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Mediterranean Geoscience Reviews

, Volume 1, Issue 1, pp 91–161 | Cite as

The phanerozoic palaeotectonics of Turkey. Part I: an inventory

  • A. M. Celâl ŞengörEmail author
  • Nalan Lom
  • Gürsel Sunal
  • Cengiz Zabcı
  • Taylan Sancar
Original Paper
  • 502 Downloads

Abstract

This paper presents an inventory of the relevant information to delineate the Tethyside sutures and the continental blocks they stitch in Turkey and to summarise their history. A palinspastic palaeogeographic/palaeotectonic interpretation is reserved for the second part of this paper and in a third paper we hope to deal with its neotectonic episode. In Turkey there are two groups of Tethyside sutures: Cimmeride sutures corresponding to the closure of the Palaeo-Tethys and its dependencies such as back-arc basins and Alpide sutures corresponding to the welding lines of the Neo-Tethyan oceans. During the Palaeozoic, the entire area of Turkey was a part of the northern margin of Gondwana-Land, along which subduction seems to have been continuous at least since the Devonian, possibly since the Ordovician. During the early Carboniferous a Lycaonian back-arc basin disrupted this margin from the western end of the country to the eastern part of the Taurus Mountains. In the latest Permian another extensional basin began opening the Karakaya Basin, which seems to have been coeval and in places confluent with the western segment of the northern branch of the Neo-Tethys in Turkey. This rifting largely dislodged the future pieces of the Sakarya Continent and the Rhodope–Pontide Fragment from Gondwana-Land. Another region of Permian extension rimmed the present-day eastern Mediterranean and extended via the Bitlis Suture in southeastern Turkey into the Zagros collision zone in Iran. Finally, the body today constituting the Sakarya Continent rifted off the southern margin of the Rhodope–Pontide Fragment opening what has been called the Intra-Pontide Ocean. These various rifting events created only two independent continental pieces: the Sakarya Continent and the Kırşehir Block, the latter originally attached to the eastern end of the Bitlis Massif. Subduction commenced during the Aptian all along the southern margin of the Sakarya Continent and Laurasia where the Rhodope–Pontide Fragment formed a Sumatra-type continental margin arc. Successive rifting events north of this arc eventually opened the Black Sea, isolating the Rhodope–Pontide Fragment from Laurasia. In the main body of the text we provide the data on which such interpretations are based and evaluate some alternative suggestions.

Keywords

Turkey Tethysides Palaeo-Tethys Neo-Tethys Collision tectonics 

Introduction

The first general synthesis of the geological evolution of Turkey in terms of plate tectonics was attempted in 1981 (Şengör and Yılmaz 1981). It was based in large part on the earlier synthesis by Ketin (1966a); the last such attempt before the kinematic tenets of the plate tectonics theory was almost completely worked out by 1969 (Wilson 1965a, b, 1966, 1968; McKenzie and Parker 1967; Le Pichon 1968; Morgan 1968; Isacks et al. 1968; McKenzie and Morgan 1969). The Şengör and Yılmaz synthesis was the fruit of an attempt to explicate, in plate tectonic terms, almost two centuries of work by numerous Turkish and foreign geologists and geophysicists on Turkish problems. When it was attempted, the pre-Mesozoic geology of Turkey and the surrounding areas was little understood and that is why it narrated the geological history of Turkey from the Triassic onwards. It turned out to be surprisingly successful, given the paucity of reliable micropalaeontological, mesoscopic structural and isotopic age data on which it was based. Despite the enormous amount of work done since then and the numerous syntheses attempted, the only major modifications to the scheme presented in Şengör and Yılmaz has been the discovery of the early Carboniferous Lycaonian marginal basin within the Menderes-Taurus Block, separation of the Kırşehir Massif from the rest of what had been called the Anatolide/Tauride Platform in Şengör and Yılmaz (1981) along an Inner Tauride Suture (Şengör et al. 1982) and the discovery of the doubling of the Kırşehir back onto itself as first pointed out by Sengör and Natal’in (1996). All other changes mainly pertain to the details of the shapes of the first- and second-order tectonic units earlier established, the timing of events and a greater recognition of the significance of strike-slip movements during the palaeotectonic episode in Turkey, i.e., before the medial Miocene.

The success of the Şengör and Yılmaz synthesis we think lay in the methodology employed: as a first step, the sutures along which various Tethyan oceans and their branches and dependencies such as marginal basins were identified and then the geological evolution of the continental fragments they surround was worked out using stratigraphy, structural geology and the record of igneous and metamorphic events that had affected them. There were then sufficient numbers of observations to do this on a period- and in many places even epoch-basis, although the age data used were mainly palaeontological. Since then there has been a very large increase in our stratigraphic and petrological/geochemical database including high-precision zircon isotopic age dates. In the face of that rich database that accumulated and a number of attempted partial refutations of the scheme presented in Şengör and Yılmaz (1981)—(e.g., to cite only some of the general syntheses: Robertson and Dixon 1984; Göncüoğlu et al. 1997; Okay and Tüysüz 1999; Altiner et al. 2000; Stampfli 2000; Moix et al. 2008; Okay 2008; Robertson et al. 2012, 2013; we found none of these satisfactory despite including many excellent observations and much high-quality data) we have decided to attempt a new synthesis. Only Ketin’s textbook on the geology of Turkey (Ketin 1983; see his pages 541–551) and Yılmaz (1990) adopted Şengör and Yılmaz’s (1981) scheme. The present paper sets forth the first product of our renewed endeavour. Its purpose was to refine the suture distribution and timing presented in Şengör and Yılmaz (1981) and to improve the precision of the history of the continental blocks separated by them. A detailed palaeogeographic story was here not attempted, as it would greatly expand the size of an already oversize paper making it difficult to accommodate in a journal. That story is reserved for the second part of this paper to be published later in which the palaeotectonic evolution will be presented in time-lapse frames showing also the palaeogeography. Finally, a third paper will present a synthesis of the neotectonics of Turkey. We here wish to show what is where today and what it was before the medial Miocene. One important difference from the Şengör and Yılmaz synthesis is that we here include part of the late Palaeozoic history of the country in our narrative in places reaching as far down as into the early Palaeozoic, although we did not include the Palaeozoic rocks in our stratigraphic columns because of the dispersed nature of their outcrops. Some unexpected clarification resulted from that regarding the geology of the western continuation of the geology of Turkey into northern Greece and Bulgaria and southern Romania.

In the following sections, we first briefly outline the methodology we follow and then describe the sutures and the continental pieces surrounded by them. In the discussion we present the main messages this paper gives and outline what are considered as remaining problems. In the following, we cite only what we thought are publications reporting critical observations. Many such observations were published by their authors many times in different places. We usually tried to cite only the earliest or the most comprehensive of those. No doubt, we probably missed some critical publications and would be grateful to those who would draw our attention to them. In the following, Fig. with a capital ‘F’ refers to our own figures; fig. with a lower-case ‘f’ refers to figures we cite from the literature. We use the correct term ‘Gondwana-Land’ for the large southern continent and ‘Gondwana’ for the eponymous historical district in India after which a number of geological entities had been named (such as ‘Gondwana rifts’). We do not include a locality map, because all of our many localities are contained in Google Earth®, where they can be viewed almost at any desired scale.

Methodology of regional geology followed in this paper

Turkey houses a number of collisional orogens within the Tethyside superorogenic complex, the product of the mainly subductive elimination of the Tethyan realm. Figure 1a–c shows schematic maps of the Tethyan realm, its place in the structure of Eurasia and Turkey’s place in it. The ‘Tethyan realm’ contains the remnants of the Palaeo-Tethyan ocean and its continental margins (the oceanic remnants of the Palaeo-Tethys are mostly in the form of ophirags,1 there are no known major Palaeo-Tethyan ophiolite nappes in Turkey), the Cimmerian Continent (Şengör 1979, 1984) and of the various branches of the Neo-Tethyan ocean, the vestiges of which include giant ophiolite nappes in the Taurus Mountains in the south of the country as well as major coloured mélange complexes including ophirags of diverse sizes and compositions all over Turkey, plus the continental margins of the Neo-Tethys.
Fig. 1

a Tectonic map of Asia showing its orogenic systems. The boxed area is the region dealt with in this paper. b The place of the Tethysides in the world and their simplified suture distribution. The boxed area is the region dealt with in this paper. c A highly schematized map showing the idealized organization of the Tethyan realm with the place of Turkey indicated

In mapping sutures, we took the older scheme by Şengör and Yılmaz (1981) and Şengör et al. (1982) as a basis and systematically went through the more recent literature plus our own published and unpublished observations to see what improvements seemed necessary. Figure 2 shows our results, which we briefly review in the next section. Figure 3 shows the principle we followed in drawing the suture map. We then circumscribe the blocks surrounded by the sutures. Figure 4 shows two cross sections across western (A–A’ and B–B’) and central (C–C’) Turkey to illustrate the presumed deep geometry of the sutures and the highly simplified internal geometries of the continental blocks delimited by them. We did not draw a cross-section across eastern Turkey, because the extensive Miocene to Quaternary volcanic cover makes it impractical to draw one, although a schematic one is given in Şengör et al. (2008a, b). Figure 5 illustrates 400 high-quality isotopic ages from the sutures documenting almost exclusively the spreading ages of the ophiolites plus some emplacement ages by dating their metamorphic soles. Figure 5 also gives a quick overview of these ages by means of a colour scale (please see the supplementary information for details of isotopic ages). Here we were able to corroborate the judgement of Şengör and Yılmaz (1981) that a continental block distribution valid from the Palaeozoic to the present was not possible in Turkey. Neither was it possible to define any other kind of tectonic unit that maintained its identity throughout the time period here considered—in a manner terranologists seem to think possible! For every tectonic unit we define below we thus specify the interval in which its definition remains valid.
Fig. 2

Suture and block distribution in and around Turkey. The blocks outside Turkey are not indicated in any detail, and the blocks in Turkey are only those of the Alpides, because before the Alpide evolution all of Turkey was a part of Gondwana-Land albeit with two marginal basins that opened and closed within it at different times as described in the text. Key to abbreviations: AM Alanya Massif, EAAC East Anatolian Accretionary Complex, MaM Malatya Digitation, MM Munzur Digitation, ND North Dobrudja. The references used specifically for this map are Alavi (1996), Ballato et al. (2011), Berberian (1983), Berberian and King (1981), Burchfiel (1980), Burchfiel and Nakov (2015), Erguvanlı (1959, 1961a, b), Guest et al. (2006), Ketin (1984), Koçyiğit (1991), Okay et al. (1996), Okay and Tüysüz (1999), Özgül (1976), Topuz et al. (2017), Sağdıç (2017), Seymen (1985), Sengör and Yilmaz (1981), Şengör (1984, 1990a, b, c), Yiğitbaş and Yılmaz (1996)

Fig. 3

Stereogram of a suture showing its geometry at depth. In our suture map (Fig. 2) we drew the lower surface shown in the stereogram, because the upper surface shows the rocks expelled from suture in the form of nappes that cover the continental fragments usually on the downgoing plate, but, much more rarely, also on the overlying plate and mask their true extent

Fig. 4

Two cross sections across western (A–A’ and B–B’) and central (C–C’) Turkey to illustrate the presumed deep geometry of the sutures and the highly simplified internal geometries of the continental blocks delimited by them. They do not show any of the extensional structures that affected them from Oligocene onwards to emphasise the palaeotectonic structures. Much less is known of the internal structure of the Kırşehir Massif than that of the Menderes and hence the structures shown are more schematic, but inspired by the sparse structural data available. The red line at the bottom of the western sections is the moho computed from 2600 receiver functions at 40 broadband seismic stations with an average spacing of 15 km. (Karabulut et al. 2013). For Central Turkey the moho depth we used is from Vanacore et al. (2013). The references used for the cross-sections are Akbayram et al. (2016b), Aksu et al. (2005), Erguvanlı (1959, 1961a, b), Göncüoğlu (1977), Gürer et al. (2016), Gutnic et al. (1979), Ketin (1956, 1963) Okay et al. (1989a, 2001, 2006), Polat and Casey (1995), Pollak (1958), Şengör (1995), Şengör and Özgül (2010), Seymen (1981a, b, 1985). In many places we also used our own unpublished observations. For the locations of the sections, see Fig. 5

Fig. 5

Distribution of isotopic ages of selected igneous and metamorphic rocks from the sutures in and around Turkey. The ages are given in the form of a colour scale so as not to make the map illegible. Many more ages from the continental pieces bounded by the sutures are given in the text in their geological contexts. The placement of the cross-sections shown in Fig. 4 are indicated with black lines

The sources for the ages are as follows: Alparslan and Dilek (2018), Babaie et al. (2006), Bonev et al. (2015a), Božović et al. (2013), Çelik and Chiaradia (2008), Çelik and Chiaradia (2008), Çelik et al. (2006,2011, 20132016), Çetinkaplan et al. (2016), Chan et al. (2007), Çörtük et al. (2016), Daşçı et al. (2015), Dilek and Moores (1990), Dilek and Thy (1992, 2006, 2009), Dilek et al. (1999), Dimo-Lahitte et al. (2001), Galoyan et al. (2009), Ghazi et al. (2003), Hässig et al. (2013), Karamata and Lovric (1978), Karaoğlan et al. (2012), Parlak et al. (2010), Khalatbari-Jafari et al. (2004), Koepke et al. (2002), Lanphere et al. (1975), Liati et al. (2004), Lytwyn and Casey (1995), Mukasa and Ludden (1987) Oberhänsli et al. (2012, 2014); Okay and Satir (2001), Okay et al. (1998, 2002, 2008), Okrusch et al. (1978), Önen (2003), Parlak and Delaloye (1996, 1999), Parlak et al. (1995, 2013), Robertson et al. (2013), Roddick et al. (1979), Rolland et al. (2010), Sarıfakıoğlu et al. (2017), Seidel et al. (1981), Sherlock et al. (1999), Smith (1993), Spray et al. (1984), Spray and Roddick (1980), Thuizat et al. (1978, 1981), Topuz et al. (2008, 2013, 2017, 2018), Uysal et al. (2015) and Yılmaz and Maxwell (1982)

There are various orders of tectonic units. Şengör (1990a, b) separated first-order tectonic units, defined as independently functioning plate tectonic elements such as (1) continents (continental margins of various kinds including magmatic arcs growing on a continental margin seem amenable to being treated as first-order tectonic units ever since the American geologist James Dwight Dana pointed out how much differently they behave from the rest of the continent, when he separated a stable interior from a less stable margin: Dana 1863, p. 198), (2) island arcs, oceanic plateaux (that are hardly ever preserved intact in the geological record, but leave behind tangible evidence of their past influence on other tectonic units; for an exception of a preserved one, see Şengör et al. 1996; Deng et al. 2017), (3) suture zones welding buoyant entities to one another upon collision (Fig. 3) or remnant oceans enclosed and preserved within a continent (such as the North Caspian Depression: Burke 1977; Şengör and Natal’in 1996), (4) entire orogens, taphrogens or keirogens. But these first-order units may contain or themselves be divided into second-order units (Fig. 6) formed by what Şengör had called the kopeogenic structures (Şengör 2003), i.e., structures of small wavelength: nappes (of both genres: Termier 1911; Tollmann 1973, has presented an exhaustive annotated list of their various kinds), strike-slip horsts (e.g., Woodcock and Fisher 1986), normal-fault bounded allochthons, regurgitation structures bounded by thrusts at the bottom and normal faults at the top (both can be oblique-throw structures), even very large slide structures often seen along Atlantic-type continental margins (Woodcock 1979), the largest examples of which give rise to sizeable non-tectonic foldbelts at the foot of Atlantic-type continental margins, such as the Perdido Fold Belt in northwestern Gulf of Mexico (e.g., Hudec et al. 2013). These second-order tectonic units chop up the first-order tectonic units and redistribute their pieces often making reconstructions cumbersome. Figure 6 shows the principles of this classification of tectonic units. Lom et al. (2018) elaborated upon the scheme presented by Şengör (1990a, b).
Fig. 6

Three diagrams showing the distinction between first- and second-order tectonic units. In a, entities I, II and III are independently functioning tectonic individuals such as island arcs, small continental platforms and major continental rafts. These are first-order tectonic units. b Shows the suturing of the first order units to one another. c Shows the creation of second-order tectonic units by redistributing by thrusting of the earlier collided first-order units (1, 2, 3). In the field, blocks separated by sutures and the later faults all would have different geological histories. Imagine the confusion that would arise if all these pieces are just called ‘terranes’. Second-order tectonic unit generation may happen in many more ways than shown here, but we hope that the principle is clearly expressed

Accordingly, we have distinguished first-order versus second-order tectonic units in our analysis, but it was not practical to show them all in our figures. Our figures show all the first-order tectonic units and some of the major second-order ones. All neotectonic deformations are ignored and thus Fig. 2 is non-palinspastic also with respect to neotectonic deformations. While reconstructing the palaeogeography for any given time in the past, the relevant tectonic units must be palinspastically restored to their proper positions. This is not always possible for two main reasons: (1) plate tectonics by its very nature destroys geological record (Dewey 1975, 1976, 2019), and (2) there are almost never sufficient amount and precision of data to make unique reconstructions of past events except in very trivial cases. Despite the great amount of work done in Turkey during the past decades, this second cause for the impossibility of making unique or even highly reliable reconstructions makes geologist’s job very difficult in most cases. However, this is no reason for despondency: as we point out below, we close the gaps in our knowledge by proposing testable hypotheses. The great explosion of the amount and quality of the work following the Şengör and Yılmaz synthesis was in part due to the testable hypotheses it contained (at this writing its Google citation record had already topped 3000) (Fig. 7).
Fig. 7

Map showing the first- and some of the major second-order tectonic units of Turkey and the locations of our stratigraphic columns (in Figs. 13, 15, 16 and 17). The rectangles in the Est Anatolian Accretionary Complex show the geological maps displayed in Figs. 10 and 11. The only locality of Topuz et al. (2017) that actually falls into the East Anatolian Accretionary Complex is in I. The other two are at on its eastern frame, in the continental fragment of northwest Iran just above the letter A (Marnelis et al. 2007)

Sutures in and around Turkey

The sutures in and around Turkey may be divided into two main classes: Palaeo- and Neo-Tethyan (Figs. 1 and 2). The elimination of the Palaeo-Tethys (largely by medial Jurassic in Turkey) gave rise to the Cimmeride Orogenic System. Thus, the sutures of the Palaeo-Tethys are the sutures of the Cimmerides. The removal of the Neo-Tethys (largely gone by the Eocene; only in the Eastern Mediterranean still open) created the Alpides and, therefore, the Neo-Tethyan sutures are the sutures of the Alpide Orogenic System. The Cimmerides and the Alpides constitute the Tethyside Superorogenic Complex (for the definitions of all these terms and concepts, see Şengör 1984, 1990a). Since the Alpides largely overprinted the Cimmerides, the Palaeo-Tethyan sutures are much more deformed and scattered into numerous second-order tectonic units created during the Alpide deformations than those of the Alpides themselves (see Fig. 2). No sutures earlier than the Tethysides have yet been identified unequivocally in Turkey with the possible exception of a Scythide suture in the Strandja Mountains. While describing the sutures below, we use the names of the Alpide continental blocks for ease of geographical reference (Fig. 2).

Cimmeride sutures in and around Turkey

Rhodope–Pontide Fragment

The Strandja Mountains [westernmost Pontides (Ketin 1966a): northwestern Turkey and southeastern Bulgaria]

Perhaps the most spectacular example of a Scythide/Cimmeride suture in Turkey has recently been mapped in some detail by Okay et al. (2001) and Natal’in et al. (2012, 2016). The uniqueness of this suture stems from the fact that a continuous development from at least the late Palaeozoic into the early Cretaceous can be observed with a clarity that leaves little to be desired. The oldest ‘basement’ seen in the Strandja Mountains in Turkey consists of Precambrian and Cambrian granitoids (consisting dominantly of granites with K-feldspar) unconformably covered by Ordovician to Permo-Carboniferous sedimentary rocks in places cut by late Carboniferous granitoids containing granites, granodiorites and tonalites. This complex is exposed in Bulgaria in the Sakar Mountains (Ivanov 1998; Gerdjikov 2005; Machev et al. 2015) and the Srednogorie/Balkan sectors (see Velichkova et al. 2004; Carrigan et al. 2005). What had been called ‘basement’ in earlier studies (Aydın 1974, 1982; Çağlayan and Yurtsever 1998) turned out to be intruded by Permian Kırklareli-type granites. The cover of this ‘basement’ consists of Triassic through Jurassic metasedimentary rocks and rare dykes of latest Triassic age (213 Ma: Natal’in et al. 2016), corresponding to the ‘Sakar-type Triassic’ in Bulgaria (Chatalov 1985).

To the west-northwest, in the Kırklareli region, there are no known Precambrian rocks, the oldest units being possibly Silurian to late Carboniferous on the basis of both intrusive rock ages and detrital zircons (Sunal and Satir 2008) from garnetiferous schists consisting mainly of pelitic protoliths. All of them are cut by late Carboniferous granitoids (311 to 315 Ma: Natal’in et al. 2012; Sunal et al. 2006; for other equivalents in Bulgaria, see Chatalov 1982). In the westernmost outcrops of the garnetiferous schists intercalated with amphibolites are Carboniferous transported zircons giving a possible Permo–Carboniferous age to these rocks cut by Permian calc-alkalic Kırklareli granites and their equivalents made up of subvolcanic granites, granite and quartz porphyries just across the border in Bulgaria (Chatalov 1992) giving ages ranging from 300 to 250 Ma (Sunal et al. 2006; Sunal and Satir 2008; Natal’in et al. 2012). All of these units were thrust northward along mylonite zones and came to rest on the units originally covering them, beginning with metaconglomerates rapidly passing upwards into metasandstones and dark shales, equivalents of the Palaeokastro Formation in Bulgaria (Chatalov 1992). Their ages are Lower to Middle Triassic (Chatalov 1992; Hagdorn and Göncüoğlu 2007; Bedi et al. 2013). A Jurassic sequence following them sits on them across a mylonitic zone. Natal’in et al. (2016) interpreted this boundary as a thrust, but we regard it as a low-angle normal fault, because it in places elides in the entire Triassic section. This normal fault may have been coeval with those that opened the Luda Kamčija Trough just to the north during the Pliensbachian–Toarcian interval in which deposition commenced with the Sinivir Formation (Tchoumatchenco et al. 1992), although Elmas et al. (2011) reported a considerable late Jurassic-early Cretaceous (40Ar/39Ar ages between 156 to 143 Ma) extension leading to the formation of an extensional metamorphic core complex in the northeastern Strandja. These two extensional events may be one protracted episode or two successive episodes of extension. Future observations on the ages along the extentional faults would solve this problem.

In the southeast, in the region of Kıyıköy, the Yavuzdere Complex, seen best in the two streams of Yavuz Dere and Arka Dere, consists of bt-ms schists, marbles, metatuffs, metamafic rocks believed initially to have been pyroxenites, and locally metacherts all intruded by a dense swarm of plutons ranging from diorites to granites of latest Carboniferous ages (311–302 Ma: Natal’in et al. 2016), which correlate with the ages of the protoliths of the metatuffs (idem). In some areas, the widespread conglomerates and quartzites were probably originally turbidites. This is similar to the probably Devonian or Triassic phyllitoid schists and diabases northwest of Gramatikovo (Kalvacheva and Čatalov 1974) and the ‘Diabase-Phyllitoid’ Formation, a part of the Sokol Formation, in Bulgaria (Chatalov 1985), although they are clearly younger than the Sokol Formation. Slightly farther to the northwest, in the Bay of Selvez, there is a lightly metamorphic wildflysch cut by late Cretaceous magmatic rocks belonging to the Alpide evolution (Natal’in et al. 2016). The age of this wildflysch is unknown, but may be anywhere between the Carboniferous (Yavuzdere equivalent?) and the early Jurassic (Lipacka equivalent?).

The Palaeo-Tethyan accretionary complex (Mahya Schist and Şermat Quartzite; Şengör and Özgül 2010) is thrust onto a basement containing Precambrian granitoids, but sparsity of outcrop does not allow the establishment of the structure of the country rock consisting mainly of phyllites (Natal’in et al. 2016). This basement extends as far to the southeast as the Çatalca region (Şahin et al. 2014) where there are no age data from the country rocks. It is unconformably covered by an Ordovician to possibly late Palaeozoic quartzite, thick marbles and marble and chert intercalations greatly reminding one of the stratigraphy of the Palaeozoic of İstanbul that spans an interval from the Ordovician to the Namurian (Şengör and Özgül 2010; Lom et al. 2016).

Figure 8 shows our interpretation of the tectonic evolution of the Strandja Mountains and its surrounding units. Sometime during the Callovian the continental rocks began overthrusting the Yavuzdere Complex with a northerly transport direction (170–160 Ma muscovite ages: Lilov et al. 2004; Sunal et al. 2011), although the presence of earlier, similar thrusting may have taken place as suggested by the onset of the Lipacka Flysch sedimentation during the medial to late Triassic (Chatalov 1979, 1997; Dabovski and Savov 1988). Thrusting continued into the Tithonian (150 Ma: Sunal et al. 2011), bulldozing also the Luda Kamčija Trough during the Bajocian–Bathonian, turning it into a foreland molasse basin that further migrated north into the Niš-Troján Trough depositing a Berriasian–Valanginian flysch here (Tchoumatchenco et al. 1989, 1992). In the Romanian part of the Moesian Platform, within the Cirligati Anticline, the Upper Triassic rocks are folded and thrust with a northerly vergence and the thrust contacts are sealed by the Jurassic, leading Tari et al. (1997, p. 1156) to state that ‘the Moesian Platform cannot be considered as an undeformed platform sequence, at least not its pre-Jurassic succession’. It is clear that the main collision here was sometime between the latest Triassic and some unspecified Jurassic although post-collisional convergence seems to have continued in the south until the early Cretaceous reaching as far north as the Niš-Trojan Trough (Tchoumatchenco et al. 1989). The isotopic ages from the Turkish Strandja between 154 and 118 Ma (Elmas et al. 2011; Sunal et al. 2011) may be related both to the waning Cimmeride events and the onset of Alpide deformations, but the Cretaceous deformations overlap with the Alpide events and it is impossible to tell when precisely the Cimmeride events here ended and the Alpide began on individual outcrops. Only the evidence of the Cirligati Anticline in Romania and the isotopic ages of the muscovites in the Strandja Mountains show that the main Cimmeride deformations probably ended already sometime during the Jurassic, most likely during the Bajocian–Bathonian interval judging from the continuation of the Strandja zone to the east described below. The statement by Cattò et al. (2018) that there is no significant Alpide overprint over the Cimmeride deformations cannot be taken seriously in view of the major Cretaceous arc magmatism that affected it and the thrusting of the entire massif over what seems to be an ophiolitic substratum in the south (see below).
Fig. 8

Two schematic and sequential cross-sections across the Strandja Mountains showing the outlines of its structural evolution as discussed in the text. To keep the figure simple, the Jurassic and early Cretaceous extensional episode in northeastern Strandja is not shown

Central Pontides

From the Strandja, the Cimmeride suture continues into the Central Pontides. North of the İstanbul Zone, there are no Cimmeride outcrops, although here the entire Jurassic and the early Cretaceous are missing, the Carnian clastic rocks being unconformably covered by the Middle to Late Cretaceous through Eocene rocks. The gap corresponds nicely to the duration of the Cimmeride collisions and the subsequent convergence. Just east of İstanbul, within the province of Kocaeli, medial Cretaceous transgression begins with the famous Hereke puddingstones (Erguvanlı 1949) of Albian age and rapidly passes into lagoonal carbonate rocks which, in the north, interfinger with latest Cretaceous andesites and basaltic andesites.

Farther east, in the core of the Çangaldağ Anticline, Şengör et al. (1980) described an ophiolite complex in the Küre region, a place of copper ore exploitation since antiquity. The Küre ophiolite and its epiophiolitic oceanic sequences plus the unconformable cover rocks had been mapped first by İhsan Ketin in 1960. They consist of cherts, black shales, phyllites, graphitic schists, and greywackes (Şengör et al. 1980, 1984). Below them are fully developed ophiolites consisting from base to top of serpentinised harzburgite separated above with a thrust contact from cumulate and isotropic gabbros, then sheeted dykes topped by pillowed and massive lavas and lava breccias covered by oceanic sediments (Ustaömer and Robertson 1993, 1995). Ustaömer and Robertson (1994, 1999) looked at the detailed geochemistry of the extrusive members of the ophiolites and wrote the following:

‘… we carried out high-precision re-analysis of 20 samples for Nb and (light rare earth) (LRE) elements (i.e., La, Ce, Nd; [see their Table 2]). The discussion below is based on the new results. The ophiolitic extrusives of the Küre lavas lie in the basalt and basaltic andesite fields in Zr/Ti versus Nb/Y plot (Winchester and Floyd 1977; [see Table 3 of U and R]). On the Ti/Y versus Nb/Y diagram (Pearce 1982), they plot in the tholeiitic field. In addition, these extrusives plot in the area of overlap of island arc tholeiite (IAT) and MORB fields, or scattered between the IAT and MORB fields in tectonic discrimination diagrams [see their Table 3].

The Küre lavas mainly plot in the back-arc basin basalt (BABB) field on a Ti versus V diagram (Shervais 1982). The La/Nb ratio of BABB is intermediate between MORB and true IAT (Floyd pers. comm., 1992). In the Y versus La/Nb diagram (fig. 4a; Floyd et al. 1991), most of the re-analysed lavas plot in the BABB, with only two in the IAT field. None of the samples plot in the MORB field in this diagram. Further information can be obtained from MORB-normalized multi-element variation diagrams. In most of these there is marked LIL element enrichment, also Nb depletion, Ce and P enrichment and a negative slope between Zr and Ti (fig. 4B ± D). These are common characteristics of supra-subduction zone eruptive tectonic settings. The geochemical evidence thus implies that the basic extrusives of the dismembered Küre ophiolite erupted in an above-subduction zone setting. Geochemical evidence of immobile major and trace elements and, in addition, pyroxene chemistry, indicate mid-ocean ridge (MOR) and volcanic arc basalt (VAB) compositions ([see their Table 3] Ustaömer and Robertson 1994), suggesting that the Küre ophiolite was generated above a subduction zone. Chrome spinel analysis also supports this interpretation (Ustaömer and Robertson 1994)’ (Ustaömer and Robertson 1999, p. 33).

From these data they correctly infer that the Küre ophiolite was generated above a subduction zone, but then, inexplicably, they ascribe it to a back arc setting of a subduction zone dipping north in the south of the Pontides, repeating an interpretation they had put forward already in 1993 only because the Devrekâni metamorphic rocks, belonging to the İstanbul Zone, are considered to thrust them from the south (Timur Ustaömer, pers. comm., 24th 2018; see Şengör et al. 1980; Ustaömer and Robertson 1993, 1995). However, the entire Küre complex is deformed by north-vergent thrusts and folds (Şengör et al. 1980, 1984; Ustaömer and Robertson 1993, 1995) and, at the time of its formation, there was no ocean to the south of it to supply the allegedly north-dipping slab (Şengör and Yılmaz 1981; Görür et al. 1983). We, therefore, think it is more reasonable to consider it having formed atop a south-dipping extensional subduction zone in a pre-arc-spreading environment. This interpretation is supported by the widespread extension along the northern margin of Gondwana-Land from the Permian to the end of the Triassic (see Şengör and Atayman 2009) in a continental margin magmatic arc setting with intrusive ages of granitois spreading in time from 237 to 221 Ma and in a geographic distribution from the northern Menderes Massif to Bolu, thus disregarding the entire Neo-Tethyan suture complex found in northwestern Turkey and showing that those sutures represent oceans which opened after the activity of the Triassic arc (Erdin Bozkurt, written communication, 12th March 2018).

The Küre Cu-sulfide ores include pyrite, chalcopyrite, bonite, covellite, sphaleratie, digenite, marcasite, tennanite and carrolite appearing at the contact between abyssal tholeiite basalts forming ophirags and black shales plus some greywackes. This rock assemblage belongs to the Upper Palaeozoic to Liassic Akgöl Mélange that also includes Triassic limestone knockers (Şengör et al. 1980, 1984; Aydın et al. 1995; Kozur et al. 2000; Okay et al. 2015). The pyritic copper deposits are seen as stockwork-disseminated ores at the upper levels of the basalts and the black shales. The mineralisation in Küre is both massive and disseminated with most of the occurrences being associated with faults near the basalt/black shale contact, although much of the mineralisation predated the faulting associated with the subduction–accretion complex formation as seen also in the classical occurrences associated with the Troodos ophiolite in Cyprus (Sengör et al. 2008a, b).

We interpret the Küre suture as a north-vergent Palaeo-Tethyan suture that followed the south-dipping Cimmeride subduction below the Devrekâni metamorphic rocks representing parts of the İstanbul Zone, an easterly continuation of the Moesian Block, i.e., a part of Gondwana-Land (Yanev et al. 2005, and the references therein). All the rocks brought into contact by the suture are unconformably covered by the conglomerates and the sandstones of the Bürnük Formation, in which İhsan Ketin had found a Dogger ammonite (Şengör et al. 1980). The zircon ages of 165–172 Ma from Küre and Çangaldağ Region (Okay et al. 2014) are thus entirely compatible with the earlier, palaeontologically determined Dogger age for the Bürnük Formation.

Immediately south of the Küre area, is the Elekdağ lherzolitic massif overlying the Domuzdağ garnet-blueschist micaschists. Originally, Şengör et al. (1984) and following them Tüysüz (1990) had interpreted these rocks also as parts of the Palaeo-Tethyan suture, but belonging to the Karakaya marginal basin (Şengör and Yılmaz 1981; Şengör et al. 1984). This was contrary to Ketin’s interpretation (Ketin 1966a), who had considered these rocks as parts of the northern Neo-Tethyan suture zone. However, Şengün (1995) argued, on the basis of stratigraphic and structural observations, that Ketin had been basically right in his assessment and Okay et al. (2006) showed that these rocks were metamorphosed at 490 °C and 17 kbar some 105 Ma ago (Albian) and were exhumed 20 Ma later (Turonian-Coniacian), thus further corroborating Ketin’s (1966a) original ascription of them to the northern Neo-Tethyan suture. We, therefore, correct the earlier interpretation by Şengör et al. (1984) and Tüysüz (1990) and discuss them below under Alpide sutures, although we dispute the interpretation of a continuous subduction from the Triassic to the Cainozoic as maintained by Şengün (1995) and Okay et al. (2006).

Eastern Pontides

In the Eastern Pontides the Palaeo-Tethyan suture is exposed in isolated, small erosional windows, which we describe below.

Ağvanis Massif. Şengör et al. (1980) had described the Ağvanis Massif mainly on the basis of Nebert (1961). Later Aral Okay and his co-workers remapped the massif and established a more detailed stratigraphy and structure (Okay 1984a; Topuz et al. 2011). The massif has an asymmetric spindle-shape delimited in the Kelkit Valley by the two branches of the North Anatolian Transform Fault forming a strike-slip duplex (cf. Şengör et al. 2005) and has an anticlinal structure the axis of which roughly parallels the two enveloping branches of the North Anatolian Fault. Okay (1984a) pointed out that the massif exposes three sequences. At the base are black metamafic rocks with local pillow structure and grey to pink metadacites in which the primary structures are mostly obliterated by intense tectonism. The present banding was formed by the deformation of the pillow structures. The black colour is imparted by an abundance of magnetite. Overlying them is a green to bluish green, grey, locally graphitic phyllite with thick metamafic and marble lenses. The top sequence is about 2 km thick and consists of dark green to bluish-green metamafic rocks and metatuffs with interbedded metadacite, marble, black calcschists and again graphitic schists. The metadacites cut the metamorphic rocks but then were folded themselves. In this sequence there is also a trondjhemite. Okay encountered a metaserpentinite in only one locality in the uppermost sequence which is formed from chrysotile and opaque minerals.

Okay (1984a) thought that the pre-Liassic age assigned by Nebert and accepted by Şengör et al. (1980) was wrong and considered the entire Ağvanis metamorphics as parts of a continuous Mesozoic arc sequence reaching into the Eocene. His assessment turned out to be incorrect, because isotopic dating of the metamorphic rocks revealed a Norian age (209 ∓ 1 Ma; Topuz et al. 2011), corroborating Nebert’s original age assignment, thus being similar to the situation not only in the Küre area, but also farther east in the Pulur Massif, Gümüşhane and the Yusufeli areas (see below). Thus, the Ağvanis metamorphic rocks have nothing to do with the much younger Pontide arc related to the subductive removal of the northern branch of the Neo-Tethys. The Ağvanis appears also to be an immature ensimatic arc formed above a Palaeo-Tethyan subduction zone.

Pulur Massif. In the Pulur Massif, within the province of Bayburt, between the Ağvanis and the Yusufeli region farther to the east, the following rocks belonging to the Carboniferous to Permian ages are exposed in two thrust sheets: the Doğankavak unit is exposed in the eponymous thrust sheet consisting of low grade, strongly deformed pile of metabasalts and phyllites with subordinate calcareous phyllites, marbles, quartzo-feldspathic schists and metacherts. This sequence exhibits a pronounced, generally northeast–southwest-striking and southeast-dipping foliation characterising north- to northwest-vergent abundant isoclinal folds. Most primary contacts have been erased by tectonism and metamorphism (Topuz et al. 2004b).

The lower unit is overlain by the Cenci Nappe thrust north to northwestward with respect to the Doğankavak Nappe. The Cenci Nappe consists of three high-grade units, namely Petekli, Pirörenler and Serenli (Topuz et al. 2004a). The lowermost Petekli subunit consists of mesocratic gneisses and aluminous melanocratic rocks with metamafic intercalations containing subordinate trondjhemite dykes (Topuz et al. 2004a, b). Above it is the Pirörenler metapelitic to metapsammitic migmatites with local amphibolite intercalations, the whole displaying a low-grade granulite metamorphism and dehydration melting. That episode is followed by nearly isothermal decompression followed by penetrative, in part mylonitic deformation in upper amphibolite-grade metamorphic conditions. Single-grain U–Pb monazite dating from the metapelitic migmatites gave 327–332 Ma ages. Cooling, as indicated by biotites, took place around 310 Ma. Above the Pirörenler subunit is the Serenli mylonitic belt formed from quartzo-feldspathic gneisses with amphibolite intercalations. Mylonitisation happened under lower amphibolite to upper greenschest conditions. Gültekin Topuz (personal communication, 26th May, 2018) thinks that the contacts between these units are all tectonic, but metamorphism has obscured the structural indications and metamorphism across them is without a break. All the contacts, including that between the Doğankavak and the Cenci are sealed by Liassic clastic rocks, equivalents possibly of the lower levels of the Bürnük Formation (Topuz et al. 2004a). The Pulur Massif as a whole makes up a part of the Ahıska Nappe lying above the Hamurkesen Nappe thrust over the Cebre relative autochthon. The thrust contacts are cut by Eocene plutons and are therefore of late Cretaceous to Palaeocene age (Topuz et al. 2004b).

In the Pulur Massif, we see the remnants of a partly deeply buried accretionary complex that surfaced before the Lias. The age of the accretionary complex itself is unknown, but cannot be much older than late Palaeozoic, as the shallow water to partly terrestrial sequences onto which it was thrust is of Permo-Carboniferous age (Ketin 1951; Okay and Leven 1996; Okay et al. 1997). We correlate this north-vergent accretionary wedge with the oceanic sequences in the Küre region and in the Ağvanis Massif and identify it as a part of the main Palaeo-Tethyan suture zone. The Petekli unit is likely a suprasubduction zone ophirag caught up within the Pulur accretionary wedge. The overlying units were more deeply buried and were subsequetly thrust on top of it.

Yusufeli Region. Şengör et al. (1980) described near Yusufeli, close to the Turkish/Georgian border region in northeastern Turkey, serpentinised ultramafics, gabbros, amphibolite gneisses and mafic volcanic rocks interlayered with slates within the deep gorge of the Çoruh River. At the junction of the Erzurum and Oltu highways, near the village of Kınalı, the uppermost part of the ultramafic and mafic rocks displays mafic pillow lavas of tholeiitic affinities. In stratigraphic contact with the pillow lavas are ribbon cherts, shales and fine sands containing Liassic fossils.

This region has been remapped in greater detail by Ustaömer and Robertson (2010), who greatly improved the stratigraphy. They separated a Demirkent intrusive complex consisting of amphibolites, mafic and tonalitic dykes. The amphibolites are the oldest rocks and exhibit complex deformation of three main phases and an isotopic age of 335 Ma (Yener Eyüboğlu, written communication, 26th May 2018) being compatible with a Carboniferous age from a tonalite intruding them (Ustaömer and Robertson 2010). The first phase produced banded amphibolites out of mafic protoliths. The second phase was associated with the intrusion of some rare tonalite dykes, gabbros and trondjhemites (see Ustaömer and Robertson 2010, fig. 9). Eyuboglu et al. (2016) studied the larger intrusions outside the Demirkent Complex: the largest of them is an isotropic olivine gabbro with only a small part in the southeast of the outcrop area with cumulate texture. The gabbro body is dated by Eyüboglu et al. (1996) at 185–181 Ma (Pliensbachian-Toarcian) and it is cut by tonalitic and granitic dykes with ages ranging from 182 to 177 Ma (all within the Toarcian). We plotted the data provided by Ustaömer and Robertson (2010), Table 2, supplementary material) from their mafic rocks in a La/Nb versus Y diagram and saw that of their nine datapoints five plot into the back arc basin basalt field, one into the MORB field, three into ocean floor basalts and one into trace element-enriched + transitional field. A Carboniferous metamorphic complex, termed the Karadağ Metamorphics, consists of banded gneisses and quartz micaschists thrust over the Demirkent complex (Ustaömer and Robertson 2010), but intruded by the younger intrusions sealing the contacts by the later Lias (Eyüboglu et al. 1996).

Ustaömer and Robertson (2010) showed that the Liassic pillow lavas and deep-water sediments were not associated with the Demirkent complex, but in fact belong to a higher slice, we think fed by the dyke system described above and indicating a Liassic extensional regime here creating a deep basin. These relationships are very similar to those in the Central Pontides: The Demirkent Complex is clearly a supra-subduction zone ophiolite with uncertain vergence (because the dips are almost all vertical). So little of it is exposed that it is hard to know its exact tectonic setting; in analogy with Küre we ascribe it to a pre-arc spreading setting. The Liassic extensional environment was probably the last gasp of a south-dipping extensional Palaeo-Tethyan subduction as also interpreted by Eyüboglu et al. (1996); the statement by Ustaömer and Robertson (2010) that ‘there is no evidence of “Palaeotethyan” ophiolites in the eastern Pontide region’ (their p. 281) is thus contradicted partly by their own observations.

The geology described by Ustaömer and Robertson (2010) is extremely similar also to the Chorchana-Utslevi zone in the Dzirula Massif in the Caucasus complete with the old basement intruded by Carboniferous granites and the unconformable Lias (Sengör 1990c), with the exception that Ustaömer and Robertson (2010) report no strike-slip displacement along the Demirkent Complex, although the extremely steep dips they show are suggestive. Limited outcrop does not allow a further characterisation of the structure and evolution of the Cimmeride suture here.

Synopsis of the Cimmeride suture within the RhodopePontide Fragment. From the Strandja in the west to Yusufeli in the east an oceanic assemblage is thrust from the south by continental nappes in the entire Rhodope–Pontide Fragment within the confines of Turkey exactly as it had originally been described by Şengör et al. (1980, 1984) and Şengör and Yılmaz (1981). Şengör and his co-workers had assumed that this oceanic assemblage consisted of ordinary ocean floor remnants (ophirags) on the basis of the sparse data then available to them. Subsequent remapping of all the localities mentioned by Şengör et al. (1980, 1984) and much petrologic, geochemical and unfortunately inadequate structural work revealed that all the oceanic rocks exposed in the oceanic assemblage, which Şengör et al. (1980) had collected under the designation ‘Küre Nappe’ turned out to be either immature ensimatic arcs or even pre-arc-spreading ophirags metamorphosed to various degrees between greenschist and granulite facies and intensely deformed invariably by north-vergent structures. In the Strandja, the İstanbul Zone overthrusts the entire oceanic assemblage in the form of a Scythide traîneau écraseur2; Şengör et al. (1980) had called this structure the İstanbul Nappe, a name we retain, but the İstanbul Nappe itself has now been shown to consist of smaller secondary units in the form of smaller nappes and strike-slip bounded sheets that formed during the collisional processes between the late Triassic and the early Cretaceous, although some of the post-medial Jurassic shortening may be ascribed to the developing Alpide shortening events.

In the eastern Pontides, the Küre Nappe is tectonically overlain by another continental unit in the Yusufeli Region (see Ustaömer and Robertson 2010, fig. 7), which Şengör et al. (1980) called the Bayburt Nappe. That designation, too, we here retain. Owing to advanced erosion, the relationship between the İstanbul and the Bayburt nappes in Central Pontides cannot be observed.

The overthrusting of the Küre Nappe by the İstanbul and the Bayburt nappes took place sometime between the Lias and the medial and late Dogger. In the Caucasus the unconformable sequences already begin with the Lias (Şengör 1990c), whereas in the Strandja and the Moesian Platform convergence continued well into the Cretaceous, but there, as mentioned above, it is difficult to distinguish between Cimmeride and Alpide convergent events. Major Cimmeride convergence, however, seems to have ceased by the medial Jurassic also in the Strandja and the Moesian Platform.

The fact that a large continental traîneau êcraseur exists in the entire Turkish Cimmerides within the Rhodope–Pontide Fragment shows that the maximum width of the ocean consumed could not have been larger than a couple of thousand kilometres, if that many, as also demanded by the Pangaean reconstructions, especially the Pangaea A2 reconstructions. As we shall argue in the companion paper mentioned above, there was also much strike-slip faulting parallel with the Cimmerides that greatly complicate the pre-late Jurassic tectonics of the Rhodope–Pontide Fragment.

The above descriptions show that the numerous papers published after the Şengör and Yılmaz (1981) synthesis purporting to prove the non-existence of a Cimmerian Continent and a separate Palaeo-Tethyan ocean from a Neo-Tethyan one in Turkey (e.g., Topuz et al. 2013b) are bereft of a sound observational basis. Where the Cimmerian continent pieces are indeed missing, it is seen that later strike-slip faulting had elided them, bringing the Küre Nappe into contact with the Neo-Tethyan accretionary complexes. Such errors have so far stemmed from inadequate mapping, insufficient attention to stratigraphy and age dating and misinterpretation of the tectonic setting. The misinterpretation of the tectonic setting has been commonly a result of too myopic a view of the areas considered without regard to their surroundings and inadequate consideration of and inappropriate comparison with the present-day tectonic environments. Hasty and poorly informed tectonic interpretations on the basis of a few isotopic ages have been unfortunately a rather common occurrence in Turkey (as in many other developing countries) since the sophisticated age-dating methods became widely available (see Şengör 2014).

Sakarya continent

The Karakaya suture Şengör and Yılmaz (1981) defined the Sakarya Continent as a continental piece delimited by the Intra-Pontide suture in the north and the İzmir–Ankara suture in the south. Within this continental fragment the Karakaya Suture extends in an east–west direction from the Biga Peninsula to Ankara and cuts across it obliquely from a more northerly position towards a more southerly one as a wide continental terrain intervenes between it and the İzmir–Ankara Suture in the west, whereas near Ankara the products of the Karakaya Suture and the İzmir–Ankara Suture are practically intermingled. As we shall see below, this is in part because the western segment of the İzmir–Ankara–Erzincan Neo-Tethyan Suture opened synchronously with the Karakaya marginal basin, possibly as parts of a complex marginal basin system. Although the vergence of subduction related structures has so far been difficult to establish due to intense reworking, limited evidence suggests that it was mainly northward. Since the description by Şengör and Yılmaz (1981) numerous studies have been devoted to the Karakaya (e.g., Tekeli 1981; Yılmaz 1981; Gautier 1984; Koçyiğit 1991; Okay et al. 1991, 1996; Genç and Yılmaz 1995; Okay 2000; the papers in Okay and Göncüoğlu 2004), but none necessitated a fundamental change in their conclusions. Especially Yılmaz’s (1981) important observations that (1) the Karakaya represents an oceanic basin that formed by rifting a pre-existing Permian carbonate platform that developed atop the Sakarya continental basement in which there are remnants of the Pan-African basement; this oceanic basin generated eclogites dated at 208 and 203 Ma by Okay and Monié (1997), (2) that the Karakaya oceanic rocks are sandwiched between “two sialic (granitic) slices” (Yılmaz 1981, p. 49) and (3) the pre-Liassic deformation was north-vergent corroborate the marginal basin character of this basin, also endorsed by later geochemical studies on its igneous products. Similar observations were later made also in the basement of the Malatya Metamorphics in the Engizek Mountains that we mention below. Genç and Yılmaz (1995) reiterate these critical observations and give further details.

The Karakaya rocks continue into the Tokat Massif, a part of the Sakarya Continent (Şengör et al. 1984), across the interruption by the Çankırı Basin (Şengör et al. 1984) where they are represented by two main assemblages studied by Yılmaz et al. (Yilmaz et al. 1997a, b) in some detail (a more recent attempt at a synthesis by Yılmaz and Yılmaz 2004, confuses the Palaeo-Tethyan and the Neo-Tethyan rocks and is of very limited use in understanding the geology of the Massif): the Karakaya Complex and a continental unit called the Amasya Metamorphic Association. The latter is exposed to the north of the Tokat Massif. It consists of a slightly metamorphosed sequence going from bottom to top from conglomerates and sandstones, dark shales containing thin lenses of dark limestone lenses containing Silurian fossils. The Amasya Metamorphic Assemblage is thrust with apparent south vergence onto the Karaya Complex that here consists of the Turhal metaophiolites and the underlying Yeşilırmak metamorphic mélange complex made up of metapelites and metalava, marble and mafic lava blocks, and pelagic limestone blocks. All of these units are unconformably covered by the Liassic Bayırköy sandstones. However, within the Yeşilırmak Complex there are thin and long, east–west striking screens of serpentinites and radiolarian cherts that give Tithonian ages (Bozkurt et al. 1997). These units were considered to be emplaced along north-vergent thrust faults during the closure of the Neo-Tethys here by Bozkurt et al. (1997), but the map pattern of these slivers (see Hakyemez and Papak 2002) suggest emplacement along right-lateral strike-slip faults of possibly a late Cretaceous to Eocene age. Thus, the present-day outcrop pattern of the Tokat Massif most likely does not reflect the original disposition of the units within it. Even the southerly vergence seems a reworking by the Alpide sutures bounding the Tokat Massif to the north and south.

Menderes–Taurus Block

After the Karakaya suture leaves the Sakarya Continent we encounter its pieces in the Tokat Massif and in the basement of the Malatya Digitation of the Menderes–Taurus Block. Yılmaz et al. (1987a) described in detail the cover rocks above the Malatya Metamorphics in the Engizek Mountains. The lowermost unit above the metamorphic basement includes dolomitic marbles and semimarbles plus recrystallised limestones of Upper Permian age based on its fossil content. These are conformably covered by pinkish-reddish, thinly-bedded calcareous shale alternating with light grey medium-bedded limestones. Upwards greenish-buff coloured shaly limestones and reddish-pinkish calcareous shales terminate what is called the Yılanova Formation that has an Upper Scythian–Lower Anisian age. In the following Karabayır Formation we see the onset of magmatism and the unit is characterised, from bottom to top, by reddish shales that have quartzite layers in its upper horizons. Here the colour turns blackish green and the following layers consist of thinly-bedded, violet-coloured, nodular shaly limestones and limestones. Above these are laterally discontinuous, coarse-grained sandstones and conglomerates. These rocks pass laterally and vertically into greenish black siltstones and shales. Influence of vulcanicity is seen in the manganese-coloured silexites and silicified claystones.

In the northern regions of the area the sequence just described contains Permian and Triassic limestone knockers and interlayered altered mafic lavas and spilites. The fossils from the matrix allowed an age assignment of Middle to Upper Triassic to the Karabayır Formation. A detailed study of the clastics of the Karabayır showed that it was fed by the Koçdağ and the Yılanova rocks, indicating that they had surfaced and were being eroded. Yılmaz et al. (1987a) interpreted the environment in which the Karabayır Formation originated as one of extension as indicated by the abundant debris flows into an unstable depositional milieu characterised by abundant mafic vulcanicity.

The following lithostratigraphic unit has even more volcanic rocks in it consisting mainly of subvolcanic diabases that are extensively altered, in a few places cut by leucocratic dykes. These rocks are within thinly-bedded shales and manganese-bearing radiolarian cherts. These rocks are intermingled with what has been called the Okkayası Formation, which is simply a coloured mélange. Its contents include Middle to Upper Triassic fossils.

The entire section described above was highly deformed by thrust faults and folds, with a northerly transport direction and both the rocks and the faults are unconformably covered by the limestones of the Engizek Formation of Middle Jurassic age. This formation begins with a basal conglomerate and rapidly evolves into shallow water deposits laid down in reefal and tidal environments.

The Malatya Digitation is the farthest east we can follow the Karakaya Suture.

Synopsis of the Karakaya Suture From the Biga Peninsula in the west to the Engizek Mountains in the east, the Karakaya suture represents an oceanic realm that opened in the latest Permian to early Triasic interval and was closed by subduction-controlled shortening and final collision before the Jurassic. In some places Lias, in others Dogger cover its deformed remnants. Blueschist debris was found in these cover rocks (Gautier 1984). The vergence has so far been difficult to establish owing to intense later reworking, but some tantalising evidence suggests in was mainly northward. In Bilecik and in the Engizek Mountains, observations clearly indicate a latest Permian rifting. Thus, it cannot be a part of the main Palaeo–Tethyan suture as has been repeatedly claimed by some; quite to the contrary, in western Turkey it may have communicated with the coevally opening northern branch of the Neo-Tethys. The original interpretation by Şengör and Yılmaz (1981) that it was a back-arc basin to the Palaeo-Tethys still seems the most appropriate interpretation of its nature.

Alpide sutures in and around Turkey

Alpide sutures are both more abundant and better exposed than the Cimmeride sutures in Turkey. The following are brief descriptions of them.

Sutures of the northern branch of the Neo-Tethys

Alpide sutures lying north of the Menderes–Taurus Block and the Kırşehir Massif were defined by Şengör and Yılmaz (1981) and Şengör et al. (1982) as the sutures of the northern branch of the Neo-Tethys within the Mediterranean Alpides (see Şengör 1984). This is the branch that continues westward all the way into the western Mediterranean. We describe them from north to south in what follows:

Intra-Pontide suture

This suture, named by Şengör and Yılmaz (1981), extends between the Gallipoli Peninsula in the European Turkey to north of Ankara and takes its name from its position within the Pontide unit as defined by Ketin (1966a), who had not yet separated the Sakarya Continent from the rest of the Rhodope–Pontide Fragment.

The westernmost outcrops of the Intra-Pontide suture are seen in the Thrace Basin and the low hills delimiting it to the south. The entire Thrace Basin was a forearc basin between the late Cretaceous, when subduction began in that segment of the northern branch of the Neo-Tethys now represented by the Intra-Pontide Ocean, and the Bartonian when unconformable sediments began covering the suture. Görür and Okay (1996) thought the age of the forearc basin was Palaeocene to Oligocene, which is unlikely, because the arc to the north had already stopped its activity during the latest Cretaceous to earliest Palaeocene and the collision was complete by the later medial-Eocene (Akbayram et al. 2016a).

Şengör and Özgül (2010) assumed that the entire Thrace Basin is underlain by a subduction–accretion complex of mainly late Cretaceous age. They based their claim on three observations: the first is that both at its northernmost and at its southernmost edges, the basin is underlain by ophiolite and ophiolitic mélange. In the north, Özaydın and Erol (1981) argued for the presence of an ophiolite slab dipping north below the Strandja Massif which was imbricated and emplaced during the late Cretaceous on the basis of gravity, magnetic and seismic profiling observations. Figure 9 is a simplified rendering of their detailed seismic profile DG-1256 that clearly shows the southerly imbrication of the ophiolite and the ophiolitic mélange. The low hills delimiting the basin to the south are made up of Eocene clastic rocks beginning with delta deposits passing upward into deeper basin facies and finally shallowing again. The age of this entire package spans an interval from the Cuisian (middle Lower Eocene) to the Oligocene (Sümengen and Terlemez 1991). They represent the southernmost edge of the same accretionary complex and the fore-arc/intramontane basin fill. Moreover, the basin is located between an active magmatic arc and the ocean basin to the south in the late Cretaceous (isotopic ages ranging from 83.5 and 70.9 Ma: Aydın 1982; Ohta et al. 1988). All these observations and the inferences made from them led Şengör and Özgül (2010) to deduce an ophiolitic mélange wedge under the Cainozoic sediments of the Thrace Basin.
Fig. 9

A simplified rendering of the detailed seismic profile DG-1256 of Özaydın and Erol (1981) that clearly shows the southerly imbrication of an ophiolite and the ophiolitic mélange along the southern rim of the Strandja Mountains in Thrace

The deeper parts of the Thrace Basin do not crop out at the surface. The oldest visible sequences are of Eocene age (Siyako and Huvaz 2007; Okay et al. 2010; Görür and Elbek 2013). Okay et al. (2010) pointed out that the Eocene sequence overlies two different stratigraphic basements where it is exposed. Slate, limestone and phyllite make up the small inliers under the Upper Eocene conglomerates and limestones north of Gulf of Saros. Okay et al. (2010) interpret these rocks as continuation of the Circum-Rhodope Belt in Greece. However, in the Circum-Rhodope Belt deformed rocks, including serpentinites, are unconforbaly covered by the Middle Jurassic Doubkon Molasse, of which no trace is seen in Turkey. It is more likely that the Circum-Rhodope Belt is cut out by a fault and does not continue into Turkey. The presence of inferred Cretaceous ophiolites north of the Gulf of Saros as shown by Özaydın and Erol (1981) supports this interpretation. Tentatively, we count the Saros outcrops as parts of the sub-Thracian subduction–accretion complex.

Farther to the southeast, just west of Şarköy, within the ‘handle’ of the oar-shaped Gallipoli Peninsula, a basement piece crops out from under Upper Bartonian to Lower Priabonian shallow marine limestones of the Soğucak Formation. In places erosional remnants of an Upper Ypresian transgressive sequence of sandstones, conglomerate and higher up sandy limestones, marls and shales of the Dişbudak Formation underlie the Soğucak Limestones. Farther south ophiolites have been drilled under the Dişbudak Formation (Şentürk and Okay 1984; Okay et al. 2010). Within the basement under the Soğucak Formation are blueschists (Şentürk and Okay 1984; Topuz et al. 2008). Topuz et al. (2008) distinguished two separate blueschist assemblages here: A serpentinite sliver with numerous dykes of metadiabase showing incipient blueschist facies metamorphism and a well-foliated and thoroughly recrystallized group consisting of blueschist, marble and metachert. The age of the blueschsits are established to be ca. 86 Ma (Coniacian–Santonian) by Rb–Sr phengite–whole rock and incremental 40Ar–39Ar phengite dating (Topuz et al. 2008). They clearly belong to the sub-Thracian subduction–accretion complex.

From the Gallipoli Peninsula, Le Pichon et al. (2014) followed the southern margin of the Thrace basin in the basement of the Sea of Marmara using the available seismic reflexion data to south of İstanbul (their fig. 1a). South of that margin the suture is largely followed by the northern branch of the North Anatolian Fault (e.g., Le Pichon et al. 2014; Şengör et al. 2014). The associated magmatic arc crops out discontinuously north of and within the city of İstanbul and its andesites and andesitic basalts have yielded ages ranging from 75 Ma to 65 Ma (Yavuz and Yılmaz 2009; Aysal et al. 2017). One microdiorite from the southeast of the city gave an age of 58 ± 1.8 Ma indicating the survival of the arc magmatism here into the late Palaeocene. Observations south of the suture north of Bursa, for example, clearly show the continuation of the collision-related shortening here into the later Oligocene or even the early Miocene.

Farther east, between the Armutlu Peninsula and the Almacık Mountains Akbayram et al. (2016a) recently described the geology of the Intra-Pontide suture, correcting a number of earlier errors. Their new paleontological and stratigraphic data show that two successive subduction–accretion complexes formed along the western and the central segments of it: (1) an Upper Cretaceous blueschist-eclogite facies metamorphic prism in the Biga Peninsula in southwestern Sea of Marmara, and (2) an Upper Jurassic–Lower Cretaceous accretionary complex metamorphosed up to greenschist-epidote amphibolite facies between the Armutlu Peninsula and Almacık Mountains. The new paleontological data indicate that collision between the İstanbul Zone and the Sakarya Continent must have occurred later than the early Ypresian. This is in contrast to the recently reported Santonian closure of the Intra-Pontide Ocean (Akbayram et al. 2013). This also corrects the Maastrichtian age ascribed to the Intra-Pontide collision earlier by Şengör and Yılmaz (1981) suggesting instead a possible medial Cuisian closure (ca. 51 Ma). This closure age can also be correlated with the similar, previously published fission-track uplift ages along the Intra-Pontide suture. Following the medial Cuisian closure of the Intra-Pontide Ocean, a wedge-shaped transtensional intramontane basin within the Thrace molasse basin opened (Leigh H. Royden, written communication, 2014) while the westerly escaping fragments tore off pieces of the Intra-Pontide suture and stacked them in the west, in the Biga Peninsula. The intervening continental slices in the form of strike-slip duplexes have given Ediacaran basement ages (protoliths of the micaschists range from 582 to 559 Ma and the metavolcanic rocks yield an age of 577 Ma) and may well represent pieces of the Rhodope–Pontide Fragment as argued by Tunç et al. (2012), but brought far from the east. This strike-slip episode may have already begun by 41 Ma ago, on the basis of using hydrogen isotopic compositions (dD) and 40Ar/39Ar geochronology of authigenic illite in clay gouge-bearing fault zones along the North Anatolian Fault northern strand, i.e., a few million years after the completion of the collision here (Boles et al. 2015). It is clear that the North Anatolian Fault strands later used these older right-lateral faults, invalidating the earlier claim of a 52 ± 1 km right-lateral offset along the North Anatolian Fault here (Akbayram et al. 2016b), in contrast to the observations of an offset an order of magnitude less farther west (cf. Şengör et al. 2014). The significance of this Eocene right-lateral movement, widespread in northern Turkey, will be discussed in the companion paper while describing the palaeogeographic evolution of Turkey and the surrounding areas.

Farther east, east of the town of Bolu, on the Bolu–Mengen road, Göncüoğlu et al. (2008) mapped parts of what is known as the Arkotdağ Mélange. A sliver in it includes a large block of massive and pillowed lavas with radiolarian chert interlayers and intra-pillow mudstones. The silicified mudstones from the upper part of an intact section yielded moderately preserved but diverse radiolarians of late Kimmeridgian to early Tithonian age. Major, trace and REE data from the tholeiitic basalts indicate generation in a mid-ocean ridge setting. Magma was probably derived from a spinel lherzolite source by 5–10% partial melting and fractional crystallization processes. The Nd isotopic data suggest heterogeneity of the source. Combined with comparative evidence from a number of similar mélanges along the inferred suture belt in NW Anatolia, it is concluded that sea-floor spreading in the Intra-Pontide Ocean continued at least from middle medial Jurassic to medial late Cretaceous. Middle Eocene Soğanlı Formation of sandstones and limestones unconformably cover the mélange.

The magmatic arc associated with the subduction of this ocean is to the north and its magmatic front migrated from north (partly beyond the present shore-line) to south between the late Cretaceous and the Palaeocene (see especially Boles et al. 2015).

When the ocean the closure of which led to the formation of the Intra-Pontide suture began opening has been discussed in some detail by Görür et al. (1983), Koçyiğit et al. (1991), Nicosia et al. (1991), Yılmaz and Kandemir (2006) and Kandemir and Yılmaz (2009) who showed that a rift topography originated here during the Sinemurian/Hettangian interval all the way from Bilecik and Mudurnu in the west to İspir–Yusufeli region in the east. They showed that on the upthrown blocks a Rosso Ammonitico facies developed, whereas in the intervening downthrown blocks turbidites commonly interlayered with volcanic rocks were laid down. The high areas (horst blocks) are seen in Bilecik, Ankara, Giresun and a less high-standing tilted block in the Yusufeli region, whereas in Mudurnu, Havza, Amasya, Niksar-Reşadiye, Gümüşhane and Bayburt regions deeply subsided blocks (grabens) are encountered. The fact that the deeply subsided rifted areas pass north of the Tokat Massif (Havza, Amasya, Niksar, Reşadiye) indicates that the Tokat Massif cannot be considered a part of the Rhodope–Pontide Fragment during the Alpide evolution and the attempt to unite the Sakarya Continent with the Eastern Pontides via the Tokat Massif into a Sakarya Zone (e.g., Okay and Tüysüz 1999) cannot be right, as Seymen had already shown in 1975.

Synopsis of the geology of the Intra-Pontide suture The Intra-Pontide suture represents an ocean that extended between the Biga Peninsula and almost to Koyulhisar east of the Tokat Massif via Ankara and that began rifting during the Sinemurian-Hettangian and was actively spreading at least during the late Jurassic. Westwards it is thought to extend into the Almopias Zone of the Vardar Suture in northern Greece. Eastward it joins the main northern branch of the Neo-Tethyan suture terminating the Sakarya Continent to the east. It began being subducted along a north-dipping subduction zone below the Rhodope–Pontide Fragment during the late Cretaceous and finally closed during the medial Eocene. The closure seems to have happened earlier in the east (before Lutetian) and later in the west (medial Cuisian). Shortening along it continued into the earliest Miocene in most places where outcrops permit an evaluation.

İzmir–Ankara–Erzincan suture

This is the main Neo-Tethyan suture in Turkey and represents an easterly continuation of the Vardar Zone in Greece and itself continues into the Sevan–Akera–Karadağ and the Zagros sutures in the east (Şengör and Yılmaz 1981). It extends from İzmir in the west via Balıkesir, Bursa, Eskişehir, Ankara, Yozgat to Erzincan and beyond (Fig. 2). Ketin (1966a) was the first to display it prominently in his influential map of the tectonic units of Anatolia. Since the publication of Şengör and Yılmaz (1981) a very large number of studies were undertaken along it and important discoveries made concerning its detailed petrography, petrology and structure, but the main interpretation as presented in Şengör and Yılmaz (1981) concerning its nature and temporal evolution remained intact.

The westernmost outcrops of the İzmir–Ankara–Erzincan suture are located in the Karaburun Peninsula and the neighbouring island of Chios. Here, Robertson and Pickett (2000) published geological maps both of the Karaburun Peninsula and the Island of Chios. A mélange complex consisting of a sandstone–shale matrix and blocks of limestones and radiolarites spanning an age from the Silurian to the Lower Triassic underlies an unconformable platform succession that reaches from the Scythian sandstones of the Gerence Formation to the Maastrichtian, which is then overthrust by the Palaeocene Bornova Mélange. The earlier mélange appears to have been originally a major debris flow deposit and its contact with the overlying platform units is everywhere tectonic although Robertson and Pickett (2000) interpret it as a sheared unconformity. The vergence in Karaburun is westward. In Chios, the mélange does not seem to contain Triassic blocks, but is overlain by deep-water Hallstatt-type pelagic rocks and varicoloured clastic rocks indicating a deep-water environment. The carbonate platform develops on top of that deeper water association.

Robertson and Pickett (2000) interpret the mélange as resulting from a Palaeozoic south-dipping subduction beneath the Menderes–Taurus Block, but there is no other indication of such a subduction in the area as Robertson and Ustaömer (2009) later also admitted. We prefer to view it as an extensional basin fill that became incorporated into the İzmir–Ankara–Erzincan suture zone. It likely represented a local basin along the northern continental margin of the Menderes–Taurus Block. The fact that it is intruded by a granodiorite, of 247.1 ± 2.0 Ma age (Akal et al. 2011) which itself is unconformably covered by the Scythian Gerence Formation strongly suggests that it was an intra-arc affair of the Palaeo-Tethys not dissimilar to the Karakaya, but clearly predating it. Somewhat later than the Karakaya Basin, the Neo-Tethyan branch here began rifting in the Anisian and then a north-facing Atlantic-type continental margin developed until ophiolite obduction and subsequent Palaeocene continental collision destroyed it (Erdoğan 1990a; Erdoğan et al. 1990; Çakmakoglu and Bilgin 2006). The westerly vergence is clearly a product of the steepened and back-thrust suture here.

From the Karaburun Peninsula, the İzmir–Ankara–Erzincan suture continues northeastward along the so-called Bornova Flysch Zone, which is nothing less than a fully developed ophiolitic suture zone. The suture is filled dominantly by a mélange with a shale-sandstone flysch matrix showing a clear turbidite structure where sedimentary structures were spared by the insense deformation. The mélange contains mainly blocks of shallow-water carbonates, some of which may reach dimensions of 20 km or so. These blocks are made up of the same limestones that one encounters within the Platform succession on the Karaburun Peninsula. The age of the matrix reaches from the Campanian to the Danian (Erdoğan 1990b). The structural vergence is mainly eastward, but in places there are also westerly verging structures in the west indicating retrocharriage.

Between İzmir and Balıkesir, the Bornova Flysch Zone also includes ophiolitic slivers. Aldanmaz et al. (2008) studied them using the LREE, LILE, HFSE and L-MREE geochemistry identifying N-MORB, E-MORB and OIB varieties. There are also pieces showing SSZ signatures indicating that the Neo-Tethyan branch here represented was a major ocean displaying a variety of tectonic environments. That the major Taurus ophiolite nappes obducted during the late Cretaceous onto the Menderes–Taurus Block were dislodged from the ocean floor by intraoceanic subduction zones (Whitechurch et al. 1984; Dilek et al. 1999) supports the presence of SSZ ophiolites in their root region.

Farther to the northeast, between Balıkesir and Bursa, Okay and Kelley (1994) and Okay et al. (1998) described the İzmir–Ankara–Erzincan suture zone just to the west of the town of Bursa, southeast of Lake Apolyont (= Ulubat). Here a peridotite traîneau écraseur with subordinate gabbros and mafic dykes lies above an accretionary complex including mélanges. Immediately below the peridotite sheet are garnet amphibolite slices with ages 101 ± 4 Ma agreeing almost perfectly with the ages of the sub-ophiolitic metamorphic slices under the Lycian Nappes in the Taurus Range in the south that gave ages of 104 ± 4 Ma (Whitchurch et al. 1984) supporting the conclusion that here we are looking at the roots of the Lycian Nappes. Another interesting similarity between the Lycian nappes and the ophiolites along the İzmir–Ankara–Erzincan suture between Balıkesir and Bursa is that they are both bereft of the tholeiitic volcanic carapace (cf. Whitechurch et al. 1984). The accretionary complex was metamorphed in blueschist facies, whereas the overlying ultramafic sheet is unmetamorphosed. In a mélange unit below the other suture rocks, but above the blueschists, consisting of recrystallised limestones, greywacke, mafic volcanic rocks and radiolarites, Globotruncana fossils were reported.

What is interesting about this segment of the suture is that the ophiolitic traîneau écraseur does not bend down northward into a root zone as one would have expected, but is cut off abruptly by a vertical fault and brought into juxtaposition against the Karakaya Complex of the Sakarya Continent, in sharp contrast to the regions farther west where a substantial continental piece intervenes between the Neo-Tethyan suture represented by the Bornova Flysch zone and the Karakaya Suture. Okay and his co-workers rightly suspected here a right-lateral strike slip fault that elides a substantial chunk of terrain and that was probably in place already right after the collision (e.g., Okay and Tüysüz 1999, p. 487). That fault is part of the much larger Eocene right-lateral strike-slip fault family that traverses entire northern Turkey close to the northern Neo-Tethyan sutures as we saw already above in the case of the Intra-Pontide Suture and the Tokat Massif. In the area just described, the early Eocene Orhaneli and Topuk granodiorites cut the ophiolitic nappe and the underlying blueschists (Harris et al. 1994).

This suture geometry, minus the fault that juxtaposes Karakaya suture and the Neo-Tethyan suture, is exposed from Orhaneli all the way to Sivrihisar, some 115 km west-southwest of Ankara. In this zone, near Orhaneli, Okay (2002b) made the remarkable discovery of some of the lowest geotherms ever recorded in the continental crust, indicating rapid subduction to substantial depths of the subduction–accretion rocks below the overriding traîneau écraseur. In addition, Topuz et al. (2018) obtained 264 to 258 Ma ages from the Karakaya ophiolites here and the greenschist metomirphism gave an age of 201 Ma, corresponding with the closure of the Karakaya marginal basin.

The northernmost outcrops of the İzmir–Ankara–Erzincan suture between İzmir and Ankara occur just south of Nallıhan where the ‘Central Sakarya Ophiolitic Complex’ is sandwiched between the Sakarya Continent to the north and the Menderes–Taurus Block to the south. Nebert et al. (1986) and Tekin et al. (2002) and Göncüoğlu et al. (2006) studied the suture here in some detail (although Göncüoğlu et al. seem unaware of Nebert et al.’s study which gives much more detailed stratigraphic, structural and petrographic data). The Dağkuplu ophiolitic mélange (Göncüoğlu et al. 2006) contains knockers of serpentinite, gabbro, blueschists, neritic and pelagic limestones and mafic rocks associated with radiolarian cherts, pelagic limestones and shales. The radiolarian assemblage here gave an age interval of Lower Berriasian to Lower Hauterivian. In one block, Tekin et al. (2002) recorded late Carnian radiolarians perhaps indicating that the ocean was open by that time, although the possibility that it might be a piece of the Karakaya mixed in with the Neo-Tethyan mélanges cannot yet be definitely discarded. Geochemical data indicate the presence of MORB, island arc tholeiite, calc-alkalic basalts and ocean island basalts. These diverse environments indicate that here too the Neo-Tethyan branch was a wide ocean accommodating different tectonic settings and that there was intra-oceanic subduction as also indicated by the basalts farther to the west, as we saw above.

The suture has a clear south vergence (see especially the structure sections in Nebert et al. 1986), but what Göncüoğlu et al. (2006) identified as the Sakarya Continent (which they inappropriately term the ‘Sakarya Composite Terrane’; equivalent to Nebert et al.’s ‘crystalline zone’) has a clear northerly vergence. The rocks here include phyllites, micaschists gneisses and quartzites and are thus not parts of the Karakaya Complex which in this region lies farther north.

Between Sivrihisar and Ankara, the İzmir–Ankara–Erzincan Suture takes a southerly course and goes on to envelop the Haymana Basin in the south as the Samsam Ridge (Görür et al. 1984; Çiner et al. 1996). Here the suture is to the south of the Samsam Ridge where it abuts against another forearc basin belonging to the Kırşehir Massif, namely that of Tuz Gölü (Görür et al. 1984). From Ankara to beyond Sivas the İzmir–Ankara–Erzincan Suture is no longer between the Sakarya Continent and the Menderes-Taurus Block, but between the Sakarya Continent and the Kırşehir Massif, an arc massif that intervenes between the Sakarya Continent and the Menderes–Taurus Block. At this stretch the suture is also two-sided with two subduction zones formerly dipping away from one another.

Here a terminological discussion is perhaps in place to forestall future confusion: in a number of recent publications, the name of the Kırşehir Massif is replaced by the ‘Central Anatolian Crystalline Complex’. This is unnecessary and confusing and results from a disrespect for the existing literature. Chaput (1931) for the first time identified a central Anatolian massif and named it, having been inspired by Philippson’s (1918) ‘Lykaonische Neogentafel’, ‘Massif de Lycaonie’. But both Philippson’s and Chaput’s locality was on the Tuz Gölü Basin and not what has later been called the Kırşehir Massif. Salomon-Calvi (1940) identified a ‘Galatisch-Lykaonisches Massif’ (in the Turkish summary of his article it is called the ‘Konya-Ankara Masifi’) that covered the same ground as Philippson’s (1918) and Chaput’s (1931) feature, but also added the western part of what is today called the Kırşehir Massif. In the same year, Paréjas (1940) excluded the regions west of Tuz Gölü from, and enlarged the regions to the northeast and east of Tuz Gölü, in what he called the ‘Massif de Kızılırmak’. Leuchs (1943) called an enlarged version of Salomon-Calvi’s entity ‘Inneranatolische Masse’, but Blumenthal (1946), at the time a geologist with the Turkish Geological Survey (the Mineral Research and Exploration Institute), wrote, in his commentary on the then new 1:800,000 geological map of Turkey ‘Halys oder Kızıl Irmak Massiv (Syn.: Kırşehir Massiv, Inneranatolische Masse)’ as the name of an entity very close to what we today call Kırşehir Massif. Egeran (1947) followed Blumenthal (1946) and in the first book called Geology of Turkey, Egeran and Lahn (1948) simply called Blumenthal’s feature ‘Kırşehir Massif’ (‘Kırşehir Masifi’; they used Egeran’s 1947, maps). By the rule of priority in scientific publications and to keep some continuity in scientific intercourse we stick to that name. Indeed, in such classic publications as Bailey and McCallien (1953) and Ketin (1956, 1959) the name Kırşehir is always used, although Ketin also uses the geographic term ‘middle Anatolian crystalline massif’ (Ketin 1956, 1959). We see no sensible reason to change the designation ‘Kırşehir Massif’ (in the light of the recent interpretations, it might have been better to call it the ‘Kırşehir Magmatic Arc’, but we prefer to use the older and the shorter name).

Koçyiğit (1991) described the classical locality of the ‘Ankara Mélange’ (Bailey and McCallien 1953) forming a part of the Samsam Ridge. He pointed out that earlier researchers, including Bailey and McCallien, had mistakenly conflated the Karakaya mélanges with those of the late Cretaceous mélanges of the Ankara Mélange. The Ankara Mélange senstu stricto, called the Anatolian Nappe or Anatolian Complex by Koçyiğit inspired by the nomenclature of Bailey and McCallien (1953), is of late Campanian age and consists mainly of chromite-bearing peridotite, serpentinite, gabbro, spilitic pillowed basalt, manganese-bearing siliceous rocks, radiolarites, debris flows of diverse origin and a flyschoid broken formation. This last has an age of Albian to early Campanian. This mélange is now disposed in northeast-southwest striking and southeast-verging thrust sheets. All these rocks are thrust onto the forearc sequence of the Tuz Gölü Basin attached originally to the Kırşehir Massif magmatic arc to the southeast during the late Maastrichtian (Görür et al. 1984; Koçyiğit 1991). The mafic rocks within the mélange indicate both ocean floor tholeiitic and ocean island alkalic basalt affinities (Tankut 1984; Floyd 1993; Rojay et al. 2001, 2004). Tankut also pointed out the presence of calc-alkalic basalts and argued for a phase of subduction-related magmatism here. Rojay et al. (2004) dated their pillow basalts palaeontologically as Callovian-Lower Aptian, an age range in agreement with other observations along the İzmir–Ankara–Erzincan Suture as we saw above.

Recently, Okay and Altıner (2016) described the equivalents of the Soğukçam Limestones. These are equivalents of the famous Maiolica in the Apennines representing a deep-water facies at the southeastern foot of the margin of the Sakarya Continent. From the late Cretaceous onwards as these rocks were being laid down an accretionary prism began growing to the southeast forming the Samsam ridge. On top of this prism continental redbeds commence sedimentation in the Alcı area going upwards into volcanic-bearing debris flow deposits clearly eroding an arc. In the Orhaniye area there is a flysch sequence atop the accretionary prism coarsening upward and passing into continental clastics to be overlain by shallow-marine reefal deposits (Ocakoğlu and Çiner 1995).

The contention by Okay and Altıner (2016) that the Sakarya continent was an arc with intra-arc basins trapping sediment to explain the Maiolica deposition cannot be true, because at the time of the Maiolica deposition the entire eastern part of the Sakarya Continent had subsided to considerable depths as already shown by Şengör and Yılmaz (1981) and that there was no arc piercing it until the late Cretaceous. Koçyiğit et al. (2003) dated the earliest arc products north of Ankara as Upper Campanian. Okay and Altıner (2016) treat the Sakarya Continent as an integral part of the entire Pontides, disregarding the presence of the Intra-Pontide Suture branch of the Neo-Tethys, which cripples their interpretation.

From Ankara, the İzmir–Ankara–Erzincan suture swings north and describes a fairly tight arc around the Çankırı Basin (which Görür et al.1984, had inappropriately called the Çorum basin) and then turns eastward to continue in an east–west direction. The Çankırı Basin is floored by ophiolitic mélange and interrupts the east–west outcrop continuity of the Sakarya Continent, which appears only on its both sides. We thus treat the entire ‘basin’ as a major allochthon lying atop the Sakarya Continent. It was recently described by Kaymakçı et al. (2000, 2003a3, 2009) who pointed out that the basin contains more than 3 km of pre-Middle Miocene section lying on top of a late Cretaceous subduction–accretion complex. Kaymakçı et al. interpret the Upper Cretaceous section as a forearc fill and the Palaeocene to Miocene section as a foreland sequence. While we agree with the former, we think a pick-a-back basin interpretation is more appropriate for its later history as the main foreland basin lay to the north of the Çankırı Basin (with respect to the Çankırı allochthon) and not within it. It is the Boyabat Basin, originally a fore-arc basin deformed with southerly verging thrusts, was redeformed by northerly verging thrusts from the medial Eocene onwards north of the Kargı Massif (Leren 2003).

The subduction–accretion complex flooring the basin consists of a mélange of serpentinised peridotites, harzburgites and pyroxenites, cumulate gabbros and trondjhemites, diabase dykes, spilites and pillowed lava, red to purple radiolarian chert, cherty limestone and reddish pelagic mudstone. The mélange is in places bereft of matrix and seems to represent a complete, but highly disrupted ophiolite nappe. The earliest forearc sediments lying on top of it are red clastics and fossiliferous limestones, much like those in the Ankara Mélange, of which the Çankırı basement is just a northerly continuation. Continuing thrusting has in places incorporated the forearc sequence into the underlying mélange and in such places the forearc sequences appear as a pseudomatrix to the mélange (Kaymakçı et al. 2010).

It seems that the northwestern margin of the basin experienced an anti-clockwise rotation and the eastern margin an opposite rotation from pre-Eocene time to the medial Miocene (Kaymakçı et al. 2003b). Kaymakçı et al. (2003b) interpret this as an indentation phenomenon, but the spreading of an allochtohonus sheet seems a more realistic interpretation, especially because it is unlikely that a mélange-dominated forearc would have the strength to indent a continental block. Moreover, the strong negative gravity anomaly that characterises the Çankırı Basin (Kaymakçi et al. 2010) supports the idea that it has a mass deficiency below it. This is consistent with the idea that it has depressed a continental piece beneath it, albeit later strike-slip tectonics opened smaller pull-apart basins within it locally augmenting the negative signature.

From the southern base of the Çankırı Basin eastward to Sivas, there have not been many studies after 1981, although this is a critically important segment of the İzmir–Ankara–Erzincan suture, because here two subduction zones dipping away from each other collided during the Eocene. One would expect to find a bivergent subduction–accretion complex. However, both the older studies by Ketin (1956, 1959), and a more recent study by Çörtük et al. (2016) found here a dominantly south-vergent ophiolitic mélange association thrusting onto the Sorgun-Yıldızeli flexural basin perched on the Kırşehir Massif to the south studied by Akçay and Beyazpirinç (2017). The Sorgun-Yıldızeli Basin extends east–west south of the İzmir–Ankara–Erzincan suture and sits on the late Palaeozoic–Mesozoic Akdağ Madeni basement (for the literature on the Akdağ Madeni basement see Yılmaz et al. 1995). It is overthrust from the north by the ophiolitic mélange (Yılmaz et al. 1995; Çörtük et al. 2016). Near Artova, north of the Sorgun-Yıldızeli Basin, the ophiolitic mélange is made up dominantly of variably serpentinised peridotites, amphibolites, garnet-micaschists, calcschists, marble, basalts, sandstones and neritic limestones. The geochemistry of the amphibolites indicates E-MORB and OIB characteristics, whereas the dolerite dykes resemble island arc tholeiites (Çörtük et al. 2016). As such the mélange very much resembles other mélanges all along the İzmir–Ankara–Erzincan suture and suggests a wide ocean with multifarious environments. All these rocks are imbricated in south-verging thrusts and the entire complex is thrust over the contents of the Sorgun-Yıldızali Basin which Akçay and Beyazpirinç (2017) called the Boğazköy Formation of Upper Palaeocene–Middle Eocene age. What is interesting about this Formation is that it contains volcanic members: calc-alkalic basalts, basaltic andesites, pyroclastic rocks, dacites and rhyolites. These volcanic rocks give isotopic ages between 57 and 45 Ma indicating here the ongoing activity of the subduction zone dipping beneath the Kırşehir arc. An intrusive gabbro has been dated at 51 Ma (Akçay and Beyazpirinç 2017). The Sorgun-Yıldızeli Basin thus has the character not of a foreland, but a hinterland basin and the north-dipping thrusts deforming it are backthrusts. Where exactly the ophiolites pertaining to the Kırşehir forearc and those pertaining to the Sakarya Continent can be demarcated is now not possible to determine, because this question has not yet been addressed by the field-workers in the area.

Farther east, the Sivas Basin extending between the towns of Sivas and Erzincan is the last segment of the İzmir–Ankara–Erzincan Suture before the suture opens up into the East Anatolian Accretionary Complex. The Sivas Basin originated as a pick-a-back basin on top of an ophiolite nappe that was thrust from the north right across the easternmost end of the Kırşehir arc (Temiz 1996). Some researchers consider this obduction to have taken place directly onto the Menderes–Taurus Block thus considering the Sivas basin to be underlain by it. We disagree with this interpretation, simply because the Inner-Tauride suture that can be followed continuously from Divriği to the Aladağ Range separates the Sivas Basin from the Munzur Mountains (see below). Moreover, in the eastern part of the Basin, the ophiolites sit with thrust contacts directly on the low grade metamorphic rocks of the Kırşehir Massif (Temiz 1996). Yılmaz (1985) studied the ophiolitic mélange here listing pillowed basalt, tuffs, agglomerates, limestone knockers, in places greywacke. In some of the reddish cherty limestone blocks Hettangian–Pliensbachian fossils were found. Elsewhere calcarenites alternating with pillow lavas yielded Upper Oxfordian–Tithonian fossils. Another sample from the vicinity of late Jurassic samples gave a Berriasian age. Yılmaz emphasised that some of these rocks are clearly found in the form of blocks yet in other places they are seen in a regular succession interbedded with the pillow lavas. Finally, Maastrichtian fossils were collected in the upper parts of the mélange matrix. Near Refahiye, serpentinites, gabbros, and marbles, greenschists and metavolcanic rocks are found in thrust contact with each other and with the underlying Sivas Basin fill of Middle to Upper Eocene.

The Refahiye ophiolite was studied recently by Topuz et al. (2013a) and they found a complex, but intact ophiolite nappe some 175 km long and 20 km wide consisting mainly of clinopyroxene-bearing harzburgite and minor dunite. This peridotite body is cut by 20 cm-thick clinopyroxene dykes and later gabbroic intrusions of various sizes. Within isotropic gabbros and their enclosing peridotites there are intrusions of trondjhemite and gabbro pegmatites disposed in dykes of 5 m to 50 cm width. Two trodjhemite samples yielded 184 and 178 Ma ages, being Pliensbachian and Toarcian, respectively. These ages are in perfect agreement with the earlier Liassic ages reported by Yılmaz (1985) on the basis of palaeontology. Topuz et al. (2013a) showed that the non-cumulate gabbros have an island-arc tholeiite signature. They concluded from this and the wide spread of the pyroxene compositions that the Refahiye ophiolite was generated above a subduction zone. They thought this was a Neo-Tethyan subduction, but the age of the Refahiye ophiolite is rather too old for a Neo-Tethyan subduction to be active; the rifting of Neo-Tethys in places was still going on as indicated above. However, the observations may be more satisfactorily explained by resorting to the peculiarities of the mantle below the very young spreading centres of the Neo-Tethys. That mantle had been extensively altered by prolonged, south-dipping Palaeo-Tethyan subduction (see above), perhaps not unlike the situation in the lavas of Bermuda (Mazza et al., 2019). This situation fits what Moores et al. (2000) called the ‘historical contingency’. Although Moores et al.’s interpretation was challenged by Metcalf and Shervais (2008) for the general case, it is indeed applicable to this specific case. The greenschist metamorphics that Topuz et al. (2013a) interpreted as accretionary complex, better fit a fracture zone interpretation, especially in view of the pre-obduction foliations seen in the igneous rocks. When the fracture zone later localised the obduction, the fracture zone metamorphics were smeared below the overriding ophiolite nappe.

Synopsis of the İzmirAnkaraErzincan Suture Since 1981 a vast number of special investigations were devoted to the İzmir–Ankara–Erzincan Suture. These were mainly petrological/geochemical in nature, supplemented in places by palaeontological studies involving mainly microfossils. There is still a distinct dearth of detailed structural mapping. However, we are now in a position of being able to say more detailed things than Şengör and Yılmaz were able to do in 1981, although, with the exception of the Ankara–Erzincan segment, nothing fundamentally different in terms of tectonic interpretation emerged from the later detailed studies. Between İzmir and Ankara, the İzmir–Ankara–Erzincan suture represents an opening sometime during the early Triassic and began closing by north dipping subduction during the late Cretaceous. Okay et al. (1991, 2012) suggested that the segment between İzmir and Balıkesir was a trench–trench transform fault, vaguely reflecting an earlier interpretation by Ricou (1980; but Okay et al. seemed unaware of their predecessor) but subsequent studies disclosed no evidence to support this interpretation. The ocean floor along the İzmir–Ankara–Erzincan Suture was generated mainly during the early Jurassic to early Cretaceous. It was a fairly wide ocean containing diverse environments of mid-ocean ridges, supra-subduction ocean floor and numerous ocean-islands; this wide ocean was the source of the major Tauride ophiolite nappes that we shall discuss below.

Inner Tauride suture

The Inner Tauride Suture was unknown to Şengör and Yılmaz (1981) and was introduced by Şengör et al. (1982) when they separated the Kırşehir Massif from the Menderes––Taurus Block. This invalidates the older concept of the Anatolides (Ketin 1966a) and renders inappropriate the designation Anatolide/Tauride platform, or block, as it had been originally introduced by Şengör and Yılmaz (1981); its continued usage by some is either a result of ignorance or indifference. The western part of the Inner Tauride Suture between Ankara and the Ulukışla Basin is hardly exposed, being buried under the vast Tuz Gölü Basin (Görür et al. 1984). The only reason that we know that there is a suture there is the presence of a blueschist belt delimiting the Tuz Gölü Basin to the west and southwest (see esp. Okay 1984b, fig. 1; Okay 1989a, fig. 4; Candan et al. 2005) and the Kırşehir magmatic arc to its east (Lünel 1985; Türeli et al. 1993; Kadıoğlu and Güleç 1996; Kadıoğlu et al. 1998; Boztuğ et al. 2009a, b; Köksal et al. 2012) with sparse basinal sediment outcrops along its eastern edge and a number of boreholes and seismic sections (Görür et al. 1984; Capraru 1991; Fernández-Blanco et al. 2013). Nowhere do we see the basement of the Tuzgölü Basin, although the sandstones in its fill almost invariably indicate ophiolitic sources in addition to arc provenance (Şahbaz 1992; Ayyıldız 2001). This is compatible with an ophiolitic substratum, but the presence of ophiolite nappes on the Kırşehir Massif makes this interpretation uncertain, as at least in its eastern parts the sediment transport directions are all from east to west, i.e., from the Massif to the Basin. (see Görür et al. 1984).

In the deepest, axial part of the basin, Upper Maastrichtian turbidites are the oldest deposits known. At this time terrestrial redbeds and evaporite lenses were laid down on the eastern periphery of the basin. These sediments pass upwards and sideways into shallow marine, richly fossiliferous sandstones, siltstones and limestones. In the Palaeocene, turbidite deposition with abundant debris flows dominated the basin. Local reefs dotted its eastern margin, but they were quickly overwhelmed by clastic influx from the Massif. Turbidite sedimentation continued during the early and medial Eocene and in the late Eocene the basin was filled and had only shallow marine to terrestrial sedimentation. This was also the time of the major evaporite deposition in the Tuz Gölü Basin (Görür et al. 1984). The Şereflikoçhisar/Aksaray Fault is viewed as an original upper slope discontinuity that originated in the late Maastrictian and consistently separated the Kırşehir Massif with a tendency to rise intermittently and the Tuz Gölü Basin with the opposite behaviour.

Collision in the Tuz Gölü Basin area took place between the Menderes–Taurus Block and the Kırşehir Massif at the latest during the Bartonian, but throughout this time there was no sediment influx from the former into the basin except in the extreme south, in the Ulukışla area (Görür et al. 1984).

Farther to the southeast, the Ulukışla suture is located between the Niğde Massif, which is the southernmost extension of the Kırşehir Massif, and the Bolkardağ metamorphic core complex of the Menderes–Taurus Block. The subduction-related Üçkapılı Granodiorite intrusion is the southernmost representative of the Kırşehir magmatic arc and is dated at 95 Ma by Göncüoğlu (1986), which later stratigraphic and structural studies corroborated (Gautier et al. 2002), in contrast to the anomalous observation by Whitney and Dilek (1997, 1998) giving a Miocene age!

The Ulukışla ‘Basin’ is a deformed forearc, a continuation of the Tuz Gölü Basin to the southeast and it was first identified as marking a suture by Oktay (1982), in which he was followed by Görür et al. (1984). Clark and Robertson (2002, 2005), Alpaslan et al. (2006) and Kurt et al. (2008) have later studied the basin, the latter two groups especially from the viewpoint of the geochemistry of its volcanic rocks. As in the Tuz Gölü Basin, the sedimentary sequences of the Ulukışla Basin sit on an ophiolitic substratum formed by the much disrupted Alihoca ophiolitic mélange that sits on the Bolkardağ basement above a blueschist cushion (Parlak 2016), equivalent to, and southeastward continuation of, Okay’s Tavşanlı Zone (Okay 1984b, 1989a). Parlak wrote the following concerning the blueschists associated with the Bolkardağ Platform and the ophiolites:

‘Glaucophane-bearing HP rocks are present within the mafic volcanic rocks of the Lower Triassic series of the Bolkar Platform (Pourteau et al. 2010) and within blocks of the ophiolitic mélange (van der Kaaden 1966). Another locality of HP rocks is exposed beneath the Kızıltepe ophiolite, a klippe above the Mesozoic carbonate platform of the Bolkardağ Platform farther south relative to the Alihoca ophiolite… A thin metamorphic sole beneath the Kızıltepe ophiolite was overprinted by blueschist-facies metamorphism during the Late Cretaceous (Dilek and Whitney 1997)’ (Parlak 2016, p. 916).

The Alihoca ophiolite itself is made up of highly serpentinised peridotites, pyroxenites, ultramafic to mafic cumulates, isotropic gabbros and subordinate basaltic volcanic rocks (Parlak 2016). Microgabbro dykes intrude all the units with the exception of the volcanic rocks and were probably their feeders. The ophiolite and the associated mélanges were conformably covered by a basal conglomerate passing upwards into cherty limestones and red pelagic limestones of the Çiftehan Formation of Campanian–Maastrichtian age (Demirtasli et al. 1984). The geochemistry of the ophiolitic rocks suggests a suprasubduction zone setting compatible with pre-arc spreading and incipient island-arc products. Hornblendes separated from a dyke cutting the gabbro-pyroxenite rocks of the Alihoca ophiolite yielded 90 Ma (Dilek et al. 1999). The Mersin ophiolites to the south are continuations of the Alihoca ophiolite, which we discuss below when describing the Menderes–Taurus Block.

The first forearc sediment fill is represented by the Maastrichtian-Lower Palaeocene Aktaştepe Formation comprising conglomerates and sandstones at the base, which Clark and Robertson (2005) interpret as fluvio-marine, although their organic geochemistry shows them to be terrestrial (Sonel et al. 2008), just like the equivalent sedimentary rocks around the Tuz Gölü Basin. These rocks pass upwards into fine-grained calcarenites, micrites and sandstone–mudstone intercalations with rare conglomerate horizons. These rocks are followed by about 2 km-thick turbidites in the southern part of the basin, containing some pillow lavas and volcaniclastic rocks, whereas towards the basin interior northwards coarse clastics are associated with abundant volcanic rocks. The entire section shallows upward exactly as in the Tuz Gölü Basin, but then re-deepens depositing turbidites. Finally, the Upper Eocene is represented by evaporites, again mimicking the development in the Tuz Gölü Basin.

The lavas and dykes reported from the Ulukışla area are medium to high potassic calc-alkalic rocks (Alpaslan et al. 2006; Kurt et al. 2008) very much resembling the volcanic rocks of the Quaternary South Italian volcanic province (Conticelli et al. 2010) and probably erupted in a similar subduction setting. Most authors interpret the tectonics of the Ulukışla area as a post-collisional extensional basin, but there is no evidence for a late Cretaceous collision. Neither the sedimentation nor the vulcanicity ceased after the Cretaceous and the main deformation was at the end of the Eocene, as Oktay (1982) had shown decades ago. The tectonic setting of the Ulukışla Basin much resembles the Weber and the Aru Troughs in the Arafura Sea, in which shortening and extension are going on at the same time owing to the vicissitudes of the accommodation of the collision front to the geometry of the foreland (e.g., Charlton et al. 1991; Pownall et al. 2016). Pownall et al. (2016) in particular have shown how the roll back of the subducting slab created the extension in the Weber forearc, having literally ‘rolled it open’ causing it to sink down to 7.2 km, the deepest part on earth outside a deep-sea trench! Such complex processes greatly complicate the subsidence patterns as Clark and Robertson (2005) noted in the Ulukışla Basin. That volcanism in such cases can even extend onto the forelands, albeit commonly in extreme alkalic compositions, is shown by the presence of the volcano Vulture in southern Italy (e.g., La Volpe et al. 1984; Melluso et al. 1994). It is thus clear that continental collision along the Ulukışla Basin took place during the medial to late Eocene as Oktay (1982) had already surmised and that the extensional scenarios are nothing more than the rolling open a forearc basin above a backrolling subduction hinge literally just before the final suturing as in the Banda Arc area. The vulcanicity resembles that associated with a similar pre-to syncollisional roll-back event Italy. Such an interpretation is much more in accord with the events in the Tuz Gölü Basin to the west and the Şarkışla area to the east.

To the east, the Ulukışla Basin is abruptly truncated by the Ecemiş Fault Zone (for a modern description of his fault zone see the papers in Ecemiş Fay Kuşağı Çalışma Grubu (EFKÇG) 2001 and Yıldırım et al. 2016) and the Inner Tauride Suture is thrown some 210 km north-northeastward to the Akkışla area, from which it continues in an east-northeast direction to the south of the Sivas Basin, beyond which it opens into the East Anatolian Accretionary Complex, thus very much following the classical Anti-Taurus. Near the town of Akkışla, the suture is characterised by narrow, east-northeast striking bands of serpentinites thrust from the northwest by what is mapped as non-descript ‘ophiolite’ (Bilgiç 2002). It is unclear whether the ophiolitic mélange exposed southeast of Sivas in the Hınzır Mountain is a part of the same suture or a part of the supra-Kırşehir ophiolites (see Aydal et al. 2008). The same is true for the late Cretaceous ophiolitic mélange described by Özaksoy and Gökten (1986) between Felahiye and Özvatan.

The ophiolites expelled from this suture form the Pozantı-Karsantı ophiolites of the Aladağ Massif. After 1981, Polat (1994), Polat and Casey (1995) and Parlak et al. (2000, 2002) studied these ophiolite nappes now consisting mainly of the mantle section including dunites with podiform chromite, harzburgites and layered cumulates consisting of dunites containing chromite (see Rahgoshay et al. 1981), wherlites, diallagites and gabbros and trondjhemites, sheeted dykes, the higher volcanic carapace having been removed during and after obduction, although diabase dykes intruding both the ultramafic foundation and the gabbro section are present (Çakır et al. 1978; Polat 1994; Polat and Casey 1995; Parlak et al. 2000, 2002). Polat et al. (1996) and Parlak et al. (2000, 2002) showed that the earliest accreted materials below the ophiolite consist of alkalic ocean island basalts metamorphosed to lower to upper amphibolite conditions. The highest temperature accretion seems to have happened 94 to 90 Ma ago (i.e., during the Turonian) within the ocean. As the ophiolite marched on, low-temperature accretion occurred below it, adding first pelagic sediments and juvenile ensimatic island arc volcanic rocks and then, as the nappes began climbing onto the continental slope of the Menderes–Taurus Block, mixed sources now including sediments derived from the continent. Polat and Casey (1995) show that the final obduction took place during the Maastrichtian, the ophiolite nappe moving from the northwest to the southeast.

North of the suture is the Sarkışla ‘Basin’ which is really part of an ensialic magmatic arc of late Cretaceous to latest Palaeocene (Thanetian) age. Gökten (1986), Özaksoy and Gökten (1986) and Gökten and Floyd (1987) studied this arc and its sedimentary cover in the Şarkışla Basin area and Clark and Robertson (2005) summarised Gökten’s (1986) results in their fig. 20. From Gökten’s and Gökten and Floyd’s work, it seems clear that here too the collision occurred sometime after the Palaeocene (late Cretaceous-Palaeocene arc vulcanicity consisting of andesite-dominated calc-alkalic rocks that developed in an ensialic arc and lively seismicity as suggested by frequent slumps, which Gökten associated with ongoing volcanic activity continued into the Thanetian). Lutetian was bereft of volcanism and a time of quiet carbonate sedimentation (also farther west: see Özaksoy and Gökten 1986). Only with the onset of the Oligocene do we see deposition in restricted basins laying down gypsum and alternating conglomerates, sandstones, siltstones, marls and even limestones giving the appearance of a typical molasse (Gökten 1986; Özaksoy and Gökten 1986). We thus take the collision to have happened sometime during the Eocene.

Eastwards the Inner Tauride Suture goes north of the Munzur Mountains and separates the Sivas Basin basement from the Menderes Taurus Block before running into the East Anatolian Accretionary Complex. East of Şarkışla, the southern border of the Sivas Basin is made up of the Güneş ophiolitic mélange (Gökten 1993) which is also called the Divriği ophiolite (İnan et al. 1993), consisting of a highly sheared serpentinite matrix containing red-coloured radiolarites, reddish pelagic limestones, gabbros, diabases and oxidised volcanic rocks. It also has large neritic limestone knockers of Upper Jurassic to possibly Lower Cretaceous age plucked from the underyling autochthon belonging to the Munzur Mountains. The ophiolites are unconformably covered by Maastrichtian limestones (İnan et al. 1993) and Lutetian turbidites and debris flows (İnan et al. 1993; Gökten 1993). The unconformable sedimentary rocks have no volcanic input (Gökten 1993). Gökten concluded that here ocean closure happened sometime during the early Eocene, although ophiolite obduction was post-early Cretaceous and pre-Maastrichtian. The Lower Oligocene sediments are entirely terrestrial (Gökten 1993; İnan et al. 1993).

East Anatolian Accretionary Complex

Ketin (1977) mapped several areas between Lake Van and the Turco-Iranian frontier demonstrating that these areas, in which Arni (1939) had discovered earlier an imbricated structure much resembling the structure of present-day subduction–accretion complexes, are underlain entirely by ophiolitic mélange and flysch. Şengör and Yılmaz (1981) took this observation and compared other areas in eastern Turkey with them to show that much of eastern Turkey was underlain by a large subduction–accretion prism, the correlative of which is still actively growing in Makran. Since then many new observations have fortified this interpretation and we summarise the geological ones in our Figs. 10a–g and 11h–k, taken from Sengör et al. (2008a, b). See for further references to new research, especially on the geomorphology, petrology and geophysics of the High Plateau, Şaroğlu and Yılmaz (1987), Yılmaz et al. (1987b, 1998), Şengör et al. (2003, 2008), Zor et al. (2003, 2007), Gök et al. (2003), Allen et al. (2004), Barazangi et al. (2006), Bektaş et al. (2007), Keskin (2003, 2007), Collins et al. (2008), Göğüş and Pysklywec (2008), Türkoğlu et al. (2008) and Zor (2008). Only one paper has so far been published disputing the subduction–accretion complex interpretation for the basement of the Eastern Turkish High Plateau (Topuz et al. 2017), but that paper presents no evidence whatever to support its rebuttal. Our Fig. 7 shows their observation points, which speaks for itself. Their argument is as much a non-sequitur as finding Palaeozoic and Proterozoic zircons (Pilot et al. 1998) in the Mid-Atlantic ridge and then concluding from this observation that the entire Atlantic is underlain by old continental crust (which Pilot et al. did not).
Fig. 10

Simplified geological maps showing various parts of Eastern Turkey where the East Anatolian Accretionary Complex crops out. They are all from Sengör et al. (2008a, b), where the sources are also given. a Sketch geological map of the accretionary complex rocks north of Beşparmak, east-northeast of Lake Van. The ‘young basalts’ are Quaternary. b Sketch geological map of the accretionary complex rocks west of Lake Erçek to the east of Lake Van. c Sketch map of the accretionary complex rocks 2 km west of Özalp between Lake Van and the Iranian frontier. Note the southerly vergence which is the opposite of the vergence farther west along the northeastern shores of Lake Van. d Sketch geological map of the accretionary complex rocks east of Lake Erçek, between the lake and Yeniçavuş. The greenschists represent a metamorphosed and thus deeper part of the accretionary complex. e Sketch map of the accretionary complex rocks at the Iranian frontier along the Kotur Creek, directly due east of the city of Van. Note the presence of a large marble and quartzite knocker and the backthrusting (retrocharriage) observed behind it. Could this be a Gondwana-Land fragment attempting to overthrust the accretionary complex as is seen north of the Bitlis Massif in the region of Gevaş (Fig. 9g)? f Sketch geological map of the area between the city of Van and Lake Erçek This map shows a part of the Tertiary (Eocene to? late Oligocene) mélange forming the younger, southerly sections of the East Anatolian Accretionary Complex. g Sketch geological map of the southern mountainous frame of Lake Van and a section near Gevaş. Notice here that the Lake is framed by an ophiolitic mélange. They are in any case accretionary complex rocks overridden by the Bitlis massif metamorphics. A clastic Eocene unit (with Discocyclina sp., Alveolina sp. and Nummulites sp.) covers both the Bitlis Massif and the accretionary complex rocks, although its clast composition is exclusively dependent on which rock type it lies. Now the Eocene is also highly tectonised and cannot be proved to have been an overlap assemblage astride the contact between the ophiolites and the Bitlis Massif

Fig. 11

Simplified geological maps showing various parts of Eastern Turkey where the East Anatolian Accretionary Complex crops out. They are all from Sengör et al. (2008a, b), where the sources are also given. h Sketch geological map of the region between Erzurum and Akdağ showing the northernmost parts of the East Anatolian Accretionary Complex, covered by the Eocene arc volcanic rocks southeast of Erzurum. All areas left white represent Neogene and Quaternary cover. Notice the presence of large marble knockers. Notice also that south of the Eocene volcanic rocks, all one sees is ophiolites and flysch bearing knockers of ophirags, pelagic limestone, schist and marble. This is a field aspect that does not change until one reaches the Bitlis Massif in the south, or the Sanandaj-Sirjan Zone in Iran in the extreme south-east. i Sketch geological map of the Akdağ region. All areas left white represent Neogene and Quaternary cover. Notice the presence of large marble knockers. The ones seen here are the largest representatives of their kind in entire eastern Turkey and northwesternmost Iran. This is the area where Topuz et al. (2017) claim to have discovered continental basement. j Geological map of the Muş Basin. Here, as in Fig. 10d above, notice the structural conformity between the structures in the cover and the accretionary complex (basement). k Geological map of the Ahlat-Adilcevaz region northwest of Lake Van. The folds in the Miocene and younger cover rocks are of flexural slip type. Notice the parallelism between the fold axial traces in the younger cover and the thrust fault traces and fold axial trends in the East Anatolian Accretionary Complex (basement), showing continuity of shortening of the soft accretionary complex basement

However, Figs. 10 and 11 are not the only collection of evidence showing the nature of the basement under Eastern Turkish High Plateau. Figure 12 shows selected GPS velocities in eastern Turkey and it is seen that the velocities do not diminish when the Bitlis suture complex is crossed, but they do so when the northern boundary of the East Anatolian Accretionary Complex is crossed, indicating that the entire are is underlain by easily deformable ‘mushy’ rocks. Moreover, the same area has a crust thinner than expected from its elevation showing that despite the great amount of shortening, as shown by the presence of the East Anatolian Syntaxis, there has been no substantial crustal thickening. This is because the area started with a thin substratum to begin with and not with a crystalline continental basement. Shortening of a normal continental basement as often depicted also by Aral İ. Okay in his various publications, would have created a geology similar to that seen in Western Turkey.
Fig. 12

GPS velocities plotted relative to Eurasia with 1σ velocity uncertainties after Reilinger et al. (2006, Fig. 2). Note that within the East Anatolian Accretionary Complex (EAAC) the velocities seem to decrease very gradually. Only after its northern boundary is crossed northwards do the velocities abruptly diminish. This suggests that the EAAC basement acts as a ‘squashy zone’. The dashed arrow in the southeast represents a vector that is plotted some 100 km too far west to show a minimum velocity for the motion of Arabia here. The velocity is a minimum, because the GPS station is located within the actively shortening Zagros foreland fold/thrust belt

The East Anatolian Accretionary Complex began growing with the onset of late Cretaceous subduction under the Rhodope–Pontide Fragment and stopped growing when its toe collided with the Bitlis Massif in the south sometime during the late Oligocene. It represents a ‘suture knot’ where the İzmir–Ankara–Erzincan, Inner Tauride, Sevan–Akera and the Zagros sutures come together.

Sutures of the southern branch of the Neo-Tethys

The sutures of the southern branch of the Neo-Tethys consist only of the Antalya and the Bitlis sutures that are parts of the Ayyubide (Antalya) (Şengör and Stock 2014) and Assyride (Bitlis) (Şengör et al. 1982) orogenic zones.

Antalya suture

Not much progress occurred concerning the Antalya suture after 1981 and we refer the reader to Şengör and Yılmaz (1981) concerning its structure and evolution. The only two papers that introduced significant novelties are that by Okay and Özgül (1984), who discovered high-pressure/low-temperature metamorphic rocks in the Sugözü Nappe sandwiched between an upper Yumrudağ Nappe and a lower Mahmutlar Nappe, both metamorphosed only up to greenschist to lower greenschist grade and that by Marcoux et al. (1987) who observed mesoscopic structures in Lower Triassic anhydrites along the boundary between the Kemer and the Alakırçay nappes. Okay and Özgül’s (1984) discovery supports the hypothesis by Şengör and Yılmaz (1981) that there was subduction involved in the emplacement of the Alanya Nappes.

The detailed observations by Marcoux et al. (1987) are consistent with southward thrusting of the upper Kemer Nappe onto the lower Alakırçay Nappe. Marcoux et al. (1987) see in this conclusion a support for Marcoux’s long-held view of a northerly derivation of the Antalya Nappes. However, it has also long been clear that the Antalya Nappes west of the Gulf of Antalya were emplaced along left-lateral, north-striking strike-slip faults (Şengör and Yılmaz 1981 and the references therein) and northeast-striking thrusts of any vergence are also compatible with that interpretation. Given the very limited observations of Marcoux et al. (1987; only two outcrops) their arguments about the regular straining of the anhydrites lose much of its strength and it seems best to reserve judgement on the general conclusion of their paper. One of the authors of that paper, Jean-Pierre Burg, later also expressed his reservations on their final conclusion (personal communication, 14th June 2018). All other observations remain compatible with the interpretation presented in Şengör and Yılmaz (1981) and Şengör and Stock (2014).

Bitlis Suture

The best exposure of the Bitlis Suture, where all of its elements can be clearly observed is in the Engizek Mountains near its western end in the Misis-Andırın Mountains. Along the southern edge of the Engizek Mountains is a narrow (500 m to 1 km wide), east–west striking belt squeezed between the Arabian Platform in the south and the Malatya Metamorphics, forming a part of the Menderes–Taurus Block, in the north. The slices within this narrow band contain ophiolitic, sedimentary and associated volcanic rocks spanning an age range from the Upper Cretaceous to the Lower Miocene. Within the imbricates, the sequences are complementary, indicating that prior to thrusting they formed one continuous sequence (Yılmaz 1993). Within the imbricated stack, there is a deep-sea sedimentary sequence of Upper Cretaceous to Middle Eocene age resting conformably above a disrupted ophiolite. The ophiolites are associated with a mélange including blocks of Middle Eocene sedimentary rocks (Yılmaz et al. 1993). The pelagic sedimentary rocks include chalks, cherts, marls and calciturbidites (Cona Group). Mafic, intermediate and felsic volcanic rocks accompany the pelagic sediments. These ‘Helete Volcanics’ (Yılmaz 1993) are regarded as slices of a former magmatic arc. The presence of this arc south of the Malatya Metamorphics is important, because it disproves the claims about the Bitlis and the Pötürge massifs, the easterly continuation of the Malatya Metamorphics, being a part of the Arabian Platform without an ocean in between (Yazgan et al. 1983; Michard et al. 1984). The westerly continuation of this arc is documented by its sedimentary products in the Oligocene to Miocene İsalı, Karataş and the Kızıldere Formations (Floyd et al. 1992) of the Misis-Andırın high in the Adana Basin, which are mélanges and turbidites.

In the Engizek Mountains, the deep-sea association is unconformably covered by late Eocene–Oligocene debris flows.

The lower thrust sheets are composed of Oligocene to Lower Miocene, rapidly deposited chaotic sedimentary rocks passing upward into a Lower Miocene shallow marine sandstone which then eventually becomes a flysch. This is regarded a lateral equivalent of the Lice Formation farther east.

North of (and above) the narrow imbricate zone are two major nappe packages

An upper nappe consists of the Engizek Metamorphic Massif, which is a westerly equivalent of the Bitlis-Pötürge massifs. Below that is the lower nappe, consisting of the remnants of a basin that began rifting during the late Maastrichtian and was closed by the late Eocene. This is the Maden Basin that probably opened as a back-arc basin that disrupted a poorly developed ensimatic arc, the erosional products of which have been encountered both in the Misis Complex and in the imbricate zone to the south. Aktas and Robertson (1984) interpreted it as a forearc basin. We do not follow that interpretation, because the Maden Basin opened by disrupting an arc; farther east it disrupted both the Pötürge and the Bitlis Massif itself (Şengör and Yılmaz 1981). Just to the northwest of the Pötürge Massif, the Gülümüsag quartz diorite is dated at 48.1 ± 1.8 Ma (early Lutetian: Yazgan et al. 1983) and is located within the area taken up by the Maden Complex rocks. This is consistent with the interpretation that the Maden Complex was opened disrupting an arc the opening of which seems to have lasted into the early Eocene. While it was opening, the arc magmatism, which had commenced some 85 Ma ago (Michard et al. 1984) shut off between the late Campanian (post-75 Ma) and recommenced about 50 Ma ago, an observation consistent with the marginal basin opening interpretation (see Yılmaz 1993). The Maden remnants overthrust a weakly metamorphic mafic lava and deep-sea sediment association (the Kızılkaya metamorphic association of Yılmaz et al. 1993) and themselves are thrust by the Berit ophiolite metamorphosed up to granulite and eclogite facies, from the north. The entire package lies above an unnamed ophiolitic mélange thrust over the Helete arc volcanic rocks.

North of the Engizek metamorphics there is yet another ophiolite across the Sürgü Fault, which Yücel Yılmaz thinks, on the basis of its straightness and steep dip, to be a strike-slip fault (see Yılmaz et al. 1997a, b). This Göksun ophiolite is stratigraphically ovarlain by an andesitic–dacitic volcanic association indicating the presence of an arc of late Cretaceous age overlain by Upper Cretaceous pelagic sedimentary rocks that are then replaced by a flysch-like clastic sedimentary sequence and shallow sea sediments of Palaeocene and Lower Eocene age (Perinçek and Kozlu 1984; Yılmaz 1993; Robertson et al. 2006).

The documentation of the PT path of the major metamorphosed units of the Engizek Mountain indicates that the continental slab was attached to a subducting oceanic lithosphere and was consequently deeply buried along a subduction zone and metamorphosed as previously suggested by Yılmaz (2010), followed by fairly rapid exhumation. The exhumation period of the continental slab began during the late Cretaceous and extended possibly to the Paleocene time.

The Berit ophiolite also underwent polyphase metamorphism, as did the overlying Engizek Metamorphic Massif that includes granulite-eclogite facies and amphibolite facies as mentioned above (Genç et al. 1993; Oberhänsli et al. 2014; Awalt and Whitney 2018). The exhumation of the continental units probably ended by the end of the Palaeocene, but exhumation of the ophiolitic slab may have lingered on possibly till the early Eocene time, because Sm–Nd (pyroxene–garnet–amphibole–whole rock) isochron ages of 52–50 were obtained from the granulite facies rocks (Karaoğlan et al. 2013).

The Engizek metamorphic massif must have reached the surface before the medial Eocene time, because the Maden Basin sequence of the Middle Eocene age was deposited above the nappe package. Therefore, the exhumation of the Berit metaophiolite and the Engizek Metamorphic Massif must have predated that event.

A similar geometry has been documented farther east, but nowhere can one see as complete a sequence as one can in the west. What is clear is that the ophiolites below the Maden Complex represent the root zone of the Cretaceous Nappes of Kızıldağ and Baer-Bassit on the Arabian Platform and the Göksun and Berit are remnants of a vast nappe that covered the entire Engizek-Bitlis-Pötürge massifs during the late Cretaceous. Şengör and Yılmaz (1981) named this nappe Yükeskova and we retain that designation here, but it is really a continuation towards the east of the Bozkır ophiolites (Özgül 1976).

Both the ophiolites north of the Engizek Metamorphics (the Göksun and the İspendere ophiolites) and those south of the Munzur Mountains now divide up the eastern end of the Menderes–Taurus Block into three digitations: these are called, from north to south, the Munzur, Malatya and the Bitlis-Pötürge. The ophiolitic sutures in between have been interpreted as remnants of small oceans that indent the eastern end of the Menderes–Taurus Block or as strike-slip sutures of the kind mentioned by Dewey et al. (1986). Robertson et al. (2013a) call the purported oceans represented by these sutures ‘unlikely oceans’. We lean in the direction of the second interpretation mentioned above, because the rock content, structure and the timing of events in the digitations are identical with one another (also see Robertson et al. 2013a and the references therein) with a single exception known to us: Zeck and Ünlü (1988) dated the Murmano quartz syenite to diorite intrusion at 110 ± 5 Ma using a Rb–Sr whole rock isochron. This pluton intruded a serpentinite complex northwest of Divriği. Since this serpentinite complex is located south of the Inner Tauride Suture, between the Munzur and the Malatya digitations, Zeck and Ünlü justifiably assumed a Cretaceous age for its emplacement and concluded that the emplacement had to predate the Cenomanian. This is some 15 to 18 Ma earlier than the emplacement of all other Neo-Tethyan ophiolites onto continental pieces, but it is nearly synchronous with the onset of intra-oceanic subduction in the case of all other Bozkır ophiolites. This leads us to think that perhaps at this spot the intraoceanic obduction front was very close to the Menderes–Taurus Block and the ophiolitic nappe made its earliest climb onto its continental margin here. This may be one piece of evidence that the Malatya Digitation was originally not where it is now, but farther east, where its northern margin may have received the ophiolites.

Synopsis of the southern Neo-Tethyan sutures The southern Neo-Tethyan sutures represent the closed parts of those sections of the Neo-Tethys that opened between the Permian and the early Jurassic between Afro-Arabia and the Menderes–Taurus Block. The ocean opened between the Bitlis/Pötürge digitation and the Arabian Platform and the Alanya Massif and the Menderes–Taurus Block sometime during the early Triassic and the Eastern Mediterranean opened, as an ocean, mainly during the early Jurassic although rifting around it seems to have commenced, in a different tectonic setting, already during the Permian. North-dipping subduction under the Bitlis-Pötürge digitation began during the late Cretaceous and led to both marginal basin opening and strike-slip faulting that may have created the transform sutures between the Munzur Mountains and the Malatya Digitation and the Malatya Digitation and the Bitlis-Pötürge Digitation. Massive ophiolite obduction occurred onto the Arabian Platform in the east and onto the Menderes–Taurus Block in the Antalya region in the west during the Turonian–Campanian interval as part of the Ayyubid orogenic events (see Şengör and Stock 2014). Continental collision occurred in the later Eocene both between the Alanya Massif and the Menderes–Taurus Block and the Bitlis-Pötürge digitation and the Arabian Platform although in the latter region, the collision progressed from the late Eocene to latest Oligocene from west to east. Convergence in the east is still active but is now distributed in the entire East Anatolian High Plateau (Sengör et al. 2008a, b).

Palaeo- and Neo-tethyan continental blocks of Turkey

During the late Palaeozoic the entire present territory of Turkey was a part of the northern margin of Gondwana-Land. Of this area, only the westernmost parts of the Strandja Mountains collided with Laurussia during the late Carboniferous, but farther east the ocean remained open. The closed ocean has been called the Rheic since the suggestion by McKerrow and Ziegler (1972) and its unclosed easterly continuation has received the name Palaeo-Tethys (Stöcklin 1974; Hsü 1977; Laubscher and Bernoulli 1977; Şengör 1979). Thus, during the late Palaeozoic the entire area of Turkey lay south of the Palaeo-Tethys as originally suggested by Şengör et al. (1980, 1982) and Şengör and Yılmaz (1981). All suggestions to the contrary have been disproven by observations that accumulated during the past two decades, especially by the evidence that showed that Moesia, the Sakarya Continent basement and the eastern Pontide basement were all parts of Gondwana-Land or at least were in its immediate vicinity. For all the blocks we present selected stratigraphic columnar sections that we think are representative.

The Arabian Platform

Only a small part of the Arabian Platform containing the middle part of the famous ‘Fertile Crescent’ is in Turkey. It is located south of the Bitlis Suture (see above) and now is the site of what Egeran (1947) called the ‘Border Folds’ and Ozan Sungurlu (in Şengör et al. 1982) suggested be renamed ‘Assyrides’, a dominantly south-vergent foreland fold-thrust belt in front of the Bitlis Suture. Being the main hydrocarbon province of Turkey, the Assyrides have been intensely studied by diverse petroleum companies. Their efforts have been summarised in an extremely useful stratigraphic lexicon by Yılmaz and Duran (1997). The following stratigraphic account is based mainly on that book. Figure 13 exhibits representative columnar sections including a number also for areas outside Turkey from Syria to permit comparisons with areas just outside Turkey.
Fig. 13

Stratigraphic columnar sections from selected areas characterising the Permian to Mesozoic geology of the Arabian Platform. We included here columns from Syria to emphasise the unity of the Arabian Platform during the time interval here illustrated (Barrier et al. 2014)

Yılmaz and Duran (1997) summarised the Palaeozoic stratigraphy in two areas: northern and southern parts of the Assyrides. Both parts are underlain by what is called the Telbesmi Formation containing sandstones, conglomerates, shales and sandstone–shale alternations showing large fluctuations in thickness ranging from about barely 12 m (in the province of Adıyaman) to more than 2.5 km in the Derik-Bedinan regions. In the latter region, the formation consists mainly of submarine volcanic rocks including spilites and andesite lavas and tuffs with subordinate volcanic breccias (Derik Volcanics). The upper contact of this formation is in places unconformable, in others conformable with the overlying Sadan Formation. Ketin (1966b) conjectured that the age of the Telbesmi would be what he called Eocambrian, but later ichnofossil finds indicate a lowermost Cambrian (Demircan et al. 2018). Later measurements with single zircon evaporation method on samples collected from the Derik yielded indeed late Ediacaran ages (A. Kröner and A. M. C. Şengör unpublished data). This was later corroborated by zircon LA-ICP-MS data giving crystallization ages of 581.4 ± 3.5 Ma (n = 7) and 559.2 ± 3.2 Ma (n = 3) for the early- and late-stage andesitic rocks, as well as ages of 569.6 ± 1.6 Ma (n = 17), 571.6 ± 1.9 Ma (n = 18), 575.4 ± 4.3 Ma (n = 6) for the rhyolites (Gürsü et al. 2015). Ketin concluded that the variations in the geometry of the Telbesmi were fault-controlled without specifying the strain regime. The overlying Sadan Formation is also dominated by clastics and has an average thickness of some 500 m with fluctuations from about 75 m in Adıyaman to a km near Hakkâri. The Sadan is followed by a carbonate section of some 200 m average thickness dominated by dolomites, indicating a gradual transgression during the medial Cambrian. The dolomites pass concordantly to sandstones and shales of the Sosink Formation of some 1000 m thickness on average indicating a renewed regression.

The section from Telbesmi to Sosink is seen in the whole of the Assyrides; afterwards no Palaeozoic Formation covered the entire area and particularly the Adıyaman-Şanlıurfa-Mardin areas remained land until the Aptian transgression. This is parallel with the persistent tendency to remain emergent of the western part of the Arabian shield during the same interval. This was the case despite the fact that sea-level was persistently higher than today until a major sea-level lowering during the Serpukhovian–Bashkirian boundary (Bilal Haq written communication in Şengör, 2015, fig. 15). We, therefore, should seek the cause of this emergence in tectonic and not in eustatic events. Both in the western and in the eastern parts of the Assyrides there were episodes of flooding in the late Devonian and the late Permian. This is surprising, because both times were times of global sea-level lowering. Especially in the late Permian, the global sea-level fell considerably below its present stand beginning with the Capitanian. The late Devonian episode of flooding during the sea-level lowering resulting from the late Devonian–early Carboniferous glaciation in Gondwana-Land (Isaacson et al. 2008) is as yet difficult to interpret, but the Permian episode was clearly related to the onset of rifting east and north of the Arabian platform heralding the opening of the southern branch of the Neo-Tethys here.

We note that a fossil specimen belonging possibly to a new taxon of the early Permian temnospondyl amphibians Branchiosauridae was found on the Arabian platform in southeastern Turkey (Hoşgör and Fortuny 2012). This may prove the existence of early Permian terrestrial sedimentation in Southeastern Turkey. The authors point out that ‘this clade is well known from the [Lower] Permian of Central Europe and Sardinia’, indicating a Pangaea-wide communication of terrestrial vertebrates between the Arabian Platform and Europe.

The Mesozoic opened with the deposition of the Çığlı Group of dominantly clastic rocks of Scythian age. However, in the eastern part, around the Nusaybin-Cizre-Şırnak-Hakkâri areas, the Triassic begins with the Yoncalı Formation of dominant carbonates passing upwards, in the upper one-fourth of the formation, into greenish-buff to violet-red coloured shales with rare thickly-bedded limestones. The rest of the Çiğli Group consists mainly of similar shales with rare intercalated limestones. In the west, the Yoncalı carbonates are lacking and the section begins with clastics. With the Anisian, dominantly dolomite deposition begins on both sides of the central high area. Towards the high, the sequences are thinner and contain evaporites. In the Jurassic, the land area of the central high is enlarged both toward the east and the west and this tendency to remain emergent continued until the Aptian Areban Formation that covered almost the entire Assyride area. This is again somewhat unexpected, because the world-wide sea-level was gradually rising after a Hettangian–Sinemurian low. Southeastern Turkey was partially flooded during that time, but continued emerging while the worldwide sea-level was rising. Between the Aptian and the Cenomanian the entire area was flooded and deposited dominantly dolomites with limestones coming in during the Cenomanian in the southern regions. Dolomite deposition predominated in the north.

During the Aptian major rifting was going on in Central Africa, in the Sirte Rift and the Marmarican Taphrogen and at this time the shelf regions of the Arabian platform began subsiding, possibly related to a stretching event (see Şengör 2009). This tendency continued until the later Cenomanian, when major basins began rapidly subsiding along the eastern margin of the Arabian Platform. This is the time, when giant ophiolite nappes were being emplaced onto the northern and eastern Arabian margins. This emplacement event was part of a large obduction front from eastern Libya to Oman and was part of a plate boundary system which Şengör and Stock called the Ayyubid orogen, Ayyubids for short (Şengör and Stock 2014). Figure 14 shows the available age data on the timing of the generation and obduction of the Ayyubid ophiolites. It is clear that everywhere the ophiolites were obducted during the Turonian to Santonian interval. Indeed, in the Assyride foreland the sedimentation in the entire area ceased after the deposition of the Derdere Formation of Cenomanian–Turonian age and did not commence until the Karababa Formation of Santonian age. Both of these formations consist almost entirely of carbonates, dominantly limestones. This limestone section goes upwards into dominant clastics that maintain their dominance until the end of the Palaeocene. Eocene is dominated by widespread carbonates of the Midyat Group that reach into various levels of the Oligocene. In the east, in the Kilis-Gaziantep regions, the Oligocene consist entirely of carbonates; the section becomes progressively more lacunar eastward with the exception of the Siirt-Mardin area. In the Miocene, coarse clastics and sandstones dominate in foredeeps in front of the impinging Pötürge-Bitlis thrusts and sedimentation ceased almost entirely in the Pliocene to resume with the alluvium of the Plio-Pleistocene.
Fig. 14

The Ayyubid orogen and the timing of ophiolite obduction along it (after Şengör and Stock 2014, where the sources data are also inidicated) In each locality the top name is that of the ophiolite nappe as commonly known in the literature. The letters in parentheses next to them indicate the inferred speed of spreading (us ultra-slow, s slow, i intermediate, f fast, vf very fast). Beneath it is the name of the formation that seals the nappe contacts and below that the age of that formation. Below the wavy line signifying an unconformity are zircon U–Pb ages. When other ages are used this is indicated. Below that, across the thrust symbol, is the age of the youngest underlying rocks. Where this is a single formation, its name is given; where not, only the youngest age of a sequence is indicated. Ophiolite genesis and obduction from Oman to Cyprus all happened synchronously within the resolution of the isotopic and biostratigraphic data we now have. The obduction events were also synchronous with the shortening all along the Syrian Arc from eastern Libya to southeastern Turkey

After the rather ‘quiet’ ophiolite obduction during the late Cretaceous, orogenic activity resumed during the medial Eocene in the western part of the Assyrides. Yılmaz (1993) summarised the tectonic evolution in the entire Assyride belt, which still remains the best account. In the extreme west, in the Amanos Mountains, the sedimentation on the foreland stopped during the early Eocene and the entire section, together with the ophiolites, obducted previously, became stacked in south-vergent nappes. Yılmaz ascribed this to continental collision; a similar conclusion was reached by Hempton (1985) farther east in the area of the Pötürge Massif. However, farther east, foreland shelf sedimentation continued until the early Oligocene, but north of the carbonate areas, the clastics of the Upper Eocene–Oligocene Çüngüş Complex began deposition. Çüngüş appears as a subduction–accretion complex but has no oceanic rocks in it, indicating that a thick flysch package was being fed into the subduction zone in a diminishing basin that became choked with it (Yılmaz 1993).

Beginning with the early Miocene, the entire foreland sequence began to be shortened in sledrunner-type thrusts and flexural folds involving the entire sedimentary package from the Ediacaran to the Miocene. Sungurlu (1974) remains the best description of the foreland structures. The Şelmo Formation is a synorogenic deposit overlying the Lice and Kapıkaya flysches, which in the medial Miocene marked the onset of foredeep formation in front of the moving allochthons. Intertonguing with the Kapıkaya and the Lice are the carbonates of the Fırat Formation that mark quiescent sites and times during the progress of the foreland deformation.

It seems clear that by the end of the Oligocene the collision between Arabia and Eastern Anatolia was in full swing. Collision seems to have progressed from the west in the medial Eocene to the east in the Miocene. This is consistent with the Miocene closure of the Zagros Ocean farther east (McQuarrie et al. 2003).

The Menderes–Taurus Block

The Menderes–Taurus Block (Fig. 7) was a vast platform with a number of canals of various, but mainly late Palaozoic ages, which is a direct easterly continuation and termination of the African Promontory of Argand (1924), also called Apulia in the more recent literature. The reason that we do not call it simply Apulia is to avoid confusion, but our readers should be aware that that is what it really is: the easterly continuation of Apulia. This whole ensemble is similar to, but was generally shallower than, the present-day Lord Howe Rise in the Tasman Sea (minimum depth is about 1000 m: Willcox et al. 2001). Özgül (1976) was the first to subdivide this platform, a first-order tectonic unit, into second-order units using their structural boundaries and internal stratigraphies. Later, Okay (1984a) extended his classification northwards, north and east of the Menderes Massif. It is their classifications that we employ in this paper to characterise the Menderes–Taurus Block (Fig. 7).

The Geyik Dağı Unit

This is the ‘autochthon’ of the Taurus Mountains and is exposed in complete and semi-windows below the various nappes of different provenance which Özgül (1976) classified under the names Bozkır, Aladağ, Antalya and Alanya. Of these, the two former bodies were emplaced from the north, whereas the latter two from the south onto the platform. Sedimentation on this platform began already during the Cambrian on a Pan-African basement, the sedimentary rocks of which even include Archaean zircons (Kröner and Şengör 1990). In the west the section is entirely missing between the Ordovician and the Triassic, while in the east, all Palaeozoic Systems are represented albeit with lacunae between some of them. The Menderes Massif is nothing but a northerly, highly deformed and metamorphic continuation of the Geyik Dağı unit. The Mesozoic units are well represented, with no significant, widespread unconformities representing significant gaps up into the Cretaceous in its southern parts, but numerous areally restricted unconformities above the Triassic and below the Liassic and Dogger depending on where one is in the central and northern parts (Özgül 1976; Moix et al. 2008). Figure 15 columns 12–29 summarise the Mesozoic–Cainozoic stratigraphy of this unit in selected regions. For additional columns, see Özgül (1976) and Şengör and Yılmaz (1981).
Fig. 15

Stratigraphic columnar sections from selected areas characterising the Permian to Mesozoic geology of the Menderes–Taurus Block (Alçiçek et al. 2007; Gökten 1976; Göncüoğlu and Turhan 1984; Önal and Kaya 2007; Özdamar et al. 2013; Özgül and Kozlu 2002; Rimmelé et al. 2004; Selim and Yanık 2008; Şenel et al 1981; Turan 2010; Yağmurlu 1992)

The Bolkardağı Unit

This unit of Özgül (1976) includes Okay’s Tavşanlı and Afyon Zones (Fig. 7). In fact, Okay (1984b, see his fig. 1) had originally named his Afyon Zone as Afyon-Bolkardağ Zone. Okay distinguished the Tavşanlı and the Afyon zones on the basis of their metamorphism: Tavşanlı Zone as a whole underwent blueschist metamorphism during the late Cretaceous, whereas the Afyon Zone had only greenschist metamorphism and its age extends into the Eocene; but recently Candan et al. (2005) showed that this zone also had earlier undergone high-pressure/low-temperature metamorphism just like the Tavşanlı Zone. Our rationale for including these into Özgül’s Bolkardağ unit is that they all come together in the Bolkardağ Region and have similar stratigraphies. The Tavşanlı Zone is essentially the northernmost bit of the Menderes–Taurus Block that stuck its margin into the subduction zone the farthest under the obducting ophiolites (the Bozkır Nappe ensemble of Özgül 1976; Okay 2011). By contrast, the lowergrade Afyon Zone metamorphism, which overprinted the earlier blueschist metamorphism, occurred in the Eocene and was related to the further imbrication of the Menderes–Taurus Block that also formed and pushed the sub-ophiolitic Lycian nappes southward burying the Menderes below them and causing its metamorphism (Şengör et al. 1984; Okay 2001). However, Rimmelé et al.’s (2003a) discovery of high-pressure metamorphic rocks in the southern parts of the Menderes Massif and the presence of similar rocks in the lowermost Lycian Nappes (Collins and Robertson 1997) in the Bodrum Peninsula (Rimmelé et al. 2003b) show that its nappes were also buried deep enough in the subduction zone to the north to generate such rocks. Rimmelé et al. (2003a, b) also report structural observations indicating a top to the north fabrics associated with retrograde metamorphism of the high-pressure rocks. We interpret this as ‘regurgitation fabrics’ of the high-pressure rocks that formed along shear zones with normal dip-slip sense and not any retrocharriage of the Lycian Nappes onto the Menderes Massif as some have suggested (see Rimmelé et al. 2003a, b, for references). We would not be surprised if the entire Menderes Massif had descended to depths of blueschist metamorphism and then retrograded during the Main Menderes Metamorphism (MMM: Şengör et al. 1984), similar to the situation in the Afyon Zone (Candan et al. 2005).

Recently, Candan et al. (2016) reported a zone of Carboniferous metagranites with ages ranging from 331 to 315 Ma. They interpreted these as products of a south-dipping Palaeo-Tethyan subduction zone, an interpretation that corroborates the earlier model by Şengör et al. (1980) and Şengör and Yılmaz (1981), although Candan et al. (2016) think that the Palaeo-Tethyan subduction had ceased by the late Permian, a conclusion which contradicts the presence of granitoids in the Menderes (and also in the Rhodope–Pontide and Sakarya Continent as mentioned above) with ages between 246 and 235 Ma (Dannat, 1997; Koralay et al. 2001; Erdin Bozkurt, written communication, 25th May 2018).

Okay’s study of the metamorphism and the structure of the Menderes Massif (Okay 1989b, 2001) for the first time clarified the details of its internal structure and showed that the Massif consists of giant, south vergent nappes of the first genre, i.e., giant recumbent folds similar to the Pennine basement nappes in the Alps that formed mostly during the Eocene thrusting of the Lycian Nappes above it (Fig. 4a). This is compatible with the earlier conclusions of Akkök (1982, 1983) and Şengör et al. (1984), who for the first time pointed out that the Menderes Massif consisted of a pile of south-vergent nappes. The fig. 10a in Rimmelé et al. (2003a) and the one modified from them in Dora (2011, his fig. 9) are identical in principle to Akkök’s (1982) fig. 3, although they seem unaware of that. The objections raised against Okay’s model for the structure of the massif (e.g., Gessner et al. 2001; Okay 2002a) are so far without a firm foundation and all subsequent observations support Okay’s model (e.g., Gülmez et al. 2018, 2019; Şengör, unpublished observations in the southern part of the Massif northeast of Söke). Another common objection, that the folds in the massif are ‘north-vergent’ (Erdoğan and Güngör, 2004; Erdin Bozkurt, personal communications 2017 and 2018), stems from confusing early normal faulting associated with blueschist regurgitation, vergence with facing and, in the western and northern parts of the Massif from the presence of some genuine backfolding. All the so-called ‘north-vergent folds’ associated with thrusting that we have seen in the Menderes Massif (mainly in its southern parts) are south-facing, similar to the early confusion encountered in the Bolkardağ Massif under the Mersin ophiolites (Parlak et al. 1996; but see Parlak and Robertson 2004). The genuine north-facing folds are early and associated with blueschist regurgitation normal-sense shear zones. Anyone who would like to get an idea of what the Menderes structure looks like, would do well to read the paper by Searle et al. (2004) on the structure of the Al Hajar Mountains in Oman around the Saih Hatat Window.

A different source of confusion concerning the structural vergences of the Menderes Massif has been the timing of the onset of lithospheric stretching in western Turkey, which is usually considered to be medial to late Miocene (Şengör et al. 1985; Şengör 1987). Later publications claimed a much earlier date, namely late Oligocene (Bozkurt and Park 1994; Çemen et al. 2006; Ring et al. 2003; Ring and Collins 2005; Seyitoğlu and Scott 1996; Seyitoğlu et al. 2004, Seyitoğlu and Işık 2015; Diniz et al. 2006). However, the pre-medial to late Miocene normal faulting is not a result of total lithosphere stretching, but the formation of extrusion wedges similar to the ones in the Himalaya (Burchfiel et al. 1992) and farther west in the Aegean (Ring et al. 2010).

The Bozkır Unit

This entirely allochtohonus unit includes the large Lycian, Hadim and the Aladağ Nappes (Fig. 7) and comprises, structurally from top to bottom, ophiolites and ophiolitic mélanges of Jurassic to early Cretaceous ages obducted during the Turonian to Coniacian interval, turbidites shed from ophiolitic highlands, radiolarites, cherty pleagic limestones, cherty, dolomitic sandstones, siltstones and neritic limestones. The deep water facies spans an age from the Jurassic to the late Cretaceous in most places, and in southwestern Turkey the ages of the associated basalts go down even into the Carnian (Sayit et al. 2015), which is compatible with the Carnian radiolarians reported from the İzmir–Ankara–Erzincan suture in the middle Sakarya area (Tekin et al. 2002). Sayit et al. (2015) noted island arc affinities in the associated basalts and concluded from this observation that subduction had already commenced within the Neo-Tethys. However, given the position of the opening of the Neo-Tethys above a long-lived Palaeo-Tethyan subduction zone (see above), their observations may be explained by what Moores et al. (2000) called historical contingency of ophiolite compositions.

The Bozkır unit began its emplacement with the ophiolite obduction during the Turonian to Coniacian interval; the collision both with the Sakarya Continent and the Kırşehir Block during the late Palaeocene to medial Eocene interval, respectively, dislodged many of the ophiolites together with the continental margin units under them in the north and transported them southward in at least two major phases until the early Miocene, picking up and stacking under them ever distal packages as the nappes marched south. The Burdigalian in the Antalya area unconformably covers their thrust contacts in their last resting place.

Between the Menderes Massif metamorphics with nummulitic protoliths and the Lycian Nappes in the Kale-Tavas Basin, Chattian deposits unconformably cover the thrust contacts (Akgün and Sözbilir 2001) indicating that the nappes had reached that point by the Oligocene. By contrast, in the frontal regions, the movements continued into the Aquitanian, implying that the thrusts post-dating the Oligocene passed under the Kale-Tavas Basin and carried the entire Menderes Massif above them together with the Lycian Nappes. Only in the Denizli region, the fact that Middle and Upper Eocene is involved in the structure of the Honaz Dağı Anticline (Okay 1989b), may possibly indicate the presence of out-of-sequence thrusting here, perhaps postdating the Kale-Tavas Basin initiation. The presence of post-Burdigalian imbrication creating a pick-a-back basin in the Acıpayam Plain (Alçiçek and ten Veen 2008), just to the southeast of Okay’s Honaz Dağı Anticline, show the presence of out-of-sequence thrusting within the Lycian nappe edifice here. In the eastern part of the Taurus belt, under the Pozantı–Karsantı ophiolite nappes, Polat and Casey (1995) showed the existence of out-of-sequence thrusting during the imbrication of the Menderes–Taurus Block in the early Cainozoic.

Farther eastward, the geology of the Menderes–Taurus Block becomes very complicated owing to mainly latest Cretaceous-early Cainozoic subduction and concomitant strike-slip faulting that redistributed the original outlines of the Block. The Block ends eastward in three digitations all separated by ophiolitic sutures (Fig. 7): (1) Munzur, separated from the Malatya Digitation by the Ovacık–Kemaliye Suture, (2) Malatya separated from the Bitlis-Pötürge by the Göksun Suture and (3) Bitlis-Pötürge, separated from the Arabian Platform by the Bitlis Suture.

Özgül and Turşucu (1984) summarised the geology of the Munzur Mountains. A thick pile of carbonates going from the Norian–Rhaetian to the Turonian–Campanian are overthrust by what they call the Ovacık ophiolitic mélange containing ultramafic and mafic rocks and pelagic sedimentary rocks with ages ranging from the Permian to the Turonian. The ophiolitic mélange was thrust over the carbonate platform in pre-Mastrichtian time from the north and was imbricated with it. In medial to late Eocene time the Munzur was separated from the Malatya metamorphics to the south (called the Keban Unit by Özgül and Turşucu 1984) by narrow fault zones that now appear as thrusts. But, peculiarly, the map view of these thrusts is fairly linear to gently curved and their strikes are, mostly, concave to the south (see Bilgiç 2002; Tarhan 2002). Along the strike to southwest, south of Kangal, a fairly tight arc of Mesozoic carbonates surrounds the equivalents of the Keban metamorphics and these carbonates are delimited to the north by narrow screens of ophiolites and ophiolitic mélange. This line of ophiolites and mélanges plus Eocene flysch and volcanic rocks we called the ‘Ovacık-Kemaliye Suture’ above. We shall return to a discussion of its real nature below, after having described another such suture below, that of Göksun between the Malatya Metamorphics and the Bitlis–Pötürge Massifs.

Yılmaz et al. (1993) mapped the central region of the Malatya Metamorphics and established that these rocks are separated from the Pötürge–Bitlis Massifs by serpentinite screens, although they were sure that the Malatya Metamorphics were nothing but a westerly continuation of the Pötürge-Bitlis Massifs. The only substantial arc rocks are the Helete volcanic rocks of late Maastrichtian to early Eocene age plus the two granitoids intruding the Göksun Ophiolite in the west (88–85 Ma: Parlak 2006) and the Baskil granitoid intruding the Kömürhan ophiolite in the east (85–82 Ma: Rızaoğlu et al. 2009).

The internal stratigraphy of the Malatya Metamorphics/Pötürge/Bitlis massifs is essentially the same as the Munzur Platform, with the difference that they are metamorphic. Ustaömer et al. (2009a, b) discovered 545 to 531 Ma ages on the Mutki granite pluton and associated intrusions in the Bitlis Massif, which they interpreted to be result of arc magmatism. This is very similar to the derik volcanic rocks from the Arabian foreland to the south and probably indicates a former unity of the two tectonic entities. Zircon LA-ICP-MS data give crystallization ages of 581.4 ± 3.5 Ma (n = 7) and 559.2 ± 3.2 Ma (n = 3) for the early- and late-stage andesitic rocks, as well as ages of 569.6 ± 1.6 Ma (n = 17), 571.6 ± 1.9 Ma (n = 18), 575.4 ± 4.3 Ma (n = 6) for the rhyolites (Gürsü et al. 2015). Blueschists in the Bitlis Massif recently gave 78 to 74 Ma, i.e., Campanian (Oberhänsli et al. 2012). All these massifs carry ophiolite nappes on top of them and in at least two places (Göksun and Kömürhan) the ophiolites seem to be ensimatic arcs, just like the rest of the large Taurus ophiolite nappes. We thus regard these massifs as nothing more than the easterly continuation of the Menderes–Taurus Block here overwhelmed entirely by ophiolites owing to the original tapering out of the Menderes–Taurus Block eastward. Moreover, their geology is extremely similar to that of the Kırşehir Massif and it seems that the Kırşehir was their easternmost member, now displaced from its original place by latest Cretaceous–Eocene left-lateral strike-slip faulting. The horizontal imbrication of the Malatya and Pötürge-Bitlis occurred at the same time along right-lateral strike-slip faults and east of the junction where the left- and the right-lateral faults meet, an extensional boundary caused the opening of the Maden Basin that was in existence between the Maastrichtian and the medial Eocene (Aktas and Robertson 1984; Yılmaz 1993; Yılmaz et al. 1993).

A peculiar observation concerning the Menderes–Taurus Block is the presence of Lower Carboniferous pelagic sedimentary rocks in places intercalated with submarine volcanic rocks and tuffs that reach from one end of the Block to the other (see Fig. 2 under ‘Lycaonian Suture’). The Karaburun Mélange of Robertson and Pickett (2000) is the westernmost representative of these important rocks. The blocks in what they call a mélange consist of competent blocks, disrupted thrust sheets of neritic and pelagic limestones, black ribbon radiolarites and pillow lavas. The undated matrix consists of a sheared shale–sandstone association and rare conglomerates. The whole mélange is underlain by a lower unit of greenschist grade shales and sandstones. Although Robertson and Pickett interpret this ensemble as a subduction mélange, there are no ultramafics or gabbros in it. More significantly, they seem unaware that Caridroit et al. (1997) found a most remarkable association of fossils in this mixture. In a slightly manganiferous micrite with radiolarites they identified, in addition, conodonts, foraminifera, ostracoda and even vertebrate microremains all of which can be dated to the uppermost Lower Tournaisian (Hastarian). They specify the depositional milieu as one of slow sedimentation in deep water and suggest that these rocks belonged once to a Palaeo-Tethys, a conclusion that Robertson and Pickett (2000) later independently echoed.

But observation of Lower Carboniferous deep water rocks associated with volcanic rocks is not confined to Karaburun. Kozur et al. (1998) reported similar rocks from the İncirbeleni Formation in the Tavas Nappe of the western Lycian Nappes. The Tavas is the lowermost of the Lycian Nappes. Since these nappes were stripped off the northern parts of the Menderes–Taurus Block, the Carboniferous deep-sea rocks must have been somehow associated with those described by Caridroit et al. (1997) and Robertson and Pickett (2000). They consist of turbidites, limestone blocks, debris flows and black radiolarites with mafic volcanic rocks exactly as in Karaburun, but they are younger, Middle to Lower Visean.

Farther east, between Kütahya and Konya, in Okay’s (1984b) Afyon Zone (within Özgül’s 1976, Bolkardağ Unit), metaclastic, recrystallised limestone, metavolcanic and metamicrogabbro blocks float in a matrix of metasandstones, metasiltstones, metaconglomerates and black cherts, the whole being termed the Halıcı Group (Özcan et al. 1988; Göncüoğlu et al. 2001; Robertson and Ustaömer 2009). Robertson and Ustaömer (2009) named these rocks the ‘Konya Complex’. The ages of the blocks range from the Silurian to Carboniferous, but Özcan et al. (1988) indicate the boundary between the Halıcı group and the overlying Permian coarse clastics as a paraconformity. Moreover, the Halıcı itself is shown to sit on the underlying Bozdağ neritic limestones with an unconformity (Özcan et al. 1988). Therefore, no deformation seems to have occurred between the Carboniferous and the Permian and Scythian red to violate shales, sandstones, conglomerates and thinly bedded conodont and foraminifera bearing limestones cover the underlying rocks with an unconformity.

The Bolkardağ unit is truncated to the east by the Ecemiş Fault and is thrown northward to the Aladağ Range (the classical Antitaurus), where the Menderes–Taurus Block is imbricated by south-vergent thrusts south of the Inner Tauride Ocean suture, as one sees farther west and as outlined above. In this nappe pile, near the Nohutluk Hill in a slice above an ophiolitic mélange nappe, Aksay (1980) reported deep water Lower Carboniferous rocks. These are widespread cherty limestones whereby the limestones seem turbiditic and the cherts are between them. The cherts appear as thin layers of 1 to 10 cm thickness, rarely as small nodules. The limestone beds are thicker, varying between 10 to 30 cm. The limestones contain echinoderm debris and foraminifers. Altered tuff and volcaniclastic debris covering some limestone bed surfaces indicate coeval local vulcanicity. There are local bituminous horizons.

These are the so far reported deep water early Carboniferous occurrences in Turkey. Göncüoğlu et al. (2007) summarized much of the data concerning the Lower Carboniferous from the central and eastern parts of the Menderes–Taurus Block and indicated that the volcanic rocks within them give a back-arc geochemical signature. The same conclusions were repeated by Robertson and Ustaömer (2009) for the central and western parts, although they interpreted the geochemical signature of the volcanic rocks simply as ‘subduction-related. All the localities indicate strong Carboniferous subsidence accompanied by both mafic and felsic volcanism, turbidite sedimentation and in places they pass without a major unconformity into the Permian. In all cases the Triassic is either thrust onto them or it sits on them with disconformity or unconformity. Although Kozur et al. (1998) compare these rocks with the ‘Hercynian flysch’ in southern Europe, they are not comparable, mainly because of the volcanic rocks they contain. Especially Özcan et al.’s (1988) observation that some mafic dykes cut also the Permian indicates ongoing extension sometime during or after the Permian and some of the tectonic contacts of the Triassic with these rocks may well have been normal faults possibly related to the opening of the Neo-Tethys. Also, two metarhyolite samples yielding weighted means 230 ± 2 Ma and 229 ± 2 Ma (Konya Ilgın area) having the same trace element geochemistry with the A-type granites may corroborate the extension here showing a bimodal vulcanicity. However, phengites in the same metarhyolites gave ~ 63 Ma (HP mineral in greenschist rocks). This is possibly related to the onset of imbrication of the Menderes–Taurus Block.

The Carboniferous and perhaps even the Permian rocks indicate an extensional basin before and partly coeval with the onset of major Carboniferous magmatism in the Menderes–Taurus Block. We interpret these Carboniferous rocks as contents of a marginal basin of the Rocas Verdes type in the southern Chilean Andes (Dalziel et al. 1974; de Wit and Stern 1981; Stern and de Wit 2003; see Dalziel and Dott 2015, for a summary of the investigations on this remarkable fossil marginal basin) that opened disrupting a continental margin arc. It seems to have opened perhaps in a ‘pre-arc spreading’ stage, but more likely in a lull in magmatism of the northern Gondwana-Land magmatic arc (see Buggisch et al. 1974) induced by the opening of the marginal basin itself, perhaps not dissimilar to the situation now seen in the Okinawa Trough (e.g., Lee et al. 1980). This marginal basin probably closed during the unfolding of the Cimmeride events to the north, but detailed structural observations tied to the stratigraphy are lacking everywhere to pin down the timing and the closure geometry precisely.

Şengör and Atayman (2009) pointed out that during the Carboniferous and Permian the entire northern margin of Gondwana-Land was a Pacific-type continental margin with an extensional magmatic arc over a long-lived, south-dipping subduction zone. Behind this arc, various extensional basins opened from Thailand to Turkey. In places, the opening of these basins turned off or diminished the arc magmatism, much like the history one knows from the western Pacific margin today. The Carboniferous marginal basin in the Menderes–Taurus Block is only a link in this long chain of basins. We here name it the Lycaonian Marginal Basin, because its longest observable outcrops occur between Kütahya and Konya, along the southwestern boundary of the ancient Lycaonian province.

Synopsis of the Menderes–Taurus Block

The Menderes–Taurus Block is the largest first-order tectonic unit of Turkey. It is the easterly continuation and termination of the African Promontory of Argand and until the opening of the Eastern Mediterranean was a part of northern Gondwana-Land. Throughout the Mesozoic it was a submarine platform, much like the present-day Lord Howe Rise in the Tasman Sea, although not as deep. Just like the Lord Howe Rise it too was fluted by deep furrows of diverse ages from the Carboniferous to the Jurassic, which were probably back-arc basins above the south-dipping Palaeo-Tethyan subduction zone. For much of its history it was a site of neritic carbonate sedimentation although its margins had pelagic, deeper water deposits that have not been preserved everywhere. Its basement is Pan-African like much of North Africa and Arabia and it experienced shortening deformations during the late Triassic and earliest Jurassic as a consequence of the destruction of the Karakaya Basin to its north. Neo-Tethys opened immediately afterwards and separated it from the units to the north. In the late Cretaceous a vast ophiolite nappe was obducted onto its northern margin, which was coeval with the Zagros and the Omani ophiolites and the obduction may very well have been continuous among them. We have used this vast nappe, which we called, following Ali Koçyiğit’s apposite designation, the Anatolian Nappe, which is a part of Özgül’s (1976) Bozkır Unit, as Ariadne’s thread to identify the now dispersed fragments of the Menderes–Taurus Block. Around the Gulf of Antalya there was also ophiolite obduction northward onto the platform accompanied by the emplacement of the Alanya Nappes. The Menderes Taurus Block collided with the Sakarya Continent and the Kırşehir Block during the medial Eocene and became internally imbricated with south-vergent thrusts. This imbrication lasted into the early Miocene. Another result of the imbrication was the generation of the metamorphic axis of the Menderes–Taurus Block going from the Menderes Massif to Bolkardağ and beyond into the Malatya and the Bitlis-Pötürge digitations. A short-lived magmatic arc, from the late Cretaceous to the late Eocene existed only in its southeastern extremity.

The Kırşehir Massif

The Kırşehir Massif is the dominant first-order tectonic unit of Central Anatolia and is almost entirely surrounded by ophiolites (Figs. 2 and 7); ophiolite klippen are found on top of it as well (Floyd et al. 2000). It has a fairly uniform stratigraphy with a Gondwanian basement consisting of biotite–muscovite–sillimanite gneisses giving Eburnean (2 Ga U/Pb ages) and very late Pan-African (460 Ma: Middle Ordovician) ages (Göncüoğlu 1982, 1986 and written communication, 1986) surmounted by a thick metamorphic pile of what used to be platform sedimentary rocks beginning with clean and dirty limestones alternating with pelitic and semipelitic layers (Fig. 16, cols. 30 and 31). This section is apparently concordantly followed by a fairly pure marble pile. The discovery of the early Palaeozoic tabulate coral Heliolites paekelmannopora and possible fragments of the graptolite genus Retiolites dated at least a part of this sequence to the Silurian (Kocak and Leake 1994), but no further sedimentation ages have so far become available from these rocks, whose metamorphism ranges from granulite to greenschist grade (Seymen 1981a, b, 1985) and the age of the Alpide metamorphism of the entire pile has been established at around 91 to 84 Ma (Whitney et al. 2003 and Whitney and Hamilton 2004). These rocks are uniformly cut by late Cretaceous gabbroic, granodioritic and granitic arc plutons. Lithological correlation with the nearby Sakarya Continent and especially the Menderes–Taurus Block has led many to surmise a similar Palaeozoic-Mesozoic age for the post-Precambrian metamorphic sequence (Necdet Özgül 1976, 1981 and personal communication, 1980; Lefebvre et al. 2011). In fact, that is why Şengör and Yılmaz (1981) erroneously considered it simply a part of a unified Anatolide/Tauride Platform. As Şengör et al. (1982), Görür et al. (1984) and Whitney and Hamilton (2004) have since emphasised, the presence of a unified Anatolide/Tauride Platform can no longer be maintained and, as urged above, the name should be dropped from usage.
Fig. 16

Stratigraphic columnar sections from selected areas characterising the Permian to Mesozoic geology of the Kırşehir massif, the Sakarya Continent and North Dobrudja (Dinu et al. 2005; Dirik et al. 1999; Rojay 1995)

The best structural data from this metamorphic sequence have been gathered by Seymen (1981a) from around Kaman, who documented an original west-vergent fold set later refolded around east-vergent folds. Most geologists who drew cross-sections across the Massif depicted only this second folding episode, although Göncüoğlu (1977) also recorded the west-vergent folds in the extreme south of the Kırşehir Massif, in the sector known as the Niğde Massif. All of this folding occurred before the intrusion of the arc plutonic rocks (Seymen 1981a; Göncüoğlu 1977). In the north, in the Akdağmadeni metamorphics, Pollak (1958) and Yılmaz and Özer (1985) recorded only southeast-vergent folds and thrusts.

The arc magmatics have ages ranging from 95 to 72 Ma with cooling ages reaching into the very earliest Palaeocene (65 Ma) (e.g., Göncüoğlu 1982, 1986; Boztuğ 2000; Köksal et al. 2004; Tatar and Boztuğ 2005; Boztuğ and Joncheere 2007; Boztuğ et al. 2007, 2009a, b). Erkan and Ataman (1981) thought that the contact metamorphic ages of the metamorphic rocks around the plutons gave cooling ages of 74 to 69 Ma. Later measurements corroborate their results (Göncüoğlu 1986; Alpaslan et al. 1996), although Boztuğ (2000) thought that this late metamorphism was related to the collision of the Kırşehir Block with the ‘Pontides’. As neither the collision was that early, nor the intrusions which he thought synchronous with the metamorphism were so uniformly late, we cannot agree with that interpretation. The magmatic arc defines an asymmetric horseshoe shape opening to the east, the axis of which is defined by the Himmetdede Thrust Zone discovered by Erguvanlı (1959, 1961a, b) and extended by Oktay (1981). Lefebvre et al. (2013) called the westernmost part of this feature the Savcılı Thrust Zone, following Oktay (1981, pp. 150–151), who thought he was the first to discover it and hence named it after the locality where he saw it. We prefer to use Himmetdede to respect Erguvanlı’s priority. On the basis of these observations Şengör and Natal’in (1996) argued that originally the Kırşehir Massif was an arc with a curvature much less than what it is today. In fact, Oktay (1981) showed that the tightening of the horseshoe continued well into the Neogene. Recently, Gülyüz et al. (2012) mapped another east–west striking thrust north of the Himmetdede Zone bounding the Çiçekdağı Basin to the north and overriding the Upper Eocene sedimentary rocks; later Lefebvre et al. (2013) corroborated the bending hypothesis of Sengör and Natal’in (1996) using palaeomagnetism; from their data it seems clear that the acquisition of the horseshoe shape occurred entirely in the Cainozoic as a result of being caught between the approaching jaws of the Sakarya Continent and the Menderes-Taurus Block, mainly in the later Eocene to early Oligocene. Göğüş et al. (2017) pointed out that this constriction provided a ready explanation for the loss of much of the lithospheric mantle beneath the Kırşehir Massif and its Neogene uplift.

There are two groups of metamafic and ultramafic rocks in the Kırşehir Massif (Floyd et al. 2000). A group of concordant metamafic rocks within the metamorphic basement consists of fine-grained amphibolites and hornblende schists the ages of which are unknown but believed to be Triassic, Jurassic or Cretaceous. Floyd et al. (2000) prefer a Triassic age and, if true, these rocks may be correlated with the Carnian mafic rocks known from the Bitlis Massif. Their chemistry is dominated by within-plate compositions indicating rifting similar to the case in Bitlis. By contrast, the metamorphosed ophiolites, disrupted to various degrees into ophirags, overlie the Massif tectonically and they contain a complete ophiolite suite when taken as a set. They seem to have formed during the 90 to 85 Ma interval (Turonian) and obducted shortly thereafter. They show the same spread of environments ranging from MORB to supra-subduction zone settings as all other Cretaceous Taurus ophiolites, but they are somewhat younger. Their outcrops are concentrated near the axis of the Kırşehir magmatic arc and they were emplaced when the arc was still active. They are thrust over the older members of the arc and the younger members cross-cut them (Floyd et al. 2000). Unfortunately, no structural data have so far been reported to indicate their direction of emplacement.

Later, the outer margins of the horseshoe was subjected to extension and a number of extensional metamorphic core complexes have formed with normal faults dipping away from the Massif, from about Kaman (Lefebvre et al. 2011) to the Niğde Massif (Whitney and Dilek 1997, 1998; Fayon et al. 2001; Gautier et al. 2002) in an environment perhaps not dissimilar to the extensional area seen in the Kai Islands in Indonesia where an arc/continent collision is now in progress (Charlton et al. 1991; Pownall et al. 2016).

Synopsis of the Kırşehir Massif

The Kırşehir Massif is an ensialic magmatic arc with a stratigraphy very similar to that of the Menderes–Taurus Block. It was rifted during the Triassic from the rest of Gondwana-Land and subduction below it began during the late Cretaceous, lasting into the Palaeocene to early Eocene. Ophiolites of late Cretaceous age were obducted onto it shortly after they were generated, but we do not know in which direction. Before the Eocene, when the collision was just about to commence, extension characterised the forearc area. Probably beginning already in the latest Cretaceous the arc was doubled back on itself creating an east-facing horseshoe shape and the tightening of the horseshoe continued into the Neogene leading to lithospheric drip and consequent uplift.

The Sakarya Continent

West of Ankara

The Sakarya Continent (Fig. 7) was identified by Şengör and Yılmaz (1981) as that strip of terrain between the Intra-Pontide Suture in the north and the İzmir–Ankara–Erzincan suture in the south. They terminated it at the island of Lesbos in the west but suggested that it may continue into the Paikon Ridge in the Vardar Zone. Later, Şengör et al. (1984) adopted this suggestion. The recent discovery of extensive calc-alkalic and island arc tholeiitic rocks of late Jurassic–early Cretaceous age in the Paikon ridge (Bebien et al. 1994; Bonev et al. 2015b) has greatly fortified this interpretation as similar rocks have since been reported also from the Sakarya Continent (e.g., Genç and Tüysüz 2010). Genç and Tüysüz (2010) interpreted their results as showing an extensional episode atop the dying phases or a recently deactivated subduction zone. Indeed, Pe-Piper and Piper (1991) similarly interpreted the thick andesites of late Triassic–early Jurassic ages from the Pindos Zone and tied the subduction signatures to the south-dipping Palaeo-Tethyan subduction zone beneath the rifting Pindos Basin. Figure 16 shows some characteristic stratigraphic columnar sections from the Sakarya Continent (cols. 32 and 33).

The most recent description of the entire Sakarya Continent was given by Yılmaz et al. (1997b) and Akbayram et al. (2016a) and we largely follow their lead in this paper adding more data wherever appropriate. The Sakarya Continent has a polymetamorphic sialic basement that yielded zircon populations with age clusterings around 2 Ga (Eburnean), 1050–950 (early Pan-African) and 550–750 (late Pan-African). Two smaller groupings are 1850 and 800 Ma (late Pan-African) (Ustaömer et al. 2012). 1850 Ma age grouping is similar to the interval 2600 to 1600 Ma from northeast Africa reported by Harms et al. (1990). All these ages indicate a north Gondwana-Land provenance and tie the basement of the Sakarya Continent firmly to Gondwana-Land as also emphasised by Ustaömer et al. (2012). A zircon measured by single zircon evaporation technique from the gneisses of the Uludağ basement in Bursa had yielded 384 Ma (Alfred Kröner and A. M. C. Şengör unpublished data). Sunal (2012), Aysal et al. (2012) and Ustaömer et al. (2012) documented, respectively, the presence of granites and granitoids with an age range of 401 to 389 and 324 to 319 Ma cutting this basement. The earlier interval is similar to the ages of the Middle Devonian volcaniclastics from the Menderes-Taurus Block (Buggisch et al. 1974) and the latter group of intrusion ages are very similar to those reported by Candan et al. (2016) from the Afyon Zone, i.e., also from the Menderes–Taurus Block and support the idea that these two entities were probably very close to one another, if not actually contiguous in the northern margin of Gondwana-Land before the Triassic.

On top of this basement a Permian transgression first brought arkoses which rapidly pass into richly fossiliferous (rich in fusulinids) neritic Permian limestones. The biofacies of these rocks is the same as those of the Menderes–Taurus Block and indicates the presence, in the words of Altiner et al. (2000) that ‘one single carbonate platform …, comprising nearly all Upper Permian marine sedimentary rocks exposed in Turkey’ (their p. 92). But then, Altiner et al. (2000) deny the validity of the observations reported by Yılmaz et al. (1997b) who argue for a transgressive contact between the Permian rocks and the Sakarya Basement. Altiner et al. (2000) claim that all the Permian in the Sakarya Continent is in the form of blocks or at least allochthonous slices. This claim is not true: In the core of the Uludağ Massif white bedded marbles, although internally sheared, rest on the basement and are overthrust by the rocks of the Karakaya Basin that contain unmetamorphosed Permian knockers (Ketin 1985). Farther north, the Değirmenbayırı Marble of Aksoy (1995) on the Marmara Island (classical Proconessos) sit on the Ortaburun Micaschists and the Abasırtı Calcschists. All the contacts seem transitional and were probably originally conformable. Although they have no ages assigned to them and Aksoy thought them pre-Permian, we think that the marbles probably represent the Permian cover of the Sakarya Continent as they are overthrust by the Karakaya-equivalent Erdek Complex, which is in turn covered unconformably by the protoliths of the classical Marmara Island marbles (Aksoy 1995), which we equate with the Mesozoic carbonate cover of the Sakarya Continent as they are in turn overthrust from the north by the ophiolitic mélange of the Intra-Pontide Suture (the Saraylar Complex of Aksoy 1995). Timur Ustaömer and his colleagues recently measured detrital zircons from the metaclastics under the Değirmenbayırı Marble and found that they were supplied from a late Palaeozoic orogen (Timur Ustaömer, personal communication, 9th June, 2018) an observation that supports our inference. Dubertret et al. (1973) noted that the ‘Fusulina limestones extend as a continuous strip from Zeytindağ (long. 27°05ʹE, 38°58ʹN) towards the north-northeast, to Bergama-Kınık, İvrindi, Balya, Danişment and Manyas. … They rest on the molasse-like Upper Palaeozoic. A basal conglomerate is seen in the road cuttings 5 km northeast of Ivrindi (long. 27°30ʹE, 39°34ʹN), close to the fork of the roads to Balıkesir and to Balya; the pebbles are elongated’ (p. 76). Whereas the outcrop is not continuous and much of the Permian is in blocks in the Triassic (see Akyürek and Soysal 1981; Altiner et al. 2000) their observation of the underlying molasse-like rocks below the Permian carbonates is in agreement with other observations on the Sakarya Continent. In the interpretation of Altiner et al. (2000), the basement of the Sakarya carrying the Permian cover (the cover part they deny) had rifted from Gondwana-Land and moved north opening behind it the basin in which the Karakaya rocks were deposited into which the Gondwana-Land Permian fell as blocks. But this is precisely what Şengör and Yılmaz (1981) and Yılmaz (1981) claimed. Altıner et al.’s problem is that they do not see the Permian cover (but do see the numerous blocks that they rightly derive from Gondwana-Land). From their descriptions we conclude that in the beginning of the Permian the Sakarya continent was still a part of Gondwana-Land, most likely the northern part of the future Menderes-Taurus Block, but in the late Permian a part rifted and drifted north above the south-dipping Palaeo-Tethyan subduction zone opening the Karakaya marginal basin.

The southernmost outcrops of the Karakaya rocks are seen in three main strips from the Gulf of Dikili to the district of Mecidiye (long. 27°39ʹE, 39°27ʹN), a part of the train of outcrops mentioned by Dubertret et al. (1973). Here Akyürek and Soysal (1980–1981) mapped Lower Triassic rocks containing numerous Upper Permian limestone blocks and subordinate mafic volcanic rocks and correlated these occurrences with the Karakaya rocks farther north. We agree with this correlation, but note that the ‘blocks’ are not randomly distributed in the Triassic schists: they seem to concentrate in three ‘belts’ and the largest chunks uniformly strike northeast-southwest. Near the district of Çamoba (long. 27°23ʹE, 39°13ʹN) Akyürek and Soysal (1980–1981) thought what they considered an autochthonous Permian showing regular bedding and all the members of what they call the Çamoba Formation in primary contacts. However, they noted that the fact that the Permian limestones show little deformation, whereas the Triassic is highly folded and thought this might imply that even this outcrop is actually a block. We disagree, because some six kilometres to the northeast, the Permian appears in long drawn-out folds in the Triassic. The structural disharmony is most likely a result of the lithological difference.

Despite the presence of Karakaya rocks here, we do not include these southern outcrops in the suture, because here it seems that the structure indicates more of an extensional basin margin that later got smashed than the oceanic basin itself. The only igneous rocks reported here are mafic volcanic rocks and those not in abundance. In only four places mappable outcrops on a scale of 1/25,000 were observed. Akyürek and Soysal (1980–1981) also made the important observation that here the Middle Triassic sits with angular unconformity on older, folded rocks. It is regrettable that their study contains no structural observations. Nevertheless, it seems as if the closure of the Karakaya Basin proceeded from south to north, corroborating the northerly vergence observations farther north.

After the Karakaya Basin closed, the Bayırköy Sandstones unconformably covered it, sealing all the thrust contacts. This sandstone unit is disconformably overlain by a neritic limestone succession, with Rosso Ammonitico layers, that reaches into the Upper Jurassic (called, not inappropriately, ‘Alpenkalkstein’ already in 1854 by the German geographer Friedrich Heinzelmann, p. 308). It is followed by the pelagic, thinly-bedded Soğukçam Limestone, displaying a Maiolica facies. Although Yılmaz et al. (1997b) inferred a gap between the Coniacian and the Santonian which they ascribed to uplift and erosion, later work by Özcan et al. (2012) closed the gap and proved continuous sedimentation. They further showed that the Jurassic limestones are unconformably overlain by the Middle Turonian and younger sediments in a Scaglia Rossa facies. This unconformity was a product of the deformation earlier reported by Genç and Tüysüz (2010). Between the Campanian and Maastrichtian, orogenic deformation is heralded by the onset of flysch sedimentation (Yücel and Soner 1991; Catanzariti et al. 2013; Akbayram et al. 2016a). The flysch basin shallowed during the Palaeocene leading to the deposition first of Lower to Middle Palaeocene shallow-marine limestones (Bargu 1982; Genç 1986; Bargu and Sakınç 1989/1990; Örçen 1992; Akbayram et al. 2016a) and the Middle to Upper Palaeocene continental red beds (Saner 1980; fig. 3). The Lower Eocene clastic sedimentary rocks lie disconformably over the Palaeocene red beds (Akbayram et al. 2016a) and unconformably over the Upper Cretaceous rocks and older units (Özcan et al. 2012). The Eocene rocks are formed from a thick and heterogeneous pile of conglomerate, sandstone, limestone, and volcanic rocks of Upper Ypresian–Lower Lutetian age (Özcan et al. 2012). North–south shortening resulting from the collision of the Sakarya Continent with the Rhodope–Pontide Fragment to the north finally led to the folding and thrusting with a south vergence of this entire package of rocks almost immediately followed by right-lateral strike-slip faulting displacing parts of the suture for some 365 km westward to the Biga Peninsula (Akbayram et al. 2016a).

East of Ankara

The Tokat Massif represents the eastern part of the Sakarya Continent. Yılmaz et al. (1997a, 1998) described the geology of its eastern part. The basement there consists mostly of the Karakaya equivalents that are unconformably covered by Liassic clastic rocks (equivalents of the Bayırköy Sandstones), which rapidly pass into neritic limestones with Rosso Ammonitico layers, equivalent of the Mudurnu Formation. These rocks are then conformably succeeded by the Maiolica Soğukçam Limestones. The Liassic cover is not everywhere present and in the western part of the Massif, Seymen (1991) reported Upper Jurassic resting with a basal conglomerate on the metamorphic rocks.

In the pre-Liassic basement of the Tokat Massif, the Amasya Group of Yılmaz et al. (1997a) represents a continental assemblage beginning at the base with quartzites and then shales, siltstones and limestones passing upward to more quartzites then greywackes, conglomerates with recrystallised limestone layers, shales and slates, then fossiliferous limestones with scattered volcanic horizons and finally more metapsammites. The entire section has been metamorphosed in greenschist facies and is unconformably covered with Lias. From the fossiliferous limestones, a Silurian fauna was reported. Farther west, Seymen (1991) thought that the Permian was unconformable on earlier metamorphics including marble blocks floating in a garnetiferous green and blueschists, phyllites and calcschists and serpentinites, but he could not actually see such an unconformity in the field. It is highly likely that what Seymen (1991) considered a pre-Permian metamorphic complex corresponds with a higher metamorphic part of the Karakaya Complex.

The Sakarya Continent ends against the Intra-Pontide suture extending from Lâdik in the west to beyond Reşadiye in the east. Along this suture zone ophiolites and ophiolitic mélanges of late Cretaceous age have been reported from Mamo Dağ south of Niksar, Köse Dağları south of Reşadiye and Karaçam Dağı (Hakyemez and Papak 2002). These occurrences appear in the form of nappes and imbricate slices atop the Tokat Massif below Eocene clastic rocks in places with volcanogenic components (Yılmaz et al. 1997a; Hakyemez and Papak 2002). The most significant observations on this suture zone were made by Seymen (1975) who documented the presence of a south-facing Atlantic-type continental margin during the Jurassic and the early Cretaceous on the Rhodope–Pontide Fragment here, which Görür et al. (1983) were able to corroborate, thus cleanly demarcating the Sakarya Continent against the Rhodope–Pontide Fragment. We shall see below that even the Precambrian zircon populations are not identical in their basements. There is thus little justification in continuing the Sakarya Continent into the Eastern Pontides in the form of a purported ‘Sakarya Zone’, as is done by some on the basis of the similarities in their Mesozoic geology; let us note that corresponding similarities also exist in the Western Pontides.

Synopsis of the Sakarya Continent

The Sakarya Continent is a continental strip that extends from the Paikon Ridge in Greece to about Reşadiye in Turkey in the east between the Intra-Pontide and the İzmir–Ankara–Erzincan sutures. It was a part of the northern Gondwana-Land margin until the Triassic (latest Permian?) rifting. A Karakaya marginal basin opened above the south-dipping Palaeo-Tethyan subduction zone to the north within the Sakarya Continent during the latest Permian to earliest Triassic interval and began closing with a northerly vergence in the medial Triassic in the south. Its closure was complete by the Lias. The Sakarya Continent lost its independence by collisions, during the early (in the south) and medial (in the north) Eocene collisions both terminating north-dipping subduction. The southern subduction zone created a medial Cretaceous to Eocene ensialic magmatic arc on top of it.

The Rhodope–Pontide Fragment

The Rhodop-Pontide Fragment (Fig. 7) is the Turkish part of what Burchfiel (1980) earlier called the Rhodope Fragment and it continues into the Sanandaj-Sirjan Zone in Iran (Şengör 1990a, b, c). It is by far the most complicated and the most contentious part of Turkish geology. Figure 17 shows some representative stratigraphic columnar sections. We also illustrate a column from the Moesian Platform (col. 35) as it is a continuation of the İstanbul Fragment. As a whole, the Rhodope–Pontide Fragment was derived from the northern margin of Gondwana-Land, but not in one piece. In Turkey, it has three distinct parts: these are, from west to east, the Strandja Zone, the İstanbul Zone and the Central and Eastern Pontides (these last two we consider a single unit for the purpose of classifying the first-order tectonic units during the Alpide evolution). They correspond to the Strandja Nappe, İstanbul Nappe and the Küre and Bayburt nappes of Şengör et al. (1980, 1984) and Şengör and Yılmaz (1981). Recent studies (Ülgen et al. 2018) showed that the İstanbul Nappe indeed overlies tectonically the Strandja Zone and originally lay to the south of it as interpreted by Şengör et al. (1980, 1984) and Şengör and Yılmaz (1981) and not strike-slipped against it as once been surmised by Okay et al. (1994). At its eastern end it lies above the oceanic rocks of the Palaeo-Tethys (Şengör et al. 1980; Şengör and Ylmaz 1981; Ustaömer and Robertson 1993, 1994, 1995), which Şengör et al. (1980) had originally gathered under the designation Küre Nappe. Figure 17 displays some characteristic stratigraphic columnar sections from the Rhodope–Pontide Fragment within the borders of Turkey.
Fig. 17

Stratigraphic columnar sections from selected areas characterising the Permian to Mesozoic geology of the Rhodope–Pontide Fragment (Dokuz et al. 2017; Gürsoy et al. 1993; Sunal and Tüysüz 2002; Tari et al. 2013; Toprak and Akyazı et al. 2017;Yılmaz 1992)

The Strandja Zone

Since the Strandja Zone is described above, we deal below only with the İstanbul Zone and the Eastern Pontides.

The İstanbul Zone

The İstanbul Zone is a fragment of the Moesian Platform (Okay et al. 1994). Its geology has been reviewed in some detail in recent years (Görür et al. 1997; Özgül et al. 2005, 2009; Şengör and Özgül 2010; Özgül 2011, 2012; Şengör 2011; Lom et al. 2016) and in what follows we summarise the contents of those reviews; for references the reader should consult them.

The İstanbul Zone is made up, at the base, of Neoproterozoic (Pan-African) middle to high-grade crystalline rocks with relicts of a former magmatic arc and continental crust, which are not observed in İstanbul itself, but crop out farther east near Zonguldak. This basement is overlain by a continuous, well-developed sedimentary sequence extending from the Lower Ordovician to the Upper Carboniferous (Namurian, i.e., top Visean + Serpukhovian + lowest Bashkirian in İstanbul area, Westphalian, i.e., Bashkirian to Kasimovian, farther east in Zonguldak). Sayar and Schallreuter (1989) collected Upper Caradocian–Asgillian ostracodes from this section and pointed out that they exhibited a faunal affinity with the Baltic region. Early Devonian ostracodes from the İstanbul area belong to the Thuringian ecotype and thus indicate affinity to the Gondwanian regions of the European Hercynides (Dojen et al. 2004). This has long been known by the identification of the İstanbul Palaeozoic fauna as being of ‘Rhenic’ type. The final members of the İstanbul Palaeozoic section consisting of Griotte-type pelagic, nodular limestones before the onset of flysch sedimentation had been earlier determined to be Upper Tournaisian to Visean, but the recent discovery of Middle Tournaisian conodonts has now pulled the lower age limit even lower (Göncüoğlu et al. 2004).

South of İstanbul, Okay et al. (2008) dated a series of granitic rocks ranging in age from 585 to 460 Ma, the oldest zircon ages coming from the cores of individual crystals leading Okay et al. (2008) to conclude that in that particular locality the intrusion age was 570 and that the older ages signified an older event. The scattering of the ages indicates nearly continuous magmatism from the Ediacaran to the later medial Ordovician (although there are no Cambrian ages where they did their collecting; as we have seen earlier, they exist in the Strandja), which is characteristic of northern Pan-African regions. Okay et al.’s (2008) claim that this provides evidence for the existence of Avalonia in Turkey is not supported by any piece of data they provide, but contradicted by the Palaeozoic stratigraphy of the İstanbul Fragment and, therefore, we do not follow it here. Both Dean et al. (1997) and Göncüoğlu (1997) concluded from palaeontological considerations that the entire İstanbul Nappe originally was part of Gondwana-Land. The Lower Carboniferous flysch records the progress of a shortening event, which then passes upwards into Upper Carboniferous molasse with coal seams that are economically exploited in the east (in Zonguldak).

East of İstanbul, between the towns of Hendek and Karasu, is a small and highly deformed and slightly metamorphosed Palaeozoic outcrop area containing a section from the Ordovician to Devonian consisting at the base of arkoses and passing upwards into shales and subordinate limestones and ending with sandstones overlain by redbeds. This section is correlated with the Palaeozoic in İstanbul and is believed to represent a part of the Palaeozoic sequence of the İstanbul Nappe closer to the continent than the İstanbul area proper. Lakova and Göncüoğlu recovered early Ludlovian palynomorphs from the Upper Silurian shale-sandstone member of the Fındıklı Formation in this sequence (Lakova and Göncüoğlu 2005).

Farther east in the Zonguldak area (the classical Heraclea Pontica) the Palaeozoic section is also similar to, but somewhat different from that in the İstanbul area. Here the section begins with Upper Silurian quartzites and reddish-greenish iron-bearing sandstones passing into Devonian neritic limestones. The Visean in the Zonguldak region is mainly dolomitic and becomes sandy and shaly just before the Namurian boundary. From the Namurian to the Stephanian the entire section is clastic with abundant and economic anthracite deposits. The entire section was folded and unconformably covered by the Permian (Ketin 1983). Kozur and Göncüoğlu (2000) thought they could identify an unconformity between the Silurian and the Devonian here with a thermal event, but the visible stratigraphic development makes such an interpretation of the field relations unlikely. However, the differences between the Palaeozoic geology of the Zonguldak region and the İstanbul region are substantial enough to assume that these two areas were probably not contiguous (see Göncüoğlu et al. 2004), although they were parts of the same platform. On the basis of the differences in the stratigraphy, Sachanski et al. (2008) indicated that the boundary between the İstanbul and Zonguldak sectors should be drawn west of Derince near the easternmost end of the Gulf of İzmit.

South of this area is the Bolu Massif that exposes the polymetamorphic and intruded basement of the Palaeozoic succession in the İstanbul Fragment, which Ustaömer and Rogers interpreted, on the basis of detailed petrological and geochemical data, as remnants of an Andean arc of pre-Ordovician age, which is in good accord with observation farther west (Ustaömer and Rogers 1999).

The Carboniferous deformation within the İstanbul Nappe led to the folding and faulting of the Palaeozoic sequence with a pronounced west vergence (present geographic orientation) in İstanbul which was intruded by an uppermost Permian granitoid and unconformably overlain by the Upper Permian to Lower Triassic red sandstones and conglomerates. The folding and thrusting are in the style of a marginal fold and thrust belt with none to low grade metamorphism (up to greenschist grade) and represents the retroarc shortening of an Andean magmatic arc here. The Permian redbeds contain ichnofossils of tetrapods and plant fossils (Gand et al. 2011). The Triassic series is best developed east of İstanbul displaying a transgressive, Alpine-type sequence, which, however, begins with a European-type Scythian (‘Alpines Buntsandstein’) with basaltic interlayers. The Jurassic sequence is absent except in the eastern end of the İstanbul Fragment (Tüysüz 2017), most likely as a result of the orogeny and uplift resulting from the closure of the Palaeo-Tethys just to the north (present geographic orientation). Cretaceous follows the Triassic here and there is a small outcrop of Lower Cretaceous shallow marine sedimentary rocks and a much more widespread Upper Cretaceous-Lower Eocene clastic rocks, carbonates and, along the northern shore of the entire fragment, andesitic volcanic rocks of latest Cretaceous age overthrust by the Palaeozoic section along a north-vergent thrust system. This was a result of the pre-Bartonian closure of the Intra-Pontide suture that here led to backthrusting towards the Black Sea during the Ypresian. The Intra-Pontide suture is the boundary between the İstanbul Zone and Sakarya Continent. During the Cainozoic, the first post-orogenic structures are Lutetian-Bartonian nummulitic limestones, which themselves are covered by a Paratethyan sequence of Miocene limestones and sandstones of mainly the Vallesian Stage. The Pliocene consists entirely of fluviatile terrestrial clastic rocks. The Pleistocene was deposited on an erosion surface which later became warped north in the east and south in the west of the Bosphorus and into which the originally fluvial valley of the Bosphorus was cut. This valley was invaded by the Sea during the Holocene and caused the refilling of the Black Sea.

In the east the thrust contacts of the İstanbul Fragment with the Küre Nappe are sealed by the Middle Jurassic sandstones of the Bürnük Formation (Şengör et al. 1980). It is more difficult to put an age on the thrusts in the west, with the Strandja Zone, but Ülgen et al. (2018) pointed out that the most reasonable estimate was medial to late Mesozoic.

Between the Palaeozoic outcrops of the İstanbul area and the eastern outliers of the İstanbul Nappe extends what Akbayram et al. (2016a) call the Kocaeli Basin, which they identified as a late Cretaceous to Palaeocene forearc depocentre. Above a medial Cretaceous puddingstone, the clasts of which come from the underlying Palaeozoic and Triassic rocks seen farther west around İstanbul (Erguvanlı, 1949), Özcan et al. (2012) reported Upper Campanian to Upper Palaeocene pelagic limestones and marls. In the southern part of the Kocaeli Peninsula there are no volcaniclastic rocks. In the north, Santonian to Campanian volcanic rocks are covered by Campanian to medial Palaeocene (Selandian) pelagic limestones of some 80 m thickness. Locally, the pelagic limestones grade into a shale-marl sequence of Thanetian-Ilerdian age. This is then overlain by a flysch with debris flows consisting mainly of Upper Campanian to Middle Palaeocene limestone blocks. All of these rocks were thrust northward by a retrocharriage over the volcanic rocks before the Lutetian which covers everything unconformably (Akbayram et al. 2016a and the references therein). This retrocharriage belt continues from İstanbul all the way to Sinop, but near Sinop the retrocharriage deformation is still active (Eşref Aylan, personal communication, 29th June, 2018).

Görür (1997), Görür and Tüysüz (1997) and Hippolyte et al. (2010) described the sequence that comes above the Upper Jurassic-Lower Cretaceous İnaltı shelf deposits consisting of a southerly thickening carbonate package in the Central Pontides: in the west, atop the eastern parts of the İstanbul Nappe the İnaltı is unconformably covered by the Ulus Formation (in the eastern part of the Ulus Basin) consisting dominantly of turbiditic clastics and it is the partial temporal equivalent of the shallower water clastic deposits farther west that sit on the İstanbul Nappe. Fossils are rare and those that have been found indicate a Valanginian to Campanian age. Tüysüz et al. (2004) estimate a Barremian to Cenomanian age, which is inconsistent with the Campanian fossils. Its unconformable relation with the İnaltı was a result of rifting. It begins at the base with clastic fans deposited in a very shallow marine environment, which gradually become a turbidite as the basin deepened. Farther east it begins directly with turbidites, but the base of the sequence is not seen (Tüysüz 2017). In the eastern outcrops extensive debris flows in the turbidites contain blocks of the İnaltı Limestones. Farther east near Azdavay and Şenpazar area, the debris flows are intercalated with volcanic rocks and the large blocks of Carboniferous, Permo-Triassic and top Jurassic-lowermost Cretaceous (Tithonian-Berriasian) blocks are found in debris flows that locally may reach hundreds of metres thickness. The blocks are in places large enough to allow coal mining in them. Tüysüz (2017) noted that the blocks were very angular and in places the debris flows exhibited no matrix; he concluded that the sources of these blocks must have been very close. At the eastern margin, the Kervansaray and Sada formations of large debris flows contain volcanic blocks and pyroclastic rocks alternating with shales giving ages of Kimmeridgian-Lower Valanginian ages. The geochemical compositions of these volcanic rocks are dominantly of MORB, ocean-island alkali basalts and island arc tholeiites. We interpret these as rift products atop the former Palaeo-Tethyan subduction zone beneath Moesia in view of the setting and the dominance of MORB and intra-plate volcanic rocks and not products of a subduction zone as Tüysüz (2017) does. The nearby Dikmen magmatic rocks consisting of andesite, trachyte and latite and small granodiorite and monzogranite intrusions associated with tholeiitic mafic lavas gave an age of 163 Ma (Callovian, Okay et al. 2014). The overall setting and the igneous rocks seem similar to the basin bounded by the Trobriand Fault in the Woodlark Rift (Little et al. 2007) opening atop a former subduction zone.

The Ulus Basin gives the impression of having started as a highly asymmetric rift basin with the main fault in the east. It was later destroyed by shortening. Since the İstanbul Nappe is considered a part of the Moesian Platform, the great similarity of the evolution of the Ulus Basin to the Nish-Trojan Trough (or Axial Basin: see Minkovska et al. 2002) invites a comparison and leads to two possible interpretations: either the Ulus is a part of the Nish-Trojan Basin, later separated during the rotation of Moesia (see below the discussion and conclusions section) or a basin similar to it formed in the same environment. In this regard, it is extremely significant that Okan Tüysüz (personal communication, 11th June, 2018) observed that the zircon populations in the Ulus Basin and in the regions east of it, atop the Küre Nappe, indicate completely different provenances. The area on which Ulus was to form was incorporated into the structure of the Pontides shortly after the Moesian Platform collided with the rest of the Pontides upon the closure of the Palaeo-Tethys before the medial Jurassic.

Central and Eastern Pontides

The only difference between the Central and Eastern Pontides is that the late Cretaceous magmatic arc axis goes offshore into the Black Sea south margin west of Sinop and thus permits a view into the pre-late Cretaceous geology of the Rhodope–Pontide Fragment here. The area the Central Pontides cover is confined between the longitude of Bartın-Karabük in the west and Sinop in the east and the Black Sea in the north and the İzmir–Ankara–Erzincan Suture in the south. In the Kargı Massif, the northernmost part of the Neo-Tethyan suture is exposed and in it the Domuzdağ Complex, consisting of an accretionary complex with the large chunk of the Elekdağ Lherzolite represents a blueschist terrain metamorphosed during the medial Cretaceous (about 105 Ma ago, i.e., during the Albian: Okay et al. 2006). A large subduction-accretion complex developed to the south that terminated its evolution upon the collision with the Sakarya continent during the late Cretaceous-early Eocene interval here.

The post-medial Jurassic cover of the central Pontide consists, north of the northern boundary of the İzmir–Ankara–Erzincan suture, following roughly the southern boundary of the Çaldağ Anticline, of a southerly thickening carbonate bank made up almost entirely of the Oxfordian to Valanginian İnaltı Limestones (Koçyiğit and Altıner 2002; Tüysüz et al. 2004, pp. 43–44; see Kaya 2014, pp. 4–12 and fig. 2, for an exhaustive review of the older literature about the İnaltı Limestones and their equivalents in northern Turkey with a useful map showing their distribution). The İnaltı Limestone sits on older units either unconformably or transitionally from the Bürnük clastic rocks with sandstones and limestones. In places the oldest members are simply reefal carbonates. Tüysüz et al. (2004) describe the İnaltı Limestones as fairly uniform, monotonous, light-coloured platform carbonates consisting of grey, white, in places darker grey, locally pinkish, medium to thickly-bedded algae-, bryozoa-, coral- and gastropod-debris-containing micrites, pelmicrites and oosparites. Tüysüz et al. (2004) place their site of deposition in tidal flats (they were able to document tidal-flat channels in them) and platforms just below the wave-base. Although spanning a similar time interval, this sequence exhibits internal stratigraphic differences from the Bilecik Group on the Sakarya Continent (which may span a longer interval from Callovian to Hauterivian: Altıner et al. 1991, pp. 29–31). Nowhere in the İnaltı there is a trace of vulcanicity except for a minor tuff horizon in the Hauterivian-Barremian boundary in the extreme east (Okay and Sahintürk 1997). To the south it sits on the rift to post-rift deposits spanning the age from the Hettangian-Sinemurian to the late Jurassic (Görür et al. 1983). It clearly represents a south-facing Atlantic-type continental margin that rifted during the Lias and was destroyed by the onset of north-dipping Neo-Tethyan subduction under the Rhodope–Pontide Fragment during the Aptian–Albian interval.

Once we step off the İstanbul Nappe eastwards onto the area of the Küre Nappe, the world above the İnaltı Limestones changes significantly. Here the dominant sequence is one of dominant sandstones passing into sandstones and shales. They have been gathered under the common designation Çağlayan Formation (Görür 1997; Görür and Tüysüz 1997; Tüysüz 2017), which begins with shallow water sandstones at the base containing reworked components of the underlying İnaltı and even Bürnük Formations. Tüysüz et al. (1990) pointed out that blocks of underying formation continued falling into the Çağlayan throughout its depositional history, except during its last phases when distal turbidites and organic-rich black shales dominated the sedimentation. Görür (1988), Görür and Tüysüz (1997), Tüysüz (2017) and Tüysüz et al. (1990) rightly concluded that Çağlayan was a syn-rift deposit, fed dominantly from highlands to its south and laid down in a restricted underwater basin-and-range topography related to the opening of the Back Sea marginal basin. Tüysüz noted that no blocks representing the İstanbul Fragment have yet been found in the Çağlayan, but blocks from the Kargı Massif were abundant. However, zircons, interpreted to have been derived from the Russian Craton are also seen (Akdogan et al. 2018), from which it was concluded that the Çağlayan rift, destined to become the Black Sea (at least its western basin) had opened between the Rhodope–Pontide Fragment and the Russian Craton plus its peripheral Scythides. The thickness of the Çağlayan is hard to estimate because of its great variability and the intense later shortening that affected it. Tüysüz gives a minimum thickness of 3 km. and an age range from the Hauterivian to Albian. The Black Sea rifting began sometime during the Çağlayan deposition and it is most likely that the switch from the shallow water members of the Çağlayan to the turbidites and debris flow deposits marked the onset of rifting, which may fall sometime into the Barremian.

This timing is very interesting: the oldest Neo-Tethyan subduction-related rocks in the Pontides are the ca. 105 Ma old blueschists from the Domuzdağ Complex (Okay et al. 2006), which fall into the Albian. For blueschists to form a depth of some 20 km is necessary. During the time interval of anomaly M0 (125 Ma) and 34y (84 Ma), the rate of head-on convergence between Africa and Europe along the longitude of the Central Pontides was about 1.1 cm/a (see Şengör and Stock 2014, fig. 18). If we assume a 45° dip for the subduction zone responsible for the 105 Ma blueschists, it would mean that we need a 28 km horizontal motion. If we further assume that the blueschists were generated when the subduction zone was activated, we get about 25. 5 Ma. This is precisely the time interval from early Barremian, i.e., when the rifting of the western Black Sea commenced, to 105 Ma ago! We shall see below that the magmatic arc magmatism also began shortly thereafter if an Aptian granite intrusion south of Samsun is not considered (Mr. Ozan Sungurlu, personal communication, 1980; as yet no description has been published).

This scenario seems to contradict the recent inference of a mid-Jurassic subduction-accretion prism in the Saka Complex (Okay et al. 2013). Okay et al. (2013) put the Saka Complex as a subduction-accretion pile at the prow of the İstanbul Fragment (see their fig. 17), whereas elsewhere it appears to the south of and below what they call the Domuzdağ Plateau of medial Cretaceous age (we see no evidence for a plateau in the Domuzdağ Complex; it is a subduction-accretion complex with blueschists as mentioned above). Thus, the Saka Complex seems an intraoceanic affair, not a part of the main Neo-Tethyan subduction. In fact, Çelik et al. (2011) came to the same conclusion about the similar medial Jurassic metamorphic rocks within the İzmir–Ankara–Erzincan suture farther west. We have expressed our reservations above about the subduction interpretation by Topuz et al. (2013a) for the Refahiye ophiolite, farther east along the İzmir–Ankara–Erzincan suture, which Okay et al. (2013) cite in support of their interpretation. We view the intraoceanic subduction events as local, possibly along bends in large transform faults within the Neo-Tethys, perhaps similar to the situations along the Macquarie Ridge (Williamson 1988) or the Gorringe Bank (Auzende et al. 1984).

Yılmaz and Boztuğ (1986) claimed that the medial Jurassic Kastamonu granitoid belt roughly between Zonguldak and Sinop, consisting of dark metaluminous to peraluminous granodiorites with biotite + amphibole, an intermediate peraluminous group with biotite + muscovite and a group of peraluminous leucoadamellite, was a product of a north-dipping Palaeo-Tethyan subduction. However, earlier, Boztuğ et al. (1984) had shown that these rocks had very largely resulted from crustal melting and the largest members (Ahiçay, Elmalıçay and Büyükçay plutons) do not even belong to the calc-alkalic association, a conclusion they later corroborated and amplified (Boztuğ et al. 1995). The claim similar to that by Yılmaz and Boztuğ (1986) for the Crimean plutons and volcanic rocks of Jurassic age (Meijers et al. 2010) suffers from the unsuitability of the subalkalic to tholeitic trends of the volcanic rocks including trachybasalts for a subduction magmatism interpretation. It seems clear that those rocks resemble, at least in part of their composition, to the alkalic volcanic rocks of Tibet, that are entirely post-collisional. They may very well be related to localised extension along strike-slip faults related to post-collisional convergence, as the widespread Dogger magmatism in the Caucasus also appears to be collision-related. We think that the original interpretation of Şengör et al. (1980) that these rocks are post-collisional, suture crossing plutons is better in accord with the observations.

To conclude an onset of major subduction already during the medial Jurassic south of the Rhodope–Pontide Fragment is, on the basis of the available observations, not justified, especially because there is no associated arc and the authors who postulate it are forced to introduce discontinuities into the subduction history, which major subduction zones tend not to have.

The end of the Black Sea rifting is marked by the deposition of the Kapanboğazı red pelagic limestones, shales and cherts of Cenomanian to Turonian age, essentially in Scaglia Rossa facies, blanketing the distal Çağlayan turbidites (Tüysüz et al. 2016). Tüysüz et al. (2016) emphasise that in these deposits there is no trace of vulcanicity; in fact, they underline that in the whole of the Pontides there is no Neo-Tethyan subduction-related magmatism before the Turonian. We entirely agree with this statement (excepting the possible Aptian granite south of Samsun) and infer that at least the western Black Sea began rifting in a pre-arc environment almost simultaneously with the onset of the north-dipping Neo-Tethyan subduction below the Rhodope–Pontide Fragment. This is in agreement with the growth of young marginal basins, during the birth of which the magmatic arc activity is greatly diminished, even where it existed previously (e.g., Letouzey and Kimura 1986). This post-rifting sequence, with a thickness of some 20 m., was then followed by about 8-km thick volcanogenic turbidites that accumulated in a fore-arc setting south of the magmatic arc, which in the Central Pontides is now almost wholly under the waters of the Black Sea. Within this post-rift sequence, Cemal Tunoğlu discovered fossils of a Mosasaurus hoffmanni at the very top of the Maastrichtian part of the Davutlar Formation, some 30 km south of the coastal town of İnebolu (Bardet and Tunoglu 2002; Tunoğlu and Bardet 2006). This finding indicates that the Rhodope–Pontide Fragment at this time was around 30°N (Bardet and Tunoglu 2002), which is perfectly compatible with other available geological data.

The Eocene collision resulted in a mostly bivergent deformation of these sedimentary rocks in a fold-and-thrust belt on both sides of which the Boyabat (in the south) and the Erfelek (in the north) peripheral flexural basins formed. Younger sedimentation was largely confined to the Sinop Peninsula in the north and consists of Paratethyan sedimentary rocks (Görür et al. 2000).

Much of the exposed geology of the Eastern Pontides consists of late Cretaceous and Eocene magmatic arc volcanic rocks. Robinson et al. (1995), Okay and Şahintürk (1997) and Yılmaz et al. (1997b) reviewed their geology. There are more recent observations on the detrital zircon populations (e.g., Akdoğan et al. 2018) some of which as yet unpublished. Yener Eyüboğlu obtained in the detrital zircons he collected from the Eastern Pontides two peaks at 2.5 and 1.7 Ga (Yener Eyüboğlu, written communication, 26th May 2018). The Precambrian ages are reminiscent of the interval 2600 to 1600 Ma from northeast Africa (Harms et al. 1990), but rather different from the basement of the Sakarya Continent, further underlining that the Sakarya Continent and the Central and eastern Pontides should not be lumped under a Sakarya Zone as has recently become fashion among some Turkish authors. The continental basement of the Eastern Pontides is largely exposed in the Gümüşhane-Bayburt region and is formed from amphibolite-grade schists, gneisses and migmatites plus the Gümüşhane granite pluton of late Carboniferous age (Yılmaz et al. 1997b). The ages of the metamorphics are unknown, detrital zircon ages scatter from the latest Devonian into the medial Jurassic, but given Yener Eyüboğlu’s recent unpublished results, Precambrian seems a significant component. On top of this polymetamorphic basement a sedimentary sequence of Upper Carboniferous age sits consisting of alternating arkoses, sandstones, conglomerates, light-coloured quartzites and dark fossiliferous limestones (Okay and Leven 1996; Çapkınoğlu 2003) with two horizons of quartz-containing lavas and horblende-biotite andesites (Ketin 1951), essentially coeval with the Gümüşhane pluton. In one of the higher limestone beds, a late Carboniferous trilobite, the pygidium of an undetermined species of Ditomopyge Newell 1931, was recently discovered (Kandemir and Lerosey-Aubril 2011). This genus has a wide distribution in Pangaea, including North America, but, significantly, does not occur in Europe north of the Alpine suture or in Asia, north of the Palaeo-Tethyan suture. Stephanian plant fossils found here indicate an Euramerian affinity (see Şengör 1990c) and some wish to take this as an indication of a Laurussian attachment for the Eastern Pontides (e.g., Okay and Şahintürk 1997, pp. 296–297). This is, however, inappropriate as a criterion of demarcation, as an Euramerian flora is known from the indisputably Gondwanian Sibumasu part of the Cimmerian Continent (Torsvik and Cocks 2017, p. 192) and Ziegler (1990) indicated a similar occurrence from the middle of Arabia (his biome 2, which he treats as a province).

The Palaeozoic sedimentary rocks of the Eastern Pontides are covered by a Liassic to early Cretaceous sedimentary package consisting of two parts: one part, a mainly volcanoclastic-dominated sequence, called the Kelkit Formation, spans the time interval from the Pliensbachian to the end of the Bajocian with a thickness that fluctuates between 1.5 to 2 km and is generally viewed as a rift facies that formed during the opening of the Neo-Tethyan ocean here (Görür et al. 1983; Okay and Şahintürk 1997). It has a wide distribution from north of the Tokat massif to Yusufeli. Okay and Şahintürk summarised its three dominant facies as follows: a turbiditic volcanoclastic sandstone shale intercalation, a widespread lithic tuff and a volcanoclastic sandstone–conglomerate unit with coal horizons. This last is interpreted as a paralic environment and is the most attractive target for finding Mesozoic tetrapod fossils in TUrkey (possibly including dinosaurs). The third dominant facies is a Rosso Ammonitico representing tops of high-standing fault blocks. As one goes southward, the Kelkit becomes more pelagic with a finer grain size. Okay and Sahintürk (1997) find the interpretation by Görür et al. (1983) that the Kelkit Formation is a rift/Atlantic-type continental margin facies problematic, because its stratigraphy is dissimilar to the opposing Tauride margin. This, however, is not entirely true, as Özer (1994) reported south of the Tan village and north of the Hamarat Hill in the Munzur mountains shaly, marly calcareous sequences containing ammonite moulds which he thinks represent a continental slope facies of the otherwise platformal, ubiquitously shallow-water Munzur Limestones of Liassic to Cretaceous age (Özgül and Turşucu 1984). As the Eriç Complex overthursting them represents an accretionary complex, it clearly picked up more northerly pieces representing the continental margin than the uniformly neritic facies of the lower and therefore more southerly nappes studied by Özgül and Turşucu (1984). Okay and Şahintürk further point out, following Şengör and Yılmaz (1981), that the Liassic faunas on both sides of the İzmir–Ankara–Erzincan suture along the Eastern Pontides are different: one belonging to Laurasia and the other to pieces rifted from Gondwana-Land and cite Bassoullet et al. (1975) and Enay (1976). Since Enay (1976) deals only with Oxfordian and younger faunas it cannot serve the purpose for which Okay and Şahintürk (1997) cite it. Bassoullet et al. (1975) indeed claim what Şengör and Yılmaz (1981) and Okay and Şahintürk say, but then point out that both Morocco and Algeria have ‘Pontic’ elements in their fossil record. Moreover, what Bassoullet et al. (1975) identify as ‘south Tethyan’ fauna also occurs in Majorca, where, during the Liassic, the rifting of the Neo-Tethys was just taking place. The faunal differences thus seem more climate- and environment-controlled than due to a Liassic oceanic separation. However, the ocean-opening in the İzmir-Ankara segment of the İzmir–Ankara–Erzincan Suture and the opening along the Inner Tauride Ocean was early Triassic. This makes it likely that the continental slope blocks in the Eriç Complex are of that age. This then still leaves open the question as to what rifted from the Eastern Pontides to open the Neo-Tethys here regardless of the faunal arguments. We shall return to this question in the discussion section below. The lively vulcanicity of the Kelkit Formation represents the last phase of the Palaeo-Tethyan subduction here which completely switched off by the end of the medial Jurassic (Okay and Şahintürk 1997). In fact, Akdoğan et al. (2018) recently discovered Triassic to medial Jurassic zircons from the Sinemurian-Pliensbachian Şenköy Formation overlying the late Carboniferous Gümüşhane Granite. That shows continuous magmatism in the Eastern Pontides at least from the Triassic to the medial Jurassic, which is not dissimilar to the situation in the Sakarya Continent and the Western Pontides. Akdoğan et al. (2018) interpret this as resulting from a north dipping subduction under Laurasia, which is impossible, because rifting of the Neo-Tethys had just begun at that time and the Eastern Pontides were still south of the Palaeo-Tethys as shown above.

Okay and Şahintürk (1997) mention another Jurassic facies in the Pulur area that begins with Liassic shallow-water limestones sitting unconformably on Carboniferous sedimentary rocks. They pass gradually into an oolitic limestone and then to a thickly-bedded cherty micrite. These rocks span the interval Liassic to Valanginian. Above them a sequence of Senonian deep-water rocks sit unconformably. As other nearby rocks exhibit a volcaniclastic sequence, Okay and Şahintürk (1997) reasonably conclude that the Pulur area must have been a high-standing block during the rifting that did not receive the volcaniclastic-bearing turbidites.

An ophiolitic mélange nappe extending from Çimendağ about 20 km northwest of Erzincan to north of Tortum was thrust onto the magmatic arc of the Eastern Pontides and now rests mostly on Neocomian pelagic limestones, but dissolved into individual klippen owing to later erosion. The blocks in the mélange have minimum ages of Aptian, corresponding to the earliest possible time of the initiation of north-dipping Neo-Tethyan subduction under the Eastern Pontides. It is surmounted by yet another nappe of entirely sedimentary derivation, consisting mainly of pelagic radiolarian limestones (Kimmeridgian to Berriasian) ending upwards in calcilutites with intercalated mafic flows of Lower Senonian age (Bergougnan 1987; Okay and Şahintürk 1997 and the references therein). These nappes appear to have been parts of an accretionary complex to the south of the Eastern Pontides and were emplaced as backthrusts to the edge of the arc massif before the Maastrichtian, when rudist-bearing limestones unconformably covered them (Okay and Şahintürk 1997 and the references therein).

The second and vigorous magmatism in the eastern Pontides commenced during the Turonian, following an episode of carbonate deposition ranging from about the Callovian to the Berriasian in the north and Aptian in the south and lasted, with interruptions, into the Oligocene. This time span is correlative with the growth of subduction–accretion complexes to the south of the Eastern Pontides and clearly resulted from the north-dipping subduction of the northern branch of the Neo-Tethys under them. Okay and Şahintürk (1997) summarised the literature as to the contents of this 2 km-thick volcanic pile: it started with basalts and andesites and evolved into dacitic and rhyolitic eruptions with tuffs, breccias and local and subordinate limestones. The magmatism here generated during the Maastrichtian Kuroko-type massive sulphide deposits and Okay and Şahintürk (1997, fig. 2) used their occurrence to define the magmatic arc axis in the Eastern Pontides during the Maastrichtian. They further speculated that this was the time when the Eastern Black Sea probably began to rift. We basically agree with this interpretation. The first cycle of magmatism in the Eastern Pontides ended during the Maastrichtian-Danian.

The cessation of magmatism coincided with the onset of renewed northward thrusting of the southern edge of the arc and the emergence of the entire arc above sea-level. Okay and Şahintürk interpret this episode of deformation and emergence as resulting from the collision of the eastern Pontides with the continental units to the south (they still use the old and by then inappropriate term of Şengör and Yılmaz 1981, ‘Anatolides-Taurides’; a correct description would have been the Kırşehir Massif and the Menderes–Taurus Block; see above). We disagree with this interpretation, because, after the first episode of vulcanicity, a second cycle began and ended largely during the Lutetian, possibly between the Ypresian and the Lutetian. This cycle was more mafic in composition than the earlier cycle and geochemically resembles island arc calc-alkalic rocks (Okay and Şahintürk 1997 and the references therein). This indicates ongoing subduction. Moreover, the onset of Palaeocene thrusting coincided with the uplift of the entire arc. Had another continent impinged on it, the arc would have subsided under the weight of the thrusts. In the combination of arc cessation, renewed backthrusting and wholesale uplift of the arc we rather see indications of ridge subduction under the Eastern Pontides during the Maastrichtian–Palaeocene interval (duration about 10 Ma). This is consistent with the subduction of fairly warm crust that is about 300 km wide astride a slowly-spreading spreading centre (see especially, Eilon and Abers 2017). Remarkably, the Palaeocene volcanic gap is not observed farther west, where in the eastern part of the Central Pontides, Mr. Eşref Aylan (personal communication, 20th June, 2018) noticed the presence of tuffs in the Palaeocene section. Moreover, east of the Munzur Mountains, just east of Erzincan, there was no continent to collide with the Eastern Pontides, as the entire East Anatolian High Plateau is underlain, as far south as the Bitlis Massif, by a large subduction-accretion complex belonging to the Eastern Pontides, in which granite intrusions continued almost into the Oligocene (see Şengör et al. 2008a, b).

After the Eocene, the Eastern Pontides remained largely above sea-level with occasional transgressions coming from the Paratethys and inundating its bays along the northern shores of the Eastern Pontides (Okay and Şahintürk, 1997). Shortening continued up to the present under the influence of first continued subduction till the Oligocene and then collision with Arabia (see especially Şengör et al. 2008a, b).

Synopsis of the Rhodope–Pontide Fragment

The Rhodope–Pontide Fragment did not have an independent identity before the opening of the Black Sea that confined it between two oceans: The Black Sea in the north and the northern branch of the Neo-Tethys in the south and therefore that designation should not be used for times before the latest Cretaceous. Until the Jurassic, the Central and the Eastern Pontide parts were segments of the northern, Pacific-type continental margin of Gondwana-Land. The Strandja and the İstanbul segments were parts of Moesia that had collided with Laurussia in the Carboniferous. The collision pushed parts of Moesia eastward into the Palaeo-Tethys, which began subducting under it. It seems that Moesia and the rest of the Rhodope–Pontide Fragment became united before the Bajocian when the Palaeo-Tethys became completely eliminated by continental collision. That collision made the Rhodope–Pontide Fragment a part of Laurasia, but by that time it had already rifted away from the pieces of Gondwana-Land that had rifted earlier, during the Triassic from the main body of the continent. A south-facing Atlantic-type continental margin formed on it, which was destroyed by the onset of the subduction of the northern branch of the Neo-Tethys under it, at the latest by the Turonian. That subduction began opening the Black Sea north of it during the Aptian to Maastrichtian interval (continuing into the Eocene) making the Rhodope–Pontide Fragment an independent first-order tectonic unit, which resembled the present-day Japan island arc. That arc collided with the Sakarya Continent, the Kırşehir Block and the easternmost tip of the Menderes-Taurus Block during the early Eocene. The shortening across the unit continued in places until the early Miocene (in the Sinop area it is still continuing!). During the medial Miocene, the North Anatolian Fault formed in most places following the Intra-Pontide Suture (Şengör et al. 2005), thus largely delimiting the Rhodope–Pontide Fragment against the first order palaeotectonic units of the Alpides to the south. However, it is better to refer to the areas north of the North Anatolian Fault as the North Anatolian Neotectonic Province (Şengör 1980; Şengör et al. 1985) and discontinue the usage of the name of the palaeotectonic unit for the time after the formation of the North Anatolian Fault, because their spatial correspondence is not perfect and they formed as a consequence of entirely different tectonic events at very different times.

Discussion and conclusions

The first sentence of the Şengör and Yılmaz (1981) paper was dictated by the late Kevin Burke: ‘The analysis of the structure of Turkey has proved exceptionally difficult in the past, because of a large concentration of a number of convergent events throughout its history.’ This statement retains its validity, but we must now add to it also the rifting and strike-slip events. Although even Archaean zircons have been obtained from the basement of the Taurides, the Precambrian record is too little exposed and what little there is has been so pervasively superposed by subsequent events that there is as yet little hope to reconstruct a palaeogeography for those remote times. We are, however, fairly certain of one thing: all of Turkey belonged to Gondwana-Land at the end of the Pan-African events. Wherever Precambrian is reached in Turkey it invariably reveals traces of Pan-African events. In the geological literature on Turkey one encounters such terms as ‘Assyntic’ in the older publications and ‘Cadomian’ in the newer; they all signify Pan-African events that pulled together Gondwana-Land and that lasted into the Ordovician (in fact, Léon Bertrand’s type Cadomian itself, in northern France, is Pan-African). In Europe there really is no clean break between the Pan-African events and the Hercynian events despite the rifting of the Avalonian magmatic arc that led to the opening of the Rheic Ocean behind it.

As far as we can see, there were no well-documented early and medial Palaeozoic orogenic deformations in Turkey, but both the Ordovician and the Devonian are represented by what seems to be arc magmatism. It is seen, albeit very patchily and commonly only by means of its zircons, in all of the Alpide first-order tectonic units of Turkey with the exception of the Arabian Platform. This activity represents a long-lived magmatic arc rimming the northern margin of Gondwana-Land.

By contrast, there are significant late Palaeozoic orogenic events, but none are Hercynian (or ‘Variscan’). This is important to emphasise, because in Turkey, every late Palaeozoic event has a tendency to be dubbed as Hercynian or Variscan. Okay and Topuz (2017) recently published a paper discussing the alleged Variscan orogeny around the Black Sea, although that orogeny never reached that far! They seem to think that anything late Palaeozoic must be Variscan. This is like saying that the late Palaeozoic events of the North American Cordillera are Variscan. This usage of the term Variscan (or Hercynian) is a remnant of Bertrand’s (1887) and Stille’s (1924) ideas, who believed that every orogeny was worldwide and synchronous. Even before plate tectonics this idea was shown not to be correct (Suess objected to it in 1909), but after plate tectonics continued usage of a terminology that has its root in a defunct theory is inadmissible.

Figure 18 shows where the problem lies in the case of Turkey. Figure 18a shows a very simplified and schematized picture of the late Paleozoic situation in the world, after the completion of the Caledonides by the partial closure of the Iapetus and while the Rheic/remnant Iapetus, Pleonic and the Palaeo-Tethyan oceans were still open. At this time, Gonwana-Land had a Pacific-type continental margin to the north with south-dipping subduction that formed what Lom et al. (2017) the called Protogonos Arc (first-born in classical Greek). That margin facing the Rheic/remnant Iapetus Ocean housed the evolving Hercynides and the other facing Palaeo-Tethys was the site of the Cimmerides. Turkey lies at the junction of the two, but on the ‘Cimmeride’ side. At the same time the Uralides were evolving at the expense of the Pleonic Ocean. Figure 18b illustrates the situation after the closure of the Rheic/remnant Iapetus and the Pleonic Ocans. Now we have completed the Hercynide and the Uralide orogens, but since Palaeo-Tethys was not yet closed, the evolution of the Cimmerides was continuing. If we call all late Palaeozoic orogens Hercynian or Variscan one can easily see the confusion that would create; one would not only throw unrelated orogens into one basket, but one would also conflate separate oceans and tear genetic connections (e.g., separating the Cimmeride events into a Hercynian and a Cimmeride episode, which cannot be justified by any means in a plate tectonic context). One cannot very well object to this by pointing out that the Rheic and the Palao-Tethyan oceans were in communication before the closure of the former and therefore should be considered a single ocean, because this would be like arguing that the Arctic and the Atlantic Oceans should all be named the same because one is the continuation of the other. We name things for descriptive facility and while doing so pay attention, as much as we can, to genetic connexions.
Fig. 18

a A very simplified and schematized picture of the late Palaezoic situation in the world, after the completion of the Caledonides by the partial closure of the Iapetus and while the Rheic/remnant Iapetus, Pleonic and the Palaeo-Tethyan oceans were still open. At this time, Gonwana-Land had a Pacific-type continental margin to the north with south-dipping subduction. That margin facing the Rheic/remnant Iapetus ocean housed the evolving Hercynides and the other facing Palaeo-Tethys was the site of the Cimmerides. Turkey, shown in grey, lies at the junction of the two, but on the ‘Cimmeride’ side. At the same time the Uralides were evolving at the expense of the Pleonic Ocean. b This figure shows the situation after the closure of the Rheic/remnant Iapetus and the Pleonic Ocans. Now the Hercynide and the Uralide orogens were completed, but since the Palaeo-Tethys was not yet closed, the evolution of the Cimmerides was continuing. The Palaeozoic parts of the Cimmerides are thus not Hercynides (or ‘Variscides’)

We stated above that the late Palaeozoic events in the Strandja Mountains were not easy to classify. This is because orogeny in the Strandja did not end with the late Palaeozoic events and continued into the Jurassic. Some of the Palaeozoic events in the Strandja resemble those in the Palaeozoic of İstanbul. This connects it with Moesia, but the Moesian Permian has typical European sedimentary rocks including evaporites (e.g., Yanev et al. 2005). Yet its continuation in İstanbul also has Permian granodiorite intrusions indicating ongoing subduction under a continent that had typical European earliest Triassic sediments, but intercalated with mafic volcanic rocks, that rapidly became Alpine-type until the Carnian and then became deformed without Jurassic being laid down on it. The way to resolve this dilemma seems to be to allow Moesia to collide with Laurussia in the Carboniferous with the Strandja at its northern margin and then be ejected into the Palaeo-Tethys, together with the Strandja, while doing an almost 180° rotation around a vertical axis, much like the units around the eastern syntaxis of the Himalaya today. This would explain how the Gondwanian Moesia with typical ‘European’ coal basins ended up being above a Palaeo-Tethyan subduction zone, yet possessing all the hallmarks of a late Palaeozoic collision. This further explains why various parts of the İstanbul Fragment have differing geologies. They were probably dissevered from one another and even rotated independently during the ejection and the general rotation. It also helps to solve the conundrum of the North Dobrudja, a fault-bounded fragment that has a typical Alpine-type Triassic amidst foreland-type Permian deposits and Permian ocean opening and pre-Liassic closure (Fig. 16, col. 34). If it was in the south, and part of the northern margin of Gondwana-Land, we can imagine it being nothing more than a piece of the Karakaya marginal basin, ripped and transported north with the rotation of Moesia. All of this probably happened before the final closure of the main Palaeo-Tethys during the medial Jurassic here.

All the Neo-Tethyan, i.e., Alpide sutures in Turkey carry evidence of rifting during the Permian to Liassic interval. We now have evidence that even an earlier basin may have opened during the early Carboniferous in northern Gondwana-Land within the future Menderes-Taurus Block. The overall extensional regime reigning in northern Gondwana-Land facing the Palaeo-Tethys during the late Palaeozoic is thus also well-documented in Turkey. The earliest rifting events following the Carboniferous marginal basin opening was during the Permian in the Antalya Nappes and around the eastern Mediterranean. However, this did not lead to ocean opening in the Eastern Mediterranean that finally opened during the Lias as a consequence of the opening of the Central Atlantic (Le Pichon et al. 2019). The Zagros sector of the Palaeo-Tethys, that began opening during the Permian, seems originally to have continued into the Inner Tauride Ocean and from there to the westernmost segment of the İzmir–Ankara–Erzincan sector. Along these oceans, deep-water conditions existed already during the early Triassic (in Oman, even in the Permian).

The Intra-Pontide Ocean rifted during the Hettangian-Sinemurian interval. The remarkable uniformity of rifting during this time in the entire Mediterranean Alpides from the Betic Cordillera to Turkey is astonishing and may indicate some connexion to the opening Atlantic all the way to Turkey, perhaps much like the Eastern Mediterranean. As such a connexion had to have been along transtensional transform fault systems because of the well-known geometry of opening in the Central Atlantic Ocean, segments of constraining bends along them may very well have triggered local sites of very early, even Liassic, subduction that are now confused with the onset of wholesale subduction of the Neo-Tethys in Turkey.

All the Neo-Tethyan continental blocks are now well-defined in Turkey. None of them had any independent existence earlier than the Triassic, one, the Rhodope-Pontide Fragment, earlier than the late Cretaceous. When discussing their provenances, it is important to specify for which time period one is speaking and one should employ only those parts of the geological record falling into that interval to make comparisons and correlations with other blocks. We carefully named each block with a tectonic designation where it was appropriate. The Rhodope-Pontide Fragment was a fragment of Laurasia until it rifted from it by the opening of various basins such as the Black Sea. The Sakarya was a true continental fragment. The Menderes-Taurus Block is so named, simply because it is nothing but a continuation of the African Promontory and calling it ‘Apulia’ may have misled some readers. Arabia is a platform. We avoided calling these ‘terranes’ because such a designation is uninformative. It would have been like calling them apples or oranges.

The Turkish Neo-Tethyan oceans closed mainly during the early Eocene, and mainly along north-dipping subduction zones (the only exceptions is the closure of the Antalya Ocean and the south-dipping subduction zone of the northern moiety of the Kırşehir Massif) although subduction in the east, under the East Anatolian Accretionary Complex, continued until the Oligocene producing ongoing magmatism in the Eastern Pontides. The final ocean closure in Turkey was preceded by large ophiolite obductions onto the Menderes-Taurus Block and onto the Arabian Platform from the north during the late Cretaceous. The northern obduction event permitted us to consider the Kırşehir Block, that carries parts of the giant ophiolite nappe, a part of the Menderes-Taurus Block until it was severed from it, transported westward with respect to it and in the process became bent into an easterly-opening hairpin. We tentatively place it to the east of the Bitlis Massif as its original continuation. The continental collision between Arabia and the rest of Turkey began in the Eocene north of Hatay and gradually progressed westward until the end of the Oligocene. Medial Miocene marked the onset of an entirely new regime in Turkey dominated by the westward escape and internal disintegration of Anatolia.

One thing we often mention in this paper, but do not illustrate, is the importance of strike-slip motion during the palaeotectonic events in Turkey. The most conspicuous examples were the massive Eocene–Oligocene right-lateral translations around the İzmir–Ankara–Erzincan suture and probably also embracing the entire circum-Black Sea area. Şengör (1990c) pointed out the presence of similar movements in northern Iran and the Turkish lateral motions. This indicates that the latter was nothing but a continuation of the former. The first concrete evidence for such motions was published by Akbayram et al. (2016a), but even in the older literature there are indications. The minimum cumulative right-lateral displacement in northern Turkey during the Eocene–Oligocene was probably more than 300 km judging from the displacements reported by Akbayram et al. (2016a). The cause of this lateral motion is unknown. It is not required by the direction of the Africa-Eurasia motion vector and seems a consequence of the collision. Another possibility is the effects of the Himalayan collision in the east.

Another significant lateral motion is the bending of the Kırşehir Block. We have no basis on which to hazard a guess as to which faults may have accomplished its translation and bending. The faults suspected to have existed in Eastern Turkey and seem to have created the transform sutures (the ‘unlikely oceans’ of Robertson et al. 2013a) at the eastern termination of the Menderes-Taurus Block were most likely also Eocene–Oligocene, but no study yet exists that addressed this question. The existence of the Murmano Pluton near Divriği (Zeck and Ünlü 1988) may be one obstacle in front of a strike-slip interpretation of the suture between the Munzur Mountains and the Malatya Digitation. Whatever comes out of future studies, it seems clear even from the existing meagre observations that the Neo-Tethyan oceans in Turkey can no longer be treated as simple open-and-shut cases.

What are the remaining outstanding problems? One problem that is frequently pointed out is the disparity in timing of the rifting between the northern margin of the Menderes–Taurus Block and the southern margin of the Rhodope–Pontide Fragment. Whereas the Taurus northern margin already had a deep-water facies in the early Triassic, the southern margin of the Rhodope–Pontide Fragment (then Laurasia) only rifted during the Hettangian–Sinemurian interval. We have difficulty understanding why this is presented as a problem, because the Sakarya Continent intervenes between the Rhodope-Pontide Fragment and the Menderes–Taurus Block and its southern margin indicates rifting even during the late Permian (e.g., Altiner et al. 2000). Therefore, it is very likely that it was the Sakarya Continent that rifted from the Menderes–Taurus Block. However, during the late Permian-early Triassic the Sakarya Continent as such did not yet exist. Instead the north-facing Cimmeride arc including pieces of the future Rhodope–Pontide Fragment and the Sakarya Continent rifted from the Menderes–Taurus Block, which itself was still a part of the northern margin of Gondwana-Land albeit with some incipient rifts in between. It thus seems that the Karakaya Basin opened almost coevally with the westernmost segment of the İzmir–Ankara–Erzincan suture and the Inner Tauride suture, but in those days the Inner Tauride suture was not facing the Kırşehir Massif to its north. This supports the original interpretation by Şengör et al. (1980) and Şengör and Yılmaz (1981) that the Karakaya was a back-arc basin of the south-dipping Palaeo-Tethys.

Another problem is the provenance and travel path of the Kırşehir Massif. The problem arises because of its metamorphic nature and the paucity of studies concerning its geological history. Almost everybody now agrees that the basement geology of the Kırşehir arc is very reminiscent of the geology of the Menderes–Taurus Block and most likely it was once a part of it. The question is where? The only available place now seems to be the eastern end of the Bitlis Massif, because everywhere else all parts of the Menderes–Taurus Block have their northern margins and the ophiolites that had been obducted from those margins. Since Kırşehir has the same stratigraphy as the Menderes–Taurus Block and a metamorphism and ophiolite obduction during the late Cretaceous, the continuation of the Bitlis eastward is for the time being the best candidate as the place of the Kırşehir. This would also give a reason to bend it into a crescent while travelling westward and then have it squeezed into a hairpin during the Eocene collisions and the following intracontinental shortening.

Finally, the Ordovician ostracodes with a Baltic affinity from the İstanbul Fragment seem to pose a problem for biogeography. All other fossils from İstanbul until the Carboniferous indicate a Gondwana-Land affinity. Baltic regions were very far away from Gondwana-Land at the time, across the Tornquist Sea; even Avalonia had not yet collided with the Russian Craton and the Caledonides. It is difficult to resort to an assumption of worldwide distribution of the relevant ostracode species, because ostracodes are notoriously provincial and thus excellent biogeographic indicators. However, Klimophores also occurs in Saudi Arabia, a good Gondwana-Land real estate and also in Thuringia (Sayar and Schallreuter 1989), which was also a part of Gondwana-Land at the time. Eochilina occurs as far as northeastern Russia, but that part of the world was even farther away from Gondwana-Land than Baltica during the Ordovician. A consolation here is that the Russian Craton and northeastern Russia were at the time separated by wide oceans and thus one cannot give up hope of finding Eochilina one day somewhere in Gondwana-Land. Piretella occurs in the Bohemian Massif, which was at the time also a part of Gondwana-Land. Therefore, we find the conclusion by Sayar and Schallreuter (1989) that their finds in İstanbul indicate a Baltic affinity premature. It would have been truly extraordinary to find a Baltic association on a piece of continent that both before and after that time belonged to Gondwana-Land.

All other problems in the geology of Turkey are of the kind that will be sorted out when more and better data in the form of detailed geological maps and rock ages are available. One problem concerning the geological work in Turkey is the recent proliferation of geochemical and geochronological papers without a solid geological foundation and generation of tectonic evolution models without any regard to actualistic consideration of present-day environments. One of the messages of the Şengör and Yılmaz (1981) paper was that the geology of Turkey cannot be understood by confining one’s attention to Turkey alone.

The second part of this paper contains the palaeogeography and palaeotectonic evolution of Turkey on the basis of time-lapse frame reconstruction maps. In a third paper we hope to deal with its neotectonic episode.

Footnotes

  1. 1.

    Ophiolites are those fragments of oceanic crust and upper mantle displaying the complete ophiolite sequence progressing from pillow lavas at the top, through sills and sheeted dikes, isotropic gabbros, cumulate gabbros, cumulate ultramafics, and ultramafic tectonites, commonly above a highly sheared garnet amphibolite sole (the Penrose definition). Such fragments are found emplaced by thrusting onto continental margins (obduction). Ophirags are fragments of ophiolites that have been dismembered in subduction zones, along collision fronts or along major strike-slip faults.

  2. 2.

    ‘Crushing sledge’, used for the uppermost, overriding thrust plates that maintain a degree of rigidity.

  3. 3.

    These two papers are mostly about palaeostress inversions. We do not think it can be done, because the striations on fault mirrors almost always record the youngest motion and they are almost impossible to date, unless one finds dateable calcite in fault steps. Also, McKenzie (1969) showed that it is not possible to obtain stress orientations from faults even when the slip on them is known. Even the remedy suggested by Gephart and Forsyth (1984) is of little help, because of the widespread local and regional to subcontinent-size anisotropies prevalent in the continental crust. Moreover, progressive deformation creates structures of contrasting types without any deviation in the continuous change of shape of rock bodies (Flinn 1962, 1965, 1994; Dewey 2002; Şengör and Bozkurt 2013). Finally, even in simply deformed areas there is always more than one episode of deformation and brittle structures are the easiest to rejuvenate. For all these reasons, we do not follow the conclusions of Kaymakçı et al. (2000, 2003a) concerning stress estimates.

Notes

Acknowledgements

We thank our colleagues in İTÜ and in some other Turkish universities, in the Turkish Petroleum Company and the Turkish Geological Survey (Maden Tetkik ve Arama Genel Müdürlüğü) for many years of exchange of ideas and information. For this paper, Aral İ. Okay, Gültekin Topuz, Demir Altıner and Erdin Bozkurt were particularly helpful about informing us of the data sources and giving us their assessments on individual problems. We also thank our editor Attila Çiner and Ali Polat from University of Windsor for their detailed reviews and Douwe van Hinsbergen from Universiteit Ultrecht, who read and reviewed an earlier version of this paper.

Compliance with ethical standards

Conflict of interest

On behalf of all authors, the corresponding author states that there is no conflict of interest.

Supplementary material

42990_2019_7_MOESM1_ESM.xlsx (57 kb)
Supplementary material 1 (XLSX 57 kb)

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Copyright information

© Springer Nature Switzerland AG 2019

Authors and Affiliations

  • A. M. Celâl Şengör
    • 1
    • 2
    Email author
  • Nalan Lom
    • 1
    • 4
  • Gürsel Sunal
    • 2
  • Cengiz Zabcı
    • 2
  • Taylan Sancar
    • 3
  1. 1.İTÜ Avrasya Yerbilimleri EnstitüsüİstanbulTurkey
  2. 2.Jeoloji BölümüİTÜ Maden FakültesiİstanbulTurkey
  3. 3.Edebiyat FakültesiMunzur ÜniversitesiTunceliTurkey
  4. 4.Department of Earth SciencesUtrecht UniversityUtrechtThe Netherlands

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