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Post-Variscan (end Carboniferous ­Early Permian) basin evolution in Western and Central Europe

´ ´ T. MC CANN1, C. PASCAL2,3, M. J. TIMMERMAN4, P. KRZYWIEC5, J. LOPEZ-GOMEZ6, A. WETZEL7, 8 9 10 C. M. KRAWCZYK , H. RIEKE & J. LAMARCHE 1 Geologisches Institut, Bonn University, Nußallee 8, 53115 Bonn, Germany (e-mail: [email protected]) 2 Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands 3 Present address: NGU, Geological Survey of Norway, N-7491 Trondheim, Norway 4 ¨ ¨ Institut fur Geowissenschaften, Universitat Potsdam, Karl-Liebknecht-Strasse 24, 14476 Potsdam, Germany 5 Polish Geological Institute, ul. Rakowiecvka 4, 00-975 Warsaw, Poland 6 ´ ´ ´ Instituto de Geologia Economica, CSIC-UCM Facultad de Geologia, 28040-Madrid, Spain 7 ¨ Geologisch-Palaontologisches Institut, University of Basel, Bernoullistrasse 32, CH-4046, Basel, Switzerland 8 Geoforschungszentrum, Telegrafenberg, 14473 Potsdam, Germany 9 PanTerra Geoconsultants B.V., Weversbaan 1 ­ 3, 2352 BZ Leiderdorp, Netherlands 10 ´ ´ ´ ´ Universite de Provence, Unite CNRS Geologie des Carbonates, Place Victor Hugo, Case 67, 13331 Marseille, France

Abstract: The Variscan orogeny, resulting from the collision of Laurussia with Gondwana to form the supercontinent of Pangaea, was followed by 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 (e.g. NE German Basin, NW part of the Polish Basin, Oslo Rift, northern Spain). Coeval transtensional activity led to the formation of more than 70 rift basins across the region. The various basins differ in terms of their form and infill according to their position relative to the Variscan orogen (i.e. internide or externide location) and to the controls that acted on basin development (e.g. basement structure configuration). This paper provides an overview of a variety of basin types, to more fully explore the controls upon the tectonomagmatic ­sedimentary evolution of these important basins.

The Permo-Carboniferous collision of the continents of Gondwana and Laurussia, termed the Variscan orogeny, led to their amalgamation and the formation of the late Palaeozoic supercontinent of Pangaea. Oblique convergence resulted in collisional processes in the Appalachians and the Urals, whereas sinistral wrench faulting caused widespread rifting of the Northern European crust (Pegrum 1984a,b; Ziegler 1990). In addition, the collapse of the thickened Variscan orogenic crust resulted in late orogenic crustal extension. The post-Variscan period was one of intense crustal re-equilibration and reorganization under an alternating transtensional and transpressional tectonic regime, and the combined effect of re-equilibration and tectonic activity controlled the kinematic patterns and subsidence of approximately 70 basins, all of which are characterized by a major strike-slip component in their deformational history. Western and Central Europe was already thermally weakened by the preceding orogeny and most of the Carboniferous­ Permian basins trace long-lived Variscan fault systems (Henk 1993). Subsequent basin evolution involved extensive, predominantly clastic sedimentation (e.g. Glennie 1990; Maynard et al. 1997), and some of the newly formed rifts became the loci of extensive intraplate magmatism (e.g. Neumann et al. 2004). In recent times, there have been a number of volumes published on the Variscan history of Europe (e.g. Dallmeyer et al. 1995; Franke et al. 2000). Although these books provide much-needed information concerning the nature and style of deformation during the Variscan orogeny, they are not really concerned with the postVariscan evolution of the region. One of the main problems with unravelling the post-Variscan deformational history of Western Europe is the localized nature of deformation. Small isolated grabens were gradually filled by predominantly locally derived sediments and/or associated volcanic and volcaniclastic rocks. The nature of the sedimentation (predominantly alluvial) constitutes another problem, as correlation is much more difficult, both within and, more importantly, between grabens. The nature of the

sedimentary record within the basins that formed as a result of Permo-Carboniferous wrench-fault activity, therefore, frequently precludes detailed investigation, as it is difficult to correlate the sedimentary record from one basin to another. In those Early Permian basins that are rich in biogenic remains it may be possible to correlate within the basin (e.g. Schneider 1989, 1996). However, the general level of biostratigraphic uncertainty leads to problems with correlation in this stratigraphic interval. It is only when basinwide crustal subsidence gave rise to the widespread Northern and Southern Permian basins, with the establishment of a unified depositional pattern across Northern and Central Europe, that areally more extensive correlation becomes possible (e.g. McCann 1998a). There are, however, many Permo-Carboniferous basins that are outside the Southern and Northern Permian basins region, and such basins provide much-needed understanding of the broader evolution of the Permo-Carboniferous transition in terms of magmatic, tectonic and sedimentary activity outside the area of the former Carboniferous foreland basin. The following summary, by nature selective, reviews the evidence for basin formation at this time and investigates a selection of these basins. Of marked interest is the similarity in terms of basin infill (sedimentary and magmatic) between the basins, despite differences in both timing and location. The aim of this study, therefore, is to provide an overview of the main areas of basin initiation that followed the cessation of Variscan orogenic activity in Western Europe. The following sections will outline the regional framework of Western Europe and provide an introduction to the most recent research in these areas, including sedimentology, tectonics, magmatic history and basin modelling.

Background geology and palaeogeographical setting

The Variscan belt is a broad (c. 1000 km) complex curvilinear feature extending across Europe and marking the zones of Variscan-age deformation (Fig. 1). Variscan orogenic activity was

From: GEE , D. G. & STEPHENSON , R. A. (eds) 2006. European Lithosphere Dynamics. Geological Society, London, Memoirs, 32, 355­388. 0435-4052/06/$15.00 # The Geological Society of London 2006.

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Fig. 1. The main tectonic elements of Central and Western Europe (modified after Ziegler 1990; Berthelsen 1992). Inset map shows the tectonic zones corresponding to volcanic belts: MZ, Moldanubian Zone; OMZ, Ossa­ Morena Zone; RHZ, Rheno-Hercynian Zone; SPZ, South Portuguese Zone; STZ, Saxo-Thuringian Zone; TTZ, Tornquist­Teisseyre Zone. Massifs: a, Armorican; b, Bohemian; bf, Black Forest; c, Massif Central; h, Harz Mountains; ic, Iberian Cordillera; rh, Rheinische Schiefergebirge; v, Vosges (after Francis 1988).

associated with the convergence of the southern continent of Gondwana with the northern continent of Laurussia (i.e. Laurentia, Baltica and Avalonia), to form the supercontinent of Pangaea and leaving a relict Palaeotethys to the east (Scotese & Langford 1995; Fig. 2). The reconstructed geometry of the Variscan orogen can be divided into a number of distinct geotectonic zones, many of which are separated by steeply dipping faults or shear zones, that exhibit a general continuity around the belt. Some of the boundaries may represent suture zones and these are marked by the occurrence of ophiolite assemblages and calc-alkaline continental arc volcanic rocks of Devonian or Carboniferous age. The high-grade core of the Variscan orogen runs through central and NW Spain, France, Germany and the Bohemian area of the Czech Republic (Figs 1 and 3). The Devonian to Early Carboniferous evolution of the Variscan orogenic system was governed by alternating tensional and compressional tectonic cycles, reflecting the development of an essentially intra-continental Pacific-type, back-arc system (Ziegler 1990). Four principal phases of deformation and exhumation may be recognized, each of which lasted between 20 and 39 Ma and was restricted in geographical extent. These phases probably resulted from the successive docking of continental lithospheric fragments along the southern margin of Laurussia. The first of these periods, the Ligerian Phase (Late Silurian­Early Devonian), was contemporaneous, but not cogenetic, with the Acadian phase of the Caledonian orogeny. The remaining three phases, Bretonian (Late Devonian­Early Carbon´ iferous), Sudetian (Visean­Early Namurian) and Asturian (Westphalian­Early Permian), are assigned to the Variscan orogeny (Warr 2000). ´ During the late Visean the Variscan orogenic system entered the Himalayan-type continent­ continent collision stage. Continued dextral-oblique convergence of Gondwana and Laurussia and the progressive closure of the Proto-Atlantic and western

Proto-Tethys oceans was accompanied by the propagation of their collision front into the Appalachian and central Mediterranean domains. At the same time, major crustal shortening was achieved in the Variscan fold belt as indicated by the occurrence of major nappe structures, in part involving basement in the Moldanubian area, the southern Massif Central and the Variscan externides (Behr et al. 1982; Burg et al. 1984; Cazes et al. 1986; Behr & ´ Heinrichs 1987). During the late Visean and Namurian, the collisional front between Africa and the Southern European margin propagated rapidly westwards and eastwards. Following the Late Westphalian ­Early Stephanian consolidation of the Variscan fold belt, convergence between Gondwana and Laurussia apparently changed from an essentially north­ south-directed collision to an east­ west convergence. During Late Carboniferous times a subduction zone developed along the Variscan Deformation Front, with oceanic crust being subducted beneath the continental Variscan system (Gast 1988). Destruction of the subducted crust led to stretching (dominantly east ­ west) and associated block faulting. Coeval counter-clockwise rotation of the southern African plate against the stable northern European craton caused wrenching along NW­ SE-trending faults (Ziegler 1990). Continental-scale dextral shears (e.g. the Tornquist­ Teisseyre fracture zone) were linked by secondary sinistral and dextral shear systems (Ziegler et al. 2006). The westward translation of Gondwana relative to Laurussia along a 3500 km long dextral megashear has been proposed, but recent work from northern Spain does not support the existence of such a major structure (Weil et al. 2001). It is, however, possible that a series of smaller shear structures were active (see the following section). The development of the Variscan orogen involved major crustal shortening and subduction of substantial amounts of supra-crustal rocks, continental and oceanic crust, and mantle-lithosphere

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From late Early Carboniferous times, compression within Central Europe was continuous for c. 45 Ma, with Variscan deformation advancing northward at a rate of c. 0.5 cm a21 (Ahrendt et al. 1983). This is only slightly less than the value of 0.79 cm a21 that characterized Alpine orogenic compression (Schmid et al. 1996). In western Germany, for example, a series of major north- and southverging, basement-cored nappes were emplaced during Namurian and Westphalian times. The widespread occurrence of low-pressure metamorphic rocks and late to post-orogenic calc-alkaline intrusive rocks in the internal parts of the Variscan orogen suggests that a significant amount of crustal shortening, accompanied by crustal delamination, subduction and anatectic remobilization of lower crustal material, and partial melting of upper mantle material, occurred (Ziegler 1984, 1986). Rough estimates of total crustal shortening yield a value of c. 150­200 km (Ziegler 1990). This is less than half of the crustal shortening recorded from the Alpine Realm during the Tertiary, which was estimated at c. 500 km (Dewey et al. 1989; Schmid et al. 1996). By latest Westphalian times, the Variscan fold-and-thrust belt of Western and Central Europe was consolidated and inactive (Ziegler 1990). However, localized tectonic activity is evidenced by diverse angular unconformities between the Westphalian and Permian deposits (e.g. NE German Basin; McCann 1999). In Germany, post-orogenic uplift of the Rheno-Hercynian thrust belt, preceding deposition of the Permian sediments, amounted along its northern margin to 2­3 km, increasing southwards to 6 km (Littke et al. 2000), with values of up to 10 km in the area of the Saar­Nahe Basin (Oncken et al. 2000).

The eastwards extent of Variscan deformation

As noted above, the collision between Gondwana and Laurussia continued to develop until late in the Carboniferous­Early Permian, at which time the intercontinental collision began to affect the northwestern part of Africa (e.g. the West African ´ ´ orogens have a Late Carboniferous to Early Permian age; Lecorche et al. 1989). As a result of the extensive collision episode, a central Pangaean mountain range was formed, extending from Mexico to Poland (Golanka & Ford 2000) and southwards to Morocco (Pique & Michard 1989). Late Carboniferous events were also marked in the Alps (e.g. Matter et al. 1987), the Carpathians (Dallmeyer et al. 1995) and the Rhodope area (Yanev 1992). Further eastwards, the situation was a more complex one, as it involved the rifting (from Late Carboniferous times onward) from Gondwana of several Cimmerian continents ¨ (e.g. parts of Turkey, Iran and Afghanistan; Sengor et al. 1984; Scotese & Langford 1995; Fig. 4). During the main phase of the ´ Variscan orogeny (late Visean to Westphalian), the collision front between Gondwana and Laurussia propagated eastward and southwestward in conjunction with the progressive closure of the Palaeotethys and Protoatlantic oceans. By Westphalian times, Gondwana had collided with North America whereas to the east Palaeotethys remained open. Therefore, the western parts of the Variscan megasuture were characterized by a Himalayan-type (continent­ continent collision) setting, whereas its eastern parts remained in an Andean-type (continent­ ocean collision) setting (Stampfli 2000; Ziegler & Stampfli 2001). The setting of the West European segment of the Variscan orogen was, therefore, transitional between a Himalayan- and an Andean-type setting (see Ziegler et al. 2006, for details). During the Stephanian and Autunian orogenic movements continued in the Appalachian ­ Mauretanides (Rodgers 1970; Michard & Sougy 1977) and in the Urals (Ivanov et al. 1977), whereas the Variscides region remained largely inactive. These latter movements were accompanied by the emplacement of a right-lateral transform fault system, which linked the southern Uralides and northern Appalachians, and which crossed Europe, where it caused the development of a complex pattern of conjugate shear faults and related pull-apart structures (Arthaud & Matte 1977; Ziegler 1978a).

Fig. 2. Early (a) and Late (b) Permian palaeogeography (after Scotese & Langford 1995).

(Ziegler et al. 1995). The convergence rate between Gondwana and Laurussia, however, was not constant during Devonian and earliest Carboniferous times, as suggested by the Early Devonian development of an extensive back-arc rift system in Western and Central Europe, which remained intermittently active until the late ´ Visean (Zieger 1990). Full-scale collision between Gondwana and the southern extension of Laurussia occurred during the late ´ Visean at a time when the back-arc extensional system that had governed the early evolution of the Variscan orogen in Western and Central Europe finally became replaced by a compressional stress regime (Ziegler 1990). Gondwana and Laurussia became increasingly coupled during the Early Carboniferous, as is evident from their joint clockwise rotation and northward drift (see Ziegler 1990, Fig. 8).

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Fig. 3. Structural sketch map of the European Variscan orogen. Three external coal basins are shown in black. C.C.S., Coimbra­Cordoba suture; L.R.S. Lizard­Rhenish suture; N.V.F., North Variscan Foreland; M.C.S., Massif Central ¨ suture; M.T.S., Munchberg­Tepla suture; O.M.S., Ossa­ Morena suture. The Iberia and Corsica­Sardinian blocks are represented in their possible Permian positions relative to Europe (after Matte 1991).

An important problem is the continuation of the Variscan orogen in SE Europe and Turkey. In Western Europe, subduction ended with the collision of NW Gondwana (i.e. Morocco, Algeria) with SW Europe (i.e. Iberia, Corso-Sardic block). Further to the east (i.e. Tunisia, Libya, Egypt) there is no evidence of subduction-related activity (e.g. intense folding, subductionrelated magmatism) at the northern margin of Gondwana during the Late Palaeozoic (Dornsiepen et al. 2001). In contrast, orogenic activity can be traced along the southern margin of Laurussia from the Alps to the Caucasus, although it appears that the closure of Palaeotethys was asymmetrical, with subduction occurring on the northern (i.e Laurussian) margin (Vavassis et al. 2000). It is probable that the active spreading ridge of Palaeotethys arrived at the southern margin of Laurussia around the Permo-Carboniferous boundary, leading to a cessation in active subduction, and the convergent margin was then transformed into a dextral strike-slip movement zone (Dornsiepen et al. 2001; Fig. 4). Stampfli et al. (2001) noted that there is a general pattern of rifting across the region of Tethys. The dynamic evolution of the region of Neo-Tethys (India, Arabia, Turkey) reveals that the Early Carboniferous rifting phase was followed by a second rifting phase in the Early Permian and the late Early Permian­ Late Triassic opening of Neotethys. A similar pattern of rift activity is recorded in the East Mediterranean area, where the initial Carboniferous rifting phase affected the southern margin of the area between Syria and Tunisia. This was followed by a Permian rifting phase and protracted Late Permian subsidence (extending to the present day). Following late-stage subduction of Palaeotethys, and in conjunction with slab roll-back (since Early Permian times) and slab detachment beneath the Variscan orogenic belt, the Cimmerian Terrane was detached from the Gondwana margin and the Mediterranean­ Neotethys Basin began to open along the Eurussian margin. The post-collisional Permo-Carboniferous rifting phase accompanied the transtensional collapse of the overthickened Variscan lithosphere in a basin-and-range fashion (Malavielle 1993). This analogy implies the presence of a still-active or transform margin in areas where the Palaeotethys was not yet closed during the Permian. The mid-ocean ridge of

the Palaeotethys was obliquely subducted beneath the active Eurasian margin during the Permo-Carboniferous (Stampfli 1996). This was responsible for the widespread Late Carboniferous and Early Permian magmatism and volcanism characterizing the southern Variscan domain.

Late to post-Variscan structures

During the main phase of the Variscan orogeny, lateral escape tectonics was accompanied by the development of synorogenic transtensional basins in the Armorica­ Iberia­ Avalonia domain. During the latest Carboniferous and Early Permian the convergence of Gondwana and Laurussia changed from oblique collision to dextral translation controlling the Alleghenian orogeny of the Appalachians. At the same time the Variscan orogen was transected by conjugate wrench faults, partly terminating in pull-apart basins (Ziegler 1990; Matte 1991). Along the northern margin of Gondwana, rifting resumed during the Early Carboniferous, reflecting an intraplate stress reorganization following the docking of the composite Hun terrane to the Laurussian margin (see Stampfli et al. 2001, for details). A number of major late to post-Variscan tectonic structures have been recognized. The Tornquist Zone, comprising a fan-shaped series of fault and fault-zone splays was activated in the western Baltic area (Berthelsen 1992). Indeed, right-lateral wrench movements along the Tornquist­ Teisseyre lineament led to the formation of the Oslo ­ Skagerrak Graben system (Ziegler 1978b; Neumann et al. 1992). According to Ziegler (1990), a major postVariscan wrench fault was active along the present Elbe Fault System (northern Germany) as part of a complex megashear zone that developed in North ­ Central Europe between the Appalachians and Uralides as a result of dextral translation of the Armorican ­ European plate with respect to the African plate (Arthaud & Matte 1977). Along the SE end of this wrench fault, geological evidence for Late Carboniferous-age dextral ductile shearing has been found in outcrops of the Elbe Zone (Mattern 1996). Scheck et al. (2002) have noted that maximum deformation occurred along the Elbe Fault System during late Carboniferous

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Teisseyre lineament. Displacements along the Bay of Biscay fault system were in part taken up in the Arctic ­North Atlantic rift, in which crustal distension continued during the latest Palaeozoic, as illustrated by the East Greenland graben system (Haller 1971). Indeed, 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).

Palaeogeography and palaeoenvironments

Lithologically, the Upper Carboniferous successions across the region are relatively monotonous. This is partly due to the similarity in terms of lithological and stratigraphical evolution between the Variscan fold-and-thrust belt and that of its foreland basins, related both to the infilling of the basins and the gradual cessation of Variscan-age tectonic activity. Only the Namurian A still shows the orogen ­foreland differentiation. During Late Carboniferous times, paralic coalfields extended across Laurussia from North America across Britain and the Rheno-Hercynian region of Germany to Silesia and the Donets region, Ukraine. Coal-forming environments were present south of the Rheic suture in the Saar ­ Lorraine Basin during late Westphalian times (possibly earlier) and continued until early Stephanian times. After a short hiatus, these conditions were re-established in mid-Stephanian times (Cleal 1984). This extensive region experienced several marine incursions and clearly most of the Westphalian deposition (largely non-marine) occurred close to sea level (Paproth 1991) on the northern (passive margin) side of the Rheic suture. In contrast, to the south of the suture mainly intermontane limnic basins (of Stephanian age) developed on the various components of the Armorican Terrane Assemblage (McKerrow et al. 2002). During the Permian, continental size and aggradation were at a maximum, as few continental fragments escaped incorporation into Pangaea. The new continent, stretching almost from pole to pole, had a significant amount of terrestrial area. A single world ocean, Panthalassa, with a semi-enclosed Tethyan Sea, dominated the marine environment. This singular continental configuration led to the development of extreme climatic conditions. Furthermore, the Permian climate was not a stable one, with evidence of profound changes throughout the duration of the period (Veevers & Powell 1987; Barron & Fawcett 1995; Parrish 1995). Climate was not the only unstable factor. Given the sheer size of Pangaea, the supercontinent was unstable from the outset. The timing of break-up was varied, beginning within the late Carboniferous immediately north of the Variscan orogen. In many places (e.g. Germany and Belgium), extension to the north of the Variscides occurred coevally with thrusting and strikeslip faulting further south (Ziegler 1990). However, the timing and extent of individual phases of extension and rifting throughout the North Atlantic (and associated) rift systems are still subject to some debate, as the dating of Late Carboniferous and Early Permian red beds is imprecise. Following the main phases of Variscan compression, thermal relaxation of the crust occurred in Early Permian times, creating the rifts and graben that allowed accumulation of the first phase of sedimentation.

Fig. 4. Schematic plate configuration and plate boundary geometry between Europe and Gondwana. (a) Early Permian. (b) Mid­Late Permian. (c) Mid­Late Triassic. (After Dornsiepen et al. 2001.)

wrenching and volcanism (and subsequent to these events). Late Palaeozoic dextral transtensional movements with estimated displacements of 60­ 120 km have been postulated. The major Variscan faults in Germany and Poland (i.e. Elbe Zone, Main Intra-Sudetic Fault, Odra Fault Zone) are considered to have been dextral shears trending NW­ SE during the late Devonian­ early Carboniferous with maximum offsets of 50 ­ 300 km (Aleksandrowski 1995; Alexandrowski et al. 1997) and dextral offsets of up to 400 km along the Teisseyre ­ Tornquist Zone (Lewandowski 1993). In late Carboniferous­early Permian times, many local and regional WNW ­ESE- to NW ­SE-oriented fault or shear zones in the Western Sudetes were reactivated under a semi-brittle regime. The main elements of the Late Variscan fracture system of northern Africa and Europe are the Agadir­ Kelvyn, Gibraltar and Bay of Biscay fault systems, as well as the Tornquist­

Stephanian­ Autunian magmatic activity

During the Permo-Carboniferous, the Variscan foreland in Europe was subjected to extensive rift-related tectonism and related extensional magmatic activity (Fig. 5). The rare occurrence of Autunian-age sediments suggests that much of the area later occupied by the Southern Permian Basin was regionally uplifted and subjected to profound erosion. This uplift, probably induced

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Fig. 5. Distribution of late Carboniferous and early Permian magmatic rocks in NW Europe (after Ziegler 1990). NGB, North German Basin; OS, Oslo Rift; WSC, Whin Sill Complex; VF, Variscan Front; RFH, Ringkøbing­Fyn High; OR, Oslo Rift; MV, Midland Valley.

by a combination of wrench-related lithospheric deformation, magmatic inflation of the lithosphere and thermal erosion of the mantle lithosphere, was coupled with periods of significant magmatic activity. Melt generation was probably related to localized divergent wrench-induced decompressional partial melting of the uppermost aesthenosphere and the lithospheric thermal boundary layer, possibly combined with the impingement of a not very active mantle plume on the base of the lithosphere (Ziegler 1996). Although the most voluminous and evolved magmatism occurred in the North German Basin (Benek et al. 1996), dated at 297­ 302 +3 Ma (Breitkreuz & Kennedy 1999), the Oslo Rift contains the most extensive and best-preserved sequences of basaltic lavas associated with this event. In the latter, Rb ­Sr age determinations indicate that the total period of magmatic activity extended from c. 305 to 245 Ma (Sundvoll et al. 1990), and the earliest basaltic magmatism appears to have been restricted to a relatively short period (305­ 290 Ma) (Neumann et al. 2002). However, recent dating suggests that these ages are slightly too young and that the duration of the various magmatic periods in the Oslo Graben could be shorter than proposed by Sundvoll et al. (1990) (see Neumann et al. 2004). The early basalts from the Oslo Graben contain the least evolved magmatic products in the region and provide information about the primary magmas and the magma sources involved in the magmatic activity. Volcanic activity, however, was widespread throughout Europe, with thick sequences being deposited elsewhere; for example, in Spain (e.g. Iberian Range, Lago & Pocovi 1984), dated at ´ 282 + 12 Ma (Hernando et al. 1980), and in the Italian Alps (e.g. Collio Basin, Breitkreuz et al. 2001), dated at 283 + 1 Ma and 281 + 2 Ma (Schaltegger & Brack 1999), Liguria and Sardinia (e.g. Cortesogno et al. 1998). Autunian-age magmatic activity also included the intrusion of the Whin Sill dolerite (northern England), some dolerite dyke swarms in the Midland Valley of Scotland and the Northumberland Basin (Francis 1978; Coward 1995), and in the Argyll area of Scotland

(Speight & Mitchell 1979), and episyenite dykes along tensional ´ structures in the Central Range of Iberia (Gonzalez-Casado et al. 1996). In these areas volcanism followed Westphalian-age inversion and predated much of the Permo-Triassic succession, suggesting that the region was underlain by hot asthenosphere. The widespread Stephanian­ Early Permian (305­ 285 Ma) alkaline intrusive and extrusive magmatism of the Variscan area and its northern foreland is mantle derived and shows evidence of strong crustal contamination (e.g. Marx et al. 1995; Benek et al. 1996; Breitkreuz & Kennedy 1999). Melt generation by partial melting of the uppermost aesthenosphere and the lithospheric thermal boundary layer was probably triggered by a rise in the potential temperature of the aesthenosphere and localized divergent wrench-induced decompression. Aesthenospheric upwelling was presumably triggered by detachment of subducted lithospheric slabs, resulting in a reorganization of the mantle convection system and the impingement of a not very active system of mantle plumes on the base of the lithosphere. There are at least two discrete tectonic settings for sedimentary basins that formed simultaneously during the Late Palaeozoic Variscan orogeny in Western Europe. Immediately adjacent to, and to the north of, the Variscan deformation front, narrow foredeep basins, termed the external basins (Variscan externides), are interpreted to have formed as a result of subsidence related to thrust loading of the crust. The related internal basins (Variscan internides) comprise those basins formed within the orogenic belt and that were structurally controlled. The following discussion of the magmatic activity within the Variscan-influenced regions of Europe will examine both basin types.

Extrusive activity (Variscan foreland­ externides)

In the foreland of the Variscan region, sedimentation and magmatism generally occurred in and close to the north- and NW-trending grabens. Widespread magmatism occurred as

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felsic to mafic volcanism, and emplacement of mafic dykes and sills. In the Oslo Graben and Central North Sea, extension, block faulting, tilting, uplift and erosion of structural highs occurred simultaneously with, and subsequent to, the initial phase of volcanic activity. As noted above, the most voluminous magmatic activity occurred in the Oslo Rift, with more than 100 000 km3 of produced melts (Neumann et al. 2004), and the NE German Basin, where more than 48 000 km3 of magmatic material was extruded. Five stages of magmatism have been recognized and dated in the Oslo Graben (Olaussen et al. 1994; Neumann et al. 2004). The pre-rift stage (i.e. between 304 and 294 Ma) comprises felsic sill emplacement in underlying Westphalian sediments. The second stage corresponds to rift initiation and is characterized by widespread basaltic volcanism (i.e. B1 basalts) representing the most primitive magmatic rocks in the Oslo Graben. The main rift stage (i.e. Stage 3) resulted in the fissure eruption of thick sequences of porphyritic trachyandesite flows (rhomb porphyries), which was accompanied by the intrusion of large amounts of syenitic magmas that have yielded ages of 292­ 298 Ma according to recent U­ Pb dating (Neumann et al. 2004). During Stage 4 rifting abated and the magmatic style changed to the development of central volcanoes and caldera collapse. Rb ­ Sr dating of associated magmatic rocks gives ages between 280 and 240 Ma (Sundvoll et al. 1990). The final magmatic stage was dominated by the intrusion of composite batholiths of trachyandesite to rhyolitic compositions. In the subsurface of Northern Germany and Poland extensive Late Carboniferous to Early Permian felsic to intermediate volcanic rocks have been cored by exploration wells (e.g. Hoth et al. 1993). In this region, centres of volcanic activity appear to coincide with the intersection of fault systems (Plein 1978) and are probably related to pull-apart structures at the termination of subsidiary wrench faults that parallel the Tornquist­ Teisseyre lineament. In Northern Germany the magmatic succession is up to 2000 m thick, and comprises predominantly calc-alkaline SiO2-rich volcanic rocks (Kramer 1977; Marx et al. 1995; Benek et al. 1996). The volcanic and pyroclastic succession of Northern Poland is broadly similar to that of NE Germany and comprises mainly acid volcanic rocks (rhyolites, rhyodacites) and locally abundant trachytes and trachybasalts (Pokarski 1988, 1989). In the northern part of the Fore-Sudetic Monocline, there is an Early Rotliegend-age volcanic succession composed of lava flows, pyroclastic deposits, and hypabyssal and intrusive rocks (Jakowicz 1994). The rocks, however, are not comagmatic, as they were derived from two sources, the mantle and crust. The parent magma was a high-aluminium basalt, but anatectic crustal fusion produced acid melts that mixed with the derivatives of the parent magma to produce poorly homogenized hybrids with intermediate compositions. The distribution of these rocks is parallel to the trend of regional tectonic units of Early Rotliegend age, and their thickness varies from a few metres to over 1500 m. In general, the volcanic rocks of the Polish Basin are limited to its western margin and overall only attain a thickness of c. 100­ 200 m (with exceptions, see above), and are, thus, an order of magnitude less than in the adjacent NE German Basin (Karnkowski 1999). In the offshore regions of Northern Europe it is more difficult to determine the extent of magmatic activity. For the Central North Sea, Denmark and Skagerrak Graben there is geophysical evidence of mafic intrusions in the middle and lower crust. Furthermore, in the Central North Sea area, the thick basalts and tuffs (,160 m) of the Inge Volcanics Formation (c. 299 Ma, Heeremans et al. 2004) are interbedded with Rotliegend-age mudstones and sandstones, whereas bimodal suites of basalt, trachyandesite and rhyolite flows occur in the eastern North Sea. However, in other parts of the Central and Northern North Sea and along the Scottish­ Irish Atlantic seaboard, the importance of the Stephanian ­ Autunian tectonism is difficult to assess.

Many of these latter areas, however, were subjected to uplift (e.g. Dunlap & Fossen 1998).

Intrusive activity (Variscan foreland­ externides)

Stephanian to Early Permian-age intrusive activity in the Variscan foreland occurred largely in the form of dolerite sills and dyke swarms. For example, the Whin Sill Complex in northern England comprises a series of sills (,90 m thick) and four major ENE-trending dyke echelons of high-Fe, sub-alkaline basalt (Johnson & Dunham 2001). The total volume of the dyke complex has been estimated at 120­215 km3; however, as it extends eastwards under the North Sea, this figure is a minimum. In the Midland Valley of Scotland, earth ­ west-trending dykes (up to 50 m wide and 130 km long) and a sill (,180 m thick) are composed of sub-alkaline to transitional basalt, with a total volume of almost 500 km3. This dyke swarm also extends c. 200 km eastwards into the North Sea (Smythe 1994). The Midland Valley of Scotland also contains Westphalian­ Early Permian-age pyroclastic volcanic rocks, vents, necks, alkaline dolerite sills, dykes and plugs of basaltic to trachytic and phonolitic composition (Francis 1992). Many of the vents contain mantle and crustal xenoliths of high-grade mafic and felsic gneisses. The mantle xenoliths probably reflect derivation from variously metasomatized mantle sources, but biotite, amphibole and/or alkali feldspar-bearing ultramafic xenoliths and some of the megacrysts may be fragments of coarsegrained intrusions that fractionated under high-pressure conditions in the upper mantle. The mafic crustal xenoliths may represent metamorphosed cumulate or underplated mafic mantle melts, whereas the felsic gneisses were derived from Precambrian to Palaeozoic crust. In the western Scottish Highlands at least 3000 Early Permian, SW- to NW-trending camptonite and monchiquite dykes, many containing mantle xenoliths, intruded the Caledonian basement (although some of these dykes may be of Tertiary age; Rock 1983). In southern Sweden a c. 70 km wide swarm of NW-trending sub-alkaline basalt and basaltic andesite dykes occurs. Individual dykes can be up to 100 m wide but the majority vary between 1 and 50 m; their total volume has been estimated at c. 4000 km3 (Obst et al. 2004). Of similar age may be the isolated, NNWtrending dolerite dykes on the SW coast of Sweden and the dolerite sills that intruded Cambrian alum shales in south ­central ¨ ¨ Sweden (Vastergotland). These dykes and sills may be related to volcanic rocks obtained from drill core in the Kattegat and offshore eastern Denmark and interpreted from seismic data as volcanic edifices (Mogensen 1994; Marek 2000). Drill core from the ¨ island of Rugen (NE Germany) contains a few basalt sills that intrude lower Palaeozoic sediments, and that are also interpreted as having an age of late Carboniferous­ early Permian (Korich & Kramer 1994).

Magmatic activity in the Variscan internides

Compared with the foreland and internal Variscides, post-tectonic magmatism in the Rheno-Hercynian foreland fold-and-thrust belt was largely restricted to the intrustion of granites (c. 295­293 Ma) in SW England and granitoids (c. 290­ 307 Ma), gabbros, and rhyolitic to basaltic sills and dyke swarms in the Mid-German Crystalline Rise and Harz area in Germany. Stephanian­ Autunian-age magmatism in the internal parts of the Variscan orogen is dominated by granitoid intrusions exposed in midcrustal metamorphic terranes. Volcanic activity, on the other hand, consisted of widespread ash-fall tuffs of distal provenance, with evidence of more pronounced activity being locally restricted. For example, the Ilfeld Basin (southern Harz region) is a small dextral pull-apart basin that contains up to c. 400 m of

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Autunian-age tuffs, ignimbrites and latitic, trachytic and rhyolitic volcanic rocks that erupted from a large number of small centres to overlie Stephanian-age sediments. To the west, the Saar ­ Nahe Basin contains Autunian-age sub-alkaline basalts and andesite flows that are associated with sub-volcanic dacite and rhyolite domes and pyroclastic deposits, and that are interbedded with fluvio-lacustrine sediments (Stollhofen 1994, 1998; Schmidberger & Hegner 1999). In the Western Mediterranean (Iberia, Pyrenees, Balearic Islands, Sardinia, Corsica, Provence), Stephanian­ Autunian-age sedimentation occurred in deep basins. Here, volcanic activity took place in at least two stages, separated by a period of strikeslip activity and granite intrusion (270­ 290 Ma). The initial magmatic stages are of predominantly calc-alkaline character, whereas later, mid-Permian-age, magmatic rocks are alkaline. In the Iberian Peninsula and the Pyrenees, deep basins were formed in which sequences of late Westphalian-C to Stephanian-age clastic sediments are overlain by Autunian-age volcanic rocks. In some half-grabens in the Pyrenees, several cycles of calc-alkaline rhyolitic to andesitic and alkaline basaltic volcanic rocks have been recognized, ranging in age from Stephanian to Early Permian. Late Carboniferous magmatic activity is evidenced by thick sequences of lavas and ignimbrites, but Permian-age ´ volcanicity has also been established (Marti & Barrachina 1987; ´ Marti 1996). The Permo-Carboniferous volcanic rocks are represented by a suite of basaltic andesite to rhyolite of calc-alkaline orogenic type, with a predominance of rhyolitic and rhyodacitic volcaniclastic deposits. Volcanic activity was essentially continuous, and without any significant change in terms of its character, from the deposition of the first Mid-Stephanian sediments to the beginning of the Late Permian sedimentation (c. 18 Ma; ´ Marti 1996). During the Late Carboniferous, volcanism is represented by lavas and pyroclastic deposits mainly associated with caldera-forming events. Indeed, at this time it would appear that each basin had an independent volcanic history, with several volcanic centres located around the basin margins. Most of these deposits appear to be remnants of intra-caldera fill. During the Early Permian only rhyodacitic and rhyolitic magmas were extruded, and volcanic activity was concentrated in one basin (i.e. Castellar de N'Hug Basin), although it is possible that other volcanic centres were also active in the ´ Central Pyrenees at this time (Marti 1996; Fig. 3). Eruptions were explosive, and generated widespread pyroclastic flows and associated pyroclastic surge and fall deposits. In the northern and central parts of the Iberian Peninsula, post-collisional magmas of mafic composition are rare. The magmatic rocks of the Massif Central area of France are limited to centimetre- to decimetre-thick layers of strongly altered air-fall tuffs (tonstein) in both the Stephanian and Autunian sequences, and minor amounts of felsic to intermediate volcanic rocks. In the Rodez Basin, c. 5 m of andesite lavas and c. 30 m of trachytic tuffs occur at the base of Autunian sandstones and are probably of early Autunian age. The Variscan basement south of the Decazeville Basin is intruded by a topaz-bearing rhyolite dome and a later leucogranite, both of crustal origin (Badia & Fuchs 1989). However, U­ Pb zircon dating of a tuff (332 + 4 Ma) and a rhyolite flow (333 + 2 Ma) at the base of, respectively, the Bosmoreau and Decazeville basins showed that some of the sediments and volcanic rocks, previously thought to be of Stephanian age, were actually deposited in the ´ late Visean (Bruguier et al. 1998).

and the external Variscides (i.e. the area of the northern foreland basin, to the north of the Variscan orogenic belt) took place within a relatively short time span (c. 300­ 290 Ma) in a dextral wrenching setting. In the internal parts of the orogen the basins tend to be small, deep and isolated. Basin infill was a relatively rapid process with the accumulation of thick sequences of Rotliegend-age aeolian and fluviatile ­ lacustrine sediments. Tectonic activity coincided with an overall change in climate to semi-arid (Stephanian­ Early Permian). As noted above, many of the basins contain volcanic rocks, and the presence of these magmatic successions aids in distinguishing the Permian sequences from the underlying Stephanian (see Plein 1995). The following section will examine a number of basins and regions across Europe to provide an overview both of the main basin types that formed at this time, and of the differences between the basins that formed in front of and within the Variscan orogen. The section is subdivided into those basins (Northern and Southern Permian basins) that were located to the north of the Variscan orogen, and those that were located within or to the south of the orogenic belt (Fig. 6).

Basins occurring within the Variscan foreland

Southern Permian Basin (SPB). The Southern Permian Basin comprises a series of connected basins extending across Northern Europe from England to Poland (Figs 1 and 6). The SPB, with a north­ south extension of 300­600 km and an east ­ west extension of c. 1700 km, developed between the northern foreland of the Variscan mountain belt and the Ringkøbing­ Fyn High. The south­ central parts of the basin are superimposed on the Variscan fold-and-thrust belts and intervening Gondwana-derived blocks (e.g. East Silesian Massif, SW part of the Malopolska Massif ) and on the Precambrian of the East European Craton. Despite this, the geometry of the SPB shows no direct relationship with either the different crustal domains or the suture zones on which it subsided. The SPB is of great economic importance, containing a number of significant hydrocarbon finds (e.g. Groningen, Salzwedel­ Peckensen). By the early 1990s almost 180 wells had been drilled into the Rotliegend of Western Germany, and in Eastern Germany more than 1500 wells were drilled (see McCann et al. 2000, for details). The succession is subdivided into two distinct units, separated by the Saalian unconformity (Schneider et al. 1995). The Lower Rotliegend is characterized by acid and intermediate volcanic rocks with only minor sediments, whereas the Upper Rotliegend is essentially sedimentary and contains only rare volcanic rocks, of a more basaltic composition (Fig. 7). In the area of the Southern Permian Basin, up to 800 m of Stephanian continental red beds, containing correlative marine bands, were deposited in a broad successor basin to the Namurian ­ Westphalian Variscan foreland basin. During the late Stephanian ­Early Permian, this basin was disrupted by predominantly transpressive wrench tectonics, as evidenced by the deep truncation of Late Carboniferous series and the conspicuous absence of deep Early Permian basins (Ziegler 1990; McCann 1999). These wrench tectonics were associated with the development of extensive volcanic fields (see above). Ziegler et al. (2006) noted that crustal thinning within the North German part of the SPB may be interpreted in terms of a `stretching' factor of 1.45. This thinning is attributed to late Stephanian­ Early Permian magmatic destabilization of the crust ­ mantle boundary that was paralleled by major thermal attenuation of the mantle-lithosphere (Van Wees et al. 2000). Southern Permian Basin: North German Basin. The North German region, forming part of the later SPB, comprises crystalline basement of varying Precambrian ages (see McCann 1998b, for details), which is covered by a thick (.12 km) sedimentary

Permian basins in Europe

Following the end of Variscan contraction in latest Westphalian times, subsequent Stephanian­ Autunian magmatic activity and basin formation in both the internal Variscides (i.e. the area within the Variscan fold-and-thrust belt and to the south of it)

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Fig. 6. Outline of the Northern and Southern Permian basins of NW Europe (after Stemmerik et al. 2000).

succession of early and late Palaeozoic-, Mesozoic-, and Cenozoic-age strata. Basin evolution commenced with the destruction of the subducted oceanic crust of the Rhenohercynian Ocean, which led to lithospheric stretching (dominantly east ­ west) and associated block faulting (Late Carboniferous­ Early Rotliegend, and Early­ Late Rotliegend; e.g. Gast 1991;

Fig. 7. Permian time scale (after Menning 1995). In addition, Menning (2001) has provided an integrated Permian time scale based on absolute dates and field indicators for the duration of permian stages. More recently, the International Stratigraphic Chart (Gradstein et al. 2004) set the Carboniferous ­ Permian boundary at 299 + 0.8 Ma (see also Stratigraphische Tabelle von Deutschland, STD 2002; Menning et al. 2005).

Kiersnowski et al. 1995; Stemmerik et al. 2000; Kockel 2002), accompanied by calc-alkaline magmatic activity (which was particularly pronounced in NE Germany), which resulted in the formation of a series of north­ south-striking grabens. These acted as a feeder rift system for sediment transport from the Variscan hinterland to the areas to the north. The Lower Saxony rift system and associated basins (e.g. Hessen Basin) were part of a major continent-separating suture parallel to the subsequent MidAtlantic Rift (Gast 1988, 1991), which can be extended ¨ ¨ northwards via the Gluckstadter Trough (Schleswig­ Holstein), Horn Graben, and the Skagerrak Graben into the Oslo Graben (Fig. 1). Analysis of the form and sedimentary infill of a series of half-grabens along an east­ west transect in Northern Germany reveals variations in both, suggesting that there were differences in both the rates and amounts of stretching. In the NE German Basin (NEGB) the angular unconformities between the the Westphalian­ Stephanian and the overlying Permian sequences reflect a period of tectonic activity (McCann 1998a). This was followed by the major Permo-Carboniferous volcanic event that heralded the onset of Permian basin evolution. Initial sedimentation (i.e. Autunian) was restricted and mostly confined to isolated basins. At the onset of the latest Rotliegend (i.e. Rotliegend II), there was a clear change in terms of the basin geometry (e.g. Rieke et al. 2001, 2003). Initially, the basin ¨ comprised two distinct sub-basins (Havel ­Muritz and West Mecklenburg basins), although it is not clear to what extent these were isolated from one another (Fig. 8). However, c. 2 Ma later there was a clearly unified depositional area across the entire NEGB, with sediment being sourced from basin margin highs, and adjacent orogenic piles, and transported towards the basin centre (McCann 1998a). The distribution pattern within the basin itself broadly resembles the models of closed-basin sedimentation as outlined by Leeder & Gawthorpe (1987) albeit with the significant difference that the NEGB was not a half-graben structure. Strongly increasing thermal subsidence modified the facies architecture as well as the basin geometry through the remainder of the Rotliegend II period, resulting in a broadly smooth topography, a decrease in sediment supply and the expansion of a playa lake environment across almost the entire basin. The NEGB shows no evidence of significant synsedimentary tectonism (e.g. Kossow et al. 2000), as described for the NW German Basin, the Dutch Basin (e.g. Verdier 1996; Geluk 2005) and the

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Fig. 8. Isopach maps of the Rotliegend-age (a) Parchim, (b) Mirow, (c) Dethlingen and (d) Hannover formations, NE German Basin (after McCann et al. 2000).

English Basin (e.g. George & Berry 1997). This absence presumably reflects the trend of gradually increasing tectonic activity in these latter areas, which may be related to the initiation of the break-up of Pangaea in Late Permian times.

Southern Permian Basin: Polish Basin. The Mid-Polish Trough

(MPT) was continuously connected with the North East German Basin to the NW (Ziegler 1990), but not with the Tethyan domain to the south (Dadlez et al. 1998; Figs 9 and 10). Rotliegend development of the main depocentre of the MPT followed a period of Westphalian wrench tectonics and Early Permian volcanic activity (within the NW MPT and the NE Germany Basin) associated with regional wrenching and generalized crustal destabilization. Immediately overlying the volcanic rocks are late Rotliegend continental sediments deposited under arid conditions (Marek 1988). The sediments are broadly similar to those deposited in the NE German Basin to the west, comprising mainly coarser-grained clastic deposits (confined to the basin margins) with finer-grained clastic deposits occupying the central parts of the basin. Depositional environments were mainly fluvial, aeolian and lacustrine (Karnkowski 1999). The extent of the Variscan orogen in Poland is still uncertain (Pozaryski et al. 1992; Dadlez et al. 1994), mainly because of the lack of reliable data documenting the regional extent and styles of orogenic deformation. However, recent studies of well data have provided new clues to the styles of deformation within the external Variscides and the possible location of the front, including the possibly significant role of strike-slip movements within the frontal part of the orogenic zone and source area for foredeep basin infill (Aleksandrowski et al. 2003; Jaworowski

2002; Mazur et al. 2003). Extension, subsidence and basin development, subsequent to Variscan orogenic activity, was parallel to the SW margin of the East European Craton and the Tornquist­ Teisseyre Zone (Figs 7 and 8). The Polish Basin was formed in the Late Palaeozoic to the north of the Variscan orogen (Dadlez 1997, 2006) and belonged to a series of sedimentary basins that developed around the margins of the East European Craton (Nikishin et al. 1996). The main depocentre of the basin is termed the Mid-Polish Trough (MPT), and this was active during the basin's evolution, beginning in the Permian and extending to the Mesozoic, as a zone of maximum subsidence with almost uninterupted sedimentation (Dadlez 1997; Kutek 2001). Indeed, the axial part of the MPT was also the locus of the Late Cretaceous inversion of the Polish Basin (see Krywiec 2002a and Lamarche et al. 1999 for details). The present base of the Permian in the MPT ranges from 3 to 8 km (Dadlez 1998; Znosko 1999). The Polish Basin developed on highly heterogeneous basement. The MPT developed along the Tornquist­ Teisseyre Zone, which marks the boundary between the East European Craton and the Palaeozoic Platform (Guterch et al. 1983; Grad et al. 1999). Integration of industry seismic reflection data with gravity and magnetic data has shown that the location of the NE MPT margin was strongly and directly controlled by the southwestern margin of the East European Craton (Krzywiec & Wybraniec 2003; Krzywiec 2004; Lamarche et al. 2003a). Indeed, the form of the MPT is considered to have partly resulted from rheological boundaries within the Trans-European Suture Zone (Stephenson et al. 2003). Furthermore, the MPT is segmented along strike into (from NW to SE) the peri-Baltic, Pomeranian, Kuiavian and

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lithospheric cooling and contraction following the cessation of the Autunian wrench faulting and magmatism.

Northern Permian Basin: Denmark. The Rotliegend sediments of

Fig. 9. Structure on the base Upper Rotliegend (below sea level) with the geographical segments of the Polish Basin. Contour lines are at 500 m interval. The position of the cross-section shown in Figure 10 is indicated (after Stephenson et al. 2003).

Malopolska segments (Dadlez 1998), which are related to ´ crustal-scale fracture zones (Krolikowski & Petecki 1995; Grad et al. 1999; Lamarche et al. 2003b; Lamarche & SchechWenderoth, M. 2003). Along the axial zone of the basin the sedimentary succession is 3 ­7 km thick (Marek & Pajchlowa 1997) including about 1500 m of Zechstein salt (Tarka 1994). Deposition of Zechstein evaporites during the Permian was limited to the SE ´ by the Grojec Fault, which was also active during the Carboniferous (Zelichowski et al. 1983).

Northern Permian Basin. The Northern Permian Basin (NPB) is an elongate east­west-oriented basin extending across the Central North Sea from Scotland to Northern Denmark (Fig. 6). The basin is bounded to the south by the fragmented Mid-North Sea­ Ringkøbing­Fyn High. To the east and NE the Sorgenfrei­ Tornquist Zone separates it from the stable Fennoscandian Shield. As a result of Mesozoic overprinting, the outlines, geometry and facies patterns of the NPB are less well understood that those of the SPB (Ziegler 1990), largely because of the lack of outcrops and limited drilling (Stemmerik et al. 2000). To the west the basin rests on Devonian clastic deposits whereas to the east it overlies Caledonian basement, on Lower Palaeozoic sediments preserved in the Caledonian foreland, and on Precambrian crystalline rocks of the Fennoscandian Shield (Hospers et al. 1985). Ziegler (1990) has noted that it is assumed that the NPB also evolved in response to

Denmark were deposited in the Danish Central Graben and the Danish­ Norwegian Basin, located north of the fragmented Mid-North Sea ­ Ringkøbing­Fyn High, and form part of the succession deposited within the Northern Permian Basin (Fig. 11; Stemmerik et al. 2000). Modelling suggests that early Permian extension resulted in crustal thinning of up to c. 1.35 and was probably related to a major Late Carboniferous ­Early Permian heating and extension phase (Frederiksen et al. 2001). In the Danish North Sea the Rotliegend is subdivided into two formations. The lowermost Karl Formation is areally widespread and is defined to include the syntectonic volcanic, volcaniclastic and sedimentary fill of the Permian half-grabens. The succession is dominated by volcanic rocks of alkaline basaltic composition and volcaniclasti sediments. With the exception of some isolated rhyolite flows (e.g. Horn Graben Rhyolite Member), the succession differs in composition from the more acid volcanic rocks that characterize the Lower Rotliegend of the SPB (Aghabawa 1993). The rhyolitic flows were extruded as silicic lava, which solidified as glass and was subsequently devitrified (Aghabawa 1993). The majority of the volcanic rocks are basic in composition, described as alkaline basalts, hawaiites and mugearites (Aghabawa 1993). Associated sediments are mainly volcaniclastic sandstones and conglomerates that were probably deposited in alluvial and fluvial settings. The overlying Auk Formation represents the postrift succession, comprising sandstones, pebbly sandstones and conglomerates with thin interbeds of silty mudstone, which was probably deposited in aeolian, fluvial and sabkha ­lacustrine settings.

Oslo Rift. The Permo-Carboniferous Oslo Rift extends northwards

from the Sorgenfrei­ Tornquist Zone and comprises three opposing half-graben segments: the southern Skagerrak Graben, the central Vestfold Graben and the northern Akershus Graben (Figs 12 and 13). The basin formed as a result of dextral transtension, which led to the development of a purely extensional regime in the Oslo Rift system and the culmination of rifting and peak magmatic activity in the Oslo Graben (Olaussen et al. 1994; Heeremans et al. 1996; Torsvik et al. 1998). The onshore part of the present-day Oslo Rift consists of a c. 400 km long and c. 35­ 65 km wide graben containing large volumes of rift-related extrusive and intrusive rocks and minor amounts of rift-related sedimentary rocks (Neumann et al. 1995, 2002). The driving force behind the initiation of rifting in the Oslo region is difficult to ascertain in terms of active (plume-related) and/or passive (lithospheric stretching) end-members (Kirstein et al. 2002). There is no evidence of crustal doming prior to magmatism (Olaussen 1981). Late Carboniferous pre-rift sediments (i.e. Asker Group) were deposited close to sea level and contain evidence of episodic marine incursions (Olaussen 1981). Rifting in the Oslo region appears to have begun in Late Carboniferous time with the formation of a shallow depression close to

Fig. 10. Geological cross-section based on well and seismic data through the Polish Trough (location shown in Fig. 9). The approximate location of the Trans-European Suture Zone (TESZ) is indicated (after Stephenson et al. 2003).

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Fig. 11. Simplified SE ­NW cross-section from the eastern North Sea­Ringkøbing ­Fyn High to the Danish­Norwegian Basin showing the distribution of Rotliegend sediments. (Note the onlap of the Auk Formation towards the Ringkøbing­Fyn High (after Stemmerik et al. 2000).

Fig. 12. Regional structural map of the Oslo Rift and adjacent areas based on maps by Ramsberg et al. (1977), Falkum & Petersen (1980), Buer (1990) and Ro et al. (1990b). FZ, fracture zone; GS, graben. The location of the seismic profile OG-7 (shown in Fig. 13) is indicated.

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Fig. 13. Seismic line drawing of profile OG-7 across the Skagerrak Graben (after Ro et al. 1990b). (See Fig. 12 for location.) TWT, two-way travel time.

sea level, in which deposited sediments are sealed by the firsterupted lava flows (i.e. B1 basalts). This suggests that rifting began with a period of lithospheric stretching and thinning prior to the onset of the main magmatic phase. Coeval intrusion of sills of intermediate to felsic character is interpreted in terms of compressional activity prior to the main phase of extension (Sundvoll et al. 1992). The initial magmatic phase (involving basalt extrusion) and vertical movement along NNW­ SSE- to north­ south-oriented faults appears to have been contemporaneous in response to ENE ­ WSW- to east­ west-oriented crustal extension (Heeremans et al. 1996). The main phase of rifting and volcanism involved large displacements (up to 3 km) along some of the basin-bounding faults (i.e. the Oslofjorden fault; Neumann et al. 2004, and references therein). This phase also coincided with the development of the Oslo Graben. Subsequent caldera formation was accompanied by changes in the magma chemistry and the increasing dominance of intrusive activity (Neumann et al. 1995), associated with a change in the orientation of structural deformation to more NE­ SW, NW ­ SE and west ­ east. The Permian Oslo Rift was located within a broad intracratonic basin during early Palaeozoic time (Ramberg 1976) containing an up to c. 4 km thick sedimentary sequence of Cambrian to late Silurian age. This succession is locally unconformably overlain by Upper Carboniferous and Lower Permian sedimentary rocks which underlie the Permian lavas (Olaussen 1981). Sedimentary rocks also occur as thin layers (,10 m) between lava flows. Significant thicknesses of fanglomerates are preserved close to the Oslofjorden fault zone in the SE part of the Oslo Graben. Here the clasts mainly comprise lavas and these deposits are indicative of synvolcanic tectonic activity (see Larsen et al. 1978, for more details). The thickness of Permian sediments, however, is secondary to that of the magmatic succession, attaining (in the Skagerrak Graben) a maximum thickness of only c. 1 km (Ro et al. 1990a).

The North Swiss Permocarboniferous Basin (NSPB) is found within the subsurface of northern Switzerland and was first discovered early in the 1980s (e.g. Matter et al. 1987; Fig. 15). Initially, it was called the Constance­ Frick Trough (Konstanz­ Frick Trog, e.g. Laubscher 1987) and later, as its continuation further to the west was recognized, the NSPB (i.e. Nordschweizer Permokarbon Trog). Formed within crystalline basement, the basin is 10­ 12 km wide and filled with ¨ .1500 m of continental clastic deposits (see Blum 1989, for ¨ overview). Supply was local (Matter 1987; Blum 1989). Carboniferous deposits have been drilled only at Weiach and appear to be restricted to a narrow, graben-like structure. The NSPB is disrupted by a series of NW ­SE-trending faults (Fig. 15). At Weiach 572 m of Carboniferous sediments were drilled, but the basement was not reached (Fig. 16). Microfloral remains are dated as Stephanian (Hochuli 1985). Later radiometric dating of zircons from ash layers confirmed the age (303 Ma in the middle of the drilled Carboniferous deposits and 298 Ma at the top; Schaltegger 1997a). The palaeogeographical position within the Variscides, the varying sediment thickness, and the dominance of crystalline basement clasts derived from local sources suggest a pull-apart ¨ origin of the NSPB (e.g. Matter 1987; Blum 1989). The basin is interpreted as a series of en echelon pull-apart basins (e.g. Diebold et al. 1992) that formed in relation to late Variscan wrench tectonics. Areally, the Permian deposits occupy a significantly wider area than those of the Carboniferous. It would appear that only Early Permian deposits were accumulated, but biostratigraphical dating using palynomorphs (Hochuli 1985) may be subject to some uncertainty because of palaeotopographic effects (e.g. Becq-Giraudon 1993). Except for Weiach, the Permian deposits rest on crystalline basement. Permian deposits ¨ were mainly correlated according to lithology (e.g. Blum 1989) and are all continental, ranging from lacustrine to alluvial fan units.

Massif Central. The Variscan Massif Central shows the characteristic tectonometamorphic evolution of classic collisional belts, with significant horizontal thrusting and progressive crustal thickening (Echtler & Malavielle 1990). The Montagne Noire forms the external, southernmost segment of the Massif Central, which may be subdivided into a crystalline metamorphic core (Axial Zone) and a tectonically overlying upper unit of strongly deformed sedimentary rocks (Figs 17 and 18). Partly interfolded ´ sediments of late Visean and Namurian age occur along the southeastern border of the thrust belt (Engel et al. 1982). Up succession, the unit is increasingly dominated by synorogenic coarse-grained turbidites and a progressively chaotic set of kilometre-sized olistolithic slabs, which are syntectonic (Engel et al. 1982). Deposition of these units was followed by a period of Westphalian ­ Stephanian extension.

Basins occurring within the Variscan orogen

Permo-Carboniferous Basin, Switzerland. The Late Palaeozoic basins in northern Switzerland formed after the main Variscan orogeny (dated in Switzerland as pre-late Westphalian), but in the French Alps, post-orogenic sedimentation commenced in the ¨ Late Namurian (Trumpy 1980). Because all these basins formed within the same time span and in close proximity to one another, they exhibit striking similarities, in that the basins, predominantly grabens or graben-like structures, tend to be small, elongate and mainly trending SW ­ NE (Fig. 14). Furthermore, Carboniferous deposits are restricted to a narrow trough whereas the basins are filled with continental clastic deposits, mainly eroded from the crystalline basement, and containing volcanic and/or volcaniclastic material.

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Fig. 14. Permo-Carboniferous basins in Switzerland north of the Penninic front. Basins close to the Penninic front (BF, ¨ Bifertengratli; GV, `Glarner Verrucano ´ Basin'; SD, Salval­Dorenez Basin) are ¨ ¨ exposed (after Trumpy & Dossegger 1972). Basins in the subsurface (NSPB) after Boigk ¨ & Schoneich (1974), Bachmann et al. (1987), Meier (1994), Thury et al. (1994) and Wetzel & Allia (2003).

Autunian areas, from France (e.g. Paris Basin

deposits in France are found in a variety of the classic region of the Massif Central, to NW Carentan Basin, Chateauneuf & Farjanel 1989), the (Bouas 1987), the northern part of the Pyrenees

(Bixel & Lucas 1983) and SE France (Chateauneuf & Farjanel 1989) (Figs 17 and 19). Stephanian­ Autunian basin formation in the Massif Central is attributed to the collapse of thickened crust, aided by a NE­ SW-oriented extensional stress field.

¨ Fig. 15. North Swiss Permo-Carboniferous Basin in map view. Compiled from data by Boigk & Schoneich (1974), Diebold et al. (1992), Meier (1994), Thury et al. (1994) and Wetzel & Allia (2003). In addition, data from boreholes have been used (see McCann et al. 2007, for details). URG, Upper Rhine Graben; RF, Rhine Fault (Rhenish Lineanent).

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Fig. 16. North Swiss Permocarboniferous Basin cross-section close to the well at Weiach (see Fig. 15), and based on interpreted seismic records (Diebold et al. 1992; after Schaltegger 1997a,b). The boundary between the lower and upper basin fill units roughly corresponds to the Carboniferous­Permian boundary.

In the Massif Central the most notable event marking the Stephanian was the appearance of numerous small limnic coalbearing basins, closely related to a network of regional faults (related to Variscan tectonics) that appeared or were reactivated at this time (Arthaud & Matte 1975). From the end of the Westphalian to Stephanian B north­ south compression continued in the Massif Central, but the Westphalian regional ductile shear faults progressively gave way to fracture deformation, responsible for the formation and evolution of the Stephanian limnic basins. The anti-clockwise rotation of north­ south-

Fig. 17. Permian basins of France (after Cassinis et al. 1995).

Fig. 18. (a) Schematic cross-section of the Montagne Noire metamorphic core complex from the southern part of the Massif Central, France. The lower plate (L.P.) comprises high- to medium-grade metamorphic, gneissose and mica schist formations whereas the upper plate (U.P.) consists of undifferentiated low-grade or non-metamorphic Cambrian to Early Carboniferous sediments. (b) The main Permo-Carboniferous basin in the French Massif Central (after Echtler & Malavielle 1990).

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oriented compression to an east­ west direction was responsible for the evolution of these regional strike-slip faults and consequently of the associated basins. At the end of the Stephanian, the final stage of east ­ west-oriented compression interrupted sedimentation and caused intense deformation of the basins located on the approximately north ­south-striking faults. In the Pyrenees, the North Pyrenean Fault, which borders the axial zone, probably underwent dextral strike-slip with horizontal displacement estimated at 150 km (Arthaud & Matte 1975). Subsequent to the period of Stephanian compression at the close of the Variscan orogeny, a new period was initiated during that intracontinental basins formed, the thickness and extent of whose deposits differed from those that accumulated during the Stephanian. The Stephanian compressive basins of the Massif Central are long and narrow with varying thicknesses of sediment, whereas the extensional Permian basins contain successions up to several thousands of metres thick and are much more extensive, occupying the sites of the future large Mesozoic basins. In France the Permian includes only continental deposits. The succession is generally subdivided into Autunian and `Saxonian' strata (both of Early Permian age) and Thuringian deposits (considered to be more or less equivalent to the Late Permian) (Cassinis et al. 1995). The typical Autunian was defined in the Autun Basin, located to the north of the Massif Central. These units comprise lacustrine calcareous and bituminous shales, coarse fluvial deposits and some volcanic ash, which overlie folded Stephanian beds (Cassinis et al. 1995). The many Stephanian­ Autunian basins in the Massif Central contain coal measures and alluvial, fluviatile and lacustrine clastic sediments deposited in an active tectonic environment; magmatic rocks are present in only small volumes (Legrand et al. 1994; Djarar et al. 1996; Allemand et al. 1997). Saxonian deposits include those that overlie the Autunian, and these units are generally separated by an unconformity attributed to a period of intra-Permian (i.e. `Saalian') deformation. The expression of this varies from changes in sedimentation to an actual angular unconformity (Cassinis et al. 1995). Permian-age sediments were deposited on a variety of strata that were deformed and metamorphosed during the Variscan orogeny, and in many basins sedimentation continued into late Permian ^ times (Chateauneuf & Farjanel 1989). Sedimentation was controlled by synsedimentary basin margin faults, and faulting caused considerable palaeorelief and the formation of basement horsts that were important sources of sediment, especially alluvial cones. Some of the larger basins began as smaller sub-basins with independent drainage systems that amalgamated during acceler´ ated subsidence (e.g. the Cevennes Basin, Djarar et al. 1996). Those basins located along approximately east-trending faults (e.g. the Saint-Etienne Basin) tend to have pull-apart geometries that are related to strike-slip movements along the basin margin faults. Those basins located along approximately north-trending faults tend to be half-grabens, some of which are strongly asymmetrical (Faure 1995; Mattauer & Matte 1998). Furthermore, the extent of soft-sediment deformation from the Saint-Affrique ´ and Cevennes basins testifies to the syndepositional, extensional to transtensional character of the basin margin faults (Legrand et al. 1994; Djarar et al. 1996). In the southern Massif Central distinct sequences are present ` within the Autunian deposits. For example, in the Lodeve Basin (Laversanne 1978) sedimentary sequences between 8 and 15 m thick occur within a unit that is c. 800 m in thickness (Fig. 19). These sequences, which are continental in origin, comprise fluviatile sandstones in the lower part, sandstones and bituminous dolomites of palustrine or lagoonal origin in the middle part, and calcareous silty floodplain mudstones and siltstones in the upper part (which also contains evaporitic precipitates). Each basin developed in a general tensional context. In many cases there is no appreciable lithological change through the entire Upper Permian sequence and, because of a lack of data, many workers

prefer to classify such rocks as Saxonian­ Thuringian units. In ` the Lodeve Basin the Upper Group is present.

Saar ­ Nahe Basin. This Permo-Carboniferous basin extends from

SW Germany into France and is filled with exclusively continental sediments (Fig. 20). The basin has a half-graben geometry, being ¨ bordered to the north by the south-dipping Hunsruck Boundary Fault (HBF), which, according to Henk (1993) is a detachment soling out at mid-crustal levels (at a depth of c. 16 km). Transtensional subsidence of the partly inverted Saar ­ Nahe Basin, which contains up to 5.6 km of Permo-Carboniferous clastic deposits, accounts for a stretching factor of .1.36. Contemporaneous extrusive activity reflects destabilization of its lithospheric system (Ziegler et al. 2006). The evolution of the Saar ­ Nahe Basin is closely related to the complex kinematics of the HBF, which is part of a prominent suture zone separating two of the main tectonostratigraphic units of the Variscan fold belt: the Rheno-Hercynian and the Saxo-Thuringian zones. The sedimentary infill of the Saar ­ Nahe Basin is dominated by continental ¨ ¨ clastic deposits (Schafer 1989; Schafer & Korsch 1998) with a significant thickness of contemporaneous volcanic rocks. Basin initiation occurred in the latest Namurian or possibly earliest Westphalian, and the absence of older Namurian sediments ´ suggests that this was an elevated area from the late Visean to the late Namurian. This period of non-deposition reflects Variscan compression, uplift, exhumation and cooling, as indicated by 320­ 335 Ma 40Ar/39Ar cooling ages from Mid-German Crystalline Rise (MGCR) basement rocks to the south (i.e. Odenwald and Erzgebirge, Werner & Lippolt 2000; Schubert & Lippolt 2000). There is an angular unconformity between the Westphalian D and Stephanian A, indicative of a reorganization of fault kinematics in the area. Major uplift in the southeastern part of the basin elevated the MGCR basement to provide a new sediment ¨ source (Korsch & Schafer 1995). The volcanic rocks are related to the crustal stretching that commenced in the earliest Westphalian, where a regional, rather than local, mechanism was responsible. The cause of this was probably dextral translation between Laurussia and Gondwana, which was about to collide with eastern to southeastern North America. Rotation and translation led to the reactivation of old lineaments (such as the northern ¨ ¨ boundary of the Saar ­ Nahe Basin, the Hunsruck Fault) and the establishment of large-scale, new fault systems in Europe (e.g. Arthaud & Matte 1977; Ziegler 1990). Magmatism led to a thermal anomaly, leading to a subsequent phase of thermal relaxation and crustal re-equilibration. The west ­east extensional stress regime that dominated during the post-Westphalian basin formation (Stollhofen 1998) was oblique to the pre-existing fault pattern; the common slip direction was maintained by transtensional strike-slip movements on the transfer faults and oblique-slip motion on the normal faults (Stollhofen et al. 1999). An angular unconformity, indicative of a reorganization of fault kinematics in the area, underlies the synrift megasequence. The Stephanian prevolcanic synrift sequence has a thickness of 3.8­ 4.7 km and was deposited over a period of c. 14 Ma. It comprises lacustrine, fluvio-deltaic and fluvial sediments with minor limestones, coals and pyroclastic fallout deposits (these last were derived from sources outside of the basin; Stollhofen et al. 1999). This is overlain by the Lower Permian volcanic synrift sequence (1.1 km thick, deposited over c. 4 Ma); during this phase widespread bimodal calc-alkaline magmatism occurred with subvolcanic intrusion of rhyolitic­ dacitic domes and basaltic to andesitic sills and dykes (von Seckendorff et al. 2004b). Magma generation can be attributed to underplating and intrusion of pulses of mantle-derived melts into the crust inducing partial melting, or perhaps a distal subduction zone (Schmidberger & Hegner 1999). The extrusive rocks are interbedded with fluvial sediments and minor lacustrine units (Stollhofen 1994).

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Fig. 19. Cross-sections through Permian basins in Central ` France: Lodeve Basin, south of the Massif Central (after Cassinis et al. 1995), Autun Basin and Blanzy­ Le Creusot Basin ` (after Bles et al. 1989). (Locations are shown in Fig. 17.)

Fig. 20. Geological interpretation of DEKORP 1C and 9N showing the half-graben form of the Saar­Nahe Basin, Germany (after Henk 1993).

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Iberian basins. The Iberian microplate was affected during the Permian by the final stages of Variscan activity and by the early Atlantic rifting (Fig. 21). These tectonic disturbances led to the development of intracratonic basins. Two phases of Variscan deformation have been described in northern Spain, an early phase consisting of two east­ west-oriented compressional events in the period from the Namurian to the Stephanian that resulted in arc-parallel folds and thrusts, and a later north ­ south-oriented phase (Sakmarian ­ Early Permian) that marked the final stage of Variscan deformation in northern Iberia and resulted in plunging fold axes of the early arc-parallel folds (Weil et al. 2000). Subsequent rifting episodes (e.g. Early Permian­ Triassic) resulted in the formation of basins with orientations ranging from NE­ SW to east ­west (Jabaloy et al. 2002). Additionally, magmatic events have been recorded. Well-dated volcanic rocks and coeval granites crop out in the Pyrenees, the ~ Iberian Range and the central part of Iberia (Munoz et al. 1986; ~ Sopena et al. 1988). The basins are filled with red beds interdigitated with volcanic and volcaniclastic rocks. Thicknesses are variable and can be as much as 2000 m. These units appear as isolated outcrops of generally small extent because they were deposited in complex graben systems. Stephanian deformation involved NE­ SW compression and horizontal extension orthogonal to the compression direction. In the Early Permian there was major volcanic activity. The fracture patterns associated with this activity in the central part of Iberia indicate extensional deformation in this ´ region (de Vicente et al. 1986; Gonzalez-Casado et al. 1996; van Wees et al. 1998) and probably also in the Pyrenees. Structural data are scarce, but appear to indicate that the central area was characterized by a NNE ­SSW- to NNW­ SSE-oriented extension

´ ´ ´ (Arche & Lopez-Gomez 1996; Gonzalez-Casado et al. 1996), although transcurrent deformation with a NE ­ SW orientation continued in the SW of the Iberian Peninsula (Herraiz et al. 1996). There are five regions in Spain where the Permian is well known: the Cantabrian Mountains, the Pyrenees, the Central System, the southern margin of Iberia and the Iberian Range (Cassinis et al. 1995). The succession may be broadly subdivided into a Lower and an Upper Group. The Lower Group, associated with the formation of small basins controlled by strike-slip faults, comprises one or more tectonosedimentary units that unconformably overlie the Stephanian or older Palaeozoic rocks. Its thickness ranges between 200 and 2000 m. Widespread andesitic and rhyolitic volcanism of calc-alkaline type is associated with this group (Lago et al. 2004). The overlying Upper Group sediments were deposited in an extensional cycle and are areally more widespread. There is a marked unconformity between the two groups. The Permo-Carboniferous succession of the Central Pyrenees has been considered to have originated from strike-slip dynamics that developed during a compressional episode at the end of the Variscan orogeny as established from facies analysis (e.g. Marti 1986, 1991; Besly & Collinson 1991) and from regional palin~ spastic reconstructions (e.g. Munoz 1992; Casas et al. 1989). In the Pyrenees, the oldest sediments comprise breccias, sandstones and coal beds, and contain Stephanian flora (Fig. 22). These are overlain by a widespread volcanic unit comprising andesitic pyroclastic deposits and volcaniclastic rocks (see Cassinis et al. 1995; Lago et al. 2004, for details), which is overlain by conglomerates and sandstones that contain a Stephanian ­Autunian flora. In the Anayet Massif, for example, there is an extensive

Fig. 21. Tectonic sketch map of the Iberian Peninsula (after Van Wees et al. 1998). Inset map provides a reconstruction of the basin-bounding faults of the Iberian Basin (after Arche & ´ ´ Lopez-Gomez 1996).

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Fig. 22. Upper Carboniferous and Permian of the Pyrenees and Iberian Ranges in Spain (after Cassinis et al. 1995). AU, Autunian macro- or microflora; ST, Stephanian macro- or microflora; TH, Thuringian microflora.

sedimentary record of Stephanian ­Permian-age deposits. Here, the succession is composed of four lithological units representing a stratigraphic transition from Stephanian to Permian. In terms of the sediments, these units represent a transition from humid to arid climatic conditions (with fluvial sediments) (Gisbert 1983). Contemporaneous with sedimentation, there were several intrusive and extrusive magmatic episodes favoured by the transtensional tectonic regime, which was also responsible for the development of small listric-fault-bounded basins containing up to 3 km of volcanoclastic and detrital rocks (Bixel et al. 1996). These mixed magmatic­ sedimentary successions were deposited in isolated basins (five of which have thus far been identified; ´ Marti 1996). The only outcrops of marine deposits of Stephanian and Permian age to be found in Western Europe (with the exception of the Eastern Alps and Sicily) occur within the Iberian Massif ´ ´ in western Spain and Portugal (Martinez Garcia 1990). Compressional tectonics characterizes the evolution of the Stephanian basins, whereas the Permian (and Stephanian C) shows a tensional environment that reveals plate-tectonic conditions suggestive of failed rifting. Two units, corresponding to different basins, have been distinguished in the Permian succession of the Cantabrian ´ ´ and Palential zones by Martinez Garcia (1983). The lowermost of these comprises alternating clastic sediments, tuffs, volcanic agglomerates, lava flows and shallow marine limestones. Overlying these are volcanic tuffs and ash up to 900 m thick and with an increasingly acid character. Thick conglomerate units attest to syndepositional tectonic activity. The unconformably

overlying succession comprises predominantly continental clastic sediments. The Permian succession of the present-day Iberian Range, eastern Spain, is very well exposed (Figs 22 and 23). These sediments are of continental origin and filled isolated, small basins that were initiated during latest Carboniferous ­early Permian ~ times (Sopena et al. 1988) and that evolved into a single basin (i.e. the Iberian Basin) during late Permian times. This period was one of tectonic readjustment of plates by transtensional faulting in the Chedabucto­ Gibraltar and Bay of Biscay areas and the development of a conjugate wrench zone, originating later in the Iberian Basin, following an ancient suture (NE ­ SE-oriented) running across the microplate (Arthaud & Matte ´ ´ 1977; Salas & Casas 1993; Lopez-Gomez et al. 2002). During early Permian (Autunian) times, a series of intermontane basins were filled in by alluvial fans, slope breccias and lacustrine deposits associated in their lower part with volcaniclastic rocks of calc-alkaline affinities (Lago et al. 1992). Lacustrine deposits at the top of the volcaniclastic succession record a change from freshwater to saline lakes, indicating a progressive aridity in Iberia. These early Permian sediments were mostly unconformably deposited on rocks of early Palaeozoic age, but also upon late Carboniferous (Stephanian B­ C) sediments. These comprise sandstones and coal-bearing successions that filled the Henarejos intermontane small basin contemporaneously with those of the Cantabrian and Palential zones and are related to the final, extensional collapse of the Variscan orogeny, which also resulted in monzogranitic magmatism and uplift in Central Iberia.

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During the late Permian, red bed successions up to 300 m thick accumulated in grabens and half-grabens that formed during a period of widespread extension within the Iberian Basin ´ ´ (Lopez-Gomez et al. 2002). The overall development shows a fining-upwards trend in most of the basins, beginning with transverse alluvial fan deposits and followed by longitudinal braided river system deposits. There was no volcanic activity during this period in the Iberian Basin, where basin boundary faults have shallow listric geometries at depths of only 13­ 14 km. The outcrops in the Iberian Range are the most widespread and the best studied. The Molina de Aragon series (Ramos 1979; Ramos & Doubinger 1979) is the most characteristic and the one that is best correlated with the Permian of Central Europe (Virgili et al. 1976, 1983; Fig. 23). The lowermost units

unconformably overlie Lower Palaeozoic rocks and consist of breccias, volcano-sedimentary rocks and thin volcanic intercalations. These are overlain by 300 m of lacustrine and fluvial sandstones and black shales. The succession ends with lacustrine, siliceous dolomites unconformably overlain by the Upper Group (Permian). The onset of sedimentation of the Upper Group is marked by an important and widespread unconformity reflecting an important change in tectonic evolution.

Geophysical investigations

Subsequent to the European Geotraverse (EGT; Blundell et al. 1992), the increase in the number and quality of seismic data

´ Fig. 23. The Palmaces­ Riba and Molina de Aragon sections from the Castilian branch of the Iberian Ranges. AS, Arandilla sandstones; HGC, Hoz de Gallo sandstone; MB, Monstsoro beds; PB, ´ Prados beds; PLC, Palmaces lower ´ conglomerates; PMS, Palmaces mudstone ´ and sandstone; PS, Palmaces sandstone; ´ PUC, Palmaces upper conglomerates; RGS, Rillo de Gallo sandstone; RSC, Riba de Santiuste conglomerates; RSS, Riba de Santiuste sandstones; VSC, volcano-sedimentary complex (after ´ ´ Lopez-Gomez et al. 2002).

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has greatly improved our understanding of the nature of the crust that underlies many of these Permo-Carboniferous-age basins, allowing us to ascertain the nature and relevance of any preexisting structures (e.g. the Tornquist­ Teisseyre Zone and the Sorgenfrei­ Tornquist Zone) to the formation of the post-Variscan basins (e.g. Polish Trough and Oslo Rift, respectively). The following section will outline some of the main advances and results that have helped to improve our understanding of the evolution of these basins.

Germany

A series of DEKORP profiles were shot in the 1980s to image the western margins of the Variscan deformation front close to the German border with Belgium (Meissner & Bortfeld 1990). Of greater interest, however, was the series of profiles shot in the 1990s, which imaged not only the North German Basin but also the major structures to the north and south of it. Prior to obtaining the deep seismic profiles, the precise nature of the North German Basin (i.e. whether it was an extensional basin or not) had been much discussed in the literature. As is clearly shown on the DEKORP 9601 profile (DEKORP-BASIN Research Group 1999), the basin is intracratonic, and was initiated above a region that had undergone significant rupturing as a result of the wrench tectonics at the Permo-Carboniferous transition. However, there is little evidence of these movements on the seismic profiles. Instead, we have a clear image of the Moho (continuous across the entire profile) and the northern (where Avalonia is obducted onto Baltica) and southern (with clear evidence of Variscan thrusting) margins (Fig. 24). The depth to the Moho is 30 km beneath the entire basin, except at the margins, where it extends to 35 km. More detailed profiles are required to reveal any evidence of Stephanian ­Rotliegend tectonic activity. However, tectonic activity in the form of continuous transpression is evidenced by the diverse angular unconformities visible between the Westphalian and Permian deposits on 3D seismic profiles and in drill core (McCann 1998a). In the Saar ­ Nahe Basin two deep seismic profiles (DEKORP 1C and 9 N) crossed parts of the basin (DEKORP Research ¨ Group 1991; Korsch & Schafer 1991). The southernmost part of the basin was also imaged by the ECORS-DEKORP 9S profile (Brun et al. 1991). These profiles reveal the asymmetric geometry ¨ of the basin bounded by the South Hunsruck Fault, which has been interpreted as having either a subvertical (Korsch & ¨ Schafer 1995) or listric (Henk 1993) geometry, although the latter interpretation is the favoured one. The upper crust beneath the basin shows a segmentation into three distinctively different reflectivity patterns. The uppermost highly reflective package represents the pre-rift sediments and Permo-Carboniferous basin fill. The underlying wedge-shaped unit lacks major reflectors and is interpreted as crystalline basement rocks of the northern Saxo-Thuringian Zone (i.e. Mid-German Crystalline Rise). Beneath is a thick, highly reflective zone, which may represent remnants of sediments and oceanic crust from the

Rheno-Hercynian Ocean (Behr et al. 1984; Franke & Oncken 1990), or an anastomosing pattern of shear zones and duplex structures (DEKORP Research Group 1991). Heat-flow modelling has been carried out in a number of areas, generally in conjunction with other forms of modelling. In NE Germany, for example, Ondrak et al. (1997) integrated the structural models of Scheck (1997) to produce a regional thermal model that allowed the determination of temperature distributions and a depth-dependent estimation of the local heat-flow conditions. Temperature gradients within the model matched the regional trends of heat-flow distributions, although the pattern is more complex. There is a clear relationship between temperature and the Zechstein salt layer, where high temperatures are related to salt margins and regions where sediments with low thermal conductivities cause local elevation of the isotherms.

Poland

The crust of central and northern Poland has been extensively studied by deep refraction and wide-angle reflection seismic studies. Older profiles have recently been reprocessed and reinterpreted (see Grad et al. 2003, for details) and integrated with the new high-quality LT7 profile (Guterch et al. 1994), the TTZ profile (Grad et al. 1999) and profiles from the POLONAISE 97 experiment (Guterch et al. 1999; Jensen et al. 1999; Janik et al. 2002). All of these seismic data have provided new information on the deep structure of the transitional area between the East European Craton and the Palaeozoic Platform in central and northern Poland. The depth to the Moho in this region is 32­ 39 km beneath the two-layered Palaeozoic Platform and 43­ 45 km beneath the three-layered crystalline crust of the East European (Precambrian) Craton. The Trans-European Suture Zone (TESZ) located between these two main crustal domains is characterized by the presence of thick (up to 20 km) complexes of low-velocity, sedimentary, metamorphic or volcanic rocks, and by a lower crust characterized by high velocities. In this region, the Moho is located at intermediate depths in the NW (30­ 33 km) and deepens to the SE. According to some interpretations, the lower crust within the TESZ area in central and northern Poland represents the attenuated Baltica margin underthrust towards the SW beneath the Avalonian accretionary wedge (Grad et al. 2002). The results of gravity modelling based on deep seismic refraction data suggest that the crustal structure of the Mid-Polish Trough (MPT), especially of its pre-Zechstein substratum, is ´ more complex than suggested by deep refraction data (Krolikowski & Petecki 1997, 2002; Petecki 2002). The upper mantle beneath the TESZ is dense, and within the upper crust a highdensity body was also identified, as well as a complex transition zone between the crust and the upper mantle. Long-wavelength gravity anomalies are associated with lateral density variations ´ within the upper mantle and lower crust (Krolikowski & Petecki 2002; Petecki 2002). Results of magnetic modelling suggest that the average depth of magnetic sources within NW Poland is of the order of 18 km, and could be correlated with the crystalline

Fig. 24. Interpreted line drawing of BASIN 9601 profile and its offshore extensions PQ2-009.1 and PQ2-005 showing the main tectonic and stratigraphic features.

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basement as evidenced by seismic data (Petecki 2002). Upper crust fault zones revealed by gravity and magnetic data present within the pre-Zechstein MPT basement played a significant role during the Mesozoic evolution of the MPT (Dadlez 1997) as evidenced by regional interpretation of seismic reflection data (Krzywiec & Wybraniec 2003) and structural modelling (Lamarche et al. 2003a). Geothermal models in Poland have been more concerned with examining the broad structural framework of the region, rather than concentrating on the basin infill succession. For example, Majorowicz et al. (2003) have shown that there is a sharp change in heat-flow data between the East European Craton and the Palaeozoic Platform. Numerical modelling of the crustal temperatures along several deep seismic profiles suggests extensive crustal­ mantle warming within the zone located between the Sudete Mountains and the TESZ (Grad et al. 2003; Majorowicz et al. 2003). This anomalous zone coincides with the location of the Dolsk Fault and the Variscan Deformation Front. High heat flow for the Palaeozoic Platform and related high temperatures of the crust coincide with the reduced crustal thickness, whereas the low heat flow of the East European Craton coincides with a higher crustal thickness. Modelling results also suggest that high mantle heat flow is required within the high heat-flow zone located within the Palaeozoic Platform (characterized by the 100 km thermal lithospheric thickness) whereas cold crust and cold mantle are typical of the East European Craton (characterized by the 200 km thermal lithospheric thickness; Grad et al. 2003; Majorowicz et al. 2003). Thermal modelling of the Polish Basin performed by Karnkowski (1999) suggested that the Rotliegend volcanism began with high geothermal anomalies in the western part of the basin. The anomalies are characterized by higher values (100­150 mW m22) during the Late Permian ­Early Triassic interval, in relation to rifting in the Polish Basin. The Late Triassic and Jurassic were a time of cooling, until the break-point in thermal evolution of the Polish Basin at the Jurassic­ Cretaceous boundary as a result of uplift and erosion, after which the heat inflow decreased.

1991). The subsequent Moho map shows that the crust is thicker east of the Oslo Graben than west of it (Kinck et al. 1991; Thybo 1997). The refraction or wide-angle reflection data of greatest relevance to the crustal structure of the centre of the Oslo Graben segment include the profiles of Tryti & Sellevoll (1977) (see Neumann et al. 1995, for details). Three distinct crustal layers were noted, including a high-velocity (7.1 km s21) lower crustal layer that had been interpreted as resulting from magmatic underplating (see above) and/or a zone of magmatic cumulates and residues (Neumann et al. 1995). Teleseismic studies carried out along the 60th parallel (i.e. parallel to the modelled section AA0 ) suggest that deepening of the base lithosphere occurs abruptly from west to east in the Oslo region (Babuska et al. 1988). These finding are confirmed by a more recent teleseismic ´ study (Plomerova et al. 2001).

Modelling post-Variscan basin development

Although numerical methods have been applied to tectonic problems for over 30 years, relatively few have been concerned with modelling the post-Variscan (i.e. Permian) basins of Europe. Most Permian basins are buried at considerable depths by Mesozoic and Cenozoic sediments, and their structure has been obscured by younger tectonic events. As a result, the acquisition of accurate data for the modelling is very often difficult. Nevertheless, the increasing collection of subsurface observations (i.e. well and seismic reflection data) has allowed for the modelling of the post-Permian subsidence history in some areas (i.e. Paris Basin, North Sea, NE German Basin, Danish Basin). In other areas, which were uplifted in post-Permian times, direct observation of the Permian basins (i.e. the Oslo Graben) allows relevant data for thermo-mechanical modelling including rheology and decompression melting to be acquired.

Models based on sediment succession ­basin analysis

Models based on subsidence analyses have been the most widely applied to reconstruct the pre-rift configuration and the rifting history of Permian basins (e.g. Van Wees et al. 1998, 2000; Frederiksen et al. 2001). These models are based upon the classical model of McKenzie (1978), and its numerous derivatives, in which rapid synrift subsidence is followed by a more extended phase of thermal subsidence. The pattern of basin subsidence is strongly dependent on the initial configuration of the lithosphere and the amount of stretching. One of the advantages of this method is that it provides information on a previous rift event from the analysis of post-rift sediments. Thus, direct observation of the synrift sedimentary infill can be bypassed (see Allen & Allen 2005, for further discussion on this topic). A variety of basins both within and to the north of the broad zone of Variscan-age deformation have been studied with this method. Basins within the deformation zone include the Paris Basin (e.g. Brunet & Le Pichon 1982; Prijac et al. 2000), the Mid-Polish Trough (e.g. Dadlez et al. 1995), the Iberian Basin (e.g. van Wees et al. 1998), and the Southern Permian Basin (e.g. van Wees et al. 2000); those to the north of it include the Danish Basin (e.g. Sørensen 1986; Frederiksen et al. 2001), the offshore arm of the Oslo Rift (e.g. Pedersen et al. 1991), and the Northern North Sea (e.g. Odinsen et al. 2000). In detail, the results differ somewhat from one basin to another, reflecting the complexity of the structure of the European lithosphere in Permian times and the interplay between different subsidence mechanisms. For example, Permian rifting and post-Permian subsidence in the Paris Basin can be explained by the collapse of the Variscan mountain chain and the slow decay of the associated thermal anomaly (Brunet & Le Pichon 1982; Prijac et al. 2000). In the Variscan foreland such a mechanism cannot be invoked.

Norway

The Oslo Graben and surrounding areas have been the subject of a number of geophysical investigations (e.g. Ro et al. 1990a,b; Kinck et al. 1991), which have revealed that there is a marked crustal thinning beneath the graben, with the amount of thinning increasing southwards. The depth to the Moho has been estimated to range from 28 to 35 km beneath the Oslo rift system (Cassell et al. 1983; Thybo 1997). Gravity data indicate the presence of a dense 90 km wide body towards the base of the lower crust, which has been interpreted as representing cumulates and gabbroic rocks in a deep crustal magma chamber (Neumann et al. 1992). Furthermore, gravity and seismic data imply a different crustal structure along the rift from in the adjacent Precambrian terrane (e.g. Wessel & Husebye 1987; Kinck et al. 1991). However, more recent data have questioned these interpretations of gravity data in the Oslo region (Ebbing et al. 2006). These data and modelling results show that the high-velocity layer present at the base of the Oslo Graben is similar to the typical high-velocity layers that are commonly found at the base of Proterozoic crusts (Durrheim & Mooney 1991), suggesting that the idea of underplating beneath the Oslo Graben is a false one. It has been noted that although Permian mafic intrusions in the crust may be present, they do not reside in a hypothetical underplate, but most probably in the middle crust (Ebbing et al. 2006). Seismic refraction or wide-angle-reflection data have been used to model variations in crustal thickness and structure in the Oslo rift and adjacent parts of the Precambrian shield (e.g. Bungum et al. 1971; Cassell et al. 1983; Ro et al. 1990a,b; Kinck et al.

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In contrast, Dadlez et al. (1995) and van Wees et al. (2000) argued for strong destabilization of the lithosphere by late Variscan wrenching involving deep fracturing and decompression melting of the underlying mantle. This interpretation explains the large thicknesses of Permo-Carboniferous magmatic rocks and post-rift sediments that accumulated in the absence of major normal faulting of the crust. Modelling results concerning the precise age and timing of the Permian rifting event differ from one study to another. This reflects differences in data coverage between the various studied areas, and differences in the methods used, but also probably different responses of the lithosphere in terms of its structure and past history. Whatever the precise mechanisms associated with the Permian event, it is clear that subsidence modelling studies agree on two major points: (1) latest Carboniferous­ early Permian rifting (290­305 Ma) was a widespread and dramatic event in Europe; (2) the thermal signature of the Permian rifting was a significant control on the subsequent Mesozoic and Cenozoic evolution of the European lithosphere. In particular, the second conclusion has serious implications for oil exploration in the North Sea area (Sørensen 1986; Pedersen et al. 1991; Odinsen et al. 2000) in terms of reassessing the reconstructed thermal history of the region and, by extension, predictions concerning levels of organic maturation. Ziegler et al. (2006) have noted that subsidence curve modelling suggests that there was a period of Permo-Carboniferous `stretching' from 300 to 280 Ma. This involved decoupled crustal extension and mantle-lithosphere attenuation. Such an assumption is compatible with the concept that during the Permo-Carboniferous re-equilibration of the crust ­ mantle boundary crustal extension played a significant, but local, role. This can be concluded from the fact that some important Permo-Carboniferous troughs (e.g. Massif Central, Bohemian Massif) do not coincide with major Late Permian depocentres, and depocentres such as the Southern Permian Basin are not underlain by major PermoCarboniferous basins (Ziegler 1990; Ziegler et al. 2006). This would suggest that during the Permo-Carboniferous tectonomagmatic cycle uniform and/or depth-dependent lithospheric extension was, on a regional scale, only a contributing factor and not the dominant mechanism of crustal and mantle-lithosphere thinning. Mechanical stretching of the lithosphere played a subordinate role whereas thermal thinning of the mantle-lithosphere and magmatic and erosional thinning of the crust dominated, providing the principal driving mechanism for the Late Permian and later subsidence of intracratonic basins.

Thermo-mechanical models

Rheological numerical models investigate the way in which materials deform in response to stresses. Rocks are considered to deform in different ways (i.e. by elasticity, plasticity or viscosity) depending on the duration of the load, the petrological composition, the temperature, and the confining pressure. The rocks can display a very complex rheology, as it is possible that various modes of deformation occur at the same time. An additional control is the depth to the brittle­ ductile transitions, which can be present within each lithospheric layer and are strongly controlled by the geotherm. Thus, rheological modelling implicitly involves thermal calculations and can include routines to determine possible melt volumes. Application of these rheological models is not straightforward, as they require various, and relatively accurate, datasets. The Oslo Rift is one of the few Permian basins in Europe where this requirement is met and, consequently, it has been the primary target for 2D numerical modelling (Ro & Faleide 1992; Pedersen & van der Beek 1994; Pascal & Cloetingh 2002; Pascal et al. 2004). The Oslo Rift presents the paradox of showing little extension (i.e. b , 1.3) in association with huge volumes of synrift magmatic rocks (i.e. .100 000 km3, Neumann et al. 2004). A

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 Beek (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 Beek (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 der 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 rifting, 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 and structural models

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 km, respectively. The modelled line is 500 km long at t ¼ 0 Ma. The model is stretched using a velocity of 1.6 cm a21. The upper panel presents the horizontal strain distributions (i.e. 1xx) 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 ¼ 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; Schafer 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 b ¼ 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 NE ­ SW 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. Schematic cross-section across the northern part of the NE German Basin showing the faulted basement comprising Permo-Carboniferous volcanic units, which were subsequently overlain 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-bearing 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 rifting 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 rifting (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).

(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 processes

The Stephanian­Autunian 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 subduction-related origin. With the possible exception of some magmatic rocks in the Alpine basement, this contradicts their intracontinental setting and the fact that the Variscan oceans had closed by mid-Carboniferous times. However, Sm­Nd 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 1Nd(t) values of the 290­300 Ma volcanic and intrusive rocks of felsic to intermediate composition: 22.1 to 26.0 for the Krkonose Basin (Ulrych et al. 2002), 22.7 to 26.1 for the Intra-Sudetic Basin (Ulrych et al. 2004); 20.8 to 27.0 for the rhyolites of the Halle Volcanic Complex (Romer et al. 2001), 24.3 to 27.5 for the granites in Cornwall (Darbyshire & Shepherd 1994), and 20.6 to 25.7 for the Saar­Nahe 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 lithospheric thinning and decompressional melting of updoming asthenosphere may have been a contributing factor in the late Carboniferous­early Permian period. Compared with the foreland, Stephanian­Autunian 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.

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 Stephanian­ Autunian 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

Magmatic ­ tectonic activity

The relatively short and widespread pulse of Stephanian ­ Autunian 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 Visean­ Westphalian crustal shortening and orogen-parallel extension,

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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-field 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 Carboniferous­early 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 (Sørensen 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-rift 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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