An overview of the evolution of the internal hellenides (Albania, Republic of North Macedonia, Greek Central Macedonia): obductions, collisions and North Aegean extension

The Dinarides, Albanides and Hellenides form a 2500 km long NW-SE orogen whose Internal (Eastern) Hellenic Zones (Brunn, 1960; Kossmat, 1924) contain two large ophiolitic belts. Along the Albanian-Macedonian W-E transect, the Pelagonian micro-continent is bordered to the west by the Radiolaritic External Ophiolitic belt (ERO, Fig. 1) (Mirdita, Subpelagonian, Maliac ophiolitic massifs) and to the east, by the Internal Radiolaritic Ophiolitic Vardar belt (IRO) (Kober, 1952), divided into the Almopias, W. Peonias (Guevgueli) and E. Peonias (Thessaloniki) ophiolitic massifs (Mercier, 1968a) bordered to the east by the Serbo-Macedonian Massif (SMM) (Fig. 1). The ophiolitic basins closed during the Mid.-Late Jurassic forming ophiolitic nappes obducted top-to-the-W; the Internal Hellenides uplifted, were eroded and covered by Mid.-Late Cretaceous transgressive marine series and, a new marine Late Cretaceous Almopias basin formed. During the Palaeocene-Early Eocene, the Late Cretaceous Almopias continental basin subducted eastward; a Paikon dacitic volcanic arc formed above the continental subduction. The Internal Hellenides submitted to collision uplifted again, were eroded and unconformably covered by the Mesohellenic, Krania and Vardar molassic basins (Fig. 1). In what follows, we analyse the Mirdita and Vardar-Axios zones, highlighting the field data that constrain the architecture and evolution of the ophiolitic basins (Sect. 2) and, the tectonics of the Late Cretaceous-Palaeocene Paikon volcanic Arc, of the Vardar Paleogene molassic basins and of the E. Peonias units (Sect. 3); we summarize the Neogene-Recent extension of the North Aegean basin (Sect. 4). The detailed magmatic evolution of the ophiolites is not analysed in the paper; Gradstein et al. (2013)’s geological time scale is used in the following.

During the Jurassic times, the oceanic basins were opened and subsequently closed by subduction and obduction.

The Albanian Mirdita ophiolitic massif (Fig. 2I and II) continues in Greece into the Kastoria-Subpelagonian ophiolitic massifs (Mountrakis, 1983, Papanikolaou, 2009) (Fig. 1). The Mirdita Massif is chosen here because it was analysed by many Albanian workers (Aubouin and Ndojaj,1964; Shallo et al., 1987; Gjata et al., 1987; Kodra and Gjata, 1989; Frasheri et al., 1996; Kodra, 2016; Aliaj and Bushati, 2019, among others), because we conducted there a French-Albanian 5-year field work program with PhD theses (Godroli, 1992; Hoxha, 1996; Dimo, 1997; Insergueix-Filippi et al., 2000) and, because the Mirdita massif better exposes the Jurassic deformations than the Subpelagonian Massif obscured by Tertiary thrusts (Mountrakis, 1983).

In the Early Triassic, the Mirdita basin opened forming a graben in the Korab Palaeozoic substratum (P_z on Fig. 2III) overlain by Werfenian-Anisian marbles (T₁–T_2a) with transitional alkaline volcanism, by a Ladinian-Mid. Liassic carbonate platform (T₂l-J₂-J₃¹) and by a Dogger-Early Malm tholeiitic Diabase-Radiolarite formation (J₃). The continental crust broke and the Mirdita oceanic crust and mantle were denudated (Nicolas et al., 1999), covered by lavas and radiolarites of Mid. Bathonian to Callovian-Early Oxfordian age (Marcucci and Prela, 1996).

During the Late Jurassic, the oceanic crust of the Mirdita basin thrust first eastward upon the Korab Massif, forming a Mirdita ophiolitic nappe; the normal faults that separated the Korab continent and the Mirdita Jurassic oceanic basin were reworked as eastward reverse faults and the Korab marbles were folded with an eastern vergence. At the base of the nappe, a micaschist sole in HP-HT metamorphic facies with T° up to 700–900°C (Dimo, 1997) suggests the origin of the ophiolites at an oceanic ridge (Nicolas et al., 1999; Maffione et al., 2015). The HP-HT metamorphism decreases downward and eastward from Granulite to Greenschist facies (Turku, 1986; Dimo, 1997) and vanishes eastward supporting a top-to-the-E thrust of the Mirdita ophiolitic nappe. The micaschist sole of the nappe is folded with an eastern vergence (Fig 3.1a), it delivered helicitic garnets (85% of 130 data) and shear criteria (75% of ∼200 δ, 6, S/C data) indicating thrusting with an eastern vergence (Godroli, 1992; Vergély and Kodra, 1995).

Subsequently, the Korab Massif covered by the supra-Korab ophiolitic nappe, thrust westward (Vergély et al., 1991; Kodra et al., 1993) (Fig. 2I) and overturned top-to-the-W the folded structures (Fig. 3.1b) with a previous eastern vergence (Fig. 3.2); the westward thrust of the Korab Massif upon the Mirdita ophiolites is coeval with the eastward underthrust (subduction) of the oceanic crust below the Massif (D₁ event); subduction is attested by the presence of ophiolitic bodies below the Korab basement within the Mali i Bardhë window (Aliaj and Bushati, 2019) (Fig. 2I). Subsequently, the basement, overthrust by the ophiolites, obducted westward upon the Mirdita ridge whose the rift basin closed progressively from south to north, by diverging thrusts with HP-HT metamorphic soles dated between 174 and 160 Ma (see Dimo et al., 1996; Vergély et al., 1998). Obduction continued westward upon the Mid. Callovian-Early Oxfordian Diabase-Radiolarite basin (J₃) (Fig. 2I) whose low-angle syn-sedimentary normal faults were reworked as reverse faults (Vergély and Kodra, 1995; Hoxaha, 1996).

Briefly, the oceanic basins opened west of the Pelagonian micro-continent during the Jurassic (Hynes et al., 1972; Mercier et al., 1975; Mountrakis, 1984; Kodra and Gyata, 1989; Vergély et al., 1991; Robertson, and Shallo, 2000). The oceanic crust thrust first eastward upon the Korab margin in HP/MP-HT metamorphic facies, not extending as far as the Pelagonian eastern margin and the Almopias ophiolites in HP-LT metamorphic facies (see Sect. 2.2). Subsequently, the ophiolites over thrusting the Korab basement, obducted westward on the Mirdita ophiolitic basin, coeval with their underthrusting below the Korab Massif, subduction is attested by the presence of ophiolites below the basement in the Korab tectonic windows. The Mirdita rift basin opened and, closed by diverging thrusts at 174–160 Ma (Vergély et al., 1998), followed by transgressive Jurassic-Early Cretaceous detrital deposits and Mid.-Late Cretaceous neritic limestones. Thus, the Mirdita Ocean (=H₄ ophiolitic belt of Papanikolaou, 2009) was situated west of the Pelagonian massif in contrast with a proposed eastern Vardar location (see Bortolotti et al., 2013; Ferrière et al., 2012 and ref herein); the “Pindos Ocean” name, currently used, is not retained because it may confuse the Mirdita-Subpelagonian Jurassic oceanic basins with the Triassic-Jurassic Pindos sedimentary basin named a long time ago (Renz, 1952; Brunn, 1956; Aubouin, 1959). The Mirdita nappe is not a far-travelled nappe rooting in the Vardar zone (Vergély, 1975; Mercier et al., 1975; Aliaj and Bushati, 2019; among others) as previously proposed (Ricou and Godfriaux, 1991; Tremblay et al., 2015, among others); conversely, it did not thrust as far as the Vardar zone as suggested by Jones and Robertson (1991).

In the Voras Mts (Fig. 4A), the Almopias basin opened in a Pelagonian Paleozoic basement (Mposkos and Krohe, 2004), overlain by a volcaniclastic-conglomeratic series, by Werfenian-Lower Anisian marbles with a Porphyrite-Radiolarite formation (Dumurdžanov et al., 1976) and by Mid.-Late Triassic to Jurassic carbonates (Mercier, 1968a; Bulle, 1973; Mavridis et al., 1977a). Below the Almopias Mavrolakkos unit (Fig. 4B), a thin slice of red radiolarites with MORB volcanics indicates the existence of an Almopias oceanic basin in the Early Carnian-Mid. Norian (=Vrissi unit of Stais et al., 1990; Saccani et al., 2015). Near Veles (N. Macedonia, Fig. 1), the Almopias Triassic-Jurassic carbonates are overlain by Late Jurassic radiolarites (Robertson et al., 2013) and, near Arnissa (pt. 4 on Figs. 4B and 5A), by schistose radiolarites and volcaniclastic deposits reworking clasts of radiolarites, pillow lavas and metamorphic rocks (Bijon, 1982; Mercier, 1968a). In the Askion Mts (Fig. 5B), ophiolitic blocks are reworked in cherts and limestones of Pliensbachian age (Mavridis et al., 1977a); they indicate an early tectonic activity (beginning of subduction?) in the Almopias basin. In the High Vermion Mt. (Braud, 1967), the ophiolites comprise harzburgitic mantle tectonics, oceanic crust and lavas of island-arc tholeiitic affinities; ophiolites are dated at 169 ± 2.4 and 173 ± 3 Ma (U-Pb zircon date, Liati et al., 2004), they are overlain by pillow lavas and radiolarites of Latest Bajocian-Early Bathonian age (168 ± 1 Ma) (Chiari et al., 2003). In the Almopias basin, they have island-arc tholeiitic affinities (pt. A6 on Fig. 4B; Georgiadis et al., 2016; Saccani et al., 2015) and, in the Vourinos massif (Brunn, 1956; Moore, 1969; Zimmerman, 1972; Beccaluva et al., 1984 among many others), they formed in a SSZ setting (Noiret et al., 1981).

The Almopias ophiolites are thrust, with at the base, slices of prasinitic mélanges (Klissochori unit, Fig. 4B) comprising exotic blocks (Fig. 6.2), metamorphosed in Amphibolite facies, retrograded in Greenschist facies. Below the Vourinos nappe, metamorphic slices are affected by metamorphism in Amphibolite/HP-LT facies with lawsonite in garnet amphibolites (Vergély, 1984) (Fig. 5B), dated at 171–169 Ma (Spray et al., 1984); the pre-alpine Pelagonian basement is also metamorphosed in HP-LT facies (Lips et al., 1998; Mposkos and Krohe, 2004; Kilias et al., 2010); the HP-LT metamorphism requires subduction of the ophiolitic units and of the eastern margin of the Pelagonian basement. The Vourinos ophiolites were subsequently overlain by non-metamorphic Callovian-Oxfordian (∼166–156 Ma) deep-water siliceous limestones, covered by Oxfordian neritic limestones (Pichon and Lys, 1976; Carras et al., 2004), the neritic limestones indicate that the oceanic crust and the siliceous limestones were emplaced by obduction near sea-level.

Below the Almopias ophiolitic nappe, the Eastern Pelagonian and Central Almopias marbles are affected by folds (B₁) (Fig. 6.1) trending N140 ± 15°E, with a western vergence (Fig. 5A) (D₁ event); the Pelagonian carbonates thrust by the Vourinos ophiolites are also affected by isoclinal recumbent folds (B₁) (Figs. 6.3, 6.4, 6.5), with a westward vergence (red arrows on Fig. 5B). The westward recumbent folds and Lizardite-Antigorite shear surfaces in the serpentinite sole below the ophiolites, support the westward obduction of the ophiolites (D₁ event) (Vergély, 1977); it occurred between 168 and 166–156 Ma, coeval with the metamorphic M₁ event of Late Jurassic-(?) Lower Cretaceous age (Mercier, 1966a).

The carbonate platform was again affected by folds, (B₂) folds trending N10°E± 20° (Fig. 7.1; 7.2) with an eastward down-slip (normal) vergence (Vergély, 1977) that results from a westward gravitational exhumation (see Chemenda et al., 1996; Boutelier et al., 2002) of the subducted platform (see cartoons on Fig. 5A) (D₂ event). In the Central Almopias units, the exhumation-extension is coeval with a volcanic activity of island-arc tholeiitic and boninitic characters (Saccani et al., 2015) formed in a horst and graben structure of Kimmeridgian-Tithonian age (Bijon, 1982, Mercier and Vergély, 1984); in the Eastern Almopias units (Fig. 4B), the volcanic activity comprises meta-dolerites and lavas of MORB tholeiitic affinity, interbedded with radiolarites (Mercier, 1968a; Bébien et al., 1980; Béchon, 1981; Sharp and Robertson, 2006; Saccani et al., 2015); it is of Callovian-Kimmeridgian to Early Tithonian age (Stais et al., 1990) and formed in a deep basin at the margin (fore-arc?) of the Paikon Volcanic Massif. Calc-alkaline N20–40°E trending dykes intruding the Eastern Almopias units indicate a strike slip tectonic regime with a WNW extension-NNE shortening (ster. A4 and A8 on Fig. 4B)

The Mirdita ophiolites did not thrust as far as the western margin of the Almopias ophiolitic basin and the ophiolites of the eastern Almopias basin subducted below the Paikon Massif; thus, the Almopias ophiolites have their roots neither in the Mirdita ophiolites (see Sect. 2.1) nor in the Guevgueli basin (see Sect. 2.4). The Almopias ophiolitic basin opened in the Mid.-Late Triassic between the Pelagonian and the Paikon-Malarupa continents. The oceanic crust was denudated around 173–170Ma and overlain by radiolarites (∼168Ma) and by non-metamorphic Callovian-Oxfordian (∼166–156 Ma) deep water limestones. The Almopias oceanic crust and the Pelagonian eastern margin metamorphosed in HP-LT facies subducted subsequently eastward below the Malarupa continent and, a Paikon Volcanic Arc formed above the subduction (see Sect. 2.3). Subsequently, the ophiolites obducted westward upon the Pelagonian margin between 168 and ∼166–156 Ma (D₁ event) and exhumed gravitationally westward coeval with a volcanic activity during the Callovian-Early Tithonian. The original location of the Vourinos ophiolites is debated between an eastern Almopias Vardar (Bernouilli and Laubscher, 1972; Vergély, 1977, among others) and a western Mirdita-“Pindos” origin (Jones and Robertson, 1991; Rassios and Dilek, 2009, among others). Because the Mirdita ophiolitic nappe in HP-HT setting (up to 700 to 900°C) did not thrust as far as the Almopias ophiolites that are in HP-LT setting (see Sect. 2.1) and, because the Vourinos ophiolitic nappe is thrust top-to-the-W upon the Eastern Pelagonian margin (Vergély, 1976), the Vourinos ophiolites originated in the Almopias basin.

The Paikon Volcanic Arc formed upon the Malarupa metamorphic massif (Voras Mts.) overlain by the Tzena marble formation (Fig. 8A) that is the eastern continuation of the Mid.-Late Triassic-Jurassic Pelagonian-Almopias carbonate platform (Fig. 4A; see Sect. 2.2). The Paikon and the Malarupa separated by the dextral strike-slip Voras fault zone, striking WSW-ENE, re-activated as a normal fault during the Neogene-Quaternary (Katrivanos et al., 2015). In the Paikon Massif, the carbonates are overlain by a thick Central Sequence comprising marbles, calcschists, arenaceous schists, basalts, rhyolites and andesites (Paleochori, Pirghos and Livadia formations; Mercier et al., 2005) (Fig. 8B) that have low-K island-arc affinities, some with boninitic or calc-alkaline characters (Bébien et al., 1994; Brown and Robertson, 2003; Saccani et al., 2015), they are overlain by Upper Jurassic-Lower Cretaceous calc-schists and marbles with Cuneolina sp..

The Western units of the Paikon Massif (1, 2, 3 on Fig. 8B) thrust top-to-the-NE upon the Central Sequence re-activating probably the W-dipping normal faults that separated the Paikon Volcanic Arc and the Almopias basin in the Jurassic. The units comprise volcanics of low-K island-arc tholeiitic-IAT, deeply weathered and transgressively overlain by Late Tithonian-Berriasian limestones with Anchispirocyclina (Iberina) lusitanica (Fig. 10.1) and/or by Lower Cretaceous detrital deposits. The Eastern Paikon units (4, 5, 6 on Fig. 8B) thrust top-to-the-W upon the Central Sequence; the Kastaneri unit (5) comprises rare prasinites and a thick (>1000 m) series of white tuffs, rhyolites and hypo-volcanic granites (Mercier, 1968a; Davis et al., 1988) of calc-alkaline characters (Bébien et al., 1994). At the top, they are interbedded with limestones passing upward to pyroclastic sandstones, tuffs and limestones with an abundant microfauna-flora of Latest Oxfordian-Early Kimmeridgian age (Mercier, 1968a), an age in agreement with the 155 ± 2 Ma crystallization age of the mylonitic rhyolites (U-Pb zircon date, Anders et al., 2005). The Kastaneri series was a rhyolitic volcanic centre located east of the low K-tholeiitic volcanic centre of the Paikon Volcanic Arc. More to the east, the Artzan-Vafiochori series (Fig. 11) is a lateral continuation of the Kastaneri series, however, metamorphosed in a LP/MP-HT facies. Via marble slices, it thrusts upon the Paleozoic Karathodhoro basement (dated at 319 ± 4 Ma; Anders et al., 2005), affected by LP/MP-HT metamorphism and anatexis and, intruded by Crd-Sil and Crd-And anatectic granitic dykes dated at 161 ± 5 Ma (Rb-Sr Bt data, Borsi et al., 1966b; Mercier, 1968b).

The Western Paikon units are metamorphosed in lawsonite-albite facies (Mercier, 1968b; M₁ metamorphic event) and the Paikon Central Sequence in the high-pressure chlorite subzone of greenschist transitional blueschist facies (Barroz et al., 1987). The pressure estimated at 6–7 kbars requires sinking of the Western units and of the Central Sequence into a subduction zone that dip eastward below the Kastaneri rhyolitic volcanic centre metamorphosed in greenschist facies. The Central Sequence is affected by NNW-SSE trending isoclinal (B₁) folds with a western vergence (ster. A1a on Figs. 8B and 10.2) refolded by (B₂) folds trending N30°E ± 15°with a top-to-the-SE vergence, due to a subsequent exhumation (Fig. 10.3). In the north facing cliff of the Tzena (Evgeni) summit (Fig. 8A), the Tzena marbles, with (S₁) schistosity, are re-folded by spectacular km-scale recumbent (B₂) folds with an eastward vergence (Mercier, 1973) (see Fig. in Annex I). The (B₂) folds formed during the westward gravitational exhumation of the subducted Paikon Arc that pre-dates the Late Tithonian-Berriasian (Vergély, 1984).

In contrast with models that rooting the Almopias ophiolites into the Guevgueli basin, subsequently thrust above the Paikon Massif (Ricou and Godfriaux, 1991; Katrivanos et al., 2013); the Guevgueli ophiolites did not thrust above the Paikon Massif during the Late Jurassic because no remnants of a MORB ophiolitic nappe are present within the Paikon Central Sequence, because the Massif was an active low-K tholeiitic island-arc at that time (Bébien, 1983, among others) and, because the Almopias and Guevgueli ophiolites are in drastically different metamorphic settings (HP-LT and low greenschist facies, respectively). During the Tertiary, it is no more possible because the Almopias ophiolites, sealed upon the Pelagonian eastern margin since Late Albian (∼100 Ma) (at pt. 4 on Fig. 4B), could not thrust over the Paikon Late Maastrichtian (75–65 Ma) limestones (Mercier and Vergély, 1995).

The submarine Paikon Volcanic Arc (W and Central island-arc tholeiitic volcanic centres plus the Kastaneri rhyolitic volcanic centre) formed upon the Malarupa massif; the W and Central volcanic products are not precisely dated; in the Voras Mts., they overly the Tzena (Evgueni) marble platform that is the eastern continuation of the Mid. Late Triassic-Jurassic Pelagonian-Almopias carbonate platform and, in the Paikon Mts. they are overlain by Late Jurassic-Early Cretaceous transgressive deposits. The Kastaneri rhyolitic centre is dated of Latest Oxfordian-Early Kimmeridgian age. The volcanic Arc formed above the top-to-the-E Almopias driven; subduction progressing, the Paikon tholeiitic island-arc centre, metamorphosed in HP-LT facies, was drawn by the load of the Almopias slab into the Paikon subduction zone, below the rhyolitic Kastaneri volcanic centre metamorphosed in greenschist facies (see Sect. 2.4); subsequently, they gravitationally exhumed westward.

The Guevgueli basin opened above the subduction of the Paikon Arc; ophiolites are non-schistose gabbros, lavas, pillow lavas of MORB tholeiitic affinities and low TiO₂ calc-alkaline dykes (Pendjerkovski, 1963; Ivanovski, 1963; Bébien and Mercier, 1977; Vergély, 1984; Ivanov et al., 1987; Saccani et al., 2008, 2015; Kukoč et al., 2012; Bortolotti et al., 2013; Bonev et al., 2018); along the western margin of the Fanos granite, they are monzo-gabbros and monzo-diorites (Fig. 11) (Mercier et al., 2005); ultrabasites are scarce.

The Guevgueli ophiolites are of Mid.-Late Jurassic age, dated at 160 ± 1.2 Ma (U-Pb SHRIMP zircon date; Zachariadis et al., 2006), 164 ± 0.5 Ma (Ar-Ar date; Božović et al., 2013); the upper pillow lavas are overlain by Late Callovian-Oxfordian radiolarites (∼165–156 Ma) (Danelian et al., 1996). The ophiolitic basin opened along and back of the Kasteneri volcanic rhyolitic centre dated at ∼155 Ma; the ophiolites are intruded by the Platania anatectic granite dated at 158 ± 4 Ma (Pb zircon date; Anders et al., 2005), by the magmatic Fanos granite (Fig. 11) (Bébien and Mercier, 1977; Christofides et al., 1990; Šarić et al., 2009) dated at 156 ± 2 Ma (7 Rb-Sr Bt, corrected dates), at 152 ± 2 Ma (K-Ar Bt date, Borsi et al., 1966a; Mercier, 1968b) and, at 158 ± 1 Ma (U-Pb Zircon date, Anders et al., 2005). The ophiolites extend northward into the Kara Dagh massif, east of Skopje (Fig. 1), where they are intruded by the Jurassic Lojane granite (Rollet, 1969). Geophysical investigations (Kiriakidis and Brooks, 1989) indicate their southward continuity below the Axios basin with the non-schistose Efkarpia, Oreokastro (Haenel-Remy and Bébien, 1985) (Fig. 13) and Sithonia-Kassandra ophiolitic massifs (Gauthier, 1984; Jung and Mussalam, 1985; Bonev et al., 2015).

The sheeted dykes of the Guevgueli ophiolites are considered to be conjugated N60° and N170° striking shear planes (11.1 on Fig. 11) (Bébien and Gagny, 1978) opened subsequently by fluid pressure; they suggest that the basin opened in a strike-slip tectonic regime; however, followed by a WNW trending tension probably due to diapiric intrusions. At the southern border of the Fanos granite, the Piyi migmatites comprise basic blocks indicating that the acid and basic magmas were emplaced synchronously (Zachariadou and Dimitriadis, 1994; Mercier et al., 2005). NNE-SSW striking aplitic dykes cutting the Fanos granite also indicate a WNW-ESE trending extension.

The upper (eastern) mb. of the Guevgueli ophiolitic massif thrusts westward upon the lower mb. with at the base, marble slices and the Karpi slice (Figs. 8B and 11); the latter comprises schistose gabbros, amphibolites and serpentinites metamorphosed in low grade HP-LT metamorphic facies (Chl + Ab + Cal + Lws + Pmp) (Vergély, 1984). They indicate a limited intra-oceanic eastward subduction (?), followed by the westward obduction of the entire ophiolitic massif upon the Paikon Ghriva (Fig. 10.4) and Kastaneri units, the latter of Early Kimmeridgian age is folded with a western vergence in greenschist metamorphic facies (ster. 11.3; 11. 4 on Fig. 11). The observed westward obduction is contrast with the westward subduction of the Guevgueli ophiolites below the Paikon Arc proposed by Ferrière and Stais (1995), not supported by field data but inferred from a model of subduction of the ophiolitic basins below their continental margins. Here, the model is non adequate; the Guevgueli ophiolites are non-schistose ophiolites, metamorphosed in low grade greenschist facies while the Paikon Volcanic Arc is folded top-to-the-W., highly schistose and metamorphosed in HP-LT facies; the structure requires the eastward subduction of the Paikon Arc below the ophiolites and the thrust of the ophiolites top-to-the-W that is easily observed all along the Paikon/Guevgueli contact (∼30 km) (Figs. 8A, 8B and Fig. 10.4).

Subsequently, the Guevgueli ophiolites exhumed and were eroded; in the Vafiochori unit (Fig. 11), the pillow lavas capped by siliceous iron oxide material or laterites (Mercier, 1968a; Xydas and,Efstrafides, 1993), are overlain transgressively by a ≈1500 m-thick series of black pelites and greywackes (Erikinos formation). The Erikinos formation is overlain by limestone lenses with Anchispirocyclina (Iberina) lusitanica of Late Tithonian-Berriasian age (Chorighi formation of Mercier, 1968a; Carras and Georgala, 1998, in Greece) and by 300m-thick neritic limestones (Demir Kapija formation of Boev et al., 2006; Bozović et al., 2013, in N. Macedonia). In N. Macedonia, a thin Erikinos series (Ivanov et al., 1987) is interbedded with Foraminifera limestones of Oxfordian-Early Kimmeridgian age (Effendiantz, 1971; Kukoč et al., 2012) and, in the Karadagh massif (Fig. 1), with Dacinella irregularis and Globigerina (Haeusslerina) helveto-jurassica of Early Kimmeridgian age (=flysch formation of Rollet, 1969). In Greece, the Chorighi formation is covered by a thick Late Jurassic-Early Cretaceous conglomeratic formation (Fig. 11) (Dubkon formation of Leuch, 1916) that reworks pebbles of red sandstones, quartzites, meta-tuffites, −dolerites and −gabbros, white marbles and Bt-Ms micaschists.

Briefly, the Guevgueli ophiolitic en-sialic basin opened at ∼164–160 Ma, within the eastern margin of the Malarupa microcontinent (Karathodhoro-Bogdanci-Pyi massifs, Fig. 11), above the subduction of the Paikon Volcanic Arc, in the back of the Kastaneri rhyolitic centre, probably activated by roll-back of the Paikon slab. Then, a deep basic magma arose and intruded the continental margins of the basin that were submitted to anatexis; an acid anatectic magma formed that was emplaced synchronously with the basic magma forming the Pyi migmatites; anatectic granitic dykes (161 ± 5 Ma) intruded the Karathodhoro basement; the anatectic Platania (∼158 Ma) and the magmatic Fanos (∼156–158 Ma) granites were emplaced into the Guevgueli ophiolites. A weak acid magmatic activity continued during the Oxfordian-Kimmeridgian (Davis et al., 1980) and small rhyolitic lava flows covered the Tithonian-Berriasian Demir Kapija limestones (Bonev et al., 2018). Obduction of the Guevgueli ophiolites occurred in the Early Kimmeridgian (∼157 Ma); it is clearly westward, upon the Paikon Volcanic Arc, by contrast with the proposed eastward obduction upon the E. Peonias units, inferred from a model of subduction of the ophiolitic basins below their continental margin.

We separate the Eastern Peonias zone (= E. Vardar z.; Papanikolaou and Stojanov, 1983) into a Continental Margin (= Circum Rhodope Belt of Kaufmann et al., 1976) metamorphosed in greenschist facies and a Magmatic Suite, metamorphosed in HP-LT facies.

The Continental Margin is an overturned sedimentary sequence (Fig. 12A) with at the base, a rhyolitic volcaniclastic series (Dimitriadis and Asvesta, 1993) of Lower Scythian age (Pirghoto formation of Ferrière and Stais 1995), dated at ≈240 Ma (U-Pb zircon method; Frei cited in Kostopoulos et al., 2001). It is covered by Anisian limestones with Brachiopods (Cuneirhynchia trinodosi) and by Mid.-Late Triassic marbles (Deve Koran formation) (pt.13.2, Fig. 13A) that thrust upon the Late Triassic-Early Jurassic detrital carbonate Valti formation (Mercier, 1968a; Stais and Ferrière, 1995). The latter thrusts upon Ladinian-Late Norian limestones and calc-schists of the Vaptisti formation (= lower mb. of the Svoula gr.; Kaufmann et al., 1976) and the Early-Mid. Jurassic Melissochori flysch (Mercier, 1968a) (= upper mb. of the Svoula gr. of Kockel et al., 1977). A 10 m-thick lava flow of MORB tholeiitic affinity is interbedded within the Vaptisti calcschists (Fig. 12A); the Metallikon unit (Fig. 12B) also comprises ∼10 m thick-sills and lava flows of MORB tholeiitic affinities, interbedded with Foraminifera and Ostracoda (Carinobardia gr. triassica alpina) limestones of Rhaetian age. The sequence thrusts a slice of Paleozoic gneisses (Fig. 13A) intruded by the Balteika calc-alkaline granites and by Ms-pegmatitic dykes dated at ∼299 to 256 Ma (Ms Rb-Sr dates, Mercier, 1968b).

The Magmatic Aspri Vrissi unit (Fig. 13) comprises two sub-units, the upper one continues SE of Thessaloniki, into the previously named “upper mb. of the Chortiatis unit”(Monod, 1964). It is an overturned series of Early to Late Triassic limestones (Barroz et al., 1987; Ferrière and Stais, 1995) interbedded at the top with andesites, red rhyolites and greywackes, covered by Late Triassic-Liassic black limestones with Ostracoda and Lagenidae (Mercier, 1968a) and Glomospira, Frondicularia cf. woodward (Kockel and Ioannides, 1979). They are overlain by a sequence of chloritic, graphitic garnet phyllites and ferruginous sandstones, interbedded with black cherts (lydian stones), covered by MORB dolerites, gabbros sills and basalt flows (Bébien et al., 1986).

The Chortiatis unit (=”Chortiatis lower mb.” of Monod, 1964) is a thick (∼ 4 to 5 km) island-arc tholeiitic suite of schistose serpentinites, meta-gabbros, diorites and granophyres of calc-alkaline characters, dated of Late Jurassic age (∼159 Ma, see below). The Thessaloniki ophiolitic unit (Fig. 13) comprises a ≈3 km-thick series of highly schistose harzburgitic mantle tectonites, schistose and cataclastic gabbros, diorites, plagiogranites, lavas, and pillow lavas of island-arc tholeiitic affinities (Ricou, 1965; Antoniades and Ioannides, 1978; Sapoutsis, 1979; Kockel and Ioannides, 1979: Vergély, 1984; Mussalam and Jung, 1986: Zachariadis et al., 2007); they are intruded by dykes with calc-alkaline characters (Schünemann, 1985; Bonev et al., 2018).

The Late Triassic-Early Jurassic Melissochori and Metallikon units of the Continental Margin (Fig. 11), metamorphosed in greenschist facies, are affected by isoclinal folds (B₁), trending N325°E to N20°E with a top-to-the-W vergence (D₁ event), superimposed by (S₂) schistosity planes trending WSW-ENE (Fig. 14.1) (D₃ event); in the Melissochori unit (ster. A28 and site 9 on Fig. 13A; D₁ event), the Jurassic (B₁) folds trending NNW to N-S have high angle dip axes, being refolded by (B₂) folds of the Early Tertiary D₃ event (Fig. 14.2a, b) (see Sect. 3.5).

The Aspri Vrissi unit of the Magmatic Suite is affected by (B₁) folds trending also NNW to N-S (D₁ event) (ster. B5–6, B11–12 on Fig. 13A), re-folded by (B₂) folds; the Chortiatis unit, highly schistose (ster. A28 and site 9 on Fig. 13A), with a diorite of Jurassic age (159 ± 4 Ma; Zircon U-Pb date, Zachariadis et al., 2006, Zachariadis 2007), is also affected by (B₁) folds of the D₁ event, metamorphosed in HP-LT facies (P = 6–7, T ≥ 300°C, Michard et al., 1994). The Thessaloniki meta-gabbros and dolerites are metamorphosed in Epidote-Amphibolite facies (Hbl green+ Ep+ Cal); the ultrabasites are highly schistose. The Mylonitic Slices (Ricou, 1965), below the Thessaloniki nappe (Fig. 13), have also flow schistosity (S₁) in HP-LT metamorphic facies [Ep+ Qtz +sericite+ Am blue −Crossite, not analysed, and post-tectonic Lawsonite, Stilpnomelane and Quartz] (Vergély, 1984) (D₁ event). The Hbl minerals are stretched in a 125–135°E direction with an south-eastern vergence (Fig. 14.3) due a subsequent gravitational exhumation (D₂ event).

Thrusting top-to-the W. of the E. Peonias Magmatic Suite and HP-LT metamorphism require its subduction eastward below the Continental Margin of the Thessaloniki basin, metamorphosed in greenschist facies (D₁ event). The metamorphic rocks are dated at 110–120 Ma, considered to be the age of the metamorphic M₁ event (Kyndonakis et al., 2015); in fact, being overlain unconformably by Aptian-Albian deposits (see Sect. 3.4, Fig. 16), they pre-date the Late Cretaceous; metamorphism affects the Jurassic diorites (∼159 Ma, see above), it is of Jurassic age, rejuvenated during the Early Cretaceous exhumation.

The Paleozoic gneissic slices, at the base of the Continental Margin sequence, intruded by calc-alkaline granite dated at ∼299 to 256 Ma, have an age similar to the age of the Vertiskos Paleozoic Ms-pegmatitic dykes (∼ 327 ± 15 to 284 ± 30 Ma; 5 Ms Rb-Sr corrected dates) and granites (300 ± 10 Ma; K-Ar date) (Borsi et al., 1964; Mercier, 1968b; Zervas, 1979). The substratum of the Continental Margin is a western continuation of Vertiskos basement, from which it was separated in the Early Triassic by graben filled by rhyolitic volcaniclastic deposits. During the Late Triassic, the Continental Margin was bordered to the west by a Mid.-Late Triassic carbonate slope and rise sequence, subsequently overlain by the post-rift Early-Mid. Jurassic Melissochori flysch (Ferrière and Stais, 1995). Cherts and MORB lavas (Mercier, 1968a) mark a change to a deep oceanic environment in the Rhaetian times. The Chortiatis and Thessaloniki magmatic rocks are the Mid-Late Jurassic crust and mantle of the E. Peonias Thessaloniki ophiolitic basin (Vergély, 1984), dated at 172 ± 5 Ma (Zircon Concordia date, Bonev et al., 2015) and 169.2 ± 1.4 Ma, 160.6 ± 1.2 Ma (Zircon Pb-Pb date, Zachariadis et al., 2006). The Chortiatis-Thessaloniki oceanic crust metamorphosed in HP-LT metamorphic facies subducted eastward below the E. Peonias Continental Margin of the basin metamorphosed in greenschist facies and, subsequently, it obducted westward upon the W. Peonias Vassilika unit (Fig. 13).

We did not analyze the SMM, simply the boundary between the SMM Vertiskos and the E. Peonias (“Circum Rhodope Belt”) units.

The E. Peonias units border the SMM (Kockel et al., 1977; Burg et al., 1995) whose basement is intruded by Permo-Carboniferous Granitoides (Himmerkus et al., 2006; Meinhold et al., 2009; Bonev et al., 2013), by Permian-Late Carboniferous granites (300 ± 10 Ma; K-Ar date) and pegmatitic dykes (∼ 327 ± 15 to 284 ± 30 Ma; 5 Ms-Rb-Sr corrected dates) (Borsi et al., 1964; Mercier, 1968b; Zervas, 1979). The basement is overlain by the Late Paleozoic detrital Examili formation (ca 300 Ma; Meinhold et al., 2009) that fills a graben (Fig. 11); it is a pre-rift (eventually syn-) sequence. The Vertiskos Sequence of the SMM comprises a Lower Vertiskos Series of metasediments interbedded with greywackes and rare marble beds and an Upper Vertiskos Series of migmatitic gneisses and orthogneisses with, between them, a mylonitic zone of meta-ophiolites of abyssal tholeiitic affinities considered to be a Vardar oceanic basin (Burg et al., 1995).

In the Early Triassic, the Late Paleozoic basements of the E. Peonias Continental Margin and of the Vertiskos units were separated by the Langhada-Doirani fault zone bordered by graben filled by rhyolitic volcaniclastic deposits, to the west, the Lower Triassic Pirghoto formation (Ferrière and Stais, 1995) of the Continental Margin, metamorphosed in greenschist facies, dated at 240 Ma (Frei cited in Kostopoulos et al., 2001) and to the east, the rhyolitic volcaniclastic Akrita formation of the Vertiskos unit metamorphosed in HP-LT (Asvesta, 1992) (Figs. 11 and 13). The rifting stage is dated by the Early Triassic Pirghoto formation, by the Middle Triassic a-type Arnea granite (Himmerkus et al., 2009) and the Volvi gabbroic body (240 Ma, Bonev et al., 2019) of the Therma-Volvi-Gomati ultramafic-mafic complex. The Langhada-Dorani fault bordered by the Akrita rhyolitic volcaniclastics (Fig. 1 in Burg et al., 1995) was a major normal fault in the Early Triassic.

The Vertiskos Massif is a large antiform sheared southward (Burg et al., 1995), it is bordered to the west by the Langhada-Doirani fault that continues NNW of the Dorani Lake, into the Dorani-Strumića reverse fault with, at the base, the HP-LT Akrita and meta-gabbro slices (Erdmanndörfer, 1924 and F and ν formations on the North Macedonia Gevgelija and Strumića geological sheets) (Figs. 11 and 16).The Akrita formation (Fig. 11) being metamorphosed at P = 7–8 kbars (Michard et al., 1994) and the Lower Vertiskos unit at P =11–14 kbars (Kostopoulos et al., 2001), probably underthrust (subducted) north-eastward below the Upper Vertiskos unit metamorphosed in lower amphibolite facies (Dixon and Dimitriadis, 1994) (Fig. 13B). Subsequently, the Lower Vertiskos unit, the HP-LT Akrita and meta-gabbro basal slices, thrust top-to-the-W upon the E. Peonias Continental Margin units probably during the Late Jurassic (D₁ event) and again in the Late Cretaceous-Palaeocene (D₃ event), as indicated by the D₁ and D₃ fold families (Figs. 14.1 and 14.2b) that affect the Melissochori unit below the thrust.

Subsequently, via the Langhada-Doirani reverse fault (striking N140°E, dipping 60–80°E) (Figs. 13A and 14.2); the entire Vertiskos massif thrusts upon the Mid.-Late Triassic Nea Santa marbles (pt. 13.2 on Fig. 13B), the Melissochori flysch and Paleogene conglomerates (Kockel et al., 1977); the Langhada-Doirani fault was re-activated as a collisional fault.

Briefly, the Langhada-Doirani-Strumića fault, considered to be a Jurassic suture between the SMM Vertiskos Massif and the Vardar (Peonias) units (Karamata et al., 1994.; Kostopoulos et al., 2001; Himmerkus et al., 2006, among others), has a complex history being a normal fault in the Early Triassic, probably a subduction zone in the Jurassic (the absence of marine deposits associated with the thin ophiolitic slices makes uncertain its interpretation in terms of oceanic basin) and subsequently, a collisional fault.

Although the analysis of the SMM is out the scope of our paper, it is noteworthy to recall that the SMM Kerdilion massif suffered magmatic pulses of Late Palaeozoic, Late Jurassic and Palaeocene ages (Himmerkus et al., 2006, 2012), similar to those that affect the Vardar Palaeozoic Massifs (see Sect. 2.5; 2.3; 2.2; 2.1); the Kerdilion massif was accreted to the Vertiskos massif via the Athos-Volvi Jurassic ophiolitic suture (Fig.1) (Dixon and Dimitriadis 1994; Burg et al., 1995; Himmerkus et al., 2006; Burg, 2012) dated at ∼145–146 Ma (Himmerkus et al., 2012). The N. Rhodope gneisses, dated at ∼138–140 Ma, obducted upon the Thrace Terrane of Kerdilion affinities via the Nestos ophiolitic suture (Turpaud, 2006; Reichman and Turpaud, 2009; Papanikolaou, 2021). These data suggest a continued eastward closure of the ophiolitic basins during the Jurassic.

Figure 15 schematically summarizes our observations and conclusions on the Late Jurassic-Early Cretaceous geodynamic evolution of the Internal Hellenides. Rifting of the Internal Hellenic basement occurred in the Early Triassic (∼240 Ma) (Fig.15.1) and breaking of the continental crust around ∼220–210 Ma (Fig. 15.2). The Pelagonian-Korab micro-continent separates the Mirdita (Sect. 2.1) and the Almopias ophiolitic basins (Sect. 2.2) and, the Malarupa-Veles micro-continent separates the Almopias (Sect. 2.2) and the Guevgueli ophiolitic basins (Sects. 2.4; 2.5) (Figs. 15.2 and 15.3); the gneissic melanges at the base of the Thessaloniki ophiolitic nappe (Sect. 2.5) (Fig. 13) are probably remnants of a continental domain (Fig. 15.4) that separated the Guevgueli and the Thessaloniki ophiolitic basins. Thus, there were three, possibly four, marginal oceanic basins (Figs. 15.3, 15.4) (Mercier et al., 1975; Ohnenstetter et al., 1979; Vergély, 1984), in contrast with a model of large Tethysian-Vardar-Maliac oceanic basin (Bortolotti et al., 2013; Ferrière et al., 2016, among others). The width of the oceanic domain of Late Jurassic-Early Cretaceous age (H₄ oceanic terranes of Papanikolaou, 2013) is unknown; however, considering a life time between the beginning of spreading and closure by obduction of every basin and, a mean arbitrary spreading rate at ∼2 cm/yr, the sum of the widths of the oceanic basins gives [∼20 + 62 + 10 + 50 Myr] * 20 km/Myr = ∼2,800 km. It is a rough estimation, because of the difficulty to precisely date the opening of the basins and because of the unknown mean spreading rate.

Major transverse faults, the Voras fault zone and eventually an Oreokastro-Langhada transverse structure, cut the Vardar units (Fig. 4B). There is no axial continuity of the Triassic-Jurassic structures on both sides of these fault zones; the Voras fault is thrust by the Guevgueli ophiolitic nappe (Fig. 8A); thus, the Voras transverse structure pre-dates the Late Jurassic D₁ tectonic event; it is probably an oceanic transform fault re-activated by the subsequent tectonic events.

The Mirdita, Almopias, Guevgueli and Thessaloniki slabs dip eastward into the mantle (Figs. 15.4; 15.5; 15.6); however, the seismic tomographic images of the Aegean mantle show only two high-velocity anomaly zones. One dipping to 1500 km-depth is interpreted to be a major slab grouping the Late Cretaceous-Tertiary slabs (see Sect. 3.2) (Van Hinsberghen et al., 2005); the other, in front of the latter, is interpreted to be due either to dragging of the retreating slab (Facenna et al., 2003) or to a piece of a broken slab grouping together the Jurassic slabs of the basins sinking into the mantle (Wortel and Spakman, 2000). The second interpretation better agrees with our analyses, sinking into the mantle being coeval with the Late Jurassic-Early Cretaceous uplift of the Internal Hellenides; if true, it means that the Jurassic slabs of the subducted Late Jurassic-Early Cretaceous oceanic basins were connected at depth to a single major slab dipping eastward.

During the Late Cretaceous- Eocene, deformation was collisional.

The Mirdita ophiolites capped by laterites, were covered transgressively by the Tithonian-Valanginian flysch, they emerged again in the Hauterivian and were covered transgressively by Barremian-Aptian shallow water limestones that pass upward to Late Cretaceous pelagic limestones and Eocene terrigenous deposits (=Subpelagonian flysch of Bizon et al., 1968).

The Tithonian-Valanginian flysch in the Mali i Bardhë tectonic window (Aliaj and Bushati, 2019) (Fig. 2.I) and the Paleogene Apulian Krasta (Pindos) and Kruja (Gavrovo) flysch in the Peshkopia window, subduct below the Korab massif (Kodra, 2016) indicating that the Jurassic ophiolitic subduction was followed by an Apulian continental subduction (=collision). The Korab nappe that overthrusts the window, is overlain by folded Late Cretaceous limestones and Eocene flysch unconformably covered by deposits of the Krania basin (pt. Kr on Fig. 2.I). The Krania limestones, dated of Late Lutetian age by Nummulites and pelagic Foraminifera (Globigerina parva, Globorotalia centralis) unconformably overly the folded Ypresian-Lower Lutetian flysch of the nappe, dating the continental collision (D₃ event) of Lutetian age (∼45 Ma) (Bizon et al., 1968).

Similarly, we propose that, in the Olympus window (Godfriaux, 1962, Godfriaux 1968) (Fig.1), the Kalipefki ophiolitic body (=W. Pelagonian ophiolites), the Late Cretaceous Ossa fm. (=Mirdita carbonates) and the Olympus marbles (= Parnassus limestones, Vergély and Mercier, 1990; considered to be Gavrovo limestones by Fleury and Godfriaux, 1975) are affected by HP-LT metamorphism dated at 55–53 and 43 Ma (Schermer et al., 1990), they subducted below the Pelagonian Massif (Vergély and Mercier, 1990). Subduction progressing, the lower part of the Apulian series continued to subduct below the Korab (Pelagonian) basement while the upper sedimentary series were thrust westward forming the Thrust and Fold Belt of the External Hellenides (Van Hinsbergen et al., 2005).

The Almopias ophiolitic slices thrust upon the Pelagonian eastern margin are sealed by Late Albian-Cenomanian transgressive deposits (pt.4 on Fig. 4B) (Mercier and Vergély, 1995); they are scoured and disconformably covered by a ∼10 m-thick Late Santonian-Early Campanian Rudist (Vaccinites atheniensis) breccia overlain by Maastrichtian neritic limestones and flysch. The western (Pelagonian) slope of the Almopias basin (Kedronas unit, Fig. 4B) is also covered by a thick heterogeneous conglomeratic series of Late Turonian- Early Campanian age that reworks pebbles and olistoliths of Pelagonian origin. The sequence is overlain by ∼150 m-thick Campanian-Maastrichtian neritic limestones that vanish eastward, replaced by pelagic limestones. More to the east (Kerassia unit, Fig. 4) the conglomeratic contribution is reduced, Albian-Cenomanian limestones with Orbitolina discoidea and O. concava, are overlain by pelagic Globotruncana limestones of Coniacian-Santonian-Campanian age that are interbedded with sandstone and conglomeratic beds. In the eastern Nea Zoi unit (Fig. 4), the ophiolites are covered by pelagic calc-schists with calcified Radiolaria and Foraminifera (Rotalipora cushmani, Prae-globotruncana stephani-turbinata) of Cenomanian-Turonian age, passing upward to schists and cherts (Mercier et al., 1984, 1987), data indicate an eastward deepening of the basin. It is noteworthy that the Late Cretaceous deposits that formed the eastern slope of the Almopias basin connecting with the Upper Cretaceous Paikon carbonate platform are missing.

The Late Cretaceous Almopias sequence is metamorphosed (M₂ metamorphic event; Mercier, 1966a; Vergély, 1977); greenschist transitional blueschist facies is evidenced in the Fe-Ni laterites covering the harzburgitic slices (Mposkos et al., 1981). The sequence is affected by (B₁) folds with a ∼N110°E trend and a SSW vergence (D₃ event) (Fig. 4B), re-folded by (B₂) folds trending N60–80°E (ster. 5a, b; D₄ arrow on Fig. 4B) with a down-slip (normal) NNW vergence, due to a Tertiary gravitational exhumation (D₄ event) (see Sect. 3.3). The Eastern Almopias radiolaritic units thrust north-eastward upon the Paikon margin (green arrow on Fig. 4B) (Brown and Robertson, 2003; Tranos et al., 2007) and are re-faulted top-to-the southwest (Fig. 17.2), subsequently, the pile of Eastern Almopias units thrusts westward upon the Central Almopias units (red arrow and ster. A.9 on Fig. 4B) and, the latter upon the Western Almopias units and the Pelagonian margin (D₅ event).

Briefly, the Late Cretaceous Almopias basin formed between the Pelagonian and Malarupa continents in the Late Aptian-Albian, upon a substratum of metamorphic rocks and obducted ophiolites. The Almopias Late Cretaceous units metamorphosed in greenschist transitional blueschist facies, plus the missing eastern slope of the basin, indicate that the Late Cretaceous Almopias units subducted eastward below the andesitic-dacitic Paikon Volcanic Arc (D₃ event) (see below). The Apulian-Mirdita and Almopias slabs dip eastward below the Pelagonian and the Malarupa continental massifs, respectively; however, tomographic images of the Aegean mantle indicate a single high-velocity zone interpreted to be an Apulian slab dipping eastward to a 1500 km-depth (Sect. 2.7, Van Hinsbergen et al., 2005; Hafkeinscheid et al., 2005). If correct, it means that the Apulian-Mirdita and Almopias slabs were connected at depth to a single major slab dipping eastward.

Several Western Paikon carbonate units dipping SW form the western flank of the Paikon Massif (Sect. 2.3); the lowers (1, 2 on Fig. 8B) comprise Tithonian-Berriasian limestones transgressively covered by Late Cretaceous dolomitic Rudist limestones. In Southern Paikon, they are imbricated E.-W. striking slices (Fig. 8B), of Jurassic marbles covered unconformably by an Early Cretaceous fluvio-deltaic conglomeratic formation, disconformably overlain by Late Albian to Early Turonian dolomitic limestones. In the upper units (3 on Fig. 8B), they pass to a Turonian- Campanian pink flysch and, to Late Campanian-Maastrichtian limestones and flysch (Mercier, 1968a; Brown and Robertson, 2003). The Paikon Kastaneri unit (5 on Fig. 8B) thrusts top-to the-W upon the Paikon Central Sequence, it is covered via minor tectonic contacts by the Ghriva unit (6, Fig. 8B) with patches of transgressive Late Cretaceous limestones (Ricou and Godfriaux, 1991). The Gola Tchouka unit (4 on Fig. 8B) (Bonneau et al., 1994) and the Gandatch Mt. (Fig. 8B) (Mercier et al., 2005) are also Late Cretaceous crystalline Rudist marbles overlying after emergence foliated marbles; they indicate that the Paikon eastern flank was also overlain transgressively by a Late Cretaceous carbonate platform, subsequently sliced during the Tertiary tectonic events. In the Voras Mts., at Duditsa site (Fig. 8A) and on the Paikon western flank (Fig. 8B), the Late Cretaceous marbles are intruded by Early Tertiary monzo-rhyolitic dykes and calc-alkaline micro-granites-tonalites (Soldatos, 1955, Mercier, 1968b).

The Western Paikon units are a pile of slices in greenschist metamorphic facies, affected by folds trending N110°E with a NE vergence (D₃ event), re-folded by (B₂) folds with a ∼NNW vergence due to a gravitational exhumation (D₄ event) (ster.A3a, b, A4 and arrows on Fig. 8B).The pile of Western Paikon units thrusts top-to-the NE upon the Central Sequence; the basal thrust trace extends northward from the southern Paikon Massif (Fig. 8B) and, from the Pinovon Massif (Galeos et al., 2003), as far as north of the Malarupa Massif; it is subsequently overthrust top-to-the-W by the Eastern Paikon Kastaneri thrust (Fig. 8A). Thus, the Western and Eastern Paikon units are successive convergent thrusts, in contrast with divergent limbs of a Paikon anticline proposed by Brown and Robertson, 2003.

The Paikon Mt., seen from far, may appear to be a tectonic window with a Jurassic Gandatch anticlinal core (Bonneau et al., 1994 and Ferrière et al., 2001’s tectonic window models); however, when crossing the mountain it is a folded syncline of Late Cretaceous Rudist marbles (Fig. 8B) (crossing the mountain is not easy, however, there is a mule path across it). The nappe supposed to surround the anticlinal core to the south is, in fact, made of a NNW-SSE striking segment of the Kastaneri sheet thrust top-to-the W (brown colour on Fig. 8B), plus E-W trending segments of the Western Paikon slices (green colour) thrust top-to-the NE. It is clearly seen on a new road at pt. 22 on Figure 8B, that the Kastaneri sheet does not connect with the Western Paikon slices to surround a core, but thrusts these slices (Mercier et al., 2005); the Kastaneri thrust continues southward between pt. 22 and 23; the Paikon Massif is not a Tertiary tectonic window.

An Early Tertiary dacitic Paikon Volcanic Arc formed above the continental subduction of the Apulian-Almopias series (D₃ event); volcanics intrudes the Late Cretaceous marbles of the Voras Mt. (Fig. 8A), they are dated at 55–48 Ma in the N. Macedonia Kopaonik massif (Rb-Sr age; Lanpherer and Pamić, 1992). In Chalkidiki, the Sithonia, Ouranopolis, Ierissos and Gregoriou granitic massifs, dated at ∼56 and 43 Ma (de Welt et al., 1989, Christofides et al., 2000: Bébien et al., 2001), are Early Tertiary plutons also associated with the continental subduction of E. Peonias-(?) SMM units. During the continental subduction, the Malarupa continental crust increased in thickness and exhumed forming a dome as shown by the geometry of the Gola Tchouka, Kastaneri and Guevgueli Paleogene basal thrust traces that surround the termination of the Paikon-Malarupa dome (Fig.8A).

The E. Peonias series of Mid.-Late Cretaceous age are not observed in Greek Macedonia; they are known in N. Macedonia; near Šopur, they are Albian-Cenomanian molasses (Maksimović and Sikošek, 1954; Mercier, 1968a) that overly transgressively the Late Jurassic Štip granite (∼155 Ma, Soptrajanova, 1967) (Fig. 16). Near Gaber village, Cenomanian-Senonian sands and limestones form a syncline that overlies unconformably the E. Vardar (Peonias) units affected by ∼NNW-SSE trending schistosity planes (Fig. 16) that continue, in N. Macedonia, the Greek Peonias units (Fig. 11) affected by the NNW (S₁) schistosity of the Jurassic D₁ event (Sect. 2.5).

WNW-ESE trending km-scale incline folds affect the Late Cretaceous Gaber deposits (Fig. 16.2), they are unconformably overlain by Vardar molassic deposits, folding results from the collisional D₃ event. Another km-scale WNW-ESE trending fold in greenschist metamorphic facies with a SSW vergence affects the E. Peonias Aspri Vrissi unit (Anilio fold, D₃ event) (Fig. 13); it refolds ∼10 cm-(B₁) folds with (S₁) schistosity trending NNW-SSE to N-S resulting from the Jurassic D₁ event (Vergély, 1984) (see Sects. 2.5, 2.4). The schistosity planes (S₁), proposed to be of Late Cretaceous age, are overlain unconformably by the Gaber Albian-Senonian deposits (Fig. 16.1) and thus, do not support the proposed Late Cretaceous gravitational emplacement of the schistose units due to the exhumation-extension of the near-by SMM metamorphic core complex (Kyndonakis et al., 2015), The (S₁) schistosity is of Jurassic age, it affects a Chortiatis diorite dated at 159 ± 4 Ma (see Sect. 2.5) and the Jurassic formations that are reworked as pebbles in the conglomerates of the Late Jurassic Erikinos fm. (see analysis in Sects. 2.4, 2.5); the (S₁) schistosity results from the D₁ tectonic event. In contrast with a model of Late Cretaceous gravitational emplacement, the E. Peonias units were first affected by the D₁ Jurassic folding, coeval with the obduction, superimposed by the Late Cretaceous-Early Eocene collision (D₃ event).

In summary, the collisional D₃ event is dated of Late Senonian-Early Eocene age (∼70/65–45 Ma), by the molassic Vardar deposits that overly unconformably the folded Cenomanian-Senonian Gaber deposits; it is synchronous with the collisional event in the Korab Massif (56–45 Ma) and in the Paikon Massif (∼70/65–45 Ma) (see Sect. 3). By contrast with a model of Late Cretaceous of gravitational emplacement related to the exhumation-extension of a near-by SMM metamorphic core complex, the E. Peonias units were folded in the Late Jurassic (D₁ event) and superimposed by Late Cretaceous-Early Eocene folds (D₃ event) (Vergély, 1984). Change from the Jurassic NNW-SSE to the Late Cretaceous WNW-ESE trending fabric (∼61 to 45 Ma) is due to the progressive change to a NNE-SSW/N-S direction of the Africa-Europe convergence (∼67/65 to 46 Ma) (Rosenbaum et al., 2002).

Following the collisional D₃ event (∼56–45 Ma), the Krania (Fig. 2) and Vardar (Fig. 11) molassic basins opened in the Vardar zone at 45 Ma.

In Greek Macedonia (Fig. 11), the substratum of the Vardar Artzan basin is transgressively covered by ∼10 m-thick reef limestones of Late Lutetian age (45 Ma) and by Priabonian marls (Bourcart, 1916; Mercier, 1960, Mercier 1968a) that are overlain unconformably by 150 m-thick Late Priabonian-Early Oligocene conglomeratic sandstones (Artzan formation). In the lower Babuna valley, near Veles (Fig. 1), a Late Priabonian-Early Oligocene transgressive series overlies the Central Vardar Veles metamorphic series (Bizon et al., 1966; Mercier, 1968a); both sites evidence a Late Priabonian D₅ event (∼36–34 Ma; Mercier, 1968a). The 1500–2000 m-thick Vardar deposits, subsided in extension; however, normal faults are poorly observed being obscured by subsequent folding and by thrusting of gneissic and serpentinite slices, the Dedeli-Rabrovo slice (Fig. 17.4; Pendzerkosski, 1963; IFP, 1967; Mercier, 1968a) and the several km-long reverse Leskovića reverse fault (Fig. 16). Subsequently, Vardar marine detrital sediments deposited again in N. Macedonia, passing upward to Paleogene limnic and continental sandstones and pelites (more than 1000 m thick) (Inst. Geol. Acad. Sci. Serbe, 1954; Mercier, 1968a); they deposited in an extensional tectonic regime as indicated by seismic profiles in the western Vardar basin, east of the N-S Verria fault (Hunt company, in I.F.P. 1967, full-page plate 31).

An important volcanic activity within and around the basins follows the collisional D₃ event (∼56–45 Ma); in the Vardar Trough, calc-alkaline latites, andesites and dacites (∼45–36 Ma) are interbedded with or covered by Priabonian molasses (Inst. Geol. Acad. Sci. Serbe, 1954; Mercier, 1960, Mercier 1968a). Micro-quartzodioritic dykes interbedded with or covered by Priabonian deposits are also observed on the Serbo-Macedonian Asogovo-Besna Massif (Boev and Yannev, 2001). A second magmatic pulse follows the Late Priabonian D₅ event; in the Vardar Trough, ultrapotassic rhyolitic and trachytic lavas and domes of Early Oligocene age (33 ± 5 Ma) outcrop in the vicinity of the Doirani Lake (Fig. 11) (Stojanov and Svešnikova, 1985; Stojanov and Serafinoski, 1990; Asvesta, 1992); along the SMM margin, latites, andesites, dacites and plutonic rocks of Early Oligocene age form the N. Macedonian Kratovo-Zletovo volcanic massif (35 to 29 Ma, Stojanov and Serafinoski, 1990; Karamata et al., 1992; Boev and Yannev, 2001). Calc-alkaline volcanism may persist in the Late Miocene-Pliocene, in the Almopias-Kožuf Massif of the Voras Mts.; they are shoshonites, trachytes, latites, andesites (Soldatos, 1955, Mercier, 1968b) of Late Pliocene age (Mercier and Sauvage, 1965), dated at 6.5 to 1.8 Ma (Kolios et al., 1980).

Briefly, the Vardar Trough is an andesitic basin, with magmatism typical of a continental Andean continental margin (D₃ event); opening of the basin was probably activated by roll-back of the Paikon slab (not excluding a possible contribution of extension due to exhumation of a near-by SMM metamorphic core complex, Dinter and Royden, 1993; Brun and Soukoutis, 2004). The high subsidence of the Vardar basin occurred in an extensional tectonic regime, however, interrupted by slicing and folding during the Late Priabonian D₅ event (∼35–30 Ma) and the Early Miocene D₆ event (∼23–20.5 Ma) (see Sect. 4.1), synchronous with the events of the External Hellenides.

Our paper concerns the North Aegean islands (N. Sporades, Lesbos, Limnos, Samothraki, Thasos) and the Macedonian (Mesohellenic, Jannitsa, Thessaloniki) and Thraki basins; the Southern and Central Aegean domain, submitted to a major N-S stretching, to roll-back and tearing of the Aegean slab are out the scope of the paper. They are addressed by numerous articles (see Van Hinsbergen et al., 2005; Ring et al., 2010; Royden and Papanikolaou, 2011; Jolivet et al., 2013, 2015 among others and ref. herein).

Subsequently to the major Late Priabonian D₅ event, weak compressional deformations affected the Paleogene Vardar basin and the N. Aegean domain (D₆ event). In Limos island, Late Oligocene-Early Miocene volcaniclastic deposits are folded and unconformably overlain by Early Miocene volcanic deposits (22.6 ± 0.7 Ma; Innocenti et al., 2009; Pe-Piper et al., 2009). In the Thessaloniki-Jannitsa basin, the Vardar Eocene-Oligocene deposits (1500–2000 m thick), observed by seismic prospection and in boreholes (Hunt company, in I.F.P. 1967), are folded and overlain unconformably by conglomerates interbedded with lignitic beds (Moskopotamos series) of Lower Burdigalian- Mid. Serravallian age (I.F.P. 1967; Benda and Steffen, 1981; Lalechos, 1986) and by Pontian conglomerates (Vallesian-Turolian, ∼11–7 Ma) (Mercier et al., 2005). In the Mesohellenic basin, compressional structures are also evidenced by unconformities between the Lower and the Upper Meteors Conglomerates fm., followed by the emersion of the latter formation (Brunn, 1956; I.F.P. 1967). The unconformities are dated of Aquitanian-lower Burdigalian (20–20.5 Ma) and of lower Langhian (16–15 Ma, Sorel et al., 1992) ages by numerous bio-stratigraphical data (see review of the data in Ferrière et al., 2013). The compressional events, dated between ∼23 and 20.5 Ma are contemporaneous with those of the Thrust and Fold Belt of the External Hellenides: the Aquitanian D₆ event with folding of the Gavrovo massif and, the lower Langhian event with thrusting of the western Ionian sheets (Brunn, 1956; Aubouin, 1959; IGSR-IFP, 1966). Strike-slip faulting (with WNW-ESE extension-NNE-SSW shortening) was active in the Mesohellenic basin during the Early Oligocene-Early Miocene (Vamvaka et al., 2006) and in the Vardar zone (Lalechos, 1986; Pamic, 2002); they are followed by E-W extensional deformations.

We do not present a detailed analysis of the Neogene-Quaternary normal faults of the N. Aegean domain that was published previously in Mercier, 1981; Lyberis, 1984; Mercier et al., 1987, 1989; Pavlides et al., 1990. In Limnos Island, the Early Miocene volcanic deposits (22.6 ± 0.7 Ma, see above) affected by the compressional deformations are cut by normal faults resulting from a WNW-ESE trending extension, followed by a NE-SW trending extension. The youngest N. Aegean deposits affected by the WNW-ESE extension are of Tortonian age in Thassos island and those affected by the NE-SW extension are of Pliocene age in Lesbos island (Mercier et al., 1989); the N-S extension is presently active as indicated by the seismic faults of the 1978 Thessaloniki earthquake (Mercier et al., 1983). In N. Macedonia, ultra-potassic rocks of the shoshinitic series, having within-plate characters, dated at 9.5 Ma and ∼5.5 Ma (Boev and Yanev, 2001), are coeval with the Late Miocene-Pliocene extension.

Analyses of the N. Aegean faults at km- to 10 m-scale, conducted by numerical inversion of striations families supposed to be parallel to the tangential stresses resolved on the fault planes, separate normal faulting events with successive N300°E ± 09°, N229°E ± 08° and N351°E ± 07°trending tensional directions (ster.11.9; 11.10; 11.11 on Fig. 11) (Mercier et al., 1989). The Neogene-Quaternary extension began around ∼21–20 Ma; the three events are approximately of Early Miocene, Pliocene-Lower Pleistocene and Mid. Pleistocene-Recent ages (Mercier et al., 1989). The N. Aegean extension is controlled by the back-arc extension of the Aegean Arc and by the penetration of the North Anatolian strike-slip fault into the basin around 10–9 Ma (Šengör et al., 2005). Transition from the WNW-ESE to the SW-NE/S-N extensional directions does not result from rotation of the North Aegean material because paleomagnetic studies have shown that no rotations have affected the area since the Lower Miocene (Koundopoulou and Lauer, 1984; Kissel et al., 1986). Paleomagnetic studies showed also that the Ionian branch of the Aegean Arc has rotated 30° anticlockwise since the Tortonian (Kissel and Laj, 1988) and the Lycian branch 50° clockwise (Kissel et al., 1985; Van Hinsbergen et al., 2005); Crete had undergone no rotations. Ajusting these branches, the arc boundary was approximately linear, trending ∼E-W during the Early Miocene (Mc Kenzie, 1972; Le Pichon and Angelier, 1979; Mercier et al., 1989; Gautier et al., 1999). The WNW-ESE extension is orthogonal to the Lower Miocene ∼N-S direction of convergence (Dercourt et al., 1986; Rosenbaum et al., 2002), a situation often observed along continental subduction margins (see for example the Peruvian Andean margin, Sébrier et al., 1988). Several hundred kms of ∼N-S stretching in the Central and Southern Aegean basins has migrated the margin of the Aegean Arc southward (see Ring et al., 2010; Jolivet et al., 2015 among others and ref. herein); we suggest that this major stretching has driven the direction of the North Aegean extension from WNW-ESE to NE-SW/N-S. Change in the extensional direction in the N. Aegean domain occurred between the Tortonian and the Pliocene, between ∼12 and 6 Ma.

Below, we resume the main results obtained by analysing the Internal Hellenides along the Albanian-Macedonian W-E transect.

• 1. Radiolaritic MORB tholeiitic basins opened in the Early Carnian-Mid. Norian (∼237–220 Ma, Sect. 2.2) in the Almopias zone and, in the Rhaetian (210–206 Ma, Sect. 2.5) in the E. Peonias zone; the continental crust broke and oceanic crust and mantle were denuded during the Mid.-Late Jurassic (∼174 Ma) (Fig. 15.1 and 15.2).

• 2. The Pelagonian microcontinent separates the Mirdita and Almopias oceanic basins (Sect. 2.1) and the Malarupa-Veles microcontinent separates the Almopias basin and Guevgueli back-arc basin (Sect. 2.6); remnants of a probable continental domain, at the base the Thessaloniki sheet (Sect. 2.5), separate the Guevgueli and Thessaloniki basins. Thus, there were three, probably four oceanic basins between the Apulian continent and the SMM western margin (Fig. 15.3;,15.4), in contrast with a proposed model of single large “Tethysian-Vardar-Maliac” oceanic basin; basins opened during a major event, in the Late Jurassic (174 to 160 Ma).

• 3. During the Jurassic, the crusts of the W. Pelagonian (Mirdita) oceanic basins first thrust upon the Korab western margin and subsequently westward upon the Mirdita basin; they subducted below the Korab-Pelagonian continent and reappeared in the Korab-Pelagonian tectonic windows (see Katsikatsos, Mercier and Vergély, 1976). The Almopias oceanic crust subducted below the Malarupa-Paikon continent and a Paikon Volcanic Arc formed above the subduction (Fig. 15.4). The Guevgueli back-arc basin (Sect. 2.3) opened at 164–160 Ma above the subduction of the Paikon Arc (Fig.15.4). Subduction progressing, the low K-tholeiitic volcanic centre of the Paikon Arc was driven by the load of the Almopias slab into the subduction zone, below the Paikon rhyolitic volcanic centre (Fig. 15.5).The E. Peonias (Chortiatis-Thessaloniki) basin opened in the Late Jurassic (Fig. 15.4) and, subducted below E. Peonias Continental margin of the basin (Sect. 2.5).

• 4. The oceanic crusts obducted upon the western margins of the basins (D₁ event) and subsequently exhumed gravitationally westward (D₂ event) (Fig. 15.5 and 15.6); the Mirdita nappe obducted between 174 and 160 Ma, the Almopias nappe between 168 and 164–160 Ma, the Guevgueli ophiolitic nappe at ∼157 Ma and the ophiolitic Thessaloniki nappe subsequent to the Late Jurassic (159±4 Ma), pre-dating the Late Tithonian-Berriasian.

• 5. The Internal Hellenides emerged in the Late Jurassic-Early Cretaceous and were overlain unconformably by Mid.-Late Cretaceous-Tertiary deposits.

• 6. In the Early Tertiary (56–45 Ma), Apulian-Mirdita units subducted below the Pelagonian continent (D₃ collisional event); continental subduction progressing, the lower part of Apulian Tertiary sequence continued to subduct and, emerged into the Pelagonian windows (Godfriaux, 1962, Katsikatsos et al., 1977; Vergély and Mercier, 1990), metamorphosed in HP-LT facies dated at 55–53 Ma and 45 Ma the upper part was thrust westward forming the Thrust and Fold Belt of the External Hellenides. The Late Cretaceous Almopias units subducted below the Malarupa continent (D₃ event) and, a Paikon Volcanic dacitic Arc formed above the subduction. The Jurassic tectonic structures were reactivated by the Late Cretaceous-Early Eocene and subsequent events, to become collisional deformations.

• 7. Following the collisional D₃ event (56–45 Ma), the Internal Hellenides were overlain unconformably by the Krania and Vardar molassic basins. The Vardar basin is an andesitic basin, it opened at 45 Ma in the back of the Paikon Volcanic Arc, above the subduction of the Apulia-W. Vardar units; it opened in an extensional tectonic regime driven by roll-back of the Paikon slab, with two volcanic calc-alkaline pulses, at 50–43 Ma and 36–30 Ma. It was subsequently submitted to the compressional Late Priabonian D₅ event (∼36–34 Ma) and to the Early Miocene D₆ event (∼23–20,5 Ma), synchronous with the tectonic events of the Thrust and Fold Belt.

• 8. Extension in the North Aegean domain is controlled by the rollback of the Aegean slab and by the penetration of the North Anatolian fault in N. Aegean; it began at ∼21–20 Ma with WNW-ESE then NE-SW/ N-S trending extensional directions. Change in extensional directions is not due to rotation of the Aegean material because paleomagnetic studies show that no rotation have affected the area since the Early Miocene. The WNW extension is normal to the ∼N-S direction of convergence; we propose that the NE-SW/N-S extensions were driven by the several hundred-km of N-S stretching in the Southern and Central Aegean basins. In N. Aegean, change from WNW-ESE to NE-SW/N-S extension occurred between the Tortonian and the Pliocene (between 12 and 6 Ma). The SMM uplifted and, was submitted to extension, not analysed in the present paper.

The field work has been supported by the C.N.R.S. (French National Centre of Scientific Research) and by I.G.M.E. (Institute for Geology and Mineral Research, Athens). We thank Prof. T. Roundoyannis (Polytechnic Univ., Athens), Ing. K. Simeakis (IGME), Dir. G. Miranovski and Dir. N. Stolić (Geoloski Zavod Skopje) for providing us geological documentation. We are grateful to our colleagues of Aristotle University of Thessaloniki for amicable welcome. Thanks to our guides “Agoyatis” Ath. Stoyiannis (Mandalos village), Chr. Soulis (Ghriva vil.), Stef. Chrisoulidhis (Chorighi vil.) that, with their mules, have accompanied us in the 60’years when there were no roads and no villages in the high mountains of Greek Central Macedonia. We are grateful to Prof. Bruce H. Purser and Frederic J. Mercier that have improved a first draft of our paper; thanks to Prof. Papanikolaou and Dr. Bonev whose reviews have greatly increase the presentation of the paper.

In the Voras Mts, the Tzena (Evgueni) carbonate platform suffered the same deformations D1 and D2 that the Paikon Central Sequence. The regional flow schistosity (S1) parallel to the stratification (S0), dips 30 to 50° WNW (ster. a) with lineations trending NNW (ster. b) (D1 event) are re-folded by spectacular km-scale recumbent folds observed in the SE facing cliff of the Tzena summit (A) (Mercier, 1973). The kilometric scale Tzena folds (B) with an eastern vergence (D2 event) (sections a-b and c-d) (D2 event) formed during the exhumation of the subducted Paikon Volcanic Arc; in the Paikon Massif, exhumation pre-dates the Latest Tithonian-Berriasian (drawn from Mercier,1968c, Mercier 1973 and Landsat images).

Bourcart, 1916; Mercier, 1960