Topography of the Variscan orogen in Europe: failed–not collapsed

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Int J Earth Sci (Geol Rundsch) (2014) 103:1471–1499DOI 10.1007/s00531-014-1014-9

ORIGInal PaPER

Topography of the Variscan orogen in Europe: failed–not collapsed

Wolfgang Franke

Received: 23 January 2014 / accepted: 6 March 2014 / Published online: 9 april 2014 © Springer-Verlag Berlin Heidelberg 2014

Melting in the HT regime extracted mafic components from Variscan and Cadomian crust as well as from Cadomian metasomatized lithospheric mantle, thus mimicking sub-duction-related magmatism. The onset of the HT regime at c. 340 Ma may also have triggered the final ascent of HP/UHP felsic metamorphic rocks.

Keywords Variscides · Palaeogeography · Orogenic topography · HT processes · lithospheric extension · Tethyan rift

Introduction

Ever since the first plate-tectonic concepts, the European Variscides have been interpreted as a collisional belt com-parable with the Himalayas and Tibet (Dewey and Burke 1973; Burg 1983). In fact, these areas have similar dimen-sions, and the Variscides—like Tibet—represent a collage of microcontinents set between two large continents (lau-russia and Gondwana). The Variscan microcontinents were separated by marine basins, at least partly floored by oce-anic crust. The widest ocean was the Rheic between ava-lonia and armorica (Cocks and Torsvik 2006, with refs.), which was closed already in Emsian time (McKerrow et al. 2000; Franke 2000). The Rheic suture was grossly retraced by the opening, from the early Devonian onwards, of the narrow Rheno-Hercynian ocean (from S-Portugal via SW-England, Germany and Silesia as far as the Istanbul terrane, Okay et al. 2001). Other seaways or narrow oceans origi-nated from Cambrian through to Silurian rifting, which spalled off armorica and neighbouring microcontinents from Gondwana mainland (review in Cocks and Torsvik 2006). The position of these basins is marked by accretion-ary belts formed during Devonian–Carboniferous collision:

Abstract The Variscan orogenic collage consists of three subduction-collision systems (Rheno-Hercynian, Saxo-Thuringian and Massif Central-Moldanubian). Devonian to early Carboniferous marine strata are widespread not only in the individual foreland fold and thrust belts, but also in post-tectonic basins within these foreland belts and on the Cadomian crust of peri-Gondwanan microcontinental frag-ments, which represent the upper plates of the subduction/collision zones. These marine basins preclude high eleva-tions in the respective areas and also in their neighbour-hood. Widespread late Carboniferous intra-montane basins with their coal-bearing sequences are likewise incompatible with high and dry plateaus. While narrow belts with high elevations remain possible along active margins within the orogen, comparison of the Variscides with the Himalaya/Tibetan plateau is unfounded. Plausible reasons for the scarcity of high Variscan relief include subduction of oce-anic and even continental crust, subduction erosion, oro-gen-parallel extension and—most important—lithospheric thinning accompanied by high heat flow and magmatism. In many areas, timing and areal array of magmatism and HT metamorphism are not compatible with a model of tec-tonic thickening and subsequent gravitational collapse. It is suggested, instead, that lithospheric thinning and heat-ing are due to mantle activities caused by the Tethys rift. The lower and middle crust were thermally softened and rendered unfit for stacking and isostatic uplift: in terms of topography, the Variscides represent a failed orogen. The HT regime also explains the abundance of granitoids and HT/lP metamorphic rocks typical of the Variscides.

W. Franke (*) Institut für Geowissenschaften der Goethe-Universität, altenhöfer allee 1, 60438 Frankfurt am Main, Germanye-mail: [email protected]

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the Rheno-Hercynian (S-Portugal, SW-England, Germany, Moravia in the Czech Republic), the Ossa Morena Zone in Iberia, the Saxo-Thuringian (best preserved in Saxony and Bavaria, but with possible equivalents in Brittany and Ibe-ria) and the Mediterranean (Galicia-Cantabria, French Mas-sif Central-Montagne noire-Mouthoumet Massif-Pyrenees, with continuations into Bohemia as well as into Sardinia,

the Carnic alps and Karawanken Mts; see Fig. 1 and the synthesis by Matte 2001).

although these basins are not convincingly documented in the biogeographic and palaeomagnetic records (see, e.g. Robardet 2003), the accretionary belts listed above contain metamorphosed sedimentary and magmatic rocks indicative of lithospheric extension, which pre-disposed

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Massif Central

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Alpine Front

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Beja/Aracena – Normannian High– Mid-German Cryst. High andaccreted Rheno-H. ocean

South Portuguese-Rheno-Hercynian

S-Portugal - Rheno-Hercynian -Sudetes Phyllite Zone

Ossa Morena – Saxo-Thuringian

"Median Massifs":Central Iberian (SW), N-CentralArmorican, Teplá-Barrandean

Allochthons of the Mediterranean Variscides

Gondwana

Variscan basement exposed

strike-slip shear sense

direction shore to basinin early Ordovician

Ordovician sections ofBuçaco (Portugal) andCrozon (Brittany)

major zonal boundariesand fault zones

terrane boundaries

Badajoz-Cordoba – SE Léon– Teplá Suture

"stranded" Armorican fragment(? NW-Saxo-Thuringia)

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Fig. 1 Tectonic and palaeogeographic subdivision of the Varisci-des. after Shaw et al. (2012, Iberia), Martínez Catalán (2011, Iberia, armorica, Massif Central), Faure et al. (2010a, b) (léon domain) and Franke 2000, Franke and Żelaźniewicz (2000, 2002) for central Europe. The Massif de Maures-Corsica-Sardinia area is interpreted, after Matte (2001), as the south-eastern (inward-facing) limb of the Bohemian orocline displaced westwards by a dextral strike-slip fault. The Rheic suture is situated within, or at the internal margin of, the narrow belt comprising the Pulo do lobo–Start–northern Phyllite Zone–basement of the Fore-Sudetic Monocline (pink). note that the seaward direction in the early Ordovician in Galicia is opposed to that in armorica and its equivalents in the SW part of the Central Iberian Zone (Buçaco-Crozon: Gutíerrez-Marco and Robardet 1990, Galicia and CIZ: Shaw et al. 2012), which implies the existence of two juxtaposed shorelines. The GTMZ allochthons are interpreted to be rooted in the intervening basin (Galicia—Massif Central Ocean, Roger and Matte (2005), possibly closed along the Malpica-Tuy suture of Shaw et al. (2012). Abbreviations (in alphabetical order): Geographic units B Buçaco, C Crozon Peninsula, F Flechtingen Hills, H Harz Mts., L lizard, O Odenwald, P Pfalz, RS Rheinis-ches Schiefergebirge, S Schwarzwald, Sp Spessart, SP Start Point, T Thüringer Wald, V Vosges. Tectonic units CIZ Central Iberian Zone, subdivided into DUF (domain of upright folds) and DRF (domain

of recumbent folds), CZ Cantabrian Zone, FM Fore-Sudetic Mono-cline, GTMZ Galicia-Tras os Montes Zone, M Moldanubian Zone, MS Moravo-Silesian Zone, NPZ northern Phyllite Zone, OMZ Ossa Morena Zone, SPZ South Portugues Zone, TB Teplá-Barrandean unit, WALZ West-asturian leonese Zone. Fault Zones BCSZ Badajoz-Cor-doba shear Zone and Suture, EFZ Elbe Fault Zone, ISF Intra-Sudetic Fault Zone, MSS Moravo-Silesian Suture, MTS Malpica-Tuy Suture, NES nort-sur-Erdre Suture, OFZ Odra Fault Zone, PTFZ Porto-Tomar Fault zone, RhS Rheic Suture, SAFZ-N and SAFZ-S South-armorican Shear zone, northern and southern branch, SHFZ Sillon Houiller Fault Zone, TAFZ Tauve-aigueperse Fault Zone. Nos. 1–25 Devonian/Early Carboniferous marine deposits in internal parts of the Variscides. 1 Cabrela basin, 2 Terena syncline, 3 Valle syncline, 4a Santos de Maimona syncline, 4b Berlanga syncline, 5 Obejo-Valse-quillo-Puebla de la Reina syncline, 6 Sierra de San Pedro, 7 almaden syncline, 8 eastern Sierra Morena, 9 Guadiato belt and Peñarroya coal basin, 10 Pedroches belt, 11 St. Georges s. loire area, 12 a,b,c Chateulin, Menez-Belair and laval basins, 13 Morvan, 14 Montagne Bourbonnaise, 15 Beaujolais, 16 Brevenne unit, 17 Chagey-Belfort area, 18 Southern Vosges, 19 Badenweiler-lenzkirch zone, 20 Saar I drilling, 21 Delitzsch-Torgau-Doberlug syncline, 22 Hainichen basin, 23 Intra-Sudetic basin, 24 Teplá-Barrandean unit (W of Elbe fault zone), 25 Teplá-Barrandean unit (E of Elbe fault zone)

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the formation of subduction/collision zones with exhumed HP and UHP metamorphic rocks. Therefore, recent “one-ocean” paleogeographic models for the Variscides (e.g. Kroner et al. 2007; Martínez Catalán 2011) appear less plausible. However, the number and width of Variscan basins are not relevant to the problem of orogenic topogra-phy and therefore not discussed any further in the present paper.

Subsequent to continental collision, Variscan orogenic topography is thought to have collapsed by thermal weak-ening of the lower crust (e.g. Malavieille et al. 1990; Burg et al. 1994; Costa and Rey 1995; Rey and O’Halloran 1997; Rey et al. 2001; ledru et al. 2001; Zulauf 1997). Heating has mainly been attributed to processes relating to plate convergence, such as thinning or loss of mantle lithosphere after subduction (see the systematic treatment in Henk et al. 2000 discussed below). However, recent studies in the metamorphic core complex of the Montagne noire (South France: Franke 2009b, c; Franke et al. 2011) have revealed that Carboniferous magmatism and high-temperature/low-pressure metamorphism in large parts of the Variscides are unrelated to Variscan crustal stacking.

In spite of some similarities, it is doubtful whether the Variscides have ever reached elevations of 5,000 m, which are widespread in the notional Tibetan equivalent. The Vari-scan foreland fold and thrust belts contain early Palaeozoic through to Carboniferous marine strata, and marine ingres-sions still occurred during deposition of late Carboniferous foreland molasse. Devonian to early Carboniferous marine sediments occur even within central parts of the orogen. For these reasons, Franke (2006, 2009b, c, 2012) has ques-tioned the existence of large Tibetan-style plateaus in the Variscides.

The first part of this paper provides a review of Devo-nian to early Carboniferous marine deposits within cen-tral parts of the Variscan orogen, where they overlie either pre-Variscan or early Variscan basement. Where data are available, derivation of syn-collisional clastic rocks is dis-cussed. Special attention is given to the Bohemian Massif. The review does not revisit the well-known Devonian and Carboniferous marine sequences of the accretionary belts listed above, which were deposited on the lower plates. However, part of these foreland rocks were later accreted to the upper plates, thus contributing to crustal thickening and uplift. Therefore, the present review also includes marine sequences deposited in the accretionary belts after the main phase of deformation, or else after significant stratigraphic breaks.

a brief paragraph discusses the sedimentary and pal-aeontological records of late Carboniferous intra-montane basins in the Variscides and their bearing on palaeo-altitude.

Possible geodynamic causes for low orogenic topog-raphy are discussed, with a focus upon anorogenic

lithospheric thinning and heating. The model derived has also important repercussions on the interpretation of “sub-duction-related” magmatism and the exhumation of HT/HP metamorphic rocks.

Correlation of the sedimentary and volcanic records with tectono-metamorphic processes is based upon the timescale of Gradstein et al. (2012). The areas addressed are marked, in Fig. 1, with numbers repeated in the text. For localities and units in the Bohemian Massif, see also Fig. 2.

Devonian through to early Carboniferous marine strata on pre‑Variscan and Variscan basement

Ossa Morena and Central Iberian zones (Iberia)

The Ossa Morena and Central Iberian Zones mainly consist of Cambrian through to Carboniferous strata deposited on a neoproterozoic (Cadomian) basement (see Dallmeyer and Martínez-García 1990; Gibbons and Moreno 2002). These sequences form part of fold and thrust belts, most of which were emergent in early Carboniferous time (Colmenero et al. 2002). Differences between the zones occur mainly during the Silurian and Devonian (García-alcalde et al. 2002).

In several synforms of the Ossa Morena Zone, Devo-nian to Carboniferous marine clastic rocks and subordi-nate carbonates unconformably overlie folded Devonian and older rocks. The Cabrela Fm. in the western part of the Ossa Morena Zone (no. 1 in Fig. 1) has been described by Pereira et al. (2012), Quesada et al. (1990) and Ribeiro (1983). It consists of early Carboniferous mudstones, greywackes, conglomerates, bimodal volcanic rocks and—near the base of the formation—limestones lenses, which have yielded Frasnian conodonts. The Cabrela fm. rests unconformably on folded and cleaved Eifelian limestones. The Terena synform, c. 80 km to the nE (no. 2 in Fig. 1), consists of terrigenous turbidites with few conglomeratic and volcanic intercalations. The clastic sequence grades upwards into shelf deposits with intercalations of late Tour-naisian and Viséan limestones, which suggests that the lower, main part of the Terena sequence is Devonian in age. Silva and Pereira (2004) have interpreted the Cabrela and Terena areas as transtensional basins. Famennian shales with limestone nodules have also been reported from the Valle syncline (no. 3 in Fig. 1). Marine deposits also occur in the early Carboniferous, above a stratigraphic break caused by erosion or non-deposition (Colmenero et al. 2002): in the Santos de Maimona (no. 4a in Fig. 1), Ber-langa (4b) and Terena-Cala synforms, late Tournaisian and/or Viséan marine carbonates associated with clastic rocks rest upon folded Variscan basement. In the Santos de Mai-mona synform, the Tournaisian-Viséan marine sequences are associated with volcanic rocks.


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a similar evolution is recorded in the south-western part of the Central Iberian Zone. The Obejo-Valsequillo-Puebla de la Reina syncline (no. 5 in Fig. 1) contains, above the mid-Devonian stratigraphic break, ?late Givetian shallow-water sandstones and Frasnian shales. Frasnian shales and late Famennian quartzites are known from outcrops in this syncline near Cordoba. late Devonian deposits are also known from the Sierra de San Pedro (no. 6), the almadén area (no. 7) and the eastern Sierra Morena (no. 8 in Fig. 1). The sequences are composed of shallow-water sandstones and shales, with important volcanic intercalations in the almadén area. S of almadén, Famennian shales and lime-stones extend upwards into the Tournaisian.

Carboniferous deposits make up the Guadiato belt (adja-cent to the late Carboniferous Peñarroya coal basin in the nE, both no. 9 in Fig. 1; see also Wagner 2004) and the Pedroches belt (no. 10 in Fig. 1, on both sides of the Pedroches batholith). These areas contain late Tournai-sian through to Viséan volcanic rocks, shales and lime-stones, with a basal clastic member in the Guadiato area. according to Quesada et al. (1990), all the Carboniferous sequences in the Ossa Morena and Central Iberian Zones belong to one basin, which straddled the tectonic boundary between these zones (Badajoz-Cordoba shear zone). How-ever, Wagner (1999, 2004) has stressed the importance of strike-slip processes for the formation of the Carboniferous basins in SW Iberia.

armorican Massif

South‑Armorican zone

Ducassou et al. (2011) have analysed Devonian sedi-mentary rocks in the St. Georges s. loire area (Upper allochthon of the South-armorican belt). In the southern part of the area (Châteaupanne unit), the neoproterozoic basement of the Mauges unit and its Ordovician cover is unconformably overlain by Emsian limestones and Emsian to earliest Eifelian immature sandstones (no. 11 in Fig. 1). This association of rocks has been overthrust by a northern (Tombeau leclerc) unit, which contains a latest Ordovician through to Emsian condensed sequence of shales and limestones. The Devonian marine strata are assigned to a back-arc basin, created by northward sub-duction of the Galicia-Brittany-Massif Central Ocean fur-ther south.

North‑central Armorica

Since the armorican Massif is well known, it is sufficient here to recall key points of its evolution (see also the review in Faure et al. 2005). The north-central armorican domain is framed by important shear zones and resembles a huge

tectonic clast embedded in a matrix of strongly metamor-phosed and deformed rocks (Fig. 1). It is bounded to the south by the high-grade metamorphic rocks of the South-armorican belt. To the nW, the armorican block abuts against the léon crystalline belt, a plausible candidate for correlation with the Mid-German Crystalline High (Faure et al. 2010b). a metagabbro from the numerous mafic rocks in the le Conquet-Penzé unit (set between the léon and the Central-/north-armorican block) has been dated, by la-ICP-MS, at 478 ± 4 Ma, which invites comparison with the Ordovician metabasic rocks of the Saxo-Thuringian rift (see reviews in Crowley et al. 2000; aleksandrowski and Mazur 2002).

The Cadomian basement of the north-central armori-can block is unconformably overlain by Cambrian through to early Carboniferous strata, weakly deformed and of very low metamorphic grade (see reviews in Doré 1994; Robardet et al. 1994; Pelhate 1994). The Devonian consists of shallow-water sandstones and shales, with intercalations of limestones, especially from late lochko-vian through to early Emsian and Givetian. The young-est Devonian beds preserved belong to the late Famenn-ian. a stratigraphic break in late Devonian time has been attributed to a “bretonic” tectonic phase, but a drop of sea level during the latest Devonian glaciation seems more likely (see the sea-level curve in Hacq and Schutter 2008). While the Devonian beds are widespread and similar in large parts of north-central armorica, early Carboniferous sedimentation is more localized in basins lined up along the north-armorican Shear Zone (from W to E: Cha-teaulin, Menez-Belair, laval, nos. 12a–c in Fig. 1). The highly individual character of the stratigraphic sequences suggests deposition in separate pull-apart basins (see e.g. Pelhate 1994; Rolet et al. 1994). The sequences consist of bimodal volcanic rocks and fluvial to lacustrine clastic rocks, partly with coal seams. Shallow-water limestones are widespread in the Viséan and locally occur also in the Tournaisian.

Morvan

apart from Variscan granites, the Variscan basement of the Morvan (no. 13 in Fig. 1) consists of anatectic gneisses with relicts of high-pressure rocks (Godard 1990). These crystalline rocks are interpreted as Eo-Variscan (e.g. Faure et al. 2009). This basement is overlain, in the south-western Morvan, by the Somme sequence. a comprehensive survey has been published by Delfour (1989), see also Mouthier in Feist et al. (1994). Conodont datings are available in Del-four and Gigot (1985). The oldest strata are Givetian lime-stones with the index brachiopod Stringocephalus burtini. Carbonates continue into the Frasnian, where they also con-tain resedimented conodonts of early- to middle-Devonian


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age. Weyer (1976) has published evidence of marine Tour-naisian deposits. The Famennian and basal Carboniferous consist of silty shales with intercalations of calc-alkaline and mafic volcanic rocks. Coarser clastic rocks make their appearance in the latest Devonian and continue through to the middle Viséan. Conglomerates contain pebbles of quartz, quartzite and black cherts. Marine conditions are documented by middle Viséan limestones. Delfour (1989) has described similar Carboniferous sequences from the Montagne Bourbonnaise (Pays d′Urfé, no. 14 in Fig. 1), to the SSW of the Morvan. Givetian–Frasnian limestones have not been detected in this area, but middle Viséan lime-stones are present.

Beaujolais and Brevenne areas

The Beaujolais area is situated to the SSE of the Morvan (no. 15 in Fig. 1). The earliest strata known are Givetian limestone pebbles in late Viséan conglomerates (Wick-ert 1988). The oldest coherent sequences occur in the le Goujet unit. Micaceous Goujet sandstones are overlain by dark limestones and calcareous sandstones, greywackes and pyroclastic deposits. The calcareous sandstones have yielded resedimented late Devonian conodonts (Vuag-nat 1984, cited in Wickert 1988) and Viséan brachiopods (lacroix 1980) and corals (Wickert 1985). a bimodal vol-canic sequence with calc-alkaline affinities occurs higher in the sequence (Sider and Ohnenstetter 1986). Wickert (1988) has described the le Goujet rocks as unmetamor-phosed and undeformed and suggested a sedimentary con-tact with the underlying affoux Gneiss of the Brevenne unit. Contrastingly, leloix et al. (1999) report bedding-parallel cleavage from these rocks and insist on a tectonic contact between the le Goujet and the Brevenne rocks. anyhow, a Devonian–Carboniferous marine basin must have existed in the root zone of the notional le Goujet allochthon.

The Brevenne unit (no. 16 in Fig. 1), which underlies the Beaujolais rocks, consists of metasedimentary rocks and a bimodal “spilite-keratophyre” volcanic sequence. Meta-morphism is of low to medium grade (Mouthier in Feist et al. 1994). The assignment of the volcanic sequence to the Devonian or early Carboniferous (see Sider and Ohnenstet-ter 1986, with earlier refs.) is supported by a U–Pb zircon age from a keratophyre at 365 ± 10 Ma (unpubl. report Milesi and lescuyer 1993, cited in leloix et al. 1999). However, Reitz and Wickert (1988) have described, from associated phyllites, organic-walled microfossils of late Proterozoic age. This conflicting information may be rec-onciled if the neoproterozoic metasedimentary rocks rep-resent the basement to the Devonian–Carboniferous cover. anyhow, further isotopic and biostratigraphic studies are necessary.

nE France/SW Germany

Belfortais and Southern Vosges

From a southern appendix of the Vosges, in the Chagey-Bel‑fort area (no. 17 in Fig. 1), Dressler (1989) has described marine strata of late Devonian to early Carboniferous age. Bituminous shallow-water limestones have yielded con-donts of the early Famennian. These are overlain by shales with holes left by weathered calcareous fossils and concre-tions (“schistes troués”). The shales have yielded Famenn-ian acritarchs and spores (Bain 1964). The overlying clastic sequence (shales, greywackes and quartz conglomerates) still contains Famennian organic-walled microfossils (Bain 1964), but also early Carboniferous conodonts (Creuzot 1983). The clastic strata are attributed by Dressler (1989) to proximal parts of submarine fans. Pebbles are predomi-nantly crystalline in the SW and more volcanic towards the nE.

The sedimentary belt of the Southern Vosges (no. 18 in Fig. 1) has recently been treated in much detail by Kre-cher et al. (2007), Krecher and Behrmann (2007), Krecher (2009) and Maass (2005). The oldest sedimentary unit (Markstein group) consists of greywacke turbidites and shales. From red and green shales with greywackes, Maass and Stoppel (1982) have reported Famennian conodonts. Two single zircon grains from a tuffaceous layer in these shales have been dated by U–Pb TIMS at 386 ± 1 and 386 ± 2 Ma (Schaltegger et al. 1996), which corresponds to a Givetian biostratigraphic age.

The Oderen Group to the SW of the Markstein sequence is attributed, by Krecher and Behrmann (2007), to a sepa-rate basin which was later emplaced in its present-day posi-tion by dextral transpression. It contains bimodal volcanic rocks extruded in basal Viséan time (U–Pb zircon from a rhyolite: 345 ± 2 Ma, Schaltegger et al. 1996). The lower part of the Oderen succession consists of mud-rich turbid-ites, with some limestone turbidites of arundian age (late early to early middle Viséan). During the Holkerian (mid-dle/late Viséan), turbidite sedimentation is replaced by nearshore and fluvial environments, which are overlain by outer shelf mudstones of the late asbian (late Viséan, V3b).

The basement underlying the Devonian–Carboniferous successions of the Belfort/Southern Vosges area is not exposed. Early Carboniferous volcanic rocks and Famen-nian carbonates in the Belfort area remind the Morvan (see above) and may be taken to suggest, by analogy, an early Variscan crystalline basement (pre-386 Ma, see above).

Southern Schwarzwald (Black Forest)

The zone of Badenweiler-lenzkirch (no. 19 in Fig. 1) is a belt of low- to very low-grade sedimentary and magmatic


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rocks set between the crystalline blocks of the Central Schwarzwald and the Southern Schwarzwald. as summa-rized in Güldenpfennig (1998), loeschke et al. (1998) and Maass (2005), it contains Ordovician and Silurian clastic metasediments, Frasnian shales and radiolarian cherts and late Devonian-Tournaisian greywacke turbidites. These are overlain by Viséan felsic volcanic and continental sed-imentary rocks.

The basement to the Badenweiler-lenzkirch zone is unknown. It cannot be represented by the Southern Schwarzwald crystallines adjacent to the south, because these rocks underwent high-grade metamorphism during the same time span in which the Badenweiler-lenzkirch sediments were being deposited (see the review in Maass 2005). as stated by Maass, the tectono-metamorphic evolution of the Southern Schwarzwald is very similar to that of the Central Schwarzwald. For these reasons, Franke (2000) has proposed that the Southern Schwar-zwald crystallines represent an eastern equivalent to the Central Schwarzwald rocks and was emplaced into its present position by a nE-trending dextral strike-slip fault (Fig. 1). a former eastern continuation of the Badenweiler-lenzkirch zone possibly exists in the hidden Permian–Carboniferous basins of northern Switzerland.

Basement of the Saar-nahe Basin (Germany)

The scientific drilling Saar I (no. 20 in Fig. 1) has encoun-tered, under the late Carboniferous and Permian clas-tic sediments of the Saar-nahe Basin, undeformed and unmetamorphosed rocks. The basement is represented by an albite granite dated at 381 ± 24 Ma (Rb/Sr whole rock, lenz and Müller 1976) and at 444 ± 22 Ma (Pb207/206 on zircon, Sommermann and Satir 1993). Franke and Oncken (1995) and Franke (2000) have assigned this granite to the magmatic arc formed by northward intra-oceanic subduc-tion of the Rheic ocean.

The granite is overlain by a middle-Devonian through to namurian sedimentary sequence (Krebs 1976; Clausen 2008): platform carbonates (Eifelian and Givetian), lime-stones, greywackes and fine-grained conglomerates with volcanic debris (Frasnian to basal Famennian), hemipelagic limestones and shales (Famennian), shales and limestones (Tournaisian) and black shales (late Tournaisian through to early namurian).

This marine sequence is in marked contrast to the Mid-German Crystalline High exposed in the Odenwald adjacent to the E, which represents an active plate mar-gin formed by subduction of the Rheno-Hercynian ocean (Engel et al. 1983; Weber and Behr 1983; Franke et al. 1995; Oncken 1997; Franke 2000).

Bohemian Massif

Delitzsch–Torgau–Doberlug syncline

The Variscan record of this area (no. 21 in Fig. 1) has recently been summarized by Gaitzsch et al. (2008a, b, 2010; see also biostratigraphic details in Heuse et al. 2010). The area is situated in the north-western part of the Saxo-Thuringian basin, to the SE of the Mid-German Crystalline High. low-grade Cadomian basement is overlain by weakly folded but unmetamorphosed Cambrian strata, which, in their turn, are overstepped by Carboniferous clastic and car-bonate sediments, with intercalations of volcanic rocks.

The Carboniferous of the Delitzsch area in the W has yielded late Viséan to early namurian floras and fau-nas. Brachiopods reported by Steinbach (1987, cited in Gaitzsch et al. 2010) indicate marine horizons, although the sequence as a whole is taken to represent alluvial fan and lake deposits. While most of the pebbles appear to be derived from local sources, metamorphic and granitoid clasts are referred to sources in the Mid-German Crystal-line High.

The Carboniferous of the Torgau-Doberlug area further E starts with greywackes, thin coal seams and carbonates (Finsterwalde fm.). The carbonates have yielded benthic faunas attributed to the asbian of the British zonation, which corresponds to the Belgian V3b and the crenistria goniatite zone of the German deep-water facies. The over-lying Doberlug and Kirchhain formations represent similar nearshore to shallow marine environments. The topmost Werenzhain fm. consists of fluvial conglomerates. Pebbles of Silurian cherts either reflect reworking of near-by units, or else could be derived from the Saxothuringian alloch-thons to the SE. Macrofloras suggest a late Viséan age also for the Doberlug, Kirchhain and Werenzhain formations (see Refs. in Gaitzsch et al. 2010).

Hainichen basin

The southern part of the Saxo-Thuringian Belt in nE-Bavaria and Saxony is characterized by the tectonic klippen of Münchberg, Wildenfels and Frankenberg, in which HP- to MP-metamorphic rocks override volcanic and sedimen-tary sequences of very low metamorphic grade (see Behr et al. 1982; Franke 1984a, b and reviews in Franke 2000; Klemd 2010). In this area (no. 22 in Fig. 1), the metamor-phic allochthon is overlain by diverse clastic sediments. The following survey of petrography, sedimentology and biostratigraphy is taken from Gaitzsch et al. (2008a, b, 2010; see also Schneider et al. 2005).

The oldest part (Striegis Fm) consists of turbidites and debris flow deposits attributed to a fan delta. There is no


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palaeontological dating. Detrital muscovites of three sam-ples record K/ar ages around 380 Ma (ahrendt et al. 2001), which could be derived from the near-by Franken-berg Klippe, but—in the opinion of this author—also from the north-western margin of the Teplá-Barrandean. Gran-ite pebbles in the Striegis formation have been dated at ca. 500 Ma (Gehmlich et al. 2000).

The clastic sediments and local coal seams of the overly-ing Ortelsdorf Formation are attributed to delta and flood-plain environments. Microfloras have been dated into the nM and nF spore zones, which corresponds to the V3b-cuIIIα and V3c-cuIIIβ-γ, respectively. a zircon Pb–Pb evaporation age of 330 ± 4 Ma has been reported from a tuff (Gehmlich et al. 2000). although an age of 330 Ma grossly corresponds to the Viséan/namurian boundary, the zircon age is compatible, within error bars, with the biostratigraphic findings.

The overlying Berthelsdorf formation consists of mud-stones, arkoses and coarse-grained conglomerates depos-ited in alluvial fans. Pebbles consist of granite, micaschist, quartzite and phyllite. K/ar ages of detrital muscovites mainly range between 333 and 324 Ma, with two sam-ples yielding 382 and 365 Ma (ahrendt et al. 2001). Peb-bles and the c. 333–324 Ma micas may be derived from the Granulitgebirge area to the n or the Erzgebirge to the S, while the Devonian mica ages point to sources in the near-by Frankenberg klippe or else at the north-western margin of the Teplá-Barrandean.

It is important to note that non-marine fish faunas have been retrieved from the top of both the Ortelsdorf and the Berthelsdorf formations. Gaitzsch et al. (2010) have argued that these fish were not able to migrate as far upstream as modern salmon and that the Erzgebirge basin was con-nected with the marine realm by a low-gradient drainage system. Hence, the Hainichen basin cannot have been high above sea level during the early Carboniferous.

Intra‑Sudetic basin

The Intra-Sudetic basin is situated in the West Sudetes, to the SE of the Izera-Karkonosze Massif (no. 23 in Fig. 1, see also Fig. 2). From the basal, north-westernmost part of the basin, Turnau et al. (2002) have retrieved miospores from mudstones intercalated among fluvial coarse-grained clas-tic sediments. The floras are not older than the TS miospore zone (equivalent to the British Holkerian and the Belgian V2b-V3a). The Paprotnia beds higher up in the sequence contain a highly diverse marina fauna including goniatites of the late asbian (Early Carboniferous IIIα crenistria goniatite zone, equivalent to the latest nM miospore zone). a bentonite layer has yielded volcanic zircons with a U–Pb SHRIMP age of 334 ± 3 Ma (see Kryza et al. 2011 for ana-lytical data and a review of earlier biostratigraphic studies).

Teplá-Barrandean block (TB) and areas accreted to it in time before c. 360 Ma

General features

The TB unit (no. 24 in Figs. 1, 2) has been taken to repre-sent, in late Devonian/early Carboniferous time, a Bohe-mian equivalent of the Tibetan high plateau with elevations of ≥5,000 m and an areal extent of only c. 175 × 100 km (see Dörr and Zulauf 2010, 2012 and references therein). This plateau is thought to have foundered, between 340 and 335 Ma, by thermal weakening of the lower crust. Subsid-ence is interpreted to have been located along a “Bohemian Shear Zone”, which bounds the TB to the SE, SW and nW. This interpretation has been questioned by Franke (2006, 2009a, b, 2012 and Franke and Żelaźniewicz 2000, 2002). The following paragraphs summarize and evaluate the doc-uments in the case, including few relevant points in Dörr and Zulauf (2012).

The TB occupies a position which is very similar to that of the north-central armorican Massif (Fig. 1). Both these massifs are bounded by Variscan magmatic and high-grade metamorphic rocks, which belong to the Saxo-Thuringian belt (to the n) and the S-armorican/Massif Central/Moldanubian (to the S). The Moldanubian is con-ventionally accepted to represent the easternmost part of this originally continuous, southward propagating accre-tionary belt (e.g. Matte 2001; Matte et al. 1990; Faryad and Kachlík 2013), which is supported by evidence laid out below. Discussion of alternative palaeogeographic concepts for the Moldanubian (e.g. Schulmann et al. 2009) is not rel-evant to palaeo-altimetry and therefore beyond the scope of this paper.

South‑central part: Barrandean Syncline

Similar to the north-central armorican block, the TB con-tains, in its south-central parts (the Barrandean Syncline in the Praha-Plzeň area), Cambrian through to late Givetian sedimentary and volcanic rocks, which rest unconformably on a Cadomian basement (see Chlupáč 1993a for a survey). The Cambrian through to early Givetian deposits record syn- and post-rift sedimentation. The Palaeozoic rocks have undergone little tectonic deformation. Metamorphism, in the Devonian rocks, occurred largely within the oil win-dow, with temperatures up to 180 °C (Suchy et al. 1996) or c. 130 °C (Glasmacher et al. 2002).

The reflection seismic line 9HR has traversed the TB c. 10 km to the W of the Palaeozoic cover, near Plzeň. according to Guy et al. (2011), the Proterozoic basement to the SW of the Barrandean Syncline is c. 18 km thick (from the top of the subduction-related rocks to a level close to the base of the Palaeozoic). The thickness of the


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Palaeozoic sequence (Cambrian through to mid-Devonian), following the cross-section (Fig. 3) in Chlupáč (1993a), amounts to c. 3,500 m. adding the average thicknesses of the individual formations (see the tables in Chlupáč 1993a) yields c. 7,300 m, and the maximum thicknesses of the formations add up to 11,000 m. Since neither maximum nor average thicknesses of the individual units will have occurred in one and the same area, the thickness derived from the cross-section is more realistic. If one generously starts from 5,000 m and adds 20 % for tectonic thickening (Franke 2012, calculated from Chlupáč 1993a), one arrives at 6,000 m of thickened Palaeozoic rocks, which, together with the 18 km of folded Cadomian basement, add up to a total of only 24 km for the deformed basement and cover of the Barrandean Syncline. Peak temperatures of the Barran-dean Devonian (≥130 °C, see above) translate, with a geo-thermal gradient of 30 °C, into a depth of burial of ≥4 km. Even if one adds several km of material eroded above the youngest (Givetian) beds, post-tectonic crustal thickness is in the order of c. 30 km. This estimate is in agreement with a continental block of average crustal thickness first extended by rifting and then partly re-thickened by weak tectonic shortening. Higher crustal thicknesses are plausi-ble for the tilted, western margin of the TB, where the Pal-aeozoic cover has been eroded.

Subduction-related rocks which underly the Cado-mian basement in the seismic and gravity cross-sections of Guy et al. (2011) cannot be regarded as the root of the TB crust, because isotopic and petrological data prove that the descent of Saxo-Thuringian crust was effected by conti-nental subduction. Subduction was active already before c. 400 Ma (the metamorphic age of the Münchberg eclogites, see the latest review in Klemd 2010), and continental UHP rocks of the Erzgebirge dated at 340 Ma have risen from the stability field of diamond (nasdala and Massonne 2000; Kotková et al. 2011). Rock units presently underlying the Cadomian basement of the TB either represent subducted Saxo-Thuringian rocks weighed down by the eclogitic root of the downgoing slab, or else were extruded by channel flow from mantle depth at or shortly after 340 Ma (see dis-cussion on potential causes for low orogenic topography below). In both cases, these rocks cannot have been part of a buoyant crustal root supporting a Tibetan-style plateau in late Devonian/early Carboniferous time. Deformation, metamorphism and uplift to the SW of the TB are younger

features (see the paragraph on the “Bavarian” domain below).

Northwestern part

Contrary to the Barrandean Syncline, the northwestern part of the TB (Teplá crystalline area) has been strongly deformed and metamorphosed up to eclogite facies between ≥380 and 365 Ma (Dallmeyer and Urban 1998, review in Hajná et al. 2012). It contains Cadomian base-ment and Cambro-Ordovician magmatic rocks relating to Saxo-Thuringian rifting (review in Hajná et al. 2012). Parts of this association of rocks are also represented in the tec-tonic klippen of Münchberg, Wildenfels and Frankenberg (survey in Klemd 2010), which were emplaced on low- to very low-grade rocks of the Saxo-Thuringian autochthon to the nW of the TB (Behr et al. 1982; Franke 1984a, b). Exhumation of pressure-dominated metamorphic rocks in the north-western TB and the Saxo-Thuringian klippen was effected, from the late Devonian onwards, by extrusion of a tectonic wedge from the subduction zone (Franke and Stein 2000). as discussed by Hajná et al. (2012), rocks of the north-western TB have also been transported towards the SE, by dextral transpression, over the low- to very low-grade south-eastern part of the TB (Barrandean Syncline, see Figs. 2, 3).

although the topographic elevation of the north-western TB is unknown, the area is well documented as the source area of the Famennian through to Viséan flysch clastics of the Saxo-Thuringian foreland basin (Franke 1984a) and the Barrandean Syncline. In early Famennian greywackes (c. 370 Ma), Schäfer et al. (1997) have detected heavy mineral associations derived from the Münchberg and related rocks and detrital zircon dated at 380 Ma. From the Tournaisian through to Viséan flysch, neuroth (1997) has reported detri-tal micas of c. 632–562, 504–492 and 410–370 Ma. The W- to WnW-derived, late Givetian clastic sediments of the Barrandean syncline contain, among older grains, detrital zircons dated at c. 390 Ma (Strnad and Michaljevič 2005). neuroth (1997) obtained, from the Givetian sediments, K/ar muscovite ages of 487 ± 10 Ma, probably derived from Ordovician magmatic rocks of the TB basement. Subsid-ence accommodating the late Givetian clastic sediments was probably caused by thrust loading from the nW.

In time around 340 Ma, HP and UHP rocks were extruded, from the Saxo-Thuringian subduction zone upwards and towards the nW, where they intruded, at c. 333 Ma, the upper crust of the foreland (DEKORP and Orogenic Processes Working Groups 1999; Franke and Stein 2000; Henk 2000 and Krawczyk et al. 2000). The present-day north-western boundary of the TB is a late fault system partly reactivated during formation of the Ter-tiary Eger Graben.

Fig. 2 Geological map of the Bohemian Massif, after Franke (2000), Franke and Żelaźniewicz (2000, 2002) with some details taken from Geological Map of the Czech Republic (Cháb et al. 2007). note that the contours of basement units to the nE of the Sudetic Marginal Fault-Bila Fault are interpreted from limited surface outcrops and drillings and that the basement highs in the Fore-Sudetic monocline are only known from deep drillings. The letovice and Staré Město units are interpreted as parts of the Saxo-Thuringian belt

◂


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South‑eastern part

Cadomian basement and Palaeozoic cover with stronger metamorphism and deformation have also been preserved at the southern margin of the TB, where they occur as roof pendants (“metamorphic islets”) within the Central Bohemian batholith (Figs. 2, 3; see Chlupáč 1989). The gneissic fabric of the Staré Sedlo and Mirotice metagran-ites (intruded at 373 ± 5 and 369 ± 4, Kosler et al. 1993; Kosler and Farrow 1994) and the foliation in the adjacent Palaeozoic rocks pre-date the contact metamorphic imprint of the batholith, which was emplaced in time between 354 and 337 Ma (review in Žák et al. 2005). With the possi-ble exception of one late magmatic body dated at 337 Ma, the Central Bohemian Batholith pre-dates formation and emplacement of the Moldanubian granulites (Gföhl unit; see again Žák et al. 2005).

Devonian MP metamorphism also occurs at the south‑western tip of the TB. In this area, the Cambrian neukirchen-Kdyně pluton (see review of Cambro-Ordo-vician magmatism in Dörr et al. 1998) is separated by a mylonite zone from the Moldanubian MP-micaschists of the Královský Hvozd to the SW (Fig. 2, see Scheu-vens 2002 for the tectono-metamorphic evolution and Reitz 1992 for the neoproterozoic/Cambrian and Silurian protolith ages). Weger and Masch (1999) studied amphi-bolite mylonites, which record southward reverse faulting and indentation of the stiff, mafic root of the neukirchen-Kdyně pluton into the rocks adjacent to the south in time around 380 Ma (K/ar hornblende). Indentation explains the curved trace of the TB boundary fault in this area. The adjacent micaschists have locally preserved similar ages (Rb–Sr white mica c. 373 Ma, Ihlenfeld et al. 1998, 1999).

This demonstrates that at least this south-western part of the TB margin originated already in early Variscan (late Devonian) times. Intercalation of ?Moldanubian (Gföhl-type) garnet pyriclasite into the shear zone (Bues and Zulauf 2000) is probably a later process, because metamor-phism of the Gföhl assemblage (around 340 Ma) clearly post-dates the late Devonian metamorphism and shearing in the surrounding rocks. The present state of the southern and western margins of the TB was achieved by still later, (semi-)brittle activities (Pitra et al. 1999).

a similar but better preserved section across the south-ern margin of the TB is exposed in the Kutná Hora unit further to the nE. It consists (in order from bottom to top) of Moldanubian rocks belonging to the Varied Group, mica schists, HP rocks of the Moldanubian Gföhl assemblage and equivalents of the Central Bohemian Pluton associ-ated with TB rocks of its frame (Fig. 3 of the present paper, Kachlík 1999). like in the neukirchen-Kdyně area, the Gföhl-type HP rocks are set between units deformed and metamorphosed in late Devonian time (the TB above, and micaschists below). The Chýnow micaschists (set between the Královský Hvozd and the Kutná Hora units, Fig. 2) are likewise associated with (overlain by?) fragments of Gföhl-type high-grade rocks.

It is obvious from the above that the South-Barrandean tectono-metamorphic belt must have occupied a much larger area and was later reduced by uplift and erosion (Fig. 3). This South-Barrandean belt is also documented in late Devonian K/ar ages of detrital micas in the Moravo-Silesian flysch (Schneider et al. 1999). The occurrence of these micas in the lower flysch reveals that the host-rocks occupied a high position in the tectonic edifice and were, therefore, eroded first. Erosion of the structurally deeper

Cadom.basement

TEPLÁ-BARRANDEAN (TB)

Central-Bohemian Batholith 350-337 Ma

granite 522 Madiorite 523

gabbro 511 Ma

Gföhl Unit

(UHP/HT)

Barrandean

very low grade

low - medium grade

(m

etam. islets)

metagranitoids

++ ++

MOL-

DANU-

B

IAN

mafic com

plex

Marianske Lazn +

++

Givetian throughCambrian

Hradec Kralové:Famennian marinetransgression

vvv518 - 513 Ma metagranitoids

380-365 Ma

Variscan overprint c. 375 Maincreasing increasingweak

++

Variegated Group

Monotonous GroupMP micaschists

(c. 380 Ma)

TB

NNW SSE

SA

XO-THURINGIAN

++

+

++

++

++

++

+

+

+

South-Bohemian Batholith 325-310 Ma

Fig. 3 Diagrammatic section across the Teplá-Barrandean. Inter-nal structure of TB after Chlupáč (1993a, 1994) and Hajná et al. (2012). Structure of the nnW margin of the Moldanubian after Bues and Zulauf (2000), Ihlenfeld et al. (1998, 1999), Scheuvens (2002), Weger and Masch (1999, SW corner of the TB) and Kachlík (1999,

Kutná Hora area). Intrusion ages of early Palaeozoic plutons after Dörr et al. (1998, 2002), Finger et al. (2009); see also review of intru-sion ages in Hajná et al. (2012). For Famennian transgression E of the Elbe fault zone, see refs. in text


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Moldanubian (including the Gföhl-type HP rocks) is doc-umented in detrital mica and pebbles in the upper (late Viséan) flysch of Moravia (Dvořák 1973; Fiala and Patocka 1994; Hartley and Otava 2001; Kotková et al. 2007; Sch-neider et al. 1999). another part of the debris might have been contained in the eroded uppermost part of the Barran-dean Syncline (see above).

Topography and thermal state

Contrary to Dörr and Zulauf (2010, 2012 and refs. therein), these tectonic and geochronological observations reveal that the Barrandean Syncline occupied, at least during mid-Devonian and probably also during late Devonian and early Carboniferous times, a low position relative to the north-western and south-eastern TB, which is due to weak collisional thickening of the Barrandean Syncline relative to the areas adjacent to the nW and SE (see the summary of events in Fig. 4). In fact, the Barrandean Syncline repre-sents the remnant, weakly deformed core of an originally much larger block, whose northern and southern margins were severely affected by crustal thickening (partly in a regime of dextral transpression), uplift and erosion. This is in accord with fission track studies on apatite from the Ordovician of the Barrandean syncline (Glasmacher et al. 2002), which reveal cooling below 80 °C at 324 ± 8 Ma (late Viséan/namurian). If these rocks had occupied the

roof of the Bohemian world already in the late Devonian, cooling should have occurred at this earlier time.

Delamination of mantle lithosphere (followed by plateau collapse and heating) is contradicted by Babuška and Plom-erová (2012), who present evidence for the preservation of pre-collisional mantle lithosphere under the TB. also, strong heating of the lower crust of the TB around 340 Ma by the ascent of asthenosphere (as to be expected from the model of Dörr and Zulauf 2010) should have left, in the tilted crustal profile of the western TB, a HT metamorphic imprint of that age, which is not detectable. In fact, this area has preserved a Cadomian MP-metamorphic profile (Dörr and Zulauf 2010), and late Devonian K/ar horn-blende ages have survived, e.g. in the southwesternmost corner of the TB (see the above text on the neukirchen–Kdyns area and further Devonian ages in the compilation of Hajná et al. 2012).

In order to explain the preservation of young, open marine sediments in the notional Bohemian high plateau, Dörr and Zulauf (2012) have referred to the Himalaya/Tibet example. However, the Tibetan plateau is characterized by widespread post-tectonic intra-montane molasse basins with terrestrial volcanic rocks of mid-Cretaceous age (Kapp et al. 2005) and fluvio-lacustrine clastic sequences extend-ing into the tertiary with thicknesses of several kms (Kapp et al. 2007), partly overlain by Miocene freshwater lime-stones (Wang et al. 2002). These conditions do not bear

Namurian

Stefanian

Westfalian

DE

VO

NIA

NC

AR

BO

NIF

ER

OU

S

Din

antia

nS

ilesi

an

Viséan

Tournaisian

Famennian

Frasnian

Givetian

Eifelian

Emsian

Pragian

Lochkovian

Late

Middle

Early

300

325

350

375

400

flysch sedimentation

(hem

i-)

pel

agic

sed

imen

tati

on

intrusion of granulitesfolding

vvvvvvvvv vvvv

SAXO-THURINGIAN FORELAND

?

intramontane basins

TEPLÁ-BARRANDEAN NW part

granulite facies metam.

deformation &MP-metamorphism

eclogitefacies

PRAHA SYNCLINE

synorogenic clastic sed.

deformation &metamorphism

hemipelagic carbonates

TEPLÁ-BARRANDEANSE-part / E of ELBE FZ

?

basal clast.

?

deformation &MP-metamorphism

Kłodzko

Železné Hóry

metamorphic islets

Hra

dec

Kra

lové

?granulite-

MOLDANUBIAN

facies metam.

Gföhl / Monot. + Var.

eclogitefacies metam.

gran. extrusion HT -metam.

mainly platform carbonates

1

2

3

4

5

6

78

13

10

11

14

12

14

Fig. 4 Correlation of tectonic and metamorphic events in the Teplá-Barrandean unit (and rocks accreted to it prior to c. 360 Ma) and units adjacent to the nW and SE. numbering of events refers to the relevant references. 1 Gaitzsch et al. (2008a, b, 2010), 2 Franke and Stein (2000), 3 Franke (1984a, b), Schäfer et al. (1997), 4 Falk et al. (1995), Franke (2000), 5 Franke and Stein (2000), Rötzler and Romer

(2010), 6 Dallmeyer and Urban (1998), 7 Klemd (2010), 8 Chlupáč (1993a), 9 Čech et al. (1989), Chlupáč and Zikmundová (1976), Zukalová (1976), 10 Ihlenfeld et al. (1998, 1999), Weger and Masch (1999), 11 Bederke (1924), Kryza et al. (1999), 12 Chlupáč (1994), 13 Chlupáč (1989), 14 review in Franke (2000)


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comparison with the early Palaeozoic marine syn- and post-rift sequences of the TB. The close causal relation between crustal thickening, uplift and erosion of upper plate rocks is much better exemplified in the alps. The austroalpine lid with its Variscan basement and Mesozoic sedimentary cover (well comparable with the TB) has only survived in the East: in this area, nS convergence has partly been com-pensated by orogen-parallel extension (Ratschbacher et al. 1989) so that crustal thickening, uplift and erosion were much less important than in the western alps.

Continuation of the TB in E Bohemia and the West Sudetes

The eastward continuation of the TB beyond the Elbe fault zone (no. 25 in Figs. 1, 2) yields definite proof of low ele-vation in late Devonian/early Carboniferous times. Contin-uation of the TB has correctly been acknowledged by Dörr and Zulauf (2010; Fig. 1), who figured Cadomian basement and its Palaeozoic cover in the greenschist-grade sequences of the Železné Mts. with the same colours as in the TB proper (see also Fig. 2 in the present paper). In this area, the Palaeozoic sequence extends upwards into the Silurian and possibly even into the early Devonian (Chlupáč 1993b, 1994). Deformation and metamorphism in the Železné Mts. represent an eastern equivalent of the South-Barran-dean tectono-metamorphic belt (see above). a drilling nE of Praha, S of Mlada Boleslav (Fig. 2), has encountered a lower Permian conglomerate with pebbles of limestones whose facies and faunas match Silurian and early Devo-nian (Pragian) limestones of the Barrandean Syncline (Zik-mundová and Holub 1965; Chlupáč 1994). This proves that also the very low-grade Barrandean strata extend eastwards beyond the Elbe fault zone.

This eastern continuation of the TB is also documented in drillings near Hradec Kralové, which have reached strongly deformed ?neoproterozoic and Ordovician rocks. The tectono-metamorphic grade and the position to the nE of the Železné Mts. suggest correlation with the South-Bar-randean tectono-metamorphic belt (see above). The base-ment rocks are unconformably overlain by a basal clastic member and Famennian through to early Carboniferous marine carbonates (Fig. 2, see Čech et al. 1989; Chlupáč and Zikmundová 1976; Zukalová 1976). This Devonian marine cover continues eastwards into the West Sudetes (see Fig. 2 and the refs. in Franke 2012). Pebbles of late Devonian to early Carboniferous marine limestones have also been described from Permian conglomerates south of the Karkonosze dome (Fig. 2; Martínek and Štolfová 2009). Marine sedimentation shortly after folding, meta-morphism, uplift and erosion was possibly brought about by crustal extension, which may have been caused by a change from transpression to transtension in the shear zones which frame the TB.

The age of Variscan deformation E of the Elbe fault zone is well constrained in the Kłodzko block. The early Variscan tectonic stack of the Kłodzko basement (alek-sandrowski and Mazur 2002) contains deformed and meta-morphosed early Givetian limestones (Hladil et al. 1999) and is unconformably overlain by the transgressive Łączna conglomerate and overlying Famennian marine limestones of the Bardo basin sequence (Bederke 1924; Kryza et al. 1999). These relationships date Variscan deformation of the Kłodzko basement as late Givetian to Frasnian, i.e. in the interval between c. 385 and 372 Ma. This is in agreement with the ar/ar ages on hornblende and white mica from the north-western part (see above) and the south-western corner of the TB W of the Elbe fault zone (380–373 Ma, Weger and Masch 1999; Ihlenfeld et al. 1998, 1999). Uplift and erosion in the Kłodzko block must have proceeded rap-idly and probably by the same process which fed the Saxo-Thuringian flysch (see above). Most of the eroded material has been transported into the forelands, i.e. radially away from the present-day core of the Bohemian Massif (see the above refs. to the Moravo-Silesian flysch and, for the Sudetes, Mazur et al. 2010).

another marker for the correlation of the TB segments on either side of the Elbe fault zone is the southern boundary fault of the TB (Central Bohemian shear zone in the present paper, Fig. 2), which is situated at the SE margin of the Cen-tral Bohemian pluton (see a detailed assessment in Žák et al. 2005). as correctly shown in Dörr and Zulauf (2010), this shear zone is continued eastwards in the Hlinsko Fault Zone (Pitra et al. 1994; Pitra and Guiraud 1996), which, in its turn, can be traced further towards Rychnov, S of the nové Město unit (Fig. 2, see also Franke and Żelaźniewicz 2002).

These observations clearly demonstrate that the TB extends beyond the Elbe fault zone (see also Chlupáč 1994; Kachlík 1999; Faryad and Kachlík 2013), and that this lat-ter fault zone is only of minor importance: Fig. 2 reveals dextral offset of c. 70 km at the south-eastern margin of the TB. Comparisons of the eastern TB with the Saxo-Thuring-ian domain—suggested by Čech et al. (1989) and Dörr and Zulauf (2012)—are not substantiated: the southern bound-ary of the Saxo-Thuringian zone is situated far to the n of Hradec Kralové, at the S flank of the Karkonosze Mts. (e.g. aleksandrowski and Mazur 2002), and late Devonian car-bonates in the Saxo-Thuringian belt are nowhere observed to overlie metamorphic basement formed in middle Devo-nian time.

The Famennian marine transgression over its eastern part reveals that the TB, in late Devonian and Early Car-boniferous times, did not represent a Tibetan-style plateau, but was positioned at least partly below sea level. If a high plateau had existed in the western TB, it should have shed large quantities of clastic sediments towards the E, which is demonstrably not the case—an aspect never explained by


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Dörr and Zulauf (2010, 2012, with refs. to earlier publica-tions by Zulauf).

With the Famennian marine transgression as an abso-lute altimeter, it becomes evident that the strong con-trast in metamorphic grade observed at the north-western (Saxo-Thuringian) and south-eastern (Moldanubian) mar-gins of the TB was not caused by subsidence of the TB, but by uplift of the surrounding areas. a tectonic model is sketched out in the paragraph on potential causes for low orogeny topography (see below).

South-western margin of the Bohemian Massif (“Bavarian” domain)

at the south-western margin of the Bohemian Massif (Fig. 2), the TB abuts, along the W-Bohemian shear zone (Zulauf 1997), against HT/lP metamorphic rocks and gran-itoids of the Oberpfalz (Fig. 2). This area has undergone the same HT/lP metamorphic episode (see below) as more south-easterly parts of the Bohemian Massif and, therefore, has traditionally been assigned to the Moldanubian tec-tono-metamorphic domain, although its palaeogeographic affinities are unknown. In the north-western part of this area, the HT/lP rocks form the core of a nnW-plunging antiform (Tanner and Behrmann 1995), which is flanked by the TB to the E and equivalent rocks of the zone of Erben-dorf Vohenstrauß to the W (ZEV, Fig. 2). Dörr and Zulauf (2010, 2012) include the Oberpfalz “Moldanubian” into the crustal domain that was first thickened and then extruded during collapse of the notional Tibetan-style plateau shortly after 340 Ma. In doing so, they ignore that the tectono-met-amorphic evolution along the south-western margin of the Bohemian Massif occurred in a HT/lP tectono-metamor-phic regime, which is clearly distinguishable, in space, time and style, from metamorphism and exhumation to the nW and SE of the TB (Franke 2000; Kalt et al. 2000).

• although two small intrusive bodies were already emplaced at c. 347 Ma (Dörr and Zulauf 2010), the main thermal event in the Bavarian domain is clearly younger. Klein et al. (2008) have listed U–Pb zircon and monazite ages from “Bavarian” magmatic and, even more important, metamorphic rocks, which reveal a sharply defined peak at c. 326 Ma (see also the review in Finger et al. 2009; Galadí-Enríquez et al. 2010). Fin-ger et al. (2007) have reviewed and acknowledged this “Bavarian” HT phase (330–315 Ma) and domain (along the SW margin of the Bohemian Massif), as a fully independent stage of the Variscan orogeny, to be dis-tinguished from a “Moravo-Moldanubian” phase (345–330 Ma) in the name-giving areas (see their Fig. 2). It is clear that metamorphic rocks and migmatites formed around 326 Ma cannot have been extruded by plateau

collapse shortly after 340 Ma, because they originated ≥10 Ma later.

• as pointed out by Finger et al. (2007), the Bavar-ian thermal event has overprinted, after intermittent cooling, south-western continuations of the Moravo-Moldanubian rock units, including high-grade Gföhl-type rocks formed around 340 Ma. Such “Bavarian” overprinting of older metamorphic phases has also been reported for more north-westerly parts of the Bavarian domain (Behrmann and Tanner 1997; Ihlenfeld et al. 1999; Siebel et al. 2012; Stein 1988; Teipel et al. 2004, 2012; Wagener-lohse and Blümel 1986). Behrmann and Tanner (1997) have characterized the HT/lP over-print in the Bavarian domain as a late orogenic thermal pulse, which essentially post-dates large-scale crustal deformation and tectonic transport. These observations support the sequence of events as documented by the isotopic ages (see the paragraph above).

• as shown in Fig. 2, the Bavarian HT domain extends from the south-western part of the S-Bohemian batho-lith even into the Frankenwald Transverse Zone (FTZ) in the north-western part of the Saxo-Thuringian belt, i.e. beyond the Saxo-Thuringian/TB suture (Franke 1984a; Kosakowski et al. 1999 and refs. therein). In a WnW-trending alignment of granite outcrops within the FTZ, Rotthaus et al. (2005) have dated the post-tectonic Henneberg granite at c. 325 Ma (U–Pb TIMS), and Kunert et al. (1998) have reported a U–Pb single grain zircon age of 320 +10/−8 Ma from the near-by Sil-berberg granodiorite. as it is apparent from Fig. 2, the Bavarian domain also extends towards the SE, so that it sticks out over and crosscuts the entire TB. This array clearly reveals that its evolution was independent from that of the latter.

• as documented in Behrmann and Tanner (1997) and Tanner and Behrmann (1995), synmetamorphic defor-mation (D3) in the Oberpfalz segment of the Bavar-ian zone (Fig. 2) occurred in a regime of compression with subordinate strike-slip. a subvertical stretching lineation with nE side-up shear sense indicates reverse faulting of the TB over the Moldanubian (instead of normal faulting in the model of Dörr and Zulauf 2010, 2012). a late increment in dextral transpression around 315 Ma has also been proposed by Galadí-Enríquez et al. (2010). according to Behrmann et al. (1994), the evolution of the Oberpfalz documents a positive thermal anomaly with short-lived “regional contact metamor-phism”, developed over a mantle high (Behrmann and Tanner 1997)—features incompatible with extrusion from an orogenic crustal root.

“Bavarian” detrital micas with K/ar ages around 320 Ma are first observed in the latest Carboniferous (Stefanian) of


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the intra-montane Weiden basin immediately W of the ZEV (Fig. 2) and occur throughout the Triassic of the Mesozoic foreland (Welzel 1991). There are no constraints on the topographic elevation of the Bavarian domain during and after its exhumation. However, all the above points consist-ently disprove exhumation of the Bavarian zone by plateau collapse shortly after 340 Ma.

Late Carboniferous/early Permian intra‑montane coal‑basins in Europe

It might be argued that high topography was attained late in the collisional history of the Variscides. It is therefore necessary to briefly recall the sedimentary and palaeonto-logical records of the late Carboniferous intra-montane basins, which are widespread in all parts of the orogen. For this purpose, it is sufficient to note that most of these basins contain coal seams, partly of economic significance.

as summarized in Colmenero et al. (2002), the oldest coal seams in the Ossa Morena Zone of Iberia have been detected in the Tournaisian of the S. de Maimona basin (no. 4a in Fig. 1), the Valdeinfierno basin (along-strike to the SE of the Berlanga basin, 4b in Fig. 1) and the Bena-jarafe basin (E of the Berlanga basin). Coal also occurs in the namurian of the Berlanga basin. In the Central Ibe‑rian Zone, coal is widespread in the Westphalian of the Sierra de San Pedro and Peñarroya basins (nos. 6 and 9 in Fig. 1). Wagner (1999, 2004) has interpreted the Carbon-iferous basins in SW Iberia as strike-slip related. Pelhate (1994) has summarized the mid-/late Carboniferous coal-bearing basins formed along the South-Armorican shear zone. a review of the strike-slip-related coal-basins in the French Massif Central is available in Echtler et al. in Feist et al. (1994). In most of these basins, deposition started in Stephanian time and often extended into the lower Per-mian (autunian). The Permian–Carboniferous Saar-Nahe basin formed S of the Rheno-Hercynian suture (no. 20 in Fig. 1) has coal seams extending from the early Westphal-ian into the lower Permian (autunian) (Schäfer 2011). The Saale basin SE of the Harz Mts. represents a north-east-ern equivalent of the Saar-nahe basin. Coal seams occur repeatedly from the Viséan through to the Stephanian (see Schneider in McCann et al. 2008; Schneider and Romer 2010). In the Bohemian Massif, coal-bearing sequences are known from the Westphalian to early Permian of cen-tral and western Bohemia nW and n of the Barrandean syncline, Fig. 2; Svoboda et al. 1966). In the Intra-Sudetic basin (no. 23 in Fig. 1, see also Fig. 2), coal seams occur from the latest Viséan through to the middle Westphalian (Dziedzic and Teisseyre 1990).

The rich faunas and floras of these coal-bearing basins are not compatible with high and dry plateaus. Therefore,

Roscher and Schneider (2006) concede alpine elevations only for limited areas and strongly argue against a coher-ent, high Variscan mountain chain. It is also important to note that the Permian–Carboniferous floras of Europe are rather uniform and do not reveal evidence of an altitudinal zonation (H. Kerp, Münster, personal comm. 2013). Since part of these floras occur in paralic foreland basins with marine ingressions, the uniformity of the floras suggests rather low topography throughout the Variscides.

Curiously, it is from such intra-montane coal-bearing basins that two older publications have reported notional evidence for glaciation: Franke (2012) has discussed and dismissed indications of a Permian–Carboniferous gla-ciation in the Massif Central (Becq-Giraudon et al. 1996). likewise, a report on striated boulders in the late Carbon-iferous Peñarroya basin in Spain (Delgado et al. 1980) has not been substantiated (andreis and Wagner 1983).

Potential causes for low orogenic topography in the Variscan belt

Tectonic causes

Thrust loading

During continental collision, the lower plate is weighed down by the load of the accretionary wedge and the sedi-mentary prism of the foreland basin. This mechanism accounts for repeated marine flooding still in the late-Vari-scan, coal-bearing molasse basins. likewise, backthrusting within the upper plate may also effect subsidence in the orogenic hinterland. This process may have aided subsid-ence of the central TB below sea level (Table 1).

Convergent thrust loading by the north-western part of the Teplá-Barrandean and by backthrusting of the Mid-Ger-man Crystalline High (Schäfer et al. 2000) may also have helped to create the Carboniferous basins of the Delitzsch-Torgau-Doberlug and Hainichen areas. Thrust loading may also be responsible for Devonian-Carboniferous subsidence of the Belfortais-Southern Vosges and Schwarzwald. These areas have either been assigned to a foreland basin relat-ing to a n-dipping subduction zone (loeschke et al. 1998; Güldenpfennig 1998) or to a retro-arc wedge impinging on a back-arc basin formed by southward subduction of the Rheno-Hercynian ocean (Krecher and Behrmann 2007).

Extension in strike‑slip zones

Collisional shortening is usually compensated by crus-tal thickening and uplift, unless these processes are counter-balanced by a component of orogen-parallel extension. Extension may take place by penetrative


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ductile transpression, which displaces material laterally, thus reducing the amount of crustal thickening and uplift. lateral extension may also be effected by discrete strike-slip faults set en-echelon, by the lateral extrusion of tec-tonic wedges bounded by conjugate strike-slip faults, or else be localized in pull-apart structures.

all these processes require space in which the laterally extending material is accommodated. This is exemplified in eastern asia and in the eastern alps, where lateral escape has been facilitated by subduction rollback to the east of the contracting orogen (Schellart and lister 2005; Scharf et al. 2013). as discussed in Franke and Żelaźniewicz (2002), orogen-parallel extensional belts and strike-slip faults in central Europe cannot have operated in their present-day position, in which they abut against the south-western mar-gin of Baltica and are, therefore, blocked (see Figs. 1, 2). However, balancing of contraction in the mid-European Variscides in time after c. 330 Ma yields shortening of con-tinental crust in the order of 1,000 km. Shortening of the extended northern margin of the Rheno-Hercynian basin makes up an important part of this balance (Doublier et al. 2012). Palinspastic restoration for the Early Carboniferous places the Variscan terranes in a position to the SE of Bal-tica, where dextral strike-slip between laurussia and Peri-Gondwanan had free play.

In the Variscides, the importance of dextral strike-slip has always been stressed, from the benchmark paper of arthaud and Matte (1977) to the endmember model of Martínez Catalán (2011), in which the complexity of the Variscan belt is explained by strike-slip re-shuffling of one single continental margin. Dextral strike-slip is wide-spread in Spain (e.g. Silva and Pereira 2004; Wagner 2004), armorica (Pelhate 1994; Rolet et al. 1994), in the Mid-German Crystalline High (Oncken 1997), in the northern part and along the northern margin of the Teplá-Barrandean (Franke 2000; Hajná et al. 2012), and along the “Moldanu-bische Überschiebung” of Suess (1903), which forms the boundary between the Moldanubian and Moravo-Silesian units (Gayer and Schulmann 2000). all these strike-slip fault zones will have helped to reduce the amount of crustal thickening.

Subsidence by lateral extrusion of a tectonic wedge has been proposed, for the basement of the Saar-nahe basin, by Weber (1995). areally limited extension may also occur in pull-apart structures even in an overall transpressive setting (see the case of the Montagne noire: Franke et al. 2011, or the Maures-Tanneron Massif, Corsini and Rol-land 2009). Pull-apart extension formed in transpressive

Table 1 Sedimentary basins discussed in the text (numbers as in text and Fig. 1) and their tectonic settings according to literature cited in the text. Settings preceded by a question mark are proposed in this paper. note that “strike-slip” is meant to also include transpression (with pull-apart windows) and transtensiont

Basin Tectonic setting

Iberia

1 Cabrela Strike-slip

2 Terena Strike-slip

3 Valle Strike-slip

4a Santos de Maimona Strike-slip

4b Berlenga Strike-slip

5 Obejo-Valsequillo-Puebla de la Reina

Strike-slip

6 Sierra de San Pedro Strike-slip

7 almadén Strike-slip

8 eastern Sierra Morena Strike-slip

9 Guadiato-Peñarroya Strike-slip

10 Pedroches Strike-slip

11 South armorica(St. Georges s. loire)

Back-arc, ? subduction ero-sion

north-central armorica

12a Chateaulin ? Waning rift, strike-slip

12b Menez-Belair ? Waning rift, strike-slip

12c laval ? Waning rift, strike-slip

Massif Central

13 Morvan arc, back-arc, ? strike-slip

14 Montagne Bourbonnaise Back-arc, ? strike-slip

15 Beaujolais Back-arc, ? strike-slip

16 Brevenne Back-arc, ? strike-slip

nE France/SW Germany

17 Chagey-Belfort 17–19:

18 Southern Vosges Retro-arc thrust loading, strike-slip

19 Southern Schwarzwald ? Foreland basin, strike-slip

Rhenish Massif

20 Saar-nahe: Devonian - early Carb.:

late Carb.:

Subduction erosionWedge extrusion, strike-slip

Bohemian Massif

21 Delitzsch-Torgau-Doberlug ? Thrust loading

22 Hainichen ? Thrust loading

23 Intra-Sudetic basin Strike-slip

24, 25 Teplá-Barrandean, pre-tectonic

(W and E of the Elbe fault zone)

Waning rift, thrust loading

post-tectonic marine sedi-ments

? Strike-slip (transtension)


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or strike-slip regimes produced sedimentary basins, parts of which were even marine. This applies to the Tournai-sian through to Westphalian basins in Iberia (Silva and Pereira 2004) and in the armorican Massif (Pelhate 1994). Wrench-related subsidence has also been considered for the Tournaisian/Viséan clastic sediments and volcanic rocks of the Southern Vosges and Schwarzwald (Echtler and Chauvet 1992; Krecher and Behrmann 2007). The Intra-Sudetic basin provides an example of a partly marine pull-apart basin formed at a transverse fault zone (Franke and Żelaźniewicz 2002).

Dextral transtension is widely accepted for the late- and post-Variscan evolution: in Carboniferous and Permian time, Gondwana was translated westwards with respect to laurussia and rotated clockwise, so that the Pangaea B sce-nario changed into Pangaea a, thus propagating the Tethys ocean westwards (e.g. arthaud and Matte 1977; Matte 1986; Henk 1999; Muttoni et al. 2003, 2009).

Subduction of continental crust

Metamorphic petrology has clearly documented ultra-high metamorphic pressure in Variscan continental rocks, including—in the Saxo-Thuringian Erzgebirge—meta-morphic diamond (nasdala and Massonne 2000; Kotková et al. 2011). These UHP rocks must have been formed at depths of ≥150 kms, which exceeds by far any plausible crustal thickness of an orogenic belt. Consequently, these rocks must have been subducted into the asthenospheric mantle, which implies that the bulk density of the descend-ing lithospheric slab effected negative buoyancy. Subduc-tion produces uplift of the upper plate only in special cases, when sediment is detached from the downgoing slab and accreted to the base of the upper plate, or if—like in the andes—the orogen is backthrusted over the upper plate (e.g. Oncken et al. 2006). a detailed discussion of uplift and extrusion models goes beyond the scope of this paper. It suffices to point out the results laid out in DEKORP and Orogenic Processes Working Groups (1999), Franke and Stein (2000), Henk (2000) and Krawczyk et al. (2000): Saxo-Thuringian rocks subducted southwards under the TB were detached from the downgoing slab at mantle depths, expelled upwards, by buoyancy and hydraulic forces, via the subduction channel and eventually into the upper crust of the foreland. Rötzler and Romer (2010) have recently supported the extrusion model and confirmed the timing of intrusion into the foreland. Franke (2006) has expanded this concept also to the southern margin of the TB (see also the extrusion models of Schulmann et al. 2009 (with refs.) and Babuška and Plomerová 2012).

The extrusion model of Dörr and Zulauf (2010, 2012) for rocks dated at 340 Ma is basically similar. However, it does not account—in spite of assurances to the contrary—for

the ultra-high metamorphic pressures observed in these rocks but proposes, instead, channel flow from the base of a thickened, Himalaya-type crustal root supporting high surface elevation (see Jamieson and Beaumont 2013 for numerical modelling). as demonstrated above, the Hima-laya/Tibet model for Bohemia is proven wrong by geologi-cal observations. It is also unnecessary for the explanation of the exhumation history observed, because felsic rocks metamorphosed in the upper mantle are even hotter and less viscous than those in a crustal root, profit from higher lithostatic pressures for their expulsion from the mantle as well as from buoyancy during their return through the man-tle, and also effect, after exhumation, a more drastic pres-sure contrast to the adjacent TB upper plate.

One might regard subduction of continental crust to mantle depth as a rare case. However, the metamorphic pressures recorded from HP und UHP rocks are always minimum pressures. Besides, evidence for HP and UHP conditions may only survive, under special conditions, in miniature strongboxes such as zircon, garnet and pyroxene grains (see review in Ernst 2006). Hence, we may expect that rocks formed under high and ultra-high pressures were originally much more frequent and more voluminous than the scarce relicts that have escaped from retrogression.

Subduction erosion

Subduction erosion is a widespread phenomenon (see review by Stern 2011). In such a scenario, the upper plate is thinned by tectonic abrasion at its base, which causes sub-sidence. Subduction erosion occurs in high-stress regimes caused by high spreading rates. Tectonic erosion is also enhanced by high topographic roughness of the subducted oceanic plate and the absence or limited thickness of over-lying sediments which otherwise lubricate the subduction zone (e.g. Kukowski and Oncken 2006). The marine (partly deep-water) sediments of the Cretaceous Intra-alpine Gosau basin have also been explained by subduction ero-sion (see e.g. Wagreich 1995; Wagreich and Decker 2001).

In the Variscides, subduction erosion has been invoked, e.g. by Oncken (1997) for the Mid-German Crystalline High and by lardeaux et al. (2001) for the Massif Central. Since there is no evidence to support high spreading rates in the Variscan oceans, one has to consider an alternative explanation: high shear stress in the subduction zone may also be effected by low negative buoyancy of young, hot and, therefore, less dense oceanic lithosphere. Hot litho-sphere is quite plausible for the oceanic corridors which separated the Variscan microcontinents (i.e. the Rheno-Hercynian, Saxo-Thuringian and Galicia-Massif Central-Moldanubian oceans), because they were young and rather narrow: they are neither documented by palaeomagnetism nor by palaeogeography (see, e.g. discussion in Franke


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2000; Robardet 2003). For these reasons, subduction ero-sion by the downgoing plate is a possible explanation for low topography in zones of Variscan subduction or early collision, although the surface topography of the downgo-ing slab and the amount of lubrification by subducting sedi-ments remain a matter of speculation.

Back‑arc spreading

numerous publications have proposed back-arc extension to explain subsidence, marine sedimentation and volcanism within the upper plate of subduction systems (e.g. Cogné 1990 and Ducassou et al. 2011 for armorica, Pin and Paquette 2002 and Matte 2007 for the Morvan-Beaujolais, Krecher et al. 2007 Krecher and Behrmann 2007) for the Southern Schwarzwald and Vosges, Schulmann et al. 2009 for the Moldanubian). However, this scenario implies old, cold and dense oceanic lithosphere whose steep descent effects subduction rollback. The limited width and young age of Variscan oceans (see the paragraph above) rather imply hot lithosphere with low negative buoyancy.

The only exception is the Rheic ocean between avalo-nia and the armorican microcontinent(s), which was closed already in early Devonian time (Franke 2000). Since this ocean is well documented by both biogeography and pal-aeomagnetism (Cocks and Fortey 1982, 2011; Fortey and Cocks 2005; Tait in Franke et al. 2004), its older, colder and denser parts could have been steeply subducted northwards (Franke and Oncken 1995), thus effecting early Devonian back-arc spreading within avalonia (Franke 2000). South-eastward subduction of the Rheic mid-ocean ridge under the northern margin of armorica might have occasioned, in Emsian time, opening of the Rheno-Hercynian (lizard-Gießen) basin (Franke 2006) as a “successor ocean” to the Rheic. at an earlier time, south-eastward subduction of the Rheic might also have produced opening of a Saxo-Thuringian back-arc basin within the north-western part of armorica.

Thermal weakening of the crust leading to “orogenic collapse”

numerous authors have proposed that the Variscan oro-gen collapsed after crustal thickening and heating, with the combined effects of a new, level crust/mantle bound-ary and a much reduced topography (Costa and Rey 1995; Rey et al. 2001). Henk et al. (2000) have pointed out wide-spread heating from c. 340 Ma onwards and provided a sys-tematic discussion of potential heat sources. The ascent of a plume and break-off of the subducted slab are considered unlikely. Heat is thought to have been advected to the crust by the reduction or loss of lithospheric mantle from the base of the upper plate. according to Henk et al. (2000),

this may be explained in several plausible ways: loss of the thermal boundary layer, delamination of mantle lithosphere or lithospheric extension. Subduction of lithosphere is also considered favourably, since it prevents the formation of thickened lithosphere, which would effectively insulate the crust against the asthenosphere. Heating by lithospheric extension is only acknowledged as a post-orogenic pro-cess (≤305 Ma). Radiogenic heating in thickened crust is accepted as an inevitable, additional cause (see also Gerdes et al. 2000).

all the causes considered most plausible by Henk et al. (2000) are directly related to subduction/collision pro-cesses. a detailed discussion of the individual aspects is not necessary, since it will be shown below that timing and areal extent of granitoid magmatism and high-temperature metamorphism from c. 340 Ma onwards do not agree with ready-made tectonic models of convergent processes.

Variscan HT metamorphism and associated magmatism: constraining facts

The following points are vital for the understanding of the HT evolution:

1. It is important to recognize zones of high heat-flow cutting across the structural trend of the orogen. One example is the nS-trending lippstadt-Ramsbeck trans-verse zone in the north-eastern part of the Rhenish Massif (Franke et al. 1995; Hoyer et al. 1974; Teich-müller and Teichmüller 1982; de Roo et al. 1992; Weber 1972). Vitrinite reflectance and illite crystallin-ity in this area reveal exceptionally high anchizonal metamorphism, which permitted ductile deformation of reef limestones and intense slaty cleavage in late Carboniferous time (ahrendt et al. 1983). The Bavar-ian Zone along the south-western margin of the Bohe-mian Massif (see the text above and Fig. 2) represents another, much larger example of a “hot” transverse zone. Both the Bavarian transverse zone and the nW-trending Elbe fault zone (Fig. 2) also contain bodies of ultra-potassic mafic melt (“Durbachites”), usually interpreted as being derived from metasomatized lith-ospheric mantle in subduction zones (see discussion and map in Finger et al. 2007).

a transverse boundary between contrasting thermal regimes has also been described by Costa (1992). The tectono-metamorphic evolution in the western part of the Massif Central (limousin) occurred during 390–370 Ma and was pressure-dominated, while the eastern part was formed at 360–340 Ma in a temperature-dom-inated regime. The boundary between these domains is marked by the nnE-trending Sillon Houiller fault,


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which transects the entire Massif Central. Interestingly, the Sillon Houiller and a north-easterly branch fault (Tauve-aigueperse fault zone, Fig. 1) mark an impor-tant boundary in the seismic anisotropy pattern of the mantle lithosphere, which is inherited from Variscan processes (Babuška et al. 2002).

2. at least some of the Carboniferous HT events are areally widespread. Simancas et al. (2003, 2006) have recorded, from the IBERSEIS reflection profile, a mid-crustal, sill-like, highly reflective band of 1–2 s thickness, which extends for 175 km length across the structural trend of the Ossa Morena and Central Iberian Zones. The authors propose that this band represents mafic intrusions result-ing from the interaction between the collisional orogen with a mantle plume, starting at about 360–350 Ma. Simancas et al. (2003, 2006) attribute, to this phase, crustal extension, bimodal volcanism and mineralization in the South Portuguese, Ossa Morena and Central Ibe-rian Zones of the Iberian Variscides.

The South Portuguese Zone has since long been acknowledged as the south-westernmost segment of the Rheno-Hercynian belt, which is continued in SW-Eng-land, the Rhenish Massif, the Harz Mts. and Moravo-Silesia (Engel et al. 1983; Kossmat 1936; lotze 1963; Oliveira et al. 1979; Suess 1888). This well-known correlation is based, among other criteria, upon mid-/late Devonian through to Viséan within-plate metaba-salts (see Floyd 1995 for a geochemical survey of the Rheno-Hercynian segment).

Geophysical profiling in the Variscides of central Europe has revealed features, which closely resemble the Iberian profile. Reflection seismic profiling in the Rhenish Mas-sif (DEKORP Research Group 1990) reveals a band of reflections at c. 5–6 s two-way travel time with vp around 7.2 (Giese et al. 1990), which merges southwards with the top of the reflective lower crust of the Saxothuring-ian and Moldanubian Zones (DEKORP Research Group 1985; Behr and Heinrichs 1987). This analogy with the Iberian seismic profile might be taken to reflect the same magmatic processes as those inferred for Iberia.

Faure et al. (2010a) have reviewed Viséan migmatites in the French Massif Central and referred them to the same magmatic phase as volcanic rocks in the armori-can Massif and in the north-eastern Massif Central. These volcanic rocks have been pointed out also in the present paper, in connection with Carboniferous marine sedimentation (see the text above). Faure et al. (2010a) envisage, in late Viséan time, a shallow, flat crust/mantle boundary and thinned lithosphere under the Massif Cen-tral. They explain these features with the loss or thinning of the lithosphere and ascent of asthenosphere to the base of the upper plate above S-dipping subduction zones.

Discussion of this plate-tectonic scenario goes beyond the scope of this paper, but Faure et al. (2010a) have cor-rectly acknowledged the large areal extent of the Carbon-iferous HT regime in France.

3. Franke (2009b) and Franke et al. (2011) have demon-strated, in the Montagne noire (S France), that HT met‑amorphism and magmatism in the S-Variscan foreland commenced already before the onset of crustal thicken‑ing and continued into the early Permian. although the new timescale by Gradstein et al. (2012) has recently shifted the Viséan/Serpukhovian (namurian) bound-ary to 330.9 Ma, the fundamental problem of the Mon-tagne noire remains: the onset of HT metamorphism (≥335 Ma) pre-dates the deposition of the youngest flysch sediments (namurian a), and the younger incre-ments of the HT evolution (c. 300 Ma and early Per-mian) clearly post-date the Variscan orogeny. Compari-son with the Pyrenees, the Massif de Maures and the Massif Central reveal that these thermal events are are-ally widespread.

The intrusion of the Meissen pluton in the Elbe fault zone (see above) tells a story similar to that of the Montagne noire, because the area belongs to the lower plate (foreland) of the Saxo-Thuringian bel, and crystallization of the Meissen monzonite (c. 335 Ma, review in Förster and Romer 2010) occurred in mid-/late Viséan time, i.e. still during or shortly after flysch sedimentation (biostratigraphy in Heuse et al. 2010). These observations rule out formation in a subduction/collision zone.

4. The timing of magmatism and HT metamorphism in the Montagne noire and neighbouring areas also reveals that these processes are not continuous, but occur in discrete pulses at ≥333 Ma, c. 308, c. 300 and in post-Stephanian time (i.e. ≤299 Ma). Similar age clusters of granite intrusions occur in the Massif Central, in the Pyrenees and in the Massif de Maures (Franke et al. 2011). It has to be pointed out that it was Schaltegger (1997) who first defined similar magmatic pulses in the Vosges, Schwarzwald and the alpine External Massifs, at 340–330, 310–307 and 304–294 Ma. The c. 300 Ma thermal pulse has been inter-preted by Costa and Rey (1995) as the time of regional thermal collapse of thickened crust and establishment of a new, shallow and flat crust/mantle boundary. It is important to note that U–Pb TIMS and ar/ar dates from alkaline and tholeiitic basalt lavas in in the Mid-land Valley of Scotland reveal quite similar pulses at c. 343, c. 335, c. 308 and 298–292 Ma (Parrish 2006). The authors relate volcanism and crustal extension to the dextral strike-slip system between Gondwana and


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laurussia (arthaud and Matte (1977), which is also responsible for dextral faulting in the Montagne noire. Schaltegger and Brack (2007) have analysed Permian volcanic rocks erupted between 285 and 275 Ma. They propose that the felsic rocks originated from heat advection during mafic underplating and that the mafic component is comparable to the subcontinental mantle of the incipient Tethys rift. The timing of magmatism in Scotland suggests that Permian magmatism only represents the youngest and most widespread incre-ment in polyphase anorogenic lithospheric heating and thinning, which commenced already in early Carbonif-erous times.

Discussion

low orogenic topography

The evidence laid out above reveals marine sediments of mid-Devonian through to early Carboniferous age, which have been deposited either on pre-Variscan basement or on exhumed early Variscan magmatic and metamorphic rocks. Together with the well-known external fold and thrust belts of the Variscan orogens with their complete sedimentary sequences, these findings rule out one coherent Tibetan-style plateau of late Devonian through to early Carbonifer-ous age in the Variscides. In late Carboniferous and early Permian times, the Variscides were transected by intra-montane, mostly coal-bearing basins whose floras and fau-nas are not compatible with high and dry plateaus.

Major thickening and surface uplift may only have occurred in rather narrow belts along crustal-scale shear zones or active margins, in which Devonian-Carbonifer-ous overstep sequences have not been detected: these are the Beja-aracena Belt in south-west Iberia and its pos-sible equivalents (central parts of the léon domain in north-western Britanny, normannian High in the English Channel and the main part of the Mid-German Crystalline High), the Galician-Castilian zone, the South-armorican zone, the Massif Central S of the Brevenne unit, the crys-talline central parts of the Vosges and Schwarzwald, the nW and SE margins of the Teplá-Barrandean block and the Moldanubian zone adjacent to the South. all these struc-tural highs are also known as the source areas of synoro-genic clastic sediments deposited in foreland basins (see Hartley and Otava 2001; Schneider et al. 1999 for Moravo-Silesia, Pastor-Galán et al. 2013 for n-Iberia and reviews in Engel and Franke 1983; Franke and Engel 1986). This is especially well documented in the Bohemian Massif, where petrography and isotopic ages of clasts and minerals can be tied to adjacent synsedimentary uplifts (Schäfer et al.

1997; Schneider et al. 1999). However, even in these areas, topography may have been low, if erosion controlled and kept pace with uplift.

The absence or scarcity of high Variscan topography is by no means trivial, because field evidence and seis-mic surveys have documented long-distance overthrust-ing of microcontinents over adjacent basins and foreland blocks: Variscan shortening of continental crust in the German-Czech transect exceeds 1,000 km (see Franke and Żelaźniewicz 2002; Doublier et al. 2012).

Plausible causes

Plausible tectonic causes discussed above comprise thrust loading, orogen-parallel extension, subduction of continen-tal lithosphere (which, in itself, does not produce buoyant uplift) and subduction erosion. Crustal thinning by back-arc extension is unlikely, because the young and hot litho-sphere of the narrow Variscan oceans is improbable to have caused subduction roll-back. Since orogen-parallel strike-slip fault systems are ubiquitous in the Variscides, orogen-parallel extension may be regarded as an important process, which counter-balanced crustal thickening and uplift.

Most marine basins are formed by net crustal extension. Extension may occur in transtensional windows within an overall transpressive regime (e.g. Montagne noire, see above), or else in transtensional settings.

Plausible tectono-thermal causes discussed above lead to thermal weakening of deeper parts of the crust. Volcan-ism as an indicator of heat advection to the crust occurs already in some Devonian basins in Iberia, and most of the Tournaisian/Viséan marine basins in Iberia and armorica contain bimodal volcanic suites, which indicate that subsid-ence was related to activities in the mantle. Devonian and/or early Carboniferous basaltic volcanism is even known from the Rheno-Hercynian foreland fold and thrust belt in South Portugal (Colmenero et al. 2002), Southwest Eng-land (leveridge and Hartley 2006), Germany (Floyd 1995) and Moravo-Silesia (Kalvoda et al. 2008).

The Tournaisian/Viséan boundary corresponds to an absolute age of c. 347 Ma. The vast majority of Variscan granitoids started to intrude at about this time. The same is true for temperature-dominated metamorphism at all crustal levels. The large areal extent of Carboniferous vol-canic rocks, granitoids and HT metamorphic rocks suggests widespread heating and thermal weakening of the crust. Heating combined with the tectonic causes discussed above suggests that the Variscan crust and lithosphere, in early Carboniferous time, were hot and of moderate thickness (or even thinned).

This weak thermo-mechanical state of the lithosphere from 350 to 340 Ma onwards has seeded the common-place model of “orogenic collapse”. However, there is no


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evidence, in the Variscides, of high topography supported by thickened crust which might have collapsed—neither before nor after 340 Ma. Besides—as laid out above—the Carboniferous to Permian HT processes occur in dis-crete pulses, partly pre-date crustal thickening, are areally widespread and partly are confined to fault zones, some of which even cut across the structural trend of the orogen. These features demand an explanation which is independ-ent from Variscan subduction/collision processes.

Model: Variscan orogeny meets Palaeo-Tethys rift

as already discussed by, e.g. Robardet (2003) and Franke (2006), there is no geological or biogeographic indication of a Palaeo-Tethys ocean separating the armorican terranes from Gondwana mainland before the Permian and the palaeomagnetic arguments for an oce-anic separation are based upon sites in australia (Hur-ley and Van der Voo 1987), which renders long-distance extrapolations uncertain. The first marine ingressions of the Palaeo-Tethys into the Variscan realm are observed in the late Carboniferous and Permian of the Carnic alps (see Krainer 1993, and the stratigraphic evidence in Schönlaub and Histon 2000). This late appearance of Tethys deposits is consistent with the palaeomagnetic considerations of Muttoni et al. (2003, 2009), who pro-posed that the transition of Pangaea “B” (with africa juxtaposed against Variscan Europe) to Pangaea “a” (with the Palaeo-Tethys separating these areas) occurred not before the Permian, and was brought about by EW-trending, large-scale dextral shear zones (see already arthaud and Matte 1977; Matte 1986). In the Pangaea “B” model, the Variscides, in Carboniferous time, were positioned at the tip of the westward prograding Tethys rift. as proposed by Franke (2009b), this position may explain lithospheric thinning, formation of mantle-derived melts advecting heat to the lower crust, high heat flow and a steep geothermal gradient. an alternative, Pangaea “a” model for the Carboniferous has recently been advocated by Domeier et al. (2012). In this sce-nario, the Variscides are not positioned at the tip of the rift, but represent the northern passive margin of Tethys. In such a position, rift-related lithospheric heating and thinning are likewise possible.

as suggested by Simancas et al. (2003, 2006), heat may have been advected by a mantle plume. although some of the Variscan marine basins were formed by thrust (or back-thrust) loading (see above), marine basins originated along-strike-slip faults require pull-apart settings (armorican Massif, Intra-Sudetic basin) or transtension (SW Iberia). It appears that the transtensional regime producing the later Pangaea a—contrary to Henk et al. (2000)—started much earlier than at 305 Ma.

Interaction between the Variscan orogen and the Tethys rift explains the wide areal extent of thermal events. Ther-mal pulses observed in the upper crust may have been controlled either by heterochronous mantle activities, or else by the intermittent activity of fault systems, or both. anyhow, major fault zones (e.g. Montagne noire, Bavarian zone, see above) have acted as conduits for heat-advecting melts. Basically, the model sketched out above extends, backwards in time, processes well known from the Permian situation in Europe and the Mediterranean realm. Henk et al. (2000) stated that primitive asthenospheric melts did not make their way to the surface before the Permian. However, the timing of basaltic magmatism in Scotland (Monaghan and Parrish 2006) reveals that pulsed mantle-derived volcanism goes back into the early Carboniferous (see above).

after an incipient phase of bulging, lithospheric thin-ning by extension will lower lithospheric thickness and surface elevation. These processes permitted the marine ingressions into the orogen listed above. Rising isotherms will have reduced the viscosity of lower to mid-crustal fel-sic rocks and also have occasioned the formation of melts. This also applies to felsic rocks transported downwards by collisional underthrusting. The HT regime will have caused thermal weakening, so that low-viscosity felsic rocks spread laterally already during or immediately after stacking. narrow belts of rocks uplifted during incipient collision have existed, e.g. in the Mid-German Crystal-line High and the margins of the Teplá-Barrandean block. However, thermal weakening will have prevented, in time after c. 340 Ma, further crustal thickening and uplift, i.e. the formation of a substantial orogenic root. Hence, it appears that large parts of the Variscides—in terms of topography—do not represent a collapsed, but a failed orogen.

Interaction of intrinsic (collisional) and extrinsic (rift-ing) processes severely impede tests by numerical model-ling. It appears that the Variscides originally represented a collage of small/cold orogens (in the sense of Jamieson and Beaumont 2013), but the high-temperature regime prevail-ing in time after c. 340 Ma—regarded as typical of large/hot orogens—was probably caused by an independent pro-cess: the Variscides are a child of two fathers.

Consequences for magmatism, magmatic geochemistry and exhumation of (U)HP metamorphic rocks

The thermal regime sketched out above probably represents the main reason for the abundance of granitoids and HT/lP metamorphism, which has traditionally been identified as a distinctive feature of the Variscides, as opposed to alpino-type orogens (see, e.g. the classical publications of Suess 1926; Zwart 1967).


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High heat flow will also allow for a higher degree of melt-ing at relatively shallow levels and therefore make available more mafic components than in a colder regime. Mafic and intermediate lithologies are ubiquitous in the andean-type lower crust and metasomatized lithospheric mantle of the neoproterozoic (Cadomian) orogen, which underlies the entire Variscides (see publications in D′lemos et al. 1990; Dörr et al. 2004). Combined with the high heat flow, these protoliths allow for a large variety of melt compositions, including melt types usually attributed to subduction zones or back-arc basins. This aspect has already been addressed by Henk et al. (2000) and—for the composition of Ordovi-cian magmatic rocks—by Ballèvre et al. (2012). Recycling of Cadomian crust and lithospheric mantle (see also Fernández-Suárez et al. 2011) may present a better explanation for the geochemical signatures in many of the “magmatic arcs” and “back-arc basins” that have been proposed in the Variscides.

It is widely accepted that the ubiquitous U–Pb zircon ages around 340 Ma from granulites in the Saxo-Thur-ingian and Moldanubian zones do not represent peak metamorphic pressures (see the refs. below). O’Brien and Vrána (1995) have shown that granulite facies overprint on a Moldanubian eclogite was extremely short-lived—a feature best explained by a rapid, mantle-controlled ther-mal pulse, rather than by slow processes such as thermal relaxation after stacking. a similar observation from the Śnieżnik massif in the West Sudetes has been published by anczkiewicz et al. (2007): prograde metamorphism at c. 381–387 Ma (lu–Hf and Sm–nd on garnet) was followed by an areally widespread granulite facies overprint at c. 340 Ma (U–Pb zircon, Sm–nd garnet). Taken together, these observations suggest that rapid heating from c. 340 Ma onwards caused a granulite facies overprint on all rocks subject to the relevant temperatures. Zircon growth around 340 Ma in felsic granulites probably occurred dur-ing decompression melting (Finger et al. 1996; Roberts and Finger 1997; Kotková and Harley 2010). This process must have reduced the viscosity of the granulites dramati-cally, so that they were able to intrude into the foreland (e.g. Franke and Stein 2000). It should be noted, however, that granulites in other parts of the Variscides may have a different origin and exhumation history, especially if they were formed in time before or after 340 Ma (lardeaux et al. 2001; Pin and Vielzeuf 1983).

Conclusions

• High altitudes in the Variscides are only possible (although not proven) in narrow belts of crystalline rocks, which represent active plate margins and shed synorogenic clastic sediments into adjacent foreland basins. Devonian and early Carboniferous marine

deposits are widespread not only in the foreland fold and thrust belts, but also in internal parts of the orogen, where they overlie neoproterozoic (Cadomian) or early Variscan basement, thus documenting low topography. For late Carboniferous and Permian times, rich floras and faunas in intra-montane basins likewise contradict high topography.

• Exhumation of HP and UHP rocks in the Bohemian Massif does not record exhumation from the base of thickened continental crust supporting a high plateau, but channel flow from the mantle.

• The absence or scarcity of high relief in Devonian and early Carboniferous time may partly be explained by orogen-parallel extension, continental subduction and subduction erosion. However, the marine deposits sug-gest net lithospheric extension, either in extensional windows (pull-apart basins) in an overall transpres-sional setting or else by transtension. Back-arc exten-sion is considered unlikely because—after the closure of the Rheic ocean in Emsian time—old, cold and dense oceanic lithosphere was no longer available.

• From c. 340 Ma onwards, strong heating of the crust prevented stacking and uplift, because the heated rocks were too weak to build an orogenic root: Variscan oro‑genic topography did not collapse, but failed. HT/lP Metamorphism, magmatism and rheological weakening have started partly before the onset of crustal thicken-ing, occurred in pulses between c. 340 and 270 Ma and were partly confined to fault zones, which cut across the orogen. These features clearly argue against heat sources which were connected, in time and space, with continental collision.

• The Carboniferous HT regime established from c. 340 Ma onwards is best explained by the influence of the Palaeo-Tethys rift, which was eating its way into the Variscan orogen and effected both extension and heat-ing of the lithosphere.

• Onset of strong heating around 340 Ma probably trig-gered the final ascent of felsic HP rocks and their intru-sion into the forelands.

• Heating also effected high degrees of melting and may have extracted mafic components from andean-type Cadomian crust and mantle lithosphere, thus mimicking arc- and back-arc settings.

• High temperatures in time after c. 340 Ma are consid-ered as the main reason for widespread granitoid mag-matism and lP/HT metamorphism, which are the trade mark of the Variscides.

Acknowledgments The author gratefully acknowledges exten-sive and highly informative correspondence with Valerian Bachtadse (Munich), Michel Ballèvre (Rennes), Michel Faure (Orléans), Frie-drich Finger (Salzburg), Ulrich Glasmacher (Heidelberg), Marc Kre-cher (Freiburg), Onno Oncken (Potsdam), Cecilio Quesada (Madrid),


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Stefan Schmid (Zürich), Jörg Schneider (Freiberg), Celâl Şengör (Istanbul), Jennifer Tait (Edinburgh) and andrzej Żelaźniewicz (Wrocław). all these colleagues have contributed good advice and references. The present version of the paper has greatly profited from reviews by M. Faure, an anonymous reviewer and helpful advice of B. Murphy as topical editor.

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Topography of the Variscan orogen in Europe: failed–not collapsed