n i c strike-slip f a u l t i n g a n d granitoid e m p l a c e m e n t (Late C a r b o n i f e r o u s to E a r l y Triassic)
The internal part of the Uralides is made up of a late orogenic strike-slip fault system (e.g. Echtler et al. 1997; Friberg et al. 2002; Hetzel & Glodny 2002) that extends for more than 700 km along the Uralides before it disappears beneath Mesozoic cover in the south and north (Fig. 7a). Throughout much of the Middle and Southern Urals this strike-slip fault system corresponds to the East Uralian zone, although the currently defined Main Uralian fault in the Middle Urals appears to be its western limit there (Ayarza et al. 2000a; Brown et al. 2002). Dating on one segment of the strike-slip fault system indicates a Late Permian to Early Triassic age (247-240 Ma) for the development of fault-related mylonites (Hetzel & Glodny 2002), and latest Carboniferous (305-291 Ma) ages for associated metamorphic rocks (Echtler et al. 1997; Eide et al. 1997). The late orogenic strike-slip fault system was extensively intruded by latest Carboniferous to Permian granitoids, first in the southern part (292-280 Ma) and then in the northern part (270-250 Ma; Fig. 7; Bea et al. 1997, 2002, 2006; Montero et al. 2000). In general, the granitoids were emplaced at a high level in the crust, at c. 12-15 km depth (Fershtater et al. 1997). These granitoids have a high SiO2 content, and are mildly peraluminous, with elevated Rb, Cs, Ba, Th and U contents (Fig. 7b), but with an unusually primitive Sr and Nd isotopic composition (Fig. 7c) (Bea et al. 1997 2002; Fershtater et al. 1997;
Fig. 6. (a) Geological map of the northern Bashkirian Anticline (location is shown in Fig. lb). (b) Simplified, balanced and restored cross-sections across the northwestern part of the Bashkirian Anticline (locations are shown in (a)) (after Brown et al. 1999). Comparing the map and the cross-sections, it should be noted how the Yurmatu anticline changes abruptly along strike into the Inzer syncline. Such a change is strongly indicative of a lateral structure (the Inzer lateral ramp in (a)). Also, the Karatau fault is an excellent example of a lateral structure across which displacement is transferred toward the foreland. (c) Schematic block diagram showing the relationships between hanging-wall structures and basement topography (after Brown et al. 1999). It should be noted that, because of problems in the projection, the Karatau fault and the Inzer lateral ramp have not been drawn in their true orientation relative to the transport direction; in reality they are somewhat oblique to the orientation shown.
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Fig. 7. (a) Schematic map of the Southern and Middle Urals outlining the late orogenic strike-slip fault system and the location of continental-type granitoids. (b) Continental-crust normalized trace element and REE plots of Permian collision-related granitoids. [3, gabbros; O, diorites; ~ , granodiorites; open crosses, granites. For Dzhabyk, additionally, crossed squares and open circles represent the Mochagi and Rodnichki quartz monzonites, respectively (from Bea et al. 2002). The average Permian granite of the Uralides has a trace element composition characteristic of continental granites, in which some trace element anomalies characteristic of arc magmas, although attenuated, are still recognizable. The only materials able to produce partial melts with this conjunction of mantle-like isotope and crust-like chemical composition are subduction-related rocks with a short crustal residence time of a few tens of million years. (c) eNd(t) V. esr(t) of Uralide continental granitoids (from Bea et al. 2002). The isotopic signature of the Permian continental-type granitoids is very primitive, with 87Sr/86Sr(t) and 143Nd/la4Nd(t) values that match the subduction granites. This feature excludes continental materials older than Silurian as a possible protolith.
Montero et al. 2000; Gerdes et al. 2002, pp. 3-19). Bea et aL (2002) interpreted this to have resulted from recycling of the older continental arc material that was deeply buried after the collision; they also interpreted Permian crustal melting to be the result of a combination of radiogenic heating of an overthickened sialic crust, from local underplating by mafic magmas, and from local accumulation of heat and fluids related to the oblique, crustal-scale strike-slip shear zones that finally assembled the Uralides. The existence of a late orogenic strike-slip fault system along the entire interior of the Uralides suggests that widespread mass transfer took place along the axis of the orogen during the late stages of its tectonic evolution. Estimates of displacement along some strands of this fault system range from a few tens of kilometres to more than 100 km (Ayarza et al. 2000a; Hetzel & Glodny 2002). The presence of high-grade metamorphic rocks near the surface at the time of granitoid generation suggests extensive exhumation of material from the lower crust and its emplacement into the upper crust. The widespread melting of deep crustal material and its subsequent emplacement in the upper crust is also indicative of mass transfer. Both these processes are also suggestive of heat transfer from the lower crust, as hot material in the form of granulites and melt ascends and is emplaced in the colder upper crust. The evolution of the melt emplacement from south to north is also suggestive either of heat transfer along strike in the
orogen, or differential heating from south to north. Much work is needed on the structure, kinematics, granitoids, and geochronology of this important strike-slip fault system before it is possible to fully understand its relevance to orogen-parallel mass and heat transfer during the late stages of the Uralian orogeny.
F i n a l c r u s t a l s t r u c t u r e ( L a t e T r i a s s i c to R e c e n t )
The ESRU (Juhlin et al. 1998; Fig. 8a), URSEIS (Berzin et al. 1996; Fig. 8b), and reprocessed Russian seismic reflection or refraction surveys provide significant new data for interpreting the crustal structure of the Uralides (Steer et al. 1995, 1998; Carbonell et al. 1996, 1998, 2000; Echtler et al. 1996; Knapp et al. 1996; Friberg et al. 2000a, 2002; Brown et al. 2002). In the Southern (URSEIS) and Middle (ESRU) Urals the East European Craton crust thickens eastward from c. 40 km to c. 48 km, and is imaged by subhorizontal to east-dipping reflectivity that can be related to its Palaeozoic and older evolution (Fig. 8). The suture zone between Baltica and the accreted terranes, the Main Uralian fault, is poorly imaged in the URSEIS section, but in the ESRU section it is imaged as a zone of east-dipping reflectivity that extends from the surface into the middle crust; it marks an abrupt change to weakly subhorizontal reflectivity in the Tagil
TECTONICS IN SOUTHERN AND MIDDLE URALS
415
Fig. 8. (a) Interpreted line drawings of the coherency filtered, depth-migrated ESRU data (after Brown et al. 2002). (See Fig. lb for location). The main suture zones that bind the tectonic units together have been interpreted to end at the Moho. Their exact location at depth cannot be unambiguously interpreted, and has therefore been shown as a zone in which they may possibly occur. The location of the UWARS wide-angle Moho is from Juhlin et al. (1998). (b) Interpreted line drawings of the coherency filtered, depth-migrated URSEIS vibroseis data (after Tryggvason et al. 2001). (See Fig. lb for location). The main suture zones that bind the tectonic units together have been interpreted to end at the Moho. Their exact location at depth cannot be unambiguously interpreted, and has therefore been shown as a zone in which they may possibly occur. The location of the URSEIS explosion-source reflection Moho (Steer et al. 1998) and the refraction Moho (Carbonell et al. 1998) are shown along with the Moho imaged in this dataset.
arc (e.g. Ayarza et al. 2000a). East of the Main Uralian fault, the Magnitogorsk (Southern Urals) and the Tagil (Middle Urals) volcanic arcs display moderate to weak upper crustal reflectivity, and diffuse middle to lower crustal reflectivity. The Moho beneath both arc complexes is poorly imaged in the reflection data, but based on refraction data is interpreted to be at 5 0 - 5 5 km depth (Fig. 8; Thouvenot e t al. 1995; Juhlin e t al. 1996; Carbonell e t al. 1998). East of the arc complexes, the wide zone of anastomosing strike-slip faulting and granitoids of the East Uralian zone is imaged in the seismic sections as clouds of diffuse reflectivity interspersed with, or cut by sharp, predominantly west-dipping reflections. In the Southern and Middle
Urals, west-dipping reflectivity of the Trans-Uralian zone extends from the middle crust into the lower crust, where it appears to merge with the Moho (EchOer e t al. 1996; Knapp e t al. 1996; Steer e t al. 1998; Friberg et al. 2000a, 2002; Brown et al. 2002). The URSEIS experiment imaged a number of sub-Moho reflections (Knapp e t al. 1996; Steer et al. 1998) that may represent deformation scars related to the Uralian orogeny. The overall seismic reflection pattern of the Uralide crust as imaged by the URSEIS and E S R U data is bivergent, perhaps representing the original collision-related crustal architecture (Fig. 9). With the exception of possible minor extensional features in the eastern part of the E S R U section (Friberg e t al. 2002), there
Fig. 9. Generalized crustal-scale structural cross-section of the Southern Urals along the URSEIS profile. The location of the Moho is from Carbonell et al. (1998) and Steer et al. (1998).
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Fig. 10. Schematicplate model for the Southem Urals, outlining the key tectonic processes over time that went into the building of the Uralides.
is little evidence in the seismic reflection fabric for large-scale extensional collapse of the Uralides. The URSEIS and ESRU sections indicate that crustal thickness and Moho topography change somewhat between the Southern and Middle Urals, although the crustal root can be seen to extend along the western volcanic axis of the orogen. Recently, Diaconescu & Knapp (2002) argued that the formation of eclogite in the root zone may have led to an isostatically balanced system that ultimately preserved the Uralide structure. However, petrophysical modelling of the Uralide crust along the URSEIS transect indicates that the root zone is made up of mafic garnet granulite and not eclogite (Scarrow et al. 2002; Brown et al. 2003), so perhaps other as yet unidentified processes have been active. The Uralide orogen records a long and complex subduction-accretion history (e.g. arc-continent collision along the margin of Baltica, Andean-type subduction beneath Kazakhstan) prior to the final collision that gave it its final bivergent architecture. The complex late orogenic history, which involved extensive wrench faulting accompanied by widespread melt generation and granitoid emplacement in the interior of the orogen (see above), probably significantly overprinted and/or reworked much of the subduction- and accretionrelated tectonic fabric, giving this zone its varied and complex reflection seismic character. For example, orogen-parallel mass transport of material, as outlined above, may account for the subhorizontal reflectivity in the lower crust imaged in the ESRU seismic reflection profile (Koyi et al. 1999).
Conclusions The Uralide orogen of Russia was one of the main orogens built during the Palaeozoic assembly of Pangaea. Unlike the Variscide-Appalachian orogenic system, which was largely rifted apart by the opening of the Atlantic Ocean or extensively overprinted by post-orogenic processes, the Uralides have been preserved intact, providing an opportunity to study the tectonic processes that went into forming this Palaeozoic orogen. Clearly, subduction and accretion processes dominated during the Mid-Devonian to Early Carboniferous, as intra-oceanic island arcs collided with Baltica. During the same time period, we interpret that Andean-type continental arc(s) were forming on the margin of Kazaldastan (Fig. 10). The Southern Urals is of particular importance in the subduction and accretion history of the Uralides because it contains one of the best preserved examples of an arc-continent collision in any Palaeozoic orogen. The state of preservation and the level of exposure allow this arc-continent collision to be compared in detail with those that are currently active around the world, providing unprecedented insight into Palaeozoic tectonic processes. With the closure of the Uralian ocean, deformation began in the western Uralides foreland fold and thrust belt and, concomitantly, deposition of the foreland basin began (Fig. 10). The Uralides foreland fold and thrust belt is distinct from most other thrust belts, in particular in the amount of shortening, the amount of basement
TECTONICS IN SOUTHERN AND MIDDLE URALS
involvement, and the along-strike structural changes. W h y the shortening is so small is not clear. Perhaps the far-field stress induced by a highly oblique c o n t i n e n t - c o n t i n e n t collision was too small to imbricate the dense crust of the island arc systems that formed the margin at that time, and merely resulted in the reactivation of earlier structures in the basement. Whatever the reason, the small amount of shortening allows the relationship between the pre-existing basement structures and changes in structural style to be correlated. At the same time as the foreland fold and thrust belt was forming, the interior part of the orogen underwent extensive strike-slip faulting, metamorphism, melt generation and emplacement, and exhumation (Fig. 10). Finally, the bivergent crustal structure of the Southern and Middle Urals reflects the crustal stacking that occurred on both sides of the orogen during the subduction and accretion stage and during the c o n t i n e n t - c o n t i n e n t collision stage.
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TECTONICS IN SOUTHERN AND MIDDLE URALS
STEER, D. N., KNAPP, J. H., BROWN,L. D., RYBALKA,A. V. & SOKOLOV, V. B. 1995. Crustal structure of the Middle Urals based on reprocessing of Russian seismic reflection data. Geophysical Journal International, 123, 673-682. STEER, D. N., KNAPP, J. H., BROWN,L. D., ECHTLER, H. P., BROWN,D. L. & BERZIN, R., 1998. Deep structure of the continental lithosphere in an unextended orogen: an explosive-source seismic reflection profile in the Urals (Urals Seismic Experiment and Integrated Studies (URSEIS 1995)). Tectonics, 17, 143-157. THOUVENOT, F., KASHUBIN, S. N., POUPINET, G., MAKOVSKIY, V. V., KASHUBINA, T. V., MATTE, Ph. & JENATTON, L. 1995. The root of the Urals: evidence from wide-angle reflection seismics. Tectonophysics, 250, 1-13. TRYGGVASON, A., BROWN, D. & PEREZ-ESTAUN, A. 2001. Crustal architecture of the southern Uralides from true amplitude processing of the URSEIS vibroseis profile. Tectonics, 20, 1040-1052. VAN dEN BEUI1000 km. 'Hard' coupling between Laurentia and Baltica is implied, in contrast to the 'soft' docking of Eastern Avalonia to the Baltica. The former is also indicated by inversional tectonic features recognized in the lowermost Devonian succession of the Pechora Basin. This study was performed within the framework of the European Science Foundation EUROPROBE programme. It is a part of the project K-104 financed by the LithuanianScience Foundationand AB Geonafta (S.S.).
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SVENNINGSEN, 0. M. 1994. The Baltica-Iapetus passive margin dyke complex in the Sarektjakka Nappe, northern Swedish Caledonides. Geological Journal, 29, 323-354. SVENNINGSEN, O. M. 1995. Extensional deformation along the Late Precambrian-Cambrian Baltoscandian passive margin: the Sarektjakka Nappe, Swedish Caledonides. Geologische Rundschau, 84, 649-664. VECOLI, M. & SAMUELSSON,J. 2001. Quantitative evaluation of microplankton palaeobiogeography in the Ordovician-Early Silurian of the northern Trans European Suture Zone: implications for the timing of the Avalonia-Baltica collision. Review of Palaeobotany and Palynology, 115, 43-68. VEJBAEK, 0. V., STOUGE, S. • POULSEN, K. D. 1994. Palaeozoic Tectonic and Sedimentary Evolution and Hydrocarbon Prospectivity in the Bornholm Area. Danmarks Geologiske UndersCgelse. Serie A, 34, 1-23. VOLOZH, Yu. A., ANTIPOV, M. P., BRUNET, M. F., GARAGASH, I. A., LOBKOVSKII, L. I. & CADET, J. P. 2003. Pre-Mesozoic geodynamics of the Precaspian basin (Kazakhstan). Sedimentary Geology, 156, 35 -58. VISHN1AKOV, I. B., SINICHKINA, A. M. & KCHNYKIN, V. I. 1997. Volhyn-Podolian margin of the East European Platform. In: SINICHKINA, A. M. (ed.) Geology and Oil Prospects of the Western Part of the East European Platform. Belaruskaya Navuka, Minsk, 565-614 [in Russian]. VISHN1AKOV, I. B., VOSANTSHUK, S. S. & GLUSHKO, V. V. 1982. Palaeotectonics and oil prospects of Lviv depression (UkSSR). In: Perspectives of Exploration in Hydrocarbon Prospective Areas of Ukraine Lviv, 13-20 [in Russian]. WILLNER, A. P, ERMOLAEVA,T., STROINK, L., E T AL. 2002. Palaeozoic amalgamation of Central Europe: new results from recent geological and geophysical investigations. Tectonophysics, 360, 5-21. WINCHESTER, J. A. & THE PACE TMR NETWORK TEAM 2002. Palaeozoic amalgamation of Central Europe: new results from recent geological and geophysical investigations. Tectonophysics, 360, 5-21. WINCHESTER,J. A., PHARAOH,T. C., VERNIERS,J., IOANE,D. ~; SEGHEDI A. 2006. Palaeozoic accretion Winchester, of Gondwana-derived terranes to the East European Craton: recognition of detached terrane fragments dispersed after collision with promontories. In: GEE, D. G. & STEPHENSON, R. A. (eds) European Lthosphere Dynamics. Geological Society, London, Memoirs, 32, 323-332. ZELAZNIEWICZ, A., SEGHEDI, A., JACHOWICZ, M., BOBINSKI, W., BULA, Z. & CWOJDZINSKI, S. 2001. U - P b SHRIMP data confirm the presence of a Vendian foreland flysch basin to the East European Craton. ESF Europrobe Meeting Abstracts. Middle East Technical University, Ankara, 98-100. ZIEGLER, P. A. 1994. Geodynamic processes governing development of rift basins. In: ROURE, F., ELLOUZ, N., ET AL. (eds) Geodynamic Evolution of Sedimentary Basins. Moscow, 19- 67. ZHEMCHUGOVA,V. A., MELMKOV, S. V. & DANILOV, V. N. 2001. Lower Paleozoic of Pechora oil and gas bearing basin (structure, origin conditions, oil and gas potential). Academy of Earth Sciences Press, Moscow [in Russian]. ZINOVENKO, G. V. 1986. Baltic-Dnestr Pericratonic Basins. Nauka i tekhnika, Minsk.
Late Palaeozoic intra- and pericratonic basins on the East European Craton and its margins R. A. S T E P H E N S O N 1, T. Y E G O R O V A 2, M.-F. B R U N E T 3, S. S T O V B A 4, M. W I L S O N 5, V. S T A R O S T E N K O 2, A. S A I N T O T 1'6 & N. K U S Z N I R 7
1Netherlands Centre for Integrated Solid Earth Sciences, Faculty of Life and Earth Sciences, Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands (e-mail: randell, stephenson @f alw. vu. nl ) 2Institute of Geophysics, National Academy of Sciences of Ukraine, Kyiv, Ukraine 3Laboratoire de Tectonique, Universiti Pierre et Marie Curie, 4 place Jussieu, 75252 Paris cedex 05, France 4Naukanaftogaz, Naftogaz of Ukraine, Uritckogo 45, 03035 Kyiv, Ukraine 5Institute of Geophysics and Tectonics, School of Earth and Environment, Leeds University, Leeds LS2 9JT, UK 6present address: Geological Survey of Norway (NGU), Leiv Eirikssons vei 39, N-7491 Trondheim, Norway 7Department of Earth and Ocean Sciences, University of Liverpool, Liverpool L69 3GP, UK
Abstract: The (Mid-) Late Devonian to Early Carboniferous was a time of widespread rifting on the East European Craton (EEC) and its margins. The most prominent basin among these and, accordingly, the best documented is the Dniepr-Donets Basin (DDB) in Ukraine and southern Russia. The DDB is associated with voluminous rift-related magmatism and broad basement uplift. Two other large, extensional, basin systems developed along the margins of the EEC at the same time: the East Barents Basin (EEB) and its onshore prolongation the Timan-Pechora Basin (TPB), and the Peri-Caspian Basin (PCB). Rifting, associated magmatism, and possible domal basement uplift are also reported elsewhere within the EEC, suggesting a common, 'active', rifting process, involving a cluster of thermal instabilities (or generalized thermal instability) at the base of the lithosphere beneath widely separated parts of the EEC by Mid-Late Devonian times. The DDB is an intracratonic rift basin, cutting across the Archaean-Palaeoproterozoic structural grain of its basement and, as such, differs from the EBB-TPB and PCB, which are pericratonic rift basins developed on reworked and juvenile crystalline basement accreted to the EEC during the Neoproterozoic. The DDB opened into a deep basin, possibly having oceanic lithospheric affinity, to the SE, in the area where it adjoins the southern PCB, suggesting the possibility that rifting led to (limited?) continental break-up in this area at this time. Post-rift compressional tectonic reactivations and basin inversion in the DDB, leading to the formation of its prominent Donbas Foldbelt segment, are related to Tethyan events (Cimmerian and Alpine orogenies) occurring on the nearby southern margin of the EEC. Post-rift compressional inversions in the PCB and TPB, which lie closer to the Urals margin of the EEC, are related to Uralian tectonics.
The Late Palaeozoic, in particular the Late Devonian, was an important time for extensional basin development on the East European Craton (EEC) and along its margins (see Fig. 1). The most prominent basin among these, and the one that received the most attention by E U R O P R O B E (Stephenson 1996), is the D n i e p r - D o n e t s Basin (DDB). This is an intracratonic rift basin with well-defined syn- and post-rift sedimentary successions within the Archaean-Palaeoproterozoic Sarmatian segment of the EEC. Two other large, extensional, basin systems developed along the margins of the EEC during the Late Palaeozoic: the East Barents Basin (EBB), mainly below sea level at present, with its onshore prolongation the T i m a n - P e c h o r a Basin (TPB), and the Peri-Caspian Basin (PCB), also in part below present-day sea level (northern Caspian Sea). Late Devonian intracratonic tiffing, associated magmatism, and possible domal basement uplift are also reported elsewhere within the EEC, on the Kola Peninsula (Kontozero Graben) and in the Vyatka Rift (Fig. 1). Whereas extensional tectonics and basin formation characterized the whole of the EEC during much of the Late Palaeozoic, the margins of the European continent had been or were shortly to be strongly affected by orogenesis during this time (Fig. 1). The main aim of this paper is to make a critical reassessment of what is actually known about Late Palaeozoic basin development on the EEC and to judge this in terms of lithospheric processes that may or may not be linked to plate boundary (convergence a n d / o r divergence) tectonic events taking place at about the same time. The more complete knowledge of the DDB is used as a point of departure in discussing these issues as they pertain to the other Late Palaeozoic basins of the EEC.
Overview of major Late Palaeozoic rift basins: architecture, magmatism and crustal structure
Dniepr-Donets Basin (DDB) The D n i e p r - D o n e t s Basin (DDB) is located in the southeastern part of the EEC along a N W - S E - t r e n d i n g axis between the present-day Ukrainian Shield and Voronezh Massif (Figs 1 and 2). It is part of the same rift basin system as the shallower Pripyat Trough to the NW (mainly in Belarus) and the inverted Donbas Basin (Donbas Foldbelt (DF), straddling the U k r a i n e Russia border; Fig. 3) and its prolongation to the SE, the presentday Karpinsky Swell (Fig. 2). The sedimentary succession of the DDB can be readily subdivided into pre-, syn-, and post-rift series, corresponding to pre-late Frasnian ( D 2 _ 3 ) , late F r a s n i a n Famennian (D3), and post-Devonian units, respectively (Fig. 4). The sedimentary thickness increases southeastwards to more than 20 km in the DF. Rifting may have begun slightly earlier in the SE, propagating northwestwards (see Stephenson et al. 2001; McCann et al. 2003). The post-rift succession is well developed, displaying evidence of multiple extensional reactivations as well as compressional tectonic events. A lack of stratigraphy of suitable age in the DF (as a result of subsequent uplift, deformation, and erosion affecting it) and, therefore, an absence of diagnostic structural relationships led to uncertainty regarding the timing and nature of post-rift tectonic events controlling its development (see Fig. 3). However, by comparing the exposed and drilled geology of the DF with seismic images from the adjacent Donets segment of the DDB (e.g. Fig. 4d), Stovba & Stephenson (1999) demonstrated that the main pre-inversion events affecting the DF
From: GEE, D. G. & STZPI-IENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 463-479. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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Fig. 1. Late Devonian palaeogeography of the EEC and surrounding areas (modified from Ziegler 1988). Pink, 'shield', areas; brown, inactive orogenic belts; grey, active orogenic belts; light blue, sedimentary platforms; green, ocean basins. Rift basins (dashed lines, labels italicized): BrS, Barents Shelf; DDB, Dniepr-Donets (Rift) Basin; KG. Kontozero Graben (Kola Peninsula); PCB, Peri-Caspian Basin; PKR, Pechora-Kolva Rift; PT, Pripyat Trough; VR, Vyatka Rift. Other abbreviations: EEP, East European Platform; UkS, Ukrainian Shield; VM, Voronezh Massif. Also shown, with boxes labelled accordingly, are the approximate locations of the maps shown in Figures 2 (DDB), 6 (PCB), and 8 (EBB-TPB). do not significantly differ from those of the DDB. This has been confirmed by subsequent structural studies (Saintot et al. 2003a,b) and by DOBREflection deep seismic profiling (Storba et al. 2005). The oldest sediments in the DDB are of Eifelian to MidFrasnian age, the so-called 'undersalt', pre-rift sediments. These were deposited in platformal terrestrial and shallow marine
environments and comprise sandstones, siltstones, clays and carbonates. These pre-rift Devonian sediments correlate with equivalent Devonian sequences of the East European Platform (Eisenverg 1988). They are characterized by homogeneous lithofacies, have an average thickness of 300-400 m, and include a series of stratigraphic gaps, the most significant being between the Eifelian and Givetian and between the Givetian and Frasnian. Thickness variations of this pre-rift succession are independent of the modern basement relief, although it is observed only locally on the rift shoulders. They were probably deposited over a much wider area, but were eroded during synrift uplift of the rift flanks. Similarly, they are locally absent atop intrabasinal structural highs developed during rifting. The marine Mid-Devonian sediments are not recorded in the southern DF area, where Eifelian to early Frasnian sediments are continental clastic deposits, transported northwards, with a few lacustrine carbonate intercalations, mainly deposited in a fluvial or delta-plain setting (McCann et al. 2003). Basal conglomerates rest unconformably on weathered Precambrian basement and are reportedly associated with fissural basaltic extrusive rocks already in the Eifelian (McCann et al. 2003). There is no evidence for the presence of a coaxial, but narrower, pre-Devonian, perhaps Riphean-aged graben underlying the DDB, as reported in much of the older literature (e.g. Chekunov et al. 1992). This was based on deep seismic sounding (DSS) velocity models, but is not observed on seismic reflection profiles recorded up to 12 s two-way travel time (TWT) (Stovba et al. 1996). No strata older than Mid-Devonian have been encountered in any of the numerous boreholes that penetrate basement beneath the Palaeozoic sediments of the DDB (Chirvinskaya & Sollogub 1980; Eisenverg 1988). Rather, the Devonian-Carboniferous succession revealed by the reflection data is much thicker than inferred from the earlier velocity models and occupies those parts of these models thought previously to represent Riphean strata (see Stovba et al. 1996). Therefore, tectonic models suggesting that a precursor Riphean rift basin was reactivated during the Devonian are no longer viable. Although modified by post-rift tectonic and especially salt movements (Stovba & Stephenson 2003), the basic architecture
Fig. 2. Tectonic map of the southern EEC, showing the extent of the Late Devonian Pripyat-DDB-DF rift basin. The dashed-line box indicates the location of the map shown in Figure 3. The locations of cross-sections shown in Figure 4 are also shown (red lines).
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Fig. 3. Cenozoic subcrop map of the Donbas Foldbelt area, showing the location of the DOBREflection profile (continuous line), shown in Figures 4e and 5. PreC, Precambrian crystalline basement; other stratigraphic labels are as in Figure 4.
of the DDB, seen in Figure 4, was developed during its Late Devonian rifting stage. High and laterally variable synrift subsidence rates, accompanied by the development of grabens and half-grabens, resulted in a wide range of local depositional environments and considerable palaeogeographic heterogeneity both in space and time, intense volcanism, and multidirectional tectonic movements (Stovba et al. 1996). Synrift deposits in the DDB reach a maximum thickness of about 4 km (e.g. Ulmishek et al. 1994; Stovba et al. 1996). They overlie the pre-rift sequence discordantly and are absent on some interior fault blocks, as a result of reduced sedimentation and subsequent erosion. Much of the lower part of the synrift sequence consists of Frasnian salt, called the 'lower salt', that alternates with clastic deposits and carbonates rocks in a complex laterally variable pattern. The depositional thickness of the Frasnian series is at least 1000 m and reaches a maximum (up to 2 km) in the axial zone of the southeastern part of the DDB. The upper part of the synrift series consists of a thinner Famennian 'upper salt' that thickens in the northwestern part of the DDB. In the southern DF, McCann et al. (2003) described the formation of half-grabens along major normal faults, filled with fluviatile continental clastic deposits and some lacustrine limestones. Short-lived but very frequent subaerial fissural extrusions (making up about two-thirds of the sequence) are always preceded by clastic input showing relief formation and erosion. Synrift volcanic and intrusive rocks, consisting of a variety of alkali basalts and their differentiates and associated pyroclastic deposits, occur in two main series of late Frasnian and late Famennian age in the DDB, attaining thicknesses of more than 2000 m (e.g. Wilson & Lyashkevich 1996). Additionally, there was widespread intrusion of tholeiitic basalt dykes, sills and stocks, which cross-cut formations ranging from Frasnian to late Famennian in age. The synrift phase s e n s u stricto terminated by the end of the Devonian and, in general, the Carboniferous and younger post-rift sedimentary fill of the DDB has the configuration of a broad syncline centred on the rift axis, overlapping the rift shoulders, and increasing in thickness towards the SE (Fig. 4). Seismic profiles published by Stovba et al. (1996) clearly demonstrate, however, that the DDB was affected during its Permo-Carboniferous
evolution by a series of post-rift extensional reactivations, generally synchronous with salt movements (Stovba & Stephenson 2003), but tectonic in origin; these occur at the end of the early Vis6an, during the mid-Serpukhovian, and during latest Carboniferous-earliest Early Permian times. Evidence of these events is visible in the regional cross-sections (Fig. 4); they have been comprehensively documented by Stovba & Stephenson (1999). The intensity of each of the Permo-Carboniferous extensional events increases in the DDB southeastwards towards the DF, where the late early Vis6an rift reactivation is clearly in evidence in the field, with uplifted and clearly rotated blocks along active normal faults, and associated magmatic activity (McCann et al. 2003). Saintot et al. (2003a) inferred a N N E - S S W extension in the DF that clearly affected the Early Carboniferous and older succession and, also, a younger transtensional stress regime thought to correspond to the latest Carboniferous-earliest Early Permian event recognized in the DDB by Stovba et al. (1996). Additional evidence of Early Permian extensional deformation along the northern margin of the DF, documented widely but generally not in published literature, was presented and discussed by Stovba & Stephenson (1999). Elsewhere in the DF, sediments of Late Cretaceous age directly overlie block-faulted and rotated Devonian and Carboniferous strata. The lack of a Permian-Early Cretaceous sedimentary record prevents a definite interpretation of the age of these faults; however, it is likely that the faulting and block rotation seen along the southwestern margin of the DF are part of the widespread phase of Early Permian extension (transtension) seen throughout the DDB (Stovba & Stephenson 1999). The Carboniferous succession is represented by continental deposits in the northwestern part of the DDB (e.g. Ulmishek et al. 1994; Dvorjanin et al. 1996; Izart et al. 1996). Elsewhere in the DDB it is characterized by continuous rhythmic sedimentation and comprises mainly siliciclastic rocks (with some clastic-carbonate sequences) deposited in shallow marine and lagoonal environments. There is little variation in the position of the basin depocentre. Only in the axial part of the southeastern DDB, where the Lower Carboniferous sequence includes marine carbonates, did the depth of deposition exceed 200 m. Exposed and drilled Lower Carboniferous sediments in the DF are mainly
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Fig. 4. Structural cross-sections through the DDB (a-d) and DF ((e); from DOBREflection), based on depth-converted versions of interpreted regional seismic reflection profiles (from Stovba et al. 1996, 2005; Stovba & Stephenson 1999). (For locations, see Fig. 2.) Light blue, salt bodies, brown, Devonian sediments, shades of grey, Carboniferous Ar-PR1, Archaean-Palaeoproterozoic; D2_3, Middle Devonian-Upper Devonian; C, Carboniferous; C1, Lower Carboniferous, (t, Tournaisian, vl, lower Visean, v2, upper Visean, s, Serpukhovian); Ca, Middle Carboniferous (Ukrainian-Russian usage: b, Bashkirian; m, Moscovian); C3, Upper Carboniferous (Ukrainian-Russian usage: e.g. Kasimovian and Gzelian); P1, Lower Permian, (as, Asselian; s, Sakmarian); Mz, Mesozoic; T, Triassic; J, Jurassic; K, Cretaceous; K2, Upper Cretaceous; Kz, Cenozoic; Pg, Palaeogene.
LATE PALAEOZOICBASINS ON THE EEC marine limestones overlain by sandy-clay deposits interbedded with thin coal and limestone beds. The (uppermost Famennian?-) Lower Carboniferous limestones were probably deposited in a very quiet shallow-water inner platform with occasional terrigeneous input. The uppermost part of this succession where it is exposed in the DF shows subaerial karstification, suggesting its emergence at the time of late early Vis~an extensional reactivation. The overlying silica-rich unit shows numerous synsedimentary deformational features such as normal faults and slumps (McCann et al. 2003). Middle and Upper Carboniferous successions are exposed throughout most of the DF and consist mainly of arenaceous-argillaceous rocks interbedded with coal and limestone. With the exception of coal beds and sandy-clay continental intercalations, most were deposited in a shallow-marine environment. Carboniferous sediments in the DDB reach thicknesses of 11 km, with the maximum depth of their base at about 15 km (Stovba et al. 1996). The present-day total thickness of Carboniferous sediments in the DF area is about 20 km based on the DOBREflection profile acquired as part of the EUROPROBE programme (Figs 4e and 5). The lowermost Lower Permian sediments are represented by monotonous sand-shale series containing rare interbeds of limestones and coals that, similar to the Upper Carboniferous units, reflect coastal-continental facies. Asselian sediments consist of five to seven layers of rock salt, separated by clastic deposits and carbonates, and also include numerous beds of gypsum, anhydrite and dolomite. The thickness of the salt layers, and per cent volume, increases upward in the section (Eisenverg 1988). The Sakmarian part of the series consists of a single salt layer probably representing redeposited Devonian salt dissolved from diapirs piercing the depositional surface in the Early Permian (Stovba & Stephenson 2003). In the southern pre-shoulder zone of the DDB the Lower Permian sequence abruptly decreases in thickness and pinches out as a result of a decrease in depositional thickness as well as subsequent erosion. In contrast, its thickness decrease towards the northern shoulder of the basin is far more gradual. There are no sediments of Early Permian age preserved within the DF, although Upper Carboniferous and Lower Permian sediments are documented beneath the eastern extension of the northern margin of the DF. A general absence of Upper Carboniferous and Lower Permian sediments in the northwesternmost part of the DDB can be explained by a decrease in the rate of post-rift subsidence within a platform-wide regime of relative sea-level fall. Elsewhere within the DDB, the basin margins, particularly the southern one, were exposed during Early Permian times whereas the axial part of the basin continued to subside (see Fig. 4). Uplift of the southern margin of the DDB was very shortlived, lasting no more than 2 - 3 Ma between the late Asselian and early Sakmarian (Stovba et al. 1996). Extensive erosion occurred, with progressively older sediments subcropping beneath the erosion surface in the direction of the Ukrainian Shield; by implication, considerable erosion of the Ukrainian Shield may also have occurred. Locally more than 2 km of Upper and Middle Carboniferous sediments were eroded at this time and during an ensuing dormant phase, which lasted until the Triassic. The widespread regional Permian unconformity observed throughout the DDB is, therefore, interpreted to be the result of the Early Permian event followed by a relative sea-level lowstand during the later Permian. Sedimentation resumed in the DDB in the Triassic, a time of tectonic quiescence, rising sea levels, and the resumption or continuation of post-rift subsidence. Most of the Mesozoic succession, comprising both marine and continental sediments, occurs throughout the area, overlying the rift axis as well as its flanks. Exceptions are the Upper Triassic and Lower Jurassic units, which occur only in the southeastern part of the DDB, and the Upper Cretaceous marls and chalks, which were eroded from large parts of the southern flank. The Upper Cretaceous succession
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was, as a whole, characterized by Chirvinskaya & Sollogub (1980) as 'close to' platform type, although subsidence coincident with the Devonian rift axis exceeds that of the marginal zones (see Fig. 4). It is up to 2000 m thick in the central part of the DDB, with maximum thicknesses of the Triassic, Jurassic and Cretaceous units being 900, 700 and 1000 m, respectively. In the vicinity of the DF, no marginal facies or developments are observed near the erosional edges of the Mesozoic successions. It is, therefore, likely that the entire area of the DF underwent post-rift subsidence during the Mesozoic and that, depending on relative sea-level variations, Mesozoic successions were deposited within its confines, but were later eroded. There is little evidence of post-rift magmatic activity in the DDB; however, this is not the case for the DF, where igneous rocks of Early Carboniferous, Early Permian and Mesozoic ages have been reported. A summary of the available geochronological data has been given by Alexandre et al. (2004). A widespread angular unconformity in the DDB developed at the end of Cretaceous-beginning of Palaeogene (Kabyshev et al. 1998). The magnitude of inferred relative uplift increases towards the Ukrainian Shield and, as during the Early Permian, its maximum occurred in the area bordering the DF. In this area, Upper Cretaceous, Jurassic and Triassic sediments were eroded (see Fig. 4). In the axial part of the southeastern DDB, there are local folds, domes and salt diapirs defining linear trends, which correspond to the trends of the main folds of the DF as seen on the Cenozoic subcrop map (Fig. 3). Within the DF itself, structural relationships determining the age of formation of folds, thrust and reverse faults can be observed only near its margins, where Lower Permian, Mesozoic and Cenozoic sediments are preserved. Stovba & Stephenson (1999) reported that no single geological section could be found in the published literature showing tightly constrained, structurally defined pre-Triassic folding or reverse faulting in the DF. In contrast, where Cretaceous sediments are present, for example along the northern margin of the DF, reverse faults and/or folds younger than the Cretaceous sediments and exposed at the surface are relatively common. Reverse faulting of Late Cretaceous age is also evident on the southern margin of the DF (Stovba & Stephenson 1999). Saintot et al. (2003a,b) determined that the palaeostress field associated with compressional structures observed in the Cretaceous sediments on the margins of the DF is identical to that recorded by the outcropping Carboniferous sediments. Thus it can be concluded that the inversion of the DDB and formation of the DF occurred mainly in the Late Cretaceous (see Stovba & Stephenson 1999; Stephenson et al. 2001). The DOBREflection profile (Fig. 5) shows that the shortening of the DF occurred at the crustal scale as a 'mega-pop-up', which involved a major detachment fault through the entire crust and an associated back-thrust (Maystrenko et al. 2003; Stovba et al. 2005). The Cenozoic section of the DDB unconformably overlies Upper Cretaceous and older series and reaches a maximum thickness of 500 m in the NW DDB (Eisenverg 1988). The Palaeogene sequence consists mainly of sands, clays and marls, and the Neogene sequence mainly of sands with clayey interbeds. Deep seismic sounding (DSS; e.g. Chekunov et al. 1992; Ilchenko 1996) and more recent wide-angle reflection-refraction (WARR) seismic studies (DOBREfraction'99 Working Group 2003) show that the amount of crustal thinning beneath the DDB increases to the SE, concurrently with increasing sedimentary thickness (see Stephenson et al. 2001). The most recent profile is DOBRE (DOBREfraction'99 Working Group 2003), crossing the inverted DF segment of the DDB (Fig. 5). The sedimentary basin itself is well-defined, overlying a main crustal layer that thins significantly beneath the main sedimentary depocentre. In turn, a high-velocity lower crustal layer thickens significantly in the same part of the profile. The shape of the sedimentary basin is asymmetric, with the steepest crystalline basement surface on the southwestern margin of the basin, whereas the asymmetry
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R.A. STEPHENSON ET AL. of the high-velocity layer displays its steepest upper surface beneath the northeastern margin of the basin. t-c) O
Peri-Caspian Basin (PCB)
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The Peri-Caspian Basin (PCB; also sometimes referred to as the Precaspian, Pricaspian or North Caspian Basin) is situated on the southeastern margin of the East European (Russian) platform and extends into the northern part of the Caspian Sea. It runs 900 km east-west and 600kin north-south, bordered on the east by the Ural Mountains and to the SE and SW by crustal terranes that have an uncertain relationship with the EEC (see Saintot et al. 2006). The sedimentary succession of the PCB is about 20 km thick (Figs 6 and 7). There is a prominent Lower Permian salt layer some 4-4.5 km thick with its base at a depth of 7-9.5 km in the central part of the basin. Overlying sediments comprise the 'postsalt' layer, which is up to 7 km thick, and range in age from Late Permian to Quaternary. What lies below (called the 'sub-salt layer', which is up to about 9 - 1 0 km thick) is characterized primarily on the basis of seismic data (e.g. Volozh 1991). The conventional view holds that four seismo-geological successions can be recognized in the 'sub-salt' section: Riphean, Lower Palaeozoic, Devonian-Lower Carboniferous and Middle Carboniferous-Lower Permian, separated by hiatuses seen as erosional unconformities in marginal areas. Even though a huge volume of seismic (reflection and refraction) data exists, the age of the sub-salt sediments in the centre of the basin is controversial because seismic correlation from the margins to the deep basin is rather uncertain, given complications arising from seismic facies, steep slopes and interruption of key seismic markers. This contrasts with the view of many authors that the reference horizons can be traced (robustly) through the entire depression (e.g. Lobkovsky et al. 1996). Thus, the oldest sediments in the basin could be as old as Riphean or as young as Devonian (see Brunet et al. 1999; Volozh et al. 2003a). Riphean ages are based on seismic velocities and fabrics, and on strata thought to be Riphean (but undated) encountered in wells on the northern and northeastern margins of the basin (e.g. Soloviev et al. 1989). It is worth noting that the now rejected (based on modern regional seismic reflection profiling (Stovba et al. 1996)) postulate of a thick Riphean sequence deep in the DDB was based on very similar arguments. According to Zonenshain et al. (1990) and others, all these sediments are of Devonian age (as in the DDB) and, as such, they could overlie Devonian-aged oceanic crust. The regional interpretation shown in Figure 7 is based on the conventional interpretation that Devonian sediments are underlain by Neoproterozoic and Early Palaeozoic successions (e.g. Volozh et al. 2003a). What can be stated with certainty is that Vendian strata occur above the Precambrian basement on the margins of the PCB, on the Russian Platform to the north and west, and on the eastern and southeastern margins, for example, in the South Emba region (Fig. 6). A sedimentary hiatus occurred in the Early Palaeozoic; probably no Cambrian series exists within most of the basin; and Ordovician-Silurian strata are limited in extent, although up to 1000 m thick in pericratonic troughs such as the South Emba (e.g. Brunet et al. 1999). Thick, more terrigenous deposits with marginal carbonate reef complexes form the Upper DevonianCarboniferous succession. The basin was more or less filled (probably accompanied by a sea-level drop) by Early Permian times and became isolated from the open sea by structural highs developed especially on its south and southeastern margins, after which the thick salt layer was deposited. The initial thickness of this layer (prior to reconfiguration by diapirism) is estimated to be about 4.5 kin, deposited mainly during the Kungurian (Early Permian). Clastic rocks of this age are present in the eastern part of the basin, sourced from the eroding Urals Mountains and
LATE PALAEOZOIC BASINS ON THE EEC
469
Fig. 6. Main tectonic units of the PCB (modified from Volozh 1991; Volozh et al. 2003) showing surrounding marginal uplifts. The light green area in the central PCB indicates the extent of the high-velocity layer at the base of the crust (eclogites?). Also shown are the locations of the Aralsor and Khobda positive gravity anomalies ('AA' and 'KA', respectively), the location of the schematic cross-section shown in Figure 7 (with position C marked by an X for reference), and the 12 and 20 km depth contours to top basement.
showing basinward progradation. Post-depositional movement has resulted in the development of about 1800 salt structures in the PCB (Volozh et al. 2003b), of various types, some related to hydrocarbon production. Permian and younger sediments in the PCB were deposited in shallow-water or continental conditions but, because of the dominance of salt movement during this time in producing local, intradiapiric depocentres, little can be said about post-Permian tectonic controls on basin subsidence. The nature of unconformities in the PCB and the timing of fault activity are poorly described in the literature and subject to some inconsistency (e.g. Brunet et al. 1999). The Frasnian (beginning in
the Givetian?) has been reported by some workers as a time of active rifting in the PCB and the F a m e n n i a n - T o u r n a i s i a n as a time of relatively stable subsidence with the formation of a topographic depression not compensated by sediments (Nikishin et al. 1996; Volozh et al. 1999). However, it is extremely difficult to document this from the existing literature because stratigraphic boundaries in the PCB itself are poorly resolved and defined because different authors use different interpretations for the same seismic horizons. Soloviev et al. (1989) reported that the Riphean U z e n - S a k m a r a graben was reactivated in mid-Devonian (Eifelian) times and that tectonic movements occurred in the
Fig. 7. Simplified sketch (after Brunet et al. 1999) of a north-south basin-crustal cross-section of the PCB (from an unpublished interpretation and compilation of the seismic line Zhambay-Uralsk by Yu.A. Volozh, V.I. Kozlov & Yu.G. Yurov); location is shown in Figure 6 (with position C marked for reference). Seismic refraction velocities are indicated in the crust and the high-velocity layer. The presence of the Riphean and Lower Palaeozoic sedimentary layers in the deep basin is based on the interpretation of seismic velocities and is not confirmed by drilling.
470
R.A. STEPHENSONETAL.
Late Devonian (Frasnian), accompanied by rapid subsidence towards the PCB. The Frasnian lies unconformably on Middle Devonian deposits in some wells on the eastern margin of the PCB (Akhmetshina et al. 1993). Abrupt changes in DevonianCarboniferous sequence thickness, limited by faults, as well as facies heterogeneity, are observed on the Astrakhan Dome (Brodsky et al. 1994), which is a structural high in the southwestern part of the PCB adjacent to the Karpinsky Swell (Fig. 6). Active faulting from the Late Devonian to the Mid-Carboniferous has been reported by Kirukhin et al. (1983) and Brodsky et al. (1994) in this area. A number of authors (e.g. Lobkovsky et al. 1996; Brunet et al. 1999) have reported tectonic subsidence curves for the PCB, but these are not well-constrained given the degree of actual borehole penetration and uncertainties of stratigraphic identification from seismic data (see Brunet et al. 1999). Backstripping of wells on the margins of the basin does not illuminate the Late Palaeozoic, and backstripping of synthetic stratigraphic columns from the basin centre are dependent on the authors' choice regarding the interpretation of key seismic horizons. However, independent of these considerations, a robust tectonic event did occur in the evolution of the basin during the Devonian, with other events dependent upon age interpretations and assumptions regarding paleobathymetry. There exists very little solid reference to volcanogenic deposits in the PCB. Volozh (1991) reported late Riphean continental volcanoclastic deposits in the exterior zone of the northwestern region, Ordovician-Silurian marine to continental volcano-sedimentary deposits in the southeastern region (Primugodzhar, South Emba) and in the early Kungurian in the SE (Primugodzhar, Koltyk-Zamstan). Shein et al. (1989) showed on a north-south cross-section volcanogenic rocks of mid-Devonian and Carboniferous (Vis6an-Bashkirian) age in the centre of the basin and of Triassic age in the south. According to Kostyuchenko et al. (2004), the presence of pyroclastic rocks of variable composition indicates volcanic activity during the Early and Mid-Carboniferous (Kalashnikov 1974; Vishnevskaya & Sedaeva 2000). Kostyuchenko et al. (2004) also reported the interpretation of Brodsky et al. (2000), from seismic data (reflection and refraction) across the Astrakhan Dome (Fig. 6), of the presence of a large basic extrusive magmatic body (about 40 km wide) lying below a depth of about 12 km, down to about 24 kin. The age of sediments overlying this proposed volcanic unit, and therefore helping to date it, is subject to exactly the same uncertainty, discussed above, relating to the PCB in general; that is, it is dependent upon the acceptance or not of the unverifiable deep seismic stratigraphy. The inferred volcanic body is overlain by sediments older than the Devonian succession according to the conventional view (e.g. Volozh et al. 2003a) and, therefore, of Ordovician-Silurian age. The alternative view, that the whole of the sedimentary succession is Devonian (e.g. Zonenshain et al. 1990), would allow this inferred magmatic body to be of Devonian age. The crystalline basement beneath the central PCB is thin, being only some 10-12 km thick, characterized by the absence of an upper crustal velocity layer (velocities 1000 Siemens), which delineate more resistive crustal units (S < 100 Siemens), characterize the electrical structure of the Shield. The most dominant conductor coincides roughly with the border between the Archaean and the Proterozoic crust in Finland, and in Russia under the Palaeozoic sediments of the East European Platform. Another branch cuts the Palaeoproterozoic Svecofennides in Southern Finland and continues through the Bothnia Belt in Finland towards the Skellefte~ region in Sweden. Prominent crustal conductors exist also around the Lapland Granulite Belt (S several thousand Siemens), in the Pechenga Belt, and in the Imandra-Varzuga Belt (Figs 1 and 3). The Caledonides also contain conductive material, at least in J~imtland, Sweden (Gee 1972; Gharibi 2000). In the following, we summarize the main features of the long, elongated conductors, although each of them also has specific features that provide information on the structure, properties and evolution of its host region (Korja & Hjelt 1993, 1998; Korja e t al. 2002). (1) The conductance of conductors is high (several thousand Siemens), which can be explained only by electronic conducting mechanisms. Hence the conductors are likely to consist mainly of graphite- and sulphide-beating rocks.
545
(2) Graphite in conducting assemblages has an organic origin, suggesting that the conductors are composed mainly of sedimentary rocks. (3) The internal structure of the conductors is complex, as is evident from airborne electromagnetic data, containing extremely conducting graphite- and sulphide-beating metasedimentary layers, hosted by resistive rocks. (4) Dipping conductors often have an association with a band of seismic reflectors. (5) Most of the conductors are located in the upper and middle crust without penetration into the lower crust. However, there are some conductors that penetrate through the entire crust (e.g. Skellefte~ and Bothnian regions; Figs 1 and 3), suggesting a transportation of coffductive sedimentary material into the lower crust. (6) Virtually all conductors represent supracrustal Palaeoproterozoic and younger assemblages. It is noteworthy that the age of most of the conductors, if known, seems to concentrate around 2.1-1.9 Ga, with the notable exception of the conductors within the Caledonides, which mainly represent Cambrian metasediments. As a part of the BEAR work, old MT data from the SVEKA profile (Korja & Koivukoski 1994) have been reinverted together with new data from 112 MT soundings along a 750 km long GGT-SVEKA profile (Fig. 1). An improved conductivity model (Fig. 4) confirms that the upper crust is highly resistive and the lower crust is conductive (S ~ 2 0 0 - 5 0 0 Siemens) beneath the Palaeoproterozoic Central Finland Granitoid Complex (CFGC), whereas the entire crust is very resistive in the Archaean Karelian Domain to the east of the Kainuu Belt. Two dipping upper and mid-crustal conductors exist at both sides of the CFGC. One set of conductors is found beneath the Tampere, Pirkkala, Hame Belts and Satakunta Rapakivi area in southern Finland. The conductors are probably caused by two separate subduction and collision processes, as their dips are towards north and south. Minor, SE-dipping conductors beneath the Ladoga-Bothnian Bay Zone, and a major SW-dipping conductor beneath the Iisalmi Archaean unit and the Kainuu (Schist) Belt can be
Fig. 4. Smooth 2D-inversionmodel of the conductivityof the crust and uppermost mantle along the GGT-SVEKAprofile (lower panel) and integrated crustal conductance of the model from surface to 60 km depth (upper panel). The main geological units and their abbreviationsare given in Figure 1. KuB, Kuhmo Greenstone Belt; LBBZ, Ladoga-Bothnian Bay Zone. Moho boundary is from Korsman et al. (1999). Figure is modifiedfrom Lahti et al. (2002).
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S.-E. HJELT ET AL.
explained by the presence of Palaeoproterozoic sedimentary rocks beneath the resistive Archaean rocks of the Iisalmi complex (Lahti et al. 2002). The lower crust is electrically rather heterogeneous in the Fennoscandian Shield. Beneath the belts of upper and middle crustal conductors, it is difficult to obtain any information on the lower crust because of the attenuation of EM fields in highly conducting overlying bodies. Reliable information on lower crustal properties can therefore be obtained only within the resistive regions. The lower crust in the Archaean Belomorides and Karelides is highly resistive, having a conductance below 10 S in many parts (Figs 3 and 4). The northwestern part of the Karelides is over 10 times more conductive (Fig. 3) than the southeastern part; yet the Archaean lower crust is in general much less conductive than the Palaeoproterozoic Svecofennian crust (Fig. 4), where lower crustal conductance is well over 100 S. The lower crust in Central Sweden is, however, highly resistive, having conductivities similar to the presumably Archaean lower crust ( < 100 S). Finally, the new results from the BEAR data from site B42 (Fig. 5) show that the Archaean middle to lower crust may have conductances of a few hundreds of Siemens, similar to the Proterozoic lower crust. In summary, there are areas of both resistive and conductive Archaean lower crust and areas of conductive and resistive Proterozoic lower crust. Consequently, there is no obvious correlation between the age (nominal age determined according to the surface lithology) and conductance of lower crustal rocks. Therefore, it is difficult to explain the enhanced lower crustal conductivity by some 'universal' causes (e.g. precipitation of carbon from CO2-bearing fluids from the mantle, or trapping of water in the lower crust). Explanations are likely to be related to 'local' tectonics; that is, to the style of subduction and following collision processes, which determine how much and where conductive sedimentary material are transported.
U p p e r m a n t l e c o n d u c t i v i t y : the B E A R p r o j e c t
The BEAR subproject was one of the key experiments of EUROPROBE's SVEKALAPKO project. The BEAR project focused on studying the electrical properties of the upper mantle
Fig. 5. Examples of 1D models of resistivity v. depth for northern Felmoscandia from BEAR data (B42) and older experiments (KIR, Jones 1982, 1983); KAR Kaikkonen et al. 1983; Korja & Koivukoski 1994; Korja et al. 2002). Left panel has a logarithmic and fight panel a linear depth scale. For comparison, a 1D model of Central Europe from Olsen (1998) and dry olivine resistivities (= 1/ conductivities) under a relevant continental geotherm _ 100 ~ from Xu et al. (2000) are shown. Moho depths (Korsman et al. 1999) are for site B42. Figure is modified from Lahti et al. (2005).
beneath Fennoscandia and thereby aimed at gaining a deeper insight into the structure, evolution and contemporary dynamics of the continental lithosphere beneath cratons and possibly deeper below lithosphere, and finally at correlating and interpreting the results jointly with other geophysical and geological data available from the Fennoscandian Shield. The BEAR experiment itself consisted of ultradeep electromagnetic sounding, with use of a Shield-wide MT and magnetometer array of simultaneous long-period recordings (Fig. 2). Time variations of the Earth's electromagnetic field were measured for 45 days at 46 MT and 20 magnetometer sites, having an average separation distance of c. 150 km. The time series data from the array recordings were processed by three methods (Varentsov et al. 2003a), which resulted in a number of different EM transfer functions (magnetotelluric impedance, tipper, horizontal magnetic, horizontal spatial gradients) for a wide period range of 10-100 000 s and for a number of remote reference approaches (Varentsov et al. 2003a). However, as mentioned above, because of the source effect (proximity of the source region of the magnetotelluric fields), the useful period range is limited to c. 10 000 s for magnetotelluric impedance data. Consequently, the following results are obtained using BEAR data from 10 s to 10 000 s and pre-BEAR A M T - M T data from 1000 Hz to 1000 s. The analysis and modelling of the BEAR data have resulted so far in the following three conclusions that will be discussed separately below: (1) conducting material is required somewhere below 100 km almost everywhere beneath Fennoscandia; (2) 1D inversion results from site B42 suggests a depth of c. 170 km for the upper mantle conductor in northern part of the Shield; (3) magnetotelluric data exhibit strong anisotropic behaviour, in particular in the central part of the Fennoscandian Shield. The 3D modelling has shown (Engels et al. 2002; Varentsov et al. 2002) that an excess of roughly 5000 S of conducting material (e.g. 50 km of 10 1)in) is required somewhere below 100 km. The 3D model used included 3D crust and 1D upper mantle. For 1D upper mantle, two model variants were used. The first was the Fennoscandian reference model (Korja et al. 2002), which is, in general, compatible with the dry olivine conductivity model (Xu et al. 2000) and the model for Central Europe (Olsen 1998). The model had no conductive layer in the upper mantle, but a monotonous increase as a result of rise in temperature. Comparison of the model responses with observations showed that the observed phases were systematically higher than the modelled phases. The second model contained a conductive layer in the upper mantle having a conductance of 5000 S. This model removed systematic bias between observations and model responses, yet this model, with a single conductive layer at a depth of 1 5 0 - 2 0 0 k m , could not produce a satisfactory fit between observations and model responses (one-third of the sites had a good fit, in another third of the sites observed phases were still higher than the model phases, and in the remaining sites they were smaller), indicating that a simple 1D model with a single conducting layer is not valid for the Shield. Dimensionality analysis of the BEAR data (Lahti et al. 2005) shows that in the northeastern part of the array, electrical structure is nearly 1D, whereas in other parts the data has rather strong 2D or 3D character. Therefore site B42 (Salla) was selected (Fig. 5) as an example, where simple 1D inversion will provide reliable information on upper mantle conductivity. For static shift correction, a commonly used method is to adjust the level of apparent resistivity curves at their long-period branch using some global reference curve obtained from magnetic data, which is not affected by static shift. An obvious reference curve would be that of Olsen (1998) for Central Europe, estimated from European geomagnetic observatory data. It turned out that phases from site B42 and Olsen's response do not coincide, which is a condition for the correction of static shift, at a common long-period interval
ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD around 10 000 s. This indicates that the upper mantle structure at depths corresponding to 10 000 s is different in these two areas and longer periods would be needed. Longer period data, however, are difficult to obtain in the BEAR experiment, as a result of the source effect, as discussed above. Alternative approaches for the correction of static shift include the correction at short period intervals using magnetic data, e.g. from time domain EM methods, or the correction by spatial averaging. The latter was used for site B42 and the static shift was corrected by averaging six apparent resistivity curves from nearby sites (Lahti et al. 2005). Following this procedure, 1D inversion of magnetotelluric data from site B42 was accomplished using two approaches (Lahti et al. 2005). The resulting models (smooth Occam model and a fivelayer model with a minimum number of layers required by the data) and their comparison with some other models are shown in Figure 5. The three main features of the model are: (1) the presence of a middle to lower crustal conducting layer; (2) an abrupt increase of conductivity at a depth of c. 170 km; (3) rather low conductivity (c. 100 [l m) of the mantle lithosphere. The cause of the middle to lower crustal conductor at site B42 is difficult to explain because, according to geological mapping, the site is located in an area of Archaean crust. The result suggests that either the Archaean lower crust can be conductive, in contradiction to previous results (e.g. Jones 1992), or an unknown process has affected the lower crust since Archaean times and made the middle to lower crust conductive. The resistivity of the mantle lithosphere is c. 100 l) m beneath site B42, although it should be noted that the actual resistivity of a layer below a conducting layer (middle to lower crust in this case) is difficult to obtain. Yet the resistivity is compatible with the results from site KIR in northern Sweden (Jones 1982, 1983), but 10 times lower than in the central part of the Shield in Karelia (KAR, Korja & Koivukoski 1994). Similarly, both at sites B42 and KIR, an abrupt increase in conductivity is detected at depths of 170 and 150 km, respectively, whereas at KAR no such interface is found in the uppermost 200 km. This indicates, as pointed out above, that there exist considerable lateral variations in the electrical properties of the upper mantle in Fennoscandia. The olivine conductivity profile (Xu et al. 2000) is shown in Figure 5. Comparison of this with the models of B42 and KIR shows that the mantle lithospheric conductivities are higher in the northern part of the Shield than predicted by the dry olivine model, in particular in the region of enhanced conductivity below 170 km at B42 and 150 km at KIR. In contrast, at KAR in the central part of the Shield, the model resistivities are in agreement with the dry olivine model, at least at a depth of 200 km. Comparison of results from Fennoscandia with those from the Canadian Shield (Schultz et al. 1993; Hirth et al. 2000; Neal et al. 2000) indicates that the mantle lithosphere in Fennoscandia is roughly 10 times more conductive than the Archaean lithosphere beneath the Canadian Shield, whereas in the central part of the Fennoscandian Shield (KAR), the resistivity is similar to that in the Canadian Shield. In summary: (1) upper mantle conductivity is laterally heterogeneous in the Fennoscandian Shield; (2) there must be a layer of enhanced conductivity in the upper mantle beneath the entire Fennoscandian Shield, which has 10-100 times higher conductivities than predicted by the dry olivine model; (3) the depth to the top of the conducting layer (or a region of enhanced conductivity) is 1 5 0 - 1 7 0 k m in the northern part of the Shield, whereas in the central part of the Shield it must be deeper than 200 km; (4) the conductivity of the mantle lithosphere, above the conducting layer, is roughly 10 times higher than the dry olivine conductivity in the northern part of the Shield, whereas in the central part of the Shield the conductivity is comparable with the conductivities of the dry olivine model. Magnetotelluric data from the BEAR array are strongly anisotropic, in particular in the central part of the Fennoscandian Shield; that is, the data yield stable geoelectric strikes (50~
547
and large phase split (30-45~ This led Bahr & Simpson (2002) to suggest that the upper mantle is electrically anisotropic in the Fennoscandian Shield. Similarly, using earlier data, Rasmussen (1988) and Korja & Hjelt (1998) have suggested that the deep crust and upper mantle might be electrically anisotropic in the Fennoscandian Shield. Three-dimensional modelling using the 3D crustal model compiled in the BEAR project (Korja et al. 2002) shows, however, that isotropic 3D crust and isotropic, layered mantle can explain nearly all the observed anisotropic features (Korja & BEAR Working Group 2003). The isotropic 3D model produces very stable strikes of c. 50 ~ NE, as observed in the field. Similarly, the 3D model produces phase splits nearly as large as observed (30-45~ The remaining part (i.e. the part that cannot be explained by the current isotropic 3D crustal and 1D mantle models) might be due to genuine anisotropy, or due to heterogeneities in the upper mantle. It is clear, however, that the observed strikes have no bearing on the azimuth of the anisotropy and that if the upper mantle is anisotropic (the unexplained part of phase split), then the anisotropy factor (proportional to phase split) is much smaller than estimated from the original phase split.
Seismological studies M o h o m a p ( D S S profiles)
Early DSS profile data indicated significant variations of the crustal thickness in the Precambrian parts of the Fennoscandian Shield. The thickness varies between 42 and 52 km but can reach greater thicknesses up to 65 km. The map of Moho topography (Fig. 6) by Luosto (1991, 1997) has become a seminal starting point for both seismic tomography and reflection studies. Malaska & Hyvrnen (2000) improved the crustal model by interpolating and smoothing the published 2D seismic models into a 3D model. Korsman et al. (1999) analysed the 160 km wide and 840 km long GGT-SVEKA transect using all existing geophysical and geological information. The transect covers the western part of the Archaean Karelian Province, crosses its boundary zone towards the Palaeoproterozoic Svecofennian arc complex, traverses the main tectonic units of the northern part of the Svecofennian complex, and ends in the Mesoproterozoic rapakivi granite area. The main conclusion was that thinner crust (with average crustal thickness of 45 km) is found in regions that have experienced one or more anorogenic extensional events, whereas the orogenic crust of the Svecofennian Domain has much greater thickness, on average 55 km, and orogenic collapse, normally producing a thinned crust, was apparently inhibited. The crust was thickened tectonically and by magmatic under- and intraplating. The entire Svecofennian crust equilibrated soon after magmatic underplating, between 1.885 and 1.800 Ga, and mafic magmatism increased the density of the crust, helping to preserve the thickened crust (Korsman et al. 1999). A systematic reanalysis of the old DSS data by Pavlenkova et al. (2001) more or less corraborated the previous findings by Luosto (1991, 1997) and Korja et al. (1993). They divided the crust into three layers with velocities of 6.0-6.4, 6.5-6.6 and 6.8-7.0 km s -1. The interface between the two upper layers is the most stable and its depth increases from 9 - 1 0 km in the Kola province to 16-18 km in southern Finland. The boundary separating middle and lower crustal layers is very stable throughout the region and is situated at depths of 27-30 km. The thickness of the lower crust varies from 10 to 12 km. In the region of thickened crust, an additional high-velocity lower crustal layer with velocities of 7.2-7.4 km s-1 was necessary to explain the observations. Reflection profiling
The BABEL seismic experiment, in 1989, revealed a great variety of structures in the crust of the Fennoscandian (Baltic) Shield.
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S.-E. HJELT E T A L .
Fig. 6. The Moho depth map of Fennoscandia drawn from data collected by Luosto (1991, 1997), Korsmanet al. (1999) and Sandoval et al. (2003). Black dots show original data points (=sampling of velocity models) of Korsman et al. (1999).
Prominent reflections in the uppermost mantle were originally interpreted to define a mantle convergence zone where Proterozoic mantle underthrust Archaean lithosphere (BABEL Working Group 1990, 1993). A southward-dipping zone of less reflectivity was interpreted as the major strain zone accommodating horizontal shortening in the crust (Snyder e t al. 1996). Snyder (2002) has reinterpreted data from the northernmost BABEL profiles and concluded that the Archaean block forms a wedge of uppermost mantle rocks embedded in a Proterozoic block. The extent of the Archaean rocks is as great as 1 0 0 - 2 0 0 k m at Moho depths, suggesting that the Archaean lithosphere is laterally more extensive at depth than at the surface. In his model, Snyder (2002) suggested that the crustal convergence was partitioned between a wedge of weaker Archaean crust, thrusting higher in the crust to the south and channel flow within the lower crust. Altogether, he preferred a shear deformation origin to a magmatic enhancement of impedance contrasts for the 'bright Moho' reflector observed on the northernmost BABEL lines. Korja e t al. (2001, 2006) have reinterpreted the BABEL lines 1 and 6, along with new marine gravity data in the central part of the Gulf of Bothnia. The two parallel, north-south lines 1 and 6 have a transparent central area flanked by reflective structures dipping away from the centre. In the northern part of the profile, bright saucer-shaped reflectors have been interpreted as post-Jotnian diabase sills (BABEL Working Group 1993) that crop out on the sea bed (Korja e t al. 2001). Otherwise, the northern part has a complicated reflectivity pattern with a weakly
reflective upper crust and highly reflective, northward-dipping structures in the lower crust. Korja e t al. (2001) explained the reflectors in the lower crust as a double Moho structure formed by under- and intraplating of previously thinned lower crust in the northern part of lines 1 and 6. They concluded that the thickness and strength of the lithosphere were great enough to prevent the heat pulse from the mantle rupturing the crust; instead, minor extension and rifting took place. In the south, extensional shear zones are seen as a band of dipping reflectors levelling out horizontally at a depth of 40 km underneath weakly reflective areas interpreted as rapakivi granite batholiths. With these new interpretations of BABEL data it has become increasingly evident that a 3D approach to the structure of the Fennoscandian lithosphere is essential. Additional seismic reflection data crossing the geological structures of the central and thick part of the Fennoscandian Shield are necessary, and the reflection results must be supplemented with other geoscientific data, most notably with potential field data. Two major reflection profiles have been studied in the eastern parts of the Shield. Along the profile 1-EV of the International Global Geosciences Transects programme (GGT), the dataset incorporates geological, gravimetric and magnetic maps, compiled along a 100 km wide zone, and seismic (CDP and DSS), magnetotelluric and interpretational structural deep sections. The transect crossed the Karelian craton, the Palaeoproterozoic Central Russian region, and Belomorian fold belts and the Svecofenian Domain.
ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD
The preliminary interpretation of profile 4B of SVEKALAPKO has been presented by Berzin et al. (2002). The profile runs almost east-west from the southern end of the White Sea to the RussianFinnish border at the northeastern end of the SVEKA profile. In the detailed CDP cross-sections obtained from wide-angle and near-vertical reflections several inclined boundaries are traced from the surface to 2 5 - 3 0 km depth. The boundaries correlate with the well-known fault zone between the Belomorian Mobile Belt and the Karelian Craton. The DSS data also show these inclined reflectors, as well as a near-horizontal boundary at a depth of 10-15 km, under a low-velocity zone. A lower crustal boundary at a depth of 30 km has no clear expression in the CDP reflectivity pattern. Strong PmP reflections from the Moho at 40 km coincide with the boundary between reflective lower crust and transparent upper mantle on the CDP section. Berzin et al. (2002) suggested that the near-horizontal crustal boundaries and the Moho are transition zones with high-velocity gradients and not sharp discontinuities. Further investigation and modelling of profile 4B are in progress. The continuation of 4B along the SVEKA line in Finland was originally a part of the SVEKALAPKO project plan, but was not measured until 2001, as line 1 (see Fig. 1) of the Finnish Reflection Experiment (FIRE). Along FIRE 1 the lower crust is weakly reflective, which has been suggested to indicate magmatic underplating in addition to tectonic thickening. The data along FIRE 1 are of good quality, and changes in the reflectivity patterns are correlatable to surface geology (Heikkinen et al. 2003).
S V E K A L A P K O seismic t o m o g r a p h y e x p e r i m e n t
The SVEKALAPKO seismic tomography experiment consisted of a network of 128 temporary stations (40 broadband and 88 shortperiod instruments). Data from 15 permanent seismic observatories were also used. The array covered the Shield from 59 to 68~ and 18 to 34~ (Fig. 2)~ The array was designed for maximum ray density of teleseismic sources at the depth range between 100 and 300 km. From August 1998 to May 1999 more than 1300 local, regional and teleseismic events were recorded. The first results of multidisciplinary seismic tomography, anisotropy and receiver function studies of the dataset have been presented (Bock & SVEKALAPKO Seismic Tomography Working Group 2001; Bruneton et al. 2002, 2004a,b; Sandoval 2002; Funke et al. 2003; Alinaghi et al. 2003; Sandoval et al. 2003, 2004; Yliniemi et al. 2004; Plomerovfi et al. 2006). One of the key targets of the experiment was the upper mantle. Variations in the crustal velocities, however, distort the teleseismic wave fronts, causing spatial travel-time variations. Unless appropriate corrections for the crustal effects are used, the latter are back-projected during the inversion and lead to artefacts in the derived structure of the upper mantle. Only the corrected teleseismic travel-time observations were inverted for mantle structure. Comparing the inversion results of the synthetic travel-time dataset, with and without crustal corrections, demonstrates the need to apply appropriate 3D crustal corrections in high-resolution regional tomography for upper mantle structure beneath the Fennoscandian Shield (e.g. Sandoval 2002; Bruneton et al. 2004b). Crustal models. Sandoval (2002) prepared a refined 3D crustal velocity model from existing DSS data; for example, Moho topography and lateral variations in average velocity as defined by Luosto (1991, 1997), Korja et al. (1993) and Korsman et al. (1999). Other geophysical information (e.g. gravity data) was not included in the model at this stage. The model was constructed by first determining the Moho interface and topography, followed by calculation of the 3D velocity structure. In the case of
549
SVEKALAPKO, the anomalous high-velocity lower crust demanded an additional step. The Svecofennian crust is thicker than the crust of the Karelian or Lapland-Kola realm. The thinnest crust, with a thickness between 38 and 42 km, surrounds the deep central Svecofennian Domain. The lack of sedimentary rocks on the surface is a major advantage when constructing a priori models. In the eastern parts of the Shield, where recent data were scarce, a uniform Moho depth of 42 km was assigned. Weighting and interpolating the data a Moho depth uncertainty of __%2 km at minimum was obtained for the highest-quality reflectors and a Moho uncertainty of +_ 10 km for the lowest-quality reflectors. The maximum crustal thickness of 64 km occurs beneath the surface contact region between the Archaean and the Svecofennian regions. A secondary maximum of crustal thickness exists beneath the western coast of Finland, with a value of 56 km. A narrow trough, with crustal thickness up to 52 km, stretches from the main maximum and becomes shallower (48 km on average) to the north. In central Fennoscandia, two interfaces in the 3D velocity model are the Moho and the upper limit of the lower crust. Sandoval et al. (2003) used a constant velocity of 8.3 km s -1 at the base of the model at 70 km and a constant value of 5.9 km s -1 at the surface, both values chosen as an average value derived from the DSS experiments. The top of the high-velocity lower crust under the central part of the Baltic Shield was defined as the depth at which the P-wave velocity reaches 7.0 km s -1 (Korja et al. 1993). Increased P-wave velocities are observed just above the Moho interface. Two velocity gradients were defined, the first between the surface and the top of the high-velocity lower crust and the second between the upper limit of the high-velocity lower crust and the Moho. The S model was derived from the P model by assuming the same Moho interface for both models and by applying Vp/V~ ratios of 1.71, 1.76 and 1.78 for the upper crust, lower crust and upper mantle, respectively (Luosto 1997; Korsman et al. 1999). The analysis of Sandoval (2002) indicated that stations situated in the centre of the SVEKALAPKO array show the largest positive delays, with 0.22 s for the P model and 0.24 s for the S model. Positive delays occur in areas with thicker crust, although the retarding effect of the crust is reduced by up to 50% by the presence of fast lower crust. Early arrivals are obtained in the surrounding areas, with minimum values of - 0 . 4 3 s for the P model and - 0 . 6 2 s for the S model. The average crustal thickness here is 42 km (7 km thicker than IASP'91), but the high-velocity lower crust increases the average crustal velocity to values close to 6.4 km s -1 (compared with 6.1 km s -1 in IASP'91). The P-wave delays relative to IASP'91 have a distribution centred at - 0 . 1 5 s with maximum and minimum delays of 0.37 and - 0 . 4 9 s, respectively. The S-wave delay distribution has a similar pattern and is centred at - 0 . 3 0 s with a maximum value of 0.38 s and a minimum delay of - 0 . 7 6 s. This is an expected result, as the S model is derived from the P model. In both distributions a secondary maximum can be observed between 0.00 and 0.20 s. This population of positive delays is caused by the points that lie in the area with thickest crust (Sandoval 2002). A further improvement of the crustal model has been prepared by Kozlovskaja et al. (2004a). They used the crustal Vp velocity model of Sandoval et al. (2003), petrophysical data on the density of bedrock in Finland and new velocity data for the crust obtained from SVEKALAPKO studies of local events, a priori, information for inversion of the observed Bouguer anomaly. A four density-layer model was obtained: the upper crust varying from 2610 to 2900 kg m -3, the middle and lower crust from 2800 to 3000 kg m -3, underlying high-velocity lower crust from 3050 to 3250 kg m -3, and the density beneath the Moho boundary varying from 3250 to 3245 kg m -3. The resulting model demonstrates that the Moho depression in central and southern Finland is not reflected in the observed
550
S.-E. H J E L T E T A L .
SVEKALAPKO P-velocity 250 km depth
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Fig. 7. P-wave velocity structure beneath the Fennoscandian Shield from high-resolution teleseismic tomography. (a) Deviations from the0 IASP'91 velocity model (in percent); (b) a schematic illustration of the main structural tectonic elements of the crust and upper mantle beneath the study area (Sandoval et al. 2002).
ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIANSHIELD
Bouguer anomaly. The depressions in central and southern Finland are fully compensated, or even overcompensated, by dense mafic rocks in the crust and hence no corresponding minimum of the Bouguer anomaly is observed. On the other hand, the Moho depression beneath the Gulf of Bothnia is compensated only in the southern part, resulting in a regional-scale minimum of the Bouguer anomaly in the northern part of the depression. The different degree of compensation may result from differences in the origin and age of the Moho boundary that is generally defined by the last major tectonothermal event in the area. Thus, formation of the thick crust and present-day Moho geometry in central and southern Finland was due to several consecutive tectonic processes during the Svecofennian orogeny between 1885 and 1800 Ga that were concluded by magmatic underplating (Korsman et al. 1999). The increased density in the upper and middle crust here resulted from both mafic magmatism and thrusting of highly metamorphosed crust toward the surface. On the other hand, the formation of the thick crust beneath the Gulf of Bothnia was most probably the result of accretion of two microcontinents (or terranes?), as proposed by Lahtinen et al. (2005), who suggested that the Svecofennian domain was formed as a result of five orogenic processes in the time period 1.92-1.88 Ga and that the whole Fennoscandian segment of the lithosphere was formed by the accretion of several microcontinents. Accretion of two crustal blocks with different densities explains the different degree of compensation of the Moho depression in the northern and southern parts of the Gulf of Bothnia. Upper mantle structure: receiver function studies. Bock & SVEKALAPKO Seismic Tomography Working Group (2001) and Alinaghi et al. (2003) stacked receiver functions to enhance converted P-to-S amplitudes. The arrival times of PS indicate a considerable thickening of crust across the Trans-European Suture Zone (TESZ) from 30 km in the German Basin to over 50 km in the Fennoscandian Shield. The change in crustal thickness across the TESZ was corroborated by previous seismic studies (EUGENO-S Working Group 1988; Gossler et al. 1999; Grad et al. 2002). The pronounced asthenosphere beneath the c. 100 km continental lithosphere of West-Central Europe abruptly terminates along the TESZ, altogether the TOR seismic experiment documents a surprisingly sharp boundary of the Baltic Shield along the TESZ in Denmark (e.g. Wilde-Pi6rko et al. 2002). The variations in Moho depths beneath the Baltic Shield ranging from 40 to 60 km and established by previous controlled-source seismic experiments are observed also in the receiver function data. The two major 410 km and 660 km upper mantle discontinuities are clearly observed both under the TOR profile and underneath the SVEKALAPKO network (Figs 8 and 9). Whereas the arrival times of converted P to S waves from 410 km and 660 km discontinuities undergo changes across the TESZ, at the southern edge of the Shield, the difference between the arrival times of P410s and P660s phases increases below the Shield. This is indicative of cooler upper mantle underneath the Precambrian Fennoscandian Shield than that of the Palaeozoic North German Basin. Across the SVEKALAPKO profile the thickness of the transition zone shows signs of increase in the eastern part whereas traces of some local anomalies can be found in the central parts (Bock & SVEKALAPKO Seismic Tomography Working Group 2001). However, generally the variations of depth to the 410 km and 670 km boundaries are small beneath the Shield. This result, together with thermal models of the lithosphere beneath the SVEKALAPKO area and thermobarometric data on mantle xenoliths in eastern Finland (Kukkonen et al. 2003), allows us to suppose that no significant temperature variations exist in the upper mantle beneath Finland. Upper mantle structure: teleseismic P-wave tomography. Using a non-linear teleseismic tomography algorithm, Sandoval (2002)
Moho? /
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551
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Fig. 8. Receiver function north- south traces from the TOR and SVEKALAPKO (19 station subset) seismic tomography arrays. Dark streaks are converted P-S waves emerging from boundaries with depth-increasing velocity. The change in distance scale and the break between the array profiles are indicated by the bold line (Bock & SVEKALAPKOSeismic Tomography Working Group 2001).
and Sandoval et al. (2003, 2004) found P-wave velocity variations of up to 4% throughout the SVEKALAPKO region. A positive velocity anomaly can be followed down to about 300 km depth beneath the centre of the array (Fig. 7) that correlates very well with the region of thickened crust. Sandoval et al. (2004) interpreted this as the signature of the deepest-reaching tectosphere beneath the Shield. The Archaean-Proterozoic suture zone does not show up as a perceptible structure in the mantle (Sandoval 2002; Sandoval et al. 2003, 2004). Because both thermal modelling of the lithosphere and receiver function studies revealed no significant temperature variations beneath the Shield, and the lithosphere has been in place since about 1.5 Ga, the high-velocity anomaly was interpreted as a continental 'keel' stabilized by compositional differences. Differences may have been created by extraction of melt during the formation of the thick lower crust in this region. However, interpretation of velocity variations revealed by teleseismic P-wave tomography in terms of composition is difficult, because the tomography reveals only relative values of P-wave velocities that cannot be directly compared with the values revealed by petrophysical studies of
552
S.-E. HJELT ETAL. Lsfitude (deg)
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The broadband part of the SVEKALAPKO seismic array showed clear signs of seismic anisotropy: time delays between the fast and slow shear split waves, and varying directions of the fast S polarization of the incoming teleseismic waves. In general, shear-wave splitting was detected at most stations and measured time delays of the slow shear waves were between 1 and 2 s, on average. Lateral variations of shear-wave splitting parameters indicate lateral variations of the anisotropic structure of the upper mantle beneath the SVEKALAPKO seismic array. The mantle lithosphere seems to consist of several large-scale domains with different orientation of anisotropy (Plomerovfi et al. 2002a,b, 2006). The splitting parameters are coherent within the individual lithospheric blocks, but vary from one block to another. According to Plomerovfi et al. (2002b), a subcrustal lithosphere of about 150 km is thick enough to accommodate this observed large-scale seismic anisotropy. Additional information about an anisotropic upper mantle structure was obtained from combined analysis of S-wave splitting parameters and direction-dependent P-wave residuals. In particular, strong anisotropy and uniform orientation of anisotropic material in the upper mantle was revealed beneath the Archaean domain. In contrast, the anisotropic pattern corresponding to the Proterozoic domain is more heterogeneous and weakly anisotropic (Plomerovfi et al. 2006; Kozlovskaya et al. 2006). Both large Archaean and Proterozoic tectonic units of the eastern part of the Shield seem to be composed of several smaller lithospheric domains with different orientation of large-scale mantle fabric that may result from different geological history.
59
O10
Fig. 9. Time domain (a) and migrated (b) sections of move-out corrected receiver functions along the TOR and SVEKALAPKOarrays, stacked in 50 km wide windows with moving intervals of 10 krn providing an 80% overlap between adjacent windows. The dark positive amplitudes represent an increase of the S-wave velocities with depth (Alinaghi et al. 2003).
mantle rocks. In addition, Sandoval et al. (2003, 2004) inferred that the anomalous upper mantle velocities are well defined horizontally, but vertically smeared both upwards and downwards. Sandoval et al. (2003) also assessed the influence of the crustal correction on the resolution of the upper mantle structure. Two inversion tests were used, first a synthetic dataset was inverted (1) for crustal and mantle structure combined and (2) for mantle structure only, after correction for crustal effects. They demonstrated a strong 'leakage' of crustal effects down to 200 km. The effects can still be noticed at 450 km depth. However, introducing crustal corrections allows significant reduction in the effect of the crust on upper mantle velocities. Therefore, improved crustal models will be essential in improving upper mantle models beneath the Fennoscandian Shield.
Upper mantle structure: seismic anisotropy. Anisotropy of seismic velocities can result for various reasons; for example, from oriented fractures in the upper crust, from alternating layers with different isotropic velocities, or because of the alignment of crystals of rock-forming minerals in a stress field. A highly heterogeneous crust can contribute only up to about 10% of the observed large-scale anisotropy. According to the analysis of Plomerovfi et al. (2001, 2002a,b, 2006) the main source of the large-scale anisotropy beneath Scandinavia has to be in the upper mantle, caused especially by its large-scale fabric owing to preferred orientation of olivine.
Upper mantle structure: surface-wave investigations. Surface-wave studies have an advantage over teleseismic body-wave tomography, because they allow estimation of the absolute values of S-wave velocity in the upper mantle that are directly comparable with values estimated by studies of upper mantle xenoliths from Finland (Kukkonen et al. 2003) and from other Precambrian areas (Weiss et al. 1999; Griffin et al. 2003). Bruneton et al. (2002, 2004a,b) used data from the 2D grid of 46 broadband stations of the SVEKALAPKO array. Fundamentalmode Rayleigh wave arrival times with periods between 10.5 and 190 s were used to investigate the S-wave velocity in the upper mantle beneath the SVEKALAPKO array. Joint inversion for the S-wave velocity model under the array and the shape of incoming wave fronts reduced the artefacts caused by structure outside the study region (Fig. 10). The results of inversion for the upper mantle seem to be very well constrained to a depth of 150 km and weakly dependent on crustal thickness (Bruneton et al. 2004a,b). A regional average 1D shear-wave velocity model for the SVEKALAPKO area to a depth of 300 km (Bruneton et al. 2004b) has S-wave velocities that are c. 4% faster than in standard Earth models. The model lacks a substantial low-velocity layer that could define the base of the lithosphere. This indicates a cold upper mantle beneath the SVEKALAPKO array and agrees with the results of Sandoval (2002), Alinaghi et al. (2003) and Sandoval et al. (2004). The 3D S-wave velocity model (Bruneton et al. 2004a) shows both lateral and vertical S-wave velocity variations ( -t- 3%) that can be explained by variations of composition of upper mantle peridotites; for example, different modal proportions of rock-forming minerals (mainly olivine and orthopyroxene) and differences in M g / ( M g + Fe) ratio (Weiss et al. 1999; Griffin et al. 2003). The model obtained by Bruneton et al. (2002, 2004a,b) is in agreement with the result of surface-wave studies (Fig. 11) by Funke et al. (2003). The 3D model has a mean crustal thickness of 52 km. It reveals positive and negative S-wave velocity variations, but no perceptible low-velocity zone in the upper 300 km. The absolute values of S-wave velocities in the upper mantle beneath the SVEKALAPKO area vary from 4.6 to 4.8 km s - t .
ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIAN SHIELD
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Similar values were obtained by Griffin et al. (2003) by analysis of xenolith samples from Proterozoic and Archaean upper mantle around the world. However, the models of Bruneton et al. (2002, 2004a,b) and Funke et al. (2003) do not show S-wave velocities that are systematically higher beneath Archaean domains than beneath the Proterozoic, as one could expect from global xenolith analysis (Griffin et al. 2003). Instead, both S-wave velocity models revealed a laterally and vertically heterogeneous structure for the upper mantle. Because the upper mantle temperatures beneath the Shield are uniform (Alinaghi et al. 2003; Kukkonen et al. 2003), the S-wave velocity heterogeneities identified by Bruneton et al. (2002, 2004a,b) and Funke et al. (2003) can be explained by either compositional variations or anisotropy of the seismic velocity. Both models revealed a layer of very high S-wave velocity (down to c. 120 kin) that is widespread beneath the boundary of
553
Archaean and Proterozoic domains (Fig. 11). The velocity in this layer (up to 4.8 km s -1) agrees with the value estimated from highly depleted lherzolite and harzburgite xenoliths from eastern Finland that contain about 70% olivine with a high Mg/ (Mg + Fe) ratio of about 0.9 (Kukkonen et al. 2003; Bruneton et al. 2004a). The velocity beneath this layer is 4.65-4.7 km sand lower than that estimated from xenoliths from Finland, which may be due to a less depleted composition. However, these values are also slightly lower than the velocity in Archaean mantle xenoliths reported by Griffin et al. (2006). This can be explained by seismic anisotropy revealed beneath the Archaean domain by Plomerovfi et al. (2002b, 2006) and Kozlovskaya et al. (2004b). Similar stratification of the upper mantle was revealed also beneath the Slave craton, where an ultradepleted upper mantle layer is underlain by a more typical depleted Archaean mantle with a higher orthopyroxene/olivine ratio. Griffin et al. (1999) proposed that the ultradepleted layer of the Slave mantle was generated in a collisional setting. In the Proterozoic area in Finland, the surface-wave studies revealed several domains with slightly varying stratified velocity structure. Generally, the velocity beneath the Proterozoic domain to a depth of 100-120 km varies from 4.6 to 4.7 km s -1, which is lower than the values estimated from xenoliths from eastern Finland for the same depth range. These values agree with those estimated from Proterozoic xenoliths worldwide (Griffin et al. 2003), which may indicate a more fertile composition. Beneath this layer, the velocity values agree well with those estimated from xenoliths from Finland (Bruneton et al. 2004b), indicating highly depleted mantle there. However, such a distribution of mantle material (e.g. a high-density fertile layer over a low-density depleted layer) would not be gravitationally stable. Therefore, the more feasible explanation is that the low velocity in the upper layer is due to contamination by the crustal material. Upper mantle structure: local event studies. Yliniemi et al. (2004) presented results of forward raytrace modelling of reflected and refracted P waves of the strongest local events registered by the SVEKALAPKO array. They reported two types of mantle reflections: subhorizontal and gently dipping reflectors below the Moho at a depth of 70-90 km, and phases originating from a depth of 100-130km. Based on the irregular character of reflectors of the first group, on their different spatial orientation and on a correlation with Moho offsets, they interpreted the boundaries of the first group as relicts of ancient subduction and collision processes. This explanation is in accord with that of Alinaghi et al. (2003), who did not identify any upper mantle discontinuities except for global ones at 410km and 660km. This can be explained by both masking effect of multiples of the Moho conversions and the irregular nature of boundaries at a depth of 70-80 km that do not produce coherent P-to-SV conversions. The position of the reflectors from the first group beneath the SVEKA profile coincides with the location of a highly depleted upper mantle layer (Fig. 11). The reflectors of the second group coincide spatially with an area of slight change of both P-wave velocity revealed by teleseismic tomography (Sandoval et al. 2004) and S-wave velocity revealed by surface-wave studies (Funke et al. 2003) (Fig. 11). Therefore it can be attributed to a lithological contact between a highly depleted upper mantle layer and more typical Archaean mantle. This boundary correlates also with the estimated depth to the lower boundary of the mechanically strong lithosphere (i.e. the depth at which the ductile strength is reduced to 50 MPa; see Fig. 11) (Kaikkonen et al. 2002). The rheological weakening and deviatoric stresses may result in reorientation of anisotropic minerals (mainly olivine and orthopyroxene) and a horizontal foliation, which would explain the high reflectivity at this depth (Weiss et al. 1999). The heterogeneous velocity structure of the subcontinental lithospheric mantle (SCLM) beneath the contact of Archaean
554
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4
-0.16 0.0 0,16 Variation of Vs (km/s)
and Proterozoic domains in the SVEKALAPKO area, together with traces of former subduction and collisional processes, suggests a complex history of formation and stabilization of the lithosphere in the region. This is in agreement with the recent analysis of xenoliths representing Archaean and Proterozoic SCLM by Griffin e t al. (2003). They proposed that formation of SCLM beneath Archaean and Proterozoic domains was the result of two very different tectonic regimes. Most of the Archaean SCLM was formed by high-degree melting at sub-lithospheric depth producing thick, highly depleted volumes of buoyant upper mantle material that formed the roots of continents. This tectonic regime seems have operated until c. 2.5 Ga, after which it ended as a result of secular cooling of the Earth. Subsequently, a regime similar to that of modern plate tectonics was established, which included moderate depletion at spreading centres, subduction and cyclic delamination and replacement. The regime was operating during the late Archaean and Proterozoic, resulting in progressive modification and a heterogeneous structure of former Archaean SCLM. C o m p a r i s o n o f s e i s m i c a n d e l e c t r i c a l models. We can compare the results of the SVEKALAPKO seismic experiment and the BEAR experiments only in the area covered by both arrays (Fig. 2) in the central Fennoscandian Shield. Such a comparison makes sense only in the case when both the seismic and electromagnetic datasets are sensitive to the same structures or properties of the SCLM. Interpretation of the SVEKALAPKO seismic data has demonstrated, however, that the velocity heterogeneities beneath the cold and stable central Fennoscandian Shield are explained mainly by compositional variations in upper mantle peridotites formed as a result of ancient tectonic processes. Such compositional variations have no effect on electrical conductivity and cannot be detected by MT data. However, the partially molten asthenosphere would decrease both seismic velocities and electrical conductivity and would have an effect upon both seismic and MT data, as is observed in young and active regions. Therefore, comparison of the results of the SVEKALAPKO seismic experiment and the BEAR experiments can be used to answer the question about the possible presence of
Fig. 11. The subcrustal structure along the SVEKA profile based on 3D models of S-wave velocity in the upper 300 km of the mantle (reproduced by courtesy of Funke & Friederich 2003). Fine black lines show intracrustal boundaries and bold black line shows the Moho boundary (Korsman et al. 1999; Kozlovskaya & Yliniemi ! 999; Kozlovskaya et al. 2004a). Black triangles indicate location of the shot points of the SVEKA profile. Bold yellow lines indicate position of the upper mantle reflectors derived by Yliniemi et al. (2004). Position of the boundary of mechanically strong lithosphere (after Kaikkonen et aL 2002) is indicated by dashed blue line. The boundary of anisotropic Archaean mantle (red dot-dashline) is adopted from Plomerovfi et al. (2002b) and Kozlovskaya et al. (2004b).
partially molten asthenosphere beneath the central Fennoscandian Shield and to explain the origin of the upper mantle conductivity in this area. Seismic velocity models obtained by various seismic techniques demonstrate that there is no low-velocity layer that can be attributed to partially molten asthenosphere down to the depth of c. 300 km. This generally agrees with the interpretation of the BEAR data, showing that in the central part of the Shield the conductivity is comparable with the conductivities of the dry olivine model. However, the MT data also indicate an abrupt increase in conductivity somewhere below 200 km in the area, which is covered by both datasets. The exact depth and geometry of this feature is uncertain, and cannot be retrieved from MT data alone. Therefore, the conducting region could be caused by partial melting, or dissolved water in olivine or the 410 km phase transition. Partial melting is less plausible because seismic methods do not detect a layer that could be associated with such phenomena. Therefore this conducting feature can be caused by dissolved water, if it proves to be shallower than 410 km. In the northern part of the Fennoscandian Shield (site B42 of BEAR and KIR of Jones 1982, 1983), the enhanced conductivity is observed at much shallower depth ( 1 5 0 k m at KIR and 170 km at B42). This enhanced conductivity cannot be explained by graphite, because the graphite-diamond transition takes place at shallower depths in the region (for further discussion, see Lahti et al. 2005). Therefore, it may be caused either by partial melting (asthenosphere) or by dissolved water. However, we cannot distinguish between these two alternatives, because the area was not covered by the SVEKALAPKO array and the velocity structure of the upper mantle is poorly known at present.
Conclusions
The analysis and interpretation of the latest large-scale seismic and EM arrays on the Fennoscandian Shield are far from completed. Work is in progress on S-wave tomography of both teleseismic and local events, for anisotropy studies and
ELECTROMAGNETIC & SEISMIC TOMOGRAPHY, FENNOSCANDIAN SHIELD
improvements of the crustal velocity models. The preparation of a 3D conductivity model of the Fennoscandian lithosphere has been painstakingly complicated, but the final tests are under way. The centre of the Fennoscandian Shield is characterized by thickened crust. This is accompanied by seismic velocity anomalies that extend to at least 250 or 300 km depth. The Pand S-wave velocities seem to be up to 4% higher than in the global Earth models for the upper mantle down to 200 km. The difference between the 410 km and 660 km arrival times increases beneath the Shield, which corroborates the interpretations of a thick, cold and early stabilized lithosphere. The depth to the 410 km and 670 km boundaries is very stable, which implies that no significant temperature variations exist in the upper mantle beneath the Shield. No evidence for a substantial mantle low-velocity layer (LVL) has been obtained so far. P- and S-wave velocity inhomogeneities in the mantle lithosphere are most probably explained by compositional variations and/or by seismic anisotropy. According to electromagnetic investigations: (1) upper mantle conductivity is laterally heterogeneous in the Fennoscandian Shield; (2) there must be a layer of enhanced conductivity in the upper mantle beneath the entire Fennoscandia Shield, which has 10-100 times higher conductivities than predicted by dry olivine; (3) the depth to the top of the conducting layer (or a region of enhanced conductivity) is 150-170 km in the northern part of the Shield, whereas in the central part of the Shield it must be deeper than 200 km; (4) the conductivity of mantle lithosphere (above the conducting layer) is roughly 10 times higher than the dry olivine conductivity in the northern part of the Shield, whereas in the central part the conductivity is compatible with the conductivities of the dry olivine model. The interpretation of BEAR and SVEKALAPKO data has demonstrated that the structure of the upper mantle beneath the shield is heterogeneous; this supports the major conclusion obtained already (e.g. from the interpretation of BABEL reflection experiments and previous magnetotelluric soundings), namely, that the structure of the Fennoscandian lithosphere, in general, is highly variable and complicated. Therefore, models of the lithosphere evolution must be revised to accommodate lateral and vertical heterogeneity. The architecture of the Fennoscandian deep lithosphere is not yet known, because of inadequate spatial sampling. The SVEKALAPKO seismic tomography array was relatively small compared with the size of the Shield, although the lateral sampling interval was small. As a result, the tomography array provided detailed images of the upper mantle, but from only a rather limited region. The BEAR array, on the other hand, covered the entire Shield, but the distance between sites was large, making it difficult to define the exact location of the borders of lithospheric units. Total coverage and denser spatial sampling is therefore required for the detailed understanding of the structure of deep lithosphere-upper mantle in this craton. Among the most interesting problems remaining is to define the structure and geometry of the transition between the cratonic and oceanic lithosphere or a transition from oceanic lithosphere to cratonic tectosphere. In addition, seismic reflection data are needed especially across the crustal structures of the thick central part of the Fennoscandian Shield, and reflection results have to be complemented with other geophysical data, most notably with potential field and electromagnetic data, to obtain improved 3D understanding of the crust and upper mantle. The top-quality seismic images provided by the FIRE project as well as joint inversion of potential field and seismic data will certainly contribute to a better understanding of the birth and structure of the Fennoscandian lithosphere.
555
Appendix Participating organizations of the SVEKALAPKO Seismic Tomography Working Group CZECH REPUBLIC Geophysical Institute of CAS, Prague GERMANY GFZ Potsdam University of Stuttgart FINLAND University of Oulu University of Helsinki FRANCE University of Grenoble University of Strasbourg NETHERLANDS Utrecht University POLAND Warsaw University Institute of Geophysics of PAS RUSSIA Kola Scientific Center RAS Apatity Institute of the Physics of the Earth Moscow St. Petersburg University Spetzgeofisika MNR Moscow SWEDEN University of Uppsala SWITZERLAND Institute of Geophysics, ETH Zurich The SVEKALAPKO Seismic Tomography Working Group consists of following individuals: U. Achauer, A. Alinaghi, J. Ansorge, G. Bock, M. Bruneton, W. Friederich, M. Grad, A. Guterch, P. Heikkinen, S.-E. Hjelt, T. Hyv6nen, E. Isanina, J.-P. Ikonen, E. Kissling, K. Komminaho, A. Korja, E. Kozlovskaya, M. V. Nevsky, N. Pavlenkova, H. Pedersen, J. Plomerovfi, T. Raita, O. Riznichenko, R. G. Roberts, S. Sandoval, I. A. Sanina, N. V. Sharov, J. Tiikkainen, S. G. Volosov, E. Wieland, K. Wyegalla, J. Yliniemi and Y. Yurov.
Participating organizations and individuals of the BEAR Working Group FINLAND Finnish Meteorological Institute, Geophysical Research, Division, Helsinki, and NurmijSxvi Geophysical Observatory, Helsinki and Nurmijarvi, Finland Team leader: A. Viljanen Team members: K. Pajunp~i~i,H. Nevanlinna University of Oulu, Institute of Geosciences, and Geological Survey of Finland, Oulu and Espoo, Finland Team leader: T. Korja Team members: S.-E. Hjelt, P. Kaikkonen, I. Lahti, I. Silvola, J. Tiikkainen, E. Kozlovskaya
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S.-E. HJELT ETAL.
GERMANY Technical University of Braunschweig, Institute for Geophysics, Braunschweig, Germany Team leader: K. Roden University of Goettingen, Geophysical Institute, Goettingen, Germany
This paper is dedicated to our SVEKALAPKO colleague Dr. GiJnter Bock, who tragically lost his life in an airplane crash in November 2002. The following members of the Working Groups have made valuable contributions to the manuscript: A. Korja, S. Sandoval, A. Alinaghi, M. Bruneton, M. Engels, W. Friederich, S. Funke, V. HaRk, V. Kobzova, A. Kovtun, N. Palshin, H. Pedersen, L. B. Pedersen, J. Plomerovfi, M. Smimov, E. Sokolova, I. Varentsov, and A. Zhamaletdinov. We wish to express our thanks to two anonymous reviewers for fruitful comments.
Team leader: K. Bahr Team member: E. Steveling
References
GeoForschungsZentrum-Potsdam, Potsdam, Germany
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The Svecofennian orogen: a collage of microcontinents and island arcs A N N A K A I S A KORJA l, RAIMO LAHTINEN 2 & MIKKO NIRONEN 2
1Institute of Seismology, P.O. Box 68, FI-O0014 University of Helsinki, Helsinki, Finland (e-mail: Annakaisa.Korja@ helsinki.fi) 2Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland
Based on an integrated study of geologicaland geophysicaldata, a tectonic model for the Palaeoproterozoic evolutionof the Svecofennianorogen within the FennoscandianShield at the northwestern corner of the East European Craton is proposed. The Svecofennian orogen is suggested to have formed during five, partly overlapping, orogenies: Lapland-Savo, Lapland-Kola, Fennian, Nordic and Svecobaltic. The Svecofennianorogen evolved in four major stages, involving microcontinentaccretion (1.92-1.88 Ga), large-scale extensionof the accreted crust (1.87-1.84 Ga), continent-continent collision (1.87-1.79 Ga) and finally gravitationalcollapse (1.79 and 1.77 Ga). The stages partly overlapped in time and space, as different processes operated simultaneouslyin different parts of the plates. In the Lapland-Savo and Fennian orogenies, microcontinents (suspect terranes) and island arcs were accreted to the Karelianmicrocontinent,which itself was accreting to Laurentiain the Lapland-Kola orogeny. The formation of the Svecofennian orogen was finalizedin two continentalcollisionsproducingthe Nordic orogen in the west (Fennoscandia-Amazonia)and Svecobaltic orogen in the SSW (Fennoscandia-Sarmatia).The collisionswere immediatelyfollowed by gravitationalcollapse. Abstract:
Orogeny is, by definition, a process of creation of mountain belts by tectonic activity (Bates & Jackson 1995). Orogenic belts are characterized by folding, faulting, regional metamorphism and igneous activity. In terms of plate-tectonic theory, orogenic belts mark sites of continent-continent or continent-island arc collision zones at convergent, destructive plate margins. At first, convergence is accommodated by subduction and later by tectonic thickening of one or both of the plates (Fig. la). When converging plates are moving at oblique angles, major strike-slip faults parallel to the subduction zone will develop. These faults may develop into transform plate boundaries where pieces of the margin are transported along strike. If such smaller fragments bordered by fault zones are later recognized in an orogen, they are interpreted to be suspect, exotic or translated terranes. Good examples are found along the west coast of North America (Jones et al. 1983). Oblique convergence also initiates transtensional regimes where microplates, including remnants of the continent and/or island arcs, are formed (e.g. Woodlark Basin in Indonesia; Hall 2002). The fate of the microplate is to collide either with the parent continent or with another continent on the other side of the nascent ocean. Good examples of the latter are found in the Tethysides (Stampfli & Borel 2002; von Raumer et al. 2003). Suspect terranes can also be formed by escape tectonics where smaller fragments of colliding plates are pushed aside from the main collision front into areas of thinner crust and lithosphere, and thus the orogen spreads laterally via major strike-slip faults. Orogenies are referred to as either collisional or accretionary (Windley 1993) depending on the dominant type of colliding plates. Collisional orogenies occur when large continental plates collide and in these the crust is mostly reworked. The formation of an accretionary orogen is more diverse, as it may involve the accretion of arc terranes formed along long-lived convergent margins, of exotic terranes split from neighbouring continents, and of oceanic seamounts. Lateral growth of continental plates mainly takes place in accretionary orogens. In both types of orogenies, the thickening of the crust and lithosphere may spread to the adjacent areas by escape tectonics. In general, all major orogenies begin in the accretionary modes at convergent margins and some of them evolve into collisional ones. Plate-tectonic theory accounts poorly for the effect of gravity and gravitational instabilities produced by the thickening of the crust. Thermal instabilities are also induced during orogenies. These anomalies are the driving force of gravitational collapse (Rey et al. 2001), whereby thicker orogenic crust is thinned, leading to the thickening of the adjacent crust (Fig. l b - d ) .
Another stabilizing phenomenon that may take place is lithospheric delamination, in which the cooler and denser parts of the thickened lithosphere detach and sink and are replaced by warmer and lighter asthenospheric material (e.g. Platt & England 1994). This process leads to increase in heat flow, magmatic underplating of the crust, and regional extension (Fig. l e - h ) . The orogenic cycle includes pre-collisional, syncollisional and post-collisional tectonic and magmatic stages in the plate-tectonic framework. Emphasizing gravity, the orogenic cycle is characterized by thickening of the crust, thermal maturation, partial melting, and syn- to post-convergence gravitational collapse (Vanderhaeghe & Teyssier 2001). The pre-collisional tectonics includes subduction of oceanic material and obduction of ophiolites. The syncollisional tectonics involves the colliding of accreted terranes or continents and associated thickening of the crust as well as lithospheric mantle, and post-collisional tectonics involves continued indentation of the colliding terranes or continents, and finally gravitational collapse. Although new continental crust is produced during the pre- and syncollisional phases of accretionary orogeny, the post-collisional phase is particularly important because, during this phase, many of the lithological associations are exhumed to higher structural levels. The major Phanerozoic orogens of the world are linear belts (Appalachians-Caledonides, Western Cordilleras, Alps, Himalayas), which have formed over long periods (of the order of 100 Ma). In detail, however, the orogens were formed in sequential short-lived (10-20 Ma) tectono-metamorphic events (such as the Acadian, Grampian and Laramide events), often referred to as separate orogenies. As in the Phanerozoic, the major Proterozoic orogenic belts were also formed over long periods and during semi-continuous processes (Windley 1993). One example of the proposed long-lasting orogenies is the Svecofennian (100 Ma; Ga~il & Gorbatschev 1987; Fig. 2). Although the Svecofennian has been classified as an accretionary orogeny (Windley 1993), the synorogenic, late-orogenic and postorogenic terminology above suggests that one continuous orogenic process formed the Svecofennian orogen. The first plate-tectonic model for the Svecofennian orogen was created by Hietanen (1975), who compared it with the Western Cordilleras of the USA. Since then, most of the plate-tectonic models have concentrated on the evolution of the ArchaeanProterozoic boundary, an important suture zone (Fig. 2; Bowes & Ga~il 1981; Koistinen 1981). Three models have been used to explain the development of the suture: (1) continental arc-continent collision (e.g. Gafil 1990; Lahtinen 1994);
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. EuropeanLithosphereDynamics. Geological Society, London, Memoirs, 32, 561-578.0435-4052/06/$15.00 9 The Geological Society of London 2006.
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(g)
Gorbatschev & Bogdanova (1993) in outlining continuations of the Fennoscandian Palaeoproterozoic lithologies beneath areas covered by Phanerozoic rocks. They also defned Fennoscandia's southern boundary towards Sarmatia. This paper focuses on the processes associated with the formation of the Svecofennian orogen. It is a complementary paper to that by Lahtinen et al. (2005), which emphasized the plate kinematics forming the Fennoscandian Shield. It is suggested here that the so-called Svecofennian orogeny involved a sequence of accretionary and collisional events or orogenies that partly overlap in time and space, but resulted in different structural grains. It is also suggested that Svecofennian orogeny began with accretionary tectonics and was followed by continental collision tectonics. An attempt is made to define the colliding exotic or suspect terranes (microcontinents), arcs and continents. The evolution of the Svecofennian orogen is divided into five sequential events, here called orogenies. An attempt is also made to recognize extensional stages alternating with collisional ones during the Svecofennian evolution.
Geological background
Fig. 1. Orogenicprocesses forming or reworking the continental crust. (a) Thickening by stacking of continental slivers (after Coward 1994); (b-d) orogenic collapse (after Rey et al. 2001); (e-g) lithosphericdelamination (after Dewey 1988); (h) magmaticunderplating.After thickening, via stacking of continental slivers or suspect terranes (a), the crust is in gravitational potential disequilibriumwith its surroundings (b). The thickened crust may be thinned by upper crustal extension and sliding of material to the sides (e), or the lower crust may flow sideways(d). If the thickenedlithosphere is denser than its surroundings (e), then it may delaminate (f), and it may be replaced by asthenosphericmaterial from the sides (g). The asthenospheric material may initiate partial melting of the upper mantle. Rising melts may cause mafic underplating of the crust.
(2) back-arc-retro-arc basin development related to NE-directed subduction, occurring further to the SW of the suture (e.g. Hietanen 1975; Ga~il, 1986); (3) strike-slip movement, where all parts of the Svecofennian orogen are considered exotic (e.g. Park 1985). Wilson (1982) suggested an Andean-type plate-tectonic model for the Swedish part of the Svecofennian orogen. The increasing number of isotopic and geochemical datasets in the 1990s allowed more detailed and more complex models for the Svecofennian evolution. Lahtinen (1994) defined three arc complexes and three collisional events at 1.91-1.90 Ga, 1.89-1.88 Ga and 1.86-1.84 Ga and Nironen (1997) presented the first kinematic model for the Svecofennian orogen. Tectonic models for the contemporaneous Lapland-Kola orogen, situated in the northeastern part of the Fennoscandian Shield, fall into two groups: (1) models with the suture zone within the Lapland Granulite Belt (LGB) and subduction towards the NE (Fig. 2; Barbey et al. 1984; Krill 1985; Daly et al. 2001; Daly et al. 2006); (2) models with the suture zone within the Imandra Varzuga-Pechenga Belt (IVB and PeB in Fig. 2) and subduction towards the SW (Berthelsen & Marker 1986a; Marker 1990). Crustal-scale geophysical data in the 1980s added the vertical dimension and inspired correlation with modem analogues (BABEL Working Group 1990). In an integrated geologicalgeophysical study, Korja et al. (1993) attempted to locate the sutures and terrane boundaries within the Svecofennian orogen and proposed mantle underplating to account for the thick crust in central Finland. Later, Korja (1995) suggested that orogenic collapse may have played a role in the crustal evolution of southern Finland. Drill-core and geophysical data guided
The East European Craton (EEC) is composed of the FennoscandJan, Sarmatian and Volgo-Uralian crustal segments (Gorbatschev & Bogdanova 1993), of which the last two are mainly covered by Phanerozoic platform sediments. The Fennoscandian segment is exposed in its northern and central parts (Fennoscandian Shield), and covered by platform sediments in the south and by the Caledonides in the west (Fig. 2). This study concentrates on the Finnish and Swedish parts of the Fennoscandian Shield, with less emphasis on the Kola area, which has been described by Daly et al. (2006). The evolution of the other parts of the EEC has been described by Bogdanova et al. (2006) and Claesson et al. (2006). Ga~il & Gorbatschev (1987) have divided the Fennoscandian Shield into the Karelian, Belomorian and Kola Provinces, Svecofennian Domain, Transscandinavian Granite-Porphyry belt, Southwest Scandinavian Domain and Caledonides (Fig. 2a). Traditionally, the Archaean bedrock in the Fennoscandian Shield includes two cratonic nuclei, the Karelia and Kola Provinces, dispersed as fragments and subsequently reassembled during the Palaeoproterozoic (Ga~il & Gorbatschev 1987). Based on the existence of the Baltic-Bothnian megashear along the Swedish-Finnish national boundary (Berthelsen & Marker 1986b), lithological differences across the boundary, and especially the existence of the ophiolite-bearing Kittil~i allochthon (KA; Fig. 2), Lahtinen et al. (2005) have proposed that another unit, the Norrbotten craton, in the western part of the Karelian Province, may be a separate block (Fig 2b). In this paper, the Karelian Province is divided into the Karelian and Norrbotten cratons and intervening Proterozoic terranes. The Karelian craton (see Fig. 5b) encompasses the eastern part of the Karelian Province and the Belomorian Province (Fig. 2a). The Karelian craton consists of Archaean granitoid-gneiss complexes and supracrustal rocks (e.g. greenstones) ranging in age between 3.2 and 2.5 Ga (Sorjonen-Ward & Luukkonen 2005). In the Palaeoproterozoic, the Archaean lithologies were intruded by layered mafic intrusions (2.5-2.1 Ga), A-type granitoids (2.52.4 Ga) and mafic dykes (2.4-1.97 Ga). Autochthonous supracrustal rocks ranging from quartzites to pelites and mafic volcanic rocks were deposited on the Archaean basement from 2.45 Ga onwards. Allochthonous younger units comprising greywackes, the c. 1.95 Ga ophiolites at Outokumpu (O) and Jormua (J), as well as the Kittil~i allochthon (KA in Fig. 2) composed partly of oceanic crust (Koistinen 1981; Kontinen 1987; Peltonen et al. 1996; Hanski & Huhma 2005), were thrust onto the craton and its cover at 1.9 Ga. The enigmatic Central Lapland Granitoid Complex (CLGC; 1.85-1.77Ga) covers large areas of the
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Fig. 2. Simplified geological map of the Fennoscandian Shield, based on Koistinen et al. (2001). The shear zones are mainly interpreted from magnetic and gravity maps (Korhonen et aL 2002). (a) major geological units of the Fennoscandian Shield, after Gafil & Gorbatschev (1987). N, Northern Svecofennian Subprovince; C, Central Svecofennian Subprovince; S, Southern Svecofennian Subprovince. (b) Archaean cratonic terranes of the Shield. Archaean units: Norrbotten craton, Kola craton and Karelian craton, including Belomorian. Palaeoproterozoic units in Kola peninsula: IA, Inari area; PeB, Pechenga Belt; IVB, Imandra Varzuga Belt; UGT, Umba Granulite Terrane; TT, Tersk Terrane. Palaeoproterozoic units in Finland: LGB, Lapland Granulite Belt; KA, Kittili allochthon; CLGC, Central Lapland Granitoid Complex; SB, Savo Belt; CFGC, Central Finland Granitoid Complex; TB, Tampere Belt; HB, Hime Belt; UB, Uusimaa Belt. Palaeoproterozoic units in Sweden: SD, Skellefte district; BB, Bothnian Basin; BA, Bergslagen area; S6B, S6rmland Basin; OJB, Oskarshamn-J6nk6ping Belt; TIB, Transscandinavian Igneous Belt. J, Jormua; K, Knaften; O, Outokumpu; R, Revsund. BBZ, Baltic-Bothnia Megashear; HSZ, Hassela Shear Zone; LBZ, Ladoga-Bothnia Bay Zone.
564
A. KORJAETAL.
Karelian craton. Palaeoproterozoic (1.9-1.8 Ga) plutonic rocks intruded large areas of Palaeoproterozoic cover sequences in Northern Sweden and Northern Finland (Haapala et al. 1987; Ohlander & Ski61d 1994; Perttunen et al. 1996; Bergman et al. 2001; Nironen 2005). The Belomorian Belt and its boundary zones were strongly reactivated in the Palaeoproterozoic (Gafil & Gorbatschev 1987; Bibikova et al. 2001). The area between the Karelian and Kola cratons comprises a mixture of Archaean, reworked Archaean and Palaeoproterozoic terranes. The following Palaeoproterozoic terranes are identified in the Lapland-Kola area (Fig. 2): Inari area (IA), Lapland Granulite Belt (LGB), Umba Granulite Terrane (UGT), and Tersk Terrane (TT) (Korsman et al. 1997; Daly et al. 2001; Koistinen et al. 2001). The Archaean Kola and Karelian cratons have also been affected by Palaeoproterozoic magmatic activity and deformation. These are expressed as calc-alkaline arc magmatism in the Inari area (1.94-1.93 Ga; Barling et al. 1997) and Tersk Terrane (c. 1.96 Ga; Daly et al. 2001) and as strong crustal reworking and metamorphism in the Lapland Granulite Belt, the Umba Granulite Terrane, and the Belomorian Belt (Bibikova et al. 2001). The largest Palaeoproterozoic rift-related belt in the Kola craton, the Imandra Varzuga-Pechenga Belt (IVB and PeB in Fig. 2; e.g. Berthelsen & Marker 1986a), which displays a long evolutionary history from 2.5 to 1.8 Ga (Melezhik & Sturt 1994), is located near or at the Kola-Karelia contact zone. Based on lithological associations, the Svecofennian Domain (Fig. 2a) has been further divided into Northern, Central and Southern Subprovinces (Ga~l & Gorbatschev 1987). At the southern rim of the Northern Svecofennian Subprovince, the oldest lithological units are a few remnants of older volcanic rocks and granites (> 1.95 Ga; Knaften (K)) south of the Skellefte district (SD; Wasstr6m 1993, 1996; Eliasson & Str~ing 1998), and 1.92 Ga tonalites interlayered with volcanites and turbidites in the Savo Belt (SB; Lahtinen 1994; Korsman et al. 1997). The 'primitive' arc complex of Lahtinen (1994), or the Savo oceanic island arc, was later intruded by synkinematic granitoids between 1.89 and 1.88 Ga, and by post-kinematic pyroxene-bearing granitoids starting at 1.885 Ga (H61tt~ et al. 1988; Nironen & Front 1992; Kousa et al. 1994). The Skellefte district (SD) in the northwestemmost part of the Northern Svecofennian Subprovince is composed of two groups of calc-alkaline metavolcanic and metasedimentary rocks intruded by a variety of granites. The older volcanic rocks (1.89-1.88 Ga) were deposited in a marine environment and the younger volcanic rocks (1.88 Ga) in a continental extensional environment. The volcanic sequences were intruded by granitoids at 1.89 Ga, 1.881.86 Ga and 1.80-1.78 Ga, and deformed in three stages, at 1.87, 1.8 and 1.79 Ga (Weihed et al. 1992, 2002; Allen et al. 1996b; Bergman Weihed, unpub, data). In the eastem part of the Central Svecofennian Subprovince, a more continental arc environment is found in the Central Finland Granitoid Complex (CFGC), comprising mainly calc-alkaline I-type granitoids (1.89-1.88Ga) with minor amounts of mafic plutonic rocks as well as remnants of deformed sedimentary and volcanic rocks. Later, the CFGC was intruded by a younger group of hybabyssal rocks as well as post-kinematic granitoids at 1.88-1.87 Ga (Huhma 1986; Elliott et al. 1998; Nironen et al. 2000; Nironen 2003). Based on Sm-Nd (~N~(1.9~ -- 1.6 to +0.6) and geochemical data, an older protolith (c. 2.1-2.0Ga) for the 1.89-1.87 Ga granitoids in the CFGC has been proposed (Lahtinen & Huhma 1997; R~im6 et al. 2001), indicating an older crustal nucleus. At the southern rim of the CFGC, calc-alkaline granitoid rocks and arc-type volcanic rocks (1.90-1.88Ga) are found in the Tampere Belt (TB), which has been interpreted as a mature arc, or to have been formed close to a continental margin. Migmatites with tonalite leucosome (1.89-1.88 Ga) south of the TB have been interpreted as remnants of an accretionary prism (K/ihk6nen 1987; Lahtinen 1994; Korsman et al. 1999).
The Bothnian Basin (BB), in the western part of the Central Svecofennian Subprovince, is composed of psammitic metagreywackes interbedded with black shales and minor mafic volcanic rocks as well as 1.89-1.87 Ga calc-alkaline granitoids. The peak of metamorphism and deformation was associated with the formation of migmatites and granites at 1.82-1.80 Ga (Claesson & Lundqvist 1995; Lundqvist et al. 1998). Inherited Archaean zircons and Sm-Nd (eNd(1.5) --8.5 to --5.7) data from Mesoproterozoic rapakivi granites (Andersson 1997; Andersson et al. 2002) indicate the existence of an older Archaean to Palaeoproterozoic crustal source beneath the BB. The Southern Svecofennian Subprovince includes the 1.901.89 Ga Bergslagen area (BA) and Uusimaa Belt (UB), partly formed in an intra-arc basin of a mature continental arc (e.g. K~ihk6nen et al. 1994; Allen et al. 1996a). Crustal-type Pb-isotopic composition in sulphides and Sm-Nd data (/3Nd(1.9) c. 0) indicate older (>2.0 Ga) crust in the southernmost part of the UB (Lahtinen & Huhma 1997; R~im5 et al. 2001). Similar results have been obtained from the BA (Valbracht et al. 1994). Less evolved island-arc volcanic rocks are found in the H~ime Belt (HB in Fig. 2; K~ihk6nen 2005). The subprovince is transected by a swarm of roughly east-west- to SW-NE-oriented shear zones. Typical lithologies are volcanic rocks with variable tectonic affinities, pelite-dominated sedimentary rocks, quartzites and carbonates. Plutonism shows age groups of 1.89-1.85Ga, 1.84-1.82 Ga and 1.81-1.79 Ga. The S-type late orogenic granites (1.84-1.82Ga) and migmatites with granite leucosome form a belt that extends from southeastern Finland (UB) to central Sweden (BB; e.g. Lundqvist et al. 1998; Korsman et al. 1999). To the south of the Bergslagen area lies the S6rmland Basin (S6B; Fig. 2) composed of several groups of juxtaposed supracrustal rocks. The sedimentary and volcanic rocks were formed in environments ranging from terrestrial to shallow water or marine (Beunk & Page 2001). Further south, volcanic and plutonic rocks, formed in a continental volcanic arc environment, are found in the 1.83 Ga Oskarshamn-J6nk6ping Belt (OJB; Fig. 2; Mansfeld 1996; Mansfeld & Beunk 2004). A possible continuation of the OJB is found in western Lithuania (Mansfeld 2001; Skridlaite & Motuza 2001). Southern and western Sweden is dominated by a c. 1400 km long, north-south-trending batholithic belt, the Transscandinavian Igneous Belt (TIB in Fig. 2; Patchett et al. 1987). Three age groups of volcanic and plutonic rocks (Larson & Berglund 1992; Ah~ill & Larson 2000) have been found: TIB1 (1.811.77 Ga), TIB2 (1.7 Ga), and TIB3 (1.68-1.65 Ga), of which TIB 1 are by far the most voluminous and constitute the southernmost part of the belt. Andersson (1991) and Gorbatschev & Bogdanova (1993) also included the Revsund granitoid intrusions in central Sweden in the TIB. Based on deep borehole samples, Sundblad et al. (1998) have suggested that the TIB continues through the Baltic Sea into the Baltic States. The Fennoscandian Shield becomes younger towards the west, where Gothian evolution took place between 1.75 and 1.55 Ga. Rocks in the westernmost part were reworked during the Sveconorwegian-Grenvillian orogeny at 1.15-0.9 Ga (e.g. Gorbatschev & Bogdanova 1993; Ah~ill & Larson 2000). Distribution of terranes within the Fennoscandian Shield There is growing evidence that Palaeoproterozoic crustal growth older than 1.92 Ga occurred in Fennoscandia. Geochemical and isotopic data (Valbracht et al. 1994; Anderson 1997; Lahtinen & Huhma 1997; R~im6 et al. 2001) as well as the occurrences of 1.95-2.1 Ga ages in detrital zircons (Huhma et al. 1991; Claesson et al. 1993; Lahtinen et al. 2002) indicate that continental nuclei, now seen as crustal domains or suspect terranes, had already started to form at 2.1-2.0 Ga.
SVECOFENNIAN OROGEN
Geophysical markers
It is suggested here that changes in the orientation of the Moho depth isolines (Fig. 3) broadly indicate terrane boundaries within the Fennoscandian Shield. The trends are north-south in the Karelian Province and Southwest Scandinavian Domain (SSD), and east-west in most parts of the Svecofennian Domain, but change to N W - S E south of the Bergslagen area (BA). The large crustal thickness variations within the Svecofennian Domain indicate a heterogeneous block structure (Korja et al. 1993). The thinnest parts have been correlated with the Mesoproterozoic Baltic Sea aborted rift (Korja et al. 2001). The inference that crustal thickness variations are related to suspect terrane geometries is supported by the existence of dipping wide-angle mantle reflections (Fig. 3). These moderately dipping mantle reflections, which are interpreted as frozen subduction zones, serve as criteria for suspect terrane identification within the Svecofennian orogen. Figure 4 shows a 1200 km long crosssection of the Svecofennian orogen along BABEL profiles (Korja & Heikkinen 2005). The seismic data suggest that the crust is composed of crustal terranes that are currently c. 100 km in width. The age of the crustal units decreases from Archaean in the north (Karelia) to late Palaeoproterozoic in the south (TIB). The Mesoproterozoic, extensional rapakivi event was superimposed on the collisional structure (Korja et al. 2001; Korja & Heikkinen 2005). The cross-section displays the end result of plate-tectonic processes that in detail are governed by the stiffer rheology of crustal indentors. Deep conductivity anomalies (Hjelt et al. 2006) are interpreted to represent closed basins between older crustal blocks or indentors. At present, they define major terrane boundaries. In a search for other types of terrane boundaries, lineaments from potential field data have been interpreted. The trends on both Bouguer anomaly and aeromagnetic maps (Korhonen et al. 2002a, b) show regional variation in intensity, anomaly patterns and lineament strikes. The changes are most apparent between the Karelian Province and the Svecofennian Domain, as well as between the Svecofennian Domain and TIB. An abrupt change also takes place at the southern margin of the Southern Svecofennian Subprovince.
Terrane outline
Based on lithological, geochemical, isotopic and geophysical data, it is suggested that there are several Palaeoproterozoic crustal fragments or suspect terranes within the Svecofennian orogen (Tables 1 and 2). These terranes are outlined in Figure 5. The older terranes ( > 1.92 Ga) that took part in the Svecofennian orogeny fall into three categories: Archaean cratonic terranes (Karelia, Kola and Norrbotten), Palaeoproterozoic (>2.0 Ga) microcontinents (Keitele, Bergslagen and Bothnia) and Palaeoproterozoic island arcs (Kittil~i (c. 2.0 Ga), Savo, Knaften, Inari and Tersk (c. 1.95 Ga)). Keitele and Bothnia are hidden and have no identified surface expressions. The approximate extent of the microcontinents at depth is shown in Figure 5a, and the lithological and isotope data defining the crustal terranes are given in Table 2. Later during the Palaeoproterozoic, additional new terranes were formed (e.g. Tampere Belt (TB), H~ime Belt (HB), Uusimaa Belt (UB)) and the abovementioned terranes were modified. The surface extent of the exposed terranes is outlined in Figure 5b. The Ume~ area is distinguished from the Bothnian microcontinent and Skellefte districts (SD) based on seismic reflection data, where the uppermost 15-25 km of crust is interpreted to be detached from the lower and middle crust (Fig. 4; Lahtinen et al. 2005). At the surface, an isotopic boundary delineating older rocks (> 1.9 Ga) in the northern part of the Bothnian Basin has been suggested by Rutland et al. (2001).
565
Lahtinen et al. (2005) correlated the rocks to the south of the Bergslagen area (BA) with those in the Baltic States that have similar NW-SE-striking geophysical patterns and lithological similarities, and called them the Svecobaltic area.
Model Lahtinen et al. (2005) suggested a plate-tectonic model accommodating the geophysical, geochemical and geological observations. In the model, it was assumed that the current extent of terranes was not significantly affected by collisions, implying that no account was taken of transportation along strike-slip faults, escape tectonics, or major internal stacking of the microcontinents. In the following, the model is elaborated with emphasis on processes forming the continental crust during each orogeny at a given time. The model distinguishes five stages: (1) continental rifting; (2) microcontinental accretion; (3) extension of the accreted crust; (4) continent-continent collision; (5) extensional collapse. C o n t i n e n t a l rifting
At the beginning of the Palaeoproterozoic, the Archaean Karelian continental crust was subject to large-scale extension leading to emplacement of mafic layered intrusions, granitoid intrusions and bimodal volcanism at around 2.5 Ga. This was followed by several generations of mafic dyke swarms between 2.2 and 1.97 Ga. Formation of oceanic crust in marginal basins is indicated by the 1.97-1.95 Ga Jormua and Outokumpu ophiolites. Juvenile oceanic island arc lithologies in the Savo Arc, at the ArchaeanProterozoic boundary zone, indicate opening of an ocean. To explain the abundant oceanic island arc lithologies in the Kittil~i allochthon (KA), a major rift is also suggested to have developed between the Karelia and Norrbotten cratons. An ocean between Karelia and Kola is proposed by island arc type lithologies in the Tersk Terrane and Inari Area. M i c r o c o n t i n e n t a l accretion
A microcontinental accretion stage is suggested to explain the complex assemblage history between 1.92 and 1.88 Ga in Fennoscandia (Fig. 6). After the rifting of the Archean continental plate at 2.1 Ga, new juvenile arcs were initiated. Arcs of different evolutionary stages started to assemble and to form microcontinents. By 1.89 Ga, several small plates carrying continental fragments had accreted to the Karelian continent from several directions. Oblique collision caused some of the shear zones at or close to the terrane boundaries. The accretionary processes and the dimensions of the accreting blocks are schematically illustrated by vertical cross-sections (see Fig. 8). The amalgamation of Fennoscandia started from the NE, where the Kola and Karelian cratons as well as intervening Palaeoproterozoic terranes (IA, TT, LGB, UGT) merged together during the Lapland-Kola orogeny (Fig. 6a). Simultaneously, at the western margin of the Karelian craton, the Lapland-Savo orogeny commenced as the result of the approaching Norrbotten craton and the Keitele microcontinent. Juvenile crust that had formed in island arcs (KA, SB, K) close to the continental margins was accreted before the final collision of the microcontinental nuclei. The main phase of the Lapland-Savo orogeny took place when the Keitele microcontinent collided with the Karelian craton (Fig. 6b). The westward growth of the continental collage continued with the docking of the Bothnia microcontinent (Figs 6c and 8a). The collision caused a change in the plate motions and led to a subduction switchover, with the onset of northward subduction at the southern edge of the Keitele-Karelia collage (TB; Figs 6c and 8b). The northward subduction ended when the southern ocean was consumed; the Bergslagen microcontinent was accreted to Keitele, starting the Fennian orogeny.
566
A. KORJA E T AL.
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Neoproterozoic Sveconorwegian orogenic belt (1.10 - 0.92 Ga) partly reworking Palaeo- to Mesoproterozoic rocks
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Fig. 3. Moho-depth map (Luosto 1997) and the surface projection of dipping mantle events, compiled from reflection and refraction studies (BABEL Working Group 1990; Abramovitz et al. 1997; Ostrovsky 1998; Bailing 2000; Heikkinen & Luosto 2000; Luosto & Heikkinen 2001).
SVECOFENNIAN
OROGEN
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Integrated geological-geophysical interpretation of the structure of the upper lithosphere along the EUROBRIDGE profiles G e n e r a l structural characteristics o f the crust
The interpretative models of the Earth's crust along the EUROBRIDGE EB'95, EB'96 and EB'97 seismic profiles presented in Figures 13 and 14 have been compiled from the overall geophysical and geological information, referred to in the previous sections. Collision between Fennoscandia and Sarmatia was decisive in determining the seismic characteristics of the lower crust and upper mantle in the study region, and the distribution of the magnetic and gravity anomalies (e.g. Garetsky et al. 2002). Pre-collisional terrane tectonics, in contrast, is reflected best by
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Fig. 12. Relationshipsbetween density, Vp and Vsfor the major geologicalunits crossed by the EUROBRIDGE'97 profile (after Kozlovskayaet al. 2004). The left panel shows the density- Vprelationship, the right the density-Vs relationship(for Vs, the axis is scaled by a factor of 1.73). The reference density-velocity relationshipsare shown by open stars. (a) OsnitskMikashevichiIgneous Belt; (b) Podolian Domain; (c) Volyn Domain and Korosten Pluton.
structures in the upper and middle crust. Tectonically, the rock belts and domains in the Fennoscandian terranes make up a number of 'thick-skinned' nappe packages, thrust towards the SSE 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 of N N W - N W - and NNE-NE-trending post-collisional faults and the markedly east-west-striking Mesoproterozoic faults. The EUROBRIDGE profiles suggest that the formation of highvelocity layers in the crust was commonly associated with detachment, whereas lateral undulations may have been shaped by complementary deformation of the whole lithosphere (Figs 8, 10, 13 and 14). Almost all of the high-velocity layers are accompanied by distinct subhorizontal seismic reflectors and mark sharp compositional discontinuities in the crust. In many cases, mafic sheet intrusions were responsible or contributed.
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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.
Complex crustal and upper mantle structures characterize the Fennoscandia-Sarmatia junction area beneath the CBSZ and part of the BPG, where the more ancient and more rigid crust of Sarmatia has been particularly strongly deformed, laminated and also altered compositionally. As modelling of the gravity and magnetic data shows, a substantial proportion of the post-collisional and later post-orogenic faults have listric configurations (Garetsky et al. 2002). In a number of cases, their flat-lying deeper parts coincide with nearly horizontal reflectors (detachment zones) at the boundary between the upper and middle crust. However, some large faults (e.g. the Minsk Fault in the CBSZ) appear to extend to the Moho (Figs 2 and 13).
Geophysical images of accretionary and collisional tectonics Terranes related to the Fennoscandian crustal segment. The PolishLithuanian terrane with the West Lithuanian Granulite Domain (WLG) and the Mid-Lithuanian Suture Zone (MLSZ) resembles in many respects the crustal province of southeastern Sweden, situated to the south of the classical Svecofennian orogen. In that province, the rocks were formed during several orogenic events at c. 1.841.82, 1.81-1.78 and c. 1.75 Ga, tending to young towards the south. The structural trends are dominantly W N W - E S E to westeast, and there exist active-margin volcanic-arc and back-arc
EUROBRIDGE
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 lithotectonic complexes are recognized fairly well in the seismic images. These indicate subduction towards the NNE (BABEL Working Group 1993; Abramovitz et al. 1997; Balling 2000). In Figure 13, steep dips towards the NW are indicated for the MLSZ and the adjoining area, whereas other studies, particularly those of metamorphism (see p. 603), have suggested that the WLG had been thrust towards the east, overriding the East Lithuanian Belt (EL). Skridlaite et al. (2003a) inferred that the widespread occurrence of high- to moderate-pressure granulites in the WLG indicated a subduction-collision tectonic regime between 1.84 and 1.80 Ga, whereas island-arc settings have been identified 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 a thickness of only c. 40 km. Similar differences also characterize the lower crust, which measures 10-12 km in the WLG, but only 5 - 1 0 k i n in the WLSZ (see Kozlovskaya et al. 2001; Yliniemi et al. 2001; Grad et al. 2003b). The thicker lower crust beneath the WLG appears largely to be due to the presence of a basal crustal layer with densities as high as 3.0 g cm -3 and thus probably 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 velocities of 6.1-6.4 and 6.5-6.8 km S - 1 , respectively, both in the WLG and the MLSZ. However, the upper crust in the WLG is largely granulitic, whereas that in the MLSZ is made up of various rocks in the granulite and amphibolite facies, such as blastomylonites, calc-alkaline metavolcanic rocks and gabbroic to tonalitic plutonic rocks. At depths of 1 2 - 1 5 k m (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 probably consists of Mesoproterozoic granitic rocks created by the remelting of the upper crust. A large body of a c. 1.46 Ga monzogranitoid rock is present close to the northern end of the EB'95 profile (Cerys 2004; Motuza 2005). A prominent feature of the crust particularly in the northwestern part of the WLG is its multi-layered structure, built up of distinctly delimited, conformable, persistent individual layers. The seismic velocities at the base of the crust are high (Vp 8.258.35 km s-~). Although this might suggest a 'platformal' type of crust in the sense of Christensen & Mooney (1995), lithological and 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 in the south. Major differences of rock composition and deep structure also exist between the eastern and western parts of the WLG (Kozlovskaya et al. 2001). Particularly worth noting is the recurrent granitoid magmatism both in the Palaeo- and Mesoproterozoic. The Lithuanian-Belarus terrane, including the EL and BPG, together with the Okolovo terrane forms a composite terrane where the crust is substantially thicker (up to c. 55 kin) and, as a whole, also denser than that in the Polish-Lithuanian terrane (Fig. 13). The principal mechanisms responsible for the development of this thick crust appear to have been collisional orogenic processes involving compression and folding, and the stacking of large piles of nappes in the junction zone between Fennoscandia and Sarmatia. Indications of tectonic thickening, thinning, folding and wedging-out of the rock units are common in the seismic profiles. With regard to the thickness of the upper and middle parts of the crust, the EL and BPG are not very different from the WLG, but no low-velocity layers appear to be present. Here, the seismic velocities vary substantially in accordance with lithological variation, but as most of the rocks are either mafic to
619
intermediate granulites or igneous rocks of similar compositions, the upper crustal P-wave velocities are mostly relatively high. They measure c. 6.25 km s -1 in the EL, 5.8-6.0 km s -~ in the BPG, and 6 . 1 - 6 . 2 k m s 1 in the largely metavolcanic Okolovo terrane. Substantially lower velocities are, naturally, found in the large, anastomosing systems of shear zones marked by blastomylonites and retrograde recrystallization of the granulites, and also the presence of metasediments. These occur particularly in the EL. Major west-dipping listric faults that could be traced to depths of 1 5 - 2 0 k m (see Aksamentova et al. 1994) have previously been found along the Grodno-Starobin seismic profile transecting the BPG and the Okolovo terrane in a westeast direction (see also Bogdanova et al. 2001b). In the middle crust, which has densities around 2.8 g cm -3 and P-wave velocities of 6.3-6.5 km s -1, granulites and TTG-type plutonic rocks appear to predominate. This part of the crust forms 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. 2 0 k m thick lower-crustal layer made up of mafic granulites with P-wave velocities of 6.8-7.1 km s -1 and densities of 2.9-3.1 g cm -3 (Fig. 10, Table 1). A remarkable feature is the southeasterly dips of this lower crustal layer, which appears to protrude into the upper mantle; its densitY3of 3.1-3.2 g cm -3 is substantially less than the 3.3-3.4 g c m - of the normal upper mantle in the region. Thus, this mantle offset-lower crustal protrusion may represent a 'fossilized' slab of subducted Palaeoproterozoic oceanic crust. The F e n n o s c a n d i a - S a r m a t i a junction. Within the CBSZ, the seismic and gravity characteristics of the crust change drastically across the major, west-dipping Minsk Fault (Figs 10a and 13). The latter extends to the Moho and, at the Earth's surface, separates two very different groups of tectonic units (Figs 1 and 2), the BPG and Okolovo terrane in the NW and the Vitebsk Domain (VG) and Borisov-Ivanovo Belt (B-I) in the SE. The P-wave velocities and rock densities in the upper crust are different on the two sides of the Minsk Fault, being 6.1-6.2 km s -1 and 2.7 g cm -3 in the NW and 5.8-6.0 km s- 1 and 2.60-2.67 g c m - 3 in the SE. This reflects the difference between the amphibolite- to granulitefacies mafic rocks of the Okolovo terrane and the granite-intruded metasedimentary gneisses and migmatites in the B - I . In the middle crust, the P-wave velocities are 6.3-6.5 km s -1 in the NW and 6.4-6.9 km s -1 in the SE, but there is apparently no corresponding difference in density values (Kozlovskaya et al. 2002). The best explanation appears to be that the higher P-wave velocities below the southeastern part of the CBSZ are due to a markedly laminated structure in a part of the crust where strongly deformed amphibolite- and granulite-facies rocks have been emplaced tectonically. In the CBSZ region, these diverse structural patterns in the upper to middle crust can be followed down to depths of 2 5 30 km, at which level the P-wave velocities reach 7.0 km s -1 and densities of 2.9-3.0 g cm -3 have been modelled. Farther down is a rather more uniform lower crustal high-velocity layer (HVLC in Figs 8, 10a and 13) with P-wave velocities of 7.27.4 km s -1 and densities of 3.0-3.1 kg m -3. These values correspond best to eclogitic granulites or garnet granulites. Similar high-velocity layers in the lower crust, with P-wave velocities between 7.0 and 7.7 km s - 1, appear to be common in Precambrian regions adjoining major tectonic sutures (Guggisberg & Berthelsen 1987; Korja etal. 1993; Korja & Heikkinen 1995; Funck etal. 2001; Hall et al. 2002). In some cases, they are associated with sizeable Moho offsets and are mostly explained as resulting from magmatic mantle underplating (Korsman et al. 1999; Funck et al. 2001 ; Hall et al. 2002). In the case of the CBSZ, the lower crustal high-velocity layer is a relatively young feature, as it appears to underlie adjacent terranes as well. These include the OMB, which adjoins the CBSZ in the SE, and, some 250 km farther SE, the part of the Volyn
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Domain that encloses the AMCG Korosten Pluton. In view of this spatial association, this high-velocity, high-density layer is most probably coeval with the c. 1.80-1.74 Ga Korosten body. Xenoliths from the lower crust (Markwick et al. 2001) and new data on the isotopic compositions of both OMB and Korosten rocks suggest that this high-velocity layer may be restitic, having been formed between 1.80 and 1.74 Ga by the successive removal of the AMCG melts from a 2.0 Ga lower crust similar to that beneath the OMB (Bogdanova et al. 2006). As indicated by the gravity-seismic modelling in Figures 10, 13 and 14, many crustal units in the CBSZ are wedge-shaped and imbricated, and are separated from each other by numerous distinct reflectors dipping in opposite directions. Structural patterns of this kind are characteristic of collisional-type crust (Meissner 1989; Cook et al. 1999) and therefore fit well with the location of the CBSZ region in the zone of collision between Fennoscandia and Sarmatia. A conspicuous major feature of the crust beneath the CBSZ and adjoining parts of the OMB and BPG is a large antiformal (domal) structure defined by convex high-velocity middle to lower crustal rock layers apparently related to the OMB (Fig. 13). As this structure coincides with the collisional zone between Sarmatia and Fennoscandia, it appears reasonable to assume that stacking and thickening of the crust with attendant metamorphism and gravitational instability must have been part of the early stages of its development (see Coney & Harms 1984). An important key to deciphering the formation of this antiform beneath the CBSZ and its evolution into a metamorphic core complex is the presence of a bulge of high-velocity material in the lower crustal core of the antiform (Fig. 13). In conjunction with other features, such as the association of the Korosten Pluton with the high-velocity lower crust, this suggests that c. 1.8 Ga magmatism could have been a major agent causing the doming and attendant metamorphism during the post-collisional stage. This evolution was associated with post-collisional extensional tectonics leading to c. 1.8 Ga AMCG magmatism in the OMB and the Lithuanian-Belarus terrane, local granulite metamorphism, intense mylonitization along the Minsk Fault, and, eventually, listric faulting and fast final uplift (Taran & Bogdanova 2003). Also, downfaulting of the edge of the OMB to expose the Borisov-Ivanovo and the Vitebsk tectonic units must have been part of this extension. The latter are connected with southwards dipping reflectors in the middle and lower crust of the OMB (Juhlin et al. 1996; Stephenson et al. 1996). The Archaean and Palaeoproterozoic crust of Sarmatia. The OMB was formed by voluminous magmatism between c. 2.0 and 1.95 Ga. In accordance with the OMB igneous mode of origin, the character of its c. 50 km thick crust is determined largely by the presence of numerous batholiths of granitic, granodioritic, dioritic or gabbroic composition. These obviously correspond to the seismic-velocity and density properties (Figs 12 and 13), which also suggest that the more felsic of the plutons dominate the 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 mantle xenoliths indicates that the mafic plutonic rocks of the OMB have been partly transformed into eclogite-like, garnet-bearing granulites with P-wave velocities of 6.8-7.0 km s -1 and correspondingly high densities. In addition, the OMB contains younger, c. 1.8 Ga, mostly syenitic to quartz syenitic intrusions, which are associated with the coeval AMCG-type Korosten Pluton farther south and define a belt of marked, more or less isometric, magnetic anomalies (see Fig. 2). The distribution of these intrusions appears to have been controlled 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 particular feature numerous major reflectors (Figs 8 and 13) that create
an overall multi-layered structure, presumably mostly caused by recurrent magmatism and tectonic deformation, especially in the vicinity of the Fennoscandia-Sarmatia junction. Some of the layering, however, must rather be due to the formation of the Devonian Pripyat-Dniepr-Donets Aulacogen (Fig. 1). As discussed by Stephenson et al. (1996), the system of listric faults in that structure coincides closely with Palaeoproterozoic wedge fabrics within the OMB, which indicates significant reactivation of Precambrian faults during Phanerozoic rifting. A well-preserved fine lamination of the crust also characterizes the OMB, presumably caused by deformation of rocks with contrasting elastic properties (Meissner & Rabbel 1999) and probably related to the latest major deformation event in the Devonian. In the lower crust, a high-velocity lower crustal layer with Vp of 7 . 2 - 7 . 4 k m s -1 exists also in the OMB, similar to that beneath the CBSZ, but substantially thinner and denser than in the latter. In the Volyn Domain (VD) with the large AMCG Korosten Pluton (KP), the crust is only 45 km thick; that is, substantially thinner than the 5 0 - 5 2 km crust in the neighbouring OMB and PD (Figs 8 - 1 0 and 14). The available seismic and gravity data suggest, however, that beyond the limits of the Korosten Pluton, the VD is similar to the OMB. All the crustal layers of the OMB appear to continue into and across that domain, extending southwards as far as the Berdichev Boundary Zone, which dips north and separates the Palaeoproterozoic VD from the Archaean interior parts of the Podolian Domain (PD). In the Berdichev Zone (BZ), the lower part of the Palaeoproterozoic crust appears to wedge out at depth, and the distinctly layered upper and middle parts are replaced southwards by seismically more uniform and less reflective Archaean crust. In this ancient crust there is a rather gradual increase of the P-wave seismic velocities with depth, from 6.1-6.2 to nearly 6.9 km s -1. The distribution of these velocities and the Vp/Vs ratios suggest a two-layered crust, but the S-wave data and the gravity-seismic modelling by Kozlovskaya et al. (2004) indicate the presence of three crustal layers. The relatively low Vp/Vs ratios of 1.69-1.74 indicate that, the Archaean crust of the Podolian Domain is much richer in quartz than the neighbouring Palaeoproterozoic crust (see also Yegorova et al. 2004). The crustal structures and boundaries in the Archaean part of the Podolian Domain dip gently towards the north, conforming to the inferred northerly dips of the Archaean-Proterozoic crustal boundary in the BZ. The grades of Palaeoproterozoic metamorphism increase southwards from amphibolite- and granulite-facies rocks to high-grade granulite-facies rocks. This suggests that the Pa|aeoproterozoic crust of the VD overlies the Archaean crust of the PD in a manner that may be a result of collisional tectonics at c. 2.1-2.05 Ga, followed by extension and the formation of a metamorphic core complex (Fig. 14). The greatest lateral and vertical variations of crustal composition and structure in western Sarmatia are associated with the 1.80-1.74 Ga multiphase Korosten Pluton. The influence of this intrusion is not restricted to its area of exposure, but extends for tens of kilometres in the surrounding region. All the crustal units have been updomed in a wide region (see Figs 8, 9 and 14). The gravity and magnetic modelling of the Korosten Pluton, employing also the data of the east-west-trending Geotraverse II seismic profile (Ilchenko & Bukharev 2001), indicates an extremely complex structure in the underlying crust (Bogdanova et al. 2004b). Whereas layered gabbro-anorthosite intrusions can be followed only to depths of less than 10 kin, granitoid rocks of various kinds, including rapakivi granites and monzonites, form flat-lying sheeted bodies at various levels of the upper and middle crust. The interlayering of these igneous sheets with earlier Palaeoproterozoic supracrustal and plutonic rocks is inferred to be responsible for the presence of low-velocity crustal layers with P-wave velocities of 6.1-6.5 km s -1 and densities varying between 2.6 and 2.8 g cm -3 (Fig. 14).
EUROBRIDGE At depths below 16 kin, mafic rocks form a single semicylindrical, lensoid body beneath the eastern part of the Korosten Pluton (see Fig. 14; note that the EB'97 profile crossed only the western half of the intrusion). This body measures c. 90 km across and extends to the high-velocity, high-density layer at the base of the crust, which underlies the VD, the OMB and even part of the PD (see above). This giant mafic body is considered to represent the feeding magma chamber of the Korosten Pluton (Bogdanova et al. 2004b). The underlying crust, with P-wave velocities of 7.4-7.8 km s -I, Vp/Vs of 1.77-1.79 and densities of 3.0-3.15 g cm -3, is therefore assumed to be mostly composed of mafic, ultramafic or eclogitic rocks, presumably representing a mixture between cumulates of the Korosten magma and restitic material (Bogdanova et al. 2006).
Aspects o f the mantle
With regard to Moho topography, the EUROBRIDGE seismic profiling and gravity modelling indicate depths varying between 40 and 55 km as well as a number of irregularities and offsets. Some of these were related to accretionary and/or collisional tectonics, or to superimposed late to post-collisional magmatism. The most obvious case of the former is the mantle irregularity beneath the Belarus-Podlasie Belt (VPG), which appears to connect with the Fennoscandia-Sarmatia collisional junction as defined by the Minsk Fault. Another, but so far less evident instance, may be the relationship between the Mid-Lithuanian Suture Zone (MLSZ) and a Moho offset beneath the West Lithuanian Domain (Fig. 13). An offset beneath the Volyn Domain (VD) may continue the Archaean-Proterozoic boundary in the Berdichev Zone, but coincides with and may have been masked by the root of the Korosten Pluton (Fig. 14). In addition, it is tempting to relate the Korosten magmatism to the extensive Moho uplift to 45 km beneath the Volyn Domain and the adjacent parts of the Osnitsk-Mikashevichi Belt (OMB) and Podolian Domain (PD). The upper mantle is rather inhomogeneous with regard to seismic velocities, the P-wave velocities ranging between 8.1 and 8.35 km s -1. This is probably due to lateral compositional variation from peridotite to eclogite. The latter composition is particularly characteristic of sites of collisional thickening in zones of deformation and fluidization of the lower crust and upper mantle (e.g. Austrheim et al. 1997). However, as mentioned above, a restitic origin of eclogites in the lowermost crust and upper mantle beneath the Korosten Pluton is also possible. Xenoliths in kimberlites of various ages and near-source alluvial placers have demonstrated that the upper mantle beneath the VD is made up of a 20 km thick layer of eclogites, underlain by garnet pyroxenites and peridotites (Tsymbal & Tsymbal 2003). According to Tsymbal & Tsymbal, the age of the mantle is Proterozoic beneath both the Volyn and the Podolian domains. A more eclogitic composition of the mantle beneath Sarmatia may be the reason why it has substantially higher Vp/V~ ratios than the mantle of Fennoscandia (1.83 and 1.72, respectively). Of particular interest in the tectonic interpretation of the lithosphere are reflectors in the upper mantle. Apart from some subhorizontal reflectors referrable to rheological and mineralogical changes with depth, the EUROBRIDGE transect also shows one distinct inclined reflector and more circumstantial evidence of several others. These may represent 'fossil' zones of subduction of oceanic as well as continental crust (see Balling 2000). The less distinct reflectors can to some extent be extrapolated on the basis of the Moho topography and compositional variation in the upper mantle, and also from lower crustal lenses of melting apparently related to post-collisional processes. The presence of a lens of lower-velocity and lower-density mantle beneath the edge of the Lithuanian-Belarus terrane where it faces Sarmatia thus suggests subduction into the mantle of a slab of Fennoscandian lower crust. Similar relationships, albeit less clearly
621
expressed, are also found in the upper mantle beneath the MLSZ and the adjoining parts of the WLG. Another zone of elevated mantle reflectivity is associated with the Berdichev Zone outlining the junction of the PD and VD. The distinct SSW-inclined mantle reflector beneath the OMB along the EB'97 profile (Figs 8 and 9) has previously been interpreted for purely geometrical reasons as the trace of a collisional boundary between Sarmatia and Volgo-Uralia (Thybo et al. 2003). An alternative interpretation (Aisberg & Starchik 2005) suggests that this reflector represents a detachment surface in the crust and upper mantle that was related to the formation of the Pripyat-Dniepr-Donets Aulacogen (PDDA) in the late Devonian. The latter interpretation accounts well for the part of the mantle with lower density and lower seismic velocity that overlies this reflector and is best explained as consisting of mafic and ultramafic igneous rocks. Support for this interpretation is provided by a Devonian (c. 380 Ma) age of lower crustal hornblendite xenoliths found in lamprophyric tuffites in southeastern Belarus (Markwick et al. 2001). Geochemically, these may represent remelting products of a garnetiferous mantle. Thus, it appears possible that the distinct flat-lying reflectors found at depths of c. 10 km below the undulating Moho can also be related to the Devonian event and mark the presence of mantle-derived melts (see Figs 8, 9 and 14). Similar conditions have also been observed along other seismic profiles running across the PDDA (Grad et al. 2003a; Maystrenko et al. 2003).
Conclusions The EUROBRIDGE traverse project has provided a new understanding of the structure and formation of the crust and upper mantle in the western part of the East European Craton. Although the results mostly concern the key region between the Baltic and Ukrainian shields, and the late Palaeoproterozoic collision of Fennoscandia and Sarmatia, they have relevance also for understanding the upper lithosphere in the entire EEC. The major conclusions are as follows. (1) The crust in the region between the Baltic and Ukrainian shields is Palaeoproterozoic and juvenile. It was formed between c. 2.0 and 1.8 Ga by accretionary plate-tectonic processes along the margins of the Archaean-earliest Palaeoproterozoic nuclei of Fennoscandia and Sarmatia. (2) Several Palaeoproterozoic terranes, related either to Fennoscandia or to Sarmatia, are recognized on the basis of their different ages, lithologies and tectonothermal evolution. They include various tectonic settings: juvenile island arcs, back-arcs and active continental margins. The Sarmatian terranes were formed between c. 2.2 and 1.95 Ga, whereas the Fennoscandian ones are, in general, younger, ranging between c. 2.0 and 1.8 Ga. Palaeomagnetic data indicate that the Fennoscandian and Sarmatian terranes belonged to different plates. (3) The complex, belt-shaped crustal structure and the fault zones that bound the various belts and domains mostly originated during the collision between the Sarmatian and Fennoscandian plates at some time between 1.85 and 1.80 Ga. However, the listric character of many faults and associated late to postcollisional magmatism, retrograde metamorphism and strong mylonitization along the inter-terrane boundaries as well as within their interiors all suggest that post-collisional extensional tectonics was of crucial importance for the following crustal development between c. 1.80 and 1.74 Ga, and even later at c. 1.71-1.67 Ga. Emplacement of the large AMCG plutons at 1.80-1.74 Ga in Sarmatia, and between c. 1.6 Ga and 1.50 Ga in Fennoscandia, substantially influenced the composition, petrophysical properties and geophysical structure of the crust and upper mantle. (4) The present major characteristics of the seismic profiles and potential fields in the Baltic-Belarus region were predetermined
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s. BOGDANOVAETAL.
by late Palaeo- and Mesoproterozoic accretionary, collisional and post-collisional geodynamics. The last, in particular, caused rearrangement of the lithospheric structure and shaped its present geophysical images. (5) Occasionally, the fault and suture zones between the Fennoscandian belts and domains can be traced throughout the entire crust (e.g. the Mid-Lithuanian Suture Zone). Displacements and offsets of the Moho boundary and various crustal layers along these zones, as well as the 'imbricate' character of the Palaeoproterozoic crust in southern Fennoscandia, allow comparison with 'thick-skinned' orogens. The offsets and irregularities of the M o h o boundary and lateral changes of petrophysical properties and compositions in the upper mantle may be interpreted as 'fossilized' Palaeoproterozoic zones of subduction and collision. This is particularly the case in the Central Belarus Suture Zone, between the Fennoscandian and Sarmatian terranes, where the crust is characterized by a pronounced tectonic layering and numerous reflectors. (6) The boundary between Fennoscandia and Sarmatia is defined by the major Minsk Fault, an extensional feature superimposed on the suture zone. Beneath the Minsk Fault, the crust was affected by doming of the collisionally stacked crustal layers, voluminous m a g m a t i s m at the base and the formation of a metamorphic core complex. (7) Subsequent rifting of the crust and the development of the Late Mesoproterozoic V o l y n - O r s h a Aulacogen was shallow and dispersed, roughly coinciding with the Central Belarus Suture Zone. Also, the Devonian rifting and the formation of the P r i p y a t - D n i e p r - D o n e t s Aulacogen did not cause substantial thinning of the c. 50 k m Palaeoproterozoic crust or its significant reworking. However, the underlying Palaeoproterozoic faults were reactivated and controlled the position of major listric faults (e.g. those outlining the Pripyat Trough). The SSW-dipping reflector beneath the northwestern margin of Sarmatia, thus, most probably represents a detachment surface bounding this aulacogen in the NE. The low-velocity upper mantle above this reflector is probably a Devonian mantle underplate. The EUROBRIDGE project (1994-2002) has been a highly successful WestEast co-operation enterprise. Despite the many economic difficulties in the East European countries involved, it produced a wealth of results and offered a unique experience to numerousjunior researchers. During its lifetime, geological and geophysical institutions, research councils and academies in 17 countries contributed financially. Particular thanks go to the Swedish Institute's Visby-Programme, the Royal Academy of Sciences in Stockholm and the INTAS organization (project 94-1664). Research and the workshops were always conducted in a warm and cordial atmosphere, with most participants feeling like members of a single family. In the above text, the section reporting the seismic results was compiled by M. Grad, A. Guterch and T. Janik, and E. Kozlovskaya authored the section on gravity-seismic modelling. L. Taran and 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 invaluable material from Lithuania, and V. Starostenko did the same for the Ukraine. S. Bogdanova was the scientific leader of EUROBRIDGE, and R. Gorbatschev coordinated the INTAS effort. The authors thank D. Kurlovich from the Belarussian State University in Minsk for help with the preparation of the GIS-formated maps presented in this paper.
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SUNDBLAD, K. & CLAESSON, S. 2000. The Precambrian of Gotland. Geofizichesky Zhurnal, 22, 136. SUNDBLAD,K., MANSFELD, J., MOTUZA, G., AHL, M. & CLAESSON, S. C. 1994. Geology, geochemistry and age of a Cu-Mo-bearing granite at Kabeliai, Southern Lithuania. Mineralogy and Petrology, 50, 43-57. SUNDBLAD,K., GULLENCREUTZ,R. & FLOOgN, T. 1998. The Precambrian crust beneath the Baltic Sea. Geofizichesky Zhurnal, 20, 121-124. TARAN, L. N. & BOGDANOVA,S. V. 2001. The Fennoscandia-Sarmatia junction in Belarus: new inferences from a PT-study. Tectonophysics, 339, 193-214. TARAN, L. N. & BOGDANOVA,S. V. 2003. Metamorphism of the Palaeoproterozoic paragneisses of the Belarus-Podlasie Granulite Belt: a prograde-retrograde evolution. Petrology (Petrologija), 11, 444-461. THYBO, H., JANIK, T., OMELCHENKO, V. D., ET AL. 2003. Upper lithosphere seismic velocity structure across the Pripyat Trough and Ukrainian Shield along the EURUBRIDGE'97 profile. Tectonophysics, 371, 41-79. TSYMBAL, S. N. • TSYMBAL, Y. S. 2003. The upper mantle composition and diamond prospects in NW Ukrainian Shield. Mineralogichesky Zhurnal, 25, 40-56 [in Russian]. VALVERDE-VAQUERO, P., D(SRR, W., BELKA, Z., FRANKE, W., WISZNIEWSKA, J. & SCHASTOK, J. 2000. U - P b single-grain dating of detrital zircon in the Cambrian of central Poland: implications for Gondwana versus Baltica provenance studies. Earth and Planetary Science Letters, 184, 225-240. VERKHOGLIAD, V. M. 1995. Age stages of the Korosten pluton magmatism. Geokhimia i Metallogenia, 21, 34-47 [in Russian]. WOOLLARD, G. P. 1959. Crustal structure from gravity and seismic measurements. Journal of Geophysical Research, 64, 1521 - 1544. WYBRANIEC, S. 1999. Transformations and visualization of potential field data. Polish Geological Institute, Special Papers, 1. WYBRANIEC, S., ZHOU, S., THYBO, H., ETAL. 1998. New map compiled of Europe's gravity field. EOS Transactions, American Geophysical Union, 79, 437-442. YEGOROVA,T. P., STAROSTENKO,V. I., KOZLENKO,V. G. & YLINIEMI,J. 2004. Lithosphere structure of the Ukrainian Shield and Pripyat Trough in the region of EUROBRIDGE-97 (Ukraine and Belarus) from gravity modelling. Tectonophysics, 381, 29-59. YLINIEMI, J., TIIRA, T., LUOSTO, U., ET AL. 2001. EUROBRIDGE-95: deep seismic profiling within the East European Craton. Tectonophysics, 339, 153-176. Z1NCHENKO, 0. V., SKOBELEV,V. M., ESIPCHUK,K. E., SHEREMET,E. M. & VERKHOGLYAD, V. M. 1990. The Korosten complex. In: SHCHERBAKOV,I. B. (ed.) Petrology, Geochemistry and Metallogeny of Intrusive Granitoids of the Ukrainian Shield. Naukova Dumka, Kiev, 134-164 [in Russian].
The Archaean nucleus of the Fennoscandian (Baltic) Shield A. I. S L A B U N O V l, S. B. L O B A C H - Z H U C H E N K O 2, E. V. B I B I K O V A 3, P. S O R J O N E N - W A R D 4, V. V. B A L A G A N S K Y 5, O. I. V O L O D I C H E V 1, A. A. S H C H I P A N S K Y 6, S. A. SVETOV 1, V. P. C H E K U L A E V 2, N. A. A R E S T O V A 2 & V. S. S T E P A N O V 1
1Institute of Geology, Karelian Research Centre, RAS, Petrozavodsk, 185910, Russia 2Institute of Precambrian Geology and Geochronology, RAS, St. Petersburg, 199164, Russia 3Vernadsky Institute of Geochemistry & Analytical Chemistry, RAS, Moscow, 117975, Russia 4Geological Survey of Finland, Kuopio, 70211, Finland 5Geological Institute, Kola Science Centre, RAS, Apatity, 184209, Russia 6Geological Institute RAS, Moscow, 119017, Russia
Abstract: Archaean supracrustal complexes, known in the Fennoscandian (Baltic) Shield, are described and discussed by analysingthe time sections 3.1-2.9, 2.9-2.75 and 2.75-2.65 Ga. Data on granitoid complexes, interrelated in time and space, and evidence for Archaean metamorphic events are classified and presented briefly. Fragments of ophiolitic and eclogitic associations have been found in Archaean rocks in the Shield. The first evidence of continental crust in the Shield is from Meso-Archaean time (3.5-3.1 Ga); isolated microcontinents, such as Vodlozero, Iisalmi and North Finland, have been identified. New continental crust was mainly generated in the 2.9-2.65 Ga interval. The geodynamic settings in which the continental crust was formed in the Mesoand Neoarchaean included subduction (ensialic and ensimatic), accretion and collisional mechanisms. The continental and oceanic crust were affected by mantle plumes.
Archaean rocks form much of the eastern and northern parts of the Fennoscandian (Baltic) Shield, and can be divided into a number of discrete crustal provinces, each of which has a distinctive history of crustal formation and reworking. From SW to NE these are the Karelian, Belomorian, Kola and Murmansk Provinces, respectively (Fig. 1). Of these, the Murmansk Province has been little affected by younger events, whereas the Belomorian and Kola Provinces both record significant thermal and tectonic reworking and amalgamation related to the Palaeoproterozoic Lapland-Kola collisional orogeny (Daly et al. 2006). Further SW, the Karelian Province was blanketed by Palaeoproterozoic intracratonic basins; however, it shows significant thermal overprinting and tectonic reworking only along its southwestern margin, as a result of accretionary processes associated with the 1.9-1.8 Ga Svecofennian Orogeny. The Karelian and Murmansk Provinces form the cratonic nuclei of the Shield, and are therefore designated here as the Neoarchaean Karelian and Murmansk Cratons (Fig. 1). The Karelian Craton contains the typical range of granite-gneiss, greenstone, paragneiss and granulite complexes characteristic of the Archaean granite-greenstone association (Glebovitskii 2005; SorjonenWard & Luukkonen 2005). In contrast, the Murmansk Craton is composed dominantly of various granite-gneisses and granitoids, within which supracrustal rocks occur only as enclaves (Radchenko et al. 1994), usually metamorphosed to amphibolite grade; relics of granulite-facies mineral parageneses have also been described from the central part of the Craton (Petrov et al. 1990). The Belomorian Province is generally understood as a mobile belt, along the eastern and northeastern margin of the Karelian Craton (Fig. 1). Discriminating between Archaean and Palaeoproterozoic processes and events in the Belomorian mobile belt has been a source of controversy, but recent studies have clarified much of this and demonstrated that high-pressure (kyanite-facies) metamorphism (Volodichev 1990; Glebovitsky et al. 1996), with associated deformation, occurred in both in the Neoarchaean and the Palaeoproterozoic (Bibikova et al. 1996, 2001). Thus, although the Belomorian mobile belt contains rock units that are similar to those of the adjacent Karelian Craton in terms of age and composition, it has a distinctly different structural architecture, being composed of large-scale intensely folded nappe complexes
(Miller & Mil'kevich 1995). This generally recumbent structural development is also evident across the transition zone from the Belomorian mobile belt northeastwards into the Palaeoproterozoic Lapland-Kola orogen (Daly et al. 2001). The Kola Province consists of the Kola-Norwegian, Keivy and Sosnovska terranes and the Kolmozero-Voronya greenstone belt (Daly et al. 2006). Each of these terranes includes greenstone, paragneiss and granite-gneiss complexes, and the K o l a Norwegian terrane consists of a granulite-gneiss complex (e.g. Avakyan 1992). The entire Kola Province has been involved to a greater or lesser extent in the Palaeoproterozoic Lapland-Kola orogeny and it is separated from the Belomorian mobile belt by the Palaeoproterozoic (2.0-1.75 Ga) Lapland-Kola collisional suture (Daly et al. 2006) collisional sutures. The Inari and Tersky-Strel'na domains within the Kola suture zone comprise both Neoarchaean and Palaeoproterozoic rocks, the latter representing juvenile crustal protolith of the Lapland and Umba granulites (Glebovitsky et al. 2001; Daly et al. 2006). In the NE, the Kola Province borders against the Murmansk Craton. In summary, much of the Archaean of the Fennoscandian (Baltic) Shield is a typical granite-gneiss association (covering about 80% of the area), with various greenstone, paragneiss and granulitic complexes. However, the Belomorian mobile belt shows a distinct tectonic pattern, with large-scale thrusting and nappe complexes, and includes two tectonic and metamorphic rock associations that are rare in the Archaean: ophiolitic (the Central Belomorian greenstone belt and the Iringora complex) and eclogite-bearing complexes (Gridino zone and Salma area). This review commences with a description of the supracrustal greenstone and paragneiss complexes according to isotopically determined age groupings, followed by a presentation of intrusive and metamorphic rock associations.
Greenstone and paragneiss complexes Palaeoarchaean supracrustal rocks are represented by metakomatiites and spherulitic metabasalts in the Volotsk unit of the Vodlozero terrane in the southeastern part of the Karelian Craton (Kulikova 1993). Although they have not yet been studied in
From: GEE, D. G. & STEPHENSOY,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 627-644. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
627
628
A.I. SLABUNOV E T A L .
(a)
Archaean greenstone complexes (letters in square indicate major greenstone and schist belts: II = ilomantsi; KB = Khedozero-Bolsheozero, Ke = Keivy; Ks = Kostomuksha, KT = Kuhmo-Suomussalmi-Tipasj&rvi; KV = Kolmozero-Voron'ya; NK = North Karelian, OI = Olenegorsky, SK = Sumozero-Kenozero, SV = South Vygozero, T = Tulppio, V = Voche-Lambina; VS = Vedlozero-Segozero; Y = Yenb)
(b)
100Km White Sea
P IN , Murmansk Craton; LKO, Lapland-Kola Orogen; BP and KP, Belomorian and Kola Provinces, respectively; KC, Karetian Craton; SO, Svecofennian Orogen.
1
Archaean eclogitebearing complex (Gr = Gridino, Sa = Salma) Paragneiss complexes (2.7-2.78 Ga; N = Nurmes)
F~
CK
White Sea
~ , ~ --
"
2.75-2.68 Ga
o ~ m
2.9-2.85Ga
~
2.8-2.75 Ga (rhomb=lringora ophiolite complex)
E o ~ n
3.1.2.9 Ga
2.9-2.85 Ga ophiolite-like
complex of Central Belomorian Greenstone Belt 2.9-2.7 Ga granitoids, including Keivy alkaline granites, from Central Karelian (CK), Kola-Norwegian (KN), Keivy (Ke), Sosnovka (So) and North Sweden terranes
33 ~ E Lapland granulite belt (Lp; mainly 2.0-1.9 Ga), including Tanaelv melange
Caledonides Phanerozoic and Neoproterozoic platform cover []~-[
Neo- and Mesoproterozoic rocks Rapakivi granites
(1.65-1.54 Ga)
Granitoids (1.85-1.75 Ga) Supracrustal rocks (2.06-1.85 Ga)
r-q
Supracrustal rocks (2.5-2.06 Ga) Tectonic mixture of Neoarchaean and Palaeoproterozoic rocks (In and TS, Inari and Tersk-Strel'na terranes, respectively)
Sanukitoids and their analogues (2.74-2.72 Ga; Tv = Tavaj&rvi massif) Granulitic complexes (2.74-2.72 Ga; Vp = i~i!~i!i!i~!~i!;~i:~Varpaisj&rvi, Vk = Voknavolok, TI = Tulos, On = Onega, Nt = Notozero)
m
2.9-2.7 Ga granitoids from Belomodan (BMB) and Kianta (Ki) terranes 3.1-2.7 Ga granitoids from Ilomantsi-Voknavolok (IV) and Vodlozero (Vo) terranes 3.5-2.7 Ga granitoids from terranes: lisalmi (li), Pomokaira (P), Ranua (R), Vodlozero (Vo; core)
illmill ~
~
"~,',~b % ",, c
Major tectonic boundary Faults: thrust (a), normal (b) and strike-slip (c)
Fig. 1. (a) Tectonic units of the eastern Fennoscandian (Baltic) Shield. (b) Schematic representation of the major geological units and structures in the eastern Shield (based on the authors' data, and Kostinen et al. 2001; Sorjonen-Ward & Luukkonen 2005).
ARCHAEAN NUCLEUS, FENNOSCANDIANSHIELD detail, they have a S m - N d isochron whole-rock age of 3391 + 16 Ma (Puchtel et al. 1991). Some of the high-grade gneises and amphibolites I of the Vodlozero gneiss complex were perhaps derived from 3.3-3.55 Ga calc-alkaline volcanic rocks (Sergeyev et al. 1990; Lobach-Zhuchenko et al. 1993). The Meso- and Neoarchaean greenstone complexes of the Shield belong to at least four generations with ages of 3.1-2.9, 2.9-2.85, 2.85-2.75 and 2.75-2.65Ga, whereas paragneiss and schist complexes have been dated at 2.9-2.85 and 2.752.65 Ga. Each of these age groupings is considered separately below.
629
belt of the Kola Province. Contacts with younger complexes are typically tectonic. The Vedlozero-Segozero greenstone belt is located on the western margin of the Vodlozero terrane in the Karelian Craton (Fig. 1). It comprises two separate complexes with different ages: 3.05-2.94 Ga and c. 2.85 Ga (Fig. 2). The older of these includes two distinct associations formed between 3.05 and 2.94 Ga, as follows. ( 1) A basaltic - andesitic- dacitic association, which is most complete in the northern part of the Hautavaara structure (the Chalka zone), has a total thickness of 2.5 km and includes pillowed, amygdaloidal and fragmental lava flows, various volcaniclastic vents and dykes, metamorphosed to amphibolite and epidote-amphibolite grade. Ages of 2995 _+ 20 Ma have been obtained from subvolcanic rocks of andesitic composition (Sergeyev 1989) and 2945 _ 19 Ma for andesitic lava (Ovchinnikova et al. 1994). These rocks belong to a normally differentiated calc-alkaline series. The more primitive volcanic rocks are rich in Cr and Ni, whereas later differentiates are enriched in Co, Zr and Y. Their Sr/Y ratio < 12, Ce/Nb ratio 20) and have no Eu anomalies. They are also commonly associated with lamprophyre dykes. (3) Subalkaline rocks occur as typically post-tectonic syenitic massifs, as at Khizhjarvi in the Karelian Craton, and vary in composition from monzodiorite to leucosyenite. They resemble sanukitoids, but are richer in alkalis and poorer in Mg.
In the Kola Province, in the southeastern part of the Kola-Norwegian terrane, this group seems to include quartz monzonites, quartz syenites and latites with an age of 2657 _ 9 Ma (Balashov et al. 1992) and aNd(t) = --0.8 (Timmerman & Daly 1995). The Porosozero polyphase monzodioritic-granitic massif, dated at 2733 + 6 Ma (Kudryashov et al. 2001), intrudes supracrustal rocks of the Kolmozero-Voron'ya belt. A very distinctive suite of granitoids in the Kola Province is represented by the alkaline granites of the Keivy terrane, which typically consist of aegirine- arfvedsonite-bearing and lepidomelanehastingsite-bearing granites (Radchenko et al. 1994). These rocks form sheet-like bodies and dykes that have been dated by the U - P b zircon method at 2630 _+ 31 and 2654 _+ 5 Ma (Belye Tundry massif, Bayanova et al. 1999; Mitrofanov et al. 2000), and at 2751 • 41 Ma (Ponoy massif, Vetrin et al. 1999a). Syenitic granites from the West Keivy massif have an age of 2674 _ 6 Ma (Mitrofanov et al. 2000). Ponoy alkaline granites are rich in Fe, poor in P and Sr, very poor in Cu, Ni, V, Cr and Co, and rich in Li (up to 1000 ppm), Zr, Nb, Y, U, Th and REE (Vetrin et al. 1999a). Their aNd(t) values are 0.1-2.9 (~;~ = 2.64-2.91 Ga), although one sample has aNd(t) = --5.9 ( t ~ = 3.62 Ga). These granites are defined as anorogenic. A final notable feature of Neoarchaean evolution in the Shield was the emplacement of the Siilinj/irvi carbonatite into the Iisalmi terrane, near the western margin of the Karelian Craton, at 2.58-2.61 Ga (Patchett et al. 1981). (4) Two-feldspar granites are ubiquitous in this age group, occurring as post-tectonic plutons and gently dipping sheets dated at 2680-2710 Ma. They have both I- and A-type affinities and aNd(t) values depend on the age of the terrane, ranging from - 0 . 4 to - 4 . 9 for plutons in the old Vodlozero terrane to 0.1 to -1.2 in the Voknavolok-Ilomantsi terrane; in the relatively young Central Karelian terrane the values are as high as +0.8 to +2.2 (Lobach-Zhuchenko et al. 2000b). In the Kola Province similar types of granitoids are represented by the monazite-bearing granites (2634 ___ 12 Ma) in the southeastern Kola-Norwegian domain (Balashov et al. 1992).
Characteristics of the Archaean metamorphism of supracrustal sequences, granulitic and eclogite-bearing complexes A r c h a e a n m e t a m o r p h i s m in the K a r e l i a n Craton
The oldest rocks of the Karelian Craton have been repeatedly and variously metamorphosed, including under high-temperature amphibolite-granulite-facies conditions. For example, 3.2-3.1 Ga granulite metamorphism has been recognized in the 3.54 Ga rocks of the Vodlozero complex, which were subsequently metamorphosed at amphibolite (2.86 Ga) and epidote-amphibolite (2.7 Ga) grade (Glebovitsky 2005, and references therein). In addition, a Neoarchaean (2.65 Ga) granulite metamorphic event was identified within this terrane (Glebovitsky 2005, and references therein). The main stages of metamorphism of rocks in the Karelian Craton coincided with the main stages of formation of greenstone, paragneiss and granitiod complexes. Supracrustal and granitegneiss complexes are usually polymetamorphic. Their P - T - t evolutionary trends typically include the following two stages: (1) early (3.0-2.75 Ga) low-pressure metamorphism; (2) late (2.72-2.65Ga) high-pressure dynamothermal metamorphism (Fig. 4a). The earlier stage is recorded throughout all greenstone belts, regardless of age, and is manifest by greenschist- to lowpressure amphibolite-facies assemblages with either andalusite or sillimanite. The later stage is more typically developed in discrete domains, particularly in zones of transpression. The metamorphic grade tends to be higher-pressure amphibolite facies with kyanite-sillimanite-bearing assemblages. However, even
ARCHAEAN
NUCLEUS,
(a)
FENNOSCANDIAN
SHIELD
637
(b) 16 P. Kbar
P. Kbar
14 12
E
~
//
~""
.2[/" /-
10 .7 Ga
~
//"
S~171-
2.73 Ga
4
2(J0
400
600
800
T, ~
2(t(/
J 4~t
4102~3~4
I 600
....
J 8o0
~1, T, ~
/?5,/6
P-T diagram for the Archaean metamorphic events in the rocks of (a) the Karelian Craton and (b) the Belomorian mobile belt (Volodichev 1990; Lobach-Zhuchenko et al. 1995). Numbers with 'Ga' show the age of the metamorphic processes. In (a), the colours indicate the evolution of different terranes of the Craton. In (b): 1, earliest granulitic metamorphism; 2, Neoarchaean granulitic metamorphism; 3, Neoarchaean eclogitic metamorphism; 4, Neoarchaean collisional metamorphism; 5, evolution of metamorphism in the western part of the Belomorian mobile belt; 6, evolution of the Neoarchaean metamorphism in the eastern part of the Belomorian mobile belt. Fields of metamorphic facies: Gr, greenstone; A, amphibolitic; G, granulitic; E, eclogitic.
F i g . 4.
within single shear zone, pressures may vary greatly, from 5 - 7 to 10 kbar. At least two stages of transpressive tectonics and metamorphism can be distinguished in the Karelian Craton. Metamorphism occurred simultaneously with development of an early regional system of shear zones (transpressional and transtensional), with granitoid (including sanukitoid) intrusions formed at depth in the crust and pull-apart basins near the surface. The evolution of second-generation shear zones was accompanied by metamorphism and the formation of subalkaline granitoids (Volodichev et al. 2002). These Neoarchaean metamorphic and tectonic processes in the Karelian Craton were accompanied by hydrothermal and metasomatic alteration processes and gold mineralization (Nurmi & Sorjonen-Ward 1993), and are presumably a consequence of the coeval Neoarchaean (2.72-2.65 Ga) collisional tectonics in the Belomorian mobile belt. Several granulite (or granulite-enderbite-charnockite) complexes are known in the Karelian Craton (VarpaisjSxvi, Tulosozero, Voknavolok and Onega complexes), in the Belomorian mobile belt (Notozero complex) and in the Kola Province (Fig. 1). They all have some features in common, in that they are dominated by enderbites of dioritic to tonalitic composition; supracrustal rocks are less abundant but include mafic and intermediate granulites, and aluminous garnet-cordierite-sillimanite gneisses. Charnockites do not occur in all granulitic complexes. The Varpaisjarvi complex, which has been studied most thoroughly, is used as an example to discuss their structure and evolution. The Varpaisj~irvi complex occurs among migmatized TTG granitoids of the Iisalmi terrane, dated at c. 3.1 Ga (Paavola 1986) and includes enderbites (with sporadic anorthositic layers) and mafic, intermediate and aluminous granulites. Dioritic to tonalitic enderbites predominate and are transitional to surrounding granitoids, or locally intrude them. Magmatic crystallization of the enderbites took place at 2.72-2.70 Ga ago; they were then metamorphosed at P = 9-11 kbar and T = 800-900 ~ probably at around 2.63 Ga, followed by decompression and cooling at 7 kbar and 700 ~ (H61tt~i et al. 2000; M/intt~iri & H61tt~i 2002). The Proterozoic Svecofennian thermal and tectonic overprint on the VarpaisjSxvi complex is probably relatively minor, whereas the adjacent Rautavaara supracrustal gneiss complex may record granulite- to amphibolite-facies exhumation and re-equilibration during Proterozoic time.
Two-pyroxene and hornblende granulites (often with garnet) occur as small lenticular bodies within the TTG rocks and enderbites. No primary protolith textures are preserved, but geochemically they can be classified into two distinct groups of basaltic and andesitic composition: (1) tholeiitic-series basalts with a chondritic Ti/Zr ratio (c. 110), flat REE patterns (sometimes poor in LREE), with tDM Na = 3.1--3.2 Ga; (2) basalts and andesites with Ti/Zr ratios < 100, enriched in LREE, and with Nd model ages of 2.7-2.9 Ga. It should be noted that the latter group is more common in the southeastern part of the complex (Hrltt~i & Paavola 2000; Hrltt~i et al. 2000). The U - P b zircon ages of the protoliths of granulite from group 1 are 3.05-3.2 Ga, whereas those from group 2 are 2.65-2.68 Ga. The enderbites and granulites of both groups were metamorphosed to granulite grade at 2.62-2.70 Ga (M~intt~iri & Htltt~i 2002). Aluminous granulites are related spatially to the Group 2 mafic granulites and have similar Nd model ages (2.7-2.8 Ga). Zircon cores usually give ages of 2.70-2.81 Ga, whereas newly-formed metamorphic rims have ages of 2.60-2.68 Ga (MS_ntt~iri & H61tt~i 2002). The S m - N d isochron ages of garnet from aluminous and marie granulites, corresponding to cooling to T ~, 600 ~ are 2.59 + 0.01 and 2.52 + 0.05 Ga, respectively (M~intt~iri & Hgltt~i 2002). The Varpaisj~irvi complex is assumed to have been formed upon accretion of two terranes differing in composition and age (M~intt~iri & H61tt~i 2002).
Neoarchaean metamorphism of Belomorian mobile belt rocks The most distinctive feature of the Belomorian mobile belt is the two superimposed Neoarchaean and Palaeoproterozoic high-pressure (including eclogite-facies) metamorphic events (Volodichev 1990; Glebovitsky et al. 1996). The Neoarchaean metamorphism occurred in different ways in the eastern and western domains of the Belomorian mobile belt (Fig. 4b). The eastern domain records a clockwise P - T - t evolutionary trend, which is particularly well expressed in rocks from the Gridino zone (Fig. 1) and the Pongoma Bay area, where prograde metamorphism culminated at eclogite-facies conditions ( P = 1 4 - 1 7 . 5 k b a r , T = 7 4 0 - 8 6 5 ~ followed by
638
A.I. SLABUNOVETAL.
multistage retrogression. This included near-isothermal decompression from 14.0kbar to 6.5 kbar and cooling from 770 to 650 ~ (Fig. 4b). Thus, the conditions correspond to high-pressure granulite-facies conditions for mafic garnet granulite assemblages, transitional to amphibolite-facies conditions (Cloos 1993). The granulite-facies conditions (P = 6 - 7 kbar, T = 750-800 ~ associated with migmatitic and intrusive enderbites, and moderate-pressure amphibolite facies were maintained at 2.72 Ga (Levchenkov et al. 1996) (Fig. 4b). At 2691 + 5 M a (Levchenkov et al. 2001) a return to higher-pressure conditions (P = 9-11 kbar, T = 650-700 ~ occurred. This metamorphic event is assumed to have been caused by transpressive tectonics during the Belomorian collisional orogeny. In contrast, the western domain (e.g. in the Lake NotozeroLake Kovdozero area) is characterized by an 'anticlockwise' P - T - t trend (Fig. 4b). The earliest event, dated at 2820 _+ 5 Ma (Bibikova et al. 2004), was moderate-pressure granulite-facies metamorphism (P = 5.5-6.5 kbar, T > 700 ~ reported from the Chupa paragneisses and from mafic granulites (LobachZhuchenko et al. 1993). A granulite complex, consisting of mafic granulites, enderbites and charnockites of calc-alkaline and Fe-tholeiitic series (Volodichev 1990), was formed later, at c. 2.72 Ga (Glebovitsky 2005). High-pressure metamorphism, up to 9-11 kbar at 650-700 ~ was accompanied by intense migmatization and occurred somewhat later, around 2.7 Ga (Volodichev 1990; Glebovitsky et al. 1996, 2000). The latter episode is common to both western and eastern domains and reflects a response to collisional processes (Searle & Rex 1989). A similar P - T evolution characterizes the northern part of the Belomorian mobile belt, where the first stage of the Neoarchaean metamorphism was at 620-680 ~ and 7.6-8.8 kbar whereas the later event occurred at 665-695 ~ and 9.7-10.6 kbar (Belyaev & Petrov 2002). An eclogite-bearing complex is present in the Gridino zone (Volodichev et al. 2004) and Salma area (Shchipansky et al. 2005) in the eastern and northern parts of the Belomorian mobile belt (Fig. 1). It can be followed as a zone of intensely migmatized mrlange from NW to SE for about 50 km and has a width of up to 10 km (Fig. 5a). The granitoid components within the eclogite complex have been transformed to granite gneisses by multiple deformational and metamorphic events and are of tonalitic-trondhjemitic composition. Relict enderbites, indicative of granulite-facies metamorphism, are also present. The mrlange comprises a chaotic mixture of diverse lithologies (Fig. 5b), including eclogites, gamet, garnet-clinopyroxene and feldspathic amphibolites, meta-ultramafic rocks, metagabbroids and zoisite rocks (meta-anorthosites), aluminous and amphibole-bearing gneisses, and marbles. They are very heterogeneous in terms of composition, the primary nature of the protolith and the metamorphic history. Archaean eclogites, consisting of omphacite (28-33 to 40% jadeite) and garnet (22-30% pyrope and 22-30% grossular; Fig. 5c), are preserved as relics among symplectitic eclogites and garnet-clinopyroxene amphibolites derived from the ecologites during retrograde decompression. The eclogites were formed at pressures of 14.0-17.5 kbar and temperatures of 740865 ~ corresponding to depths to 60-65 km. The U - P b age of zircons from the eclogites and symplectite eclogites is 2720.7-+- 5.8 Ma (Bibikova et al. 2003b; Volodichev et al. 2004). The morphology of the analysed zircon is characteristic of high-pressure granulites and eclogites (Bibikova 1989). The mrlange zone is cut by post-tectonic trondhjemite veins dated at 2701.3 • 8.1 Ma (Bibikova et al. 2003b) and gabbro-norite dykes (2.4 Ga). The eclogites correspond petrogeochemically to tholeiitic-series mafic rocks (47-51% SiO2, 1.38-4.3% Na20 + K20) (FeO*/ MgO = 0.5-2.5). REE abundances are 2 - 1 2 times that of chondrites, and they typically show flat or poorly fractionated REE patterns (La/Smy = 0.99-1.8; Ga/YbN = 0.77-1.17). Compositions
correlate with the metabasalts (amphibolites) from the ophiolite complexes, including from the Central Belomorian greenstone belt (Slabunov 2005). The recognition of Archaean crustal eclogites in the Belomorian mobile belt clearly provides a strong argument in favour of possible deep subduction during the Neoarchaean, to explain such high pressures under a relatively low geotherm (see Yardley 1989).
M e t a m o r p h i s m in the K o l a P r o v i n c e a n d in the M u r m a n s k Craton
The available data (Avakyan 1992; Belyaev & Petrov 2000) show that in the granulitic zone of the Kola-Norwegian terrane (the Kola Province) maximum pressure estimates for cordierite-free parageneses from sillimanite gneisses are 6.2 _+ 1.2kbar at T = 700 ~ whereas in the zone of transition to amphibolite facies, peak pressures were 5.2 • 0.9 kbar. Cordierite-bearing paragneses show lower pressures: 4.5 _+ 0.6 kbar at 700 ~ For rocks with kyanite-sillimanite assemblages in the transition zone, P = 5.3 kbar and T = 580 + 20 ~ Multiple (retrograde) metamorphism in the granulit,~-facies zone took place under amphibolite-facies conditions at P : 3.5 + 0.5 kbar and T = 590 ~ R b - S r isotopic data (Avakyan 1992) suggest that the first thermal event within the Kola granulite-facies paragneisses, prior to the granulite metamorphism, occurred at 2880 + 45 Ma (Isr : 0.7005 _+ 0.0004). The age of the earliest granulite metamorphism is estimated at 2.83 Ga (Bibikova 1989; Balashov et al. 1992), whereas the age of metamorphic zircon from later shear zones, also formed under granulite-facies conditions, is 2648 ___ 18 Ma (Balashov et al. 1992). Metamorphic zircons from the granulite-facies paragneisses, sampled near Lake Pulozero, suggest that yet another metamorphic event occurred at 2724 +_ 49 Ma, followed by retrogression under amphibolitefacies conditions at 2640 + 20 Ma (Pozhilenko et al. 2002). The last recorded Archaean tectonometamorphic event is related to emplacement of discordant pegmatite bodies at 2556 ___27 Ma (Balashov et al. 1992). At least three metamorphic events can be distinguished (Belyaev & Petrov 2000) in the Kolmozero-Voronya greenstone belt. Early metamorphism (2.83-2.76Ga) took place under epidoteamphibolite- and amphibolite-facies conditions at low temperature and low pressure (T = 460-560 ~ P = 2.5-4.3 kbar). A second event (2.76-2.68 Ga) occurred at the same temperatures, but at higher pressures (T = 470-530 ~ P = 3.9-5.8 kbar). During the final stages of Archaean evolution (2.68-2.52Ga), even greater temperatures and pressures were attained in late shear zones (T = 530-640 ~ and P = 6.0-8.5 kbar). In rocks of the Murmansk Craton, the earliest metamorphic events occurred under granulite-facies conditions, at temperatures up to 750 ~ and pressures of 4 - 6 kbar (Petrov et al. 1990). Subsequent metamorphic reworking took place under amphibolitefacies conditions and was accompanied by migmatization.
Summary of the main stages of crust formation The earliest 'sialic crustal nuclei' were the 3.5-3.1 Ga Vodlozero terrane in the southeastern Karelian Craton and two small blocks in central and northern Finland, at the western margin of the Karelian Craton (Figs 1 and 6). In the period 3.1-2.95 Ga, new crust accreted around the Vodlozero terrane and, to a lesser extent, in the northwestern part of the Karelian Craton. This stage is reflected most completely in the Vedlozero-Segozero greenstone belt of the Vodlozero terrane (Fig. 6), which provides the oldest evidence of orogeny in the Shield (Fig. 6). An ensialic (mature) volcanic arc and a
ARCHAEAN NUCLEUS, FENNOSCANDIAN
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Fig. 5. (a) Location of the Gridino eclogite-bearing complex. (b) Schematic geological map of the southeastern part of Stolbikha Island (by A. I. Slabunov & O. S. Sibelev, in co-operation with O. I. Volodichev). (c) Photomicrograph showing a thin section of eclogite with homogeneous non-zoned garnet crystals (Grt; Prp26 = pyrope content of Grt) and omphacite content (Omp32; number shows jadeite content). Secondary alteration is represented by formation of plagioclase (P122; number shows anorthite content), diopside (Di16; number shows jadeite content) and pargasitic hornblende (Prg-Hbl).
deep-water back-arc basin evolved there between 3.05 and 3.0 Ga (Svetov 2005). Mantle plumes gave rise to oceanic-plateau type basalts and komatiites, which were obducted later onto the continent (the Sumozero-Kenozero greenstone belt), and a subcontinental-plateau basalt type in the South Vygozero greenstone belt (Arestova et al. 2003). The first record of crustal growth in the Voknavolok-Ilomantsi terrane also dates from this time ( ~ not more than 3.1 Ga). In the period 2.95-2.85 Ga, the continental crust grew mainly in the southeastern part of the present Karelian Craton, adjacent to the Belomorian belt and in the Kola Province. In the western Vodlozero terrane, subduction produced a volcanic arc with calc-alkaline series volcanic rocks (the Vedlozero-Segozero
greenstone belt) The formation of the oldest two-feldspar granites records the final stages of evolution of this accretionary event. However, the major sialic crustal formation event related to subduction took place somewhat later (2.88-2.82 Ga), as recorded in the Sumozero-Kenozero greenstone belt at the margin of the Vodlozero terrane and in the North Karelian greenstone belt in the Belomorian belt. There are fragments of c. 2.88 Ga oceanic crust in the Belomorian mobile belt. Continental crust was formed simultaneously in the central part of the Karelian Craton, in the Murmansk Craton and Kola Province (~dM is commonly < 2 . 9 Ga). In the period 2.82-2.75 Ga, a new island-arc system was formed and began to evolve along the northeastern boundary of
640
A.I. SLABUNOV ET AL .
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Fig. 6. Correlation of Archaean greenstone, paragneiss, magmatic complexes and metamorphic events in the eastern Fennoscandian (Baltic) Shield (terranes: CK, Central Karelian; Ii, Iisalmi; IV, Ilomantsi-Voknavolok; Ke, Keivy; Ki, Kianta; KN, Kola-Norwegian; P, Pomokaira; R, Ranua; Vo (VS) Vodlozero (Vedlozero-Segozero greenstone belt); Vo (Core), core of the Vodlozero.
the Karelian Craton with the Belomorian mobile belt. It has survived as fragments of various volcanic rocks (including boninites and adakites) and supra-subduction ophiolites in some of the greenstone belts in both the Karelian Craton and the Belomorian mobile belt. In the western part of the Karelian Craton, supracrustal complexes were generated (the Kuhmo-Suomussalmi-Tipasj/irvi and Kostomuksha belts). Some workers (Luukkonen 1988; Lobach-Zhuchenko et al. 2000a) have interpreted them as riftogenic, whereas others (Piirainen 1988; Puchtel et al. 1998; Kozhevnikov 2000; Samsonov et al. 2001) have considered them in terms of accretionary-collisional and mantle plume tectonic processes. Much of the present Karelian Craton had thus been formed by 2.75 Ga, when this new system accreted to the older amalgamation. Between 2.75 and 2.65 Ga, the growth of the crust continued in the Belomorian mobile belt and culminated in the collision of microcontinents, which gave rise to an orogen with thick continental crust (Glebovitsky 2005; Slabunov 2005). During the final stage of subduction and the start of collision (c. 2.72 Ga) eclogites of the Belomorian mobile belt were exhumed. The intrusion of the 2.73-2.70 Ga Notozero intrusive complexes and their analogues, which correspond geochemically to active continental margin rocks, occurred in the zone between the Karelian Craton and the Belomorian mobile belt, and in the central Karelian Craton, pullapart basins evolved. They were filled with sediments, and felsic and intermediate volcanic rocks, and are now represented by the Ilomantsi, Khedozero-Bolshozero and Gimoly belts. At 2.742 . 7 0 G a subalkaline and sanukitoid granitoids were intruded during closure of these basins. The Kola Province became a collage of exotic terranes (such as the Kola-Norwegian and Keivy terranes) by the end of the Neoarchaean.
In summary, the formation of the Archaean sialic crust of the Fennoscandian (Baltic) Shield can be understood in terms of the subduction and collision of lithospheric plates (Gafil & Gorbatschev 1987; Mints 1998), with an additional influence from mantle plumes (Lobach-Zhuchenko et al. 1999; Puchtel et al. 1998; Arestova et al. 2003). The architecture of the Archaean continental crust can be attributed to accretionary-collisional processes, with at least four major phases of accretion and one collisional event being recognized (Fig. 6). The SVEKALAPKO Project of the EUROPROBE programme initiated this work. This paper is a contribution to programmes O N Z - 6 "Geodynamics and mechanisms of lithosphere deformation" and O N Z - 8 "Isotope systems and isotope fractionation in natural processes". We acknowledge financial support from the Russian Foundation for Basic Research (RFBR) (grants 00-05-64295, 00-05-64701, 03-05-64010, 03-05-64501 and 06-05-64876).
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KULIKOVA, V. V. 1993. Volotsk suite: a Lower Archaean stratotype of the Baltic Shield. Karelian Research Center, Petrozavodsk [in Russian]. LEVCHENKOV,O. A., LEVSKY,L. K., NORDGULEN,0., eraL. 1995. U - P b zircon ages from Scrvaranger, Norway, and the western part of the Kola Peninsula, Russia. In: ROBERTS, D. & NORDGULEN, O. (eds) Geology of the Eastern Finnmark-western Kola Peninsula Region. Norges Geologiske UndersCkelse, Special Publications, 7, 29-47. LEVCHENKOV, 0. A., VOLODICHEV, 0. I., ZINGER, T. F., YAKOVLEVA, S. Z., MAKEEV, A. F., SHULESHKO, I. K. & KEVLICH, V. I. 2001. Age of metamorphism of aluminous gneisses from the Pon'goma Inlet Region of the Belomorian Mobile Belt, Baltic Shield. Doklady Earth Sciences, Moscow, 377(1-2), 168-171. LEVCHENKOV, O. A., ZINGER, T. F., DUK, V. L., YAKOVLEVA,S. Z., BAIKOVA, V. S., SHULESHKO, I. K. & MATUKOV, D. I. 1996. U Pb zircon age of the hypersthene diorites and granodiorites of Pon'gom-Navolok Island (Baltic Shield, White Sea tectonic zone). Earth Science Sections, Moscow, 349(1-2), 852- 854. LOBACH-ZHUCHENKO, S. B., CHEKULAEV, V. P., SERGEYEV, S. A., LEVCHENKOV, O. A. & KRYLOV, I. N. 1993. Archaean rocks from southeastern Karelia (Karelian granite-greenstone terrane). Precambrian Research, 62, 375-388. LOBACH-ZUCHENKO, S. B., BIBIKOVA, E. V., DRUGOVA. G. M., ET AL. 1995. Archean magmatic rocks in the lake Notozero area of the Northwestern Belomorian: isotopic geochronology and petrology. Petrology, Moscow, 3(6), 593-621. LOBACH-ZHUCHENKO, S. B., CHEKULAEV, V. P., STEPANOV, V. S., SLABUNOV, A. I. & ARESTOVA, N. A. 1998. The White Sea foldbelt--Late Archaean accretion- and collision-related zone of the Baltic Shield. Doklady Earth Sciences, Moscow, 358(1-4), 34-37. LOBACH-ZHUCHENKO,S. B., ARESTOVA,N. A., CHEKULAEV,V. P., ETAL. 1999. Evolution of the South Vygozero greenstone belt, Karelia. Petrology, Moscow, 7(2), 160-176. LOBACH-ZHUCHENKO, S. B., ARESTOVA, N. A., MIL'KEVICH, R. I., LEVCHENKOV, O. A. & SERGEYEV, S. A. 2000a. Stratigraphy of the Kostomuksha Belt in Karelia (Upper Archaean) as inferred from geochronological, geochemical, and isotopic data. Stratigraphy and Geological Correlation, Moscow, 8(4), 319-326. LOBACH-ZHUCHENKO, S. B., CHEKULAEV, V. P., ARESTOVA, N. A., LEVSKII, L. K. & KOVALENKO,A. V. 2000b. Archaean terranes in Karelia: geological and isotopic-geochemical evidence. Geotectonics, Moscow, 34(6), 452-466. LOBACH-ZHUCHENKO, S. B., ROLLINSON, H. R., CHEKULAEV,V. P., ET AL. 2005. The Archaean sanukitoid series of the Baltic Shield: geological setting, geochemical characteristics and implications for their origin. Lithos, 79, 107-128. LUUKKONEN,E. 1988. The structure and stratigraphy of the late Archaean Kuhmo greenstone belt, Eastern Finland. In: MARTTILA,E. (ed.) Archaean Geology of the Fennoscandian Shield. Geological Survey of Finland, Special Papers, 4, 71-96. LUUKKONEN, E. 1992. Late Archaean and Early Proterozoic structural evolution in the Kuhmo-Suomussalmi Terrane, Eastern Finland. Annales Universitatis Turkuensis Series, 78, 37. MANTT~.RI,I. & HOLTTA,P. 2002. U - P b dating of zircons and monazites from Archaean granulites in Varpaisj~irvi, Central Finland. Precambrian Research, 118, 101 - 131. MIL'KEVlCH, R. I. & MYSKOVA,T. A. 1998. Late Archean metaterrigenous rocks of the Western Karelia: lithology, geochemistry and provenances. Lithology and Mineral Resources, 33(2), 155-171. MILLER, Yu. V. & MIL'KEVICH, R. I. 1995. Folded nappe structure of the Belomorsk Zone and its relation to the Karelian granite-greenstone belts. Geotectonika, Moscow, 6, 80-93 [in Russian]. MILLER, Y. V., MYSKOVA,T. A. & MIL'KEVICH, R. I. 2002. Supracrustal rocks in the tectonic windows of the marginal part of the Karelian Craton (Northwestern Belomorian Province). Geotectonics, Moscow, 36(1), 11-23. MINTS, M. V. 1998. Archaean miniplate tectonics. Geotectonics, Moscow, 32(6), 427-443. MINTS, M. V., GLAZNEV, V. N., KONILOV, A. N., ETAL. 1996. The early Precambrian of the northeastern Baltic Shield: paleogeodynamics, crustal structure and evolution. Nauchny Mir, Moscow [in Russian].
ARCHAEAN NUCLEUS, FENNOSCANDIAN SHIELD
MITROFANOV, F. P. & POZHILENKO, V. I. (eds) 1991. Archaean VocheLambina geodynamic experimental area, Kola Peninsula. Academy of Sciences of the USSR, Kola Branch, Apatity [in Russian]. MITROFANOV, F. P., ZOZULYA, D. R., BAYANOVA, T. B. & LEVKOVICH, N. V. 2000. The world's oldest anorogenic alkali granitic magmatism in the Keivy structure on the Baltic Shield. Doklady Earth Sciences, Moscow, 374(7), 1145-1148. MUTANEN, T. & HUHMA, H. 2003. The 3.5 Ga Siurua trondhjemite gneiss in the Archaean Pudasjarvi Granulite Belt, northern Finland. Bulletin of the Geological Society of Finland, Espoo, 75(1-2), 51-68. MYSKOVA, T. A., GLEBOVITSKY, V. A., MILLER, Yu. V., L'vov, A. B., KOTOV, A. B., KOVACH, V. P. & ZAGORNAYA, N. Y. 2003. Supracrustal sequences of the Belomorian mobile belt: primary composition, age and genesis. Lithology and Mineral Resources, Moscow, 11(6), 3-19 [in Russian]. NIEMINEN, J. 1998. The polymictic volcaniclastic conglomerates at Kelloj/irvi. MSc thesis, University of Turku [in Finnish]. NURMI, P. & SORJONEN-WARD, P. (eds) 1993. Geological development, gold mineralization and exploration methods in the late Archean Hattu schist belt, Ilomantsi, Eastern Finland. Geological Survey of Finland, Special Papers, 17. OVCHINNIKOVA, G. V., MATRENICHEV, V. A., LEVCI-IENKOV, O. A., SERGEYEV, C. A., YAKOVLEVA, S. Z. & GOROKHOVSKY, B. M. 1994. U - P b and Pb-Pb isotopic studies of felsic volcanics from the Hautavaara greenstone structure, Central Karelia. Petrology, Moscow, 2(3), 266-281 [in Russian]. PAAVOLA, J. 1986. A communication of the U - P b and K - A r age relation of the Lapinlahti-Varpaisj~irvi area, central Finland. Geological Survey of Finland Bulletin, 339, 7-15. PATCHETT, P. J., Kouvo, O., HEDGE, C. E. & TATSUMOTO, M. 1981. Evolution of continental crust and mantle heterogeneity: evidence from Hf isotopes. Contributions to Mineralogy and Petrology, 78, 279-297. PETROV, V. P., BELYAEV, 0. A., VOLOSHINA,Z. M., BALAGANSKY,V. V., GLAZUNKOV, A. N. & POZHILENKO, V. I. 1990. Endogenic regimes of metamorphism in the Early Precambrian. Nauka, Leningrad [in Russian]. PIIRAINEN, T. 1988. The geology of the Archaean greenstone-granitoid terrane in Kuhmo, eastern Finland. In: MARTTILA, E. (ed.) Archaean geology of the Fennoscandian Shield. Geological Survey of Finland, Special Papers, 4, 39-51. POZHILENKO, V. I., GAVRILENKO,B. V., ZHIROV, D. V. & ZHABIN, S. V. 2002. Geology of mineral areas of the Murmansk region, Kola Science Centre, Apatity [in Russian]. PUCHTEL, I. S., ZHURAVLEV,D. Z., KULIKOVA,V. V., SAMSONOV,A. V. & SIMON A. K. 1991. Komatiites from the Vodla Block, Baltic Shield: a window to the early-Archean mantle? Doklady Akademii Nauk SSSR, 317(1), 197-202 [in Russian]. PUCHTEL, I. S., HOFMANN, A. W., MEZGER, K., JOCHUM, K. P., SHCHIPANSKY, A. A., & SAMSONOV, A. V. 1998. Oceanic plateau model for continental crustal growth in the Archaean: a case study from the Kostomuksha greenstone belt, NW Baltic Shield. Earth and Planetary Science Letters, 155, 57-74. PUCHTEL, I. S., HOFMANN, A. W., AMELIN, YU, V., GARBE-SCHONBERG, C. D., SAMSONOV, A. V. & SHCmPANSKY, A. A. 1999. Combined mantle plume-island arc model for the formation of the 2.9 Ga Sumozero-Kenozero greenstone belt, SE Baltic Shield: isotope and trace element constraints. Geochimica et Cosmochimica Acta, 63(21), 3579-3595. RADCHENKO, A. T., BALAGANSKY, V. V., BASALAYEV, A. A., BELYAYEV, 0. A., POZHILENKO, V. I. & RADCHENKO, M. K. 1994. An explanatory note on geological map of the north-eastern Baltic Shield on a scale of 1:500000. Kola Science Centre, Apatity. RAYEVSKAYA,M. B., GOR'KOVETS,V. Y., SVETOVA,A. I. & VOLODICHEV, O. I. 1992. Precambrian stratigraphy of Karelia. In: RYBAKOV,S. I. & STENAR M. M. (eds) Reference sections of Upper Archean deposits. Karelian Research Center, Petrozavodsk [in Russian]. RUZH'EVA, M. S., MATRENICHEV, V. A., VREVSKY, A. B., PIN'KOVA, L. O. & MYSKOVA, T. A. 2002. Kolmozero-Voron'ya-Uragubsky greenstone belt. In: GOLUBEV, A. I. (ed.) Mantle plumes and
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metallogeny. Geological excursion guidebook. Karelian Research Center, Petrozavodsk, 63-76 [in Russian]. RYBAKOV, S. I. & GOLUBEV, A. I. (eds) 1999. Metallogeny of Karelia. Karelian Research Center, Petrozavodsk [in Russian]. SAMSONOV, A. V., BERZIN, R. G., ZAMOZHNYAYA, N. G., SHCHIPANSKII, A. A., BIBIKOVA, E. V., KIRNOZOVA, T. I. & KONILOV, A. N. 2001. Early Precambrian crust-forming processes in NW Karelia, Baltic Shield: evidences form geological, petrological and deep seismic (4B Profile) studies. In: Deep structure and crustal evolution of the eastern Fennoscandian Shield: Kem'-Kalevala reflection profile. Karelian Research Center, Petrozavodsk, 109-143 [in Russian]. SAMSONOV, A. V., BOGINA, M. M., BIBIKOVA, E. V., ET AL. 2005. The relationship between adakitic, calcalkaline volcanic rocks and TTGs: implications for the tectonic setting of the Karelian greenstone belts, Baltic Shield. Lithos, 79, 83-106. SERGEYEV, S. A. 1989. Geology and isotopic geochronology of Archaean granite-greenstone complexes in Central and Southeast Karelia. PhD thesis, St. Petersburg State University [in Russian]. SERGEYEV, S. A., LEVCIqENKOV,O. A., ARESTOVA,N. A. & YAKOVLEVA, S. Z. 1983. U - P b isotopic age of the Semch gabbro-diorite intrusion. Izvestiya Akademii Nauk SSSR, Seriya Geologicheskaya, 4, 15-21 [in Russian]. SEARLE, M. P. & REX, A. J. 1989. Thermal model for the Zanskar Himalaya. Journal of Metamorphic Geology. 7, 127-134. SHCH1PANSKY, A. A., SAMSONOV, A. V., BOGINA, M. M., SLABUNOV, A. I. & BmII3.6 Ga) rocks that have been identified worldwide are dominated by so-called grey gneisses; that is, highgrade orthogneisses of mainly tonalitic composition (Windley 1976). The oldest U - P b zircon age from a magmatic rock has been obtained from the Acasta tonalitic gneisses, Slave Province, Canada. These have been dated to 4.03 Ga (Bowring et al. 1989; Bowring & Williams 1999), but some authors have suggested that these zircons are xenocrysts and thus do not date the host rock.
VOLYN DOMAIN PODOLIAN DOMAIN
B
KIROVOGRAD DOMAIN MIDDLE DNIEPR DOMAIN AZOV DOMAIN
EOARCHAEAN! PALAEOARCHAEAN! MESOARCHAEAN NEOARCHAEA Greenstone belts
PALAEOPROTEROZOIC
TTG (tonalite-trondhjemite-granodiorite) rocks
Intracratonic dfting (Banded Iron Formations)
mmmmmmmmmmi Late/post tectonic magmatism
Metasediment:s
~,
~
AMCG (anorthosite-mangedte-charnockite-granite) magmatism
high-grade metamorphism and deformation
Fig. 4. Compilationof rock ages and periods of high-grade metamorphismin the various domains of the Ukrainian Shield, based on previouslypublishedresults and our own partly unpublishedresults from the Podolian, Volyn and Ros-Tikich domains referred to in the text.
652
S. CLAESSON ETAL.
r Shield
Balt
25 ~ 52 ~
33 ~
, . _
Ukrainian
Shield
39 ~
50 ~ :! .....
i,
....
i
(b)
48 ~
2s~ 0
200
km
~p~
46040 '
29 ~
39 ~
3.2-3,0 Ga Middle Dnicpr domain
Undivided Archacan and l}alaeoproterozoic rocks
3.7- 2,8 Ga Azov and Podolian domains re~orked in the Palacoprotcrozoic
Palaeoproterozoic suture zones
2,2~2.1 G~ Tctercv - Ros orogenic bell
1,8-1.7 (ia gabbro-anorthosite-granitic plutons
"'"
"
Major boundaries or'domains anti belts
The most studied and best preserved area of ancient rocks occurs in West Greenland, where the oldest Am~tsoq tonalitic gneisses have been dated using several isotopic systems ( U - P b , S m - N d , R b - S r , P b - P b , L u - H f ) to 3.6-3.9 Ga (e.g. Black et al. 1971; Baadsgaard 1973; Nutman et al. 1996; Whitehouse et al. 1999). Even older ages, >_3.8 Ga, have been obtained for volcanic rocks from the Isua greenstone belt in Greenland (Baadsgaard et al. 1984; Nutman et al. 1997). Similar ages have been obtained for the Saglek gneisses in Labrador (Schi6tte et al. 1989; Bridgwater & Schi6tte 1991), for granulites from the Napier complex in Antarctica (Black et al. 1986), for tonalitic gneisses and greenstones from Swaziland in South Africa (Compston & Kr6ner 1988; Kr6ner et al. 1996), for gneisses from NE China (Liu et al. 1992), and for gneisses and anorthosite inclusions from western Australia (Kinny et al. 1988). The Palaeoarchaean Ukrainian rocks have a strong metamorphic overprint and complex deformational histories, which make more detailed interpretation difficult. In both the Podolian and the Azov Domain, the 3.65 Ga rocks appear to have been affected by high-grade metamorphism, in the late Archaean at c. 2.8 Ga, and again in the Palaeoproterozoic at c. 2.0 Ga. However, the Azov and Podolian domains are separated by the Middle Dniepr and Kirovograd domains, where the oldest known rocks are dated at 3.1-3.2 Ga, and the Palaeoproterozoic overprint on the Middle Dniepr Domain was very mild. There is no evidence that the temporal similarities in the evolution of the Dniestr-Bug formation and the Novopavlovsk complex reflect a common history. On the contrary, kinematic indicators in the fault systems of the Ukrainian Shield (Gintov 2004) show that the Archaean tectonic
Fig. 5. (a) Geological sketch map of the East European Craton (EEC), as in Figure 1. (b) Schematic geological map of the Ukrainian Shield, indicating its subdivision into highly reworked Palaeoarchaean core regions in the Podolian and Azov domains, a major Mesoarchaean granite-greenstone terrane that is largely unaffected by younger metamorphism or deformation in the Middle Dniepr Domain, and a large region dominated by juvenile crust, accreted in the Palaeoproterozoic, which includes the Volyn Domain, the northern part of the Ros-Tikish Domain and the Kirovograd Domain. Regions of major post-tectonic magmatism in the Volyn, Kirovograd and Azov domains, and Palaeoproterozoic suture zones, are also shown. The abbreviations are the same as in Figure 1.
evolutions to the east and to the west of the Kirovograd Domain were independent. Despite the striking age similarities between the Palaeoarchaean and younger rock complexes in the western and eastern parts of Sarmatia, tectonic correlation must therefore be conducted cautiously. The two parts of the shield were juxtaposed not earlier than 2.1-2.0 Ga, probably concomitantly with the collision between Sarmatia and Volgo-Uralia. In view of the lack of evidence for a common history, we consider that the temporal similarities between the Azov and Podolian domains indicate that 3.65, 2.8 and 2.0 Ga events were periods of tectonic activity in Sarmatia in a more general sense.
Palaeoproterozoic accretion
Our results emphasize the important role of Palaeoproterozoic geodynamics for the assembly of the Archaean components and the formation of the dominant structure in western Sarmatia. This applies even to its oldest, Archaean parts in the Podolian Domain. Palaeoproterozoic processes have also been important in the evolution of the Azov Domain in the east. The only parts that have not been significantly affected by these Palaeoproterozoic processes are the granite-greenstone terranes in the Middle Dniepr Domain. The multiple magmatism and granulite-facies metamorphism in the Podolian Domain may reflect active-margin tectonic settings both in the Neoarchaean at c. 2.8 Ga and in the Palaeoproterozoic between 2.1 and 2.0 Ga. The eastern part of the Ukrainian Shield, in particular the Azov Domain, should be studied further, not only to give more
ACCRETION AND REWORKING OF THE UKRAINIAN SHIELD
information about the oldest crust, but also to correlate Sarmatia and Volgo-Uralia, and to clarify the relationships with the Palaeoproterozoic in the western parts of the Shield and the role of tectonic processes along the S a r m a t i a - V o l g o - U r a l i a margin. This paper is a result of collaboration within the framework of the EUROBRIDGE project (EUROPROBE/ILP/ESF). Funding by INTAS, the Swedish Institute and the Royal Swedish Academy of Science helped E.B. to visit the Laboratory for Isotope Geologyin Stockholm to carry out isotope work. We thank the laboratory staff for analytical help and support, and R. Gorbatschev and D. Gee for constructive comments on the manuscript. Grants from the Swedish Research Council to S.B. are acknowledged. This is NORDSIM Contribution 133.
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Index Figures are indicated in italic and tables in b o l d Aar massif 131,134 Acadian deformation 293, 297, 300 Acadian event 324, 356 accretion 323-329, 551 Proterozoic 12, 510 accretionary complex 69, 70, 75, 265 Anatolia 396-400 Magnitogorsk arc 408, 410, 411,413 accretionary margin, East European Craton 291-306 accretionary prism 348, 349, 351 Oman 236-238, 240-242, 250 acritarch, Vendian 508 Adana Basin 270 Adria terrane 85, 93 Adriatic indenter 192, 193 Adriatic Moho 133, 134, 135, 138, 141 Adriatic plate 130, 136, 139, 192, 193 Aegean back-arc basin 171, 177, 272 ~ g i r Sea 86, 87-88 Aeolian Islands volcanism 174, 175, 179 African plate 358 convergence 223-225, 263, 265, 267-269 plate boundary 272 age see chronology, isotope and zircon Agnostus [trilobite] 513 Akershus Graben 365, 366 Albanian ophiolites 242-244 Alboran, volcanic province 172-176, 227 Alcapa terrane 193-195, 198, 199-202 alkaline magmatism 167, 172, 198, 362 volcanism 147, 150, 152-155, 158 Alleghanian orogeny 57, 59 ALP seismic experiment 143 Alpide orogeny 389, 390 Alpine deformation 1, 3-4, 129-143, 227, 281 thermal-mechanical controls 113-123 Alpine deformation front 98, 100, 102, 356 Alpine Orogen 51,277 Alpine subduction 171 - 172 Alpine Tethys 50-51, 59-62, 65, 75 Alpine-Mediterranean geodynamics 180-182 Alps 29-30, 129, 130, 131 geophysical data 21, 29-30 lithosphere thickness 24 Amazonia 326, 328, 568, 570 Amerasia Basin, opening 507, 508 amphibolite 243, 327, 351,395 East European Craton 293, 296, 300 Anatolia 64, 73, 246, 271,389-396 accretionary complex 396-400 chronostratigraphy 391, 397 seismic profile 271 volcanism 154, 177 Anatolian Fault 171 Andean-type collision 357 Anglo-Brabant Deformation Belt 301 anorogenic magma see alkaline magmatism anorogenic volcanic province 147, 151 Antalya Basin 73, 252, 268, 271 Apenninic arc 171 Apulia terrane 85, 130, 153 Arabian plate convergence 264 Arabian Platform 389 Archaean crust 600-601 Archaean Fennoscandian Shield 627-640 accretion 627 correlation 640 granitoids 635-639 greenstone complexes 627-634 metamorphic rocks 629-634, 640 metamorphism 636-639 provinces 628
archaeocyathids 89 Arctic Caledonides 509 Arctic, magnetic map 514 Armorica 89, 93 Armorica Massif, Quaternary folding 118 Armorican Archipelago 297 Armorican microplate (Franconia) 335 Armorican Terrane 85, 323, 325, 333 deformation 327-328 Armorican Terrane Assemblage 300 Asturian Phase 356 Atlantic opening 50, 223,224 Aubrac Cenozoic volcanism 152 Austroalpine nappes 135, 199 Autun Basin 370, 371 Avalonia 3, 57, 293, 296, 323 East European Craton, 324-329 emplacement mechanism 326-327 reconstruction 83-85, 89, 333, 460 soft collision 324 transect 301, 302, 305 triple plate collision 20 Avalonian suture 291 Thor-Tornquist 46
BABEL upper mantle project 545, 546-548, 554, 565 participants 555-556 seismic line 580 working group 15,299, 313, 315 back-arc basin 4, 5, 20, 192, 207 Aegean, 171, 177, 272 Avalonia 327, 328 extension 45, 48, 52, 170, 176 Guevgueli 248, 389, 400 Pannonian 191,200 Rheno-Hercynian 357, 339, 357 rift 30, 57, 62, 70, 238 topography 123 Baer-Bassit ophiolite 244, 247 Baikalian orogeny 510 Balkan orogen 59-60, 328, 399-400 Balkan suture 68 Balonia 293 Baltic Basin 458, 460 Baltic Shield see Fennoscandian Shield Baltica 84, 85-86, 89 East European Craton 294 Eurasia 507-509, 514-515, 521-536 mantle 2, 5, 24 rotation 447 transect 301, 303 Barentsian Caledonides 507, 509 basanite intraplate 147, 150, 152, 154, 161, 162 subduction-related 172, 176, 177 basement, East European Craton 482, 484-486, 510 Bashkirian Anticline 411,412, 413 BASIN seismic profile 375, 378 basins, east Mediterranean 263-273 bathymetry 264, 268, 269 Bay of Biscay, subduction 50-51, 75 B6k6s Basin 197, 199 Belomorian terrane 528 Belomorides, Archaean crust 546 Benioff zone 175, 177 Betic Cordillera 154, 225, 226, 227 Bey~ehir ophiolite 245,246 BIRPS seismic profile 29 bituminous shale 370 Black Forest, magmatism 148, 151,338 Black Sea orogeny 399-400
blueschist 237 Cenozoic 199, 390 Cretaceous 65, 398 Palaeozoic 397, 408 Triassic 72 Bohemian Massif 296, 323, 328, 333, 338, 339 Cenozoic 50, 97, 100, 107, 117, 123 volcanic activity 51, 148, 149, 153, 154 Bohemian Terrane 298 boninitic lava 243, 409, 410 Bouguer anomaly 227-230, 551 Brabant Massif 324 Brenner fault 201 Bresse Graben 97, 99, 103, 120 Brianqonnais domain 61, 62, 130 Bruno-Silesian Promontory 323, 325, 327 Bruno-Silesian Terrane 296 Budva domain 75 Burgundy Transfer Zone 51, 52, 120 Cadomian crust 455, 457, 460 Cadomian orogeny 3, 86, 323, 389, 516 Calabrian arc 265, 272 calc-alkaline magmatism 333 Balkan Terrane 328 Cenozoic 147, 152, 154, 168, 169, 176 Betic-Rif province 172-174, 178 Magnitogorsk arc 408, 410 Variscides 356, 357, 361-363, 370, 380 Calcareous Alps 59-60, 72 Caledonian Deformation Front 24, 301 Caledonian orogeny 85, 89, 93, 327, 453 Eurasian Arctic 507-516 Caledonian suture 380, 508 Caledonides 1-5, 19, 20, 25 crustal domain 43, 45-47 geophysical data 21 lithosphere thickness 24 ~amlik granodiorite 395, 400 isotope ratios 396 Campania, volcanism 175, 179 Cantabrian Mountains, Permian 351,372, 373 Cantal, Cenozoic volcanism 152 Cappadocia ignimbrite 177 carbonate platforms and ophiolites 240-242, 245, 249 carbonates 245, 251,400 Lower Palaeozoic 454-455,457, 512 carbonatite 147, 151, 174, 181 Carmel Fault 267 Carnic Alps, Variscan 334, 336, 337 Carpathian arc 178, 191, 194, 210-213, 216 lithosphere thickness 21, 23, 29 Pannonian region 173, 176, 180 Western domain 59-60, 170 Catalan Coastal Ranges 224, 226, 227 CELEBRATION project 6, 200, 313, 314 Cenozoic basins east Mediterranean 247, 268-272 Cenozoic magmatism 103-104, 149, 155, 172 volcanism 152 Central European Rift System 20, 21, 23, 30-31 Central Russia Rift System 19, 20, 24 Channel Basin 50 Chios accretionary complex 397 chronology 496 Anatolia accretionary complexes 397 Archaean greenstone 640 Carboniferous-Jurassic 74 Cenozoic Pannonian Basin 195 Cenozoic volcanism 150, 172 Cretaceous, Oman ophiolite 236 Dramala ophiolite 242
From: GEE, D. G. & STEPHENSON,R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 655-662. 0435-4052/06/$15.00 9 The Geological Society of London 2006.
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656
chronology (Continued) Evia ophiolite 241 Greater Caucasus 285 Lycian ophiolite 245 Mersin ophiolite 245 Palaeozoic, East European Craton 452 Permian, Germany 363 Pontic terranes 391 Pyrenees 373 Cimmerian cycle 5, 57, 59, 277, 278, 358 tectonics 400, 494-498 Ciudad Roderigo Basin 226, 227 coal 359, 369, 372, 373, 392 Carboniferous 335-336, 338 collision structure 138-139, 140, 240 collision tectonics 4, 5, 6, 20, 25, 51 arc-continent 357, 407-410 Avalonia 327, 328 Baltica 407-409, 416 continent-continent 43, 99, 167, 356 Gondwana-Laurussia 355, 357, 358 iberia 345-351 Plio-Quaternary 253 Proterozoic 579-597 Pyrenees 224 soft 85, 170 Svecofennia 568-569 Variscides 99, 345-351 compression 498 Greater Caucasus 281-282, 283 Iberian Peninsula 224-227 conductivity 544-547 Conrad discontinuity 134 controlled source seismology 130 convergence 390, 410, 443, 454-457 convergence rate 50, 51, 167 Cordilleran-type ophiolite 237, 247-253 Crete Basin 272 Crimea, Triassic to Jurassic 496 cross-sections see transects crust 24, 130-136 Archaean 639-640 properties 542-546, 549-550 Proterozoic 548, 606-608, 613, 614 Russia 521-536, 585-589 shortening values 100 strength 116-117, 1 2 3 - 1 2 4 crustal evolution 43-53, 365,380, 481-482 crustal lamination 103-104 crustal structure, East European Craton 291,300 transects Avalonia- Rhenohercynian 302 Caledonide 304 Variscan 303, 305 see also delamination crustal thickness 18, 116, 131,200, 209 map 45, 46, 47, 48, 49 Southern Permian Basin 102-104 crustal wedge 195, 351,348 crystalline basement 391,474, 529, 531,533 East European Craton 292, 298 Cycladic domain 64, 73 Cyclopyge [trilobite] 328 Cyprian arc 266, 268, 272 Danish graben 104, 365, 376 Danube Basin 198, 199, 326 Dead Sea Rift 264, 268 debris flow 251,409 Balkans 240, 242-244, 246, 247, 249 DEKORP basin research group East European Craton 299, 300, 301, 304, 305 DEKORP seismic profile 24, 103, 314, 371,375 delamination 273, 285, 351,379, 561 Alps-Mediterranean 167, 169, 170, 174, 175, 179, 181 rifting 101, 102, 154 mantle 97, 574
INDEX
Denmark, TOR project 314 density variation 18, 18, 32 density, Baltic Shield 32 diamond 100, 337, 408 diatremes 151 Dinaride-Hellenide passive margin 59 Dinarides 75 Dinarides, volcanism 173, 176-177 Dnestr Basin 449, 457, 460 Dnieper-Donets-Pripyat rift 21, 22, 27 -28 lithosphere thickness 24 Dniepr-Donets Basin 380, 463-468 DOBRE reflection profile 28, 104, 465-468, 474, 475 Franconian Platform 104 Dobrogea Platform 326 Dobrogea terrane 298 Dobrudzha Trough 72 Dowsing Lineament 297 Dramala ophiolite see Pindos ophiolite Drinova Thrust 326 Dublin graben 45 Duero Basin 226-228, 231 dunite 444, 446 dyke swarm 4, 18, 361 Earth model IASP 138 earthquake 120, 124, 284 hypocentres 213 tomography 139 TOR project 314, 315, 319 East Barents Basin 470-472 East European Craton 1- 8, 11, 277, 389 accretion 293-299, 323-329 break-up 299 correlation 451, 452 geodynamic evolution 599-621 geophysical interpretation 616-621 gravity map 295 gravity model 609-615 magnetic map 294 named basins 450 passive margin 327, 328 Proterozoic crust 407, 422, 606-609 rift basins 463-476 southeast margin 481-498 subsidence curves 453 sutures 299-300 tectonic map 459, 464, 600, 601 terrane provenance 291-293 terranes 602-606 western accretionary margin 291-306 terranes and sutures map 292 East European Platform 11, 17-19, 20 geophysical data 21, 27-28, 117 lithosphere thickness 24 Ebro Basin 51,226, 227, 231 eclogite 103, 327, 351,638 Anatolia 390, 397 Bohemia 298 Caspian 470 Uralides 422, 515 Variscan 335, 337, 338 ECORS, Alps seismic profile 24, 129-132, 135, 139 lithosphere slab geometry 141-143 ECRIS see European Cenozoic Rift System Ediacaran fauna 86 EEC see East European Craton effective elastic thickness 213, 215, 219 Eger Graben 97, 149, 152 EGT seismic experiment 305 Eifel plume project 313,320 Eifel volcanic field 150, 155 elastic-plastic plate model 119 Elbe Fault 358 Elbe Line 314, 317 Elbe Lineament 291,301
electromagnetic conductivity 542-547, 554, 555 techniques 543-544 see also EUROPROBE enriched mantle component 181 Entomozoe [ostracode] environment change, Eocene-Oligocene 283-284 Eocene volcanic activity 148 Eocimmerian 59, 69, 70, 498, 278-280 Eo-Variscan orogeny see Caledonian Eratosthenes Seamount, continental crust 266, 268, 397 ERCEUGT group 303 EREGT working group 29 erosion 155, 216, 217, 337, 339 Erzgebirge diamonds 337 ESRU 408, 414, 416, 421,422 data processing 431-433 seismic profile 24, 26 seismic reflection data 427-440 Etna 153, 174 EUGENO working group 313,314, 315 East European Craton 299, 300 Eurasian Arctic 507-516 EUROBRIDGE 17 EUROBRIDGE profiles 599-621 European Cenozoic Rift System 51, 97-108 lithosphere 113-115, 117, 123-124 map of faults 98, 100 reflection lines, location 100 sedimentary basins 98, 102 European Geotraverse 11, 26, 117, 374 refraction profile 117, 129, 130, 131, 133, 134 EUROPROBE 11, 27-28, 129, 192, 203, 207 Dniepr-Donets Basin 421,463 station map 543 SVEKALAPKO 521-536 tomography experiment 549-554 Urals 407 evaporite 94, 227, 231,529 Messinian 264, 268, 270 Palaeozoic 365, 370, 375, 379, 467, 471 Evia ophiolite 240, 241 extension 4, 99, 333,568 in compressional setting 191,195 Greater Caucasus 280-281 Iberian Peninsula 223-224, 230 post-collision 170 Variscan 103, 333 extinction event, Permian 83 extrusion tectonics 191, 194, 202, 217 Fallotaspsis [trilobite] 513 far-field effects 381,458, 460 faults 46, 47 FENNOLORA seismic profile 313, 319, 541 Fennoscandia 541-556 tomography experiment 543, 549-554 Fennoscandian Shield 114, 117, 542, 582 crustal structure 522-529 geophysical studies 21, 23, 32, 541-556 lithosphere 24, 579-593 map 7 provinces 628 seismic data 12-15, 20 tectonic map 580, 628 terranes 602-605 Finike Basin 271 fish fauna, Old Red Sandstone 512 flexural downloading 210-212, 283, 285 flood basalt 94 flower structure 140 flysch 424 Devonian-Carboniferous 335-338, 348, 350 fold-and-thrust-belt 347, 348-350 folding and lithosphere strength 117-119 folding, Greater Caucasus 281-282 folds 217, 218, 282, 283 fore-arc basin, Palaeotethys 69-70 forebulge, Caledonides 455, 457 foreland basin 45, 59, 449
INDEX
fossils, Gondwanan affinity 292, 297 Fourier' s law 117, 318 Franconian Platform 107, 335 gabbro, Silurian 443-447 Galahetes [trilobite] 513 Galatia volcanic province 154 garnet, age 335, 348 geochemistry 157-160, 173 Betic-Rif province 172-176, 178 Fennoscandia 585-589, 627-636 primitive magma 158-160 subduction-related magma 168-170, 181 GEON center 523, 531 geophysical data 21-22 Alps 28-29 Fennoscandia 12-15, 20, 21, 23 lithosphere properties 19-24 lithosphere thickness 24 geophysical interpretation 11 - 33 Baltica 521-536 East European Craton 313- 320, 616- 621 Fennoscandia 541-556, 581-583 Permian basins 374-376 upper mantle 156-157 Trans-European Suture Zone 18, 19 GEORIFT 27-28 geotherms 16, 19, 26, 31,410 Germany 314 Cenozoic volcanism 155, 156 Variscan orogen 333-340 Germany, north, transect 301 Gfrhl Suture 300 Giessen Ocean 99, 327 Giessen ophiolite 297 Gissen-Harz Basin 45, 102 Giudicarie belt 136, 139, 140, 143 glacial deposits, Ordovician 89, 329 glaciation 86, 92, 94 Gloria Fault 223 Gltickst~idler Trough 363 Gondwana 5, 57, 83, 99, 355 accretion 323-329 break-up 328, 391 collision with Bohemia 336, 337 convergence 356, 379, 381 map showing named units 58-59 reconstruction 84, 87-91 tiffing 293 Variscan 333 Goniatites [ammonoid] 338 graben 97, 151,266, 267, 270, 363 granite 18, 65, 393 Variscides 70, 338, 339 granitoid plutonism, Pontides 395-396 granitoid, Fennoscandia 545,635-639 granitoid, Uralian 411,414 granulite 103, 243, 298, 337, 338 Lapland Belt 528, 579, 581,582, 585-589 Lapland and Umba terrane 591 Lithuanian 602-603 Pontides 395, 398 Urals 425 graphite as conductor 542, 544, 545 gravitational collapse 171, 174, 203, 217 Fennoscandia 569-570, 573, 592-593 gravity 18, 27, 32, 253 Danube Basin 198 East European Craton 600-615, 616 maps of Europe 3, 7, 18, 295 NW Russia 523 Poland 375-376 Timan Range 531 Uralides 26-27 Greater Caucasus chronology of tectonic events 285 compression (mid Jurassic) 281-282, 283 compression, Cenozoic 283 Eo-Cimmerian 278-280 extension (early Jurassic) 280-281
inversion, Cenozoic 283-286 rate of shortening 284 post-rift succession 282-283 structure 278 topography 278 and tectonics 284 greenschist 243, 327, 395, 398 Fennoscandia 627-639 Neoproterozoic 510 Grenville orogeny 3,532, 564 Guadalquivir Basin 226, 227 Guidicaria Fault 193, 201 Gulf of Lyons, opening 225 gypsum 282, 511, 513 Halloporina [bryozoa] 513 Hanseatic terrane 57 Hatay ophiolite 244, 247, 25 l Hawasina Complex 241,246 Haybi Complex 241, heat flow 12, 19, 25, 26, 27, 208, 214 modelling 375 values 15, 17, 19 Hebediscus [trilobite] 513 Hegau volcanic field 151, 159 Hellenic arc 85, 93, 171, 176-177, 191 Helvetic nappes 129 Hercynian suture zone 30 Hessian Graben 97, 114 Himalayan-type collision 43, 99, 167, 356 H6d-Mak6 Basin 197, 199 Holstein-Horn Graben 363 Holycross Mountains 304 Palaeozoic sequence 295-296 Hun superterrane 57, 58 Hungary Cenozoic volcanism 155 Hyblean Platform 269, 270 hydrocarbons 6, 219, 362, 377 hydrothermal vent communities 93 Iapetus Ocean 87-89, 460 opening 293 subduction 296 Iapetus Suture 20, 45, 83, 85 East European Craton 299, 304 Iberia lithosphere thickness 229-230 Iberian basins 372-374, 376 Iberian Massif 345-351 structure 346, 349 Iberian Peninsula 223-231 compression, Cenozoic 224-227 mantle plume 351 Mesozoic extensional basin 223-224, 230 Neogene basins 227 inversion 230-231 tectonic map 226 topography 223,227-231 Variscan basement 228 Iberian pyrite belt 349, 351 IBERSEIS seismic profile 20, 346-348, 351 Iceland plume 50 ILIHA seismic profile 20, 25 imbricate thrust 227 Indonesian archipelago 570, 574 inselberg pattern 195-196, 196 Insubric Line 131,135, 171 intermontane basins 227, 231,373 intraplate lithosphere 113-119 Intrapontide Ocean 67 inversion 122, 208, 209, 281 Cenozoic 283-286 Iberian 230-231 Laramide 155 Moscow Basin 460 Pannonian Basin 195-196, 208, 214, 216, 219 Polish Basin 378-379 Tethyan 50 Tornquist 319 Ionian Basin 171,263, 264-265 IPL-ALCAPA project 207
657
iron and seismic velocity 12, 32 iron formation, Archaean 632, 634 Iskenderun Basin 270 island arc 407, 421 Mid-German High 333 isotope age 443 East European Craton 293 Magnitogorsk intrusives 409 Maksyutov Complex 422 Sakarya Zone 396 see also zircon isotope chemistry 158-160 isotope measurement 444 isotope ratio 169, 175, 178, 180 Kazda~ metamorphic rocks 396 Isparta Angle 250, 251 Israel, graben and horst 266 Istanbul Block 326-327, 391-394, 400 deformation 394 succession, Mesozoic 394 Italy, volcanism 174-176, 178 Cenozoic 153, 155 Ivrea Zone 131,133, 135, 138, 139, 141 Izanca 64-66 Izhma Domain 532, 533 Izmir-Ankara ocean 67, 73, 75 accretionary complex 397, 400 rift 72 suture 64, 171, 177, 238, 398 Jura Mountains, compression 120 Kachkanar massif 445 Kaiserstuhl, volcanic activity 151, 155 kamafugite 181 Kanin peninsula 531 Kara magnetic anomaly 513-514, 515, 516 Kara Terrane 86 Karaburun accretionary complex 397- 398 Karakaya Basin 72 Karakaya Complex 238 Karelia Domain 12, 541,545, 546 Karelian Craton 579 Kasimlar basin 59 Kazakhstan 422 Kazakhstan arc 58 Kazakhstan plate 26, 409-411,416 Kazakhstania 84, 86, 87, 94 Kazbek volcano 284 Kazda~ metamorphic rocks 395, 396 Kempersai massif 443 Khanty-Mansi ocean 58, 71 Khoreyver Domain 533 Kimzha graben, seismic profile 530 Kipchak arc 58 Kir~ehir Massif 389 Kola deep borehole 524 Kolva deep well 471 Kujandaspis [trilobite] 513 Kuloy graben, seismic profile 530 Kumba gabbro, age 446 Ktire 62, 63, 64, 71 Ktire complex 397-398 Kttre ophiolite 238 Kytlym dunite 446-447 lacustrine deposits Iberia 226 Oligocene 226 lamproite 174, 176, 177, 178 lamprophyre 153 Lapland-Kola orogen 579-593 cooling age 592 Laramide basin inversion 155 Larnaca Basin 270 Latakia Basin 270 Laurasia 399, 400
658
Laurentia 5, 24, 293, 507, 508, 509 palaeoequator 89 reconstructions 83, 84, 85, 86 Laurussia 57, 58, 93, 99, 392, 400 collision 355, 356, 379 palaeogeography 394 passive margin 327, 328 reconstructions 85, 86 Leshukona-Pinega Rift 531 leucities 150, 176 Levant margin, seismic profile 267 Levantine Basin 266-268, 270 Ligeran Phase 356 Limagne Graben 97, 103, 114 Linosa, magmatism 153 Linosa, seismic profile 269 listric fault 199 LITHOPROBE 15, 19 lithosphere 472, 474, 475 conductivity 545 folding 117-119, 124 profiles 23 strength 117-119 lithosphere structure, Svecofennia 561-562, 571-573 lithosphere thickness 24, 31, 97 East European Platform 27-28 Iberia 229-230 Pannonian Basin 201,202 south Europe 25-26, 27-28, 31 lithosphere, Baltica 459-460, 542-546 lithosphere, East European Craton 291 lithosphere, Fennoscandia 579-593 lithospheric mantle composition 31-32, 117 Lizard 45, 99, 333 MORB 71 ophiolite 394 peridotite 293, 297, 300, 327 Lizard-Rhenish suture 358 London-Brabant Massif 45, 117 strength of lithosphere 123 Lower Rhine Graben 97 Lower Rhine Lineament 297 Lviv slope, carbonate 456 Lycian domain 63 ocean 66, 67, 75 ophiolote 245, 246, 247 Lyciam nappes 64 lydites 70 Lysogory terrane 85 magma geochemical characteristic 157-160 subduction-related 158, 167-203 magmatic fields Permo-Carboniferous 76, 48-49, 339 Oligocene 49 magmatic underplating 547, 548, 551 magmatism, age 172 Carboniferous 99 Carboniferous-Permian 97, 101-102, 104 Cenozoic 4, 31, 114 Devonian-Carboniferous 6, 348, 349 Permo-Triassic 5 magmatism, Armorican Terrane 327 magmatism, Cenozoic intraplate 147-162 age 150-155 and basement uplift 155-157 geochemistry 157-160 geodynamic setting 147-150 source 160-161 magmatism, Greater Caucasus 280-282 magmatism, Variscan 359-362, 380-381 distribution map 360 foreland 359 Germany 361 Iberia and west Mediterranean 362 Massif Central 362 Scotland Midland Valley 361
INDEX
Variscan foreland (externides) 362-367 Variscan internides 367-374 Whin sill 361 magnetic anomaly 532, 533 Massif Central 48 Paris Basin 48 Saar-Nahe 103 magnetic map Arctic 514 East European Craton 601 Europe 2 NW Europe 294 NW Russia 522 magnetite 634 magnetotelluric (MT) studies 29 Magnitogorsk arc 407-411,413 Magnitogorsk block 27 Magnitogorsk-Tagil island arc 422, 424-425 Main Caucasian Thrust 278, 279 Main Uralian fault 27 seismic interpretation 436-437 suture 407, 414, 421,443 Maksutovo Complex, radiometric age 409 Maksyutov Complex, subduction 421-422 Maladiodella [trilobite] 513 Maliac 62, 63, 64 back-arc basin 70 ocean 73 Malopolska Terrane 85, 293, 295 Malta Trough 269 seismic profile 270 Mamonia Complex 252 mantle conductivity 545 mantle convection 156, 162 mantle diaper 147, 162 mantle discontinuity 553 mantle events, Svecofennia 566 mantle experiments 542 mantle model, East European Craton 611 mantle peridotites 552 mantle plume 30, 46, 51 Archaean 632 Cenozoic 115, 122, 123, 157 rift-related 97, 101, 103 subduction-related 175, 181, 182 Mesozoic 360, 380, 584 Palaeozoic, East European Craton 473 mantle plume dynamics 345-355 mantle properties 542, 546-547 mantle reflector 15, 566 mantle structure 11-33, 549, 551 mantle temperature 12-14 mantle wedge 410 marginal ocean sequences 71-75 Massif Central 101,336-338, 367-370 geophysical data 21 lithosphere thickness 24 lithosphere strength 123 Permo-Carbioniferous trough 107 uplift 149 Variscan basement 152, 155 volcanic activity 31, 51, 148, 152 mechanical strength of lithosphere 115-119 Mediterranean basins 52 Mediterranean Ridge deformation front 263, 264, 265, 265 accretionary wedge 264-265 Mediterranean, geophysical model 28-29 mrlange 327, 410 east Mediterranean 239, 243,245,246 collapse of platform successions 240, 242 Tanaelv 592 Meliata 62, 63, 64 Meliata-Hallstatt ocean 72 melilite 151, 152, 161, 174, 181 melting curves 160 Menderes Massif 390 Mersin ophiolite 245, 247 Mesozoic basins, east Mediterranean 263-268 Messinian salt 264, 268, 270
metamorphic sole/ophiolite 239, 243, 244, 246 metamorphism 61, 63, 65 Armorican Terrane 327, 329 Fennoscandian 584, 636-639 Maksutovo Complex 408-409 Pontides 395-396 Variscan 57, 328, 333,337-339 meta-sediments, Archaean 631-634 metasomatism 170 meta-volcanics and intrusions Archaean 629-634, 640 Mezen Basin 522, 523-525, 528-530, 531 Mid-German Crystalline High, Variscan 297, 333, 335 cooling age 370 Mid-German Crystalline Rise, Cenozoic 99, 100 Midland Valley, Scotland 45 Midlands Microcraton 296, 301 mid-ocean ridge basalt 158, 173 trace element 169 mid-ocean ridge basalt, Palaeotethys 69, 70 Mid-Polish Trough 375, 376 Milankovi6 cyclicity 196 mineralization 6 Miocene volcanic activity 149, 150 Mobil Search, seismic survey 15 Moesia terrane 85, 298-299 Moesian Platform 298, 326, 391,392 Moho depth 348, 566, 609 Adriatic 133, 134, 135, 138, 141 depth controls 43-45 East Barent Sea 471,474 Fennoscandia 547, 548 Liguria 133, 138 NE Baltica 523, 528, 531,533 Pannonian Basin 199, 200 petrology 103 post-Variscan 97, 98, 101,375, 376 seismic 104 temperature 15, 102 topography 131 - 134 Uralides 416 Moho offset 131,133, 138, 139, 141 Tornquist Zone 300 Moho reflector, TOR data 318- 319 Mohorovirid discontinuity see Moho molasse 455 Carboniferous 335, 336 Cenozoic 193 Permian 400 Moldanubian ocean 336, 337, 338, 339 Moldanubian Terrane 100, 293,298, 305 Moldova carbonate platform 455, 456 MONA LISA Working Group 19, 20, 25 East European Craton 297, 299-301 Mont Dore, Cenozoic volcanism 152 Monte Vulture, magmatism 153, 174 Moravian Suture 300 Moravicum nappe 326 Moscow Basin 449, 454, 457 inversion 460 isopachs 455 Mugodzhar-Khanty-Mansi ocean 71 Mut Basin 270 Nagorskaya drilling project 508 nappes Alpine and Carpathian 51 Austroalpine 135, 199 Helvetic 129 Lyciam 64 Moravicum 326 Pindos 68 Semail 69 NARS seismic profile 20 Neogene basins, Iberian Peninsula 223, 226, 227 Neogene uplift 113, 114 Neoproterozoic accretion 484-486 Neoproterozoic subduction 392
INDEX
Neotethys 5, 59, 104, 130, 399 active margin 249, 251,252 closure 283 evolution 70-71 mrlange 244, 245 nomenclature 66-69 opening 358, 359 ophiolite emplacement 237, 238, 248 reconstructions 61-68 rift 50 rotation 253 nephelinite 150, 151, 152, 161 Nevadella [trilobite] 513 New Red Sandstone 94 NFP transect, Alps 134, 136, 139, 143 seismic profiles 130, 132, 135, 142, 131 Nicholsonella [trepostom] 513 Nordaustlandet Terrane 507-508 Normannian complex 300 North Aegean Trough 272 North African plate margin 247 North Anatolian Fault 272 North Atlantic 120 rift system 59 opening 51-52 uplift 123 North Danish Basin 47-50 North Danish-Polish Trough 117 North German Basin 356, 360, 378 North Kara Terrane 508, 510-513, 515 sedimentary succession 512 suture 511 North Sea Rift System 29, 50, 117, 123 North Sea subsidence 52 Northern Permian Basin 49, 355, 363, 365 Northumberland graben 45-46 Norwegian-Greenland Sea 508 Novaya Zemlya 509-510 nuclear explosion, Murmansk-Kizil 523 obduction Apulian plate 153 Beja-Acebuches 348, 351 Lizard complex 300 Saxo-Thuringian ocean 338 Urals 423, 424 Occam model 547 Ocean Island Basalt 157, 172, 174, 177, 179, 181 oceanic spreading 487-490 octupole component 86 Old Red Sandstone facies 512 Scandian orogeny (Silurian) 455 Oligocene volcanic activity 149 Oligocene-Miocene continental deposits 226, 227 olistolith 71 olistostrome 72, 227, 283, 398 see also debris flow deposit olivine conductivity 547 Oman-type ophiolite 235-238, 240, 244, 253 sedimentary cover 244 ophiolite 4, 71, 72, 398 Anatolides 63 Arabian 69 Archaean 565, 627, 632-634, 640 Baer-Bassit 271 Balkan 238-244 Bay of Islands 235 Croatian 73 Guevgueli 248 Palaeozoic 70, 99, 286, 356 Uralides 515 ophiolite obduction Apulian plate 153 Beja-Acebuches 348, 351 Dinarides 59, 75 ophiolite protolith, Bohemia 298 ophiolite, east Mediterranean 235-254 active margin (Cordilleran-type) 237, 247-253 collision-trench 235-237
passive margin (Oman-type) 235-237, 238, 240, 244 volcaniclastic sediments 249, 250 orogenic chains 224 orogenic magma 168, 178, 181 orogenic wedge 191,203 Oslo Graben 47-48 Oslo Rift 365-367 geophysics 376 Palaeozoic 27-28 post-Variscan reactivation 380 rheological model 377, 378 sedimentary fill 363 volcanism 360-361 PACE network 300 Pagetiellus [trilobite] 513 Palaeocene volcanic activity 148 palaeogeography Baltic Basin, Silurian 458-459 Devonian 464 East European Craton depocentres 451 Laurussia 394 Oman ophiolite 239 Palaeotethys 398, 399 Permian 355-359 Tethys 399 Palaeolenus [trilobite] 513 palaeomagnetism 193 Neotethys rotation 253 plate reconstruction 83-94 Palaeoproterozoic basement 389 Palaeotethys 5, 57, 58, 94, 99 accretionary complex 75 closing 358, 359 forearc sequence 69-70 mid-ocean ridge basalt 69 nomenclature 66-69 opening 333, 337, 339, 359, 357 ophiolite emplacement 237-238, 239 palaeogeography 398, 399 reconstructions 60-68, 91-93 sequence in Iran 69 subduction 59, 281 Palaeozoic orogens 19-28 Panafrican deformation see Cadomian orogeny PANCARDI project 192, 207 Pangaea 57, 86, 104, 355,359, 416 break-up 5, 50, 75, 266, 363 formation 293 reconstruction 92-93, 94 Pannonian Basin 4, 29-30, 52, 117 crustal thickness 200 extrusion tectonics 202 geophysical data 21, 23 lithosphere strength 123-124 lithospheric structure 200-202 lithosphere thickness 24, 201, 202 magmatism 179, 181, 195, 196-198 MORB diagram 173 pre-Neogene basement 194 seismic reflection section 197, 199 stratigraphy 195, 196-198 tectonic framework 192-193, 195 topography 215 volcanic deposits 195 volcanism, Cenozoic 152, 154 Pannonian, basin evolution 207-208 back-arc extension 170 deformation 198-200, 214-216 depositional environment 195, 196 formation 191-203 inversion 195-196, 208, 214, 216, 219 rifting 208-209 Pannonian-Carpathian system lithosphere strength 213-215 subsidence 208, 209, 211,212 stretching 209, 210, 214, 216-217
659
thermomechanical modelling 207-219 Pannotia supercontinent 296, 328 break-up 299, 303 Pantelleria, magmatism 153, 174 Panthalassic Ocean 86, 87, 359 Paphlagonian Ocean 71 Paradoxides [trilobite] 326, 513 Paratethys stages, 195 Paris Basin 50, 104, 119, 376 magnetic anomaly 49 subsidence curve 106, 107 Parnassus block 239 partial melting 158 and conductivity 542 and velocity 156 Partnach Basin 72 passive margin 6, 120, 123,359 Arabia 237 Caledonian collision 320 East European Craton 327, 328, 475-476 east Mediterranean 267, 272 Eurasian Arctic 509 inversion 122 Neoproterozoic 313 ophiolites 235-237, 238, 240, 244 Ordovician 486 peri-Tornquist 449-454, 457 passive rift 27-28, 30 inversion 208, 209 Peaceful Nuclear Explosion Profile Quartz 17, 26 Pechora Basin 470-472, 522-525, 531-534 igneous rocks 533 sedimentary cover 533 Pelagian block 268, 269, 270, 272 delamination 273 Pelagonian carbonate platform 240, 241 Pelgonian terrane 59, 390 Peltura [trilobite] 513 Penninic front 368 Periadriatic Line 200 magmatism 171, 172, 177-179 Periadriatic Lineament 140 Peri-Caspian Basin 451-452, 456-457, 468 -470 peridotites 30, 158, 160 mantle temperature 15 peri-Uralian basins 449-451,456-457 Permian basin evolution 355-381 Permian basins basin fill 355, 362 chronology 363 development 376-380 France 367-370 Iberian succession 373, 374 isopach maps 364 magmatism 358 marine deposits 373 modelling and basin history 376-379 palaeogeography 355-359 rifting 379-380 syn-rift sequence 71 Switzerland 367, 368 Variscan foreland (externides) 360, 362-367 Variscan internides 367-374 Permian orogens 5 Permian peri-glacial fauna 71 Peronopsis [trilobite] 326 Perunica (Bohemia) 85, 89, 293 petrogenesis, Cenozoic magma 178-182 Phlegrean Fields, volcanism 175 phonolite 152 Piedmont-Ligurian ocean 129 Pindos 63-67, 70 domain 73 nappe 68 ophiolite 238, 239, 242, 243 Pindos Zone 390 plate margin, Anatolia 389
660
plate reconstruction 57, 498 Armorican Terrane Assemblage 325 Cambrian 84, 87-93 Carboniferous 91, 92 Cretaceous 225 Gondwana 84, 87-91 Neotethys 60-68, 83-94 Ordovician 88-89 Palaeotethys 60-68, 91-94 Pangaea 92-93 Permian 5, 92, 93 Permian-Triassic 359 Precambrian 485 Silurian-Devonian 90, 91 Tethys 6 0 - 6 7 Urals 416 Vendian 84 plate tectonic evolution 208 Precambrian 15 Svecofennia 571-573 Urals 424 Variscides 334, 336, 351 plateau basalt 152 platform carbonates Palaeozoic 450 and ophiolites 240-242, 245, 249 platinum-bearing belt, Urals 443-447 playa lake 363 Plinian eruption 150 Pliocene volcanic activity 149 plume activity 379 plume induced magmatism 328 plume-related structures 20 geophysical data 21 Po basin 29 Po Plain 131, 153 Podlasie basin 449 Poland, seismic profile 375-376 transect 303 POLAR seismic profile 583 Polish Basin 364-365 cross-section 365 inversion 378-379 Polish Trough 49-50 POLONAISE seismic survey East European Craton crustal model 304, 304 Fennoscandia 6, 26, 313, 314 Trans-European Suture Zone 318, 319 interpretation 375,608-615 Pontes Basin 226 Pontide volcanic arc 170 Pontides 63, 93, 389-390, 391-396 molasse, end-Variscan 400 suture 393 Pontides, Karakaya Complex 238, 249, 250 pop-up structure 226 potential field data 521-522 Precambrian 422 lithosphere 11-17 see also Archaean, Neoproterozoic and Proterozoic Precordillera Terrane 87 Proterozoic crust, East European Craton 606-608 Protoatlantic 99 pull-apart basin 101, 195 pull-apart structure 361 Cornwall 380 Pyrenees 65 Carboniferous-Permian rocks 373 fold and thrust belt 225 geophysical data 22 lithosphere thickness 24 rift 75 volcanism 372 pyroclastic deposits 153, 154 Permian 361 Racha-Lechkhumy Fault 278, 279, 283 radiolarites 238, 240, 243 Ran Ocean 86
INDEX
Rayleigh wave 313, 315, 317, 552, 553 Rechnitz window 193 reconstruction see plate reconstruction Rheic Ocean 57, 58, 293, 327, 333, 349 suture 46, 83, 88-89, 327, 359 Rheic Suture 297, 299, 300 Rhenish Massif 97, 336 volcanic centre 51 Rheno-Hercynian Belt 45 Rheno-Hercynian Ocean 58, 71,334, 375 subduction 336, 338 Rheno-Hercynian Shelf 45 Rheno-Hercynian suture 48, 103 Rheno-Hercynian terrane 85, 90, 93, 94 rheology 116, 123,213-215,218, 319 Rhine Graben 30, 150, 154 lithosphere thickness 24 Rhine Rift System 117 Rhinish Massif, volcanism 148, 150, 155 Rhodes Basin, seisnfic profile 272 Rhodope Massif 398, 399 rhyolite dome 362 rift basin 49-59 Norwegian 120, 121 Permo-Triassic 5, 7 rifting 104-105, 333, 488-490 Caucasus 278, 280-282 Cenozoic 99 and magmatism 149-150, 154 Cretaceous-Palaeocene 50-51 East European Platform 18 Fennoscandia 583-585 Levantine Basin 266-267, 272 and lithosphere strength 121 maps 47-48 Mesozoic 266 Neotethyan 264, 270, 271 Palaeozoic 27-28, 463-476 Pannonian Basin 208 Strait of Sicily 268-270, 272 Triassic 223, 400 Urals 450-451,457 Variscan 358, 363, 376-380 Ringkcbing-Fyn High 313, 314, 319 Roccamonfina, mantle-source contamination 179 Rockall-Faeroe Bank 51 Rockall-Faeroe Trough 50 Rodinia 3, 7, 296 break-up 299, 303, 449-454, 457 Roer Valley Graben 120 roll back 35,273 Carpathians 210 Pannonian Basin 191, 193, 202, 203 Roman province 174 Romanian Terrane 298 Rondonian event 323, 326, 328 Rondonian-type crust 296, 305 rotation Baltica 447 East European Craton 293 Neotethys 253 Troodos 251 Rotliegend 363, 365, 366 depth to base (Upper) 365 volcanism 376 Russia, northwest geological provinces 525, 528 geophysical provinces 522 geophysical studies 541 state geophysical company 524 Saalian unconformity 362 Saar-Nahe Basin 49, 370, 371 Saar-Nahe Trough 103 sabkha 365 Sakarya domain 72 Sakarya Zone 328, 389-391, 395-397, 400 Sakmara arc 447 salt see evaporites
Sardinia magmatism calc-alkaline magmatism 178, 179 intraplate 153, 155 Sarmatian terrane 605-606, 617 map 7 Saros Trough 272 Saxo-Thuringian Basin 100, 335 ocean closure 336, 338, 339 Saxo-Thuringian Suture 300 Saxo-Thurngian terranes 297, 305 Scandian event 293 Scandian orogeny 455, 460, 509 Scandinavian basin 453-454, 455 Scandinavian Caledonides 516 schist as conductor 544 Schmidtiellius 296 Scythian Platform 3, 5, 7, 277, 279, 391,400 sea-floor spreading 584-585,593 east Mediterranean 272 sedimentary basins 470-472 East European Craton 449-460 subsidence curves 450, 453, 454 Iberia 225, 226-227 sedimentary sequence 493, 490-493 Eurasian Arctic 508, 511-513 Permian basins 362-374 Permo-Triassic 491 Proterozoic 529, 530, 531,533 sedimentary succession, post-rift Greater Caucasus 282-283 seismic activity, Vrancea 201,213, 216 seismic data/studies 6 East European Craton 15-17, 608-615 Fennoscandian Shield 547-554 Iberia 346-348, 351 NW Russia 523-525 Strait of Sicily 269-270 Urals 27 Variscides 25 seismic profile Antalya Basin 271 Baltic Basin 456 Fennoscandia 582 Ionian Basin 265 Levant margin 267 Linosa Trough 269 Malta Trough 270 Pannonian Basin 197, 199 Rhodes Basin 272 Skagerrak Graben 367 Urals foreland 429 seismic velocity and temperature 115 seismic, deep sounding Caucasus 278 Semail nappe 69 Semail Ocean, subduction 71 Semail ophiolite 235,244 serpentinite 123,243, 251 Urals 407, 423 Severnaya Zemlya archipelago 511, 512, 513, 515 shear wave velocity 6 Shelvian deformation 297 soft collision 301 Shemshak basin Iran 59 shield volcano 150 shoshonite 99, 181 Siberia 507, 508 Siberian terrane 86 Siberian traps 94 Sibumasu Terrane 87 Sicily intraplate magmatism 153, 155 transform fault 270 Silurian collision 328 Sitia microcontinent 73 Skagerrak Graben 363, 365, 366, 367 slab detachment 28, 29, 47-49, 52 Alpine arc 139, 143 Alpine 157, 167, 176, 181 Cenozoic rift 97, 101
INDEX
Iberia 51 magmatism 169, 175 Pannonian 52, 191,192, 196, 202 Permian 379 Rheno-Hercynian 99, 103 seismicity 170 Variscides 351,358 slab extrusion, Urals 423 slab roll-back 57, 59, 60, 62, 167, 171 see also roll-back slope stability 219 soft collision 291,293, 297, 301 Solenopleura [trilobite] 513 Sorgenfrei Line 47 Sorgenfrei-Tornquist Zone 291,301,300 South Hewett Lineament 297 South Portuguese Zone 351 Southern Pennine Basin 49, 355, 359, 363, 376 crustal thinning 102-104 Spain, Cenozoic volcanism 153, 155 spinel lherzolite (xenolith) 161 Sporades Trough 272 Srednogorie arc 61, 62-63 Strandja Massif 389, 391,395, 397 accretionary complex 398 stratovolcano 152, 177 strength and deformation 115-119 in rifts 121 strength lithosphere Pannonian-Carpathian 213-215 strength map 118, 208 strength models 116 stress field 123,214, 216, 218-219, 284 map 115 strike-slip faults 284 Pyrenees 370 Stromboli, volcanism 174, 175, 179 structure, Greater Caucasus 278 structure, Pannonian Basin 194 subduction 358 Cenozoic 29-30 Carboniferous 489 Palaeotethyan 69 subduction and conductive material 546 subduction and ophiolite emplacement 235, 236 active margin (Cordilleran-type) 247-253 mid-ocean ridge 237, 238, 240, 243 suprasubduction zone 237, 239-240, 245 subduction polarity 167,201,203,327, 349 Bohemia 298 Western and Eastern Alps 139-140, 143 subduction processes 167-203 active zones 170, 175 zones and migration 170 subduction relicts 553 subduction zones Avalonia 329 Calabrian-Hellenic arcs 264 Cyprus 247 European plate 134 Fennoscandia 585-589 Kazakhstan 26 Palaeoproterozoic 12, 15 Uralides 421-424 Variscides 338, 339 subduction zone, properties of 19 subduction-related magmas 178 subsidence curve 105, 106-107 Norwegian margin 122 Pannonian- Carpathian 211, 212 Pechora Basin 450, 453 subsidence modelling 105 Sudetian Phase 356 sulphide deposits 349, 351,634 as conductor 542, 545 suture 45, 99, 459 Belarus 605 Bohemian-Moldanubian 45 east Mediterranean 237, 238 Iberia 348, 351
Uralian 1,414, 415, 443 Variscides 333, 334 suture zones 356, 357 suture, East European Craton 299-300 Svecofennia, geological map 563 Svecofennian orogen 561-574 Svecofennian Province geophysical data 21-22 lithosphere thickness 24 SVEKALAPKO 313, 320, 521-536, 579 participants 555 Sweden, TOR project 314 Switzerland, Permo-Carboniferous basin 367 Tagil arc 422, 424-425 arc-continent collision 407-408, 410 crustal structure 414-415 platinum-bearing belt 421 seismic 437 Taimyr Terrane 510, 515, 516 geology map 511 Tajo Basin 226-227, 228, 231 Tauern Window 136, 140, 193, 200, 201 Tauric-Anatolian plate 64 Tauride carbonate platform 245, 246 Tauride ophiolite 247, 248, 249, 252 Tavas Nappe 69 tectonic evolution 11-33,488 tectonic map Anatolia 392, 393 east Mediterranean 236, 237, 266, 390 Europe 1, 12 Fennoscandia 580-581 Greater Caucasus 280 Urals 422 tectonic units 483 tectono stratigraphy Evia ophiolite 240, 241 Lycian ophiolite 245, 246, 247 Mersin ophiolite 245, 247 Oman ophiolite 236 Teisseyre-Tornquist Zone 27, 33,314 East European Craton 291,300 Tekirova ophiolite 251 temperature, mantle 12-14 terrane analysis 293-294 terrane defined 389 terrane reconstruction see plate reconstruction terranes 84, 498 Caledonide 1 East European Craton 291-293,602-606 Fennoscsandian Shield 564-565,568, 569 terranes, timing of break-up Vendian-Permian 84 TESZ see Trans-European Suture Zone Tethyan collision zone 154 Tethyan inversion 50 Tethyan Ocean 223,396-399 closure 400 evolution 65-75 Tethyan oceans 359 age of 398 reconstructions 6 0 - 6 7 Tethyan ophiolite 239, 248 Tethys rift system 59 Tethys suture 68, 277 thermal age 208, 215 thermal anomaly 156, 376 thermal data 16, 17 Uralides 26 thermal decay curve 99 thermal destabilization 101, 103, 107 thermal evolution and tiffing 120 thermal model 15, 16, 551 lithosphere thickness 24, 26, 27-28, 31 southern Europe 30 Variscides 25 thermal regime Bohemian Massif 337 Palaeoproterozoic terranes 603
661
thermal sag basin 49-50, 107 thermal structure 116 thermal subsidence intra-cratonic basins 44-45, 50 Pannonian Basin 195 Rotliegend 363 thermal subsidence and rifting 104-105 thermal thickness 30 thermal thinning 124 thermal-mechanical controls on Alpine deformation 113-123 thermo-mechanical model Pannonian- Carpathian system 207 - 219 Permian basins 377 thermotectonic age 118-119 thickness of lithosphere 12, 16, 20, 24, 156 Variscides 25 thickness, Variscan crust 103, 104 thinning of crust 44, 49-50 tholeiite 150, 153, 162, 174, 243 Archaean 630-634 tholeiitic basalt 172, 421 Thor Suture 313, 314, 317, 319, 320 East European Craton 297, 299-300, 301,304 Thor-Tornquist suture 45 thrust faults 284 thrust systems, Iberia 224, 226 Thulean flood basalt province 122 Timan Range 523, 524, 525, 530-531 seismic profile 533 Timanian Ocean 449 suture 452 Timanian Orogeny 3-4, 86, 532 Timanide Orogen 470, 507-516, 530 Timanides 1, 4, 5, 6 Tisza-Dacia terrane 193, 194, 200, 202 extrusion tectonics 202 tomography 6, 11, 13, 16 Fennoscandia experiment 549-554 high-resolution teleseismic 129-143 lower lithosphere structure 137-138 mantle model 136, 137 thermo-mechanical controls 119 see also TOR project topographic map East European Craton 483 Europe 114, 148 southeast 483 Greater Caucasus 278 Iberian Peninsula 228 Netherlands 120 topography and tectonics 337 Alpine deformation 113, 114, 123 Greater Caucasus 284 Iberian Peninsula 223, 227-231 Pannonian Basin 215, 216 Variscan 338 TOR project 313-319, 320 earthquakes 314, 315, 319 P-wave travel time residuals 316 intraplate magmatism 156 lithosphere thickness 24, 31 seismic experiment 551, 552 seismic, Pannonian Basin 200-202 seismograph location 314 subduction and magmatism 171, 176, 177, 180 tomography experiment 18, 32-33 Tornquist Ocean 83, 88, 89 Tornquist Sea 293,296 closure 299, 301 Tornquist Zone 325, 358 inversion 319 subduction 326-327 TOR data 314, 316, 317, 318, 320 Tornquist-Sorgenfrei Line 49 Tornquist-Teisseyre Zone 377, 379 trace element ratio 158, 160 trace elements, data sources 157
662
TRANSALP transect, Alps 117, 131, 136, 138-143 seismic profile 130, 132, 135 working group 29, 200, 201 transects Adriatic microplate 136 Alps 131, 132, 141, 142 Avalonia- Rhenohercynian 302 Caledonide 304 Carpathian-Libyan 74 Danube Basin 198 Dniepr-Donets Basin 466, 468 East European Craton 301, 302, 303 Fennoscandian Shield 567, 573, 582, 617, 618 German Variscides 305, 335 Greater Caucasus 279, 280 Iberian Massif 346, 347, 350 Lod~ve Basin 371 Massif Central 369 Montagne Noire 369 North Sea 304 Pechora Basin 472, 510, 534 Peri-Caspian Basin 469 Polish Trough 365 Sarmatia 617 southern France, Permian Basin 371 Svecofennia 567 Timan Range 534 Urals 410, 412, 413, 415, 422 Variscan 303, 305 Trans-European Suture Zone Baltica 516 East European Craton 291 Trans-European Suture Zone 11, 83, 364 defined 1-2 location 5, 7, 12 dispersal, peri-Tethyan 299 gravity 17 seismic velocity 18 crustal thickness 18, 22 crystalline crust 375 lithosphere properties 18-24 lithosphere, deep 313,319, 362 Palaeozoic accretion 323 transform fault, Strait of Sicily 270 transpression 6 Variscan 345-351 Transylvanian Basin 202, 215- 216 Cenozoic volcanism 152 trapdoor basin 49, 101 trench-passive margin collision 247 Triassic basin, isopach map 105 trilete spores 93 trilobite provinciality 89, 298 trilobites 328, 513 Troodos ophiolites 235, 247, 252, 253 rotation 251 Tulcea Terrane 298 Turkey 389-410 Tyrrhenian Basin 272 Ukrainian Shield 11,463 ultramafic complex 443-447 unconformity Bathonian 281 Kan'on River 513 Oligo-Miocene 283 Pechora Basin 450 Vendian 515 underplating 45, 574 magmatism 584 Upper Rhine Graben 51, 52, 97, 103, 107 Cenozoic tectonism 114, 120 Urach province 151, 159 Uralian Ocean 515, 516 subduction 509 Uralian orogeny 86, 93, 476
INDEX
Uralide orogen 444, 507 Uralides 4, 6, 24, 26-27 Maksyutov Complex 421-422 subduction and collision 421-425 Uralides tectonism 409-417 arc-continent collision 407-409 crustal features 434-436 folds and thrust belt 411-413,416-417 granitoid emplacement 411,413-415 ocean closure 414-416, 443-447 seismic profile 439 strike-slip faulting 411,413-416 subduction 409-411 Urals 26 geophysical data 21 lithosphere thickness 24 Urals foreland basin 533 Urals seismic reflection survey 427-440 crustal root 439 crustal structure 433-438 mantle reflectors 438-439 URSEIS seismic survey 407, 408, 411,421,422 interpretation 414-416 seismic profile 20, 26 Valais ocean (Alpine Tethys) 129 Valence graben 114 Valencia Trough 226, 227, 230, 231 Vardar 64-66 Vardar Ocean 170, 248, 253 subduction 59, 61-63, 65, 67, 75 Mesozoic-Cenozoic 176 volcanism 247, 248 Vardar zone volcanism 247-253 variation diagrams 158-160, 161 Carpathian-Pannonian region 173, 180 Variscan cycle 43, 45, 57 Variscan Deformation Front 98, 100, 356, 363, 375, 376 Variscan foreland 305 Variscan orogen 97, 333-340 Cenozoic Rift System 97-101, 104 Balkans and Black Sea 399-400 Iberia 345-351 Variscan orogeny, 277, 325, 327, 355-359 East European Craton 297 compression rate 357 final stage 372 fracture system 359 lithosphere 379 magmatism 359-362 Massif Central 367 molasse 335, 336, 400 plate configuration 359 reactivation 103 suture 357 tectonic map 356 unroofing 101, 102 Variscan plate model 334, 335, 339 Variscan terranes boundaries 102 collision chronology 348-350 cross-section 350 depositional history 348 Variscan thrusts, Germany 303 Variscides (Hercynian) 20, 93, 333, 334, 337 development 3-6 geophysical data 21 Himalayan-type and Andean-type 101 location map 1 lithosphere thickness 24, 19-28 seismic model 25 thermal model 26 Variscides orogen 97 terrane 100 VARNET seismic profile 24 Velay, Cenozoic volcanism 152 velocity 531-532 TOR data 313-319, 320
velocity indicating partial melting 156 velocity model EUROBRIDGE 612 Fennoscandia 549-550, 553, 554, 555 velocity variation, P-wave 137, 138, 140, 142 Veneto province 153, 154 vent, volcanic 154 Veporic nappe 75 Vestfold Graben 365, 366 Vesuvius, volcanism 175 Viking Graben 30 micro seismic activity 379 Vivarais, Cenozoic volcanism 152 Vogelsberg volcanic complex 150, 155, 162 volcanic arc 168 volcanic fields 150-155 in Europe 148, 149, 151, 168 Greater Caucasus 280-282, 285 Iberia 230, 231, 360, 372 Mid-German Crystalline High 327, 329 Norwegian margin 122 Oman 240 Polish Basin 364 Vardar zone 247 volcanicity 349-351 Cenozoic 154-155 chronology 150 Permian 365, 376 Tremadoc 513 Permo-Carboniferous 102 Variscan 335, 336, 360-361,362 volcanoes, active 153, 174, 175,280 Volgo-Uralia 7 Vosges magmatism 148, 151,338 Vosges-Black Forest arch 52, 53, 97 Voykar massif 443,447 Vrancea slab 196, 202 Vulcano, volcanism 174 Vulsini, mantle source contamination 179 wedge structures, lower crust 134-135 wedges, orogenic 62 Wessex Basin 50 West Ny Friesland Terrane 507 Westerwald volcanic activity 150, 155 Whin Sill Complex 361,381 Wilson cycle 277, 584 wrench faults 268, 380, 416 Carboniferous-Permian 45, 48 Cenozoic Rift System 97, 99, 100-102 map of Europe 48, 49 xenolith 12, 13, 15, 381,590 anorthosite 302 Cenozoic Rift System 103 crustal 361,380 garnet 196 lherzolite 25, 161,553 mantle 31, 156, 361,551,554 ultramafic 172, 176 Xystridura [trilobite] 513 Zagrab Fault 193, line 201 Zechstein salt 365, 375, 379 Zechstein Sea 59, 94 zircon age 326, 327, 333,336-338, 362 Anatolia 390, 392, 395, 396 Archaean 564, 632, 635 Dnestr 457 East European Craton 292, 293, 296, 297 Fennoscandia 583, 586, 587, 588 Proterozoic 532 see also isotope age zircon analysis, U-Pb 443-447 zircon, detrital (Palaeozoic) 513 zircon, U-Pb isotope data 445, 446 zircons, inherited 380