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Sedimentary Geology 229 (2010) 207–223

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Sedimentary Geology

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Mineral chemical and geochronological constraints on the age and provenance of theeastern Circum-Rhodope Belt low-grade metasedimentary rocks, NE Greece

Guido Meinhold a,b,⁎, Thomas Reischmann c, Dimitrios Kostopoulos d, Dirk Frei e, Alexander N. Larionov f

a Institut für Geowissenschaften, Johannes Gutenberg-Universität, Becherweg 21, D-55099 Mainz, Germanyb CASP, University of Cambridge, West Building, 181a Huntingdon Road, Cambridge CB3 0DH, UKc Geologischer Landesdienst, Hessisches Landesamt für Umwelt und Geologie, Rheingaustr. 186, D-65203 Wiesbaden, Germanyd Department of Mineralogy and Petrology, National and Kapodistrian University of Athens, Panepistimioupoli Zographou, Athens 15784, Greecee Geological Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, DK-1350 Copenhagen K, Denmarkf Centre of Isotopic Research, A. P. Karpinsky All Russian Geological Research Institute (VSEGEI), 74 Sredny Prospect, St. Petersburg, 199106, Russia

⁎ Corresponding author. CASP, University of CamHuntingdon Road, Cambridge CB3 0DH, UK. Tel.: +44 1

E-mail address: [email protected] (G. Meinhold

0037-0738/$ – see front matter © 2010 Elsevier B.V. Aldoi:10.1016/j.sedgeo.2010.06.007

a b s t r a c t
a r t i c l e i n f o
Article history:Received 20 March 2009Received in revised form 1 June 2010Accepted 3 June 2010Available online 11 June 2010

Editor: G.J. Weltje

Keywords:Mineral chemistryZircon geochronologyHellenidesCircum-Rhodope BeltThraceGreece

In north-eastern Greece themid-greenschist faciesMakri Unit and the anchizonalMelia Formation belong to theeastern Circum-Rhodope Belt that forms the uppermost tectonostratigraphic unit of the Rhodope metamorphicnappe pile. The two metasedimentary successions had different source areas, although they now lie in closeproximity in the Rhodope Massif. The U–Pb isotopic ages of detrital zircons from a metasandstone of the MakriUnit analysed using LA–SF–ICP–MS and SHRIMP-II gave age clusters at ca. 310–290 Ma and at ca. 240 Ma formagmatic zircons, which may have been derived from Carboniferous–Permian basement rocks of the ThraciaTerrane (Lower Tectonic Unit of the RhodopeMassif) that subsequently underwent Triassic rifting. The youngestdetrital zircon grains found so far indicate that themetasedimentary succession of theMakri Unit, or at least partsof it, cannot be older than Late Triassic. By contrast, clastic sedimentary rocks of theMelia Formation contain theprimary detrital mineral assemblage of epidote, zoisite, garnet, and phengitic mica, which is absent in the MakriUnit, and clearly points to metamorphic rocks being the major source for these sediments. U–Pb analyses ofdetrital zircons gave a prominent age cluster at ca. 315–285 Ma formagmatic zircons. Inherited cores indicate theinvolvement of Pan-African and Late Ordovician–Early Silurian crustal sources during Late Carboniferous–EarlyPermian igneous event(s). Moreover, U–Pb detrital zircon geochronology indicates that the Melia Formationcannot be older than latest Middle Jurassic. We suggest that the Melia Formation was deposited in front of ametamorphic nappe pilewith Rhodopean affinities in Tithonian or Cretaceous times. Both theMakri Unit and theMelia Formation have been tectonically juxtaposed fromdifferent sources to their present location during Balkanand Alpine orogenic processes.

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1. Introduction

The Internal Hellenides are an integral part of the Alpine–Himalayan orogenic belt in south-eastern Europe and comprise,from west to east, the Pelagonian Zone (including the Attic–CycladicMassif), the Vardar Zone, the Serbo–Macedonian Massif and theRhodope Massif (Fig. 1). The latter two massifs have collectively beentreated in the literature as the hinterland of the Hellenic orogen.Kauffmann et al. (1976) described a narrow arcuate belt ofmetasedimentary rocks intermittently fringing theHellenic hinterlandto the west, south and east, and termed it the Circum-Rhodope Belt(CRB). The CRB drapes the western border of the Serbo–MacedonianMassif and continuing through the southern tip of Athos peninsula and

Samothraki Island tectonically overlies the Rhodope Massif mainly inthe south and east. Other workers have also assigned mafic andultramafic rocks of the Chalkidiki, Evros and Samothraki ophiolites tothis belt (e.g., Biggazzi et al., 1989; Magganas et al., 1991; Magganas,2002). The CRB is thought to continue both to the northwest intoformer Yugoslavia (lake Doirani area) and to the northeast into SEBulgaria and NW Turkey (Strandja Massif) (e.g., Kauffmann et al.,1976; Bonev and Stampfli, 2008, 2009; and references therein). Brunand Sokoutis (2007) recently presented a tectonic model suggesting a30° dextral rotation of the Chalkidiki block in the Cenozoic that led toexhumation of the Rhodope metamorphic core complex. Such arotation would support the Circum-Rhodope Belt concept of Kauff-mann et al. (1976). In contrast with all previous workers, Ricou et al.(1998) rejected the concept of a Mesozoic CRB covering the RhodopeMassif. According to these authors, the rocks of the CRB belong to twodistinct greenschist facies belts; they also maintained that the coversequences of the Rhodope cannot be older than Upper Cretaceous. Fordistinction purposes only, we use here the term western CRB for the

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Fig. 1. Map showing the main geotectonic zones of Greece and western Turkey (modified after Meinhold et al., 2008b).

208 G. Meinhold et al. / Sedimentary Geology 229 (2010) 207–223

western part of the belt (equivalent to the “western greenschists” ofRicou et al., 1998) and eastern CRB for its eastern part (equivalent tothe “roof greenschists” of Ricou et al., 1998). Thewestern CRB includesthe Melissochori Formation (former Svoula flysch of Kauffmann et al.,1976). The eastern CRB (“Greenschists” in Fig. 2) can be subdividedinto at least two tectonostratigraphic units (e.g., von Braun, 1993): themid-greenschist facies Makri Unit and the tectonically overlyinganchizonal Melia Formation (Fig. 3). Based on compilation ofpublished data, Pe-Piper and Piper (2002) concluded that the MakriUnit comprises Jurassic clastic sedimentary rocks and dismemberedoceanic crust that have been thrust over the Upper Tectonic Unit of theRhodope Massif. The Makri Unit was, in turn, overthrust by the MeliaFormation, which mainly includes sedimentary rocks, ranging in agefrom Early Jurassic to mid-Cretaceous (Pe-Piper and Piper, 2002).Bonev and Stampfli (2003) recently interpreted the eastern CRB as anUpper Jurassic–Lower Cretaceous allochthonous subduction–accre-tion complex of Vardarian intraoceanic island-arc origin, which wasstrongly reworked by Cenozoic extensional and strike–slip tectonics.

The strong divergence of opinions (e.g., Kauffmann et al., 1976;Papadopoulos et al., 1989; von Braun, 1993; Dimadis and Nikolov,1997; Pe-Piper and Piper, 2002) encouraged us to revisit the easternCircum-Rhodope Belt. The provenance of the metasedimentarysuccessions remains unknown, whereas their stratigraphic ageshave been the subject of a long-lasting debate. In this study, wepresent, for the first time, detrital mineral chemistry and U–Pb ages ofdetrital zircons ofmetasedimentary samples from the eastern Circum-Rhodope Belt. Our data shed new light on the stratigraphic age,original provenance and position of the studied metasedimentaryrocks and help to understand the assembly of the Eastern Mediter-ranean during the Mesozoic.

2. Geological setting

The RhodopeMassif in north-eastern Greece and southern Bulgariaconstitutes the innermost part of the Hellenides (e.g., Jacobshagen,1986) and is considered to be an Alpine nappe stack of metamorphic

Fig. 2. Geological map of north-eastern Greece (modified after Bonev et al., 2006). The locations of the study areas are indicated (see also Figs. 4 and 5).

209G. Meinhold et al. / Sedimentary Geology 229 (2010) 207–223

units derived from both continental and oceanic crust (e.g., Burg et al.,1996; Ricou et al., 1998; Barr et al., 1999). In a simplified view, twomajor tectonic units can be distinguished in the Greek part of theRhodope, consisting mainly of varying amounts of orthogneisses,paragneisses, mica schists, marbles, calc-silicate rocks, amphibolitesand eclogites (e.g., Papanikolaou and Panagopoulos, 1981; Mposkos,1989; Burg et al., 1996; Barr et al., 1999; and references therein).Geochronological studies have shown that the Lower Tectonic Unit(Thracia Terrane) is dominated by Carboniferous–Permian orthog-neisses whereas the Upper Tectonic Unit (Rhodope Terrane) predom-inantly consists of uppermost Middle–Upper Jurassic and LowerCretaceous orthogneisses (e.g., Turpaud, 2006; Cornelius, 2008;Turpaud and Reischmann, 2010). The two Terranes are separated bya major NW–SE trending, northeast-dipping thrust fault zone (Nestosthrust: Papanikolaou and Panagopoulos, 1981), which was recentlyinterpreted as a suture zone (Nestos suture: Turpaud, 2006; Turpaud

and Reischmann, 2010). It is worth mentioning that acidic rockssimilar in composition and age to those of the Thracia Terrane havebeen documented from the southern Sredna Gora Zone in Bulgaria(Carrigan et al., 2005) and the Strandja Massif in NW Turkey (Sunal etal., 2006). Structural and petrologic data have shown that the rocks ofthe Greek Rhodope experienced polyphase Middle Jurassic deforma-tion up to high-pressure (HP) conditions (e.g., Mposkos, 1989; Burg etal., 1996; Liati et al., 2002; Liati, 2005; Bauer et al., 2007). Evidence forultrahigh-pressure (UHP) metamorphism was locally found in theCentral and Eastern Rhodope (e.g., Mposkos and Kostopoulos, 2001;Liati et al., 2002; Perraki et al., 2006; Cornelius, 2008). Parts of theRhodope Massif are tectonically overlain by rock successions thatexperienced nomore than greenschist facies metamorphism (e.g., vonBraun, 1993) and were assigned to the eastern CRB (see above). Innorth-eastern Greece they include the Makri Unit and the MeliaFormationwhich are the focus of this study. Cenozoic sedimentary and

image of Fig.�2

Fig. 3. Interpretative tectonostratigraphic columns for the eastern Circum-Rhodope Belt in the Makri region (a) and Melia region (b) compiled after data in von Braun (1993) andown field observations (this work).


igneous rocks, both volcanic and plutonic in origin, occupy large partsof the study area to the south of the eastern Rhodope (Fig. 2). Isotopicdating has suggested Oligocene to Early Miocene ages for the igneous

rocks (e.g., Innocenti et al., 1984; Christofides et al., 2004). Theintensive Cenozoic magmatic activity was associated with crustalextension following the thickening/uplift of the Hellenic orogen.

image of Fig.�3


2.1. Makri Unit

TheMakri Unit occupies the larger part of the eastern CRB in the Evrosregion of Greece and corresponds to the Phyllitserie of von Braun (1968)and the Makri Unit of Papadopoulos et al. (1989), respectively. It is wellexposed and comprises a lower metasedimentary series and an uppermeta-volcano-sedimentary series (e.g., von Braun, 1968, 1993; Papado-poulos et al., 1989; Figs. 2 and 4). The metasedimentary series consists ofpale grey conglomerates, pale to dark grey sericite schists, calcschists andphyllites, and greyish limestones andmarbles. The conglomerates containclasts of quartz, quartzite, gneiss, and limestone (von Braun, 1968).Despite themetamorphic overprint itwaspossible to identify fossils in thelimestone clasts. They include mainly corals, probably of the genusCalamophyllia, as well as rotalidae and fragments of hexacorallia (vonBraun, 1968). Furthermore, corals similar to the genus Oppelismilia(Maratos and Andronopoulos, 1964a) from the uppermost Triassic andthe problematic Tubiphytes (Maratos and Andronopoulos, 1964b) havebeen described. Papadopoulos et al. (1989) have suggested that theconglomerates are the basal part of the metasedimentary seriesunconformably overlying the metamorphic basement rocks of theRhodope Massif. They interpreted the chalky and limestone- (marble-)dominated succession to be the upper part, pointing to a shallowmarineenvironment. The contact with the underlying RhodopeMassif, however,is clearly tectonic (von Braun, 1993; Ricou et al., 1998). Kopp (1969) andPapadopoulos et al. (1989) assumed a continuous stratigraphy, whereasvon Braun (1993) presented evidence that the Makri Unit is a tectonicmélange. The tectonic overprint and the scarcity ofmarker horizonsmakean estimationof its original thickness difficult. VonBraun (1968) assumedan effective thickness of around 2000 m. Dimadis and Nikolov (1997)presented a general stratigraphic column with a thickness of ∼480 m inthe area of Makri village. The rocks of the Makri Unit were metamor-phosed under mid-greenschist facies conditions (D1 episode) with thedevelopment of asymmetric, sub-isoclinal folds (Ioannidis et al., 1998). D2open folds, unaccompanied by metamorphism, overprinted the D1structures (Ioannidis et al., 1998).

The stratigraphic age of the metasedimentary series of the Makri Unithas been a matter of debate over several decades and is still a matter ofdiscussion. Fossil findings predominantly range in age from (LatePalaeozoic) Early Triassic to Late Jurassic (e.g., von Braun, 1968, 1993;Papadopoulos et al., 1989). Ioannidis et al. (1998) suggested that

Fig. 4. Geological map of theMakri area (modified after Papadopoulos et al., 1989) showing sMakri Unit (von Braun, 1993). Note: Sediments of the Melia Formation are possibly tectonicthe map since their occurrences are not well described.

sedimentation of the Makri Unit started with conglomerates inCarboniferous times and continued with deposition of limestones, shalesand marls into the Late Triassic–Early Jurassic. Dimadis and Nikolov(1997) described an ammonite specimen from sericite schists, ∼1500mnorth of Dikela village that was determined by Nikolov as Pseudosubpla-nites (Pseudosubplanites) cf. combesi (Le Hegarat) indicating the presenceof early Cretaceous (Berriasian) strata in the Makri series. Based on thisfinding, they proposed that the age of the Makri Unit spans from theTriassic to earliest Cretaceous. Von Braun (1993) however cautionedagainst tectonic imbrications between rocks of the Melia Formation andtheMakriUnit thatmayexplain thefindof theBerriasian ammonite northof Dikela (see von Braun, 1993) and suggested a pre-Cretaceous age forthe Makri Unit metasediments on the basis of the remainder of thebiostratigraphic data.

The eastern CRB comprises a variety of mafic and ultramafic rocks.Intrusive but also unclear relationships between these rocks and theMakri Unit metasediments were reported in the literature, as compiledby Pe-Piper and Piper (2002). For example, Magganas (2002) pointedout that small serpentinite and gabbro bodies in theMakriUnit probablyrepresent olistoliths. Regardless of their tectonostratigraphic position,the mafic and ultramafic rocks include gabbroic cumulates, metagab-bros and serpentinised harzburgites, which correspond to the “LowerMetavolcanics” of Magganas (1988, 2002) and Magganas et al. (1991),whereas greenschists of volcanic origin and pillowed or massivemetavolcanic and metapyroclastic rocks correspond to the “UpperMetavolcanics” of Magganas (1988, 2002) and Magganas et al. (1991).The latter group of rocks belongs to the volcano-sedimentary series (so-called “Greenschist Series”) of theMakri Unit. Dykes and plagiogranitesdescribed by Bonev and Stampfli (2005) correspond to the mafic lavasunderlying theMelia Formation and the stock-like plagiogranite body atDidymotycho further east at theGreek–Turkish border respectively. Thedykes together withmafic massive and pillow lavas aremembers of theDrimos–Melia Unit (Papadopoulos et al., 1989). Plagiogranites andmafic lava flows display mutually intrusive relationships near Didymo-tycho. Since these lavas are underlainbyphyllites that strongly resemble(in terms of lithology and isoclinal folding) those in the “GreenschistSeries” of the Makri Unit, the phyllites can be correlated with the MakriUnit. TheDidymotycho lavas in turn apparently correspond to themaficlavas underlying the Melia Formation, although age constraints are stilllacking.

ample localities of theMakri Unit. The Aliki limestone is always in tectonic contact to theally imbricated with the Makri Unit (von Braun, 1993). However, they are not shown in

image of Fig.�4


Magganas (1988) considered themafic andultramafic rocks aspartsof an incomplete and dismembered ophiolite complex that he namedthe Evros ophiolite. Geochemical signatures classify the mafic rocks asmid-ocean ridge basalt (MORB) and island-arc tholeiite (IAT) whereastransitional IAT to MORB affinities also exist that are consistent with anisland-arc/back-arc tectonic setting (Magganas et al., 1991; Magganas,2002; Bonev and Stampfli, 2005). Furthermore, boninitic affinity isevident in somemafic lavas, metavolcanic rocks and dykes of the Evrosophiolite in Bulgaria (Bonev and Stampfli, 2008, 2009). An intraoceanicvolcanic island-arc tectonic setting was recently suggested for maficrocks from the Bulgarian part of the eastern CRB (Bonev and Stampfli,2003, 2008, 2009). Previous apatite fission-track ages of 140±46 Maand 161±31Ma from gabbros near Petrota village (Biggazzi et al.,1989) provided cooling ages for the Evros ophiolite below ∼120 °C.Recently, sensitive high-resolution ion microprobe (SHRIMP-II)analyses of zircons from a gabbroic sample nearMaronia gave a precisemean 206Pb/238U age of ca. 169±2 Ma that was interpreted as thecrystallisation age of the gabbro (Koglin et al., 2007).

Middle and Upper Eocene deposits, including conglomerates,nummulitid-bearing limestones, sandstones and marls with tuffa-ceous layers, rest unconformably on the mid-greenschist facies rocksof theMakri Unit (Kopp, 1965; von Braun, 1968), thus constraining theupper time limit of sedimentation andmetamorphism for those rocks.

2.2. Melia Formation

TheMelia Formationmainly crops outwest and northwest ofMeliavillage (Figs. 5 and 6) and corresponds to the “terrain de transition” ofViquesnel (1868), the “Grauwackenserie” of Wirth (1940) and theDrimos–Melia Formation of Dimadis et al. (1996) respectively. Thelatter authors were the first to use the lithostratigraphic termFormation and we will keep using it here since it is a mappablesedimentary succession. We have, however, dropped the constituentpart “Drimos” as it refers to a small abandoned village. Note that theMelia Formation is underlain bymassive lavas, pillow lavas and dykes,which are affiliated to the Evros ophiolite (e.g., Magganas, 2002).Papadopoulos et al. (1989) suggested that the Melia Formation coversunconformably theMakri Unit; von Braun (1993), however, presentedconvincing evidence that the two sedimentary successions areseparated by a tectonic contact. The Melia Formation dominantly

Fig. 5. Geological map of the Melia area (modified after von Brau

consists of ∼1000-m-thick siliciclastic rocks, including dark greyshales, fine- to coarse-grained sandstones and minor conglomeratelayers (e.g., von Braun, 1968, 1993; Kopp, 1969). The sandstone bedscan reach up to 2 m in thickness and are apparent inmany places (e.g.,along the main road north of Loutros village). The bedding has ageneral dip of ∼30–40° to the south and southeast. In general, theMelia Formation has a flysch-like character (von Braun, 1993).

The age of the Melia Formation is still a matter of discussion.Trikkalinos (1955) found a tectonically deformed ammonite (stepha-noceratid or perisphinct) near Melia village that was assigned to theMiddle–Late Jurassic transition (Callovian–Oxfordian). Kopp (1965)reported a bivalve mollusc of the genus Pecten of the Early or MiddleJurassic from the drilling site “Evros 1” near Ardanion village.Papadopoulos et al. (1989) suggested a Late Jurassic to early Cretaceousage for the Melia Formation. Dimadis et al. (1996), however, reportedtwo thin-shell bivalve specimens from a locality NW of Melia identifiedasHalobia superbaMojsisovicz,which led them to conclude that at least apart of the Melia Formation is Late Ladinian–Early Carnian in age,although this species is characteristic for the Late Carnian–Early Norian(McRoberts, 1993). Curiously, a year later, Dimadis and Nikolov (1997)suggested that the Melia Formation cannot be older than earlyCretaceous with the upper time limit of sedimentation being mid-Cretaceous.Unfortunately, it is verydifficult for us, andmostprobably forother workers as well, to test the validity of these fossil findings and thecorrectness of their stratigraphic assignment to the Melia Formation. Inthe course of this study, an ichnofossil, most likely Ophiomorpha rudis(seeUchman,2009),was foundwestofMelia (Fig. 6b). TheOphiomorpharudis ichnosubfacies is present from the Tithonian onwards, with anincrease of abundance after the Cenomanian and an apparently declineafter the Eocene (Uchman, 2009). Furthermore, near Melia, coalifiedterrestrial plant fragments were found on bedding surfaces (Figs. 6c, d).Unfortunately, the lack of characteristic features necessary for anaccurate age determination (V. Wilde pers. comm. 2005) hindered usfrom obtaining a biostratigraphic age for these plant fragments.

Kopp (1969) and von Braun (1993) mentioned that the clasticsuccession of the Melia Formation is gradually overlain by the Alikilimestone, although the latter has never been described from the areaaround Melia. Therefore, the relationship between the Aliki limestoneand the Melia Formation is still an open question. Nonetheless, themicritic matrix of limestones south of Aliki contains calcalgae and

n, 1993) showing sample localities of the Melia Formation.

image of Fig.�5

Fig. 6. Melia Formation west of Melia village (40°57′07.0″N, 26°06′12.1″E). (a) Sole marks (flute casts) on the sole of a sandstone bed. (b) Ophiomorpha rudis ichnofossil (see Uchman,2009) on the sole surface of a siltstone bed. (c, d) Terrestrial plant fragments on bedding surfaces.


foraminifera of the families Miliolidae and Textularidae (von Braun,1968). The Miliolidae have been identified as Quinqueloculine sp. andMassilina sp. and the Textularidea as Trocholina sp.; these indicate anearly Cretaceous age for the Aliki limestone (von Braun, 1968; Kopp,1969). Thus, if the gradual transition into the Aliki limestone iscorrect, the upper time limit for deposition of the clastic sedimentaryrocks of the Melia Formation seems to be early Cretaceous; however,at any rate, the upper limit for deposition of the Melia Formation isclearly set by the unconformably overlying Middle Eocene clasticdeposits (see Kopp, 1965; von Braun, 1968). A basal Middle Eoceneconglomerate, cropping out west of the now extinct Drimos village,consists, among others, of clasts of grey limestone similar to thoseexposed near Aliki. The limestone clasts contain foraminifera such asLenticulina, Miliolidae, Orbitolina, Textularia and Pseudocyclammina aswell as gastropod molluscs of the genus Nerinea (Kopp, 1965; vonBraun, 1968). The occurrence of Orbitolina indicates that uppermostLower Cretaceous–lowermost Upper Cretaceous limestones wereexposed in their source area during the Paleogene. Nowadays,however, no rocks of such an age are preserved in the Evros region.Interestingly, a calcareous clastic succession of such an age hasrecently been discovered in the Oreokastro area of the eastern VardarZone (Meinhold et al., 2009). This suggests that sediments charac-teristic of the Lower to Upper Cretaceous transition were once morewidespread in northern Greece than nowadays preserved.

3. Samples and methods

A total of twenty siliciclastic sedimentary rocks including sand-stones, siltstones and phyllites (metapelites) were collected from theMakri Unit and the Melia Formation for whole-rock major- and trace-

element geochemical analyses, detrital mineral chemistry and zircongeochronology. Sample localities (including geographic coordinates)and the analytical data referred to in this paper are included in theaccompanying Supplementary data (see Appendix A). Methodologiesof the mineral chemical analysis and zircon geochronology aredescribed below; the methodologies of the whole-rock geochemicalanalysis are outlined in Meinhold et al. (2007) and Nehring et al.(2008). Moreover, whole-rock major- and trace-element geochemicaldata are only given as Supplementary data (see Appendix A) and arenot used for further discussion because of the possibility of secondarymodification due to diagenesis and metamorphic overprint.

Sandstones of the Makri Unit are poorly sorted by size and pre-dominantly consist of angular to subangular fragments of quartz (bothmono- and polycrystalline), white mica and Na-rich feldspar (albite)(Fig. 7a).Minor chlorite, hematite and calcite (in varying amounts) arealso present. Plagioclase is often sericitised. Some samples contain asecondary matrix of sericite (e.g., sample R27). Accessory mineralsmainly include zircon, rutile, tourmaline and opaques.

Sandstones of the Melia Formation are also poorly sorted by sizeand dominated by angular to subangular fragments of quartz (bothmono- and polycrystalline), feldspar (mainly albite) and white mica(Fig. 7b). Furthermore, detrital minerals are epidote, zoisite, garnetand chlorite, which are characteristic for sandstones of the MeliaFormation but absent in sandstones from the Makri Unit. Plagioclasegrains commonly display patchy alteration due to sericitisation andreplacement by calcite. The psammites of the Melia Formation arechiefly clast-supported; primary matrix, if present, consists mainly ofsericite. Accessory minerals include zircon, rutile, titanite andopaques. It is noteworthy that in outcrops near Melia villageterrestrial plant fragments are macroscopically visible (Figs. 6c, d).

image of Fig.�6

Fig. 7. Thin section photomicrographs of samples used for zircon geochronology taken in plane light (left) and with crossed polarizers (right). (a, b) Sample R27 from theMakri Unit.(c, d) Sample R15 from the Melia Formation.


At other localities, e.g., along the main road from Loutros to Dadia, noplant fragments could be found macroscopically. However, organicmaterial (plant cuticula) could be identified in thin sections.

3.1. Mineral chemistry

Several studies have shown that microprobe analysis of detritalmineral phases can potentially reveal the identity of source rocks(e.g., Morton, 1991; von Eynatten andGaupp, 1999;Mange andMorton,2007; Meinhold et al., 2008a). The chemistry of detrital feldspar, mica,and garnet was determined in polished thin sections using an electronmicroprobe (JEOL JXA 8900RL) equipped with 5 wavelength-dispersivespectrometers at the University of Mainz. Operating conditions were15 kV acceleration voltage with a beam current of 8 nA and a beamdiameter of 5 μm for feldspar. Mica was analysed with 15 kVacceleration voltage, a beam current of 12 nA and a beam diameter of2 μm.Anaccelerating voltage of 20 kVand a beamcurrent of 20 nA,witha beam diameter of 2 μm were used for garnet. Natural and syntheticmaterials were used as standards. The PRZ procedure was applied tocalculate concentration units. Cations of feldspar, mica, and garnet werecalculated stoichiometrically, based on 8, 11, and 12 oxygens performula unit, respectively.

3.2. Zircon geochronology

U–Pb dating of detrital zircons has proved to be a powerful tool insedimentary provenance studies. In the absence of fossils and otherstratigraphic data, the youngest detrital zircon grain in a sedimentaryrock provides a maximum limit to the age of deposition (e.g., Fedo et

al., 2003; and references therein). U–Pb zircon geochronology wascarried out on two samples, one from the Makri Unit, the other fromtheMelia Formation. Each samplewas crushed using a hydraulic pressand rotary mill and sieved to a grain-size smaller than 0.5 mm. Theheavy mineral fraction was separated using a Wilfley table, a Frantzisodynamicmagnetic separator and heavy liquids (methylene iodide).Final purification was achieved by hand-picking under a binocularmicroscope. Zircon grains were set in epoxy resin mounts, sectionedand polished to approximately half their original thickness. Prior tothe analyses, cathodoluminescence (CL) images were obtained for allgrains in order to study their internal structure and to target specificareas within them, e.g., growth structures and inherited cores.

Measurements were carried out in two different sessions using twodifferent facilities, namely SHRIMP-II and LA–SF–ICP–MS. SHRIMP-IImeasurements were performed at the Centre of Isotopic Research, St.Petersburg, Russia, following the methods described in Meinhold et al.(2008b). LA–SF–ICP–MSmeasurements were performed at the Geolog-ical Survey of Denmark and Greenland, Copenhagen, Denmark,following the techniques described in Frei and Gerdes (2009) andMeinhold et al. (2010). Accuracy and reproducibility of the appliedLA–ICP–MSmethodwere tested by collecting and processing U–Pb dataon 6 zircon grains previously dated by SHRIMP-II. Furthermore, chips ofthe TEMORA 1 zircon standard (Black et al., 2003) were treated asunknown during the LA–ICP–MS analytical session.

Unless otherwise stated, 206Pb/238U ages are used for zircon grainsyounger than 1.2 Ga whereas older grains are quoted using their 207Pb/206Pb ages (e.g., Meinhold and Frei, 2008). Concordia plots andprobability density distribution diagrams were produced using theprograms Isoplot/Ex (Ludwig, 2003) and AgeDisplay (Sircombe, 2004).

image of Fig.�7


Ages reported in the text are given at the 2-sigma level. The GeologicalTime Scale (GTS) of Gradstein et al. (2004) was used as stratigraphicreference for data interpretation.

4. Results

4.1. Mineral chemistry

4.1.1. FeldsparThe analysed detrital feldspar grains from the Makri Unit and the

Melia Formation show slight differences in their chemistry that isillustrated in a ternary diagram showing the three end-membercomponents NaAlSi3O8 (albite, Ab), KAlSi3O8 (orthoclase, Or) andCaAl2Si2O8 (anorthite, An). Feldspar grains from the Makri Unit aremainly albite (Ab97–100An0–3Or0–2). By contrast, feldspar grains from theMelia Formation show a trend towards oligoclase (Ab5–99An0–15Or0–92).Notably, two grains plot in the field of anorthoclase and one in the fieldof orthoclase (Fig. 8).

4.1.2. White micaThe chemical composition of detrital white mica in metasediments

from theMakri Unit and the Melia Formation revealed a broad range inSi content from ∼3.01 to 3.44 a.p.f.u. and in Al content from ∼2.03 to2.85 a.p.f.u. (Fig. 9a). The fact that the analyses do not fall along themuscovite–celadonite mixing line but form a trend sub-parallel to itsuggests the presence of ferrimuscovite substitution. All analysedwhitemicas are K-rich with K contents of ∼0.70 to 0.95 a.p.f.u.; paragonite isnot present. Si contents N3.0 a.p.f.u. suggest the presence of phengiticmica (Fig. 9b).

4.1.3. GarnetBesides epidote and zoisite, garnet is also an abundant detrital heavy

mineral in themetasedimentary rocks of theMelia Formation. In general,the chemical variability of detrital garnet from sedimentary rocks is auseful provenance indicator (e.g., Morton, 1985, 1991; Takeuchi, 1994;von Eynatten and Gaupp, 1999; Morton et al., 2004; Mange and Morton,2007). However, secondary processes can substantially modify garnetgeochemistry or lead to total garnet dissolution and this possibility has tobe considered when dealing with garnet provenance analysis (e.g.,Morton, 1991; Morton and Hallsworth, 1999). These processes could act,for example, during burial through circulation of high-temperature porefluids (e.g., Morton, 1991; Morton and Hallsworth, 1999). Furthermore,the stability of garnet is also controlled by its composition, with high-Cagarnets being less stable than low-Ca garnets (e.g., Morton, 1987). Garnetis especially characteristic ofmetamorphic rocks of awidevariety of types,and it can be also found in magmatic rocks such as granites, pegmatites,acidic volcanic rocks, kimberlites and somemetasomatic rocks (e.g., Deeret al., 1992).

Fig. 8. Ab–An–Or ternary diagram of detrital feldspar from the Makri Unit (sample R27,n=50) and the Melia Formation (sample R15, n=60) analysed in polished thin sections.

Fig. 9.Mineral chemistryofdetritalmica from theMakriUnit (sampleR27,n=40)and theMelia Formation (sample R15, n=34) analysed in thin sections. (a) Binary plot of Sicontent vs. Al content; (b) Binary plot of Si content vs. Mg content after von Eynatten andGaupp (1999); the field for phengitic mica formed under HP metamorphic conditions isaccording to Massonne and Schreyer (1987).

Detrital garnets in metasedimentary rocks of the Melia Formationoccur generally as discrete grains and have mainly an angular shape.Most grains were analysed at the core with some grains additionallyanalysed at the rim. Calculation of end-member garnet compositionsrevealed that detrital garnets from the Melia Formation are rich inalmandine component [42–74 mol% almandine (Alm, XFe), 3–39 mol%grossular (Grs, XCa), 2–37 mol% pyrope (Prp, XFe) and 1–23 mol%spessartine (Sps, XMn)], the exception being two grains with highgrossular content (Fig. 10a). Core–rim analyses indicate an almosthomogeneous composition of single grains (Fig. 10b). Secondary

image of Fig.�8

image of Fig.�9

Fig. 11. Concordia diagram showing the results obtained by LA–SF–ICP–MS analyses ongrains of the TEMORA 1 zircon standard.


alteration or primary zonation is negligible relative to the overallvariability of garnet composition considered here. Morton et al. (2004)distinguished threemajor garnets types for source rockcharacterisation,namely garnets with low-Ca and high-Mg contents (Field A in Fig. 10),Fe–Mn-rich garnets, with variable Ca and low Mg contents (Field B inFig. 10), and garnets with high-Ca and high-Mg contents (Field C inFig. 10). According to this classification, most detrital garnets from theMelia Formation (excluding the two Grs-rich grains) fall in the Fields C(60%) and B (26%).Most of the type B garnets areMn-rich (XMnN5), butdo not plot close to the Fe+Mn apex. Type A garnets form only 14% ofthe population.

4.2. Zircon geochronology

To test the precision and accuracy of the LA–SF–ICP–MS techniqueapplied, thirteen analyses have beenperformedon chips of the TEMORA1zircon standard (Blacket al., 2003). The results are shown in Fig. 11. All theresults are concordant U–Pb ages and define a concordia age of 416.7±3.5 Ma (2σ), which is well within the range of the ID-TIMS analyses of416.8±1.1 Ma (spike uncertainty included) reported by Black et al.(2003).

Fig. 12 is a plot of LA–SF–ICP–MS vs. SHRIMP-II U–Pb ages obtainedon the same zircon grains. Some LA–SF–ICP–MS analyses were setexactly onto the SHRIMP-II craters, whereas others were positioned onthe diametrically opposite side of the zircon grain (Fig. 13). In general,U–PbagesobtainedbyLA–SF–ICP–MSare in goodagreementwith thosepreviously obtained by SHRIMP-II, which confirms previous studiesdemonstrating the successful application of LA–ICP–MS to U–Pb zircondating (e.g., Gerdes and Zeh, 2006). Nevertheless, significant differencesin the ages obtained by the two methods on the same spot (e.g., 216±7 Ma and 161±10 Ma, spot R15-79, Fig. 13) are occasionally observed.This is attributed to the different depth sampled by the two methods(∼30 μm for LA–ICP–MS vs. ≤3 μm for SHRIMP-II), implying thatLA–ICP–MS dating is more sensitive to inhomogeneities within a singlezircon crystal. Where both SHRIMP-II and LA–SF–ICP–MS ages are

Fig. 10. Ternary diagramof the end-member proportions of detrital garnet from theMeliaFormation analysed in polished thin sections. Themain composition is Alm42–74Gr3–39Py2–37Sp1–23, excluding two garnet grains with a very high grossular component. (a) Onlyanalyses at the grain core are shown. (b) Analyses of core–rim pairs are shown. Note thatrims are closest analysis to cores in all cases. FieldsA, B andC are fromMortonet al. (2004).

available for the same spot only the LA–SF–ICP–MSdatawill be used forinterpretation since they form the dominant data set of this study.

4.2.1. Makri UnitSample R27 was collected from ametasandstone succession cropping

out on the main road from Makri to Dikela (40°51′10.0″N, 25°44′25.3″E; Fig. 4). Most detrital zircons from this sample are clear and colourlessand have a subhedral shape; few are euhedral or rounded. Some zirconsare rich in mineral inclusions (e.g., apatite, quartz); others arecharacterised by fractures filled with iron hydroxides and appear dirtybrown in parts. Gas or melt channels similar to those illustrated by Corfuet al. (2003) have been found in one crystal. The length of single zirconcrystals varies between 100 and 300 μm. Most of the analysed zirconshave clear oscillatory zoning patterns in CL images (Fig. 13) and appear tobe magmatic in origin; only a few crystals exhibit weak or no zoning.Inherited cores are sparse. Altogether, 35 single zircon grains have been

Fig. 12. Concordia diagram showing the results obtained by LA–SF–ICP–MS (filledsymbols) and SHRIMP-II analyses (open symbols) on the same zircon grains. Error ellipsesin concordia plot represent 2σ uncertainties. All grains are shown in CL images on Fig. 13.Thedata showthatU–Pb ages obtainedbyLA–SF–ICP–MSare inverygoodagreementwiththosepreviously obtained by SHRIMP-II,which confirmsprevious studies (e.g.,Gerdes andZeh, 2006; Frei and Gerdes, 2009) demonstrating the successful application of LA–ICP–MSto U–Pb zircon dating.

image of Fig.�10

image of Fig.�11

image of Fig.�12

Fig. 13. CL images of representative zircon grains from analysed samples with location of analysis spot and corresponding 206Pb/238U age (±2σ), the exception being spot R15-68with corresponding 207Pb/206Pb age (±2σ). Note: SHRIMP-II analyses are shown by a white circle with ages given in italics. The remaining are LA–SF–ICP–MS analyses. Letter–number code above the ages: sample-spot. The scale bar represents 30 μm in all images.


image of Fig.�13


analysed for their U–Pb ages using SHRIMP-II (7 analyses) and LA–SF–ICP–MS (31 analyses). Grains that were analysed by both methods areshown in Figs. 12 and 13 and are marked accordingly in theSupplementary data (see Appendix A). Detrital zircons from sample R27show a polymodal age distributionwith amajor cluster between 310 and290Ma, aminor cluster at ca. 240 Ma, and single ages at 376, 343, 262 and214Ma (Fig. 14). The youngest grains are rounded (R27-23) and euhedral(R27-24) and have 206Pb/238U ages of 214±6Ma and 233±6Ma,respectively, thus giving a maximum age of deposition for themetasandstone sample R27 and hence for the Makri Unit. Zircon grainsof Middle Triassic age are predominantly euhedral. The (100) prism ofthese zircons dominates over the (110) prism with only the (101)bipyramid being present, corresponding to types P4 and P5 of Pupin(1980). Such a morphology is typical of an igneous origin and suggestscrystallisation temperatures of about 800–850 °C in an alkaline magma.Only a few crystals of Carboniferous age are sufficiently euhedral to usethem for determining their morphological type. There seems to be twodifferent groups of zircon types with Carboniferous ages. In both groups,the (101) prism of the euhedral zircons dominates over the (100) prism,or they occur in equal proportions. In the first one, however, the (211)bipyramid dominates over the (101) bipyramid, or the latter issubordinate, corresponding to types Q1, Q2, Q3, S6 and S11, whereas inthe second only the (101) bipyramid is present, corresponding to types

Fig. 14. Concordia and probability density distribution and histogram plots showing the resufor detrital zircons from samples R27, Makri Unit (a, c) and R15, Melia Formation (b, d). Erroplots: Dark grey shaded— zircon ages with 95–105% concordance, pale grey shaded— zirconn=number of analyses.

G1, P1 and P2. Such morphologies are typical of an igneous origin andsuggests crystallisation temperatures of about 650–750 °C in a metalu-minous magma for group I and in a subalkaline magma for group II.

4.2.2. Melia FormationSample R15was collected from a sandstone succession cropping out

at Melia (40°57′07.0″N, 26°06′12.1″E; Fig. 5). Most detrital zircons ofthis sample are clear and colourless and have a subhedral to roundedshape. Some zircons are characterised by fractures filled with ironhydroxides and thus appear dirty brown in parts. The length of singlezircon crystals varies between 110 and 300 μm. CL images reveal twotypes of detrital zircons. Zircons of type I have inherited coreswith clearoscillatory zoning patterns suggesting a magmatic origin, and youngerovergrowths of brighter luminescence, which exhibit no zoning(Fig. 13). The latter typifies zircon growth and/or recrystallisation inhigh-grade metamorphic rocks (e.g., Corfu et al., 2003). The zirconovergrowths in sample R15, however, have Th/U ratios that are higherthan the value generally accepted as the lower limit ofmagmatic values(≥0.4 vs. 0.1 respectively; see Teipel et al., 2004, and references therein)suggesting a non-metamorphic origin or dry metamorphic conditions.Zirconsof type II donot have inherited cores. They showclear oscillatoryzoning patterns or a weak zoning and appear to be magmatic in origin(Fig. 13). Altogether, 33 single zircon grains have been dated by the

lts obtained by LA–SF–ICP–MS (grey symbols) and SHRIMP-II analyses (black symbols)r ellipses in concordia plots represent 2σ uncertainties. Probability density distributionages with N5% discordance; Only data with≤10% discordance are shown. Abbreviation:

image of Fig.�14


U–Pb technique using SHRIMP-II (7 analyses) and LA–SF–ICP–MS (36analyses). Grains that were analysed by both methods are shown inFigs. 12 and 14 and are marked accordingly in the Supplementary data(see Appendix A). Detrital zircons from sample R15 show a polymodalage distribution with a major cluster between 315–285 Ma and minorclusters around 550 and 450 Ma (Fig. 14). The Pan-African and LateOrdovician–Early Silurian ages belong to inherited cores. The youngestgrain is subhedral and has a late Middle Jurassic age (concordantSHRIMP-II age of 161±10 Ma; slightly discordant LA–SF–ICP–MSage of163±5 Ma; see Fig. 12), thus giving a maximum age of deposition forthe sandstone sample R15 and hence for the Melia Formation. It isnoteworthy that one grain has an inherited core with a highlydiscordant 206Pb/238U age of ca. 2388Ma and a magmatic rim of ca.315 Ma (Figs. 13 and 14b). In a concordia diagram, a regression throughthe core and the rim yields a discordia with a lower intercept at 312.8±7.6 Ma and anupper intercept at 3100±19 Ma (Fig. 14b). The latter ageis in good agreementwith the 207Pb/206Pb age of ca. 3059 Ma that has tobe considered as a minimum age due to the effect of Pb loss. In general,lower intercept ages do not necessarily have a geological meaning andshould therefore be handled with care (Mezger and Krogstad, 1997). Inthis case, however, the lower intercept of the zircon rim is a concordantage andcanbe interpretedas the timeof secondary zircon crystallisationfrom a magma. The upper intercept reflects the time of primary zirconcrystallisation, and is interpreted as the age of formation of themagmatic protolith that survived as the inherited core and henceindicates the presence of a minor Mesoarchaean component.

5. Discussion and conclusions

5.1. Makri Unit

Mineral chemistry indicates that detrital feldspar is dominantlyalbite. A prominent content of detrital albite could reflect sedimentsupply from spilites (Na-metasomatised basalt by circulating seawater/brines) where it is the most characteristic mineral (Deer et al., 1992).Such a source, however, should also be seen in the whole-rockgeochemical signature (e.g., elevated Cr and Ni values), which is notthe case for the analysed metasedimentary rocks of the Makri Unit (seeAppendix A). The source of detrital albite has to be sought probably inmetapelitic rocks of the chlorite and biotite zones since albite is thestable plagioclase in these zones (Deer et al., 1992). Minor input mayhave come from magmatic albite. Metasomatism of detrital K-feldsparfollowing lithification has to be kept in mind as well. In general, theapplication of detrital feldspar composition to provenance analysis is oflimited significance.

The erosion of rocks that have formed under different temperatureandpressure conditions canexplain the large spread in Si contents of thedetrital K-rich white micas from the Makri Unit. Si contents N3.3 forsomemicas suggest that thesemineralswere formed undermedium- tohigh-pressure metamorphic conditions, depending on their primarymineral assemblage and crystallisation temperature in the source rocks(Massonne and Schreyer, 1987; Parra et al., 2002). The low-Si phengiticmica may have derived from magmatic or low-pressure metamorphicrocks. The abundanceof intermediate tohigh-Si phengiticmicapoints tomedium- to high-pressure rocks exposed in the source area during thetime of deposition. Recycling of detrital mica from older sedimentarysuccessions is also possible. The compositional range of the detritalwhitemicas could also be explained by retrograde overprinting ofwhitemicas derived from primary high-pressure metamorphic rocks. At anyrate, the chemistry of detrital mica clearly indicates supply of detritusderived from high-pressure metamorphic rocks during the time ofdeposition of the Makri Unit.

The prominent cluster of detrital zircon ages at ca. 310–290Masuggests significant input from Upper Carboniferous and Lower Permianmagmatic rocks. Similar ages have been reported as protolith ages fororthogneisses from the Thracia Terrane (e.g., Turpaud, 2006; Cornelius,

2008; Turpaud and Reischmann, 2010), which constitutes the closestpossible source area. Therefore, it seemsplausible to assume that Thracianorthogneisses or their equivalentswere themajor contributors of detritusto the clasticmetasedimentary succession of theMakri Unit. The source ofthe second cluster of detrital zircon ages at ca. 240 Mamay also be soughtin the Rhodope since Cornelius (2008) recently described detrital zirconsof Middle and Late Triassic ages from a garnet–kyanite gneiss collectedabout 16 km NE of Komotini. However, this sample also contains detritalzircons of Late Jurassic to Cretaceous ages, and zircon rims even yieldedPaleogene ages, which are absent from the Makri Unit sample. Anothergarnet–kyanite gneiss collected north of Xanthi contains detrital zirconswith ages between 205±8Ma and 223±10Ma (Liati and Gebauer,2001). Metamorphic rims have an age of 138±6Ma and show lead lossdue to a younger metamorphic event probably during the Eocene (Liatiand Gebauer, 2001). This garnet–kyanite gneiss encloses eclogite pieceswith a protolith age of 139±4Ma and similar lead loss history (Liati andGebauer, 2001). Although it is unlikely that these high-grade metasedi-ments have supplied detrital material for the clastic sediments of theMakri Unit, their protoliths may have. It can also be speculated that theMakri Unit sediments are time and facies equivalent deposits to theprotoliths of these high-grade metasediments. Possible source rockssupplying Triassic zircons may also have been amphibolitised eclogiteswith an age of about 245 Ma from the Central Rhodope (Liati, 2005) andTriassic igneous rocks of 240 Ma and 249Ma from the Serbo–MacedonianMassif in SW Bulgaria (Zidarov et al., 2004; Peytcheva et al., 2005).Moreover, zircons from the voluminous Arnea suite, which comprisespredominantly A-type granitoids of Middle to early Late Triassic age(Himmerkus et al., 2009a) may have supplied detritus but this is thoughtunlikely since Silurian ages from the host country rocks (VertiskosTerrane: Himmerkus et al., 2006, 2007, 2009b; Meinhold et al., 2010) areabsent in the age spectrum of detrital zircons from the metasandstone ofthe Makri Unit. Finally, the youngest detrital zircon grain found so farindicates that the metasedimentary succession of the Makri Unit, or atleast parts of it, cannot be older than Late Triassic.

5.2. Melia Formation

Detrital feldspar grains from theMelia Formation include, in additionto albite, subordinate amounts of oligoclase andminor anorthoclase andorthoclase. The large range in feldspar composition observed reflects thelarge range inpossible source rock lithologies. As itwas the casewith thedetrital feldspar in the Makri Unit, the source of detrital albite has to besought probably in metapelitic rocks of the chlorite and biotite zonessince albite is the stable plagioclase in these zones (Deer et al., 1992).The detrital mineral assemblage of epidote, zoisite and garnet togetherwith phengitic mica indicates that metamorphic rocks were the majorsource for clastic metasedimentary rocks of the Melia Formation. Thepresence of oligoclase suggests derivation from acidic to intermediateigneous rocks; orthoclase and anorthoclasemayhavebeenderived fromacidic igneous rocks (Deer et al., 1992). Erosion of rocks that have beenformed under different temperature and pressure conditions canexplain the large spread in Si content of detrital K-rich white micasfrom the Melia Formation. The low-Si phengitic mica may have beenderived from magmatic or low-pressure metamorphic rocks. Theabundance of intermediate to high-Si phengiticmica points tomedium-to high-pressure rocks exposed in the source area during the time ofdeposition. Recycling of detrital mica from older sedimentary succes-sions is another possibility. The compositional range of the analyseddetrital white micas could also be explained by retrograde overprintingofwhitemicas derived fromprimary high-pressuremetamorphic rocks.Whatever the case, the chemistry of detrital mica clearly indicatessupply of detritus derived from high-pressure metamorphic rocksduring the time of deposition of the Melia Formation. Detrital phengitecould have been derived from paragneisses, mica schists (metapelites)and orthogneisses of the Rhodope Massif since phengite is abundant inthese rocks (e.g., Mposkos, 1989; Mposkos and Liati, 1993).


Mostof theanalysedgarnets fromtheMelia Formationare almandine-rich with minor pyrope and grossular components. According to theclassificationofMortonet al. (2004),mostof thegarnets (60%)plot infieldC of Fig. 10, and were sourced from high-grade mafic rocks and quartz–biotite gneisses. An input from high-grade mafic rocks, however, seemsmost unlikely since accompanying minerals such as amphibole and sodicclinopyroxene are not present as a framework component in thesuccession of the Melia Formation, although it cannot be excluded thattheir absence is just a result of weathering. In general, almandine garnetswithpyrope contents from10 to40 mol%occur in epidote–amphibolite togranulite facies gneisses (e.g., Takeuchi, 1994). Field B garnets comprise26%of the assemblage.Most of themareMn-rich (XMnN5), butdonotplotclose to the Fe+Mn apex, suggesting that they are most likely derivedfrom low- tomedium-grademetasedimentary rocks (Mange andMorton,2007), rather than from granitoids. Type A garnets form only 14% of thepopulation. Their ultimate source is suggested to be high-grade (granulitefacies) metasedimentary rocks or charnockites (Morton et al. 2004).However, they are also reported from intermediate-acidic igneous rocks(Mange and Morton, 2007; and references therein). Taking into accountthe accompanying heavy mineral assemblage, the detrital garnets fromthe Melia Formation have probably been derived from mica schists andgneisses, which experienced higher greenschist- to lower amphibolite-facies regional metamorphism. The two grossular-rich garnets (Type Dgarnet according to Mange and Morton, 2007) were most likely derivedfrom contact or thermally metamorphosed impure calcareous rocks(Deer et al., 1992).

It becomes evident from the above that themain sources of theMeliaFormationweremetasedimentary rocks. Therefore, it seemsplausible tosuggest that most of the detrital zircons have also been derived fromsuch rocks and thus represent recycledmaterial, although at presentwedo not have unequivocal evidence for such an assumption. Theprominent cluster of detrital zircon ages at ca. 315–285 Ma suggestssignificant input from Upper Carboniferous and Lower Permianmagmatic rocks. The inherited cores indicate the involvement of Pan-African and Late Ordovician–Early Silurian crustal sources during LateCarboniferous–Early Permian magma genesis. Acidic igneous rocks ofsimilar composition have been documented from the southern SrednaGora Terrane in Bulgaria (Carrigan et al., 2005). In this area, zircons fromleucogranites of Late Carboniferous and Early Permian age contain 80–90% inherited cores with ages of ∼900–600 Ma, ∼450 Ma and ∼400–300 Ma (Carrigan et al., 2005). Furthermore, Upper Carboniferous–Lower Permian orthogneisses containing zircons with abundant EarlyPalaeozoic and Pan-African inherited cores have recently beendiscovered in the Thracia Terrane of the Eastern Rhodope (Cornelius,2008). Note that an input from the Serbo–Macedonian Massif in SWBulgaria, which contains metagranites of Late Ordovician age (Zidarovet al., 2003;Macheva et al., 2006), and also from theVertiskos Terraneofthe Serbo–Macedonian Massif in Greece, which contains Late Ordovi-cian–Silurian ages (Himmerkus et al., 2006, 2007, 2009b; Meinholdet al., 2010), seems unlikely because no Late Carboniferous–EarlyPermian ages have been detected so far overprinting these EarlyPalaeozoic ages in that area. Therefore, it seemsplausible to assume thatleucogranites from the Sredna Gora Terrane or protoliths of orthog-neisses from the Thracia Terrane or equivalent rocks contributedsignificant amounts of detritus to a clastic sedimentary successionprobably duringpost-Early Permian–Triassic erosion. After its formationthis succession was overprinted under upper greenschist- to loweramphibolite-facies metamorphic conditions before it was uplifted andexposed at the surface to supply detritus to the clastic succession of theMelia Formation in pre-Eocene times. Note that most of the rocks of theRhodope Massif experienced their high-grade metamorphic overprintfrom the Jurassic onwards, and theirfinal exhumation to the surfacewasin Cenozoic times (e.g., Reischmann and Kostopoulos, 2002; Liati, 2005;Bonev et al., 2006; Bauer et al., 2007). Nonetheless, under conditions offast exhumation (as fast as subduction; Rubatto and Hermann, 2001;Baldwin et al., 2004), it is possible to bring high-grade metamorphic

rocks from great depths (N70 km) back to the surface in less than5 million years (Baldwin et al., 2004). With this in mind, it cannot beexcluded that some high-grade metamorphic rocks from the RhodopeMassif had alreadybeen at the surfaceonly a fewmillion years after theirmetamorphic overprint, thus supplyingdetritus to the clastic successionof the Melia Formation. The youngest detrital zircon grain (ca. 160 Ma)found so far indicates that the rocks of the Melia Formation cannot beolder than latestMiddle Jurassic. This grainmay have been derived fromorthogneisses of theRhodope Terrane since these rocks areMiddle–LateJurassic in age (Turpaud, 2006; Turpaud and Reischmann, 2010). Inputfrom the Evros ophiolite, which contains also felsic magmatic members(plagiogranite, felsic dykes) and associated volcano-sedimentarymaterial (e.g., Magganas, 2002; Bonev and Stampfli, 2008, 2009), maybe another possibility. However, detrital input of felsic material fromthis ophiolite would have been accompanied most likely by detritalinputof (ultra)maficmaterial. This should be reflected in thepresence ofdetrital chrome spinel and elevated whole-rock geochemical values ofCr and Ni (cf. Meinhold et al., 2009; Meinhold and BouDagher-Fadel,2010), which is not the case here (see Appendix A). At the present stateof knowledge, no further constraints can be put on the age and sourcesof the Melia Formation sediments.

5.3. Palaeotectonic implications

5.3.1. Makri UnitThe sedimentary successions of theMakri Unit andMelia Formation

are quite different from each other in terms of stratigraphic age andprovenance, although they now lie in close proximity in the RhodopeMassif. Furthermore, the Melia Formation does not unconformablyoverlie theMakri Unit. The contact between them is tectonic, as alreadyemphasised by von Braun (1993). The youngest detrital zircon grainfrom theMakri Unit indicates that at least part of it cannot be older thanLate Triassic. The conglomerates, sandstones, siltstones and limestones(nowmarble) probably belonged to a sedimentary succession that wasdeposited undermarine conditions (e.g., Papadopoulos et al., 1989). Theeuhedral crystal shape of many zircon grains and the narrow range ofCarboniferous zircon ages coupled with the rarity of pre-Carboniferousdetritus indicate proximity to a Carboniferous source area of relativelyhomogeneous composition. At thepresent state of knowledge, the lattermay have been basement rocks from the Thracia Terrane or equivalentrocks (Fig. 15a). The upper time limit of deposition is difficult toconstrain, but it may be given by the ∼170 Ma-old Evros ophiolite. Wesuggest the following scenario for the Makri Unit: continental shelfsedimentation (=deposition of rocks of the Makri Unit), rifting, oceanbasin formation with later intraoceanic subduction, ophiolite emplace-ment (of the Evros ophiolite) and ocean basin closure. Implicit to theabove scenario is that the Makri Unit is older than ca. 170 Ma (theformation age of the Evros ophiolite) but younger than ca. 214 Ma, i.e. itis probably latest Triassic and/or Early Jurassic in age. Younger fossilfindings described from the Makri Unit (e.g., Dimadis and Nikolov,1997) may be explained by tectonic imbrications with lithologies fromthe Melia Formation (see von Braun, 1993).

5.3.2. Melia FormationU–Pbdetrital zircon geochronology indicates that theMelia Formation

cannot be older than latest Middle Jurassic. Dimadis and Nikolov (1997)even assume that it cannot be older than early Cretaceous with the uppertime limit of sedimentation being mid-Cretaceous, although neitherstratigraphic nor radiometric data exist in support of this age. However,the upper time limit of sedimentation and metamorphism of the MeliaFormation is clearly set by the unconformably overlying Middle Eoceneclastic deposits (see Kopp, 1965; von Braun, 1968). Our new data suggestthat the clastic sedimentary succession of the Melia Formation receivedlarge amounts of detritus frommetasedimentary rocks andminor igneousrocks, which probably experienced higher greenschist- to lower amphib-olite-facies regional metamorphism. The sedimentary input of terrestrial

Fig. 15. Schematic sketches illustrate the depositional settings of sedimentary rocks from the Makri Unit and the Melia Formation (see text for explanation). The Makri Unit wasdeposited at a passivemargin, probably the Thracia Terrane in (latest Triassic and) Early Jurassic times. TheMelia Formation was deposited in front of ametamorphic nappe pile withRhodope affinities probably in latest Jurassic or Cretaceous times.


plant fragments indicates proximity to a source area with vegetation. Theangular to subangular shape of the framework components, the poorsorting, and the high amount of unstable lithoclasts (plagioclase) indicaterapid erosion and depositionwithout significant rounding or sorting. Thisand the narrow range of zircon ages suggest deposition in a rapidlysubsiding sedimentary basin accompanied by rapid uplift in the sourcearea, which was most likely the Rhodope Massif (Fig. 15b). The input ofmetamorphic detritus may correspond to microplate reorganisation andprogressive uplift of a part of the Rhodope metamorphic complex. Takencollectively, the existing data suggest deposition in a developing forelandbasin in front of a metamorphic nappe pile with Rhodope affinitiessometime between the latest Jurassic and Eocene, probably during theTithonian or Cretaceous. Whether deposition of the Melia Formation isalso related to the Middle–Late Jurassic intraoceanic Vardarian volcanic-arc environment, which existed to the south of the Serbo–Macedonian–Rhodope Massifs (e.g., Bonev and Stampfli, 2008, 2009), is still an openquestion. Although it maybe suggested that the siliciclastic turbidites ofthe Melia Formation were deposited in a trench in front of theaforementioned Middle–Late Jurassic volcanic-arc, we would cautionagainst this, as long as no other data for the stratigraphic age of thesesediments are available.

In order to gain a regional perspective, it is important to discuss theCircum-Rhodope Belt sensu Kauffmann et al. (1976) and how thesedimentary rocks of the Makri Unit and those of the Melia Formationare incorporated. As already mentioned in the introduction, Ricou et al.(1998) rejected the concept of a Mesozoic Circum-Rhodope Belt as thesedimentary cover of the RhodopeMassif. According to those authors, therocks of the CRBbelong to twodistinct greenschist facies belts, and are notolder than late Cretaceous. Here we only discuss the metasedimentaryrocks since they are the focus of this study. According to the datapresented herein and those in Meinhold et al. (2009) regarding the ageand provenance of metasedimentary rocks from the western CRB, theCircum-Rhodopebelt sensuKauffmannet al. (1976)has tobe re-evaluated(cf. Ricou et al., 1998). Only the sedimentary succession of theMakri Unit

resembles that of the Melissochori Formation (former Svoula flysch) ofthewestern Circum-Rhodope Belt, as already noted by Pe-Piper and Piper(2002), although detrital zircon ages show that the MelissochoriFormation had a different source area, with a prominent cluster of zirconages at ca. 370–290Ma, and subordinate clusters at ca. 520–450, 655–545Ma and 2100–1780Ma (Meinhold et al., 2009). Nonetheless, theMakri Unitmay represent a time and facies equivalent of theMelissochoriFormation that was deposited at quite some distance from the latter in aNeotethyan oceanic basin along a margin exposing mainly Carboniferousigneous basement rocks and limestones of a Triassic carbonate platform(Fig. 15a), probably of the Thracia and Strandja terranes. The detritalzircons with Triassic ages can be related to the widespread Triassic rift-related magmatism observed in many terranes of the Eastern Mediter-ranean (e.g., Pe-Piper and Piper, 2002; Zidarov et al., 2004; Liati, 2005;Peytcheva et al., 2005; Himmerkus et al., 2007, 2009a; Meinhold et al.,2009, and references therein). Rifting led to formation of severalintracontinental basins some of which merged together and developedinto oceanic basins (e.g., Stampfli and Borel, 2002; Stampfli et al., 2003,and references therein). In stark contrastwithKauffmannet al. (1976),wewant to re-emphasise that the sedimentary rocks of the Melia Formationcannot be correlated with those of the Melissochori Formation. The latterhave quite different source areas and sedimentation ages (Meinhold et al.,2009). Structural studies provided evidence that the Makri Unit firstlyunderwent N- to NW-directed thrusting (von Braun, 1993). Similarfeatures were described by Bonev and Stampfli (2003) from Mesozoicschists of the eastern CRB in Bulgaria and were ascribed to an earlythrusting event, which could represent deformation in a subduction–accretion complex. North-directed Late Jurassic thrusting has also beendescribed from the Strandja nappes (about 60 km east from the easternCircum-Rhodope Belt) involving greenschist to lower amphibolite-faciesmetamorphosed Carboniferous–Permian igneous and metamorphicbasement and its Triassic–Middle Jurassic cover succession (Okay et al.,2001;Natal'in et al., 2005). FollowingBonev and Stampfli (2003), it seemsreasonable to suggest that the Makri Unit together with the Evros

image of Fig.�15


ophiolitewere thrust over theRhodopeprobably during the Late Jurassic–Early Cretaceous. This may correspond to the collision event that wasresponsible for the Late Jurassic–Early Cretaceous N-directed Balkanorogeny (e.g., Bonev and Stampfli, 2003, 2008, 2009). However, thetectonic position of the Melia Formation within this scenario is still anopen question. What is clear is that the sedimentary succession of theMelia Formation was derived from a metamorphic source (a nappe pile)with Rhodopean affinities in post-Middle Jurassic times. Later, the rocks ofthe eastern Circum-Rhodope Belt together with the underlying basementof the Rhodope were involved in late-orogenic extensional tectonicsduring late Alpine history in Cenozoic time.

Acknowledgements

This study was supported by the German Research Foundation (DFG)and the state of Rhineland-Palatinate through the Graduiertenkolleg 392“Composition and Evolution of Crust and Mantle”. Laboratory facilities attheMax-Planck-Institute for Chemistry inMainz, at the Centre of IsotopicResearch in St. Petersburg and at the Geological Survey of Denmark andGreenland in Copenhagen are gratefully acknowledged. We thank P.Haughton for his comments on an early version of this manuscript and N.Bonev for his final review, which improved the quality of this paper andhelped us to clarify the interpretations and their formulation. This paper ispublished with the permission of the Geological Survey of Denmark andGreenland.

Appendix A. Supplementary data

Supplementary data associated with this article can be found, in theonline version, at doi:10.1016/j.sedgeo.2010.06.007.

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