Timing of tectonic emplacement of the ophiolites and terrane paleogeography in the Hellenides

Lithos 108 (2009) 262–280

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Timing of tectonic emplacement of the ophiolites and terrane paleogeography in the Hellenides Dimitrios Papanikolaou Department of Geology, University of Athens, Panepistimioupoli Zografou, 15784 Athens, Greece

a r t i c l e

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Article history: Received 20 December 2007 Accepted 8 August 2008 Available online 22 August 2008 Keywords: Ophiolite obduction Aegean Lesvos Crete Volvi Vardar

a b s t r a c t The timing of tectonic emplacement of the ophiolites is analyzed in the four oceanic terranes of the Hellenides (H2, H4, H6, H8). The criteria for this analysis are based on: a) the post-emplacement sedimentary cover or intrusive rocks, b) the syn-emplacement tectonostratigraphic formations and c) the youngest rocks involved in the structure of the autochthon and the allochthon unit in each case. The timing becomes younger towards the more external tectonic units of the Hellenides with: (i) Late Eocene–Oligocene age in the external ophiolite belt of the Pindos–Cyclades oceanic terrane H2, (ii) Late Jurassic–Early Cretaceous age in the internal ophiolite belt of the Vardar/Axios oceanic terrane H4 , (iii) Post-Liassic–pre-Late Jurassic age in the ophiolites of Lesvos– Circum Rhodope oceanic terrane H6 and (iv) Pre-Late Jurassic age in the ophiolites of Volvi–Eastern Rhodope terrane H8. An ophiolite obduction model can be applied, with the ophiolitic nappes always emplaced on top of pre-Alpine continental terranes with Mesozoic shallow-water carbonate platforms. The geometry of the continental terranes drifting during the Mesozoic within the Tethys Ocean controls the number and dimensions of the Tethyan oceanic basins. Where a continental terrane dies out, the two adjacent oceanic basins merge into one larger basin. This seems to be the case of the Pelagonian terrane (H3), which is terminated north of Skopje, where the Pindos oceanic basin (H2) merges with the Vardar/Axios oceanic basin (H4). © 2008 Elsevier B.V. All rights reserved.

1. Introduction: the ophiolite belts and the Tethyan oceanic basins of the Hellenides The Hellenides represent a characteristic segment of Tethyan paleogeography and have been used as a model for the development of geosynclines in early geotectonic syntheses (e.g. Aubouin, 1965). Their paleogeographic organization has been rather complex with alternation of ridges, comprising shallow-water carbonate platforms and furrows with pelagic sediments and volcano-sedimentary formations including mafic rocks. The presence of ophiolites has been considered during the “pre-plate tectonics” period as the documentation of the axial zones of the eugeosynclinal area on both sides of the Pelagonian metamorphic massif in the so-called Sub-Pelagonian zone along the external western margin and the Vardar/Axios zone along the internal eastern margin. Nevertheless, the Vardar/Axios ophiolites have been considered as the main axial zone dividing the Hellenides/ Dinarides orogenic system to the west from the Alpidic system of the Balkanides and the Carpathians in the east (e.g. Kossmat, 1924). Since the development of the plate tectonics theory and the interpretation of the ophiolites as ancient oceanic crust the Hellenic ophiolites have been re-evaluated as major paleogeographic elements dividing oceanic basins within Tethys (e.g. Smith, 1971; Dewey et al., 1973). The complexity of Tethys with more than one ophiolite belt

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occurring within the pelagic sediments, alternating with shallowwater carbonate platforms was evident from the beginning of the application of the new theory and it was obvious that the simple Atlantic type oceanic model with two passive continental margins could not be accepted in the eastern Mediterranean. Several tectonic and paleogeographic models have been proposed with a different number of oceanic basins along the belt and also different types of geodynamic settings such as ophiolites formed within spreading centers along a mid-ocean ridge or above supra-subduction zones etc. (for extensive reviews and discussion see Robertson and Dixon (1984), Smith (1993), Robertson (2002, 2004), and Smith and Rassios (2003). The concept of tectonostratigraphic terranes has been applied in the eastern Mediterranean aiming to locate the micro-continents existing within the Tethyan belt. Pre-Alpine continental crust of Precambrian and/or Paleozoic age covered by Mesozoic–early Cenozoic shallow-water carbonate platforms occurring within the Tethyan belt form the continental terranes. The ophiolites and related pelagic sediments of the intermediate oceanic basins separating the microcontinents have been sutured and their obducted relics of oceanic crust form the evidence for the oceanic terranes. Both continental and oceanic terranes have been analyzed and mapped at 1/2500 000 scale during 1987–1995 within IGCP project 276 (Papanikolaou and Sassi, 1989; Papanikolaou and Ebner, 1997). The timing of tectonic emplacement of the ophiolites is a characteristic feature of each oceanic basin that may be related to its closure. Additionally, the timing helps to correlate major tectonic

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events along neighboring segments of the Tethyan belt (e.g. Dinarides/ Hellenides–Pontides/Taurides) by correlating the timing of formation of ophiolite suture zones and of the closure of adjacent oceanic basins. The available data on the timing of tectonic emplacement for each oceanic terrane in the Hellenides are presented in this paper, following their relative position from south to north, together with a discussion of the long-lasting processes of alternative periods of plate convergence and (micro-) collision between Eurasia and Africa. 2. The ophiolites within the terrane tectonostratigraphy of the Hellenides The paleogeographic organization of the Hellenides comprises four oceanic and five continental tectonostratigraphic terranes (Papanikolaou, 1989a, 1997; Papanikolaou et al., 2004) (Fig. 1). The symbols H1–H9 characterize each of the terranes of the Hellenides (H) and the odd numbers correspond to the continental terranes (H1, H3, H5, H7,

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H9) whereas the even numbers to the oceanic terranes (H2, H4, H6, H8). The present day remnant of Tethys Ocean in the Eastern Mediterranean, which is being subducted below the Hellenic arc during the past few millions years, is the southernmost Tethyan oceanic basin adjacent to the African passive margin that can be regarded as a future sutured terrane (H0). The four oceanic terranes of the Hellenides are: Pindos– Cyclades H2, Vardar/Axios H4, Lesvos–Circum Rhodope H6 and Volvi– Eastern Rhodope H8. They form ophiolite suture zones within the Hellenide nappe structure and represent allochthonous fragments of former oceanic basins that closed throughout Mesozoic and Cenozoic time. The distinction and location of the ophiolite outcrops of Greece for the four oceanic terranes is given on the map of Fig. 2. The classification of some ophiolite outcrops is not certain in cases where the available information is not conclusive due to lack of age determination, cover by recent sediments, younger tectonism destroying primary structures etc. Continental blocks with pre-Alpine basement rocks covered by thick shallow-water carbonate platforms of Mesozoic to lower

Fig. 1. Terrane map of the Hellenides (after Papanikolaou, 1997, modified).

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Fig. 2. The main ophiolite outcrops of the Hellenides are distinguished in four ophiolite belts corresponding to the oceanic terranes of H2, H4, H6 and H8 (after Papanikolaou, 1989a, modified). The location of the geological maps describing the tectonic structure of each terrane is shown in quadrangles (Figs. 4, 5, 6 for H2 ; Figs. 7, 8 for H4 ; Fig. 9 for H6 ; Fig. 10 for H8).

Cenozoic age occur in between the oceanic terranes. These continental terranes are: external carbonate platform of the Hellenides H1, internal carbonate platform of the Hellenides H3, Paikon–Lesvos autochthon H5, Rhodope autochthon (Pangeon–Kerdylia) H7 and Rhodope allochthon (Sidironero–Vertiskos) H9 (Fig. 1). The tectonic boundaries, the tectonostratigraphy and the criteria for the identification of the Hellenic terranes have been analyzed by Papanikolaou (1989a, 1997) and Papanikolaou et al. (2004). The general characteristic is an older history of orogenic events recorded in the terranes of the inner part of the Hellenic arc with a stepwise migration of younger events recorded in those of the external part. This implies that all tectonostratigraphic formations and geodynamic events are younger in the external part of the arc in the Ionian islands, Crete and the Dodekanese islands becoming older towards the central core in Rhodope. Thus, synorogenic flysch deposits are late Tertiary in the more external units, early Tertiary in the middle and late Cretaceous in the internal units. Several flysch type formations of Jurassic age are also found in several internal units. The end of shallowwater carbonate platform sedimentation follows the same pattern

with Eocene–Oligocene limestones occurring in the external platform (H1), Late Cretaceous limestones in the western part of the internal platform (Parnassos unit) and Late Jurassic–early Cretaceous limestones in the main part of the internal platform (H3), late Triassic– Liassic limestones in the Lesvos autochthon platform (H5) and older than Cretaceous marbles in the Pangeon unit (H7). The same pattern of younger ages is observed in the ophiolites and related pelagic sequences of the oceanic terranes. Thus, pelagic sedimentation ends in late Cretaceous–early Tertiary in the Pindos, in the Late Jurassic in the Maliac, and in the late Triassic-Lias in Lesvos allochthon and Circum Rhodope (H6). Deformation, metamorphism and magmatic/volcanic arc activity correspond to the orogenic migration with late Tertiary events in the external tectonometamorphic belt (Crete–Peloponnesus, involving terranes H1 and H2), early Tertiary events in the medial belt (Olympus–Cyclades, involving terranes H1, H2, H3 and H4) and late Mesozoic in the internal belt (Rhodope s.l., involving terranes H6, H7, H8 and H9) (Papanikolaou, 1984, 1997; Papanikolaou et al., 2004). The overall tectonic polarity of the Hellenides involved constantly northward subduction of the Tethyan paleogeographic elements since

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the Jurassic, which resulted in the closure of the Tethyan oceanic basins and the accretion of the intermediate continental blocks (Papanikolaou, 1989a, 1993, 1997). This is reflected in the development of active margins on the northern margins of the oceanic terranes and passive margins along their southern margins (Papanikolaou, 1989a; Stampfli et al., 1991). Tectonic emplacement of the ophiolites over the adjacent continental terranes was generally directed to the south and took place during the last stages of the closure of each oceanic basin, representing successive phases of ophiolite obduction. This general concept of westward migration of the subduction zones towards the external part of the evolving orogenic arc of the Hellenides does not necessarily imply that the direction of tectonic emplacement of the ophiolites was always towards the west (following the present day orientation of the arc in continental Greece). Several cases with opposite sense of ophiolite obduction have been reported for the late Jurassic tectonism (Hynes et al., 1972, in Orthrys; Vergely, 1976, in Vermion Mt, northern Greece, Robertson et al., 1991, in the SubPelagonian, Rassios and Smith, 2000; Rassios and Moores, 2006a,b, in Vourinos, Robertson and Shallo, 2000; Dilek et al., 2005, in Albania). The eastward tectonic emplacement is in several cases deduced from the present day westward dip of the ophiolite nappe and in some cases from microstructural observations within the ophiolite nappe (usually the peridotites) and not from combined structural analysis of the paleo-Alpine tectonic units involving also the deformation below the ophiolite nappes. However, this eastward direction of tectonic emplacement is opposed to the contemporaneous westward direction of tectonic emplacement of the Vardar/Axios ophiolites (Mercier,

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1968; Vergely, 1976; Aubouin, 1977; Papanikolaou, 1989a). A discussion of the tectonic emplacement of ophiolites onto both margins of the Pelagonian platform from the Pindos ocean in the west and the Vardar/Almopia ocean in the east has been recently presented by Sharp and Robertson (2006). The simple tectonic model implies that the Almopia ophiolites in the east are the same unit with the Pelagonian ophiolites in the west (e.g.Vourinos) both representing parts of the same ophiolite obduction of the Vardar/Axios ocean over the Pelagonian platform towards the west (Papanikolaou, 1984, 1989a, 1997). The paleotectonised areas with ophiolite nappes on the SubPelagonian and Almopia units have been repeatedly deformed and rotated throughout the Cenozoic and thus, primary tectonic orientations may differ considerably on a local scale (see later tectonic orientations and directions of tectonic transport in eastern Sterea, western Macedonia, Lesvos). 3. Timing of tectonic emplacement of the ophiolites in the Hellenides The timing of tectonic emplacement of the ophiolites is based on criteria that include: a) The dating of the post-emplacement sedimentary cover which seals the tectonic contacts between the ophiolites and the underlying geological formations. This is usually manifested by the stratigraphic unconformities characterizing each tectonic event. Especially in those cases where the upper crust has been denudated the dating of the post-emplacement intrusive rocks may also provide valuable age constraints. b) The dating of the syn-

Fig. 3. (a) Schematic geological profile across the Northern Pindos ophiolites (H2) from Konitsa to Grevena, showing the Eocene timing of tectonic emplacement on the External Hellenides Platform (H1) (after Papanikolaou, 1986a, modified). (b) Schematic representation of the upper Cretaceous transgression over the internal Hellenides, showing the timing of tectonic emplacement of the H4 ophiolites and related pelagic sequences of the Maliac unit over the units of the internal carbonate platform H3 in late Jurassic–early Cretaceous (after Papanikolaou, 1986a, modified). The Sub-Pelagonian unit is not metamorphosed whereas the Almopia unit is metamorphosed together with the Paleozoic basement of the Flambouron and Kastoria units.

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emplacement tectonostratigraphic formations developed along the front of the advancing ophiolite nappes, producing flysch type deposits with ophiolite detritus and mélanges. c) The dating of the youngest geological formations involved in the structure of the allochthon and the autochthon units. On the basis of the above criteria the major distinction of the early Tertiary tectonic emplacement of the Pindos–Cyclades ophiolites (Fig. 3a) from the late Jurassic–early Cretaceous emplacement of the Vardar/Axios ophiolites (Fig. 3b) was made by Papanikolaou (1986a, 1989a, 1997). Confusion and misunderstanding result quite often result from the mixing of the age of the ophiolites with the age of their tectonic emplacement. Thus, Jurassic ophiolites tectonically emplaced in late Jurassic may be reactivated also in late Eocene time. Additionally, the documentation of a late Eocene age of tectonic emplacement of an ophiolite does not imply the age of the ophiolite which may be of Triassic or Cretaceous time. Another problem arises from the fact that an age of an ophiolite determined usually indirectly by its amphibolitic metamorphic sole is considered generally as dating the whole oceanic crust and not a particular segment of the ocean floor. In several cases the dating of an oceanic basin is restricted to the ages obtained by the ophiolite complexes neglecting the information that can be obtained by the pelagic sedimentary sequences of the oceanic basins and their margins. The available radiometric dates of ophiolites sometimes are restricted to a narrow range of 5–10 myr (e.g. 169– 172 myr for the western Mirdita ophiolites and 163–172 myr for the eastern Mirdita ophiolites of Albania according to Dilek et al., 2008). This is sometimes considered as the age of the entire oceanic basin instead of the obvious assumption that it may represent only the age of a small strip of the oceanic basin, which has been obducted whereas the rest of the oceanic stripes of the ocean floor covering a major time span of 30 or 50 or 100 myr and more has been subducted before the final obduction during the closure of the basin. Thus, in this paper ages obtained from the ophiolites are taken less into account than the ages obtained from the surrounding paleogeographic formations and the ages obtained by the synchronous geodynamic events. The present day oceanic remnant of the Eastern Mediterranean (H0) is in the stage of subduction beneath the Hellenic arc during only the last few millions of years (late Miocene–Pliocene) just before the final collisional stage with Africa (Finetti et al., 1990; Mascle and Chaumillon, 1998). The present rate of convergence of the Hellenic arc towards Africa (~40 mm/yr) as measured by GPS data (e.g. McClusky et al., 2000) implies that within 8–10 million years the Ionian basin of the East Mediterranean will collide with the Cyrenaica promontory of Africa in Libya and ophiolite obduction may take place towards the south onto the continental margin. This would be similar to what occurred in the Turkish/Syrian border during early Tertiary (e.g. Al-Riyami et al., 2002; Barazangi et al., 2006) during the collision of the Anatolian segment with Arabia. 3.1. External ophiolite belt of the Pindos–Cyclades tectonostratigraphic terrane (H2) The ophiolites of the Pindos–Cyclades oceanic basin (H2) form the external ophiolite belt of the Hellenides, which is observed from the Greek/Albanian border in the northern Pindos Mt and Koziakas Mt through Argolis peninsula in eastern Peloponnesus, into central Crete and the Dodekanese islands in Karpathos and Rhodes (Papanikolaou, 1989a; 1997) (Fig. 2). The tectonic emplacement of the Pindos– Cyclades oceanic basin (H2) over the External Carbonate Platform of the Hellenides (H1) occurred in early Cenozoic time (Papanikolaou, 1989a). Characteristic tectonic profiles can be observed for the Northern Pindos ophiolites (e.g. Brunn, 1956), which overlie the Paleocene–Eocene flysch of the Pindos unit (Figs. 3a and 4). The ophiolite nappe is covered unconformably by molasse-type sediments of the Mesohellenic basin of Oligocene–Miocene age. The Krania Formation of Eocene age (Brunn, 1956) occurring only in a small part

of the western margin of the molassic basin is syntectonic, with folds and thrusts showing a compressive deformation similar to that of the ophiolites and the Pindos flysch. On the contrary, the Oligocene Eptachorion formation has no compressive deformation and covers unconformably the western margin of the molassic basin from the area west of Kalabaka to the north Greek border with Albania (Papanikolaou et al., 1988). Numerous outcrops of upper Cretaceous– Paleocene pelagic limestones are observed within the upper part of the Pindos flysch below the basal thrust of the ophiolite nappe (Fig. 4). Some outcrops of pelagic limestones are also observed within the ultramafic rocks of the basal parts of the ophiolite nappe, as seen in some spectacular tectonic lenses on the road from Kalabaka to Metsovon near Orthovouni. Another outcrop concerns an exotic block at Tragopetra at about 2 km before the Katara pass on the road from Kalabaka to Metsovon exhibiting Upper Cretaceous limestones, red Paleocene pelites and Eocene conglomerates with ophiolite detritus. The stratigraphic data and facies of these blocks observed within the upper levels of the Pindos flysch and at the base of the ophiolite nappe indicate their provenance from the tectonic unit of Western Thessaly/ Beotia (Thymiama facies of Upper Cretaceous pelagic limestones and red Paleocene pelites at the base of the Eocene flysch (Papanikolaou and Sideris, 1979). The Pindos flysch and the ophiolite nappe are thrusted as a composite nappe to the west on the upper Eocene– Oligocene flysch of the Ionian unit (Fig. 3a). The Ionian flysch towards its upper part presents the characteristics of wildflysch with olisthostromes and abundant olistoliths, breccia and conglomerates. Thus, the timing of tectonic emplacement of the external ophiolite belt (H2) in Northern Pindos is middle–late Eocene. A similar age of tectonic emplacement is observed in several other outcrops of the H2 ophiolites such as in Crete, where the ophiolite nappe overlies the Eocene flysch of the Pindos/Ethia unit and is covered by post-Alpine sediments of middle–upper Miocene age (Fig. 5) (e.g. Bonneau, 1984). In this area located on the southern slopes of Psiloritis Mt a major east–west trending and south-dipping extensional detachment fault separates the relative autochthon unit of Mani (known also as Plattenkalk or metamorphic Ionian unit) in its footwall from the topmost nappe of the ophiolites and the Asteroussia metamorphic rocks in its hanging wall. The structural omission is 8–10 km of tectonostratigraphic units involving the Tripolis unit (together with its Permo-triassic basal formations) and the Pindos/Ethia unit together with some small outliers of local units (Arvi, Miamou, Vatos etc.). The upper nappes bearing the ophiolite outcrops at their topmost position are overthrusted onto the upper Eocene–Oligocene flysch of the Pindos/Ethia unit, which is overthrusted together with the underlying nappes onto the slightly metamorphosed Oligocene flysch of the relative autochthon Mani unit. Thus, the period of tectonic emplacement of the external ophiolites in Crete was between the late Eocene and late Oligocene. In Crete it is interesting to note the occurrence of the Arvi unit, bearing Upper Cretaceous basalts of MORB affinities (Robert and Bonneau, 1982), in between the underlying Pindos/Ethia pelagic sequence and the overlying composite nappe of the Asteroussia metamorphics and the ophiolites. This is because Arvi unit may represent pieces of the last oceanic crust formed within H2 before its subduction and closure in the Eocene. Outcrops of the Arvi unit with the characteristic upper Cretaceous–Paleocene basalts are known also both from Peloponnesus at Aderes Mt of southeastern Argolis peninsula (Papanikolaou, 1989b) and at Parnon Mt around Aghios Vassilios and from central Sterea along the frontal parts of the Vardoussia chain in Kerassia–Milia (Fig. 2) (see also Robertson and Degnan, 1992). Geochronologic data from the H2 ophiolites of Karpathos and Rhodes, whose tectonic position is similar to that of the ophiolites in Crete, have yielded a late Cretaceous age of around 90 myr (Koepke et al., 2002). In the Cycladic metamorphics the H2 ophiolites are involved in the structure of the blueschists, whose metamorphism has been

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Fig. 4. Geological map and cross section of the Northern Pindos ophiolite nappe at its eastern outcrops north of Koziakas Mt. (based on the geological maps of IGME at scale 1/50 000, Koumantakis, 1980). The timing of tectonic emplacement of the H2 ophiolites is shown by the Eocene age of the underlying Pindos flysch and the Oligocene age of the transgressively overlying molasse of the Mesohellenic basin.

dated as Eocene–Oligocene (e.g. Schliestedt et al., 1987). Two distinct tectonic zones of ophiolites occur within the Cycladic blueschists: a) A lower zone observed at the base of the blueshists nappe of the Northern Cyclades over the relative autochthon platform of Almyropotamos in Southern Evia and over the equivalent platform of Kerketefs in Western Samos. b) An upper zone at the top of the Northern Cyclades unit underlying either the Makrotandalon–Ochi blueschist unit (Southern Evia and Northern Andros) or the Southern Cyclades unit in Samos (Papanikolaou, 1986b, 1987). The diversity of the ophiolites found interlayered within the Cycladic blueschists was shown by analytical data from southern Evia, Tinos, Syros and Naxos (Katzir et al., 2007). A characteristic example of the lower ophiolite zone occurs in southern Evia, where outcrops of serpentinised

ultramafics are embedded within amphibolites and amphibolitic schists, which represent meta-basic tuffs and volcanics (Fig. 6). These mafic and ultramafic rocks are found at the base of the pelagic sequence of the Northern Cyclades nappe (Papanikolaou, 1984, 1986b, 1987) locally also known as the Styra unit (Katsikatsos, 1979), which includes at its higher horizons cipolinic marbles and calcschists with intercalations of mica-schists and meta-sandstones. The relative autochthon of the Almyropotamos unit lies below the Northern Cyclades nappe. The Eocene flysch on top of the Almyropotamos shallow-water carbonate platform (Dubois and Bignot, 1979) dates the tectonic emplacement of the ophiolitic and related meta-pelagic rocks of H2 in the Late Eocene–Oligocene. Upper Miocene continental sediments with fossils of mammals seal the

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Fig. 5. Geological map and cross section of central southern Crete, showing the ophiolite nappe on top of the Asteroussia metamorphics nappe and the underlying Pindos/Ethia nappe (based on the geological maps of IGME at scale 1/50 000, Bonneau et al., 1984). These higher nappes occur on the hanging wall of an E–W extensional detachment fault whereas at its footwall the relative autochthon of Mani (metamorphosed Ionian) crops out. Middle–upper Miocene sediments cover the primary Eocene–Oligocene thrusts.

tectonic contact between the two units in the area of Almyropotamos (Fig. 6) but lower–middle Miocene sediments with lignite deposits are known from Aliveri in central Evia (Katsikatsos et al., 1981) and thus, the timing of tectonic emplacement is established between late Eocene–early Miocene. In conclusion, the ophiolites of the Pindos–Cyclades oceanic terrane H2 have been emplaced over different parts of the External carbonate platform of the Hellenides terrane H1 during late Eocene– early Oligocene. It is noteworthy that structures of the upper crust are observed in Pindos, Argolis, Crete and Dodekanese islands with nonmetamorphic units related to the ophiolite emplacement and structures of the lower crust are observed in the tectonic windows of the Cyclades where the ophiolites are metamorphosed and emplaced on blueschist type metamorphic rocks.

3.2. Internal ophiolite belt of the Vardar/Axios tectonostratigraphic terrane (H4) The ophiolites of the Vardar/Axios oceanic basin form a second ophiolite belt observed within the units of the internal Hellenides (Papanikolaou, 1989a, 1997) (Fig. 2). This internal ophiolite belt (H4) dominates in the ophiolite outcrops observed in eastern continental Greece such as the Almopia, Vourinos, Orthrys, Kallidromon, Central Evia ophiolites. The tectonic emplacement of the Vardar/Axios oceanic basin (H4) over the Internal Carbonate Platform of the Hellenides (H3) occurred in Late Jurassic–Early Cretaceous (e.g. Aubouin, 1977). This is observed in several outcrops of eastern Sterea and Thessaly, where an unconformity of upper Cretaceous (usually Cenomanian) is present. This unconformity has been used as a criterion for the distinction of

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Fig. 6. Geological map and cross section of the area between Almyropotamos and Styra in southern Evia, showing the occurrence of the H2 ophiolitic rocks at the base of the Styra/ Northern Cyclades blueschist nappe on top of the Mesozoic–Eocene Almyropotamos relative autochthon unit (H1) (based on the geological maps of IGME at scale 1/50 000, Katsikatsos, 1990).

the Internal Hellenides from the External Hellenides (Renz, 1955; Brunn, 1956). The Vardar/Axios oceanic basin is observed below the unconformity with the ophiolite nappe and the underlying Maliac unit, first described in eastern Orthys Mt (Ferriere, 1976), which comprises a pelagic volcano-sedimentary sequence of early Triassic– early Cretaceous age (Fig. 3b). In the area of eastern Sterea between the villages Tragana and Pavlos there are extensive outcrops of ophiolites emplaced over the Triassic–Jurassic formations of the Sub-Pelagonian unit (Fig. 7). The stratigraphy of the Sub-Pelagonian unit comprises a thick shallowwater carbonate platform of early Triassic–middle Jurassic age overlain by the late Jurassic schist-hornstein formation. The tectonic contact between the two paleotectonised tectonic units is covered by the unconformably overlying upper Cretaceous strata. The transgres-

sive sequence comprises a basal conglomerate of Cenomanian age, neritic limestones with rudists of Turonian–Senonian, transitional pelagic limestones with Globotruncanas of Maastrichtian–Danian and Paleocene–Eocene flysch. The Alpine deformation affects the transgressive Cretaceous–Eocene sequence creating thrusts and folds and also the paleotectonised units by refolding the previous structures of late Jurassic–early Cretaceous age. Deformation pattern, intensity of deformation and direction of folding differ in these two orogenic events. Thus, nearly isoclinal folds are observed in the upper levels of the Sub-Pelagonian Jurassic sequence resulting to a repetition of the “schist-hornstein” formation. The direction of folding and thrusting below the unconformity in this area is east–west to eastnortheast– westsouthwest. Instead, the upper Cretaceous–Eocene sequence above the unconformity shows only open folding and thrusting in

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Fig. 7. Geological map and cross section of the area of Eastern Chlomon Mt, where the ophiolites of H4 are tectonically emplaced on the Sub-Pelagonian unit of H3 (based on the geological map of IGME at scale 1/50 000, Maratos, 1965). The Cenomanian unconformity seals the tectonic contact and the paleo-Alpine structures.

the northeast–southwest direction with tectonic transport towards the northwest (Fig. 7). This Alpine structural trend is almost perpendicular to the usual Alpine trend which is in the northwest– southeast direction with tectonic transport towards the southwest. The reason is probably post-Eocene rotations occurring in neotectonic blocks of central continental Greece. The age of the unconformity sealing the ophiolite nappe may be older than Cenomanian, as in the case of Paros Island in the Cyclades,

where it is reported as Barremian (Papageorgakis, 1968; Papanikolaou, 1980). The presence of ophiolite detritus in the “schist-hornstein” formations of the Sub-Pelagonian unit (Renz, 1955; Papanikolaou, 1990) and in the upper Jurassic–lower Cretaceous flysch of the Beotian/ Western Thessaly unit (Celet et al.,1976; Papanikolaou and Sideris,1979) indicates that ophiolite obduction of H4 took place since late Jurassic. Some small outcrops of late Jurassic sediments have been reported below the Cretaceous transgressive sediments in the western

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Pelagonian margin but also in Almopia along the eastern Pelagonian margin (Sharp and Robertson, 2006). Thus, the duration of tectonic emplacement of the H4 ophiolites seems to extend throughout late Jurassic–early Cretaceous over a considerable time span of about 40 myr. The ophiolites observed in northern Greece along the internal ophiolite belt are emplaced over metamorphosed platform sediments belonging to the Almopia Unit (or Metamorphic Pelagonian), whereas those cropping out in the southern parts of the belt over Sterea and Evia are emplaced over non-metamorphic platform sediments belonging to the Sub-Pelagonian Unit (Fig. 3b). The transgressive Upper Cretaceous– Eocene sediments are observed on top of the ophiolite nappe, the

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pelagic Maliac sequence and the underlying carbonate platform either metamorphosed (Almopia) or not (Sub-Pelagonian). In northern Greece the distinction between the paleo-Alpine and the alpine structures is more important, because the paleo-Alpine orogeny has involved tectono-metamorphic events and the transgressive Cretaceous sediments are found on top of metamorphic rocks. Thus, in the area of Kozani the marbles of the metamorphosed Triassic–Jurassic carbonate platform and the overlying ophiolite nappe have been intensively folded with isoclinal folds (Vergely, 1976). This is illustrated by the isoclinal synforms formed by the ophiolite outcrops within the Almopia marbles (Fig. 8). The general structural trend of this paleo-Alpine structure in this area is

Fig. 8. a) Geological map of the Aghios Dimitrios area north of Kozani at the southern slopes of Vermion Mt (based on the geological map of IGME at scale 1/50 000, Anastopoulos et al., 1980). The ophiolites of H4 are observed tectonically overlying the Almopia metamorphosed carbonates and the primary tectonic contact is isoclinally folded below the Cenomanian unconformity. A tertiary overthrust has duplicated the upper sequence (Cretaceous) and has created the Vermion nappe. b) Geological section through Southern Vermion Mt showing the contrast in structural style of the non-metamorphosed Cretaceous cover and the metamorphosed units of Almopia (H3) and Axios ophiolites (H4).

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east–west and the axial planes of the folds dip to the north indicating a southward tectonic emplacement of the ophiolites within a deep ductile shear zone. Alpine compressive deformation of the overall structure during Middle–Late Eocene is observed, with decollement and internal thrusting of the Upper Cretaceous–Eocene sequence forming the Vermion nappe (Brunn, 1956). Thus, Alpine orogeny has duplicated the transgressive sequence on top of the paleo-Alpine structures of the isoclinally folded Almopia metamorphosed sequence and ophiolite nappe (Fig. 8). In conclusion, the timing of tectonic emplacement of the Vardar/ Axios ophiolites of terrane H4 is late Jurassic–early Cretaceous. It is noteworthy that an area of upper crust has been preserved in Orthris and Sterea–Evia (non-metamorphic Sub-Pelagonian) together with an area of lower crust in Thessaly and western Macedonia (metamorphic Pelagonian/Almopia). 3.3. Ophiolites of Lesvos–Circum Rhodope belt (H6) The ophiolites belonging to the Lesvos–Circum Rhodope oceanic basin are found in several dispersed outcrops in central Macedonia, western Thrace and north Aegean islands (Papanikolaou, 1989a, 1997) (Fig. 2). The tectonic emplacement of the ophiolites of the Lesvos– Circum Rhodope oceanic basin is poorly dated because there are not transgressive sediments post-dating the tectonic emplacement of the ophiolites. Additionally, the age of the tectonically underlying metamorphosed sediments is not easily determined. In Lesvos, the ophiolites of terrane H6 are found tectonically emplaced over the continental terrane H5, which comprises a shallow-

water carbonate platform of Upper Paleozoic–Upper Triassic age (Fig. 9a) (Papanikolaou, 1999). Characteristic neritic facies with Productus sp. have been described at the lower horizons of the carbonate platform (Hecht, 1972; Katsikatsos et al., 1986). However, the presence of Megalodon sp. near the village Moria at the upper horizons of the carbonate platform within an ancient marble quarry, indicates an upper Triassic or (?) Liassic age for the top of the marbles. The phyllitic rocks overlying the Upper Triassic marbles probably represent a Liassic flysch (Fig. 9b). The presence of these upper Triassic–lower Jurassic sediments at the top of the shallow-water carbonate platform of the Lesvos autochthon beneath the ophiolite nappe indicates a post-Liassic age of tectonic emplacement. The Lesvos allochthon comprises two different lithologies, which are genetically and stratigraphically linked: the ophiolite complex and a volcano-sedimentary pelagic sequence (Fig. 9b). In the area of Amali peninsula in southeast Lesvos the sediments are observed in stratigraphic succession overlying the ophiolite complex and the basaltic metalavas, whereas in the area of Komi in central Lesvos the volcano-sedimentary sequence is inverted (Fig. 9c). The Triassic age of the metasediments documented by conodonts in the area of Komi within the Lesvos allochthon confirms an early Jurassic age of the obduction and indicates that the tectonic event may be related to the Cimmeride orogeny (Papanikolaou, 1999) documented in the neighboring outcrops of northwestern Minor Asia (Sengor et al., 1984). The asymmetry of the tectonic structures within the geological formations of terrane H6 indicates a direction of tectonic transport towards the southeast. The ophiolites of Lesvos comprise ultramafic rocks with minor gabbroic dikes and an amphibolite sole (Migiros et al., 2000).

Fig. 9. a) Simplified geological map of southern Lesvos, showing the tectonic position of the Lesvos ophiolites within the Lesvos allochthon (H6) over the Permo-Triassic carbonate platform of the Lesvos autochthon (H5). b) Stratigraphic columns of the two terranes and c) Tectonic profile (based on data by Hecht, 1972; Katsikatsos et al., 1986; Papanikolaou, 1999).

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They show a metamorphic degree of 6–7 kbars and a temperature range of 1100–650 °C with two phases of rodingitisation whereas their chemical affinities indicate a marginal basin regime (Migiros et al., 2000; Hadzipanagiotou et al., 2003). An age of 153–158 myr has been determined by Hatzipanagiotou and Pe-Piper (1995) in amphibolites at the base of the Lesvos ophiolites from the area of Vatera. Unfortunately, no geological formations older than the lower Miocene volcanics of northern Lesbos exist and thus, the timing remains quite uncertain from Lias to Early Miocene. However, the upper Liassic unconformity is observed both in the allochthon unit of the neighboring island of Chios (Besenecker et al., 1968) and in the Cimmerides in Turkey (Sengor, 1984). The Circum Rhodope belt, sensu Kauffman et al. (1976), is a complex unit incorporating both neritic and pelagic sequences, belonging to different paleogeographic environments of Triassic to lower– middle (?) Jurassic age. A Triassic shallow-water carbonate platform, as the outcrops around Doubia in Chalkidiki, represents a distinct lower unit, over which pelagic sequences with siliceous sediments and volcano-sedimentary formations are observed, together with

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basic and ultrabasic rocks, representing a dismembered ophiolite suite. Thus, the lower unit occurring within the Circum Rhodope belt, which comprises metamorphosed shallow-water Triassic carbonates may belong to the continental terrane H5, equivalent of the Lesvos autochthon and Paikon. Ophiolite outcrops occur in Chalkidiki, Evros and Samothraki (Kockel et al., 1977; Magganas, 2002; Tsikouras, 1994). Flysch type sediments with turbidites in the Upper Triassic and Liassic are known from several localities (e.g. the Melissochori Formation, Mercier, 1968). A marginal basin-volcanic arc origin of the meta-basic rocks cropping out at the eastern part of the Circum Rhodope belt has been proposed (Magganas et al., 1991; Magganas, 2002). Blueschistfacies assemblages have been reported within some tectonic layers of the belt in central Chalkidiki, possibly representing an Eo-Hellenic HP/LT metamorphic event (Michard et al., 1994). The exact timing of ophiolite obduction remains questionable. In Chalkidiki peninsula some outcrops of upper Jurassic carbonate sediments and molasse occurring on top of ophiolitic rocks and other formations of the belt (Kockel et al., 1977) indicate a pre-upper Jurassic tectonic emplacement of the H6 ophiolites. Overall it is true that the distinction

Fig. 10. Simplified geological map of the Chalkidiki area, showing the Volvi ophiolites (H8) at the base of the Vertiskos basement nappe (H9) on top of the underlying Kerdylia relative autochthon (H7) (based on Kockel et al., 1977). The late Jurassic granites intruding the Vertiskos gneisses and the Volvi ophiolites indicate a pre-late Jurassic tectonic emplacement.

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between the Circum Rhodope units and subunits from those of the Peonia units and subunits occurring towards the west is rather difficult, due to the Alpine tectonic imbrication that confuses the overall geometry of the early Mesozoic orogenic events. Thus, pre-late Jurassic events of the Peonia and associated units related to the Vardar/Axios domain may be easily confused with late Triassic–early Jurassic events of the Cimmeride events possibly occurring in the Circum Rhodope domain. In conclusion, the timing of tectonic emplacement of the ophiolites of terrane H6 is between late Lias and late Malm. Thus, it may correspond to the Cimmerides but also partly to the paleoAlpine orogenic event involving also the tectonic emplacement of the Vardar/Axios ophiolites. The geological formations and the structures preserved belong to the lower crust, with low-grade metamorphic rocks occurring within the outcrops of the oceanic terrane and below its nappe structure over the relative autochthon units of terrane H5.

3.4. Ophiolites of Volvi–Eastern Rhodope (H8) The ophiolites of Volvi–Eastern Rhodope oceanic basin crop out in two different areas of the internal tectonometamoprhic belt: in the Serbo-Macedonian massif in Chalkidiki and in the eastern parts of the Rhodope massif around Komotini (Papanikolaou, 1989a, 1997) (Fig. 2). Both areas lie within the internal tectono-metamorphic belt of Rhodope (sensu lato) and comprise metamorphosed ophiolitic rocks, whose protoliths are not always easy to determine. Thus, besides the serpentinised peridotites representing the ultramafic part of the ophiolite suite, which are usually easily recognized, the remainder of the mafic igneous assemblages of the ophiolite complex, like gabbros, diabases and basalts, have been transformed and metamorphosed to amphibolites, amphibolite gneisses and amphibolite schists with only some textural relics of the primary igneous rocks. The Volvi ophiolites (Dixon and Dimitriadis, 1984) are observed at the base of the Vertiskos basement nappe, which overlies the Kerdylia

Fig. 11. a) Schematic representation of the paleogeographic organization of the terranes in the Hellenides with reference on the location of the main tectonic units of the Hellenides belonging to each terrane (after Papanikolaou 1997, modified). b) Simplified paleogeographic sketch of the terranes in the Hellenides with extension towards the terrane geometry of the Dinarides (based on Karamata et al., 1997). It should be noticed that: i) Oceanic basins H2 and H4 merge north of the Pelagonian terrane H3. ii) Oceanic basins H4 and H6 merge north of the Paikon block of H5. iii) Oceanic basins H6 and H8 merge on both sides of the Rhodope terrane (Pangeon–Kerdylia) H7. Thus, Tethys Ocean comprised all the area of oceanic basins H2–H4–H6–H8, among which continental terranes of various dimensions drifted northwards from Gondwana to Eurasia.

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reported (Kozhoukharova, 1980; Kozhoukharov et al., 1988). Eclogites and relics of HP/LT metamorphic assemblages have been also reported in the southern parts of Rhodope in Greece (Liati and Mposkos, 1990; Liati and Seidel, 1996). The overall tectonic position of the Volvi–eastern Rhodope ophiolites (H8) is a tectonic wedge between the underlying Pangeon and Kerdylia basement units, which carry a shallow-water carbonate platform on top of a gneissic core (H7) and the overlying basement nappes of the Vertiskos and Sidironero units (H9). The age of the tectonic emplacement of the H8 ophiolites along with the Vertiskos and Sidironero basement nappes of H9 on the H7 relative autochthon of Pangeon and Kerdylia remains uncertain. The strong Cretaceous tectono-metamorphic event detected in the Rhodopean units (Wawrzenitz and Mposkos, 1997) indicates a pre-Cretaceous age of tectonic emplacement without other age constraints. A Paleotethyan origin of these ophiolites has been proposed, involving a late Triassic–Liassic age of tectonic emplacement based only on correlations with the welldated outcrops of northwest Turkey within the Cimmerides (Sengor et al., 1984). The late Jurassic intrusion of granitoids, like the previously reported Arnea granite, indicates a pre-late Jurassic emplacement. In conclusion, the tectonic emplacement of the Volvi–eastern Rhodope ophiolites of terrane H6 has occurred before the late Jurassic. The preserved structures related to the ophiolite emplacement and the adjacent tectonometamorphic units belong to the lower crust without any evidence from sedimentary sequences overlying the emplacement structures. 4. Discussion

Fig. 11 (continued).

unit (Kockel et al., 1977) (Fig. 10). The ophiolite and meta-ophiolite rocks of Volvi border the underlying Kerdylia unit, which at its top and all along its western margin is characterized by marbles, covering a thick core of gneisses of pre-Alpine continental crust (Himmercus et al., 2006). Eclogites found within the meta-ophiolite suite at Nea Roda indicate a high pressure metamorphic event of about 12 kbars and 530 °C, overprinted by greenschist facies metamorphism (Dimitriadis and Godelitsas, 1991). An imbricated meta-sedimentary sequence is observed in the area of Nea Madytos, south of Lake Volvi, overthrusted by the Arnea granite (Sakellariou, 1989). Both Kerdylia and the overlying composite nappe of the Volvi ophiolites and Vertiskos basement rocks are found on the hanging wall of the Miocene Strymon detachment (Dinter and Royden, 1993), whose footwall is made of the Pangeon marbles — the relatively autochthonous unit of Rhodope (Papanikolaou and Panagopoulos, 1981). The existence of Late Jurassic granitoids, like the Arnea granite (de Wet et al., 1989) intruding the Vertiskos nappe, indicates that ophiolite obduction of H8 occurred before the establishment of the back-arc volcanism/ magmatism of the paleo-Alpine orogeny related to the closure of H4 (Aubouin, 1977; Papanikolaou, 1989a). In the eastern parts of Rhodope, the ophiolites are observed on top of the gneisses and related rocks, which can be correlated with the higher parts of the gneissic core of the Pangeon unit, from the base of the upper tectonic unit of Sidironero, both known from the central part of Rhodope (Papanikolaou and Panagopoulos, 1981). In the area of Soufli near Evros River there are several outcrops of ophiolites comprising serpentinised peridotites, gabbros, gabbro-pegmatites, amphibolite schists and amphibolite gneisses (Maratos, 1960). Similar meta-ophiolitic rocks are known from the northern parts of the Rhodope metamorphic complex in Bulgaria, where serpentinites, meta-gabbros, meta-diabases and amphibolitised eclogites have been

The paleogeography of the Hellenides within the Tethys ocean is controlled by the dimensions of the continental terranes bearing the shallow-water carbonate platforms (Fig. 11a). The extension of the oceanic basins developed in between the continental blocks can be estimated only on the basis of the dimensions of the tectonic units corresponding to pelagic basins such as the Pindos and Maliac units, which is only a minimum estimate for each basin. It is remarkable that the number of the oceanic basins may be different across adjacent parallel segments of Tethys, where a continental terrane wedges out. This, for example, might be the case of the oceanic basins H2 and H4, which merge into one basin north of the Pelagonian terrane (H3), which terminates a few tens of kms north of Skopje (Fig. 11b). Within the terrane structure of the Dinarides (distinguished in seven terranes D1–D7) there is a 60 km long zone where the Vardar composite terrane (D3) comes in contact with the terrane of the Dinaric ophiolite belt (D6) (Karamata et al., 1997, Figs. 1 and 2). The Pelagonian continental terrane (H3 in the Hellenides and part of D5 in the Dinarides) terminates south of this zone, whereas the Drina–Ivanjica terrane (part of D5 in the Dinarides) starts north of this zone (Karamata et al., 1997, Figs. 1 and 2). Thus, a link between the Vardar/Axios oceanic basin (H4 and D3) and the oceanic basin of the Pindos–Cyclades/ Dinaric ophiolites belt (H2 and D6) exists in between the Pelagonian and Drina–Ivanjica terranes (Fig. 11b). The latter may explain the confusion caused on the timing of tectonic emplacement of the ophiolites of the Vardar/Axios oceanic basin (H4), which is reported as late Jurassic–early Cretaceous in the south, but Eocene in the north (D3), where however, it is linked also to H2–D6, which as described earlier has closed in the middle-late Eocene. The same type of paleogeographic connection can be observed between the oceanic basins of H4 and H6, north of the Paikon terrane H5 and also between the oceanic basins of H6 and H8 on both sides of the Rhodope autochthon terrane H7 (Fig. 11b). Thus, the paleogeography of the Hellenides can be regarded as a segment of the Tethys ocean with a number of sub-basins developed in between the continental terranes drifting northwards since late Paleozoic– Triassic. The dynamic nature of this paleogeography implies that

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mid-ocean ridges may develop usually during the first stages of opening of oceanic sub-basins following initial rifting. At a later stage supra-subduction zone ophiolites may be developed during the beginning of northward subduction of each basin under the accreted continental terrane. Thus, the mixture of different type of ophiolite environments reported for several ophiolites in the Hellenides (Beccaluva et al., 1984; Magganas et al., 1991; Smith, 1993, 2006; Migiros et al., 2000; Saccani and Photiades, 2004; Rassios and Moores, 2006a,b; Dilek et al., 2008) is expected for almost every oceanic sub-basin.

The timing of tectonic emplacement of the ophiolites as described earlier becomes younger from north to south, following the general model of northward subduction of the Hellenides beneath the European margin. Thus, we can distinguish: (i) present day subduction of the Eastern Mediterranean basin beneath the Hellenic arc and trench, which was initiated sometime during late Miocene. (ii) Late Cretaceous to early Tertiary subduction of H2 with tectonic emplacement on H1 during late Eocene–Oligocene. (iii) Late Jurassic subduction with tectonic emplacement of H4 on H3 during late Jurassic–early Cretaceous. (iv Pre-late Jurassic subduction and subsequent tectonic

Fig. 12. Schematic tectonic profiles of the four ophiolite belts of the Hellenides within the terrane structure. The age of tectonic emplacement of the ophiolites is shown in each case by the indication of the ages of the involved pre-, syn-and post-emplacement geological formations. The model of ophiolite obduction is demonstrated in all four cases with the four oceanic terranes (H2, H4, H6 and H8) emplaced southwards from Jurassic to Oligocene over the four continental terranes (H1, H3, H5 and H7). The difference between the deep crustal structure of terranes H7–H8–H9 in Rhodope and the rest more shallow structures of terranes H6–H5, H4–H3 and H2–H1 should be noticed.

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Fig. 12 (continued).

emplacement of H6 on H5 and of H8 on H7 with questionable correlation to the Cimmeride orogeny and to the closure of Paleotethys. The general tectonic setting of the Hellenic ophiolites is an obduction model of the oceanic areas over the more external continental terranes covered with shallow-water carbonate platforms

(Fig. 12). As described earlier in all cases of ophiolitic rocks belonging to the oceanic terranes H2, H4, H6 and H8 the obducted ophiolites together with the related pelagic sequences are observed above shallow-water carbonate platforms of the continental terranes H1, H3, H5 and H7 respectively. The pre-obduction period of each oceanic

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basin can be determined from the available chronological data of the ophiolitic rocks as well as from the stratigraphic data of the sedimentary sequences belonging to the same paleogeographic unit. Thus, the Pindos–Cyclades ophiolitic rocks are associated with the Upper Triassic–Eocene Pindos unit as well as the Cretaceous–Eocene Arvi unit (Fig. 12, 1). The Vardar/Axios ophiolites are associated with the Upper Permian–lower Cretaceous Maliac unit (Fig. 12, 2). The Lesvos–Circum Rhodope ophiolites are associated with the Triassic volcano-sedimentary pelagic sequence of the Lesvos allochthon and to the Triassic–Liassic pelagic sequences described in the Circum Rhodope belt (Fig. 12, 3). The Volvi–eastern Rhodope ophiolites are probably associated to the Nea Madytos meta-sedimentary sequence of undetermined age (Fig. 12, 4). The obducted ophiolites are only small pieces of the overall oceanic crust of each basin in the Tethys Ocean and the age determination of an ophiolite nappe does not cover all the duration of the ophiolite formation in the oceanic basin of its provenance. Thus, the external ophiolite belt of the Hellenides comprises mainly Jurassic rocks in Northern Pindos, both Jurassic and Cretaceous rocks in Crete and mainly Cretaceous rocks in the Dodekanese islands (Koepke et al., 2002; Liati et al., 2004). Nevertheless, the tectonic setting all along the external ophiolite belt is the same and the ophiolites are always found on top of the Pindos sedimentary pelagic sequence emplaced in the Eocene–Oligocene. Cretaceous ophiolites with early Tertiary tectonic emplacement dominate the NeoTethys realm eastwards along the Anatolides–Taurides-northern Arabian margin (Moix et al., 2008). The geodynamic conditions of ophiolite formation in each basin may be different and the ocean ridge spreading model may be accepted in some cases whereas the supra-subduction model may be accepted in other cases. However, the geochemical and related petrologic criteria of the ophiolitic rocks in the case of subduction models have to be supported also by the geology and tectonics of the adjacent sedimentary units. Thus, the absence of any tectonic activity throughout the Cretaceous period in the external Hellenides (from the Parnassos unit of the internal carbonate platform to the Paxos unit of the External carbonate platform) does not support westward subduction models but ocean ridge spreading in the area of the Pindos– Cyclades oceanic basin and/or subduction to the east under the more internal parts of the arc. Stratigraphic unconformities sealing the tectonic contacts between the allochthon ophiolites and the autochthon platforms show the progressive younger tectonism from the internal to the external parts of the Hellenic arc system. Thus, Oligocene–Miocene molassic sequences postdate the tectonic emplacement in the H2 external ophiolites belt (Fig. 11,1), Upper Cretaceous–Eocene meso-autochthonous sediments postdate the tectonic emplacement in the H4 internal ophiolite belt (Fig. 11,2) and Upper Liassic–Malm sediments postdate the tectonic emplacement of the ophiolites in the H6 Lesvos–Circum Rhodope belt structure (Fig. 11,3). However, in the Rhodope structure there are no younger sediments to date the end of the H8 ophiolite obduction, but only the late Jurassic granitic intrusions, which postdate the tectonic contacts among H7, H8 and H9. This might be related to the fact that the observed structures within the internal tectono-metamorphic belt in Rhodope belong to lower crustal levels and shallow sedimentary structures have been eroded. 5. Conclusions The tectonostratigraphic analysis of the formations related to the tectonic emplacement of the Hellenic ophiolites, which was based on geological maps and sections from typical outcrops for each oceanic terrane, showed a successively older age of ophiolite obduction in the Hellenic arc from its external to its internal part. The timing was determined as late Eocene–early Oligocene for the emplacement of the Pindos–Cyclades terrane H2, late Jurassic–early Cretaceous of the

Vardar/Axios terrane H4, post-Liassic–pre-late Jurassic of the Lesvos– Circum Rhodope terrane H6 and pre-late Jurassic of the Volvi–eastern Rhodope terrane H8. The terrane paleogeography of the Hellenides is characterized by a segment of the Tethys ocean which comprises a number of oceanic sub-basins separated by continental terranes of Gondwana origin drifting to the north. This general geodynamic process implies a relatively short duration of opening of each oceanic basin, followed by northward subduction and final ophiolite obduction during its closure. This intra-Tethyan plate tectonics process may result in the co-existence of the MOR and SSZ ophiolites in the oceanic terranes of the Hellenides. Thus, drifting of continental terranes and opening of oceanic basins is followed by continental terrane accretion and subduction of oceanic basins and end up with the closure of the oceanic basins and the subsequent ophiolite obduction. The ophiolite obduction model has been documented in all four oceanic terranes of the Hellenides with ophiolites and related pelagic sediments found on top of the shallow-water carbonate platforms. Deep geodynamic phenomena with tectonometamorphic and tectonomagmatic events are observed in all cases of ophiolite emplacement, whereas shallow structures of the upper crust are missing in several cases due to erosion and/or tectonic denudation. Differences in the timing of closure and suturing of oceanic basins along the Tethyan belt are due to the continental terrane geometry. Thus, where a continental terrane dies out, the two adjacent oceanic basins merge into one larger basin as this may explain the merging of the Vardar/Axios basin with the Pindos basin north of Skopje, where the Pelagonian continental terrane is terminated. Acknowledgements I thank my collaborators Emm. Vassilakis, I. Papanikolaou and L. Gouliotis for stimulating discussions and technical assistance with the drawings of this paper. I also thank A. Robertson, G. Migiros, Y. Dilek and an anonymous reviewer for their constructive comments during reviewing. References Al-Riyami, K., Robertson, A.H.F., Dixon, J., Xenophontos, C., 2002. Origin and emplacement of the late Cretaceous Baer–Bassit ophiolites and its metamorphic sole in NW Syria. Lithos 65, 225–260. Anastopoulos, I., Koukouzas, K., Faugeres, L., 1980. Geological Map of Greece 1:50.000, Sheet Kozani. I.G.M.E., Athens. Aubouin, J., 1965. Geosynclines. Developments in Geotectonics, vol. 1. Elsevier. 335 pp. Aubouin, J., 1977. Alpine tectonics and plate tectonics: thoughts about the Eastern Mediterranean. Europe from Crust to Core. J. Wiley, pp. 143–158. Barazangi, M., Sandwol, E., Seber, D., 2006. Structure and tectonic evolution of the Anatolian plateau in eastern Turkey. Geol. Soc. America, Sp. Paper 409, 463–473. Beccaluva, L., Ohnenstetter, D., Ohnenstetter, M., Paupy, A., 1984. Two magmatic series with island arc affinities within the Vourinos ophiolite. Contrib. Mineral. Petrol. 85, 253–271. Besenecker, H., Durr, S., Herget, G., Jacobshagen, V., Kauffmann, G., Ludke, G., Roth, W., Tietze, K.W., 1968. Geologie von Chios (Agais). Geol. Paleontol. 2, 121–150. Bonneau, M., 1984. Correlation of the Hellenide nappes in the south-east Aegean and their tectonic reconstruction. Geol. Soc. London, Sp. Publ. 17, 517–527. Bonneau, M., Jonkers, A., Meulenkamp, J.E., 1984. Geological Map of Greece 1:50.000, Sheet Timpaki. I.G.M.E., Athens. Brunn, J., 1956. Contribution a l'etude geologique du Pinde septentrionale et d'une partie de la Macedoine occidentale. Ann. Geol. Pays Hellen. 7, 1–358. Celet, P., Clement, B., Ferriere, J., 1976. La zone beotienne en Grece: implications paleogeographiques et structurales. Eclogae Geol. Helv. 69, 577–599. de Wet, A.P., Miller, J.A., Bickle, M.J., Chapman, H.J., 1989. Geology and geochronology of the Arnea, Sithonia and Ouranoupolis intrusions, Chalkidiki peninsula, northern Greece. Tectonophysics 161, 65–79. Dewey, J., Pitmann, W.C., Ryan, W.B.F., Bonnin, J., 1973. Plate tectonics and the evolution of the Alpine system. Bull. Geol. Soc. America 84, 3137–3180. Dilek, Y., Shallo, M., Furnes, H., 2005. Rift–drift, seafloor spreading, and subduction tectonics of Albanian ophiolites. Int. Geol. Rev. 47 (2), 147–176. Dilek, Y., Furnes, H., Shallo, M., 2008. Geochemistry of the Jurassic Mirdita ophiolite (Albania) and the MORB to SSZ evolution of a marginal basin oceanic crust. Lithos 100 (1–4), 174–209. Dimitriadis, S., Godelitsas, A., 1991. Evidence of high pressure metamorphicm in the Vertiskos Group of the Serbomacedonian massif: the eclogite of Nea Roda, Chalkidiki. Bull. Geol. Soc. Greece 25/2, 67–80.

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