3.a. Adria-derived units

3.a.2. Boeotian Unit

The Boeotian Unit is represented by a thick Triassic–Jurassic carbonate series topped by a Cretaceous flysch. This geotectonic zone, characterized by pelagic sedimentation since late Middle Jurassic time (Celet & Clément, Reference Celet and Clément1971; Celet, Clément & Ferrière, Reference Celet, Clément and Ferrière1976; B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983; Robertson et al. Reference Robertson, Clift, Degnan and Jones1991; Thiébault et al. Reference Thiébault, Fleury, Clément and Dégardin1994; Danelian & Robertson, Reference Danelian and Robertson1998; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015), is classically located between the Pelagonian Platform, to the east, and the Parnassus Platform, to the west.

The stratigraphy of the Boeotian Unit is substantially comparable to that of the external Pelagonian Platform until the Middle Jurassic. A differentiation occurred during Middle Jurassic times when the Boeotian platform started drowning (B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983; Celet, Clément & Ferrière, Reference Celet, Clément and Ferrière1988). Drowning in the Boeotian Zone was synchronous with broad doming of the Parnassus and westernmost Pelagonian platforms that suffered emersion and laterization (b1 laterite horizon; Papastamatiou et al. Reference Papastamatiou, Tataris, Vetoulis, Katsikatsos, Lalechos and Eleutheriou1962b , Reference Papastamatiou, Tataris, Kounis, Maragoudaxis, Monopolis, Albadakis and Koukouzos1971; Maratos, Rigopoulos & Athanassiou, Reference Maratos, Rigopoulos and Athanassiou1967). After the emersion event shallow-water sedimentation resumed in the Parnassus Zone while the western borders of the Pelagonian Platform were characterized by progressive deep-water conditions, thus representing a transitional zone to the evolving Boeotian Basin to the west (Fig. 4).

Similarly, the limestone sequence cropping out in the Koziakas area (Central Greece; Fig. 1) is regarded as a transition from the Pelagonian Platform towards the Boeotian and Pindos pelagic basins to the west (R. Ardaens, unpub. Ph.D. thesis, Univ. Lille, 1978; Jaeger, Reference Jaeger1980; Ferrière, Chanier & Ditbanjong, Reference Ferrière, Chanier and Ditbanjong2012).

Pelagic sedimentation in the Boeotian Zone started with radiolarite and siliceous shale, locally pre-dated by cherty limestone, and finally followed by the turbidite deposits of the ‘Boeotian Flysch’ (Celet & Clément, Reference Celet and Clément1971). The primary contact between the Boeotian Flysch and the underlying radiolarite and limestone basement is rarely preserved because this horizon represented a preferential site of décollement during the Cretaceous and Tertiary deformation phases. However, Danelian & Robertson (Reference Danelian and Robertson1998) reported a primary contact near Makariotissa monastery, where the Triassic to Lower Jurassic shallow-water limestone of the Boeotian Zone is covered by a pelagic sequence composed of limestone and marls of Oxfordian–Kimmeridgian age, which in turn pass upwards to a shale and radiolarite sequence dated to the middle Oxfordian – lower Tithonian and finally to the uppermost Jurassic–Cretaceous Boeotian Flysch (Clément, Reference Clément1971, Reference Clément1972; B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983; Terry & Mercier, Reference Terry and Mercier1971; Richter, Müller & Mihm, Reference Richter, Müller and Mihm1996; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015).

The Boeotian Flysch was subdivided according to its stratigraphic features, age and composition into two distinct turbiditic systems: the Lower Boeotian Flysch (LBF) of Tithonian–Aptian age and the Upper Boeotian Flysch (UBF) of Cenomanian–Coniacian age (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). The LBF was further subdivided into three lithostratigraphic units depicting, as a whole, a thickening– coarsening-upward trend. A composite provenance from oceanic-derived units with minor coarse- and fine-grained inputs from dismantling carbonate platforms is clearly recognizable (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). Ophiolite detritus was derived from the ophiolitic nappe obducted onto the Pelagonian Zone (Celet, Clément & Ferrière, Reference Celet, Clément and Ferrière1976; Richter, Müller & Mihm, Reference Richter, Müller and Mihm1996; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015), whereas carbonate inputs derived from the neighbouring Parnassus Platform and from ephemeral carbonate platforms set onto the advancing ophiolite nappe (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015).

The LBF and UBF are separated by a sedimentary hiatus of middle Aptian – late Albian age (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). Near Koumaritsi village (Mount Iti, Figs 2, 5) the lowermost UBF sediments seal, through an unconformity surface, deformations of the underlying LBF turbiditic system (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015).

Figure 5.

(a) Geological map of the Koumaritsi area (modified after Marinos et al. Reference Marinos, Papastamatiou, Maratos, Melidonis, Andronopoulos, Tataris, Betoulis, Katsikatsos, Maragoudakis and Lalechos1967 and Richter Müller & Mihm, Reference Richter, Müller and Mihm1996), see Figure 2 for location. (b) Sketch structural cross-section of the Mount Iti area (modified after Wigniolle, Reference Wigniolle1977). (c) Cross-section across the Koumaritsi area.

The UBF in the studied area is represented by two lithostratigraphic units depicting a fining- and thinning-upwards sequence of Cenomanian–Coniacian age (Richter, Müller & Mihm, Reference Richter, Müller and Mihm1996; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). From the Santonian the UBF sequence shows in turn increasing breccia layers composed of neritic limestone, laterite and ophiolitic material that are then unconformably capped by post-nappe sequences consisting of Upper Cretaceous shallow carbonate (Second Mesoautochthonous Cycle; Wigniolle, Reference Wigniolle1977; Jacobshagen, Reference Jacobshagen1986; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). Steuber et al. (Reference Steuber, Gotzes, Raeder and Walter1993), although not differentiating the UBF succession from the Mesoautochthonous Complex, recognized a gradual transition from the deep-water clastic successions (namely the UBF) and the shallow-water carbonate succession of the Aptian–Coniacian portion of the Mesoautochthonous Complex (First Mesoautochthonous Cycle, see Section 3.c; Fig. 4).

Although deciphering palaeocurrents is complicated by superimposed tectonic events it is possible to assume a provenance from northern sectors for UBF turbidite currents.

3.b. Ocean-derived units

Obduction is the most widely accepted emplacement mechanism for the ophiolitic bodies widespread throughout the Dinaric–Hellenic orogen. A major constraint for this kind of process is the presence at the base of the ultramafic mantle section of slices of HT–LP metamorphic rocks (amphibolite sole). Metamorphism is assumed to have been acquired by underplated units soon after the inception of intraoceanic thrusting of a young and exceptionally hot ophiolite nappe originated near the spreading axis (Spray et al. Reference Spray, Bébien, Rex, Roddick, Dixon and Robertson1984; Carosi et al. Reference Carosi, Cortesogno, Gaggero and Marroni1996; Searle & Cox, Reference Searle and Cox1999; Dimo-Lahitte, Monié & Vergély, Reference Dimo-Lahitte, Monié and Vergély2001; Myhill, Reference Myhill, Wakabayashi and Dilek2011).

Continuing thrusting of the ophiolite nappe was also linked with the production of a ‘tectono-sedimentary mélange’ supplied by both the ophiolite thrust sheets and from off-scraping of pelagic shaly-siliceous sediments from the overthrust ocean bottom. Major slide block and tectonic slivers originated from the ophiolite thrust sheets and scraped-off sediments from the underplating of both oceanic and continental margin units were also embedded in the mélange matrix. Today, ophiolites and obduction-related mélange are arranged along two main belts separated by the Pelagonian continental units and its continuation to the north (i.e. the Korab Unit in Albania and Drina–Ivanjica Unit in Serbia; Fig. 1). However, this distinction is not evident in Bosnia where the Drina–Ivanjica and Durmitor units are overthrust by a unique and continuous ophiolite nappe (Schmid et al. Reference Schmid, Bernoulli, Fügenschuh, Matenco, Schefer, Schuster, Tischler and Ustaszewski2008).

In the studied area the Ophiolitic Unit is mostly represented by sheared shaly-sandy matrix chaotically enveloping rock fragments of both oceanic and continental origin (Ophiolitic Mélange) and dismembered portions of the ophiolite nappe mainly represented by serpentinized harzburgite and lherzolite. The main ophiolite occurrences are located in Mount Iti, the Kallidromon Mountains and near Aliki Village on the Corinth Gulf (Figs 2, 6). Large slivers of amphibolite sole were observed within the ophiolitic mélange in a small abandoned quarry near Agios Georgios village and in small blocks east of Agia Triada village (Fig. 6). Amphibolite sole underlies the ophiolite nappe of Mount Iti (northwest of Pyra village) and in the Gerania Mountains (Kaplanis et al. Reference Kaplanis, Koukouvelas, Xypolias and Kokkalas2013).

Figure 6.

Tectonic–geological map of Boeotia (location in Fig. 2). New data and observations integrate and modify former mapping by Papastamatiou et al. (Reference Papastamatiou, Tataris, Vetoulis, Katsikatsos, Lalechos and Eleutheriou1962b , Reference Papastamatiou, Tataris, Kounis, Maragoudaxis, Monopolis, Albadakis and Koukouzos1971), Maratos, Rigopoulos & Athanassiou (Reference Maratos, Rigopoulos and Athanassiou1965, Reference Maratos, Rigopoulos and Athanassiou1967) and B. Clément, unpub. Ph.D. thesis, Univ. Lille (1983). DSTF – Distomon Fault; ATRF – Agia Triada Fault; ALKF – Aliki Fault.

3.c. Upper Cretaceous sedimentation (Mesoautochthonous Complex)

A thick sequence of Aptian to Upper Cretaceous shallow-water carbonates sealed older deformations recorded in the tectonic stack composed by the ophiolitic, Pelagonian and Boeotian units. These sediments, known as the Mesoautochthonous Complex (Jacobshagen, Reference Jacobshagen1986) or sediments of the ‘Cretaceous transgression’ (Photiades et al. Reference Photiades, Carras, Bortolotti, Fazzuoli and Principi2007; Fazzuoli et al. Reference Fazzuoli, Menna, Nirta, Carras and Principi2008), are widespread in all of the Hellenic–Dinaric Belt with a similar tectonic position and sedimentation age. The Mesoautochthonous Complex also unconformably covers tectonized remnants of an older carbonate platform possibly sedimented above the ophiolite nappe during obduction and seals the syn-obduction deformations they recorded (Bortolotti et al. Reference Bortolotti, Carras, Chiari, Fazzuoli, Photiades and Principi2004a ; Fazzuoli & Carras, Reference Fazzuoli and Carras2007). The existence of syn-obduction Late Jurassic to Valanginian carbonate platforms is testified to throughout the Dinaric–Hellenic orogen by tectonized remnants above ophiolites (Bortolotti et al. Reference Bortolotti, Carras, Chiari, Fazzuoli, Photiades and Principi2004a ; Schlagintweit et al. Reference Schlagintweit, Gawlick, Missoni, Lein, Hosha, Sudar, Ercegovac and Grubic2006; Photiades et al. Reference Photiades, Carras, Bortolotti, Fazzuoli and Principi2007) and, in most of the cases, by clasts and slide blocks resedimented, after total dismantling episodes, within the ophiolitic mélange, or within Late Jurassic – Early Cretaceous foreland basins (e.g. the Pogari Formation in Serbia, Firza Flysch in Albania, Lower Boeotian Flysch in Greece; Blanchet et al. Reference Blanchet, Durand Delga, Moullade and Sigal1970; Schlagintweit et al. Reference Schlagintweit, Gawlick, Missoni, Hosha, Lein and Frisch2008; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015).

Based on their age, substratum and main sequence boundaries, we propose a subdivision of the Mesoautochthonous Complex into two main sedimentary cycles. Sedimentation above the ophiolite nappe and the Pelagonian Unit generally starts in Aptian/Albian–Cenomanian times above a level of karstic or redeposited laterites lasting, without regional unconformities, until Coniacian–Santonian times (First Mesoautochthonous cycle, MES₁; B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983; Photiades et al. Reference Photiades, Carras, Bortolotti, Fazzuoli and Principi2007; Fazzuoli & Carras, Reference Fazzuoli and Carras2007). A younger sedimentation cycle of late Santonian – Maastrichtian age, also underlined by an unconformity, is present above the ophiolites and the UBF (Second Mesoautochthonous cycle, MES₂; E. Wigniolle, unpub. Ph.D. thesis, Univ. Lille, 1975; Steuber et al. Reference Steuber, Gotzes, Raeder and Walter1993).

In Boeotia and Attica, stratigraphic contacts of the MES₁ were described in the Avlona area (northern Attic region), above ophiolites (B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983), and in the Aliartos (Fig. 2) and Mandra-Megara areas (Elefsina Gulf, Attica region) above the Pelagonian carbonates (B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983; Steuber et al. Reference Steuber, Gotzes, Raeder and Walter1993). The MES₁ is coeval and partially heteropic with the UBF (Steuber et al. Reference Steuber, Gotzes, Raeder and Walter1993) that in turn unconformably seals the deformed Tithonian–Aptian LBF (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) (Fig. 4).

Primary contacts of the MES₂ above the UBF were preserved in Mount Iti (E. Wigniolle, unpub. Ph.D. thesis, Univ. Lille, 1975) (Fig. 5) and above ophiolites located northeast of Levadhia (Dionysos–Pavlos area; Steuber et al. Reference Steuber, Gotzes, Raeder and Walter1993; Fig. 2). In the Levadhia and Jerolekas (northern Parnassus; Fig. 2) areas the basal primary contact of the MES₂ is underlined by ophiolite-bearing conglomerate deposits with abundant laterite clasts and a rudist fauna of Santonian–Campanian age (Steuber, Raeder & Walter, Reference Steuber, Raeder and Walter1994; Steuber, Reference Steuber1999; Morycowa & Marcopoulou-Diacantoni, Reference Morycowa and Marcopoulou-Diacantoni2002). Conglomerates unconformably cover the lower Cenomanian MES₁ limestone after an episode of emersion and laterization (Steuber, Raeder & Walter, Reference Steuber, Raeder and Walter1994).

Similarly to the Parnassus Unit, reddish marls and shale of middle–late Paleocene age underlie the siliciclastic turbiditic sedimentation of late Paleocene–Eocene age (Pelagonian Flysch) above the Campanian–Maastrichtian limestone of the MES₂ (Wigniolle, Reference Wigniolle1977; Richter, Müller & Risch, Reference Richter, Müller and Risch1996).

4.a. Siliceous deposits and mass flow deposits on top of the Jurassic carbonates of the Pelagonian Unit (Iti and Kallidromon areas)

In the Iti area, E. Bléhaut (unpub. Ph.D. thesis, Univ. Lille, 1975) and E. Wigniolle (unpub. Ph.D. thesis, Univ. Lille, 1975) described several sections of detritic material stratigraphically topping the Pelagonian limestone. On the whole, the sequence consists of Upper Jurassic neritic limestone passing upwards to siliceous limestone, radiolarite and shale (c. 10 m thick), thin-bedded turbidites (c. 20 m thick) and finally decametric levels of ophiolite-bearing cohesive debris flow deposits.

Three samples rich in radiolarian fauna (samples G03, G04, G05; Fig. 7) were collected on the northern flank of Tsikeri Mountain (about 2 km east of Pyra village; 38° 44′ 57.1″ N, 22° 17′ 54.2″ E; Fig. 2), in an overturned sequence characterized by a thin-bedded turbidite stratigraphically overlain by metre-thick cohesive debris flow deposits containing ophiolite material. The turbidite strata are geometrically overlain, through a tectonic contact, by siliceous limestone and bedded limestone of the Pelagonian Platform. In turn the mass flow sequence tectonically overlies the Parnassus Flysch with interposing slices of Boeotian Flysch (Fig. 5b). Unfortunately, the samples are always characterized by poor preservation of the radiolarian forms hampering dating. Nevertheless, the common presence in these samples of the nannofossil Watznaueria barnesiae and the lithostratigraphic similarities with other well-dated sections (see Baumgartner & Bernoulli, Reference Baumgartner and Bernoulli1976) allow us to tentatively attribute a Late Jurassic – Early Cretaceous age to these deposits.

Figure 7.

Schematic lithostratigraphic columns of the sampled units. Samples in bold are analysed in this work; samples in italic were analysed in Nirta et al. (Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). 1 – bedded limestone; 2 – massive limestone; 3 – cherty limestone; 4 – shale and fine-grained limestone; 5 – calcarenites; 6 – radiolarite and shale; 7 – shale; 8 – fine-grained sandstone and shale; 9 – coarse-grained sandstone; 10 – cohesive and granular debris flow deposit; 11 – ophiolite-bearing mass flow deposits and slide blocks; 12 – laterite. Dating of radiolarian chert from the Pelagonian Unit is expressed as the minimum age range.

In the Kallidromon Mountains foothills, 1.4 km west of Xyliki village, siliceous limestone and shaly-radiolarite strata top a well-bedded Pelagonian limestone succession. This is interrupted by the thrust of the above ophiolitic mélange and ophiolite nappe (Fig. 6 and cross-section A–A′ in Fig. 8). Samples were collected in the siliceous limestone (GR63, GR64, GR66) and in the shaly-radiolarite section (GR65, GR67, GR68) (Fig. 7).

Figure 8.

Regional geological cross-sections across the Boeotia region (see Fig. 6 for location and legend). The basal portion of the Lower Boeotian Flysch cropping out in the Domvrena–Makariotissa area is described in the circle: a – Triassic – Lower Jurassic neritic limestone; b – Middle–Upper Jurassic pelagic limestone and siliceous limestone; c – Upper Jurassic radiolarian cherts and shale; d – Upper Jurassic ophiolite-bearing mass flows, shales and marlstones (Lower Boeotian Flysch) (modified after B. Clément, unpub. Ph.D. thesis, Univ. Lille, 1983). D1, D2, D4, D4_bt – tectonic contacts with deformation phase attribution (see Section 5.c for details); F4_bt – D4_bt deformation phase megafold.

The following radiolarian assemblage was obtained from sample GR66: Archaeodictyomitra apiarium (Rüst), Archaeodictyomitra aff. apiarium (Rüst), Archaeodictyomitra cf. apiarium (Rüst), Archaeodictyomitra sp., Campanomitra cf. tuscanica (Chiari, Cortese & Marcucci), Cinguloturris (?) sp., Crococapsa sp., Loopus cf. blabla (Schaaf) sensu Dumitrica, Immenhauser & Dumitrica-Jud (Reference Dumitrica, Immenhauser and Dumitrica-Jud1997), Protunuma sp., Thanarla patricki gr. (Kocher), Transhsuum sp., Xitus cf. normalis (Wu & Li), Xitus (?) sp., Zhamoidellum ovum Dumitrica, Zhamoidellum cf. ovum Dumitrica, Zhamoidellum cf. ventricosum Dumitrica. This assemblage allow the attribution of a middle Callovian – early Oxfordian to late Kimmeridgian – early Tithonian age (UAZ. 8–11; UA zones after Baumgartner et al. Reference Baumgartner, Bartolini, Carter, Conti, Cortese, Danelian, De Wever, Dumitrica, Dumitrica-Jud, Goričan, Guex, Hull, Kito, Marcucci, Matsuoka, Murchey, O'Dogherty, Savary, Vishnevskaya, Widz, Yao, Baumgartner, O'Dogherty, Goričan, Urquhart, Pillevuit and Wever1995a ) based on the occurrence of Archaeodictyomitra apiarium (Rüst) with Zhamoidellum ovum Dumitrica (ranges after Baumgartner et al. Reference Baumgartner, O'Dogherty, Goričan, Dumitrica-Jud, Dumitrica, Pillevuit, Urquhart, Matsuoka, Danelian, Bartolini, Carter, De Wever, Kito, Marcucci, Steiger, Baumgartner, O'Dogherty, Goričan, Urquhart, Pillevuit and De Wever1995b ; Smuc & Goričan, Reference Smuc and Goričan2005) (Fig. 9).

Figure 9.

Scanning electron micrographs of radiolarians from sample GR66 (scale bar = 50 µm). (a) Archaeodictyomitra apiarium (Rüst); (b) Archaeodictyomitra aff. apiarium (Rüst); (c) Campanomitra cf. tuscanica (Chiari, Cortese & Marcucci); (d) Crococapsa sp., (e) Loopus cf. blabla (Schaaf) sensu Dumitrica, Immenhauser & Dumitrica-Jud (Reference Dumitrica, Immenhauser and Dumitrica-Jud1997); (f) Protunuma sp.; (g) Thanarla patricki gr. (Kocher); (h) Xitus cf. normalis (Wu & Li); (i) Zhamoidellum ovum Dumitrica; (j) Zhamoidellum cf. ventricosum Dumitrica.

4.b. Lower Boeotian Flysch

Biostratigraphic investigations in the Levadhia area allow the refinement of the ages of the lithostratigraphic Units A and B of the Lower Boeotian Flysch (LBFa and LBFb; see Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015 for details). The lowermost portion of unit LBFa, detached from its substratum, was sampled 4.5 km west of Levadhia (Fig. 10). According to Bornemann, Aschwer & Mutterlose (Reference Bornemann, Aschwer and Mutterlose2003), the nannofossil assemblages in samples GR69–70 (Fig. 7) allow the attribution of a middle–late Tithonian to early Berriasian age to the older sediments of the LBF on account of the dramatic increase of nannoconids (from 6 % to above 40 %) and the contemporary decrease of Conusphaera mexicana (from above 30 % to 1 %) and Polycostella beckmannii (from above 7 % to below 1 %) (see Table 1). The samples collected in unit LBFb in different outcrops located along the Levadhia–Arachova road (samples GR09, GR22–23, GR25, GR38–39, GR54; Fig. 7) allow us to constrain an age spanning from late Valanginian to the Hauterivian–Barremian boundary, based on the presence of Rucinolithus wisei, Calcicalathina oblongata, Cruciellipsis cuvillieri, common nannoconids and Lithraphidites bollii (zones CC3–CC5 p.p.; see Table 2).

Figure 10.

Geological map of the Levadhia area (location in Fig. 6).

4.c. The Mesoautochthonous Complex

This unit was sampled northwest of Levadhia, in overturned sections linked with imbricate thrusts exposed from the Distomon Fault to Cheronia village (Figs 10, 11).

Figure 11.

Geological cross-section along the Helicon range – Kopais plain transect (Levadhia area). On the right is reported the schematic structural profile of the Exarchos area. See Figure 10 for location and legend.

In the Cheronia area the stratigraphically lowermost levels of the Mesoautochthonous Complex are represented by platform material resedimented in a slope environment with a reddish bauxitic matrix (sample GR04; Fig. 7). The fossil assemblage consisting of Orbitolina cf. conica (D'Archiac), Ticinella cf. roberti Gandolfi and abundant Radiolitidae, Caprinidae, bryozoa and echinoderm fragments allows the attribution of a late Albian age to the penecontemporaneous bioclasts. This datum is in good accordance with the contemporaneous subaerial exposure and laterization of the Parnassus Platform and of wide areas of the more external Pelagonian Platform. The sampled section possibly represents the base of the First Mesoautochthonous sedimentary cycle (MES₁).

The carbonate levels sampled in another overturned section exposed further west, near the Distomon Fault (DSTF in Fig. 6) and just above the tectonic contact with the underlying Pelagonian Flysch (Fig. 12), are attributed a late Campanian – Maastrichtian age on the basis of the following microfauna assemblage: Siderolites calcitrapoides Lamarck, Cuneolina pavonia D'Orbigny (fragment), Lepidorbitoides sp., Orbitoides sp., Rugoglobigerina sp., Griphoporella? sp., fragments of Corallinaceae, Radiolitidae, stromatoporoids and echinoderms (sample GR46; Fig. 7). Upwards, the sedimentation shifts towards pelagic conditions with deposition of marls and marly limestones. A single sample (GR45) collected in the marly sequence gave the following nannofossil assemblage: Coccolithus pelagicus, Chiasmolithus sp., Chiasmolitus danicus, Bomolithus sp., Ericsonia subpertusa, Ericsonia robusta, Fasciculithus sp., Fasciculithus billii, Fasciculithus janii, Fasciculithus pileatus, Fasciculithus tympaniformis, Fasciculithus ulii, Heliolithus sp., Prinsius bisulcus, Sphenolithus sp., Sphenolithus anarrhopus, Sphenolithus moriformis, Thoracosphaera sp., Toweius pertusus, Toweius rotundus. This association allows the attribution of a late Paleocene age to this pelagic transition (Selandian–Thanetian Zone CNP8, corresponding to zones NP5 p.p.–NP6). As a whole, this sequence is possibly attributable to the upper part of the Second Mesoautochthonous sedimentary cycle (MES₂).

Figure 12.

Overturned section of the Mesoautochthonous Complex (Levadhia area; 38° 28′ 33″ N; 22° 44′ 10″ E).

4.d. Parnassus Flysch

Samples were collected in the lower part of the pre-flysch facies near the contact with the carbonate substratum at several sites on Mount Parnassus and in the northern Helicon range (GR25a–27a, GR31–33, GR35) (Figs 4, 6, 7). Nannofossil assemblages in the pre-flysch facies are attributable to the latest Paleocene – earliest Eocene. On Mount Parnassus, late Thanetian – early Ypresian zones CNP11 and CNE2, corresponding to zones NP9–NP10, were documented by the occurrence and absence of Fasciculithus spp., respectively (see Table 3). The early Eocene was recorded in the Helicon range, where late Ypresian Zone CNE5, corresponding to Zone NP13, was documented by the common occurrence of the first specimens belonging to Noelaerhabdaceae (Cyclicargolithus floridanus, Dictyococcites sp. and Reticulofenestra dictyoda; see Table 3). The transition to frank flysch sedimentation occurred during middle–late Ypresian zones CNE4 and CNE5 (corresponding to zones NP12 and NP13) characterized by the presence of common specimens referable to Toweius and Noelaerhabdaceae (see Table 4).

The obtained results are in agreement with the diachronous establishment of pelagic conditions in the Parnassus Zone after the Late Cretaceous tectonic events that levelled the seafloor bathymetry before flysch deposition (Gregou, Reference Gregou1996).

4.e. Pelagonian Flysch

The Pelagonian Flysch was sampled in several sites in the hills located between Levadhia and Cheronia villages and nearby Davlia village (Fig. 6). In this area the Pelagonian Flysch is both exposed in its primary position at the top of the limestone of the Mesoautochthonous Complex or tectonically below it. In general, the Mesoautochthonous Complex and Pelagonian Flysch are overthrust southwards, through out-of-sequence thrusts, above the Boeotian and Parnassus units.

Nannofossil assemblages from the reddish marls at the base of the flysch sequence (pre-flysch facies; samples GR06–07, GR42–44, GR48, GR50–51a/b; Fig. 7) indicate a late Paleocene – early Eocene age (see Table 5). The Selandian/Thanetian Zone CNP8 (corresponding to zones NP5 p.p.–NP6) was documented by the occurrence of Heliolithus kleinpellii. The Thanetian Zone CNP9 (corresponding to Zone NP7 p.p.) was established on the presence of Discoaster cf. mohleri and the absence of H. kleinpellii, while the late Thanetian Zone CNP11 (corresponding to Zone NP9) was recognized by the presence of Fasciculithus lillianae. The early Ypresian Zone CNE3 (corresponding to zones NP10 p.p.–NP11) was identified by the presence of Girgisia gammation (Toweius? gammation) and Sphenolithus radians. Biostratigraphic results point to an early Eocene age for the shaly-arenaceous and conglomeratic flysch facies (samples GR08, GR47, GR49); Ypresian zones CNE4 and CNE5 (corresponding to zones NP12 and NP13) were recognized, respectively, by the occurrence of Tribrachiatus orthostylus and the common presence of specimens belonging to Noelaerhabdaceae (see Table 6).

5.a. Data collection and methods

The structural analysis aimed to discriminate structures related to the different stages of the tectonic evolution of the area. The relative timing of deformation was constrained by cross-cutting relationships observed in all rock units, both at the map and outcrop scale. Mesoscopic structural observations were collected at several sites within the study area, focusing on the major tectonic contacts. Fault-slip data were analysed in order to derive the palaeostress field associated with the stages of regional deformation. Palaeostress analysis was performed using the Fault software developed by Caputo & Caputo (Reference Caputo and Caputo1988) that calculates the orientation of the stress tensors σ₁, σ₂, and σ₃ (where σ₁ is the maximum compressional stress and σ₁>σ₂>σ₃) through the dihedron angle inversion method of Angelier & Mechler (Reference Angelier and Mechler1977). Basic limitations and assumptions of the palaeostress analysis were taken into account. Table 7 presents the parameters of the reduced stress tensors as well as the estimated shortening directions derived from the structural analysis of fold axes. In the next sections, the tectonic implications of these parameters will be interpreted and the results will be integrated into the regional geodynamic evolution of the study area.

Table 7.

Mesoscopic structural field data from sites of measurement of the study area

The σ₁, σ₂, and σ₃ axes have been computed applying the dihedron angle method to fault populations (Angelier & Mechler, Reference Angelier and Mechler1977) through the ‘FAULT’ software of Caputo & Caputo (Reference Caputo and Caputo1988). Fold axes, grouped according to both tectonic event and rock unit, allowed a rough estimate of the maximum horizontal palaeostress axis. Abbreviations: OPH – Ophiolitic Unit: mélange; PLG – Pelagonian Unit limestone; PLF – Pelagonian Flysch; LBF – Lower Boeotian Flysch; UBF – Upper Boeotian Flysch; PAR – Parnassus Unit limestone; PAF: Parnassus Flysch; BEO – Boeotian Unit limestone; MES – Mesoautochthonous Complex limestone.

The results of the field-based structural work are synthesized in a new geological–structural map and cross-section of the Levadhia region (Figs 10, 11) and a tectonic sketch map and two cross-sections across the whole studied area (Figs 6, 8). Some of the collected orientation data are summarized in equal-area plots.

5.b. Tectonic setting and major tectonic structures

The studied area is characterized by a complex structural evolution, deriving from the superposition of several tectonic phases, related to obduction, continental collision, subducting slab retreat and gravitational collapse of over-thickened crust (e.g. Bernoulli & Laubscher, Reference Bernoulli and Laubscher1972; Robertson & Dixon, Reference Robertson, Dixon, Dixon and Robertson1984; Jacobshagen, Reference Jacobshagen1986; Mercier et al. Reference Mercier, Sorel, Vergély and Simeakis1989; Le Pichon et al. Reference Le Pichon, Chamot-Rooke, Lalleman, Noomen and Veis1995; Ricou et al. Reference Ricou, Burg, Godfriaux and Ivanov1998; Brun & Faccenna, Reference Brun and Faccenna2008).

Tectono-metamorphic events accompanied this long-lasting evolution (Fig. 4). Middle Jurassic HT–LP metamorphism is recorded at the base of the obducting nappe (Roddick, Cameron & Smith, Reference Roddick, Cameron and Smith1979; Spray & Roddick, Reference Spray and Roddick1980; Dimo-Lahitte, Monié & Vergély, Reference Dimo-Lahitte, Monié and Vergély2001) then followed by three major tectono-metamorphic events in the eastern Pelagonian Zone: (i) Early Cretaceous upper-greenschist to blueschist facies metamorphism and anatexis events (Barton, Reference Barton1976; Yarwood & Dixon, Reference Yarwood and Dixon1977; Schermer, Lux & Burchfiel, Reference Schermer, Lux and Burchfiel1990; Lips, White & Wijbrans, Reference Lips, White and Wijbrans1998; Schenker et al. Reference Schenker, Burg, Kostopoulos, Moulas, Larionov and Von Quadt2014); (ii) Late Cretaceous blueschist facies metamorphism (Lips, White & Wijbrans, Reference Lips, White and Wijbrans1998; Lips, Wijbrans & White, Reference Lips, Wijbrans, White, Durand, Jolivet, Horváth and Séranne1999); (iii) Paleocene–Eocene blueschist to greenschist facies metamorphism (Schermer, Lux & Burchfiel, Reference Schermer, Lux and Burchfiel1990; Lips, White & Wijbrans, Reference Lips, White and Wijbrans1998; Lips, Wijbrans & White, Reference Lips, Wijbrans, White, Durand, Jolivet, Horváth and Séranne1999). The latter tectono-metamorphic event is commonly related to the final westward emplacement of the Internal Hellenides upon the External Hellenides and the development of the Tertiary fold and thrust belt and high-pressure belt in the Olympos–Ossa, Pelion and Cyclades areas. Finally, since Eocene–Oligocene times the prolonged post-orogenic extensional tectonics drove exhumation of the previously stacked crustal and oceanic nappes in metamorphic domes (Lister, Banga & Feenstra, Reference Lister, Banga and Feenstra1984; Jolivet et al. Reference Jolivet, Brun, Gautier, Lallemant and Patriat1994; Jolivet & Brun, Reference Jolivet and Brun2010; Burg, Reference Burg, Skourtsos and Lister2012).

The generally NNW-striking structural fabric of the Hellenic Belt is deflected to an ESE-trend confined to the inner (eastern) border of the Parnassus Unit. In the studied area, the Jurassic thrusting of the ophiolite nappe onto the Pelagonian Platform, partly remobilized by Late Cretaceous and Tertiary tectonics, is preserved only in the Kallidromon Belt, north of the Parnassus Unit (Fig. 6 and profile A–A′ in Fig. 8).

North of the Parnassus Unit, mainly top-to-the-S thrusts are developed forming imbricate piles made of ophiolites, the Pelagonian Unit, Mesoautochthonous sequences, and Pelagonian and Boeotian flysches. South of the Kopais Plain the Parnassus Unit and carbonate basement of the Boeotian Unit are locally overthrust to the NW above the Eocene Pelagonian Flysch (Fig. 6 and cross-section B–B′–B″ in Fig. 8). The tectonic framework is particularly complicated in the Levadhia area where a new geological map and cross-section were produced (Figs 10, 11). The zone is characterized by a series of mainly S-verging thrust sheets variously superposing the Boeotian Flysch, the Mesoautochthonous Complex and the Tertiary Pelagonian and Parnassus flysches and, further southwards, by the triangular zone between S- and N-verging thrusts near the northeastern border of the Parnassus Platform (Fig. 11). Strike-slip and normal faults dislocate the previously described structures. Main dextral strike-slip faults are represented by the NNE-trending Distomon (Fig. 13b, c, d) and Aliki faults (DSTF and ALKF in Fig. 6), whereas a main sinistral strike-slip fault is represented by the NNW-trending Agia Triada Fault (ATRF in Fig. 6).

Figure 13.

(a) UBF unconformably sedimented above the deformed LBF (Mount Iti area – 38° 47′ 5.4″ N; 22° 22′ 53.3″ E). (b) Panoramic view showing the overthrusting of the Mesoautochthonous Complex above the Pelagonian Flysch. In the background is evidenced the trace of the Distomon Fault (DSTF). (c) View towards the south within the core of the DSTF (38° 29′ 33.4″ N; 22° 41′ 47.9″ E). (d) Dextral strike-slip striae along the DSTF fault plane (38° 29′ 33.4″ N; 22° 41′ 47.9″ E). (e) Fault gouge between limestone of the Pelagonian Unit and Upper Jurassic ophiolite-bearing turbidites at the contact with the underlying Parnassus Flysch (Mount Iti area – 38° 43′ 53.6″ N; 22° 20′ 27.1″ E). (f) D4 tectonic contact between the LBFa (hangingwall) and UBFb (footwall) along the Levadhia–Arachova road-cut (38° 26′ 26.4″ N; 22° 48′ 44.5″ E).

5.c. Mesoscopic structural analysis

Structural and palaeostress studies concentrated on all units cropping out in the study area and on key outcrops of the Pelagonian Unit and Boeotian Flysch in the Mount Iti area. Mesoscopic and macroscopic superposed folding phases, some of which linked to thrusting, affected the Boeotian Flysch and Tertiary flysches, while the carbonate-dominated units only registered large-scale folding and brittle structures. The structural analysis combined with the stratigraphic records enabled us to distinguish a succession of six major deformational events, hereafter referred to as D1, D2, D3, D4, D4_bt and D5, directly related to the evolution of the study area. Brittle deformation related to Triassic rifting processes and older deformations are beyond the scope of this paper.

5.c.1. D1 phase deformations

The oldest deformations recognized in the area are related to the Middle–Late Jurassic obduction of the Maliac–Vardar Ocean lithosphere onto the eastern Adria continental margin (i.e. the Pelagonian Zone). Mesoscopic deformation related to this phase was recognized in the Pelagonian and Boeotian units. Deformations are differentiated in time and space according to relative position with respect to the advancing ophiolite nappe shifting from a normal stress regime (D1a sub-stage) to a compressive stress regime (D1b sub-stage). This subdivision is clearly evident in the Pelagonian Unit while the Boeotian Unit, in this phase, registered only deformations of the D1a sub-stage.

The D1a sub-stage is characterized by widespread syn-sedimentary deformations relative to different and subsequent episodes that affected the Upper Jurassic clastic deposits sedimented on top of the Pelagonian Platform and the LBF deposited onto the Boeotian carbonate sequence. Gravity-induced deformations such as slumping events, the formation of local wedge-shaped strata and repeated angular unconformities, as well as syn-sedimentary faults, are common in these deposits. Faults (Fig. 14, sites 10N and 750N) are characterized by cumulated episodes of faulting during sedimentation with variable throw ranging from some centimetres to half a metre (Fig. 15a). Syn-sedimentary faulting is particularly evident in the Upper Jurassic clastic sediments of the Pelagonian Unit up to the upper Tithonian–Hauterivian portions of the LBF (i.e. the LBFa and LBFb lithostratigraphic units). Stress inversion of the scarce fault-slip data collected in the LBF in the Levadhia area and in the ophiolite-bearing thin-bedded turbidites on top of the Pelagonian carbonates in the Mount Iti area results in an ENE–WSW direction of extension (Table 7; Fig. 14, sites 10N and 750N).

Figure 14.

Distribution of mesostructural fault-slip data, fold axes and S-C structures relative to D1, D2 and D3 deformation phases. White arrows indicate direction of extension (σ₃) and black arrows indicate direction of compression (σ₁) calculated through inversion of fault-slip data. Black arrows of diagrams relative to folds and S-C structures are also indicative of the sense of tectonic movement. π-diagram on D1b fold axes in the OPH possibly indicates axis of successive deformation phases. All diagrams are equal-area lower hemisphere projections; density contours at 1% intervals. UBF – Upper Boeotian Flysch; LBF – Lower Boeotian Flysch; PLG – Pelagonian Unit; OPH – Ophiolitic Unit, mélange.

Figure 15.

(a) Overturned syn-sedimentary fault showing cumulated episodes of faulting during sedimentation (ophiolite-bearing thin-bedded turbidites on top of the Pelagonian limestone; Iti area – 38° 44′ 56″ N; 22° 17′ 53″ E); the inset diagram shows cumulated offset on the fault. (b) S-C structures in the mélange sandwiched between the Pelagonian Unit and ophiolite nappe in the Kallidromon area (38° 41′ 08″ N; 22° 37′ 44″ E). (c) Overthrusting (D1 phase of deformation) of the Ophiolitic Unit above and the Pelagonian Unit. The tectonic contact was successively tilted by a large-scale anticline related to the D4 deformation phase to simulate a normal fault (Exarchos area, see Figs 8 and 11; 38° 35′ 53″ N; 22° 55′ 35″ E. (d) S3 disjunctive cleavage superimposed on the limb of F2 fold in the LBF_b (Levadhia–Arachova road-cut – 38° 27′ 37″ N; 22° 45′ 46″ E). (e) S3 cleavage folded by an open F4 fold in LBF_b (Levadhia–Arachova road-cut – 38° 27′ 37″ N; 22° 45′ 36″ E). (f) Interferences of F4_bt open folds onto F4 tight fold near the back-thrust of the Parnassus units above the Parnassus Flysch (Levadhia area – 38° 25′ 57″ N; 22° 47′ 252″ E).

Since Barremian time, faulting in the LBF gave way to gravity-triggered deformations (slumps, disrupted layers, angular unconformities, erosional surfaces). The lack of any observable water-escape structures suggests that these deformations were induced by large-scale gravitational instability in the sedimentary basin rather than small-scale liquefaction of sediments due to increased pore-water pressure.

The most striking deformation related to the D1b sub-stage is the overthrusting of the ophiolite nappe with an underlying amphibolite sole and tectono-sedimentary mélange onto the Pelagonian Platform. In the studied area, the amphibolite sole always occurs as dismembered blocks embedded in the mélange matrix. Contacts of the Ophiolitic Unit onto the Pelagonian Unit were observed on the southern flank of the Kallidromon Mountains (Fig. 6). The footwall is characterized by highly sheared mélange showing pervasive scaly cleavage. A planar fabric is mainly represented by clay minerals enveloping lithic domains consisting of basalt, chert, serpentinite and sandstone. Well-developed kinematic indicators (S-C fabrics, asymmetric folds, asymmetric boudins) record a consistent top-to-the-SW sense of shear (Fig. 14, site 891I; Fig. 15b).

D1a syn-sedimentary faults were probably related to the lithospheric bending triggered in the Pelagonian Unit, first, and Boeotian Unit, later, by the obducting ophiolite nappe (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). Finally, D1b deformations marked the tectonic emplacement of the ophiolites and mélange onto the Pelagonian Platform (Fig. 15c) and the increasing gravity-triggered deformations in the Barremian–Aptian LBF (LBFc; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015).

5.c.2. D2 phase deformations

The earliest folding event (F2) developed in the LBF is characterized by pervasive close to sub-isoclinal mesoscopic non-cylindrical folds with approximately similar geometry (class 2 of Ramsay, Reference Ramsay1967). Sandstone and limestone strata in the F2 fold limbs are often affected by boudinage with both symmetric and asymmetric geometries. Veins associated with the D2 deformational phase are oriented at a high angle to the bedding and filled with sparitic calcite. Despite the high strain ratio, which characterizes the F2 folds and the associated small-scale thrusts, these structures are rarely associated with cleavage. Fold axes and axial planes are strongly reoriented by successive folding and faulting, depicting a very dispersed pattern in the stereonet projections (D2 folds in Fig. 14). A tentative subtraction of the successive folding event (F3) seems to point to an original N–S-trending axes, whereas analysis of bedding-cleavage relationships, fold asymmetry and direction of younging point to a western vergence (Fig. 16).

Figure 16.

F2 + F3 + F4 interference patterns in the LBF. Fold assemblage is dissected by D4 (reverse) and D5 (normal) faults (38° 27′ 37″ N; 22° 45′ 46″ E). Black arrows show younging direction.

D2 is possibly associated with renewed westward translation of the Ophiolitic Unit onto the Pelagonian Unit. This thrusting event in the Kallidromon zone cross-cuts the scaly cleavage acquired during the D1 phase of deformation and is associated with small-scale mesoscopic folds deforming the S1 foliation. Fault-slip data (Table 7, site 885I in Fig. 14) result in a NE–SW shortening direction, which is consistent with the major WNW–ESE-trending and N-dipping thrusts leading the ophiolites onto the Pelagonian Unit.

5.c.3. D3 phase deformations

D3 consists of a folding event (F3) well developed and exposed in the LBF and UBF. It is characterized by gentle to close folds with generally parallel geometry (class 1B of Ramsay, Reference Ramsay1967) in the LBF and close to tight folds with mostly similar geometry in the shaly-silty strata packages of the UBF. Axial plane foliation S3 is marked by disjunctive cleavage that is well developed only in the shaly layers of the turbiditic sequences (Fig. 15d, e). Isolated F3-fold hinge zones are common in the shale-dominated lithologies of the UBF.

F3 fold axes are scattered, as a result of the post-D3 deformations (Fig. 17). However, an original E–W trend of the F3 axes and the main N-dipping direction of the axial planes indicate a southward facing direction (F3 folds in the LBF and UBF of Fig. 14). Mostly N-dipping mesoscopic thrusts associated with folds were observed in the overturned limbs of tight folds (Fig. 17). Superimposed F2 + F3 interference results mostly in a type-3 pattern (Ramsay, Reference Ramsay1967).

Figure 17.

F3 + F4 interference pattern in the LBF. The overturned limb of the main F3 fold and the associated thrust plane are re-folded by F4, and rotated to simulate a low-angle normal fault (38° 27′ 29″ N; 22° 53′ 26″ E).

Widespread E- to NE-dipping thrusts cross-cut D1 and D2 planar structures in the Ophiolitic Unit and Pelagonian Unit. However, it was not possible to clearly distinguish in the field D3 structures from brittle events associated with younger compressive phases (D4–D4_bt).

5.c.4. D4 phase deformations

Mesoscopic and macroscopic structures related to this deformation phase are the most frequent and well developed in all the tectonic units.

F4 mesoscopic folds are mostly developed in the Pelagonian and Parnassus flysches and in both the LBF and UBF. Deformations in the Boeotian Flysch (LBF + UBF) are represented by gentle to tight isopach F4-folds associated with a disjunctive axial plane foliation S4 (Figs 15e, 16). In the Tertiary Pelagonian and Parnassus flysches, F4 folds are represented by close to tight structures with mostly similar geometry and with a weak continuous axial plane cleavage particularly evident only in the hinge zones. Axial plane attitude is generally upright to steeply inclined in the Boeotian Flysch and gently inclined to recumbent in the Tertiary flysches (Fig. 15f). Brittle shears commonly tear F4 hinge zones both in the Boeotian Flysch and in the Tertiary flysches.

Fold axes are gathered around a ENE–WSW trend in both the Tertiary flysches and Boeotian Flysch, plunging preferentially towards eastern sectors (F4 folds in the LBF, PAF and UBF of Fig. 18). Dispersion of F4 axes in the Tertiary flysches and the UBF appears to be caused by successive tectonic deformations, whereas most of the dispersion observed in the LBF seems to be also controlled, and locally forced, by the pre-existing structural fabric. Pole to axial plane processing on a π-diagram of F4 structures of the Pelagonian and Parnassus flysches points to an average axis trending N100. F4 axial planes steeply dip and gently dip towards northern sectors in the Boeotian Flysch and Tertiary flysches, respectively. The folds’ asymmetry and its relationships with bedding polarity, along with slip indicators of fold-related brittle shears clearly point to a SSW vergence direction.

Figure 18.

Distribution of mesostructural fault-slip data and fold axes relative to D4 and D4_bt deformation phases. White arrows indicate direction of extension (σ₃) and black arrows indicate direction of compression (σ₁) calculated through inversion of fault-slip data. Black arrows of diagrams relative to folds are also indicative of the sense of vergence. All diagrams are equal-area lower hemisphere projections; density contours at 1 % intervals. UBF – Upper Boeotian Flysch; LBF – Lower Boeotian Flysch; PLG – Pelagonian Unit, limestones; PLF – Pelagonian Flysch; PAR – Parnassus Unit, limestones; PAF – Parnassus Flysch; OPH – Ophiolitic Unit, mélange.

Asymmetric open megascopic F4 folds approximately E–W-trending and verging towards the south are well developed in the Mesoautochthonous carbonate sequences. These megastructures are bounded by thrust surfaces commonly isolating overturned fold limbs (Figs 11, 12, 13b). Mesoscopic fold interference patterns F2 + F4 and F3 + F4 observed in the LBF and UBF are mostly of type 3 (Ramsay, Reference Ramsay1967) (Figs 16, 17). Along the Levadhia–Amphissa road – 5 km east of Levadhia – D4 brittle reverse faults and open folds overprint F3 close folds, tilting and locally reactivating the D1 syn-sedimentary normal fault systems of the LBF (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) near a D4-thrust juxtaposing the LBFa onto the UBF (Fig. 13f).

Differences in the fold geometry among the Boeotian Flysch and Tertiary flysches are clearly linked to differentiation in structural levels, shallower in the Boeotian Flysch and deeper in the Tertiary flysches. In fact, the F4 folding event is associated with the westward overthrusting of the more internal tectonic units (Pelagonian, Boeotian and Ophiolitic units) above the Pelagonian Flysch, and the Parnassus Flysch in Tertiary times.

The overthrusting of the Pelagonian Unit and Ophiolitic Unit above the Boeotian and Parnassus units represents the major tectonic contact associated with this compressive phase. The superposition of the ophiolites, Pelagonian carbonates and slices of Boeotian Flysch above the Parnassus Flysch on Mount Iti and in the central Parnassus area (Amfiklia–Polydrosos area, Fig. 6) represent klippen related to this tectonic phase (Fig. 3, cross-section A–A′ in Fig. 8, Fig. 13e).

In the Kallidromon area and north of Exarchos village on the northern border of the Kopais plain, the Jurassic tectonic contact of the ophiolite nappe above the Pelagonian carbonates (D1 tectonic phase) is folded and cut by S-facing D4 structures (Figs 6, 8).

The Mesoautochthonous carbonate sequence is mostly detached from its substratum represented by the Pelagonian and Ophiolitic units, and overthrusts the Tertiary Pelagonian Flysch (Figs 12, 13b). During this tectonic phase, the progressive deformation enucleated N-dipping out-of-sequence thrusts in the Pelagonian Flysch and Boeotian Flysch mostly controlled by the structural buttress represented by the Parnassus Platform. The Triassic–Cretaceous Parnassus carbonates and Tertiary flysch are also deformed by N-dipping thrust structures.

Thrust structures are generally N-dipping, with angles varying from 10° to 50°; steeply dipping back-thrusts are less common. Major thrust structures are commonly accommodated by N–S-trending strike-slip faults, generally left-lateral.

The palaeostress analysis carried out on reverse and strike-slip faults related to the D4 phase indicates a NE–SW-trending shortening direction (Table 7; Fig. 18).

5.c.6. D5 phase deformations

The D5 deformational event consists of both well-developed regional extensional faults and transtensive (D5_T) faults that cross-cut all the previous structures. Inversion of fault-slip data points to NE–SW and NW–SE extension directions, accommodated by NNE-trending dextral strike-slip faults (Table 7; Fig. 19). They in part represent reactivated ancient normal structures possibly developed during the early rifting stage preceding obduction, successively affected by both compressional and extensional stress regimes (Table 7, sites 852D, 854SN, 932D and 963D in Fig. 19). The Distomon (Fig. 13c, d) and Aliki faults (DSTF and ALKF in Fig. 6) represent the main transverse structures affecting the studied area.

Figure 19.

Schematic diagram illustrating areal distribution of mesostructural fault slip data and fold axes relative to D5 and D5_T deformation phases. White arrows indicate direction of extension (σ₃) and black arrows indicate direction of compression (σ₁) calculated through inversion of fault slip data. All diagrams are equal-area lower hemisphere projections. UBF – Upper Boeotian Flysch; LBF – Lower Boeotian Flysch; PLG – Pelagonian Unit, limestones; PAR – Parnassus Unit, limestones; PAF – Parnassus Flysch.

5.c.7. Timing of deformation

The timing of deformation in the studied area is constrained by the age of syn-orogenic sedimentary units and by the Aptian regional unconformity underlying the Mesoautochthonous Complex and the UBF.

In the Pelagonian Unit the D1a sub-stage is dated as latest Jurassic as it is coeval with sedimentation of the ophiolite-bearing deposits topping the carbonate and pelagic succession (Figs 15a, 20). Syn-sedimentary extension-related deformations in the LBF were clearly observed only in the LBFa and in the lower part of the LBFb subunits allowing us to tentatively limit the age of the D1a sub-stage to the Tithonian–Valanginian time span.

Figure 20.

Inferred ages of the Late Jurassic to Eocene deformation phases (D1 to D4) in Central Greece.

The D1b sub-stage deformations often follow but are partly coeval with the D1a deformations recorded in the more external portion of the Pelagonian Unit. The D1b deformations affected the Pelagonian Unit after sedimentation of radiolarite and ophiolite-bearing deposits and do not affect the LBF, allowing the attribution of a Kimmeridgian–?Valanginian age to the deformations. Valanginian is also the youngest age of the deformed platform carbonate units partly preserved above the ophiolite nappe (namely the Munella Platform above Mirdita ophiolites in Albania Fig. 1; Schlagintweit et al. Reference Schlagintweit, Gawlick, Missoni, Lein, Hosha, Sudar, Ercegovac and Grubic2006) and successively sealed by the mildly deformed Mesoautochthonous Complex. This suggests that obduction-related deformations lasted at least until Valanginian time.

The UBF and First Mesoautochthonous Cycle sealed the D2 deformations acquired by the LBF and Pelagonian Unit, respectively. In the Koumaritsi area a little-deformed succession of the UBF (oldest age: Albian; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) unconformably covers the LBFc subunit (youngest age: lower Aptian; Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) that shows evidence of faulting and F2-related boudinage (Fig. 13a). The oldest age of the First Mesoautochthonous Cycle lying onto deformed ophiolites and the Pelagonian Unit is late Aptian in the more internal sector (Vermion–Vourinos area, Fig. 1; Photiades et al. Reference Photiades, Carras, Bortolotti, Fazzuoli and Principi2007) and Albian in the studied area. Considering the above data, it is possible to attribute an early to middle Aptian age to the D2 deformations.

The D3 deformations were clearly detected only in the LBF and UBF whereas it was not possible to ascertain their existence in the Mesoautochthonous Complex. On the other hand, they do not affect the Pelagonian and Parnassus flysches. As a consequence, the age and significance of this deformation phase remains so far ill-defined. Tentatively, it is possible to attribute a Santonian (youngest age of the UBF) to Paleocene age (oldest age of the Parnassus and Pelagonian flysches) to the D3 phase.

The D4 and localized D4_bt deformations took place after the Pelagonian and Parnassus Flysch sedimentation (early Eocene) and were later sealed by the oldest post-orogenic sediments known in the area (Serravallian lacustrine deposits of the Thiva Basin; Mettos, Ioakim & Rondoyanni, Reference Mettos, Ioakim and Rondoyanni2000).

6.a. From rifting to the passive margin stage

The rifting stage, accompanied by transitional to alkaline volcanism, occurred during early–Middle Triassic time and was widely documented in the Pelagonian succession (Baud et al. Reference Baud, Jenny, Papanikolaou, Sideris and Stampfli1991; Capedri et al. Reference Capedri, Toscani, Grandi, Venturelli, Papanikolaou and Skarpelis1997; A. De Bono, unpub. Ph.D. thesis, Univ. Lausanne, 1998). Furthermore, events of alkaline to transitional rift-related volcanism are also recorded in the Pindos stratigraphy (Pe-Piper, Reference Pe-Piper1982; Pe-Piper & Piper, Reference Pe-Piper and Piper1991).

In the Othrys Massif and in the Argolis Peninsula (Fig. 1) the transition from the Pelagonian Platform to the more distal eastern marginal series (West Maliac Margin Units) and up to the oceanic crust is well documented. The onset of oceanization in the Maliac–Vardar basin is testified to by the oldest radiolarites (Anisian) deposited on top of MOR basalts (Bortolotti et al. Reference Bortolotti, Chiari, Kodra, Marcucci, Marroni, Mustafa, Pandolfi, Prela, Principi and Saccani2006; Bortolotti et al. Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013; Ferrière et al. Reference Ferrière, Chanier, Baumgartner, Dumitrica, Caridroit, Bout-Roumazeilles, Graveleau, Danelian and Ventalon2015). By contrast, since Early Jurassic time the Pindos and Ionian geotectonic zones are interpreted as pelagic intracontinental basins separated by an intervening carbonate platform represented by the Gavrovo Unit (Fig. 1); the scattered outcrops testifying to the existence of a rifting stage in the Pindos Basin are then not corroborated by evidence of oceanization.

During the passive margin stage, east of the Pindos Basin, only the Pelagonian Zone could be clearly identified as an isopic zone while the Boeotian and Parnassus zones are not yet differentiated and shared with the Pelagonian Zone a similar stratigraphy dominated by neritic carbonates of from Late Triassic to early Middle Jurassic age (Fig. 4).

In the proposed model, the Late Triassic geotectonic setting of the transition from the Adria continental margin to the Maliac–Vardar Ocean can be summarized as in Figure 21a.

Figure 21.

Proposed geodynamic model and tectonic evolution for the obduction and continental collision in the Hellenic belt. See text for discussion.

6.b. Obduction – D1 tectonic phase

Intraoceanic subduction in the Maliac–Vardar Ocean is testified to by the occurrence of ophiolitic sequences developed in a supra-subduction environment (Beccaluva et al. Reference Beccaluva, Coltorti, Saccani and Siena2005; Saccani et al. Reference Saccani, Photiades, Santato and Zeda2008b ; Bortolotti et al. Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013) and by a subduction-related continental magmatic arc and back-arc basin towards the European margin (i.e. the Paikon arc and Guevgueli back-arc basin; Mercier, Reference Mercier1966; Ferrière & Stais, Reference Ferrière and Stais1995; Brown & Robertson, Reference Brown and Robertson2003, Reference Brown and Robertson2004; Saccani et al. Reference Saccani, Bortolotti, Marroni, Pandolfi, Photiades and Principi2008a ; Bortolotti et al. Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013) developed since early Middle Jurassic time (Fig. 22a). These magmatic events imply the existence of an E-dipping subducting plane since Early Jurassic time (Bortolotti et al. Reference Bortolotti, Marroni, Pandolfi and Principi2005).

Figure 22.

Schematic reconstruction of the evolution of subduction-related magmatism accompanying the closure of the Maliac–Vardar Ocean from Middle Jurassic time until Late Cretaceous post-collisional collapse. The Middle–Late Jurassic step is redrawn after figure 10 in Bortolotti et al. (Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013). Not to scale.

During Middle Jurassic time, as a consequence of the continuing convergence between Adria and Europe, tectonically decoupled oceanic lithosphere started thrusting westwards (Bébien et al. Reference Bébien, Ohnenstetter, Ohnenstetter and Vergély1980; Spray & Roddick, Reference Spray and Roddick1980). This event led to metamorphism at the sole of the overriding slab and mélange formation after off-scraping of the facing oceanic sediments and erosion of the advancing slab. At the Middle–Late Jurassic transition, the overriding slab approached the eastern margin of the Adria continental margin (i.e. the Pelagonian Zone). In late Middle Jurassic time, the advancing ophiolite nappe caused a rapid drowning of the distal continental margin (i.e. West Maliac Margin Units) that shifted from carbonate deposition to deep-sea siliceous deposition and then to ophiolite-bearing mass flow deposits. At the same time, to the west the continental margin experienced emersion episodes in the western Pelagonian Zone and in the Parnassus Zone (b1 bauxite horizon) (Figs 4, 21b). Emersion is correlated with the lithospheric forebulge induced by the load of the ophiolite nappe to the east (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015). The bending was accommodated by drowning of a large portion of the former Pelagonian Platform and development of the Boeotian Basin (Fig. 21b). Differentiation of the Parnassus and Pelagonian zones with interposition of the Boeotian Basin was a diachronous event. In fact, while some platform portions rapidly drowned to form the future Boeotian Basin (Aliki and Makariotissa sections; Danelian & Robertson, Reference Danelian and Robertson1998), some other portions necessarily represent a continuous shallow-water environment to allow the migration of the allochthonous b1 bauxites from the ophiolites to the east throughout the western Pelagonian Platform and up to the future Parnassus Zone. An obduction-related extensional regime is registered by the syn-sedimentary faulting in the mass flow deposits on top of the Jurassic carbonates of the Pelagonian Unit and in the LBF throughout latest Jurassic time (D1a sub-stage; Fig. 21b). Final overthrusting of the ophiolites and associated mélange progressively shifted into a compressional tectonic regime in the Pelagonian Platform (D1b sub-stage; Fig. 21b). The time–space distribution of the ophiolite detritus supplied by the obducting units in the facing continental margin and top-to-the-SW kinematic indicators within the tectonic foliation S1 clearly point to an E-towards-W translation of the obducted units. This sedimentation versus obduction trend is clearly recognizable throughout Late Jurassic time and culminated with the beginning of sedimentation of the LBF in the Boeotian Basin. At this time, a new emersion episode was recognized in the Parnassus Platform (b2 horizon), which could possibly reflect a tectonically forced regression, induced during the climaxing phase of obduction (Fig. 21c). Obduction could be traced up to Valanginian time, as suggested by the age of the youngest ophiolite-derived clastic deposits on the top of the Pelagonian Platform and the youngest syn-obduction carbonate platform set on the top of the ophiolite nappe (i.e. the Munella Platform; Schlagintweit et al. Reference Schlagintweit, Gawlick, Missoni, Lein, Hosha, Sudar, Ercegovac and Grubic2006).

The youngest evidence of oceanic crust in the Maliac–Vardar Ocean is of latest Jurassic age (Bortolotti et al. Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013 and quoted references). Similarly, back-arc magmatism between the Paikon arc and the Eurasian margin (Guevgueli back-arc basin), induced by the E-directed subducting slab of the Maliac–Vardar Ocean (Saccani et al. Reference Saccani, Bortolotti, Marroni, Pandolfi, Photiades and Principi2008a) since Middle Jurassic time, continued until Late Jurassic time (Danelian, Robertson & Dimitriadis, Reference Danelian, Robertson and Dimitriadis1996; Kukoč et al. Reference Kukoč, Goričan, Kosir, Belak, Halamic and Hrvatović2012) and was also associated with the coeval emplacement of calc-alkaline granitoids both in the back-arc oceanic crust and in the nearby Eurasian continental margin (e.g. Fanos granite; Anders et al. Reference Anders, Reischmann, Poller and Kostopoulos2005; Šarić et al. Reference Šarić, Cvetkoviç, Romer, Christofides and Koroneos2009). Moreover, arc-related calc-alkaline granitoids, now found as orthogneisses in the Rhodope Massif, are reported to intrude the Eurasian margin further east until Late Jurassic – Early Cretaceous times (Turpaud & Reischmann, Reference Turpaud and Reischmann2010) (‘Rhodopian intrusions’ in Fig. 22b). Therefore, we assume that E-dipping subduction of the Maliac–Vardar Ocean continued throughout Early Cretaceous time. However, a high convergence rate between Adria and Europe during this period (Rosenbaum, Lister & Duboz, Reference Rosenbaum, Lister and Duboz2002; Bertrand, Guillou-Frottier & Loiselet, Reference Bertrand, Guillou-Frottier and Loiselet2014) could have induced upper plate compression that prevented upward migration of magmas necessary to trigger widespread volcanic activity and, at the same time, could have facilitated sequestration of melts at deep to moderate crustal depths.

6.c. Continental collision: D2 tectonic phase

Ophiolite-derived mass flow gliding events progressively increased in the upper portion of the LBF of Barremian–Aptian age (LBFc). The sedimentation in the Boeotian trough was stopped by an intense tectonic event that produced the most pervasive deformations in the LBF. The beginning of sedimentation of the UBF of Albian age sealed D2 tectonic deformations in the underlying LBF (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) allowing the attribution of an Aptian–Albian age to this tectonic event. In the same time span compressional deformation likely controlled the Parnassus Platform emersion soon after the deposition of allochthonous bauxitic laterites (b3 horizon) derived from degradation of penecontemporaneous weathering crusts developed on the more internal units (i.e. Pelagonian and Ophiolitic obducted units) (Fig. 21d). The stratigraphic gap corresponding to the third bauxite horizon (b3) is the longest one since it extended from part of the Aptian to at least the lower Turonian (> 20 Ma; Combes et al. Reference Combes, Fourcade, Masse and Philip1981; Carras & Tselepidis, Reference Carras, Tselepidis, Solakius and Kati2001). The stable and relatively high sea level record from Valanginian until Late Cretaceous times (Haq, Reference Haq2014) allows ascribing the b3 emersion episode to the far-field effects of compressional tectonics.

This prominent Aptian–Albian tectonic phase was also recorded by upper-greenschist to blueschist metamorphic events in the inner Pelagonian Zone (Barton, Reference Barton1976; Yarwood & Dixon, Reference Yarwood and Dixon1977; Schermer, Lux & Burchfiel, Reference Schermer, Lux and Burchfiel1990; Lips, White & Wijbrans, Reference Lips, White and Wijbrans1998) (Fig. 4), and by nappe stacking and high-pressure conditions in the Rhodope Massif (130–115 Ma; Burg, Reference Burg, Skourtsos and Lister2012). Furthermore, the ophiolitic units of the EOB, mostly represented by island arc and back-arc basin ophiolites (Ferrière & Stais Reference Ferrière and Stais1995; Saccani et al. Reference Saccani, Bortolotti, Marroni, Pandolfi, Photiades and Principi2008a ; Bonev et al. Reference Bonev, Marchev, Moritz and Collings2015), emplaced eastwards above the Eurasian margin units previously affected by Late Jurassic high-pressure metamorphism (Bonev et al. Reference Bonev, Marchev, Moritz and Collings2015 and quoted references). This emplacement was postdated by unconformable sedimentation of Lower Cretaceous carbonates in the Circum-Rhodope Belt (i.e. westernmost part of the Rhodope Massif; Mercier Reference Mercier1968; Stais & Ferrière, Reference Stais and Ferrière1994; Ivanova, Bonev & Chatalov, Reference Ivanova, Bonev and Chatalov2015). This testifies that at the Late Jurassic – Early Cretaceous transition the magmatic arc and back-arc oceanic basin located to the west of the Eurasian margin (i.e. the Guevgueli back-arc ophiolites and the Paikon arc ophiolites) had already ceased their magmatic activity and were emplaced above deformed and metamorphosed units of the Eurasian margin (Bonev et al. Reference Bonev, Marchev, Moritz and Collings2015; Saccani et al. Reference Saccani, Chiari, Bortolotti, Photiades and Principi2015; Ferrière, Baumgartner & Chanier, Reference Ferrière, Baumgartner and Chanier2016) (Figs 21c, 22b).

Finally, anatectic melts of late Aptian age (Schenker et al. Reference Schenker, Burg, Kostopoulos, Moulas, Larionov and Von Quadt2014) were recognized in the eastern Pelagonian rim.

All these events, along with the lack of evidence of ophiolites younger than Late Jurassic in the Albanian–Hellenic sector of the orogen, allow interpreting the D2 deformation phase as a manifestation of the final continental collision of the Pelagonian Margin (Adria Plate) with the Rhodope Massif (European Plate). In this interpretation the LBF sedimented in a temporal window placed between the ongoing obduction and continental collision (Fig. 21c).

Structural analysis of D2 structures points to a westward tectonic transport. The observable tectonic relationships between the LBF and the inner tectonic units allow explaining the D2 deformation as being the result of the activation of a sole thrust at the base of the Pelagonian Platform and rejuvenated westward tectonic translation of the overlying ophiolitic thrust sheets.

However, the occurrence of Late Cretaceous ophiolites in the Kozara Mountains (northern Bosnia and Herzegovina; Karamata et al. Reference Karamata, Olujić, Protić, Milovanović, Vujnović, Popević, Memović, Radovanović, Resimić-Sarić, Karamata and Janković2000; Ustaszewski et al. Reference Ustaszewski, Schmid, Lugovic, Schuster, Schaltegger, Bernoulli, Hottinger, Kounov, Fügenschuh and Schefer2009, Reference Ustaszewski, Kounov, Schmid, Schaltegger, Krenn, Frank and Fügenschuh2010) testifies that the collision in the northern sectors of the belt occurred in latest Cretaceous – Paleocene times (Schmid et al. Reference Schmid, Bernoulli, Fügenschuh, Matenco, Schefer, Schuster, Tischler and Ustaszewski2008).

Two main hypotheses exist concerning the timing of collision in the Hellenic–Dinaric chain and final closure of the Maliac–Vardar Ocean: latest Cretaceous–Tertiary throughout the whole orogen (Robertson & Dixon, Reference Robertson, Dixon, Dixon and Robertson1984; Robertson et al. Reference Robertson, Clift, Degnan and Jones1991, Reference Robertson, Trivić, Đerić and Bucur2013; Pamic, Reference Pamic2002; Schmid et al. Reference Schmid, Bernoulli, Fügenschuh, Matenco, Schefer, Schuster, Tischler and Ustaszewski2008) or latest Jurassic – Early Cretaceous (Papanikolaou, Reference Papanikolaou1997, Reference Papanikolaou2009, Reference Papanikolaou2013; Meinhold et al. Reference Meinhold, Kostopoulos, Reischmann, Frei and BouDagher-Fadel2009; Bortolotti et al. Reference Bortolotti, Chiari, Marroni, Pandolfi, Principi and Saccani2013; Schenker et al. Reference Schenker, Burg, Kostopoulos, Moulas, Larionov and Von Quadt2014), at least in the Albanian–Hellenic sector. To date, both hypotheses could be considered valid; however, the evidence so far collected in Albania and Greece and the differences between the southern and northern sectors of the belt (see discussion in Saccani et al. Reference Saccani, Chiari, Bortolotti, Photiades and Principi2015) lead us to prefer a diachronous closure model explaining earlier collision in the south and later in the northern sectors.

This diachronous closure could alternatively be explained as the consequence of variable widths of different oceanic sectors, possibly separated by lithospheric lineaments (e.g. the Scutari–Pec fault zone, Fig. 1; Dercourt, Reference Dercourt1967; Kissel, Speranza & Milicevic, Reference Kissel, Speranza and Milicevic1995), or due to oblique subduction (Smith & Spray, Reference Smith, Spray, Dixon and Robertson1984).

6.d. The Late Cretaceous post-collisional tectonics

After the climax of the collisional event in the Boeotia region, the sedimentation resumed with the flysch deposition in the Boeotian Basin (i.e. UBF) and the neritic carbonates of the Mesoautochthonous Complex above the ophiolitic thrust sheets and Pelagonian Unit (Fig. 21d). Both the UBF and the first cycle of Mesoautochthonous sequences are underlined by unconformable contacts marked by lateritic crusts, coarse-grained mass flow deposits and erosional surfaces. The high proportion of ophiolite debris in the UBF along with common calcareous mass flows and lateritic debris (Nirta et al. Reference Nirta, Moratti, Piccardi, Montanari, Catanzariti, Carras and Papini2015) implies still active tectonics along the eastern edge of the Boeotian Basin that triggered drowning of ephemeral platforms, emersion episodes and catastrophic collapses along the shelf break. The sedimentary environment of the UBF could be that of a foreland basin, while the Mesoautochthonous Complex contemporaneously sedimented in the subsiding hinterland sectors after the tectonic climax related to collision (D2 phase).

The correlation of the UBF with the coeval siliciclastic turbidite deposits of the Pindos Zone (Premier Flysch du Pinde; Aubouin, Reference Aubouin1957; Fleury, Reference Fleury1975; Maillot, Reference Maillot1979; Wagreich et al. Reference Wagreich, Pavlopoulos, Faupl and Migiros1996; Neumann, Reference Neumann2001) is also proposed. A similar composition and palaeocurrent data of the UBF and ‘Premier Flysch du Pinde’ (Neumann, Reference Neumann2001) point to a common provenance from the internal (eastern) zones that produced turbidite currents then deposited in the basinal areas of the Boeotian Zone and those bordering the eastern Pindos Basin.

There is evidence of an older deformation phase preceding the ubiquitous Tertiary events that gave rise to the older deformations observed in the UBF and cross-cuts D2 structures of the LBF (D3 deformation phase).

This deformation phase is poorly constrained because of the paucity of explicative outcrops in the UBF and resistance in developing mesoscale deformations in the Mesoautochthonous Complex. However, we tentatively attribute to this deformation a Late Cretaceous age as no D3 deformations were recognized in the Tertiary flysches.

D3 deformations, showing westward vergence, could be linked with a renewed westward translation of the Internal Hellenid units related to the ongoing Africa–Eurasia convergence. Progressive westward thrusting was coupled with syn-orogenic exhumation of portions of the eastern Pelagonian Zone as testified to by available geochronological data (T. Most, unpub. Ph.D. thesis, Eberhardt-Karls-Universität Tubingen, 2003; Schenker, Fellin & Burg, Reference Schenker, Fellin and Burg2015). Exhumation of metamorphic terrains of the Pelagonian Zone was linked with an uplifted dome, possibly through a regional antiformal structure (Kilias et al. Reference Kilias, Frisch, Avgerinas, Dunkl, Falalakis and Gawlick2010; Schenker et al. Reference Schenker, Burg, Kostopoulos, Moulas, Larionov and Von Quadt2014), that was successively eroded and supplied both continental-derived metamorphic and ophiolitic debris to the adjacent Mesoautochthonous basins (T. Most, unpub. Ph.D. thesis, Eberhardt-Karls-Universität Tubingen, 2003; Photiades et al. Reference Photiades, Carras, Bortolotti, Fazzuoli and Principi2007; Schenker et al. Reference Schenker, Burg, Kostopoulos, Moulas, Larionov and Von Quadt2014) (Fig. 21d).

Sedimentation of the Mesoautochthonous Complex was synchronous with the development of an arc-related calc-alkaline magmatism in the upper plate, the Eastern Srednogorie and Panagyurishte area (Bulgaria). The northward continuation of these magmatic districts as far as Romania concurs to form the Apuseni–Banat–Timok–Srednogorie magmatic–metallogenic belt (ABTS; Jankovic, Reference Jankovic1997; von Quadt et al. Reference von Quadt, Moritz, Peytcheva and Heinrich2005; Zimmerman et al. Reference Zimmerman, Stein, Hannah, Koželj, Bogdanov and Berza2008; Richards, Reference Richards2015).

A wide debate exists on the mechanism guiding the magma generation in the ABTS and its development with respect to the onset of continental collision, with different models reflecting divergences in the palaeogeographic reconstructions. Summarizing, the proposed models include those favouring pre-collisional subduction-related magmatism, eventually accompanied by slab rollback (Lips, Reference Lips, Blundell, Neubauer and von Quadt2002; von Quadt et al. Reference von Quadt, Moritz, Peytcheva and Heinrich2005; Chambefort & Moritz Reference Chambefort and Moritz2006; Zimmerman et al. Reference Zimmerman, Stein, Hannah, Koželj, Bogdanov and Berza2008; Gallhofer et al. Reference Gallhofer, Quadt, Peytcheva, Schmid and Heinrich2015) or syn-collisional slab tearing/slab break-off (Neubauer, Reference Neubauer, Blundell, Neubauer and von Quadt2002) and post-collisional orogenic collapse (Berza, Constantinescu & Vlad, Reference Berza, Constantinescu and Vlad1998; Bojar, Neubauer & Fritz, Reference Bojar, Neubauer and Fritz1998; Willingshofer, Neubauer & Cloetingh, Reference Willingshofer, Neubauer and Cloetingh1999; Iancu et al. Reference Iancu, Berza, Seghedi, Gheuca and Hann2005). In particular, some authors (Lips, Reference Lips, Blundell, Neubauer and von Quadt2002; Gallhofer et al. Reference Gallhofer, Quadt, Peytcheva, Schmid and Heinrich2015) considered the continuous subduction of oceanic crust and slab steepening/rollback as the only mechanism that guided the evolution of the whole ABTS during Late Cretaceous time. They assumed that a new oceanic subduction started in the Maliac–Vardar Ocean in Albian time, after the Early Cretaceous tectonic phase (D2). In this hypothesis, melt production started as soon as the subducting slab reached the partial melting zone and continued throughout Late Cretaceous time fuelled by a continuous subduction until continental collision in latest Cretaceous time (Gallhofer et al. Reference Gallhofer, Quadt, Peytcheva, Schmid and Heinrich2015). On the other hand, recent research along the Alpine–Himalayan belt has revealed that many magmatic deposits previously assumed to be subduction related are in fact syn- or post-collisional (Shafiei, Haschke & Shahabpour, Reference Shafiei, Haschke and Shahabpour2009; Moritz et al. Reference Moritz, Márton, Ortelli, Marchev, Voudouris, Bonev, Spikings and Cosca2010; Imer, Richards & Creaser, Reference Imer, Richards and Creaser2013) and sometimes occur long (30–40 Ma) after the end of oceanic lithosphere subduction (Hou et al. Reference Hou, Yang, Qu, Meng, Li, Beaudoin, Rui, Gao and Zaw2009, Reference Hou, Zhang, Pan and Yang2011).

In our interpretation, the evolution of the ABTS possibly follows the diachronism in the collision as testified to by the different ages of the older magmatic events (and beginning of extension) varying from Cenomanian–Turonian in the south (Eastern Srednogorie and Panagyurishte area, Bulgaria) up to Santonian–Campanian (Banat and Apuseni regions, Romania) in the north (see data compilation in Gallhofer et al. Reference Gallhofer, Quadt, Peytcheva, Schmid and Heinrich2015). We interpret the Mesoautochthonous Complex and the Eastern Srednogorie and Panagyurishte magmatic arcs (roughly corresponding to the Hellenic and Albanian sectors of the belt) as connected to regional extension in the upper plate following an initial slab steepening started soon after the continental collision (i.e. Albian), followed by warming of the stalled slab and its successive break-off (Wortel & Spakman, Reference Wortel and Spakman2000; Neubauer, Reference Neubauer, Blundell, Neubauer and von Quadt2002) and widespread orogenic collapse from late Cenomanian–Turonian times (Fig. 22c,d). Numerical modelling shows that the delay in time between the onset of collision and final slab break-off can be 20 Ma (Duretz, Gerya & May, Reference Duretz, Gerya and May2011; van Hunen & Allen, Reference van Hunen and Allen2011) in the case of old, and therefore cold and strong, subducting slabs. At the time of collision onset (c. 115 Ma) the subducted Maliac–Vardar oceanic lithosphere was represented by old and strong Triassic to Middle Jurassic crust, thus the final break-off could be postulated at c. 95 Ma (i.e. late Cenomanian–Turonian), which corresponds with the age of the comagmatic sediments in the Eastern Srednogorie area (Nachev & Dimitrova, Reference Nachev and Dimitrova1995). The formation of the arc magmatism was facilitated by reduced plate convergence starting from early Late Cretaceous time (Rosenbaum, Lister & Duboz, Reference Rosenbaum, Lister and Duboz2002; Bertrand, Guillou-Frottier & Loiselet, Reference Bertrand, Guillou-Frottier and Loiselet2014). Furthermore, we assume that after the Aptian–Albian continental collision, a slow subduction continued with the involvement of the Pelagonian thinned crust (Stampfli, Marcoux & Baud, Reference Stampfli, Marcoux and Baud1991; Robertson et al. Reference Robertson, Clift, Degnan and Jones1991). Collision and subduction of the portions of thinned continental crust concur in reducing the lateral force acting on the slab causing a rapid slab steepening enhanced by its vertical gravitational force (Fig. 22c). This mechanism may have favoured orogenic collapse in the upper plate and easier ascent of fertile magmas to the upper crust (Berza, Constantinescu & Vlad, Reference Berza, Constantinescu and Vlad1998).

6.e. Tertiary crustal shortening and thickening: D4 and D4_bt deformation phases

During late Paleocene time, ubiquitous shaly pelagic sediments (pre-flysch lithofacies) unconformably covered the tectonic stack levelling the sea floor morphology (Fig. 21e). The following turbiditic sedimentation onset in early Eocene time was almost synchronous in the different geotectonic zones. On the other hand, terminal ages are diachronous showing a younging trend towards the western sectors. In the studied area, the Pelagonian Flysch lasted until early Eocene time, the Parnassus Flysch until early–middle Eocene time and the Pindos Flysch until late Eocene time. This trend reflects the W-verging tectonic transfer related to late-stage thrusting that remobilized the previously formed tectonic stack above the Parnassus Platform (D4 deformation phase). The age of overthrusting of the internal tectonic units above the Parnassus Platform is constrained by the terminal age of the Parnassus Flysch (early–middle Eocene). Moreover, during Lutetian time, ophiolite slide blocks and debris flow deposits glided in the Pindos Flysch in the Etolia region just in front of the Vardoussia slope (C. Beck, unpub. Ph.D. thesis, Univ. Science and Technology, Lille, 1975; Beck, Reference Beck1980; Figs 2, 21f). We locate the provenance area of these ophiolite-bearing mass flows in the Iti ocean-derived klippe located on top of the Parnassus Unit, thus dating the age of overthrusting of the Internal Hellenid units above the Parnassus Platform to middle Eocene (Lutetian) time.

Extensive nappe stacking and continental crust thickening characterized the Tertiary compressive phase. Protracted convergence between Adria and Eurasia during the D4 tectonic phase led to the final accretion of the Pelagonian Unit and its decoupling from the subducting Adriatic slab (Jolivet & Brun, Reference Jolivet and Brun2010) and consequent shifting of the compressive front to the more external southwestern portions of the belt. As a result, the Pindos and Gavrovo zones were progressively involved in the fold and thrust belt and metamorphosed to form the Cycladic Blueschist Complex since middle Eocene time (Jolivet & Brun, Reference Jolivet and Brun2010 and references therein) (Fig. 21f).

The last compressional event recognized in the study area (D4_bt) is localized on the northeastern rim of the Parnassus Platform that acted as a tectonic buttress locally conditioning the tectonic migration. Along the northern border of the Helicon mountain range, the Parnassus Cretaceous limestones thrust northwards folding and cross-cutting the previous D4 structures. Furthermore, the Boeotian Zone limestone sequence was squeezed between the Pelagonian and Parnassus units guiding it to thrust northwards above the previously stacked units including the Eocene Pelagonian Flysch (Figs 6, 8).

6.f. Aegean extensional regime: D5 deformation phase

The previously stacked units were then dissected by both high- and low-angle faults under an extensional stress field with a horizontal ENE-oriented minor principal (σ3) stress axis (Fig. 19). Extension inverted some D4 thrusts as low-angle fault zones (see Figs 16, 21f) or as regional detachments (Papanikolaou, Gouliotis & Triantaphyllou, Reference Papanikolaou, Gouliotis, Triantaphyllou, Hinsbergen, Edwards and Govers2009). Clear evidence within the studied area is represented by the tectonic contacts at the base of the internal units overthrust above the northern flank of the Parnassus Platform (Kranis & Papanikolaou, Reference Kranis and Papanikolaou2001; Fig. 6).

The extensional regime, whose main regional effect is the opening of the Aegean Sea and exhumation of the Cycladic Blueschist Complex, started in northern Greece in late Eocene time (Gautier et al. Reference Gautier, Brun, Moriceau, Sokoutis, Martinod and Jolivet1999; Lips, White & Wijbrans, Reference Lips, White and Wijbrans2000; Jolivet et al. Reference Jolivet, Faccenna, Goffé, Burov and Agard2003; Jolivet & Brun, Reference Jolivet and Brun2010) and progressively moved southwards while accretion continued at the subduction zone. Aegean extension is commonly considered a consequence of the slab retreat of Adria with respect to Eurasia (Le Pichon & Angelier, Reference Le Pichon and Angelier1979; Meulenkamp et al. Reference Meulenkamp, Wortel, van Wamel, Spakman and Hoogerduyn Strating1988; Jolivet et al. Reference Jolivet, Faccenna, Huet, Labrousse, Le Pourhiet, Lacombe, Lecomte, Burov, Denèle, Brun, Philippon, Paul, Salaün, Karabulut, Piromallo, Monié, Gueydan, Okay, Oberhänsli, Pourteau, Augier, Gadenne and Driussi2013).

In the studied area, transition from a compressional to extensional regime was accompanied by clockwise block rotation accommodated by NNE-trending dextral transversal faults developed at the regional scale (e.g. Distomon (DSTF) and Aliki (ALKF) fault zones; see Fig. 6). Block rotations emphasized the northward back-thrust of the Boeotian Unit with the activation of the ATRF zone (Fig. 6) with left-lateral movement. Clockwise block rotations were possibly accentuated since c. 5 Ma by the interaction of the right-lateral North Anatolian Fault system (Armijo et al. Reference Armijo, Meyer, King, Rigo and Papanastassiou1996, Reference Armijo, Meyer, Hubert and Barka1999) that also enhanced extension in the North Aegean Sea, Euboea–Sperchios Graben and Corinth Basin (Goldsworthy, Jackson & Haines, Reference Goldsworthy, Jackson and Haines2002).