39Ar age constraints on the Mesozoic low-grade schists of the Circum-Rhodope Belt in the eastern Rhodope-Thrace region, Bulgaria-Greece

Journal of Geodynamics 52 (2011) 143–167

Contents lists available at ScienceDirect

Journal of Geodynamics journal homepage: http://www.elsevier.com/locate/jog

Alpine tectonic evolution of a Jurassic subduction-accretionary complex: Deformation, kinematics and 40 Ar/39 Ar age constraints on the Mesozoic low-grade schists of the Circum-Rhodope Belt in the eastern Rhodope-Thrace region, Bulgaria-Greece Nikolay Bonev a,∗ , Gérard Stampfli b a b

Department of Geology, Paleontology and Fossil Fuels, Sofia University “St. Kliment Ohridski”, BG-1504 Sofia, Bulgaria Institute of Geology and Paleontology, University of Lausanne, Anthropole, CH-1015 Lausanne, Switzerland

a r t i c l e

i n f o

Article history: Received 29 August 2010 Received in revised form 30 December 2010 Accepted 30 December 2010 Available online 5 January 2011 Keywords: Circum-Rhodope Belt Mesozoic low-grade schists Tectonics 40 Ar/39 Ar geochronology Eastern Rhodope-Thrace region Bulgaria Greece

a b s t r a c t Deformation of the Circum-Rhodope Belt Mesozoic (Middle Triassic to earliest Lower Cretaceous) low-grade schists underneath an arc-related ophiolitic magmatic suite and associated sedimentary successions in the eastern Rhodope-Thrace region occurred as a two-episode tectonic process: (i) Late Jurassic deformation of arc to margin units resulting from the eastern Rhodope-Evros arc–Rhodope terrane continental margin collision and accretion to that margin, and (ii) Middle Eocene deformation related to the Tertiary crustal extension and final collision resulting in the closure of the Vardar ocean south of the Rhodope terrane. The first deformational event D1 is expressed by Late Jurassic NW-N vergent fold generations and the main and subsidiary planar-linear structures. Although overprinting, these structural elements depict uniform bulk north-directed thrust kinematics and are geometrically compatible with the increments of progressive deformation that develops in same greenschist-facies metamorphic grade. It followed the Early-Middle Jurassic magmatic evolution of the eastern Rhodope-Evros arc established on the upper plate of the southward subducting Maliac-Meliata oceanic lithosphere that established the Vardar Ocean in a supra-subduction back-arc setting. This first event resulted in the thrust-related tectonic emplacement of the Mesozoic schists in a supra-crustal level onto the Rhodope continental margin. This Late Jurassic-Early Cretaceous tectonic event related to N-vergent Balkan orogeny is well-constrained by geochronological data and traced at a regional-scale within distinct units of the Carpatho-Balkan Belt. Following subduction reversal towards the north whereby the Vardar Ocean was subducted beneath the Rhodope margin by latest Cretaceous times, the low-grade schists aquired a new position in the upper plate, and hence, the Mesozoic schists are lacking the Cretaceous S-directed tectono-metamorphic episode whose effects are widespread in the underlying high-grade basement. The subduction of the remnant Vardar Ocean located behind the colliding arc since the middle Cretaceous was responsible for its ultimate closure, Early Tertiary collision with the Pelagonian block and extension in the region caused the extensional collapse related to the second deformational event D2 . This extensional episode was experienced passively by the Mesozoic schists located in the hanging wall of the extensional detachments in Eocene times. It resulted in NE-SW oriented open folds representing corrugation antiforms of the extensional detachment surfaces, brittle faulting and burial history beneath thick Eocene sediments as indicated by 42.1–39.7 Ma 40 Ar/39 Ar mica plateau ages obtained in the study. The results provide structural constraints for the involvement components of Jurassic paleo-subduction zone in a Late Jurassic arc-continental margin collisional history that contributed to accretion-related crustal growth of the Rhodope terrane. © 2011 Elsevier Ltd. All rights reserved.

1. Introduction The Alpine collisional system in the Aegean region amalgamates various continental and oceanic affinity crustal domains in

∗ Corresponding author. Tel.: +359 2 9308 431; fax: +359 2 944 64 87. E-mail address: [email protected]fia.bg (N. Bonev). 0264-3707/$ – see front matter © 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.jog.2010.12.006

the orogenic belts of the Hellenides and Balkanides (Fig. 1, inset) that have interacted during the Mesozoic geodynamic evolution in the Tethys realm (S¸engör et al., 1984; Robertson and Dixon, 1984; Dercourt et al., 1993; Robertson et al., 1996; Stampfli and Borel, 2002, 2004; Stampfli and Hochard, 2009). This interaction involved Permo-Triassic rifting and opening of oceanic back-arc basins following the Paleotethys closure, ridge and supra-subduction zone ophiolites formation and construction of Jurassic arc systems, ophi-

144

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 1. Synthetic tectonic map of the eastern Rhodope-Thrace region (modified from Bonev and Stampfli, 2008, using data by Papadopoulos and Anastasiadis, 2002) showing the high-grade basement units, the Mesozoic low-grade unit and large-scale structures. Available geochronology for the distinct units is indicated. Kinematic data are generalized from Bonev (2006a,b), Bonev and Beccaletto (2007) and Bonev et al. (2010a). Inset: The Aegean tectonic framework and extension (dashed) of the Circum-Rhodope Belt (CRB). Framed: studied areas of the Mesozoic low-grade unit presented in separate maps as figures.

olite obduction and Alpine orogenic events that have ultimately sutured the crustal-scale continental blocks (Okay and Tüysüz, 1999; Stampfli and Hochard, 2009). Among these continental blocks (terranes), the Rhodope Massif together with the SerboMacedonian Massif, represents major crustal entities bounded in the south by Jurassic ophiolites and related Triassic-Jurassic low-grade sedimentary successions that limit the eastern Vardar suture zone. These low-grade to weakly metamorphosed Mesozoic sequences from the Chalkidiki Peninsula, together with similar units that occur on Samothraki Island in the Aegean Sea and extend into the eastern Rhodope-Thrace region of Bulgaria and Greece, were designated as the Circum-Rhodope Belt (CRB) (Kauffmann et al., 1976; Kockel et al., 1977; Papanikolau, 2009) found around the Serbo-Macedonian and the Rhodope massifs (Fig. 1, inset). Open questions still concern the significance of the CRB as a crustal unit of vast exposure in the northern Aegean region. Among them are the relationships within the Vardar suture zone, the regional-scale extent of this belt, the structural–temporal relationships to adjacent Serbo-Macedonian and Rhodope massifs, with which the CRB virtually interacted during the Mesozoic-Tertiary subduction and collisional history (Michard et al., 1998; Ricou et al., 1998; Brown and Robertson, 2004; Meinhold et al., 2009, 2010). The studies on the CRB in the Chalkidiki Peninsula have

shown a complex Alpine tectonic evolution. The construction of latest Middle-Late Jurassic arc system and related supra-subduction type ophiolites (Jung and Mussallam, 1985; Bébien et al., 1987; Michard et al., 1998) was followed by ophiolite obduction on both the Pelagonian margin and the Serbo-Macedonian margin in Late Jurassic-Early Cretaceous times (Mercier, 1966; Ferrière and Stais, 1995; Michard et al., 1998; Brown and Robertson, 2004) and strong overprint of the Tertiary extensional and strike-slip tectonics (Kilias et al., 1999; Koukouvelas and Aydin, 2002). In the eastern Bulgarian Rhodope Massif, the Jurassic ophiolites and associated Mesozoic low-grade schists of the CRB have experienced Tertiary crustal extension in the hanging wall of SSWdirected detachments (Bonev and Stampfli, 2003, 2008; Bonev, 2006a). They were not affected by S-directed shearing observed in the underlying units, but preserved NW to NNE-directed and earlier internal shear deformation. Only recently it became evident that the low-grade schists were involved in such NNE-directed Late Jurassic thrust emplacement, together with tectonic slices torn off the high-grade metamorphic basement in Bulgarian territory (Bonev et al., 2010a). The corresponding low-grade schists in the Greek part of the Thrace area generally lacked details on the kinematics of Jurassic and Tertiary deformations and radiometric constraints on its temporal spread up to now, which hampers

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

establishing regional-scale deformational events and tectonic evolution. New structural information from the low-grade schists in the eastern Rhodope-Thrace region would provides important insight to this still poorly understood deformational history. Establishing the deformational events in the eastern Rhodope-Thrace region will further facilitate large-scale correlation with the deformational history of the CRB in the Chalkidiki Peninsula and the northerly adjacent Strandzha Massif, which is also known to have been involved in Oxfordian to Barremian top-N orogeny (Okay et al., 2001). Arc-continental margin collision was shown as a prominent mechanism in the crustal growth and subduction-accretion influenced orogenic build-up (e.g. Searle et al., 1987; Alvarez-Marron et al., 2000; Condie, 2007; Cawood et al., 2009). The occurrence of accretionary complexes related to paleo-subduction zones are confined to the metamorphic terrains presently located in the hinterland of the orogens. In this regard, adjacent and within the Rhodope Massif and the Carpatho-Balkan segment of the Alpine belt, the subduction-accretion processes appear as an important mechanism for crustal growth because of the development of arc systems in the Tethyan realm during the Mesozoic evolution (Ricou et al., 1998; Barr et al., 1999; Robertson, 2002; Csontos and Vörös, 2004; Bonev and Stampfli, 2008; Schmid et al., 2008; Turpaud and Reischmann, 2010). Therefore, the identification of structural record and kinematics related to deformational events of the Mesozoic low-grade schists is crucial for the reconstruction of the tectonic history of Alpine orogenic build-up not only for the Rhodope massif. It also provides important clues for the crustal assembly and collisional history at the Eurasian plate during the Mesozoic-Tertiary evolution. The present paper deals with the deformation pattern of greenschist-facies to weakly metamorphosed Mesozoic schists of the CRB in the eastern Rhodope-Thrace region of southern Bulgaria and northern Greece (Fig. 1). The goals of this study are (i) to document structures and kinematics, which will provide insights into their deformation history and tectonic evolution, and (ii) to present geochronologic constraints for the timing of this evolution. We present new data on the structural development and 40 Ar/39 Ar mica ages for the low-grade schists. The obtained results, together with other available data, are discussed in terms of the Jurassic subduction-accretionary history and Tertiary collisional-extensional evolution of the low-grade schists, which is reconstructed in geodynamic context.

2. Geological outline 2.1. General The Rhodope Massif occupies a large region in southern Bulgaria and northern Greece and forms a crystalline terrane that constitutes a major tectonic zone of the Alpine system in the northern Aegean region (Fig. 1, inset). To the north, it is separated by the Maritsa strike-slip fault from the Sredna Gora Zone that, after early Cretaceous orogeny, has developed as a Late Cretaceous continental volcanic arc. To the southwest, together with the crystalline Serbo-Macedonian Massif, it is limited by the CRB and the Vardar (Axios) Suture Zone against the inner zones of the Hellenides. The Rhodope Massif is dominated by a metamorphic basement comprising pre-Alpine and Alpine (e.g. Lips et al., 2000; Liati, 2005) units of continental and oceanic affinities, intruded by Late Cretaceous to Early Miocene granitoids (Meyer, 1968; Del Moro et al., 1988; Dinter et al., 1995; Peytcheva et al., 1999; Marchev et al., 2006). Late Cretaceous-Palaeocene to Miocene-Pliocene sediments (Ivanov and Kopp, 1969; Zagorchev, 1998; Boyanov and Goranov, 2001) and Late Eocene-Oligocene volcanic and volcano-

145

sedimentary successions (Innocenti et al., 1984; Harkovska et al., 1989; Yanev and Bardintzeff, 1997) represent Tertiary syn- to posttectonic cover sequences. Earlier interpretations considered the Rhodope Massif as a Precambrian cratonic block, largely unaffected by Alpine deformations (e.g. Kozhoukharov et al., 1988). In recent synthesis, the Rhodope Massif is regarded as an Alpine syn-metamorphic nappe complex assembled by southward ductile thrusting in the hanging wall of a north-dipping Cretaceous subduction zone located in the Vardar Zone (Burg et al., 1996; Ricou et al., 1998). This Cretaceous nappe stacking was coeval with the amphibolite-facies metamorphism and was followed by syn- and post-orogenic Tertiary extension (Burg et al., 1996; Dinter, 1998; Krohe and Mposkos, 2002; Bonev, 2006a; Bonev et al., 2006b; Bonev and Beccaletto, 2007). The eastern Rhodope Massif low-grade Mesozoic schists considered in this study are part of the CRB and represent the uppermost metamorphic pile (Fig. 1). 2.2. Regional geology of the eastern Rhodope-Thrace region The tectonostratigraphy of the eastern Rhodope-Thrace region comprises four units in a structurally ascending order (Fig. 1, Bonev, 2006b): (i) a lower high-grade basement unit of continental affinity composed of orthogneisses with Variscan protolith ages (Peytcheva and Quadt, 1995; Liati, 2005; Cornelius, 2008; Turpaud and Reischmann, 2010) having an origin mostly from a continental arc S-type granitoids (Bonev et al., 2010b), (ii) an upper lithologically heterogeneous (intercalated various gneisses, schists, amphibolites and marbles) high-grade basement unit of continental-oceanic affinity, enclosing lenses of metaophiolites and metamafic rocks of oceanic ridge and supra-subduction zone tholeiitic-boninitic signature (Kolcheva and Eskenazy, 1988; Mposkos et al., 1989; Haydoutov et al., 2004; Bonev et al., 2006c). The radiometric determinations of the metamafic rocks revealed Neoproterozoic (572 Ma) and latest Paleozoic protolith ages (288 Ma) and Variscan metamorphic ages (350–300 Ma) (Carrigan et al., 2003; Bauer et al., 2007), while gneisses have Late Carboniferous and Late Jurassic igneous ages (Cornelius, 2008), (iii) a low-grade (i.e. greenschistfacies, very low-grade to unmetamorphosed) unit built of Mesozoic schists (e.g. Jaranov, 1960; Von Braun, 1968; Kopp, 1969) and associated Jurassic Evros ophiolite (Magganas et al., 1991), and (iv) a sedimentary-volcanic unit of Tertiary cover sequences (Boyanov and Goranov, 2001) related to extensional and collisional events. Both high-grade basement units are tectonically juxtaposed mostly along flat-lying Tertiary extensional detachments, bounding two large-scale metamorphic domes – the Kesebir-Kardamos and the Byala reka-Kechros domes that dominate the regional tectonic pattern (Bonev, 2006b; Bonev et al., 2006a; Bonev and Beccaletto, 2007). South-directed ductile thrusts involving metaophiolite slices in the Byala reka dome are prominent structures related to syn-metamorphic nappe stacking and crustal thickening in the region (Burg et al., 1996; Bonev, 2006a). The high-grade basement units have a complex Alpine ultrahigh-high to low-pressure and temperature tectono-metamorphic history. This history largely overlaps with Cretaceous nappe stacking in the high-grade basement that is temporarily bracketed by zircon SHRIMP geochronology to have occurred between ca. 149 and 61 Ma (Liati, 2005 and references therein). Recent data suggests that the ultrahigh and high-pressure metamorphism likely occurred before earliest Middle Jurassic (<171–160 Ma) granulitefacies high-temperature overprint (Bauer et al., 2007). 40 Ar/39 Ar hornblende and mica ages define a cooling history between 45 and 36 Ma (Lips et al., 2000; Mukasa et al., 2003; Bonev et al., 2006b, 2009) that accompanies shallow crustal levels Middle-Late Eocene extensional exhumation of the high-grade basement units follow-

146

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

ing amphibolite-facies metamorphism and related SSW-directed Cretaceous ductile nappe stacking. 2.3. Geology of the Mesozoic low-grade unit in the eastern Rhodope-Thrace region The low-grade Mesozoic schists form the uppermost metamorphic tectono-stratigraphic unit that occurs in the hanging wall of the extensional detachments or appear in isolated inliers under the Tertiary cover sequences (Fig. 1). This Mesozoic low-grade unit crops out in scattered areas, resting tectonically on the high-grade basement as remnants of a former nappe of wide regional extent (Goˇcev, 1979; Boyanov et al., 1990; Von Braun, 1993) that has been reworked by Tertiary extensional tectonics (Bonev and Stampfli, 2003; Bonev and Beccaletto, 2007). The ophiolitic rocks that also are part of the low-grade unit are known as the Evros ophiolites and have the characteristics of a supra-subduction ophiolite with arc signature. They contain boninitic and holeiitic to calc-alkaline affinities plutonic and extrusive suites related to an Early-Middle Jurassic marginal basin-arc system evolution (Magganas et al., 1991; Magganas, 2002; Bonev and Stampfli, 2008, 2009). Because the known occurrences of the Mesozoic low-grade schists refer to geographic names and/or areas usually equated with the distinct units within its section, we keep in the following description the subdivisions widely used in the literature (e.g. Papadopoulos et al., 1989; Ricou et al., 1998; Dabovski et al., 2002; Pe-Piper and Piper, 2002) (see Fig. 1 for the main units detailed below and depicted in Figs. 3, 5 and 8). In the Thrace area of northeastern Greece, the Mesozoic lowgrade unit is subdivided into the Makri unit and conformably overlying Drimos-Melia unit (Papadopoulos, 1980, 1982) (see Figs. 1, 3 and 5). The contact between both units is considered also as a thrust fault (Von Braun, 1993). The stratigraphic section of the Makri unit includes a lower metasedimentary series conformably overlain by an upper metavolcanic-sedimentary series, which are intercalated with and covered by limestone horizons. The metasedimentary series (i.e. carbonate and clastic sequences) has shallow-marine platform type slope-rise characteristics (Papadopoulos et al., 1989). The metavolcanic-sedimentary series is dominated by greenschists (chlorite-actinolite-tremolite-epidote-sericite schists) derived from mafic to acid lavas and pyroclastics, with occurrences of scarce serpentinite bodies. According to Kopp (1969) the upper limestone horizon, known as Lower Cretaceous “Aliki limestones” (Maratos and Andronopoulos, 1964) lies unconformably on the upper greenschist series, implying a pre-Cretaceous age of deposition and greenschist-facies metamorphism of the Makri unit. One of the lowest limestone horizons yielded Upper Triassic corals (Maratos and Andronopoulos, 1964) and the metasedimentary series supplied Tithonian-Beriassian ammonites (Dimadis and Nikolov, 1997). The detrital zircons in the sandstones cluster at ca. 310–290 Ma and 240 Ma, providing at least a Middle Triassic deposition age for the stratigraphically lower levels of the unit (Meinhold et al., 2010). A gabbroic body belonging to the intrusive suite of the Evros ophiolite, namely the Petrota complex, occurs in the Tertiary graben superimposed on the Makri unit (Fig. 1). A zircon SHRIMP-II crystallization age of 169 ± 2 Ma was reported for the Petrota gabbro (Koglin et al., 2007), whereas apatite fissiontrack ages between 160 and 141 Ma for the gabbro (Biggazzi et al., 1989) reflect its cooling or uplift/emplacement history at shallow crustal levels. Another Evros ophiolite intrusive suite gabbroic body occurs at Agriani (Bonev and Stampfli, 2009). The Drimos-Melia unit consists of massive mafic lava flows, pillow lavas and sheeted-like dykes intercalated in a flysch succession (Papadopoulos, 1982; Papadopoulos et al., 1989; Bonev and Stampfli, 2005), which yielded Middle-Upper Trias-

sic (bivalve Halobia; Dimadis et al., 1996) and Middle-Upper Jurassic (Callovian-Oxfordian ammonite; Trikkalinos, 1955) biostratigraphic ages. The detrital zircons in the flysch cluster at ca. 315–285 Ma and the youngest grain at ca. 160 Ma provides a latest Middle Jurassic maximum age of deposition (Meinhold et al., 2010). Middle Eocene (Lutetian) conglomerates, Late Eocene-Oligocene limestones, sandstones and marls (Kopp, 1965) are unconformable sedimentary cover unit onto the Makri and Drimos-Melia units. The cover unit includes also thick Late Eocene-Oligocene up to early Miocene volcanic and volcanic-sedimentary successions (Christofides et al., 2004). The Oligocene Maronia granodiorite intrudes the Makri unit (29 Ma, Del Moro et al., 1988). In the eastern Bulgarian Rhodope, the low-grade Mesozoic unit is exposed to the north within the Kulidzhik nappe (Boyanov, 1969), and in the Mandritsa area (Boyanov et al., 1990) to the south (see Figs. 1 and 8). The low-grade sequence of the Mandritsa area can be simply subdivided into two main units of decreasing metamorphic grade structurally upward section following a revised subdivision as given by Bonev and Stampfli (2008): (i) a greenschist unit; and (ii) an overlaying mélange-like volcano-sedimentary unit (see also Fig. 8). The greenschist unit overlies the lower unit of high-grade basement along the Tertiary Byala reka extensional detachment (Bonev and Stampfli, 2003; Bonev, 2006b). The mafic rock lenses and greywacke knockers-bearing marble horizon (Bonev, 2005a) occurs at the base and is presently assigned to this unit. This horizon has been previously attributed to the underlying upper high-grade basement unit (Boyanov et al., 1990; Bonev and Stampfli, 2003, 2008). The structural criteria (see section below) and recently obtained hornblende inverse isochron 40 Ar/39 Ar age of 156.6 ± 0.6 Ma from an amphibolite intercalation in analogous marbles northwesterly confirm this stratigraphic assignment, implying an imprint of a latest Middle-Late Jurassic upper greenschistfacies metamorphism (Bonev et al., 2010a). Up-section greenschist unit is dominated by intercalated quartz-chlorite schists, actinolite schists, actinolite-chlorite ± epidote schists, whose protoliths were mafic lavas and pyroclastics. Phyllites and pellitic micachlorite ± garnet schists also occur. The transition to the overlaying mélange-like volcanosedimentary unit is gradual, and marked by up-section increase of metavolcanic rock types. Thus, the mélange-like volcanosedimentary unit overlie depositionally the greenschist unit, and locally through thrust contacts. Sheet-like basalt to andesite ophiolitic lava flows occur essentially in this unit, together with boninitic dykes intruding basal marble horizon of the greenschist unit below the mafic lavas. The mafic lavas are intercalated with Early Jurassic (latest Pliensbachian to pre-Bajocian) radiolarian chert layers (Tikhomirova et al., 1988). The overlaying turbiditelike sediments consist of conglomerates, greywackes, siltstones and black shales, which succession contains in an olistostromic context reworked Upper Permian and Middle-Upper Triassic shallowwater carbonate clasts and blocks embedded in shale-silty and volcanogenic matrix (Boyanov and Trifonova, 1978; Boyanov and Bodurov, 1979; Trifonova and Boyanov, 1986; Bonev, 2005b). The mélange-like unit is very low-grade to unmetamorphosed. The NNE-directed Kulidzhik nappe imbricates Jurassic mafic extrusive rocks fully comparable to those in the Mandritsa area and the orthogneisses of the lower unit of the high-grade basement in the nappe allochthon, with a Late Jurassic (40 Ar/39 Ar 156.6–154.2 Ma white mica ages) cooling history (Bonev, 2006a; Bonev et al., 2010a). These data indicate a Late Jurassic thrustrelated tectonic emplacement and Middle Jurassic age (bracketed between the crystallization of the mafic lavas and cooling of the Kulidzhik allochthon) of the greenschist-facies metamorphism. Middle-Upper Eocene clastics and Oligocene sedimentary and vol-

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

canic rocks are unconformable cover successions in the Mandritsa area and the Kulidzhik nappe (Dimitrova et al., 2000; Boyanov and Goranov, 2001). Structurally allochthonous, the Mesozoic low-grade unit is regarded as forming thrust sheets of a complex nappe with inferred north-vergent movement involving also the high-grade basement (Boyanov et al., 1990; Von Braun, 1993). Preliminary structural data from the Mandritsa area have documented top-to-the SSW tectonic transport on the basal detachment related to extensional unroofing of the high-grade basement (Bonev and Stampfli, 2003; Bonev, 2006b), whereas in both the Mandritsa area and the Kulidjik nappe strong internal deformation with top-to-the-NW to NE-directed shear fabric in the greenschists was related to thrust emplacement accompanying accretion of the Mesozoic low-grade unit to the Rhodope margin (Bonev and Stampfli, 2003; Bonev, 2006a; Bonev et al., 2010a). Metamorphism in the low-grade unit varies spatially, structurally and statigraphically upwards in the section of the aforementioned units or areas. Higher grades of greenschist-facies metamorphism reaching transitional epidote-amphibolite facies conditions are established northwards, with almandine garnet and biotite in pelitic rocks and actinolite- to magnesium hornblende and epidote in metabasic rocks in the Kulidzhik nappe (Bonev et al., 2010a). The same assemblage is reported from the Mandritsa area (Boyanov et al., 1990; Bonev and Stampfli, 2008). In the Makri unit, chlorite-actinolite-white mica-sericite assemblage is common for metasedimentary rocks, whereas the mafic lavas of the Drimos-Melia unit are characterized by prehnite-pumpellyite facies assemblage of ocean-floor metamorphism and related alteration (Papadopoulos et al., 1989; Magganas, 2005). Many limestone horizons, especially in the Makri unit, are slightly recrystallized preserving fossil content and are generally weakly metamorphosed to unmetamorphosed. 3. Structural record The structural study was conducted within the largest outcrop areas of the Mesozoic low-grade unit in the eastern RhodopeThrace region, which systematically underlie the Evros ophiolite extrusive suite in the Makri-Maronia-Dokos and Mandritsa-Micro Derion-Metaxades-Dydymotichon areas, and the outcrop area near the Fanari village (Fig. 1). Systematic field and thin section observations on the geometry and overprinting relationships of the structures and fabrics at the map scale and along selected transects were combined with kinematic analysis to assess tectonic transport direction and deformational history experienced by the rocks during their Alpine tectonic evolution. The following description refers to standard structural terminology of sequentially developed planar, linear and fold structures conventionally related to deformational phases coupled with the kinematics of displacement during the tectonic events or episodes (i.e. S1 , S2 , L1 , L2 , F1 , F2 , D1 , D2 etc.). 3.1. Fanari area East of the village Fanari, an isolated coastal outcrop area exposes a flysch succession (Fig. 1). The flysch presents a typical rhythmic shale-sandstone alternation, with the occurrence of decametre to hectometre-thick sandstone-gravel horizons (Fig. 2a). Graded bedding and channelized flow provide evidence for deposition in an outer to middle fan environment from high density turbidite currents. The Fanari flysch strongly resembles in terms of lithology the flysch succession of the Drimos-Melia unit further east. The flysch is strongly ductilely deformed, with a structural pattern dominated by folds. An earlier fold generation includes rare

147

small-scale isoclinal F1 folds of chlorite and white mica-defined foliation S1 that parallels the sedimentary layering S0 . The F1 folds are coaxially refolded into the hinges of F2 folds as is apparent from the nearly uniform thickness in the profile of folded layers (Fig. 2b). F2 folds predominate and show a range of style and geometry. The F2 folds are moderately inclined, with axes plunging to NE and have developed an axial-planar slaty cleavage S2 that fans across the fold profile (Fig. 2a, stereoplot). The F2 folds are close to tight parallel folds in stiff sandstone strata, whereas in ductile shale strata a refraction of penetrative axial-planar cleavage S2 occurs (Fig. 2c). The F2 folds acquire the geometry of chevron folds in thin sandy layers of sandstone-gravel horizons of the flysch, where S0 //S1 //S2 in the fold limbs (Fig. 2d). The F2 folds have pronounced NW vergence and sheared hinges along the metre-scale hinge-propagating thrusts, which also propagates in the forelimbs. The hinge and within limbs thrusts have accommodated north-northwest directed displacement related to progressive overturning of F2 folds, together with additional movement along the S2 cleavage planes (Fig. 2d and e). Rarely observed mineral lineation along the F2 folds profile parallels the hinge thrust propagation direction. Field and in thin section observations indicate only weak (sub-) greenschist-facies metamorphism of the flysch as evidenced by quartz-chlorite-white mica/sericite assemblage, when shales become phyllitic in intensely ductilely sheared domains. Overall, the asymmetry of folds and related shear structures in the Fanari area indicate top-to-the-N-NW directed tectonic transport (Fig. 2a). 3.2. Makri-Maronia area The Makri-Maronia area exposes the Makri unit and forms the largest outcrop of the Mesozoic schists, located at the Aegean Sea coast (Fig. 1). Previous structural investigations have identified earlier B1 folds with NE-SW-oriented axes and later B2 folds with WNW-ESE and NNE-SSW-oriented axes (Kopp, 1969; Meyer, 1969). Subsequently, a group (Alx 1) of NE-SW to E-W-trending tight to sub-isoclinal folds that are syn-metamorphic with respect to the grenschist-facies metamorphism, a group (Alx 2) of open parallel folds with axial directions WNW-ESE, and a group (Alx 3) with NESW to NW-SE-oriented large open to very open folds were reported (Patras et al., 1989). In the Makri area, the internal structure of the Makri unit exhibits a dominant large-scale F2 fold pattern (Fig. 3). The metamorphic section displays a regional foliation S1 that is defined by the planar alignment of the phylosilicates. This main foliation represents ubiquitous schistosity or compositional layering that parallels primary bedding S0 , particularly when centimetrescale quartz and metre thick-limestone layers are intercalated with quartz-chlorite-muscovite-sericite schists and phyllites. The foliation S1 varies in orientation due to intense F2 folding (Fig. 3), with locally complex refolding patterns. Scarce slightly inclined isoclinal F1 minor folds with NW-SE oriented axes are refolded around axes oblique to parallel to the hinges of minor F2 folds (Fig. 4a). The foliation S1 is transposed in the hinges of asymmetric tight to isoclinal NW-vergent F2 minor folds with NE-SW trending sub-horizontal axes and steeply SE-dipping axial surfaces (Fig. 3, stereoplots). There, a discontinuous spaced axial planar cleavage S2 is marked by the alignment of phylosilicates and crystallizations of calcite and quartz aggregates forming gashes along S2 cleavage planes. The cleavage S2 typically parallels S0 //S1 in the limbs of F2 folds and fans around upright to inclined metre-decametre scale and inclined open F2 kilometre-scale antiforms and synforms, displaying a characteristic parasitic geometry (Figs. 3b and 4b). Due to shear deformation and progressive fold overturning, the F2 folds acquire the geometry of strongly inclined to recumbent folds in which shearing has been accommodated by slip along the

148

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 2. Fanari area. (a) Geological map of the Fanari area. Stereoplot: lower hemisphere projection; (b) hand sample showing the F1 –F2 fold generations and associated bedding S0 and cleavage S1 in the flysch; (c) outcrop-scale expression of F2 folds and associated S2 cleavage. Note displacement along the S2 cleavage planes in the direction of fold overturning; (d) metre-sale F2 fold cut by forelimb propagation thrust. Note S2 cleavage refraction in pelitic strata; (e) closer view of boxed area in (d) showing asymmetric boudinage of sandstone layer involved in thrusting consistent with the displacement on hinge to forelimb propagating thrust.

S2 cleavage planes. Internal NW-vergent decametre-scale thrusts are associated with F2 fold development, especially in the forelimbs of F2 folds, where they are demonstrated by small-scale shear planes (Fig. 4a and d). F2 hinge propagating thrusts also exist at the same scale. The geometric and structural features of F2 folds indicate their origin by fault-propagation folding. The mineral elongation lineation L1 , which in places becomes a stretching lineation, is delineated by streaking chlorite-micas aggregates, and trends NNW-SSE to NE-SW with shallow to moderate plunges (Fig. 3, stereoplots). The NNW-trending more steeply plunging lineation L1 apparently parallels hinges of minor F1 folds, whereas the NESW-oriented shallow plunging lineation L1 is seen folded around hinges of F2 folds. This feature is particularly evident in competent carbonaceous strata, where lineation L1 lays oblique to parallel across the hinges of F2 minor folds (Fig. 4b and c). The lineation

L2 is typically a S1 /S2 intersection lineation that parallels the F2 folds hinges (Fig. 3, stereoplots). A shallow plunge of the NE-SWoriented L1 lineation similar to that of the hinge lines of F2 minor folds is geometrically consistent with fold amplification-related progressive reorientation. The kinematic indicators associated with the lineation L1 depict top-to-the-NW ductile shear, which is topto-the-NE-directed when associated with L2 and F2 development (Fig. 4e and f). In the Maronia area, the greenschists display mostly flat-lying metamorphic sections of the regional foliation S1 or compositional layering of intercalated quartz-albite-chlorite-epidote and chlorite-muscovite-sericite schists, and phyllites, marbles and recrystallized limestones of both series of the Makri unit (Fig. 5a and b). Strongly oriented mica flakes and chlorite aggregates within foliation S1 define the mineral stretching lineation L1 , which is ori-

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

149

Fig. 3. Makri area. (a) Structural map of the Makri area (geological base simplified after Papadopoulos, 1982). Stereoplots, lower hemisphere projection. (b) Cross-section on line indicated in (a).

ented generally N-S with shallow plunges. NNW-SSE trending F1 minor folds are very rare. Intense folding has transposed locally the foliation S1 in the hinges of asymmetric tight to isoclinal ENE-WSW to E-W trending and NNW-vergent F2 folds with a south-dipping axial planar cleavage S2 , which delineate the S1 /S2 intersection lineation L2 that parallels F2 fold hinges (Figs. 5 and 6 Figs. 5c, 6a and 6b). The F2 folds have developed at high angle to the lineation L1 , with axes having an orientation nearly orthogonal to the low-angle N-dipping shear bands depicting strong shear fabric (Figs. 5c, 6c and 6d). These structural elements associated with kinematic indicators demonstrate top-to-the-NNE-directed ductile shear (Figs. 5a, 6c and 6d). Although down-dip shear fabrics in the Maronia area may originate either through contraction or extension, when taken collectively with the shear fabrics and associated thrust and fold structures in the Fanari and Makri areas, the shear fabrics despite its flat-lying attitude stand for an origin in contractional setting that shortened the metamorphic datum. The shearing developed under greenschist-facies conditions, as indi-

cated by common crystallizations of chlorite and white mica in the shear bands and chlorite infilling of albite strain shadows is observed at the outcrop scale and in thin section. The geometric relationships in the structural pattern of the Makri unit in the Maronia area suggest a deformation phase, in which localized shear deformation yielded intrafolial transposition of the planar-linear fabric during progressive deformation. The shear deformation and F2 folding are also concentrated within poorly exposed tectonic contacts with the overlying mafic lavas of the Drimos-Melia unit, where the acquisition of penetrative foliation S1 defines the likely thrust contact below these lavas, turned into greenschists at the base. 3.3. Dokos area The area near Dokos exposes the low-grade unit of the Mesozoic schists in fault contact with the extensional detachment in the southwestern flank of the Byala reka-Kechros dome (Krohe

150

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 4. Field photographs of the Makri area. (a) Refolded F1 minor fold around the hinges of asymmetric NW-vergent minor F2 folds with developed spaced axial planar S2 cleavage. Small-scale thrusts propagate into F2 fold hinges; (b) metre-scale F2 fold delineated by competent recrystallized limestone layer, with minor parasitic folds; (c) close-up view of boxed area in (b) showing relationships of the lineation L1 with F2 fold axes. Note slightly curvilinear (noncylindrical) character of F2 folds; (d) decametre F2 fold with thrust propagated into the hinge; (e and f) asymmetric boudinage of quartz layers associated with NW vergent F2 folds indicating NNW–NE-directed ductile shear, respectively tectonic transport direction. The expression of mineral lineation L1 is seen in the rock fragment in the lower right of the photograph.

and Mposkos, 2002; Bonev and Beccaletto, 2007), where it occurs in the hanging wall and below the Tertiary cover sequences (Figs. 1, 7a and 7b). No details on the lithology and structures were available for the area prior to this study. The metamorphic section is lithologically similar to that observed in the Makri unit, and therefore considered the northern extension of the latter unit. In Dokos area, the metasedimentary series mostly consisting of shales corresponding to the lower stratigraphic levels of the Makri unit is conformably overlain by the greenschists of the metavolcanic-sedimentary series with intercalated thin limestone horizon. The small remnant of massive mafic lava flow can be considered as a counterpart of the extrusive suite of the Drimos-Melia unit. The internal structure of the low-grade schists is dominated by map-scale F2 folds within the framework of the extensional hanging wall allochthon (Fig. 7b). The earliest F1 fold generation was not identified in the area. The regional foliation S1 is folded into various scale close to tight NW-vergent F2 folds with NE-SW trending and variably plunging hinge lines (Fig. 7a,

stereoplots). The limbs of F2 folds are sheared by metre-scale shear planes that displace the layering in the direction of fold asymmetry and are associated with metre-scale NNW-directed thrusts developed in the forelimbs of decametre-scale antiformsynform pair (Fig. 7b). The F2 fold development and related shear deformation are further evidenced in thin section, where smallscale thrust duplexes display displacement in the direction of fold overturning and are associated with occasionally developed axial planar cleavage S2 , locally sub-parallel to the foliation S1 (Fig. 7c). The hinge lines of F2 folds trend at high angles to WNW-ESE to N-S oriented mineral stretching lineation L1 defined by strongly elongated actinolite needles and chlorite aggregates (Fig. 7a, stereoplots). This lineation pattern clearly results from F2 folding, although the hanging wall W-E lineation trend spatially close to the detachment may result from the detachment-related displacement late rotation/reorientation. The lineation L1 is associated with kinematic indicators showing top-to-the-NNW tectonic transport that has been identified away from the intensely folded

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

151

Fig. 5. Maronia transect of the Makri unit. (a) Structural map of the Maronia area (geological base simplified after Papadopoulos, 1982). (b) Cross-section on line indicated in (a). (c) Stereoplots, lower hemisphere projection.

domains. Deformation–crystallization relationships indicate white mica, chlorite and garnet crystallizations associated with the deformational phase creating planar and linear (S1 –L1 ) fabrics at the onset of NNW-directed ductile shearing, which growth continued during the deformation related to F2 folding in greenschist-facies. This is evidenced by prolonged white mica crystallization within the matrix, along the shear bands and around pre- to late tectonic garnet porphyroblasts containing matrix mineral inclusions (Fig. 7d). The subsequent deformation relates to the extensional shearing and detachment faulting juxtaposing the hanging wall low-grade unit against the lower high-grade basement unit of amphibolite and higher metamorphic grade (Bonev and Beccaletto, 2007). The detachment is marked by hydrothermally altered rocks in a shallow south-dipping meters-thick cataclastic zone. Different metamorphic grades and orientations of structural patterns, especially highly oblique lineation patterns in the high- and low-grade basement units, unequivocally distinguishes the late extensional deformation, which the low-grade unit has suffered passively within the hanging wall. Overall, the Dokos area shows structural paragenesis of an early deformation associated with development of main structural fabrics (S1 –L1 ), dominant F2 folding and weak (S2 –L2 ) fabrics, both in greenschist-facies reaching the garnet zone. The late deformation is related to extensional detachment

faulting inducing hanging wall reorientation of earlier structural fabrics.

3.4. Mandritsa area The Mandritsa area, located at the Greek-Bulgarian border, represents the western continuation of the low-grade rocks exposed in the area of the Metaxades and Micro Derion villages and farther east at the town of Dydimothycho near the Greek-Turkish border (Fig. 1). Structurally, the low-grade rocks form the hanging wall metamorphic succession of the extensional detachment bounding the eastern flank of the Byala reka-Kechros dome (Figs. 1, 8 and 10a). The footwall mylonites of the underlying basal extensional detachment display top-to-the-SSW ductile–brittle tectonic transport (Bonev, 2006b; for details on extensional fabrics) and NNE-SSW oriented brittle extension of high-angle fault mode in the hanging wall, with Middle-Late Eocene 40 Ar/39 Ar mica ages documenting the cooling history of the footwall (38–36 Ma) and hanging wall (40 Ma) (Fig. 8; Bonev et al., 2009, 2010a; Márton et al., 2010). Here, we provide details on the kinematics and structural pattern of the Mandritsa area following our earlier work (Bonev and Stampfli, 2003).

152

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 6. Field photographs of the Maronia area transect. (a) Aspect of the regional foliation S1 , the geometry of F2 folds and associated axial-planar cleavage S2 . (b) Closer view of the relationships between F2 folds and intersection lineation L2 from the boxed area in (a). (c) NNE-dipping shear bands (half arrows) associated with ␴-type clast derived from the ductile deformation of a quartz layer. (d) Same as in (c), but the clast is ␦-type.

The regional foliation S1 in the greenschists has the typical expression of a ubiquitous schistosity or compositional layering defined both macro- and microscopically by the planar alignment of chlorite and actinolite, as well as marked by fine alternation of quartz-albite ± white mica and actinolite ± chlorite-rich layers. The foliation S1 , generally sub-parallel to the bedding S0 , varies in attitude due to map scale folding generally around NE-SW trending axis (Fig. 8, stereoplots). The mineral stretching lineation L1 trends dominantly NNW-SSE to NW-SE with shallow to moderate plunges. The lineation L1 is reoriented towards the NE-SW direction parallel to the lineation trend in the footwall lower high-grade basement unit at the vicinity of the NE-SW trending corrugation fold of the extensional detachment (Bonev and Stampfli, 2003). The small-scale F1 folds in the greenschists vary from close to isoclinal, with axes oriented NW-SE to NE-SW in a way similar to the lineation trend (Fig. 8). They fold the bedding/schistosity (S0 –S1 ) often showing parasitic geometry consistent with fold overturning (Fig. 9b). Rare refolding patterns of this early and scarce fold generation F1 is observed parallel to the hinges of F2 folds (Fig. 9a). Occasionally, a spaced cleavage S2 is developed axial-planar to tight or isoclinal F2 folds with NE-SW axial directions, implying that these folds developed subsequently to the F1 folds. The weak S1 /S2 intersection lineation L2 then parallels F2 fold hinges. The geometry, style and orientation of NE-SW trending folds relate them to the F2 fold generation described in the previous areas in this study. At least in the vicinity of the detachment and its corrugation fold, the F2 are influenced, i.e. rotated towards parallelism with the late extensional NE-SW oriented shear direction of the displacement along the detachment (Fig. 8). All described folds have a consistent northward asymmetry. Associated with the lineation L1 , sense-of-shear criteria indicate a dominant top-to-the-

NNW–N shearing in greenschist-facies metamorphic conditions (Figs. 8 and 9). The marbles that bear greywacke knockers and mafic lenses at the base of the low-grade unit unequivocally demonstrate top-to-the-NNW-directed shear fabrics depicted by the same fold style and asymmetry of shear structures that are analogous structurally upwards in the section of the Mandritsa area (Fig. 10b–e). The relationships of metamorphic crystallizations and deformational shear fabrics in thin sections revealed syn-kinematic growth of the actinolite laths incorporating matrix quartz grains that invariably are associated with the crystallization of the chloritewhite mica and quartz aggregates omnipresent in the greenschists (Fig. 10b–d). The garnet porphyroblasts also grew syntectonically, preserving incorporated matrix mineral inclusions, showing progressive growth in the F2 fold hinges and portraying an offset along the unequivocally NW-directed shear bands (Fig. 9e and f). Mineral assemblages consisting of Qtz + Act + Chl + Ms ± Ep ± Grt ± Bt, both in the metabasic and metasedimentary protoliths of the low-grade schists, indicate greenschist-facies metamorphism with temperatures generally below 500 ◦ C (e.g. Spear, 1993), with maximum grade of the transitional epidote-amphibolite facies, where the presence of epidote is largely due to Ca-rich mafic protolith of the greenschists. 3.5. Metaxades-Micro Derion-Didymothycho area The Metaxades area represents an immediate extension of the Mesozoic schists of the Mandritsa area. The greenschists at the Micro Derion village and the Didymotycho town are isolated from the Mandritsa-Metaxades areas, and occur separately (Fig. 1). In the Micro Derion-Metaxades area, the structure of the low-grade schists consists of an eroded ductile thrust sheet.

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

153

Fig. 7. The Dokos area. (a) Structural map of the low-grade sequence near the village Dokos. Stereoplots, lower hemisphere projection. (b) Cross-section on line indicated in (a) and decametre-scale F2 fold. Black half arrow indicate shear plane in the fold limb, whereas white half arrow depicts minor thrusts. (c) Microphotograph of minor F2 fold with developed axial planar cleavage S2 and micro duplexes (half arrows) along the sheared limb. (d) Ductile shear fabric in garnet-bearing white mica-chlorite schist depicting NW-directed shear bands (half arrows), rotated garnets (Grt) and mica “fish” (Ms).

Field and petrographic observations indicate the presence of an analogous to the Mandritsa area metavolcanic succession, including intercalated chlorite schists, actinolite schists, actinolitechlorite ± epidote ± white mica schists. The mica-chlorite defined regional foliation S1 displays a folded pattern on map-scale, an antiform-synform pair and in the hanging wall of the basal ductile thrust (Fig. 11a and b). The thrust surface is mostly wellexposed in the northern area, where this tectonic contact is

overprinted by high-angle faults related to the late extensional deformation (Fig. 11d and e), as documented further west in the extensional detachment in the Dokos area and the upper highgrade basement–low-grade unit contact nearby the Mandritsa area. Minor folds are not well developed in the Metaxades-Micro Derion area due to the predominance of the mafic protolith lithologies, although rare north-vergent close to tight folds can be observed in places. The mineral, and locally the stretching lineation L1

154

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 8. Structural map of the Mandritsa area (modified from Bonev and Stampfli, 2003, 2008).

trend NW-SE, and associated shear sense indicators unequivocally demonstrate top-to-the-NNW ductile shear sense in greenschistfacies metamorphic conditions (Fig. 11a and c). At Didymoticho, the low-grade schists underlie as a thin horizon of intercalated Makri unit specific lithologies (i.e. phyllites, shales, limestones) the mafic lavas of the Evros ophiolite, and occur in basal ductile thrust contact with the upper unit of the high-grade basement (Fig. 12a). The internal deformation of the low-grade schists and within the zone of thrust contact is expressed by NE-vergent tight to isoclinal occasionally intrafolial minor F1 folds in the W to S and NW-dipping foliation S1 containing mostly SW and rarely NW plunging mineral lineation L1 . These structural elements are associated with kinematic indicators showing top-to-the-NNE and NW-directed tectonic transport in greenschists-facies conditions as demonstrated by metamorphic crystallizations of chlorite-white

mica-calcite associated with the ductile shear fabrics (Fig. 12b). In contrast, the underlying upper unit of the high-grade basement displays a structural pattern of mostly N-dipping regional foliation and W-ENE trending mineral lineation that associates with scarce E-SE plunging and SSW-vergent minor folds. The attitudes of structural elements in the low- and high-grade basement units suggest that the ductile thrust surface is openly folded in a way similar to the Metaxades-Mikro Derion area. Kinematic indicators in the high-grade basement demonstrate top-to-the-SE and SW-WSW-directed ductile shearing under amphibolite-facies conditions (Fig. 12c). Alteration of the alkali feldspar to epidote and white mica in quartz-feldspatic high-grade basement gneisses testify of the lower grade amphibolite-facies or retrogression of this event, which is significantly higher than the metamorphic grade of the overlying low-grade schists. The structural elements in both

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

155

Fig. 9. Field and microphotographs of the greenschists in the Mandritsa area. (a) Refolded pattern of F1 folds in F2 fold hinges; (b and c) quartz-chlorite-actinolite schist with asymmetric actinolite laths including matrix quartz grains ± chlorite, depicting N-directed ductile shear. Note strain caps and shadows, suggesting that shearing outlasted actinolite growth; (d) quartz aggregate and associated shear bands (half arrow); (e and f) deformation–crystallization relationships of garnet porphyroblasts with respect to F2 folding and ductile shearing. Internal pattern of the inclusions of chlorite and quartz indicates synkinematic garnet growth concomitant with the deformation. Abbreviations: Ms, muscovite; Grt, garnet; Chl, chlorite; Act, actinolite; Qtz, quartz; Cal, calcite.

units are overprinted by brittle shears and small-scale faults that obviously relates to the late deformation manifested by high-angle faulting during Tertiary crustal extension.

laser-probe dating on micas from the Mesozoic schists was performed in this study.

4.1. Analytical procedures and samples 4.

40 Ar/39 Ar 40 Ar/39 Ar

geochronology

mica ages (154–157 Ma) from the high-grade basement orthogneisses in the Kulidzhik nappe allochthon overlying the greenschists and a hornblende age (apparent age of 189 ± 13 Ma resulting from excess Ar, with inverse isochron of 156.6 ± 0.6 Ma) in the marbles of the westernmost exposure of the Mandritsa area (see Fig. 8) constrain the Late Jurassic age of the greenschist up to epidote-amphibolite facies metamorphism coeval with the northdirected thrusting of the Mesozoic schists in the eastern Bulgrian Rhodope (Bonev et al., 2010a). In order to provide further temporal constraints and to refine the low-temperature history, 40 Ar/39 Ar

The 40 Ar/39 Ar determinations were carried out on mica separates in three micaschist samples. Micas were separated from the 250–500 ␮m sieve fraction using standard magnetic and density separation methods. A final hand-picking for the purity of the mineral concentrates was carried out under a binocular microscope. The concentrates were thoroughly cleaned in baths of acetone, distilled water and ethanol before being packaged in an aluminium foil envelope. A small fraction of each mineral separate was wrapped in aluminium and placed in aluminium discs in a package together with 28.00 Ma sanidine from the Taylor Creek rhyolite (Duffield and Dalrymple, 1990) used as the neutron flux monitor, with ±0.5%

156

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 10. Mandritsa area. (a) Cross-section on line indicated in Fig. 8; (b) style of F1 folding in the marbles at the base of the low-grade schists; (c) shear deformation of the same marbles, depicted by asymmetric boudinage and N-dipping shear bands (half arrow); (d) outcrop-scale NNW-directed thrusts in the high-grade basement in immediate contact below the basal marbles; (d) asymmetric boudinage of quartz layer in structurally higher greenschist levels.

1. The samples were then sent to the McMaster nuclear reactor (Hamilton, Ontario, Canada) for fast-neutron irradiation in the core position 5C. The samples were irradiated for 30 h at 2 MW power level in three-step increments. The samples were analyzed by incremental heating experiments in which the isotopic composition of the argon gas in each heating step was measured by MAP 215-50 mass spectrometer in the Geochronology Laboratory at the Massachusetts Institute of Technology (Boston, USA). Additional details on analytical procedures, corrections and uncertainties can be found in Hodges et al. (1994). The sample locations of the analyzed mica separates are shown in Figs. 5 and 8. Sample GR-114 represents a mylonitic micaschist from the Maronia area of the Makri unit. It consists essentially of quartz, chlorite and white mica. The mica flakes and chlorite define the penetrative foliation S1 and the lineation L1 , folded by F2 folds (Fig. 6). Sample M-9 is quartz-chlorite-white mica schist taken at the base of the Mandritsa unit from the greenschists immediately above basal marbles that in turn overlies the mylonites of ductilebrittle shear zone tracing the extensional detachment (Fig. 8). The mineral assemblage in this sample also includes calcite, disseminated magnetite, rare zircon and apatite. Analyzed white mica is represented by flakes that define the foliation S1 and lineation L1 associated with top-to-the-NW shear fabric in the greenschists (see Fig. 2 in Bonev and Stampfli, 2003). A single white mica generation in the metamorphic mineral assemblage was identified in

thin sections. Quartz grains show low-temperature intracrystalline deformation expressed by bulging recrystallization and development of incipient subgrains. Sample M-10 is a quartz-micaschist located structurally and stratigraphically up-section from sample M-9, both from the Mandritsa area. The mineral assemblage in modally decreasing proportion consists of quartz, biotite, muscovite, epidote, plagioclase and chlorite. Accessory mineral phases include opaque minerals and apatite. The analyzed mica is represented by partly chloritized biotite flakes that define a foliation S1 together with aligned and stretched epidote crystals. Chemically, the biotite contains 5.519–5.562 Si per formula unit and the white mica is muscovite (6.194 Si per formula unit). 4.2. Results and interpretation The analytical results of incremental heating experiments are listed in Table 1 and are presented as age spectra in Fig. 13, all with a corresponding 2 analytical error. Incremental heating experiments on a biotite and muscovite separates from the samples of mica-bearing low-grade metamorphic rocks resulted in a well-defined flat age spectra (Fig. 13). The results of the 40 Ar/39 Ar incremental heating experiments on white mica in sample GR114 reveal a plateau age of 41.56 ± 0.30 Ma that is close within the error to the total fusion age (40.63 ± 0.46 Ma) and normal isochron age (42.32 ± 0.22 Ma) (Fig. 13a). The incre-

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

157

Fig. 11. Metaxades-Micro-Derion area. (a) Geological map of the Metaxades-Micro Derion area; (b) cross-section on line indicated in (a); (c) asymmetric ␴-type clast derived from boudinaged quartz layer, depicting NW-directed ductile shearing; (d) asymmetric boudinage of aplitic gneiss intercalated with the amphibolites in the upper high-grade basement unit implicated in N-directed thrust contact with hanging wall overlying basalt lava flow; and (e) reactivation of this thrust contact by the overprinting late steep normal faults in the same direction.

mental heating experiments on muscovite from micaschist sample M9 resulted in a plateau age of 39.64 ± 0.11 Ma derived from ten concordant steps, whose age is very close to the total fusion age (39.90 ± 0.10 Ma) and indistinguishable, respectively, from normal and inverse isochron ages (39.75 ± 0.27 Ma; 39.76 ± 0.27 Ma) (Fig. 13b). The results from biotite in sample M10 defined a plateau age of 42.10 ± 0.83 Ma derived from six concordant steps, with a total fusion age of 36.35 ± 0.87 Ma (Fig. 13c). We interpret these new 40 Ar/39 Ar ages from the low-grade metamorphic rocks as dating the closure of the argon isotopic system in micas below the relevant temperatures (350 ± 30 ◦ C and 300 ± 20 ◦ C in white mica and biotite, respectively; e.g. Harrison et al., 1985; McDougall and Harrison, 1999). These ages therefore represent cooling ages of the rocks. In the case of the Mandritsa

area the samples apparently define cooling history of the hanging wall of the extensional fault system. 5. Discussion and interpretation 5.1. Deformational history The presented structures and kinematics and 40 Ar/39 Ar ages, combined with earlier published data (Bonev et al., 2010a), together with the knowledge on composition and origin of the ophiolites (Magganas et al., 1991; Bonev and Stampfli, 2008, 2009) and their age (Koglin et al., 2007) and available sedimentary constraints, provide regional-scale constraints on the deformational history of the Mesozoic low-grade schists. The tectonic history

158

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 12. Didymotycho area. (a) Geological map of the Didymotycho area (modified from Bonev and Stampfli, 2009); (b) photograph of thin section in the low-grade schists horizon showing NNE-directed ductile shear fabrics, minor isoclinal F1 folding and metamorphic crystalliazations in phyllitic shale of the low-grade schists horizon at the thrust contact. Half-arrows depict shear bands; (c) asymmetric ␴-type alkali feldspar clast showing SW-directed ductile shear in lower grade amphibolite-facies conditions. Note alkali feldspar alteration to epidote (centre) and white mica (strain caps). Abbreviations: Ms, muscovite; Kfs, alkali feldspar; Qtz, quartz; Cal, calcite; Ep, epidote.

of biostratigraphically proven Middle-Upper Triassic, Jurassic to earliest Lower Cretaceous (e.g. Tithonian-Berriasian, Makri unit) low-grade metasedimentary rocks and Early-Middle Jurassic ophiolites occurred in two deformational events D1 –D2 . These two events in the Mesozoic schists, however, have relationships with the deformational events in the high-grade basement as briefly discussed below to provide a general outline of the succession of deformational events in low- and high-grade basement in the region as a whole. The earlier deformational event D1 followed the deposition of protoliths of the metasedimentary rocks and the magmatic evolution of the supra-subduction zone ophiolites, whose metamorphic equivalents are represented by the Mesozoic low-grade schists. It has been shown that the earlier deformational event that is coeval with the greenschist facies metamorphism occurred in the

tectonic context of NNE-directed thrusting in Late Jurassic time (154–157 Ma) as documented by the Kulidzhik nappe (Bonev et al., 2010a). The internal deformation of all studied units and localities presented here provides additional regional-scale support to the structural and kinematic framework of this D1 north-directed thrusting event. Although reworked by the later D2 deformational event (see below), the basal tectonic contact of the Mesozoic low-grade schists preserves in places the initial thrust-related geometry. The structural pattern of bulk north-directed fold asymmetry associated with small-scale thrust faults propagation and the kinematics of ductile shear under greenschist facies conditions in the same direction are compatible with thrust-related deformation during D1 event. The development of F1 –F2 fold generations is consistent with a single phase of progressive deformation in which the transposition of bedding into the main planar-linear fabrics (S1 –L1 )

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

159

Table 1 40 Ar/39 Ar analytical data. 37Ar(ca)

38Ar(cl)

39Ar(k)

0.006460 0.001101 0.003443 0.001311 0.001733 0.000540 0.001177 0.000399 0.000470 0.000275 0.000535 0.000232 0.000667 0.000230 0.000759 0.000287 0.000853 0.000365 0.000637 0.000202 0.000296 0.000106 0.000467 0.000438

0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.000000 0.001325 0.000300 0.002377

0.000799 0.000215 0.000639 0.000445 0.001012 0.000490 0.001120 0.000794 0.001506 0.000961 0.001200 0.000532 0.001051 0.000391 0.000798 0.000497 0.001056 0.000598 0.001350 0.000715 0.000876 0.000138 0.000144 0.000111

0.235290 0.086855 0.403801 0.244766 0.722501 0.303658 0.814276 0.501083 1.002119 0.555080 0.860748 0.318256 0.676076 0.214473 0.550891 0.227568 0.687631 0.333335 0.893187 0.441715 0.616571 0.054282 0.054646 0.018272

1.970095 0.779911 3.619695 2.121883 6.167739 2.570318 6.942930 4.272037 8.538288 4.698367 7.287642 2.688014 5.703785 1.808878 4.650571 1.922704 5.830531 2.814101 7.608935 3.769838 5.283823 0.460087 0.476612 0.158726

39.22 42.03 41.96 40.59 39.98 39.65 39.93 39.93 39.90 39.65 39.66 39.56 39.52 39.51 39.54 39.57 39.71 39.54 39.90 39.97 40.13 39.70 40.84 40.68

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.69 0.92 0.31 0.34 0.26 0.26 0.20 0.21 0.18 0.21 0.19 0.24 0.20 0.33 0.20 0.35 0.20 0.24 0.19 0.23 0.22 1.05 1.07 3.40

50.70 70.40 77.86 84.32 92.04 93.85 94.92 96.99 98.07 97.97 97.55 97.19 96.34 96.06 95.09 95.47 95.54 95.99 97.26 98.11 98.05 93.35 77.36 55.00

2.18 0.80 3.73 2.26 6.68 2.81 7.53 4.63 9.26 5.13 7.96 2.94 6.25 1.98 5.09 2.10 6.36 3.08 8.26 4.08 5.70 0.50 0.51 0.17

13.193 3.758 23.986 12.228 115.987 17.099 340.463 42.000 666.036 68.014 158.316 17.189 108.314 9.517 51.086 10.702 89.234 23.369 675.376 33.711 101.200 20.074 89.336 3.767

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

20.952 4.627 40.355 17.459 526.236 27.255 4041.291 99.626 12,602.928 236.869 824.748 26.309 495.171 11.974 134.819 14.367 330.070 46.573 14,560.758 73.425 474.017 43.111 838.352 4.688

0.000082 0.000029 0.000031 0.000036 0.000009 0.000012 0.000015 0.000001 0.000011 0.000003 0.000003 0.000013 0.000015 0.000011 0.000007 0.000007 0.000005

0.000878 0.000397 0.000465 0.000733 0.000050 0.000150 0.000886 0.000302 0.000444 0.000178 0.000017 0.000073 0.000020 0.000006 0.000149 0.000072 0.000059

0.000063 0.000059 0.000065 0.000049 0.000051 0.000049 0.000063 0.000025 0.000025 0.000063 0.000018 0.000027 0.000041 0.000044 0.000000 0.000000 0.000004

0.036353 0.029855 0.035799 0.037232 0.039648 0.036422 0.036685 0.032028 0.029368 0.025100 0.021170 0.019491 0.021842 0.019704 0.010181 0.001823 0.000746

0.279018 0.241977 0.299711 0.314748 0.347008 0.322262 0.325902 0.289375 0.263300 0.227633 0.192508 0.174392 0.198230 0.178953 0.089154 0.013214 0.004022

35.99 37.98 39.22 39.60 40.98 41.42 41.59 42.29 41.97 42.45 42.56 41.88 42.48 42.51 41.00 34.00 25.35

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.32 0.69 1.10 0.99 1.29 0.37 0.68 1.18 1.42 3.01 3.24 3.72 2.66 2.12 3.53 19.83 103.20

91.65 96.21 96.68 96.39 98.95 98.63 98.37 99.62 98.44 99.36 99.26 97.47 97.57 97.88 97.51 86.83 74.77

8.39 6.89 8.26 8.59 9.15 8.40 8.46 7.39 6.78 5.79 4.88 4.50 5.04 4.55 2.35 0.42 0.17

20.282 36.882 37.686 24.898 390.268 119.009 20.281 52.008 32.402 69.115 599.336 131.210 529.207 1645.066 33.475 12.347 6.245

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

25.149 67.461 80.463 34.848 10,015.790 1510.084 16.199 366.797 61.364 383.787 1002.729 734.636 17,660.700 96,452.246 51.920 39.373 30.174

0.22930 0.04146 0.01004 0.00759 0.00311 0.00565 0.00687 0.00501 0.00825 0.00290 0.00119

0.13814 0.06686 0.01616 0.01701 0.00941 0.01673 0.02399 0.05215 0.02114 0.03122 0.01889

0.05317 0.01108 0.00333 0.00236 0.00062 0.00158 0.00176 0.00117 0.00251 0.00079 0.00063

6.47649 4.36954 1.01980 0.85477 0.29514 0.57803 0.70281 0.63784 0.95440 0.39287 0.14496

25.89 41.56 42.36 44.11 40.70 43.98 42.96 48.20 43.90 47.12 53.58

± ± ± ± ± ± ± ± ± ± ±

1.83 0.90 2.32 2.52 6.82 3.27 2.28 3.18 2.13 7.28 10.66

19.26 58.77 58.35 61.78 55.69 59.45 58.86 66.64 62.30 67.49 68.01

39.43 26.60 6.21 5.20 1.80 3.52 4.28 3.88 5.81 2.39 0.88

22.973 32.022 30.926 24.619 15.371 16.932 14.355 5.993 22.126 6.167 3.760

± ± ± ± ± ± ± ± ± ± ±

0.441 1.689 4.566 1.821 3.472 2.199 2.368 0.284 2.110 0.627 0.529

Sample GR-114 7F021D.01 870 ◦ C 7F021D.02 930 ◦ C 7F021D.03 980 ◦ C 7F021D.04 1020 ◦ C 7F021D.05 1060 ◦ C 7F021D.06 1090 ◦ C 7F021D.07 1120 ◦ C 7F021D.08 1150 ◦ C 7F021D.09 1170 ◦ C 7F021D.10 1190 ◦ C 7F021D.11 1210 ◦ C 7F021D.12 1230 ◦ C 7F021D.13 1260 ◦ C 7F021D.14 1290 ◦ C 7F021D.15 1320 ◦ C 7F021D.16 1350 ◦ C 7F021D.17 1380 ◦ C Sample M-10 5F035J.1 5F035JJ.1 5F035JJ.2 5F035JJ.3 5F035JJ.4 5F035JJ.5 5F035JJ.6 5F035JJ.7 5F035JJ.8 5F035JJ.9 5F035JJ.10

970 ◦ C 1050 ◦ C 1050 ◦ C 1120 ◦ C 1170 ◦ C 1220 ◦ C 1270 ◦ C 1320 ◦ C 1390 ◦ C 1470 ◦ C 1670 ◦ C

√ √ √ √ √ √ √ √ √ √

√ √ √ √ √ √ √ √ √ √ √

√ √ √ √ √ √

is associated with an initial minor scale isoclinal folding F1 that has been further amplified due to strain increment and accentuated by the F2 folds that rule the fold pattern. The F2 folds deform planar fabrics S1 , often mylonitic, overprinting this earlier foliation with regional-scale non-penetrative S2 –L2 fabrics in the F2 hinges. Both fold generations are clearly the result of the same deformation phase affecting in a progressive way the metamorphic layering causing its overturning even to recumbent F2 fold geometry. In this way F1 folds can be regarded as structures marking the inception of a deformational phase progressing with the F2 folds development, both formed under same greenschist facies metamorphic grade. These overprinting deformational features vary across the studied area exhibiting nearly orthogonal to the kinematic direction F2 folding with local S2 cleavage development in the Maronia

16.19667 17.61827 4.19125 3.66038 1.16506 2.46812 2.93018 2.98834 4.06775 1.79878 0.75604

40Ar(r) (%)

39Ar(k) (%)

K/Ca ± 2

36Ar(a)

Sample M-9 7A021A.01 7A021A.02 7A021A.03 7A021A.04 7A021A.05 7A021A.06 7A021A.07 7A021A.08 7A021A.09 7A021A.10 7A021A.11 7A021A.12 7A021A.13 7A021A.14 7A021A.15 7A021A.16 7A021A.17 7A021A.18 7A021A.19 7A021A.20 7A021A.21 7A021A.22 7A021A.23 7A021A.24

550 ◦ C 600 ◦ C 650 ◦ C 700 ◦ C 730 ◦ C 750 ◦ C 770 ◦ C 790 ◦ C 810 ◦ C 830 ◦ C 850 ◦ C 870 ◦ C 890 ◦ C 910 ◦ C 930 ◦ C 960 ◦ C 990 ◦ C 1020 ◦ C 1050 ◦ C 1080 ◦ C 1110 ◦ C 1150 ◦ C 1200 ◦ C 1400 ◦ C

40Ar(r)

Age ± 2 (Ma)

Incremental heating

area, the coaxial F1 refolding parallel to the F2 hinges in the Fanari area, which is coaxial to slightly oblique in the Makri unit and Mandritsa area. Coaxially refolded folds have been shown to be result of a single deformation (Ramsay, 1967). Progressive shearing and folding and even overprinting relationships during a single phase of progressive deformation are known from numerous examples (e.g. Williams, 1972, 1985; Hobbs et al., 1976; Ramsay and Huber, 1987; Passchier and Trouw, 1996). The variation of the kinematic direction from NW to N in the Mesozoic low-grade schists is apparently a regional-scale primary feature of the thrust sheet that has transported them on top of the high-grade metamorphic basement. Such changes of the tectonic transport direction are pertinent to the thrust sheets as they have been documented in several natural examples of thrust tectonics (Merle and Brun, 1984; Harris, 1985;

160

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Mesozoic schists, occurred in Tertiary times through extensional detachment faulting. Noteworthy is that the Cretaceous nappe staking event represents apparently the second deformation event (i.e. D2 ) for the high-grade basement. The D2 deformation (which for the high-grade basement is D3 ) caused warping of the hanging wall Mesozoic schists by the development of late-stage detachment surface open corrugation folds that parallels NE-SW extension direction and brittle deformation assisted by upper crustal highangle faulting (Bonev and Stampfli, 2003; Bonev, 2006b; Bonev et al., 2010a for details). The D2 deformation also caused reorientation of D1 planar-linear fabrics in the low-grade unit of the Mezosoic schists in the vicinity of the detachments. This implies that the Mesozoic schists have experienced D2 deformation passively in the hanging wall of the extensional system being progressively thinned by brittle faults assisting its removal from the uppermost metamorphic pile. Unequivocal evidence for relating D2 deformation to an extensional tectonic context comes from the direct juxtaposition by the detachments of the low-grade unit built by the Mesozoic schists onto the lower high-grade basement unit. This feature depicts an extension-related omission in the tectonostratigraphy of the upper high-grade basement unit, which the detachments cut out from the metamorphic pile. The obvious examples in the case are basal extensional tectonic contacts of the Mandritsa area, the western Mikro Derion area and especially the Dokos area (Figs. 7, 8 and 11a). The D2 event is very well radiometrically constrained by 40 Ar/39 Ar cooling ages in the hanging wall of the extensional detachment system, as well as complemented by the biostratigraphic constraints of the oldest Palaeocene-Middle Eocene clastic deposits of the hanging wall sedimentary cover unit unconformable onto the Mesozoic schists (e.g. Kopp, 1969; Papadopoulos, 1980, 1982; Dimitrova et al., 2000; Boyanov and Goranov, 2001). The 40 Ar/39 Ar ages derived from this study clearly confirm temporal spread of the D2 event in the Middle Eocene at ca. 42–40 Ma for the Mesozoic low-grade schists. These ages are undistinguishable from the available 40 Ar/39 Ar geochronology in the hanging wall of extensional detachment system (Bonev et al., 2009, 2010a; Márton et al., 2010) (see also Fig. 8), thus they provide a regional-scale temporal consistence of the D2 event. 5.2. Geodynamic context of the tectonic evolution

Fig. 13. 40 Ar/39 Ar age spectra of the samples dated in the study. For samples location and description see text and Figs. 5 and 8.

Dietrich and Durney, 1986; McClay, 1992). However, in the case of the Dokos and Mandritsa areas, a late-stage passive reorientation of the thrust transport direction towards an extension-related transport direction is fairly obvious in the hanging wall of extensional system, assisted by the displacement on the detachments and associated steep faults (Figs. 7 and 8a; e.g. Bonev, 2006b; Bonev and Beccaletto, 2007; Bonev et al., 2010a). Thus, based on structures and kinematics developed in same metamorphic grade, we relate on a regional-scale the D1 event to Late Jurassic north-directed thrust emplacement of the low-grade schists onto the Rhodope high-grade metamorphic units that contain Permo-Carboniferous igneous and likely high-grade Variscan metamorphic basement. The subsequent deformational event D2 , following the Cretaceous event of SSW-directed nappe stacking in the high-grade metamorphic basement which visibly has no record in the

The paleotectonic reconstructions of the Tethys realm in the eastern Mediterranean region (Stampfli, 2000; Stampfli et al., 2001; Stampfli and Borel, 2002, 2004; Stampfli and Hochard, 2009) have shown Late Permian-Triassic rifting and back-arc basin openings along the Eurasian plate margin related to the closure of the Paleotethys ocean leading to the widening of the Neotethys ocean. The rifting phase on this plate margin has created, adjacent to the Rhodope promontory, the Meliata and Maliac oceanic basins, whose life spans from the Triassic rifting-spreading history to the terminal Jurassic closure. The passive margins and immediately adjacent the ocean floor of these basins are well known in the Triassic-Jurassic paleogeography of the Hellenides-Dinarides and the Carpathians (e.g. Kozur, 1991; Channell and Kozur, 1997; Csontos and Vörös, 2004; Schmid et al., 2008; Papanikolau, 2009). The ophiolitic magmatic components in the Mesozoic low-grade unit (see Fig. 1), both the Evros ophiolite in Thrace (Magganas et al., 1991; Magganas, 2002; Bonev and Stampfli, 2009) and equivalent lavas and metavolcanics of the Kulidzhik nappe and the Mandritsa area in the Eastern Bulgarian Rhodope (Bonev and Stampfli, 2008; Bonev et al., 2010a), have been related to the Jurassic marginal basin-island arc evolution near the Rhodope continental margin. The southward directed intra-oceanic subduction of the Meliata-Maliac oceanic lithosphere has created the Early-Middle Jurassic Eastern Rhodope-Evros arc system in the overriding plate and established the Vardar Ocean in a back-arc

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

161

Fig. 14. Tectonic model (corresponding section is given with a bar) and Late Jurassic-Early Cretaceous 155 Ma geodynamic reconstruction (after Stampfli and Hochard, 2009). Symbols in the legend in this figure and Figs. 15 and 16: 1, passive margin; 2, magnetic or synthetic anomalies; 3, seamount; 4, intraoceanic subduction; 5, mid-ocean ridge; 6, active margin; 7, active rift; 8, inactive rift (basin); 9, collision zone; 10, thrust; 11, suture. Abbreviations in this figure and Figs. 15 and 16: AA, Austro-Alpine; Adr, Adria; Ana, Anatolides; AnT, Antalia; Ap, Apulia; Apu, Apuseni; Atl, Atlantic ocean; Bdg, Beydaghlari; BDu, Bosnia-Durmitor; Bet, Betic; Bri, Brianc¸onnais; BS, Black Sea; cRh, Circum-Rhodope; Dal, Dalmatian; EPt, eastern Pontides; Gos, Gossau; GTP, Gavrovo-Tripolitza-Pindos; Hat, Hatay; Hel, Helvetic; IzAn, Izmir-Ankara ocean; Kab, Kabylies; Kar, Karst; Lig, Ligurian; Lys, Lycian; Moe, Moesia; Men, Menderes; NCA, North Calcareous Alps; Pan, Panormides, Pel, Pelagonia; Pen, Penninic, Vahic ocean; Pie, Piemontais ocean; PIM, Paxi-Ionian-Mani; Rho, Rhodope; Rif, Rif; Sak, Sakarya; Sic, Sicani; SJa, Slavonia-Jadar; SPi, Sitia-Pindos; Sre, Srednogorie; SS, Sanadaj-Sirjan; Tau, Taurus; Tis, Tisia; Tor, Talea-Ori; TrD, Transdanubian; Tro, Troodos; Tus, Tuscan; UMr, Umbria-Marches; WCa, West Carpathian.

tectonic setting in that upper plate (Bonev and Stampfli, 2003, 2008). The magmatic history of the Eastern Rhodope-Evros arc system occurred between ca. 190-170 Ma and is bracketed by Early Jurassic radiolarian chert layers intercalated with basalts in the Mandritsa unit (Tikhomirova et al., 1988) and the crystallization of the Petrota gabbro at 169 ± 2 Ma (Koglin et al., 2007), the latter reaching shallow crustal levels at ca. 160–141 Ma (Biggazzi et al., 1989). In the north, towards the Rhodope margin, the cessation of the magmatic activity of the Eastern Rhodope-Evros arc system in the latest Middle Jurassic is indicated by the terminal deposition of trench-proximal flysch (younger detrital zircon of 160 Ma, Meinhold et al., 2010) intercalated with the Evros ophiolite in the Drimos-Melia unit (see Fig. 1). In the south, the Late Jurassic (154–155 ± 7 Ma, Tsikouras et al., 1990; 159.9 ± 4.5 Ma, Koglin et al., 2009) MORB-type Samothraki back-arc ophiolite (Tsikouras and

Hatzipanagiotou, 1998) behind the Eastern Rhodope-Thrace arc system supports its terminal magmatic evolution, when younger oceanic crust formation continued outboard and oceanward. Continued subduction and north-westward slab retreat (roll-back) of the Meliata-Maliac ocean led to the westward propagation of the Vardarian arc system (e.g. 164–158 Ma Paikon arc, Anders et al., 2005; nearby 160 ± 4 Ma Kassandra-Sithonia ophiolite, Zachariadis, 2007) in the Vardar Ocean, which totally replaced the MeliataMaliac oceanic basin by the Late Jurassic time (Fig. 14). The post-magmatic evolution of the Eastern Rhodope-Evros arc system relates to the deformation that affects the arc units together with the underlying sediments of the Rhodope continental margin. The latter are best demonstrated by the lithologies of the Makri unit and those at the base of the Mandritsa area. The D1 deformational event of northward tectonic emplacement of the thrust

162

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Fig. 15. Tectonic model (corresponding section is given with a bar) and Late Cretaceous (84 Ma) geodynamic reconstruction (after Stampfli and Hochard, 2009).

sheet built by the Mesozoic schists documents this history when the arc system arrived at the Rhodope margin. On a regionalscale we therefore relate the D1 event to arc-continental margin collision in the Late Jurassic time (e.g. Bonev and Stampfli, 2003, 2010). This tectonism is responsible for greenschist up to epidoteamphibolite facies metamorphism of the Eastern Rhodope-Evros arc units and associated sediments and its emplacement onto the Rhodope margin. The Late Jurassic Kulidzhik nappe in the leading edge of foreland propagating thrust stack also imbricates fragments of the lower unit of the high-grade basement (Bonev et al., 2010a; Fig. 14), implying the underplating of parts of the Rhodope margin pre-Variscan and older continental crust. The metamorphic grade of the Mesozoic schists indicates that the Late Jurassic thrust sheet represents a shallow crustal element of the tectonically emplaced arc-trench system. The unmetamorphosed Late Permian and Triasic detrital clastics occurring in the stratigrafically uppermost mélange-like unit of the Mandritsa area represent recycled Rhodope continental margin-derived upper crustal material that has been incorporated into the subduction-accretionary complex of the arc-trench system (e.g. Bonev and Stampfli, 2003). The low-temperature fission-track ages of the Petrota gabbro and the biostratigraphic age of the Aliki limestones are consistent with Early Cretaceous shallow crustal level emplacement and the terminal D1 event tectono-metamorphic evolution of the Makri unit. This tectonic scenario, however, opens questions for the origin of some lithological components and the tectono-metamorphic

imprint of the Late Jurassic arc-margin collisional event in the upper high-grade basement unit underlying the thrust sheet of the Mesozoic schists in the eastern Rhodope-Thrace region. The metamorphic grade of the Mesozoic schists does not match the eclogite relics and UHP assemblages in the upper high-grade basement unit, where mantle section metaophiolite bodies and lenses (harzburgites, dunites etc.) are spatially associated to the sites of the geochronological record for the UHP and eclogite facies events (see Fig. 1). Moreover, the mantle section of the Eastern RhodopeEvros arc-related ophiolites is lacking. One explanation could be that at least part of the metaophiolites within the upper highgrade basement unit may represent the missing mantle section of the Jurassic Eastern Rhodope-Evros arc-related ophiolites. The crustal assembly of the eastern Rhodope metamorphic pile above the continental affinity orthogneissic lower high-grade basement unit involved in the Late Jurassic nappe stack also requires an explanation. A plausible scenario is that different crustal level accretion occurred of the distinct parts of the Eastern Rhodope-Evros Jurassic arc system to the Rhodope contintal margin. In the case, the extrusive part of this arc together with the associated trench-slope sediments were thrust emplaced onto the margin reaching shallow crustal level, whereas its mantle section experienced deep crustal underthrusting in the footwall beneath the shallow subductionaccretionary Jurassic orogenic wedge. Because the UHP-HP relics are found only in local narrow areas, a possible channel flow in the subduction zone expulsed and/or exhumed UHP-HP rocks to mid-crustal depth beneath the upper crustal thrust sheet of the

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

163

Fig. 16. Tectonic model (corresponding section is given with a bar) and Middle Eocene (40 Ma) geodynamic reconstruction (after Stampfli and Hochard, 2009).

Mesozoic low-grade schists. The age constraints for (ultra) highpressure metamorphism (<170–160 Ma, Bauer et al., 2007; 149 Ma, Liati, 2005) both in the metamafic rocks (i.e. metaophiolites) and hosting metapelites in the upper high-grade basement unit are spatially and temporarily consistent with the latter scenario. Reported magmatic crystallization ages (159 ± 16 Ma, Peytcheva et al., 1998; 160–150 Ma, Cornelius, 2008) and metamorphic ages for the subsequent overprint of at least two medium-pressure type amphibolite facies events at 115 Ma and 79 Ma (Bauer et al., 2007) in the upper high-grade basement unit fit the proposed explanation. This scenario also accounts for the contrasting record of the metamorphic pressures and temperatures in the upper high-grade basement unit and the Mesozoic low-grade unit of the eastern Rhodope-Thrace metamorphic basement. Bonev et al. (2010a) outlined the full temporal overlap and regional consistence of the Late Jurassic north-directed thrusting manifested by the Kulidzhik nappe with respect to the Late Jurassic-Early Cretaceous thrusting of the Strandzha nappes in the same direction and with equivalent metamorphic grade (e.g. Okay et al., 2001; 170–160 Ma K/Ar ages, Lilov et al., 2004; 162–140 Ma Rb–Sr ages and 165–157 Ma 40 Ar/39 Ar ages Natal’in et al., 2005). More recently it became evident that the units limiting the SerboMacedonian Massif to the north (in contact with the CRB) were also involved in NE directed thrusting around 140–136 Ma (Kounov et al., 2010). Altogether the tectonic data unequivocally imply a Balkan orogen-wide Late Jurassic-Early Cretaceous nappe stacking event having the corresponding geochronological record of the north-northwestward thrust tectonics and blueschist metamorphism around 160–150 Ma overprinted in the greenschist facies

at 105 Ma in the type locality of the Meliata unit in the Carpathians (Dallmeyer et al., 2008). The Vardarian island arc system collided also with the Pelagonian margin (opposite to the Rhodope margin) partially closing the Vardar Ocean (Figs. 14 and 15 in its left part), and provoked a large-scale Late Jurassic-Early Cretaceous ophiolitic obduction in the Hellenides-Dinarides (e.g. Bernoulli and Laubscher, 1972; Baumgartner, 1985; Papanikolau, 2009). The structural evidence indicates that the Mesozoic schists obviously were not involved in the SSW-directed ductile thrusting that created the Late Cretaceous Rhodope syn-metamorphic nappe complex of the high-grade metamorphic basement. Undisputable evidence is that the metamorphic grade of the Mesozoic schists is limited to the uppermost greenschist facies and that the tectonic transport direction of the described Balkan belt-north vergent Late Jurassic thrusting event is opposite to that of the Vardar zone-south vergent syn-metamorphic nappe complex (Ricou et al., 1998). Bonev and Stampfli (2003) have explained this feature with the supra-crustal position acquired by the Mesozoic schists after their Late Jurassic thrust emplacement onto the Rhodope continental margin. Furthermore, the Mesozoic schists stayed in the same position within the upper plate of the Late Cretaceous subduction setting, when subduction reversal towards NNE in the Vardar Ocean (closing the post-Jurassic remnant of this ocean) occurred beneath southward trench-directed pilling up of syn-metamorphic nappes giving birth of the Sredna gora volcanic arc behind the nappe stack (e.g. Boccaletti et al., 1974) that span magmatic evolution between 92 and 78 Ma (Von Quadt et al., 2005) (Fig. 15). The subsequent Tertiary tectonic evolution of the Mesozoic schists is easily recognised because of the straightforward linkage

164

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

of the D2 deformation with middle-upper crustal extension dominated in the region at that time (Fig. 16). The extension started in Maastrichtian-Paleocene to Early Eocene times as indicated by youngest Ypresian supra-detachment syn-tectonic hanging wall clastics in syn-orogenic (i.e. syn-collisional, syn-Vardar Ocean terminal closure) context largely concomitant with still continued deeper level syn-metamorphic ductile thrusting (Bonev et al., 2006a; Bonev and Beccaletto, 2007). The 40 Ar/39 Ar geochronologic results for the Mesozoic schists in this study document Middle Eocene cooling history of the hanging wall. In the case of 40 Ar/39 Ar ages obtained from the Mandritsa area, the shear heating in a ductile-brittle shear zone beneath the brittle detachment and overburden of the Eocene sedimentary package (now preserved >1 km in thickness, e.g. Boyanov and Goranov, 2001) prevented cooling of the Mesozoic schists in the hanging wall and closure of argon isotopic system in micas until 42–40 Ma. Burial history by progressive loading of younger Middle Eocene to Oligocene sediments over the earliest Eocene sedimentary package hosting the Late Eocene (ca. 36.5–35 Ma) gold deposits persisted into Oligocene times as indicated by fission-track ages around 33-30 Ma (Márton et al., 2010). In the Maronia area, the 40 Ar/39 Ar mica age most convincingly indicates the burial history of the Mesozoic schists, which resided below the sedimentary cover package before the closure of the argon isotopic system at 41.6 Ma. It is worth nothing that in the considered region the Thrace basin has accumulated a ∼4.8 km-thick Middle Eocene sedimentary fill (Ivanov and Kopp, 1969), with occurrence of the Mesozoic low-grade schists inliers under sediments of same thickness and age in its southern part (Okay et al., 2010). The new 40 Ar/39 Ar mica geochronology indicates that the Mesozoic schists were the uppermost metamorphic unit that cooled earliest at 42–40 Ma in the hanging wall of the extensional system in the eastern Rhodope-Thrace region, with younger 40 Ar/39 Ar mica cooling ages in the range between 35.5 and 38.1 Ma in the footwall (Bonev et al., 2006b, 2009), the latter consistent with ongoing progressive mid-crustal level extension and exhumation.

cates a south-dipping subduction along the Jurassic continental margin of the Rhodope terrane. On the basis of timing constraints placed by 40 Ar/39 Ar geochronology in the Kulidzhik nappe (Bonev et al., 2010a), the low-temperature history of the Petrota gabbro (Biggazzi et al., 1989) and the Lower Cretaceous Aliki limestones (Maratos and Andronopoulos, 1964) sealing the Makri unit (first author unpublished data), the D1 event took place in Late Jurassic times. This Late Jurassic collision event is not an isolated orogenic phenomenon solely for the eastern Rhodope-Tharce region, as it can be extended to the scale of the Alpine orogen having a corresponding record in the Carpathians, in the Balkan units (Georgiev et al., 2001) northwards and in the Strandzha massif to the east (Okay et al., 2001). Thus, the described collisional event for the Mesozoic schists witnesses for the Late Jurassic Balkan orogeny. The deformation hiatus during the Cretaceous characterizes the thrust sheet of the Mesozoic schists, which stayed in supra-crustal upper plate position when subduction reversal occurred towards the north of the remnant Vardar Ocean. This position of the Mesozoic schists was inherited from the previous deformation D1 . In this way, being in the upper plate highest metamorphic pile, the Mesozoic schists escaped the overprint of the Cretaceous compressional tectonic and metamorphic processes that are conventionally attributed to amphibolite and higher grade SSW-directed nappe staking in the Rhodope high-grade basement units. Late deformational event D2 the Mesozoic schists experienced passively in the hanging wall of the Tertiary extensional system. This hanging wall deformation led to warping into open folds of the Jurassic thrust sheet, its brittle fracturing and faulting and extensional removal, which additionally was accompanied by progressive load of rather thick sediments from the beginning of the Tertiary. The timing of the D2 is bracketed by 40 Ar/39 Ar geochronology indicating hanging wall cooling below ca. 350–300 ◦ C in the Middle Eocene (42–40 Ma) and the stratigraphic constraints of the overlying sedimentary cover. At that time, burial history by thick sedimentary package and continued extensional ductile-brittle shear along the detachments prevented closing of argon isotopic system.

6. Conclusions Regional-scale structures and kinematics indicate that the deformational history of the Mesozoic low-grade schists occurred in two episodes. The main deformational event allows for the assessment of the dip of a previously poorly understood paleosubduction zone that has been involved in the tectonic history and orogenic build up of the Alpine Rhodope terrane. The earlier Late Jurassic deformational event D1 created the north-directed structural pattern and geometrically compatible fold generations during the progressive deformation and ductile shear fabrics coeval with greenschist up to epidote-amphibolite facies metamorphism. This deformation followed the Middle Triassic-Middle Jurassic depositional history of the metasedimentary protoliths near the continental margin of the Rhodope terrane and the Early-Middle Jurassic magmatic evolution of the Eastern Rhodope-Evros arc system. The southward intra-oceanic subduction of the Meliata-Maliac oceanic lithosphere created the arc system in the upper plate and established in a back-arc setting the Vardar Ocean. North-directed kinematics of internal ductile shearing and fold asymmetry associated with propagating small-scale thrusts and duplexes, the nappes and locally preserved individual thrust surfaces collectively relate this D1 event to a thrust-related deformation directed towards the Rhodope margin. The thrusting involved the arc units containing the Mesozoic low-grade schists onto the pre-Jurassic and Jurassic high-grade igneous and metamorphic basement of the Rhodope terrane, which was realized in the course of Eastern Rhodope-Evros arc-Rhodope continental margin collision. The structural evidence for this D1 event indi-

Acknowledgments This study results from post-doctoral research grant from AUF to N.B., which is gratefully acknowledged. Field work in Greece was supported by the University of Lausanne, Switzerland. Comments and suggestions by two anonymous reviewers and streaming editorial comments were highly appreciated, and helped to better shape of the manuscript. References Alvarez-Marron, J., Brown, D., Perez-Estaun, A., Puchkov, V., Gorozhanina, Y., 2000. Accretionary complex structure and kinematics during Paleozoic arc-continent collision in the southern Urals. Tectonophysics 325, 175–191. Anders, B., Reischmann, T., Poller, U., Kostopoulos, D., 2005. Age and origin of granitic rocks of the eastern Vardar Zone, Greece: new constraints on the evolution of the internal Hellenides. Journal of the Geological Society, London 162, 857–870. Barr, S.R., Temperley, S., Tarney, J., 1999. Lateral growth of the continental crust through deep level subduction-accretion: a re-evaluation of central Greek Rhodope. Lithos 46, 69–94. Bauer, C., Rubatto, D., Krenn, K., Proyer, A., Hoinkes, G., 2007. A zircon study from the Rhodope metamorphic complex N-Greece: time record of a multistage evolution. Lithos 99, 207–228. Baumgartner, P.O., 1985. Jurassic sedimentary evolution and nappe emplacement in the Argolis Peninsula (Peloponesus, Greece). Mémoires de la Société Hellevétique des Sciences Naturelles, vol. 99. Birkhäuser, Basel, 111 pp. Bébien, J., Baroz, F., Capedri, S., Venturelli, G., 1987. Magmatismes basique associes a l’ouverture d’une basin marginal dans les Hellenides internes au Jurasssique. Ofioliti 12, 53–70. Bernoulli, D., Laubscher, H., 1972. The palinspastic problem in the Hellenides. Eclogae Geologicae Helvetiae 65, 107–118.

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167 Biggazzi, G., Del Moro, A., Innocenti, F., Kyriakopoulos, K., Manetti, P., Papadopoulos, P., Norelliti, P., Magganas, A., 1989. The magmatic intrusive complex of Petrota, west Thrace: age and geodynamic significance. Geologica Rhodopica 1, 290–297. Boccaletti, G., Manetti, P., Peccerillo, A., 1974. The Balkanides as an instance of back-arc thrust belt: possible relation with the Hellenides. Geological Society of America Bulletin 85, 1077–1084. Bonev, N., 2005a. Siliciclastic knockers in marbles of the Eastern Rhodope, Bulgaria: significance for depositional environment and tectonic setting. Annuaire de l’Université de Sofia, Livre 1 Géologie 98, 29–42. Bonev, N., 2005b. Foraminifers from the exotic Late Permian limestone pebbles in the Mesozoic low-grade sequence of the eastern Rhodope, Bulgaria: paleogeographic and paleotectonic consequences. Neues Jahrbuch für Geologie und Paläontologie Monatschefte 7, 385–403. Bonev, N., 2006a. Structural and geochemical studies on amphibolite and greenschist-facies rocks in the Kulidjik river valley, eastern Rhodope, Bulgaria: preliminary results. Neues Jahrbuch für Geologie und Paläontologie Abhandlungen 239, 161–181. Bonev, N., 2006b. Cenozoic tectonic evolution of the eastern Rhodope Massif (Bulgaria): basement structure and kinematics of syn- to post-collisional extensional deformation. In: Dilek, Y., Pavlides, S. (Eds), Postcollisional Tectonics and Magmatism in the Mediterranean Region and Asia. Geological Society America Special Paper 409, pp. 211–235. Bonev, N., Beccaletto, L., 2007. From syn- to post-orogenic Tertiary extension in the north Aegean region: constraints on the kinematics in the eastern RhodopeThrace, Bulgaria-Greece and the Biga Peninsula northwest Turkey. In: Taymaz, T., Yilmaz, Y., Dilek, Y. (Eds.), The Geodynamics of the Aegean and Anatolia. Geological Society, London, pp. 113–142 (Special Publications 291). Bonev, N., Burg, J.-P., Ivanov, Z., 2006a. Structural evolution of an extensional gneiss dome—the Kesebir-Kardamos dome, E. Rhodope, Bulgaria-Greece. International Journal of Earth Sciences 95, 318–340. Bonev, N., Marchev, P., Singer, B., 2006b. 40 Ar/39 Ar geochronology constraints on the Middle Tertiary basement extensional exhumation, and its relation to ore-forming and magmatic processes in the eastern Rhodope (Bulgaria). Geodinamica Acta 19 (5), 267–282. Bonev, N., Peychev, K., Nizamova, D., 2006c. MOR-vs. SSZ-origin of metamafic rocks in the upper high-grade basement unit of the eastern Rhodope: geochemical diversity and tectonic significance. In: Proceedings of Annual Conference of the Bulgarian Geological Society, “Geosciences 2006” , pp. 181–184. Bonev, N., Márton, I., Moritz, R., Spikings, R., Marchev, P., 2009. Tectonic, magmatic and hydrothermal processes associated with Tertiary crustal extension in the basement domes of southeastern Rhodope Massif, Bulgaria. Geophysical Research Abstracts, 11, paper EGU-5304, 2 pp. Bonev, N., Moritz, R., Márton, I., Chiaradia, M., Marchev, P., 2010b. Geochemistry, tectonics, and crustal evolution of basement rocks in the eastern Rhodope Massif, Bulgaria. International Geology Review 52 (2–3), 269–297. Bonev, N., Spikings, R., Mortiz, R., Marchev, P., 2010a. The effect of early Alpine thrusting in late-stage extensional tectonics: evidence from the Kulidzhik nappe and the Pelevun extensional allochthon in the Rhodope Massif, Bulgaria. Tectonophysics 488, 256–281. Bonev, N.G., Stampfli, G.M., 2003. New structural and petrologic data on Mesozoic schists in the Rhodope (Bulgaria): geodynamic implications. Comptes Rendus Geosciences 335, 691–699. Bonev, N., Stampfli, G., 2005. Compositional diversity of the Evros ophiolite, Thrace, northeastern Greece: field occurrences, preliminary petrologic and geochemical data on plutonic sequence and tectonic implications. In: Proceedings of Annual Conference of the Bulgarian Geological Society, “Geosciences 2005” , pp. 28–31. Bonev, N., Stampfli, G., 2008. Petrology, geochemistry and geodynamic implications of Jurassic island arc magmatism as revealed by mafic volcanic rocks in the Mesozoic low-grade sequence, Eastern Rhodope, Bulgaria. Lithos 100, 210–233. Bonev, N., Stampfli, G., 2009. Gabbro, plagiogranite and associated dykes in the supra-subduction zone Evros ophiolites, NE Greece. Geological Magazine 146 (2), 71–92. Bonev, N., Stampfli, G., 2010. Comment on “Geochemistry, petrogenesis and tectonic setting of the Samothraki mafic suite, NE Greece: trace element, isotopic and zircon age constraints” by N. Koglin, D. Kostopoulos & T. Reischmann [Tectonophysics 473, 53–68 (doi:10.1016/j.tecto.2008.10.028)]. Tectonophysics 483, 413–419. Boyanov, I., 1969. Notes on the Kulidjik nappe. Bulletin of the Geological Institute of Bulgarian Academy of Sciences, series Geotectonics 18, 159–165 (in Bulgarian, abstract in French). Boyanov, I., Bodurov, K., 1979. Triassic conodonts in carbonate breccia within the low-grade metamorphic rocks of the East Rhodopes. Geologica Balcanica 9, 97–104. Boyanov, I., Goranov, A., 2001. Late Alpine (Palaeogene) superimposed depressions in parts of Southeast Bulgaria. Geologica Balcanica 31, 3–36. Boyanov, I., Ruseva, M., Toprakcieva, V., Dimitrova, E., 1990. Lithostratigraphy of the Mesozoic rocks from the Eastern Rhodopes. Geologica Balcanica 20, 3–28 (in Russian, with English Abstract). Boyanov, I., Trifonova, E., 1978. New data on the age of the phyllitoid complex from the Eastern Rhodopes. Geologica Balcanica 8, 3–21. Brown, S.A.M., Robertson, A.H.F., 2004. Evidence for Neotethys rooted within the Vardar suture zone from Voras Massif, northernmost Greece. Tectonophysics 381, 143–173. Burg, J.-P., Ricou, L.-E., Ivanov, Z., Godfriaux, I., Dimov, D., Klain, L., 1996. Synmetamorphic nappe complex in the Rhodope Massif. Structure and kinematics. Terra Nova 8, 6–15.

165

Carrigan, C., Mukasa, S., Haydoutov, I., Kolcheva, K., 2003. Ion microprobe U–Pb zircon ages of pre-Alpine rocks in the Balkan, Sredna Gora and Rhodope terranes of Bulgaria: constraints on Neoproterozoic and Variscan evolution. Journal of Czech Geological Society 48, 32–33. Cawood, P.A., Kröner, A., Collins, W.J., Kusky, T.M., Mooney, W.D., Windley, B.F., 2009. Accretionary orogens through Earth history. In: Cawood, P.A., Kröner, A. (Eds.), Earth Accretionary Systems in Space and Time. Geological Society, London, pp. 475–504 (Special Publications 318). Channell, J.E.T., Kozur, H.W., 1997. How many oceans? Meliata Vardar, and Pindos oceans in Mesozoic Alpine paleogeography. Geology 25, 183–186. Christofides, G., Pécskay, Z., Elefteriadis, G., Soldatos, T., Koroneos, A., 2004. The Tertiary Evros volcanic rocks (Thrace, northeastern Greece): petrology and K/Ar geochronology. Geologica Carpathica 55 (5), 397–409. Condie, K.C., 2007. Accretionary orogens in space and time. In: Hatcher Jr., R.D., Carlson, M.P., McBride, J.H., Martínez Catalán, J.R. (Eds.), 4-D Framework of Continental Crust. Geological Society America, pp. 145–158 (Special Memoir 200). Cornelius, N.K., 2008. UHP metamorphic rocks of the Eastern Rhodope Massif, NE Greece: new constraints from petrology, geochemistry and zircon ages. PhD Thesis. Johannes Gutenberg University, Mainz, 173 pp. Csontos, L., Vörös, A., 2004. Mesozoic plate tectonic reconstruction of the Carpathian region. Paleogeography, Paleoclimatology and Paleoecology 210, 1–56. Dabovski, C., Boyanov, I., Khrischev, Kh., Nikolov, T., Sapunov, I., Yanev, Y., Zagorchev, I., 2002. Structure and Alpine evolution of Bulgaria. Geologica Balcanica 32, 9–15. Dallmeyer, D.R., Neubauer, F., Fritz, H., 2008. The Meliata suture in the Carpathians: regional significance and implications for the evolution of high-pressure wedges within collisional orogens. In: Siegesmund, S., Fügenschuh, B., Froitzheaim, N. (Eds.), Tectonic Aspects of the Alpine-Dinaride-Carpathian System. Geological Society, London, pp. 101–115 (Special Publications 298). Del Moro, A., Innocenti, F., Kyriakopoulos, C., Manetti, P., Papadopoulos, P., 1988. Tertiary granitoids from Thrace (northern Greece): Sr isotopic and petrochemical data. Neues Jahrbuch für Mineralogie Abhandlungen 159, 113–135. Dercourt, J., Ricou, L.-E., Vrielynck, B., 1993. Atlas Tethys. Palaeoenviromental Maps. Gauthier-Villars, Paris. Dietrich, D., Durney, D., 1986. Change of direction of overthrust shear in the Helvetic nappes of western Switzerland. Journal of Structural Geology 8 (3–4), 389– 398. Dimadis, L., Papadopolous, P., Goranov, A., Encheva, M., 1996. First biostratigraphic evidence for the presence of Triassic at Melia (Western Thrace, Greece). Geologica Balcanica 26, 37–40. Dimadis, L., Nikolov, T., 1997. An ammonite find in the Makri unit (Berriasian, southeast Rhodopes, northeast Greece). Comptes Rendus de l’Academie Bulgare des Sciences 50, 71–74. Dimitrova, E., Goranov, A., Boyanov, I., 2000. New fossil evidence for the Lower Paleogene (Krumowgrad Group) in the East Rhodope Mountains, SE Bulgaria. Geologica Balcanica 30 (3–4), 59–63. Dinter, D.A., 1998. Late Cenozoic extension of the Alpine collisional orogen, northeastern Greece: origin of the north Aegean basin. Geological Society of America Bulletin 110, 1208–1230. Dinter, D.A., Macfarlane, A.M., Hames, W., Isachsen, C., Bowring, S., Royden, L., 1995. U–Pb and 40 Ar/39 Ar geochronology of the Symvolon granodiorite: implications for the thermal and structural evolution of the Rhodope metamorphic core complex, northeastern Greece. Tectonics 14, 886–908. Duffield, W., Dalrymple, G.B., 1990. The Taylor Creek rhyolite of New Mexico: a rapidly emplaced field of lava domes and flows. Bulletin of Volcanology 52, 475–487. Ferrière, J., Stais, A., 1995. Nouvelle interprétation de la suture téthysienne vardarienne d’après l’analyse des séries de Péonias (Vardar oriental, Hellénides internes). Bulletin of the Geological Society France 166, 327–339. Georgiev, G., Dabovski, C., Stanisheva-Vasileva, G., 2001. East Srednogorie-Balkan rift zone. In: Ziegler, P.A., Cavazza, W., Robertson, A.H.F., Crasquin-Soleau, S. (Eds.), Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basins and Passive Margins, vol. 186. Mémoires du Muséum National d’Histoire Naturelle, Paris, pp. 259–293. Goˇcev, P., 1979. The position of Strandja in the Alpine structure of the Balkan Peninsula. Review of the Bulgarian Geological Society 40, 27–76. Harkovska, A., Yanev, Y., Marchev, P., 1989. General features of the Paleogene orogenic magmatism in Bulgaria. Geologica Balcanica 19, 37–72. Harris, L., 1985. Direction changes in thrusting of the Schistes Lustres in Alpine Corsica. Tectonophysics 120, 37–56. Harrison, T.M., Duncan, I., McDougall, I., 1985. Diffusion of 40Ar in biotite: temperature, pressure and compositional effects. Geochimica and Cosmochimica Acta 49, 2461–2468. Haydoutov, I., Kolcheva, K., Daieva, L., Savov, I., Carrigan, Ch., 2004. Island arc origin of the variegated formations from the east Rhodope Bulgaria—implications for the evolution of the Rhodope massif. Ofioliti 29, 145–157. Hobbs, B.E., Means, W.D., Williams, P.F., 1976. An Outline of Structural Geology. Willey, New York, 563 pp. Hodges, K.V., Hames, W.E., Olszewski, W., Burchfiel, B.C., Royden, L.H., Chen, Z., 1994. Thermobarometric and 40 Ar/39 Ar geochronologic constraints on Eohimalayan metamorphism in the Dinggye area, southern Tibet. Contributions to Mineralogy and Petrology 117, 151–163. Innocenti, F., Kolios, N., Manetti, P., Mazzuoli, R., Peccerillo, A., Rita, F., Villari, L., 1984. Evolution and geodynamic significance of Tertiary orogenic volcanism in northeastern Greece. Bulletin of Vulcanology 47, 25–37. Ivanov, R., Kopp, K.-O., 1969. Zur tektonik des trakischen Altertiar-Beckens. Geotektonische Forschungen 31 (1–3), 117–132. Jaranov, D., 1960. Tectonics of Bulgaria. Technica, Sofia, 283 pp.

166

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167

Jung, D., Mussallam, K., 1985. The Sithonia ophiolite: a fossil oceanic crust. Ofioliti 10, 329–342. Kauffmann, G., Kockel, F., Mollat, H., 1976. Notes on the stratigraphic and paleogeographic position of the Svoula formation in the Innermost Zone of the Hellenides (Northern Greece). Bulletin de la Societe Géologique de France 18, 225– 230. Kilias, A., Falalakis, G., Mountrakis, D., 1999. Cretaceous-Tertiary structures and kinematics of the Serbomacedonian metamorphic rocks and their relation to the exhumation of the Hellenic hinterland (Macedonia Greece). International Journal of Earth Sciences 88, 513–531. Kockel, F., Mollat, H., Walther, H.W., 1977. Erlauterungen zur geologicschen karte der Chalkidiki und angrenzender Gebiete 1/100.000 (Nord Griechenland). Bundesanstalt für Geowissenschaften und Rohstoffe, Hanover, 119 pp. Koglin, N., Reischmann, T., Kostopoulos, D., Matukov, D., Sergeev, S., 2007. Zircon SHRIMP ages and the origin of ophiolitic rocks from the NE Aegean region, Greece. Geophysical Research Abstracts 9, paper 06848. Koglin, N., Kostopoulos, D., Reischmann, T., 2009. Geochemistry, petrogenesis and tectonic setting of the Samothraki mafic suite, NE Greece: trace-element, isotopic and zircon age constraints. Tectonophysics 473, 53–68. Kolcheva, K., Eskenazy, G., 1988. Geochemistry of metaeclogites from the Central and Eastern Rhodope Mts (Bulgaria). Geologica Balcanica 18, 61–78. Kopp, K.-O., 1965. Geologie Thrakiens III: Das Tertiär zwishen Rhodope und Evros. Annales géologiques des pays Helleniques 16, 315–362. Kopp, K.-O., 1969. Geologie Thrakiens VI: Der C¸oban Da˘g (Frenk Bunar) westlich von Alexandroupolis. Geotektonische Forschungen 31, 97–116. Koukouvelas, I.K., Aydin, A., 2002. Fault structure and related basins of the North Aegean Sea and its surroundings. Tectonics 21, 10-1–10-17. Kounov, A., Seward, D., Burg, J.-P., Bernoulli, D., Ivanov, Z., Handler, R., 2010. Geochronological and structural constraints on the Cretaceous thermotectonic evolution of the Kraishte zone, western Bulgaria. Tectonics 29, doi:10.1029/2009TC002509 (TC2002). Kozhoukharov, D., Kozhoukharova, E., Papanikolaou, D., 1988. Precambrian in the Rhodope massif. In: Zoubek, V., Cogné, J., Kozhoukharov, D., Kräutner, H.G. (Eds.), Precambrian in Younger Fold Belts—European Variscides, the Carpathians and Balkans. John Willey and Sons, Chichester, pp. 723–778. Krohe, A., Mposkos, E., 2002. Multiple generations of extensional detachments in the Rhodope Mountains (northern Greece): evidence of episodic exhumation of high-pressure rocks. In: Blundell, D.J., Neubauer, F., von Quadt, A. (Eds.), The Timing and Location of Major Ore Deposits in an Evolving Orogen. Geological Society, London, pp. 151–178 (Special Publications 204). Kozur, H., 1991. The Evolution of the Meliata-Hallstatt ocean and its significance for the early evolution of the Eastern Alps and Western Carpathians. Palaeogeography, Palaeoclimatology and Palaeoecology 87, 109–135. Liati, A., 2005. Identification of repeated Alpine (ultra) high-pressure metamorphic events by U–Pb SHRIMP geochronology and REE geochemistry of zircon: the Rhodope zone of Northern Greece. Contributions to Mineralogy and Petrology 150, 608–630. Liati, A., Gebauer, D., Wysoczanski, R., 2002. U–Pb SHRIMP-dating of zircon domains from UHP garnet-rich mafic rocks and late pegmatoids in the Rhodope zone (N Greece). Chemical Geology 184, 281–299. Lilov, P., Maliakov, Y., Balogh, K., 2004. K/Ar dating of metamorphic rocks from Strandja massif, SE Bulgaria. Geochemistry, Mineralogy and Petrology 41, 107–120. Lips, A.L.W., White, S.H., Wijbrans, J.R., 2000. Middle-Late Alpine thermotectonic evolution of the southern Rhodope Massif Greece. Geodinamica Acta 13, 281–292. Magganas, A.C., 2002. Constraints on the petrogenesis of Evros ophiolite extrusives, NE Greece. Lithos 65, 165–182. Magganas, A., 2005. Subgreenschist to greenschist facies metamorphism of metavolcanics of Circum-Rhodope Belt in Thrace. Bulletin of the Geological Society of Greece 37, 78–89. Magganas, A., Sideris, C., Kokkinakis, A., 1991. Marginal basin-volcanic arc origin of metabasic rocks of the Circum-Rhodope Belt, Thrace, Greece. Mineralogy and Petrology 44, 235–252. Maratos, G., Andronopoulos, B., 1964. Nouvelles données sur l’age des phyllites du Rhodope. Bulletin of Geological Society of Greece 6, 113–132. Marchev, P., Singer, B., Andrew, C., Hasson, S., Moritz, R., Bonev, N., 2003. Characteristics and preliminary 40 Ar/39 Ar and 87 Sr/86 Sr data of the Upper Eocene sedimentary-hosted low-sulfidation gold deposits Ada tepe and Rosino SE Bulgaria: possible relation with core complex formation. In: Eliopoulos, D.G., et al. (Eds.), Mineral Exploration and Sustainable Development, vol. 2. Millpress, Rotterdam, pp. 1193–1196. Marchev, P., von Quadt, A., Peytcheva, I., Ovtcharova, M., 2006. The age and origin of the Chuchuliga and Rozino granites Eastern Rhodopes. In: Proceedings of Annual Conference of the Bulgaria Geological Society, “Geosciences 2006” , pp. 213–216. Márton, I., Moritz, R., Spikings, R., 2010. Application of low-temperature thermochronology to hydrothermal ore deposits: formation, preservation and exhumation of epithermal gold systems from the Eastern Rhodopes, Bulgaria. Tectonophysics 483, 240–254. McClay, K.R. (Ed.), 1992. Thrust Tectonics. Chapman and Hall, London, 447 pp. McDougall, I., Harrison, T.M., 1999. Geochronology and Thermochronology by the 40 Ar/39 Ar Method, second edition. Oxford University Press, Oxford, 269 pp. Meinhold, G., Kostopoulos, D., Reischmann, T., Frei, D., BouDagher-Fadel, M.K., 2009. Geochemistry, provenance and stratigraphic age of metasedimentary rocks from the eastren Vardar suture zone, northern Greece. Palaeogeography, Palaeoclimatology and Palaeoecology 277, 199–225.

Meinhold, G., Reischmann, T., Kostopoulos, D., Frei, D., Larionov, A.N., 2010. Mineral chemical and geochronological constraints on the age and provenance of the eastern Circum-Rhodope Belt low-grade metasedimentary rocks, NE Greece. Sedimentary Geology 229, 207–233. Mercier, J., 1966. Paléogéographie, orogénèse, métamorphisme et magmatisme des zones internes des Hellénides en Macédonie (Grèce): vue d’ensemble. Bulletin de la Societe géologique de France 8, 1020–1049. Meyer, W., 1968. Alterstellung des Plutonismus im Sdteil der Rila-Rhodope-Masse. Geologica et Paleontologica 2, 177–192. Meyer, W., 1969. Die Faltenachsen im Rhodopen-Kristallin östlich des Strimon (Nordost-Griecheenland). Geotektonische Forschung 31, 86–96. Michard, A., Feinberg, H., Montigny, R., 1998. Supra-ophiolitic formations from the Thessaloniki nappe (Greece), and associated magmatism: an intra-oceanic subduction predates the Vardar obduction. Comptes Rendus de l’Academie des Sciences Paris 327, 493–499. Merle, O., Brun, J.-P., 1984. The curved translation path of the Parpaillion nappe (Frensh Alps). Journal of Structural Geology 6, 711–719. Mposkos, A., Perdikatsis, B., Liati, A., 1989. Geochemical investigation of amphibolites from eastern and central Rhodope. Bulletin of Geological Society of Greece 23 (2), 413–425. Mposkos, E.D., Kostopoulos, D.K., 2001. Daimond, former coesite and supersilicic garnet in metasedimentary rocks from the Greek Rhodope: a new ultrahighpressure metamorphic province established. Earth and Planetary Science Letters 192, 497–506. Mposkos, E., Wawrzenitz, N., 1995. Metapegmatites and pegmatites braketing the time of HP-metamorphism in polymetamorphic rocks of the E. Rhodope, northern Greece: petrological and geochronological constraints. Geological Society of Greece Special Publication 4, pp. 602–608. Mukasa, S., Haydoutov, I., Carrigan, C., Kolcheva, K., 2003. Thermobarometry and 40 Ar/39 Ar ages of eclogitic and gneissic rocks in the Sredna Gora and Rhodope terranes of Bulgaria. Journal of Czech Geological Society 48, 94–95. Natal’in, B., Sunal, G., Toraman, E., 2005. The Strandja arc: anatomy of collision after long-lived arc-parallel tectonic transport. In: Sklyarov, E.V. (Ed.), Structural and Tectonic Correlation Across the Central Asia Orogenic Collage: North-eastern Segment. Guidebook and Abstract volume of the Siberian Workshop IGCP-480, Irkutsk, Russia, pp. 240–245. Okay, A.I., Ozcan, E., Cavazza, W., Okay, N., Less, G., 2010. Basement types, Lower Eocene series, Upper Eocene olistostromes and the initiation of the southern Thrace basin, NW Turkey. Turkish Journal of Earth Sciences 19, 1–25. Okay, A.I., Satir, M., Tüysüz, O., Akyüz, S., Chen, F., 2001. The tectonics of the Strandja Massif: late-Variscan and mid-Mesozoic deformation and metamorphism in the northern Aegean. International Journal of Earth Sciences 90, 217–233. Okay, A.I., Tüysüz, O., 1999. Tethyan sutures of northern Turkey. In: Durand, B., Jolivet, L., Horváth, F., Séranne, M. (Eds.), The Mediterranean Basins: Tertiary Extension Within the Alpine Orogen. Geological Society, London, pp. 475–504 (Special Publications 156). Passchier, C., Trouw, R.A.J., 1996. Microtectonics. Springer-Verlag, Berlin, 289 pp. Papadopoulos, P., 1980. Geological Map of Greece, Scale 1:50 000 Sheet Maronia. Institute of Geology and Mineral Exploration, Greece. Papadopoulos, P., 1982. Geological Map of Greece, Scale 1:50 000 Sheet Ferai-PeplosAinos. Institute of Geology and Mineral Exploration, Greece. Papadopoulos, P., Anastasiadis, I., 2002. Geological Map of Greece, Scale 1:200 000 Sheet Rodopi-Thrace. Institute of Geology and Mineral Exploration, Athens, Greece. Papadopoulos, P., Arvanitidis, N., Zanas, I., 1989. Some preliminary geological aspects on the Makri unit (Phyllite series), Peri-Rhodope Zone. Geologica Rhodopica 1, 34–42. Papanikolau, D., 2009. Timing of tectonic emplacement of the ophiolites and terrane paleogeography in the Hellenides. Lithos 108, 262–280. Patras, D., Kilias, A., Chatzidimitriadis, E., Mountrakis, D., 1989. Structural analysis of deformation episodes and correlation with metamorphic events in the Rhodope massid and the phyllitic series of Alexandroupolis area, North Greece. Geologica Rhodopica 1, 131–138. Pe-Piper, G., Piper, D.J.W., 2002. The Igneous Rocks of Greece. The Anatomy of an Orogen. Gebrueder Borntraeger, Berlin-Stuttgart, 573 pp. Peytcheva, I., Ovtcharova, M., Sarov, S., Kostitsin, Y., 1998. Age and metamorphic evolution of metagranites from Kessebir reka region, Eastern Rhodopes—Rb–Sr isotope data. XVI CBGA, Abs., Vienna, 1 pp. Peytcheva, I., Kostitsin, Y., Salnikova, E., von Quadt, A., Kamenov, B., Klain, L., 1999. Alpine evolution of the magmatism in the West-Rhodopes: Rb–Sr and U–Pb isotope data. Journal of Conference Abstracts 4, 470. Peytcheva, I., Quadt, A.V., 1995. U–Pb zircon dating of metagranites from Byala Reka region in the east Rhodopes, Bulgaria. Geological Society of Greece Special Publication 4, pp. 637–642. Ramsay, J.G., 1967. Folding and Fracturing of Rocks. McGraw Hill, New York, 568 pp. Ramsay, J.G., Huber, M.I., 1987. The Techniques of Modern Structural Geology: II. Folds and Fractures. Academic Press, London, 700 pp. Ricou, L.-E., Burg, J.-P., Godfriaux, I., Ivanov, Z., 1998. The Rhodope and Vardar: the metamorphic and the olistostromic paired belts related to the Cretaceous subduction under Europe. Geodinamica Acta 11, 285–309. Robertson, A.H.F., 2002. Overview of the genesis and emplacement of Mesozoic ophiolites in the Eastern Mediterranean Tethyan region. Lithos 65, 1–67. Robertson, A.H.F., Dixon, J.E., 1984. Introduction: aspects of the geological evolution of the Eastern Mediterranean. In: Robertson, A.H.F., Dixon, J. (Eds.), Tectonic Evolution of the Eastern Mediterranean. Geological Society, London, pp. 1–74 (Special Publications 17).

N. Bonev, G. Stampfli / Journal of Geodynamics 52 (2011) 143–167 Robertson, A.H.F., Dixon, J.E., Brown, S., Collins, A., Morris, A., Pickett, E., Sharp, I., Ustaömer, T., 1996. Alternative tectonic models for the Late Palaeozoic-Early Tertiary development of Tethys in the Eastern Mediterranean region. In: Morris, A., Tarling, D.H. (Eds.), Paleomagnetism and Tectonics of the Mediterranean Region. Geological Society of London, pp. 239–263 (Special Publication 105). Schmid, S.M., Bernouli, D., Fügenschuh, B., Matenco, L., Schefer, S., Schuster, R., Tischler, M., Ustaszewski, K., 2008. The Alpine-Carpathian-Dinaric orogenic system: correlation and evolution of tectonic units. Swiss Journal of Geosciences 101, 139–183. Searle, M.P., Windley, B.F., Coward, M.P., Cooper, D.J.W., Rex, A.J., Rex, D., Li, T., Xiao, X., Jan, M.Q., Thakur, V.C., KumarF S., 1987. The closing of Tethys and the tectonics of the Himalaya. Geological Society of America Bulletin 98, 678–701. S¸engör, A.M.C., Yilmaz, Y., Sungurlu, O., 1984. Tectonics of Mediterranean Cimmerides: nature and evolution of the western termination of Palaeo-Tethys. In: Robertson, A.H.F., Dixon, J. (Eds.), Tectonic Evolution of the Eastern Mediterranean. Geological Society, London, pp. 77–112 (Special Publications 17). Spear, F.S., 1993. Metamorphic Phase Equilibria and Pressure–Temperature–Time Paths. Mineralogical Society of America Monograph. Book Crafters, Inc., Chelsea, MI, USA, 789 pp. Stampfli, G.M., 2000. Tethyan oceans. In: Bozkurt, E., Winchester, J.A, Piper, J.D.A. (Eds.), Tectonics Magmatism in Turkey Surrounding Region. Geological Society, London, pp. 1–23 (Special Publications 173). Stampfli, G.M., Mosar, J., Favre, P., Pillevuit, A., Vannay, J.-C., 2001. Permo-Triassic evolution of the western Tethyan realm: the Neotethys-east Mediterranean basin connection. In: Ziegler, P.A., Cavazza, W., Robertson, A.H.F., CrasquinSoleau, S. (Eds.), Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basins and Passive Margins, vol. 186. Mém. Mus. natn. Hist. Nat., Paris, pp. 51–108. Stampfli, G.M., Borel, G.D., 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters 196, 17–33. Stampfli, G.M., Borel, G., 2004. The TRANSMED transects in space and time: constraints on the paleotectonic evolution of the Mediterranean domain. In: Cavazza, W., Roure, F., Spakman, W., Stampfli, G.M., Ziegler, P. (Eds.), The TRANSMED Atlas: The Mediterranean Region from Crust to Mantle. Springer Verlag, Berlin, pp. 53–90. Stampfli, G.M., Hochard, C., 2009. Plate tectonics of the Alpine realm. In: Murphy, J.B., Keppie, J.D., Hynes, A.J. (Eds.), Ancient Orogens and Modern Analogues. Geological Society, London, pp. 89–111 (Special Publications 327). Tikhomirova, L.B., Boyanov, I., Zagorchev, I., 1988. Early Jurassic radiolarians from the Eastern Rhodopes: a revision of the age of Dolno Lukovo Formation. Geologica Balcanica 18, 58.

167

Trifonova, E., Boyanov, I., 1986. Late Permian foraminifers from rock fragments in the Mesozoic phyllitoid formation of the East Rhodopes, Bulgaria. Geologica Balcanica 16, 25–30. Trikkalinos, J.K., 1955. Über das Alter der vortartiären Schichten des Gebietes von Alexandropoulis-Didymotichon, Westthrazien. Annales géologique des Pays Hellenique 6, 81–82. Tsikouras, B., Pe-Piper, G., Hatzipanagiotou, K., 1990. A new date for an ophiolite on the northeastern margin of the Vardar Zone, Samothraki, Greece. Neues Jahrbuch für Geologie und Paläontologie Monatschefte 11, 521– 527. Tsikouras, B., Hatzipanagiotou, K., 1998. Petrogenetic evolution of an ophiolite fragment in an ensialic marginal basin, northern Aegean (Samothraki Island, Greece). European Journal of Mineralogy 10, 551–567. Turpaud, P., Reischmann, T., 2010. Characterization of igneous terranes by zircon dating: implications for UHP occurrences and suture identification in the Central Rhodope, northern Greece. International Journal of Earth Sciences 99 (3), 567–591. Von Braun, E., 1968. Die mesozoischen Hüllgesteine der SE-Rhodopen in Westthrazien (Griechenland). Geologisches Jahrbuch 85, 565–584. Von Braun, E., 1993. The Rhodope question viewed from eastern Greece. Zeitschrift der Deutschen Geologishen Gesellshaft 144, 406–418. Von Quadt, A., Moritz, R., Peytcheva, I., Heinrich, C., 2005. Geochronology and geodynamics of late Cretaceous magmatism and Cu–Au mineralization in the Panagyurishte region of the Apuseni-Banat-Timok-Srednogorie belt, Bulgaria. Ore Geology Reviews 27, 95–126. Wawrzenitz, N., Mposkos, E., 1997. First evidence for Lower Cretaceous HP/HT—metamorphism in the eastern Rhodope, North Aegean region, NE Greece. European Journal of Mineralogy 9, 659–664. Williams, P.F., 1972. Development of metamorphic layering and cleavage in low grade metamorphic rocks at Bermagui, Australia. American Journal of Science 272, 1–47. Williams, P.F., 1985. Multiple deformed terrains—problems of correlation. Journal of Structural Geology 7, 269–280. Yanev, Y., Bardintzeff, J.-M., 1997. Petrology, volcanology and matallogeny of Palaeogene collision-related volcanism of the Eastern Rhodopes (Bulgaria). Terra Nova 9, 1–8. Zachariadis, P., 2007. Ophiolites of the eastern Vardar zone, N. Greece. PhD Thesis. Johannes Gutenberg University, Mainz, 131 pp. Zagorchev, I.S., 1998. Pre-Priabonian Paleogene formations in southwest Bulgaria and northern Greece: stratigraphy and tectonic implications. Geological Magazine 135, 101–119.