The Miocene of the Hatay area, S Turkey: Transition from the Arabian passive margin to an underfilled foreland basin related to closure of the Southern Neotethys Ocean

Sedimentary Geology 198 (2007) 93 – 124 www.elsevier.com/locate/sedgeo

The Miocene of the Hatay area, S Turkey: Transition from the Arabian passive margin to an underfilled foreland basin related to closure of the Southern Neotethys Ocean Sarah J. Boulton a,⁎, Alastair H.F. Robertson b a

School of Earth, Ocean and Environmental Sciences, University of Plymouth, Drake Circus, Plymouth, Devon, PL4 8HS, UK b School of GeoSciences, University of Edinburgh, Grant Institute, West Mains Road, Edinburgh, EH9 3JW, UK Received 17 May 2006; received in revised form 9 November 2006; accepted 6 December 2006

Abstract Miocene sediments of the Hatay area document the distal, southerly, part of an “underfilled” peripheral foreland basin related to the closure of the Southern Neotethys. The basin is underlain by latest Cretaceous and Eocene shallow-marine carbonates, interpreted as the youngest part of the Arabian passive margin succession. Unconformably overlying Lower Miocene conglomerates, sandstones and palaeosols (up to 250 m thick) represent braided river deposits derived from an uplifted area to the south. During the Middle Miocene, carbonate sediments accumulated in a variety of shallow-marine environments dominated by a northward-sloping ramp, with non-marine, peritidal, lagoonal and coral reef deposition in local areas. The Upper Miocene succession is mainly deeper-water hemipelagic marl with clastic interbeds. Messinian evaporites were deposited near the depocentre. The Early Miocene coarse fluvial sediments above the regional unconformity reflect the development of a flexural bulge, or reactivated basement structure. The Middle Miocene carbonates and Upper Miocene marls are interpreted as the lower and middle units of an “underfilled trinity”, considered typical of peripheral foreland basins. Similar Lower Miocene coarse clastic sediments and MidMiocene shallow-water carbonates unconformably overlie the Arabian platform succession c. 75 km to the north, where the succession continues upwards into thick shallow-marine to non-marine terrigenous sediments (upper unit of the “trinity”), sourced from the overthrust load. Comparable coarse clastic sediments are absent from the distal foreland basin in the Hatay area, probably reflecting its more southerly position, structural barriers in the foreland and the westward bypassing of sediment towards the Mediterranean Sea. Regional convergence halted by the end of the Miocene and was followed by strike–slip and westward “tectonic escape”. As a result the Miocene foreland basin was dismembered to form the Plio-Quaternary transtensional Hatay Graben. This like many other foreland basins shows some features which are not accommodated by ideal models of foreland basin evolution. © 2006 Elsevier B.V. All rights reserved. Keywords: Peripheral foreland basin; Tethyan suture zone; Arabian platform; S Turkey; E Mediterranean; Shallow-marine carbonates; Braided river sediments

1. Introduction

⁎ Corresponding author. Fax: +44 1752 233117. E-mail address: [email protected] (S.J. Boulton). 0037-0738/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.sedgeo.2006.12.001

Peripheral foreland basins represent the flexural response of the lithosphere to the loading of an overriding plate during collision (Stockmal and Beaumont, 1987). The sedimentary fill records an interplay between

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regional tectonics, eustasy and normal sedimentary processes (Beaumont, 1981; Allen et al., 1986). Numerous case histories of foreland basins of different age and location have been published, notably for the Appalachians (Hiscott et al., 1986), Alps (Pfinner, 1986; Sinclair, 1997) and the Apennines (Ricci-Lucci, 1986). Also, much effort has gone into the modelling of foreland basins using various techniques (e.g. Allen et al., 1986; Stockmal and Beaumont, 1987; Sinclair, 1997; Allen et al., 2001; Rafini and Mercier, 2002; Clevis et al., 2004). Most field-based studies and the resulting models have focused on the proximal part of foreland basins, where the effects of lithospheric loading are pronounced close to the overthrust load, and where deep-water ‘flysch’ and late-orogenic ‘molasse’ predominate. However, proximal parts of foreland basins are commonly deformed or partly obscured by the overthrust load. Also, the relative effects of tectonics versus eustatic sea-level change and normal (i.e. autocyclic) sedimentation may not be easy to untangle in such deep-water “flysch” basins. Less emphasis has been placed on the distal (i.e. “cratonward”) parts of foreland basins. Such areas commonly include varied shallow-marine sediments that provide sensitive indicators of the interplay between tectonics, eustasy and normal sedimentary processes. Such distal foreland areas may be located up to several hundred kilometres from the overthust load, represented by the front of a

mountain belt. Thus, their recognition as part of a foreland basin system may not be obvious, especially in areas affected by later tectonics or erosion. Here, we focus on the Miocene sedimentation of the Hatay area of south-central Turkey, which we interpret as part of the distal (“cratonward”) area of a foreland basin related to the final closure of the Southern Neotethys. This is an excellent area to study processes affecting the distal part of a peripheral foreland basin and its development through time. The sediments were deposited on the northern edge of the Arabian platform, and include a wide variety of marine, to non-marine sediments, which are well exposed and relatively undeformed. Previously, the Miocene sediments of the Hatay area were not distinguished from the Plio-Quaternary sediments of the Hatay Graben, which formed in a contrasting tectonic regime (Boulton et al., 2006). In this paper, we will first describe and interpret the facies and palaeoenvironments of the Miocene sediments of the Hatay area. We then assemble salient information and interpretations that allow these sediments to be interpreted as the distal part of a foreland basin, in the light of evidence from other areas. We also evaluate our results in the light of the regional collisional setting of the Arabian (African) and Tauride (Eurasian) plates and take account of theoretical models of foreland basin evolution. The sedimentary evidence (field and microfacies data) from the Hatay area is used to support a

Fig. 1. Neotectonic map of the Eastern Mediterranean showing the major tectonic lineaments. Box indicates location of the inset showing detail of the Hatay Graben and vicinity. Letters on the figure: HG Hatay Graben; KR Karasu Rift; AP Amik Plain; NAFZ North Anatolian Fault Zone; DSFZ Dead Sea Fault Zone; EAFZ East Anatolian Fault Zone. Modified from Boulton et al. (2006).

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three-stage model for the sedimentary development of a distal foreland basin related to closure of the Southern Neotethys ocean, with implications for other comparable areas. Existing theoretical models for peripheral foreland basin evolution may need to be modified to take account of evidence from the Hatay area. 2. Regional setting Northward subduction was active during the Late Cretaceous within the Eastern Mediterranean region (Şengör and Yılmaz, 1981; Robertson and Dixon, 1984) resulting in the progressive closure of the Southern Neotethys. North of the Hatay area, the closure of the Southern Neotethys resulted in the southward emplacement of ophiolites, notably the Hatay and Baer–Bassit ophiolites, onto the Arabian Platform (Yılmaz, 1993; Robertson, 2002). The remaining Southern Neotethys oceanic basin closed during Early Cenozoic time and suturing was complete by the Mid-Miocene. The suture zone to the north is represented by the Misis–Andırın Complex (Misis Complex of Kelling et al., 1987), to the

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north of the Hatay region (Kozlu, 1997; Robertson et al., 2004; Fig. 1), and, further east, by the Bitlis suture, which extends through SE Turkey into Iran (Aktaş and Robertson, 1984; Hempton, 1987; Yılmaz, 1993). The Misis–Andırın Complex is interpreted as an accretionary prism related to southward closure of the Southern Neotethys during Early Cenozoic time (Robertson et al., 2004), culminating in continental collision. Thick turbiditic sediments of Early Miocene age are exposed structurally beneath the Misis–Andırın Complex and have been interpreted as the proximal part of a peripheral foreland basin (Gökçen et al., 1988; Derman et al., 1996; Robertson et al., 2004; Kelling et al., 2005). Similar Miocene foreland basin facies occur eastwards through southeastern Turkey (i.e. Lice Formation; Aktaş and Robertson, 1984) and extend into Iran (Sharland et al., 2004). After the end of the Miocene, there was a regional switch to the westward ‘tectonic escape’ of Anatolia towards the Aegean Sea (Şengör et al., 1985; Reilinger et al., 1997). Large-scale sinistral strike–slip faulting dominated the tectonics of the region during Plio-

Fig. 2. Simplified geological map of the Hatay area showing the places and locations of sedimentary logs discussed in the text (modified from Boulton et al., 2006).

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Fig. 3. Sedimentary log of the sedimentary succession exposed in Harbiye Gorge (see log 1 on Fig. 2).

Quaternary time (i.e. Arpat and Şaroğlu, 1972; Şengör et al., 1985; Mart and Rabinowitz, 1986; Barka and Kadinsky-Cade, 1988; Lyberis, 1988). The changed regional setting was reflected in the development of the

Plio-Quaternary Hatay Graben, which is superimposed on the previous foreland basin setting. By contrast, in areas further west, in the vicinity of the Eastern Mediterranean Sea (e.g. Cyprus area), part of

Fig. 4. Key to the symbols used in the stratigraphic logs.

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the Southern Neotethys remained in a pre-collisional, or incipient collisional setting during Miocene to Recent time (Robertson, 1998, 2000), but is outside the scope of this paper. 3. Previous work In the area studied, the Mesozoic Arabian platform succession was tectonically over-ridden by the Hatay Ophiolite during the latest Cretaceous (Çoğulu, 1974; Delaloye et al., 1980; Delaloye and Wagner, 1984). This ophiolite was emplaced southwards during Campanian to Maastrictian time (Tinkler et al., 1981), related to partial closure of the Southern Neotethys to the north of the Arabian passive continental margin (Yılmaz, 1993; Robertson, 1998, 2000, 2002). Upper Cretaceous carbonates were deposited unconformably on the emplaced ophiolite. Tinkler et al. (1981) and Pişkin et al. (1986) reported an extensive conglomerate below the Cenozoic sedimentary cover sequence. This was suggested to be Maastrictian in age and the result of erosion along a fault scarp. Parlak et al. (1998) do not mention this conglomerate but describe the basal transgressive formation as shallow-marine calcarenites, marls and clayey limestones. Upper Cretaceous shallow-marine carbonates are known to pass transitionally upwards into Late Palaeocene–Middle Eocene calcarenites and limestones rich in microfossils (Pişkin et al., 1986). This was followed by a phase of deformation associated with folding and local thrusting (Tinkler et al., 1981). The carbonates form part of a widespread Early Cenozoic Arabian carbonate platform (Gvirtzman et al., 1989). Tinkler et al. (1981) and Pişkin et al. (1986) mention the presence of an extensive conglomeratic horizon composed of limestone and ophiolitic debris at the base of a Miocene sedimentary sequence. Parlak et al. (1998) also describe Middle–Late Miocene sediments in the Karasu Rift, a separate neotectonic structure to the northeast of the Hatay Graben (Fig. 1, inset), as lying discordantly on these conglomerates. Tinkler et al. (1981) and Pişkin et al. (1986) describe Middle Miocene limestones as being composed of reef-derived material, transitional upwards to Late Miocene marl with intercalated sandstone and mudstone beds. The Quaternary to recent setting of the area has received limited attention (Pirazzoli et al., 1991; Erol and Pirazzoli, 1992; Rojay et al., 2001). Recently, Boulton et al. (2006) have summarised the Late Cretaceous to Recent tectonic-related development of the Hatay area based on new evidence and concluded that the Miocene sediments accumulated in a

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regional foreland basin setting. These authors established that Miocene sedimentation within the Hatay region predated the formation of major faulting within the Plio-Quaternary Hatay Graben. The Hatay Graben (Figs. 1, 2) was interpreted as a Plio-Quaternary transtensional feature with a pronounced topographic relief. This graben post-dates the final collision of the Arabian (African) and Tauride (Eurasian) plates and was influenced by the westward tectonic escape of Anatolia towards the Aegean. The present paper builds and extends this preliminary account of the area. 4. Stratigraphic context Although, Dubertret (1939, 1953) carried out early studies, a stratigraphy was not developed until Atan

Fig. 5. Sedimentary log of the type section of the Balyatağı Formation near Enek (map reference Antakya-P36-d2 0247339/40060370247796/4006188; see log 2 on Fig. 2).

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Table 1 Age, lithology and microfossil data for the sediments of the Hatay Graben Formation name

Age

Lithology

Samandağ

Pliocene

Marl and sandstone

Vakıflı Member Nurzeytin

Messinian Serravalian/ Tortonian Langhian

Marl, limestone and sandstone Limestone

Aquitanian/ Burdigalian Lutetian Lutetian

Conglomerates and palaeosols Limestone and marl P10 Acarinina bullbrooki b Limestone P10 Acarinina bullbrooki b

Sofular Balyatağı Kışlak Okçular Kaleboğazı a b c

Planktonic foraminifera zone

N9 Orbulina universa b N9 Orbulina suturalis N8 Praeorbulina glomerosa curva b N7 Globigerinoides trilobus b

Late Cretaceous Limestone and sandstone

Selected microfossils Globerinoides ruber, Globorotalia scitula, Globigerinoides trilobus, Globigerinoides sacculifer a Globoquadrina altispira Orbulina universa, Hastigerina sp., Orbulina suturalis a Praeorbulina gloerosa curva, Orbulina suturalis a Globergerinoides triobus, Globergerinoides ruber a Acarinina bullbrooki, Morozovella spirulosa a Morozovella aragonensis, Globigerina ineqispira a Globotruncana arca, Globotruncana gansseri, Globotruncana mayaroensis c

Boulton et al. in press. Şafak, 1993. Pişkin et al., 1986.

(1969), divided the Miocene sediments of the area as a whole into two formations, the Yazır and Enek formations, with the Enek Formation consisting of two

members. Atan (1969) also divided Palaeocene to Eocene limestones into two formations (Figs. 2–5; Table 1); named the Okçular and Kıslak formations.

Table 2 Summary description and interpretation of limestone facies of the Hatay Graben Facies name

Description

Hard, white/cream microbial limestone; bedding thickness b1 m. Wavy lamination, fenestral porosity and desiccation breccia common. Vertical burrows in some horizons and local chert nodules. Lime mudstone Cream coloured lime mudstone with occasional ophiolitc clasts. Some beds are silicified preserving microbial laminations. Packstone–rudstone Medium-grained, white packstone–rudstone. No sedimentary features. Bioclastic material common, especially large benthic foraminifera (i.e. Nummulites) and oncolites. Wackestone–rudstone Coarse bioclastic marly limestones, generally wackestones and rudstones. Fossil material is variably fragmented and dominated by bivalves and gastropods (e.g. Ostrea, Pecten), echinoids (Clypeaster sp., Echimolampas sp., Schizaster sp., Psammechinus sp.), oncolites and coral. Beds are 1–6m thick and lack sedimentary structures. This limestone is generally soft and grades into a harder bioclastic limestone in each cyclic unit (wackestone–packstone). Wackestone–packstone Wackestone–packstone containing the same faunal assemblage as the wackestone– rudstone; additionally, rare, apparently in situ coral heads are present, together with numerous oncolites. The base of this facies is transitional with wackestone–rudstone facies but the upper surface is sharp and irregular. Microbial mudstone Cream coloured, fine-grained limestones with irregular bedding surfaces and beds of variable thickness from a couple of centimetres to ∼1 m. These beds are irregular, often laterally discontinuous, and contain laminations and ripples. Reefal wackestone Rubbly wackestone–packstone containing fragmented and intact bivalves (e.g. Pecten sp., Mya sp.), gastropods (e.g. Conus sp., Architectonica simplex), echinoids (e.g. Opissaster polygonalis), coralline fragments; additional shelly material and rounded ophiolitic pebbles. Reef rudstone Coralline rudstone 5–10 m in thickness. This limestone has an asymmetrical dome-like morphology. Reefal packstone Packstone–rudstone (2 m thick), mainly composed of fragments of coralline algal and bivalves. Benthic and planktic foraminifera are also present. Microbial limestone

Interpretation Intertidal low-energy environment Intertidal deposition with continental influence Shallow/open marine high energy environment Sub-tidal, part of the peritidal cyclothem.

Shallow-marine, part of the peritidal cyclothem.

Lagoonal

Reef talus found adjacent to patch reefs

Small patch reef Shallow-water back reef facies

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Table 3 Summary description and interpretation of Miocene mudstone facies of the Hatay Graben Facies name

Description

Mudstone

Interpretation

Red to reddish brown mudstone and marl with caliche nodules. Parallel and convolute lamination in upper parts of beds. Rootlets common. Pebbly mudstone Fine-grained interbedded mudstone varies in colour and hardness; some beds are soft and dark reddish brown, whereas others are hard and pink, to white, in colour. Common isolated “floating” pebbles, dominantly serpentinite, are generally b0.5 cm in size and sub-rounded, to sub-angular. Some mudstones are pebbly, with the pebble content generally decreasing upwards. Pink mudstone Fine-grained, pink–grey, carbonaceous mudstone that overlies wackestone– packstone facies in each cyclothem (Table 2). The beds are 50 cm–2 m thick and generally lack sedimentary structures. Small gastropods, roots and other plant material are present locally. Some beds exhibit a mottled appearance due to the presence of caliche. Some beds contain little or no caliche, whereas this is well developed in others. Marl Rip-up clasts of marl occur in some of the harder, unlaminated beds. A regularly repeating sequence of bedded marl and marly limestone follows; beds are ∼ 50 cm thick and bedding surfaces are highly irregular. Diagenetic chert nodules are abundant, some apparently infilling vertical burrows. In addition, a dewatering pipe, ∼ 1.5 m long, was observed. Sandy mudstone Pink to white, mottled mudstones with wavy laminations and occasional small lenses of sandy material. Bedding thickness varies from 5–20 cm. Interbedded mudstone Pink–white nodular lime mudstones, brown mudstone (with desiccation cracks) and thin interbeds of nodular chalk; rippled cream-coloured mudstone is also present. Laminated mudstone Reddish brown mudstone, irregular laminations and laterally variable in thickness (∼ 20 cm thick). No rootlets. Foraminiferal marl Light to medium grey marl, generally fossiliferous with numerous foraminifera, mostly planktic but occasional benthic. Detrital plant material is also often present.

Atan's (1969) lithostratigraphic terminology continued to be used in the Hatay Graben (Selçuk, 1981; Pişkin et al., 1986) until Şafak (1993) used microfossils (planktic foraminifera and ostracods) to divide the Miocene succession into five new stratigraphic units; the Balyatağı, Sofular, Tepehan, Nurzeytin and Vakıflı formations. Recently, Boulton et al. (in press) have revised the stratigraphy of the area, using a combination of sedimentary analysis, new biostratigraphic evidence (using planktic foraminifera and ostracods) and Srisotope dating; this stratigraphy is used here (Table 1). 5. Arabian platform succession The pre-Miocene succession is now summarised to aid comparisons with the Miocene sediments that will be discussed in more detail. 5.1. Late Cretaceous and Middle Eocene — Lutetian 5.1.1. Facies description The base of the latest Cretaceous Kaleboğazı Formation (Fig. 2) generally overlies the eroded and

Soil formation in a semi-arid climate Subaerial overbank flood plain deposits or crevasse splays.

Poorly developed sub-aerial palaeosols

Shallow-marine

Sub-aqueous low energy deposition Sub-aqueous low energy deposition possibly lagoonal or peritidal Palaeosol Deep, open marine

leached top of the Maastrichtian Hatay Ophiolite. A typical section of Upper Cretaceous and Eocene limestone is exposed along Harbiye Gorge (east of the town of Harbiye; Fig. 2). The base of the Kaleboğazı Formation is composed of microbial limestone facies (Tables 2 and 6). The upper part of the formation is mudstone (Tables 3 and 6), matrix-supported conglomerate (Table 4), sandstone and lime mudstone facies (Fig. 3, Table 2) with conglomerate and coarse litharenite predominating. A slight angular discordance (∼ 7°) exists between these basal beds and the overlying Middle Eocene (Lutetian) Okçular Formation. However, this may be a local artefact of a small fault at the edge of the logged section. Palaeogene rocks are absent from this area, although Palaeogene limestone is present further south (near the town of Senköy, Fig. 2) suggesting that this area then formed a palaeohigh. The base of the Lutetian Okçular Formation is composed of thinly bedded (5–20 cm), microbial limestone facies (Table 2) with occasional lenses and horizons of red mudstone. This lithology is commonly karstified, which obscures primary sedimentary features. The interbedded lime–mudstones and mudstones pass upwards into a

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Table 4 Summary description and interpretation of Miocene conglomeratic facies of the Hatay Graben Facies name

Description

Matrix-supported conglomerate

Interpretation

Matrix-supported conglomerate, poorly sorted, sub-angular to rounded clasts. Discontinuous conglomerates Laterally discontinous matrix conglomerate with sub-angular Gmm/Gmg to rounded; generally poorly sorted clasts. Clast size varies greatly; in the lower part of the formation clasts can be N1 m in size, whereas higher in the succession the maximum clast size is b10 cm. Weak grading in some beds. Clast-supported conglomerates Gcm Massive clast-supported conglomerate with sub-angular to rounded, generally poorly sorted clasts. Conglomerates with clast Clast-supported conglomerate with pebble imbrication imbrication Gh and parallel lamination. Conglomerates with cross-bedding Gp Clast-supported conglomerate, poorly sorted with large-scale, low-angle cross-bedding. Basal conglomerate Laterally discontinuous conglomerate of variable thickness (up to ∼ 5 m), clast-supported and polymict; lithoclasts are well-rounded and occasionally bored. The clasts include limestones and serpentinite. The conglomerates commonly contain fossil material, e.g. bivalves, gastropods and coral (Fig. 9A). Red conglomerate Red, massive conglomerate, ∼ 5 m thick, poorly sorted, angular to sub-angular and clast supported. Serpentinitic conglomerate Clast-supported conglomerate beds (1–3 m thick) composed of sub-angular, to sub-rounded, serpentinite clasts up to 40 cm long. Limestone conglomerate Matrix-supported conglomerate, N4 m thick, with sub-angular to sub-rounded clasts of marl and limestone and a marl matrix (Fig. 12A).

packstone/rudstone facies (Tables 2 and 6). In places, the Eocene strata are folded into disharmonic folds and dips locally reach 90°. 5.1.2. Interpretation of the Maastrichtian and Lutetian successions The sedimentary features show that the lower part of the Maastrichtian succession accumulated in an intertidal zone. Modern carbonates can be classified in terms of the amount of exposure during a tidal cycle, the ‘exposure index’ (Ginsburg et al., 1977), based on the occurrence of common sedimentary structures. The presence of irregular algal lamination indicates that, where this is present, these sediments have an exposure index of 40–90%, corresponding to the lower intertidal zone. Palaeosols higher in the succession were deposited in a continental setting (i.e. 100% exposure) that was probably semi-arid in view of the presence of nodular caliche. Thin, generally structureless layers of ophiolite-derived clastic sediment are seen as being alluvial. During the Lutetian, an intertidal environment of deposition was maintained. Mudstone horizons are interpreted as immature soils, in view of their mottled colour and lack of fossil material; thus, the exposure index was higher during the Lutetian (Okçular Forma-

Debris-flow deposits and alluvial sands. Debris flows (viscous)

Debris flows (internal bed load turbulent) Longitudinal bedforms, lag deposits, sieve deposits. Transverse bedforms High energy erosional deposit related to rapid relative sea-level rise.

Sheet flows on an alluvial fan Sub-aqueous deposition in channels Debris flows

tion) than the Maastrichtian (Kaleboğazı Formation), reaching 100% at times. The presence of large benthic foraminifera (e.g. Nummulites) indicates that the limestones in the upper part of the succession were deposited in an open-marine setting near the shelf-edge, or on the upper slope (Sartorio and Venturini, 1988), since Nummulites thrived in water depths of 20–100 m (Saller et al., 1993). Local accumulation of oysters (possibly reworked) at one location (near Bözlü, Fig. 2) suggests that water depths were relatively shallow, to restricted at certain times and places. An overall marine transgressive setting is applicable to the Lutetian, with fluctuating and locally variable water depths, Haq et al. (1987) and Miller et al. (2005) indicate relatively constant Lutetian water depths suggesting the presence of local topographic or tectonic controls in this area. The Upper Cretaceous and Eocene limestones contain chert nodules, often orientated parallel to bedding planes. Nodular chert of diagenetic origin is commonly present within shelf limestones of the Arabian margin throughout the Middle East, including Southern Turkey, Syria, the Levant, Cyprus and also beneath the Mediterranean Sea, as indicated by ocean drilling (Robertson, 1998; Rosenfeld and Hirsch, 2005). Sponge spicules are a likely silica source in relatively shallow shelf limestones. However, diatoms and radiolaria are

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additional potential sources of the silica, related to upwelling and high productivity in a continental slope setting adjacent to the deeper Neotethys to the north. 5.2. Early Miocene — Aquitanian to Burdigalian The sequence that we relate to a foreland basin origin begins with coarse clastic facies. 5.2.1. Facies description The Middle Eocene succession, discussed above, is unconformably overlain by mainly clastic sediments, represented by the Early Miocene Balyatağı Formation (Fig. 2). The type location is on the main road from Antakya to Altınözü, near the village of Enek (Figs. 2, 5, 7). The base of the formation rests with an angular unconformity on the underlying Eocene Okçular Formation. The Balyatağı Formation is composed of interbedded conglomerates (Fig. 6A) and mudstones (conglomerates and pebbly mudstone facies; (Tables 3, 4 and 6), up to 175 m thick. The conglomerate contains ophiolitic and limestone clasts. The composition of the matrix ranges from micrite to fine terrigenous material of the same composition as the larger clasts, to pale-coloured, weathered material, mainly derived from serpentinite.

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A representative log was measured along the type section (Figs. 5–7). The strata are laterally discontinuous (Fig. 7), such that a single logged section cannot be wholly representative of the formation. However, the log shows typical features, notably thick conglomerate beds that fine upwards (from a clast size often N 30 cm), to mudstone. In the lower part of the formation, conglomerates exhibit lateral continuity with an apparently sheet-like geometry. Above, conglomerate beds in the mid to upper part of the formation are more discontinuous and lens-shaped (b 2 m thick, 10′s–100′s m in extent; Fig. 7). Pebbly sandstone and mudstone occur between these higher conglomerate beds. Near the top of the formation, sedimentary structures are more common, also less mudstone is present. Above, there is an abrupt transition to limestones of the Middle Miocene Sofular Formation. The Balyatağı Formation also outcrops ∼ 6 km southwest in Harbiye Gorge (Fig. 3), where it again unconformably overlies Eocene facies; the base of the formation is composed of bioclastic coarse pebbly sandstone facies (Table 5, Fig. 6B). This ∼ 6 m thick interval lies on an irregular surface of Eocene limestone. Above this basal interval, thick conglomerate fine upwards into pebbly mudstone (Table 3), as

Fig. 6. A) Close up of clast-supported conglomerates of the Balyatağı Formation; B) Close-up of clast with a microbial coating; from the base of the Balyatağı Formation at Harbiye; finger is pointing to the clast; C) View of the Early Miocene to Middle Miocene boundary at Kesecik; note that the Early Miocene clastics in the foreground are poorly vegetated — very pale colour in the photograph. The Middle Miocene forms the top of the hill in the centre of the field of view with the boundary marked by the line of increased vegetation; D) View of the Sofular Formation at Kozkalesi showing the lower part of the peritidal cyclothem sequence.

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Fig. 7. Panoramic photograph and field sketch of the type section at Balyatağı Tepe showing the lateral discontinuity of conglomerate beds characteristic of the central part of this formation.

observed at Balyatağı Tepe. The bed thickness is typically N2 m. The maximum clast size varies between beds but is generally b 40 cm. Clast composition is mixed, with both limestone and serpentinite. The matrix is mostly composed of cream-coloured, weathered serpentinite. The fine-grained mudstone facies at Harbiye Gorge is well lithified and varies from pink to white. It often

contains “floating” clasts up to 5 cm in size, although mostly b0.5 cm, these clasts were derived from the subjacent ophiolite. In places, these mudstones are nodular, with abundant caliche. Eight of the sections studied exhibit palaeocurrent indicators (Boulton et al., 2006). These are mostly imbricated clasts and occasional cross-bedding. Locations in the northeast exhibit palaeocurrent directions to the

Table 5 Summary description and interpretation of Miocene sandstone facies of the Hatay Graben Facies name

Description

Interpretation

Coarse pebbly sandstone Bioclastic calcarenites Pebbly sandstone

Coarse pebbly sandstone containing shallow-marine debris, including bivalves (e.g. Ostrea sp.), oncolites and pebbles with microbial coatings. Fine- to coarse-grained, rich in bioclastic material (e.g. gastropods, bivalves, echinoids and benthic foraminifera). Sedimentary structures are rare apart from planar laminations. Bed thickness is variable (b1.5 m) and some beds are lenticular. Clasts are sub-angular to subrounded, b20 cm in size, and dominated by ophiolitic rocks, although reworked micrite and conglomerate clasts are also present. The coarse, to pebbly, sandstone beds are cross-bedded. Variable thicknesses (10 cm–2 m); unfossiliferous, but exhibits sedimentary structures, e.g. parallel lamination, mud intraclasts and flute casts. Small iron-stained nodules are present within the underlying or overlying marl at some horizons. Thin, b50 cm thick, calcarenite beds, typically packages interbedded with marl. These beds have sharp bases and fine upwards; sedimentary structures include parallel lamination, cross-lamination, ripple marks, flute casts and marl intraclasts associated with small-scale slumped horizons. Widely exposed in the northerly part of the basin as thick litharenite. Three massive beds are separated by thin marl or chalk and are rich in marl intraclasts in specific horizons. Other sedimentary structures are not present although oyster shells are occasionally seen.

Shallow-marine

Calcarenite

Calcarenite 2

Litharenite

Shallow-marine High-energy beach and shoals between reefs. Distal turbidites

Channelised sandbodies

Grain flows

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Fig. 8. Sedimentary log of the type section of the Sofular Formation near the village of Kozkalesi — log 3 on Fig. 2.

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Table 6 Summary of the microfacies identified through microscopy of the sediments of the Hatay Graben Formation

Brief descriptor

Upper Cretaceous Kaleboğazı Fm.

Ferruginous palaeosol

Microfacies description

Rounded, opaque pellets possibly iron-rich pesoids; leaching from these clasts has stained the micrite cement. Textures vary across the sections; the concentration of clasts varies and an area of wavy laminations, possibly algal in origin, separates areas of high and low concentration. In addition to the opaque material there are microcrystalline quartz grains (possibly chert) and rarer mono- and polycrystalline quartz clasts. Upper Cretaceous Algal limestone with The matrix is composed of dense micrite and the pore spaces have a minor infill of sparite or possibly Kaleboğazı Fm. fenestral porosity dolomite (crystals have a distinctive diamond shape). No siliciclastic material; some bioclastic material may have subsequently been replaced. Contains planktic foraminifera and ostracod fragments. Upper Cretaceous Cream packstone Rich in bioclasts: algal material, including red coralline algae, coral, gastropod, bivalves (oysters), Kaleboğazı Fm. echinoid plates, crinoid ossicles, planktic forams, benthic forams, bryozoa, calcispheres, rare ostracod fragments. Rare silt and limestone clasts. Micrite matrix with sparite infilling porosity. Eocene Okçular Fm. White grainstone/ Rich in bioclasts: benthic foraminifera (Nummulites, Miliolidae, Discocyclinidae), planktic foraminifera, packstone echinoid spines, bivalves, coralline algae, bryozoa. Also contains clasts of older limestone containing bioclastic material (gastropods, algae, foraminifera). Drusy sparite cement with some micrite, sparite overgrowths and poikilotropic textures. Eocene Okçular Fm. Wackestone Composed of very dense micrite rich in bioclastic material: planktic forams (Morozovella sp, Globergerinidae, Globorotalidae), benthic foraminifera (including fragments of Discocyclina), bivalve fragments, ostracods, calcispheres. More planktic foraminifera present than benthic. Some siliciclastic material is mostly serpentinite (sub-angular) and monocrystalline quartz. Heamatite infills some of bioclastic material. Lower Miocene Conglomerate Micritic groundmass has a blotchy appearance suggesting that this could be partly microbial, peloidal Balyatağı Fm. or reworked in some way. Poorly sorted sub-angular to sub-rounded clasts, composed of serpentinite, partly recrystallised limestone, intramicrite, quartz, plagioclase, opaques, altered basic volcanic glass (palagonite). There is no bioclastic material. Lower Miocene Brown greywacke Maximum clasts size up to 7–8 mm; poorly sorted and angular to sub-rounded. Very fine micrite Balyatağı Fm. matrix contains some siliciclastic material; Calcite grains ∼ 40%, serpentinite, quartz, feldspar, opaque minerals ∼ 10%, iron-rich pellets. Contains no bioclastic material. Middle Miocene Wackestone Contains peloids, bioclasts (benthic and planktic foraminifera, coralline algae and encrusting algae, Sofular Fm. gastropods, bivalves) and siliciclasts (olivine, quartz, serpentinite, chert, plagioclase). Porous, although some porosity has been infilled by secondary sparite. Middle Miocene Wackestone Micrite matrix with isopachous fringes developed around some grains. Porosity variable, some infilled Sofular Fm. by sparry calcite cement. Bioclasts; coral, planktic and benthic foraminifera, including Operculina. bivalves including Ostrea sp., algal material (reworked), echinoid fragments, ostracods, bryozoa, coral. Clasts of micrite. Middle Miocene Packstone Grainsize 1–2 mm in a micrite matrix; the rock has a high porosity partly infilled by sparite, forming Sofular Fm. radial fringes on the grains. Clasts are well-rounded and bioclasts are fragmented. Bioclasts: red algae (coralline, possibly lithothamnium) forms the majority of grains, bivalves, Ostrea, benthic (e.g. Alveolina sp.) and planktic foraminifera (e.g. Globigerina sp.), echinoid fragments (with optically continuous overgrowths). Lithic fragments: quartz-rich silt, serpentinite, chloritised basalt, bioclastic limestone rich in planktic foraminifera. Grains: monocrystalline quartz (medium-sized fluid inclusions), olivine, peloids, intraclasts. Middle Miocene Mudstone/ White with a few opaque grains. Composed almost entirely of very fine micritic cement (N80%); Sofular Fm. wackestone contains some siliciclastic material including: detrital quartz grains (showing some evidence of subgrain rotation (1–5%)), clinopyroxene, epidote, chlorite, serpentinite. Rare calcite crystals are present and a small bioclastic component (mainly planktic foraminifera). Most of the rock is micrite, possibly of peloidal original. Sparite is also present, infilling pore spaces (coarsening into the centre), sparry calcite occurs around the clasts and is noticeably coarser around them. Middle Miocene Litharenite Grainsize ∼1 mm; clast-supported; grains sub-angular to sub-rounded and poorly supported. Little Sofular Fm. micrite; carbonate cement b15%. Calcite cement overgrowths on some grains. Clasts: Serpentinite (N70%), micritic limestone, siltstone. Grains: reworked foraminifera, monocrystalline quartz, feldspars, chlorite, calcite, bivalves. Lithic fragments N98%/Quartz b1%/Feldspar b1%. Middle Miocene Matrix supported Grainsize 1–3 mm, well-rounded and moderately well-sorted. Matrix-supported. Micritic matrix Sofular Formation litharenite composes ∼ 10% of the rock, with sparite infilling porosity and forming overgrowths on bioclastic material. Clasts: peloids/micrite, biomicrite limestone, serpentinite, ferruginous chert, rare quartz. Bioclasts (∼ 90%); coralline algae (probably reworked), echinoid fragments (with optically continuous sparite overgrowths), benthic and planktic foraminifera, bivalve fragments.

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Table 6 (continued ) Formation

Brief descriptor

Microfacies description

Middle Miocene Sofular Fm.

Grainstone

Middle Miocene Sofular Fm.

Calcilitharenite

Middle Miocene Sofular Fm.

Mudstone

Upper Miocene Nurzeytin Fm.

Litharenite

Upper Miocene Nurzeytin Fm. Upper Miocene Nurzeytin Fm.

Mudstone

Upper Miocene Nurzeytin Fm.

Lithicwacke

Bioclasts: algal material (50–60%), benthic foraminifera (20%), echinoid fragments (5%) with optically continuous overgrowths. Micrite clasts (peloids?) The rock is totally reworked; darker materials are remains of reworked microbial mats; these are well rounded. Coarse-grained orange sandstone, poorly sorted with subangular to rounded clasts. N75% of clasts are serpentinite; micritic matrix. Bioclasts: Ostrea, benthic foraminifera. Reworked clasts: coralline algae, benthic and planktic foraminifera. Original porosity ∼ 5%, now reduced to ∼ 2%, sparite in pore spaces; micrite envelopes around clasts. Brown, very fine-grained but contains some rare clasts b0.5 cm; very poorly sorted. Micrite (N75%) with subangular clasts. Clasts are generally very small including serpentinite, quartz, basalt. Massive, no microfossils. Grainsize ∼ 0.5 mm; poorly sorted; grains angular to sub-rounded; clast-supported. Matrix composed of very fine-grained black siliciclastic material. Clasts: serpentinite, micrite/peloids, limestone (micrite and sparite), chert (inc. radiolarian), basalt (laths of plagioclase), siltstone, calcite (peloids), feldspar. Grains: monocrystalline quartz (some contain thin trails of fluid inclusions and occasional undulose extinction), polycrystalline quartz, biotite and muscovite, plagioclase (altering to sericite), microcline, calcite crystals, chlorite, opaques. Some samples contain bioclasts e.g., benthic and planktic foraminifera; bivalve fragments. Lithic fragments 75–85% /Quartz 10–20% /Feldspar ∼5%. Composed mostly of micrite with grains of planktic and rare benthic foraminifera; ostracods; small quartz crystals, muscovite, small hematite grains. Diamond-shaped crystals likely to be dolomite. Graded, grainsize b1 mm, micrite matrix with some sparite infilling porosity in bioclasts. Bioclasts (∼90%); echinoid plates (with optically continuous calcite overgrowths), planktic and benthic foraminifera, bivalve fragments, gastropods, ostracods, bryozoa. Lithoclasts are present in some samples, e.g. micrite, cherty mudstone, siltstone, limestone. Very fine-grained sandstone; angular to sub-rounded grains; poorly sorted with a calcite matrix N15%. Clasts of bioclastic limestone, serpentinite, micritic limestone. Grains of plagioclase, monocrystalline quartz (with undulose extinction, fluid inclusions and subgrain rotation), polycrystalline quartz, biotite and muscovite, hypersthene, opaques, chlorite, glauconite, bioclasts (foraminifera).

Packstone

north or northeast, whereas by contrast, locations in the southwest exhibit a greater variety of orientations, from south, west/northwest, to east. Measurements were not made on the northwestern basin margin as palaeocurrent indicators were not observed there. 5.2.2. Interpretation of the Lower Miocene succession The basal unconformity, of probable Oligocene age, represents a regional feature separating the folded MidEocene limestone from undeformed Lower Miocene conglomerates. The sediments of the Balyatağı Formation are fluvial in origin. Such fluvial conglomerates are commonly lenticular, with poorly developed cross-bedding and sandstones exhibiting lenticular and laterally continuous bed morphologies (Tucker, 1991). Palaeosols are common in fluvial sequences under appropriate climatic conditions (Collinson, 1996). Using a standard nomenclature for fluvial sediments (Miall, 1978, 1985, 1996) the conglomerates low in the succession can be identified as facies Gmm, Gcm and Gmg, typical of sediment gravity flows (i.e. facies association SG). These sediments form sheet-like conglomerates, resulting from infrequent catastrophic flows (Collinson, 1996).

The conglomerates and sandstones high in the succession are identified as facies Gcm, Gh, Gp, Sm, indicative of deposition from stream flows (Collinson, 1996). These facies suggest deposition in channels (facies association CH), or from gravel bars (facies association GB); these are commonly associated with sediment gravity-flow deposits (SG). The lenticular shapes of many of the gravel deposits are also suggestive of channel forms. Bedding planes become better defined towards the top of the formation; grainsize also decreases. Crossbedding, parallel lamination and pebble-imbrication are all present, as in facies Gh, Gp, Sl, Ss, St. These facies are suggestive of deposition on laterally accreting macroforms (LA) or downstream-accreting macroforms (DA) that are characterised by internal lateral accretion (i.e. 3rd-order surfaces of Miall, 1996). Palaeosol horizons (P) are interbedded with the conglomerates and sandstone near the top of the formation. Some braided river systems lack significant palaeosols, but some ancient systems, such as the Escanilla Formation in the Spanish Pyrenees, contain N40% palaeosols (Bentham et al., 1993); these are interpreted as overbank deposits that were mainly deposited from unconfined flows during times of overbank flooding.

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High proportions of poorly developed palaeosol have been interpreted to indicate high rates of subsidence and sediment accumulation (Kraus and Middleton, 1987). The Balyatağı Formation is, therefore, interpreted as a gravel-dominated river with the characteristics of a braided system. The basal sediments of the Balyatağı Formation are dominated by sediment gravity flows, with gravel bars and channels becoming more common upwards. Sediment gravity flows are characteristic of gravel braided rivers (Miall, 1996); sinuosity was low but braiding was high. Gravel bar and sandy bar elements increase upwards. Although also found in gravel braided rivers with sediment gravity flows, gravel bars and subordinate sandy bars are characteristic of shallow gravel braided Scott-type rivers (Miall, 1978, 1996). These consist of horizontally bedded gravels filling stacked, shallow channels, suggesting an evolution in river type over time. By contrast, where the sediment load is a mixture of sand and conglomerate, a variety of bed forms can develop, as seen in the uppermost part of the formation. The presence of gravel bar and sandy bar elements with lateral and downstream accreting elements, and the presence of palaeosols, suggest deposition from a deep gravel, braided, Donjek-type river (Miall, 1978, 1996). Floodplain deposits are more likely to be preserved due to the greater topography generated by these mixedsediment rivers compared to gravel-dominated ones. Sinuosity in this type of river is low to intermediate and braiding is intermediate to high. We, therefore, infer that there was a progression in river type over time, from a braided-river dominated by sediment gravity flows to a shallow, then to a deep gravel-bed braided river. This progression may variously reflect a decrease in the size of the sediment bedload and a corresponding increase in stream power, a more uniform discharge, a change in basin topography, or greater incision of the river network. The overall control could be change in base-level (tectonic) or climatic change. A possible source of the Lower Miocene ophiolitederived sediments was the ophiolitic Baer–Bassit Massif exposed in northernmost Syria and the border with Turkey (e.g. Al-Riyami et al., 2002), which is consistent with palaeocurrent data (Boulton et al., 2006). The Hatay and Baer–Bassit ophiolites were emplaced as a regionally extensive thrust sheet during the latest Cretaceous (Robertson, 2002). Thus, ophiolitic detritus may have been derived from additional ophiolitic outcrops that then existed between the Hatay and Baer–Bassit areas. Ophiolites are not now exposed between these two areas as they are covered by younger sediments, or eroded.

5.3. Middle Miocene — Burdigalian/Langhian The succession passes transitionally upwards into a marly carbonate succession of Early to Middle Miocene age (Fig. 6C), known as the Sofular Formation, for which the type section is located near the village of Kozkalesi (Fig. 2). The base of the succession in this area is composed of repeated, cyclic units of marine carbonates, N200 m thick (Figs. 6D, 8). The main facies are a basal conglomerate, wackestone–rudstone, wackestone–packstone and a pink mudstone (Tables 2–4 and 6). 5.3.1. Interpretation of the type section Approximately 25 discrete depositional cycles are present, which we relate to rise and fall in relative sealevel. The basal conglomerate (Table 4; Fig. 9A) formed in response to a rapid sea-level rise, initiating carbonate deposition under shallow-marine conditions. The lower limestone beds (wackestone–rudstone; Tables 2 and 6) above the basal conglomerate horizon contain abundant reworked shallow-marine fauna but lack sedimentary structures; and are characteristic of sub-tidal sediments. Above this, the wackestone–packstone facies contains a similar faunal assemblage suggesting continuing shallow-marine conditions. Rare, near-basal, oysters and coral heads (that appear to be in situ) are likely to record slightly deeper-water conditions; i.e. a maximum flooding event. Later, the environment regressed to subaerial, as indicated by the presence of palaeosols (i.e. pink mudstone facies; Tables 3 and 6). Palaeosol development was, in turn, halted by a renewed transgression that initiated another cycle of carbonate deposition. The repeated deepening-upward, then shallowingupward carbonate cycles are interpreted as peritidal cyclothems within the interior of a carbonate platform (e.g. Ginsburg, 1975; Pratt et al., 1992). Their formation may be due to autocyclic controls such as progradation of the shoreline (e.g. Calcare Massiccio Formation, Apennines; Coliacicchi et al., 1975) to create the necessary accommodation space, or allocyclic controls such as tectonics, or eustatic sea-level change (Wright and Burchette, 1996). Comparable cyclicity in the Late Burdigalian to Langhian of the Mut Basin of centralsouthern Turkey (northwest of the Hatay area) is explained mainly in terms of shorter and longer periodicity eustatic sea-level fluctuation, controlled by polar ice volume changes (Bassant et al., 2005). A similar eustatic sea-level influence is likely for the Hatay area. Bio-accumulations, palaeosols and conglomerate horizons decrease in abundance, until eventually giving way to constant shallow-marine deposition (i.e.

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Fig. 9. A) Close up of a conglomerate lag marking the base of a clyclothem from the Koskalesi section of the Sofular Formation; note the presence of bored pebbles and various bioclasts; tape measure is 6 cm wide; B) Fine-grained basal mudstones of the Sofular Formation along the Asi River section; C) Lower part of the carbonate mudstone succession with two prominent red palaeosol horizons; GPS unit is 10 cm long; D) Bored upper surface of the carbonate mudstone facies.

oncolitic wackestone/rudstone) in the upper part of the succession. The oncolites and other shallow-marine fossil material are reworked. Benthic:planktic foraminiferal ratios (Boulton et al., 2006) suggest a water depth of up to 200 m (Meschede et al., 2002). This, in turn, implies overall upward-deepening of the basin. Normal faults are known to have existed during deposition of these sediments (Boulton et al., 2006), indicating that the carbonate slope and platform were influenced by active tectonics. 5.3.2. Comparative successions Additional facies exposed elsewhere add considerably to our interpretation of the Mid-Miocene palaoenvironments. 5.3.2.1. Non-cyclic shallow-water facies. Limestones exposed on the coast, near Çevlik (Fig. 2), closely resemble the upper part of the type succession at Kozkalesi (wackestone–rudstone, Tables 3 and 6), but a cyclic sequence is absent beneath this. A cliff exposure shows large-scale erosional surfaces (Fig. 10). In addition, the limestone is faulted such that no single complete sequence is exposed. However, good exposures of the base and top of the limestone sequence can be observed locally. The base of the formation is partially exposed on the coast, where coarse bioclastic

wackestone overlies serpentinite; however, the actual contact with the ophiolite is not exposed. Beds higher in the succession in the vicinity of this coastal area comprise bioclastic wackestone interbedded with laminated and rippled marl (Table 3). 5.3.3. Interpretation: reworked shallow-marine carbonate sediment The facies in the coastal exposures consists of shallow-marine carbonate; i.e. bivalves, gastropods, echinoids and microbial carbonate, material that was later reworked offshore. Outcrop-scale erosional features (Fig. 10) are interpreted as slump scars. Similar but larger scale features have been described, for example from an upper slope setting of the margin of a Jurassic carbonate platform (Sumeini Group) in Oman. These erosional features were termed intraformational truncation surfaces (Watts and Garrison, 1986). The presently studied limestones are interpreted as deposits on a carbonate slope adjacent to a carbonate platform (e.g. at Kozkalesi). Any slumped material was carried downslope and is not now exposed. The carbonates of the upper part of the Middle Miocene Sofular Formation represent a deeper-water facies than the underlying cyclothem facies, where present. A relative sea-level rise caused the shallow-water carbonate platform to drown and presumably move landwards.

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Fig. 10. Panoramic view of the exposure of the Sofular Formation near Çevlik along a relict sea cliff (Map reference Mersin-P35-c3 0764250/ 4001300). The cliff is ∼50 m high and the slide scars cut 10–20 m down into the underlying stratigraphy; the direction of these scars is not clear due to the orientation of the exposure.

5.3.3.1. Variable facies. Facies variation is exemplified by three locations on the Asi River (Fig. 2). On the north side of the river the Middle Miocene sequence directly overlies serpentinite. The base of this sequence is associated with a red, massive conglomerate (Table 4), ∼ 5 m thick. The clasts are 100% serpentinite. These coarse clastic sediments are overlain by several metres of very fine-grained mudstone and lime mudstone (sandy mudstone; Table 3). This facies is overlain by fine-grained microbial mudstones (Table 2), forming a more than 40 m-thick sequence. Near the base of the limestone sequence three horizons of laminated mudstone (Table 3) are present (Fig. 9B). Elsewhere the red conglomerate facies again comprise the basal sediments overlying the ophiolite, followed by ∼ 8 m of fine-grained sediments (sandy and interbedded mudstone; Table 3), (Fig. 9C). These sediments pass upwards into fine-grained microbial mudstone (Table 2). The upper part of the succession, ∼ 100 m thick, is exposed further north (Fig. 11), as microbial mudstone (Table 2), with algal laminations, rare small gastropods and fenestral porosity. The upper bedding surface of this particular lithology is bored (Fig. 9D); there is also an angular discordance with the overlying beds (a few degrees). Above this disconformity the sediments exhibit an increased clastic component. Microbial mudstone

and calcarenites are interbedded with serpentinite conglomerate (Table 4) and occasional thin beds of mudstone and carbonate (Fig. 12A). The lower part of this sequence is variable; whereas lithologies become more uniform upwards. 5.3.4. Interpretation The structureless nature and the presence of immature, angular clasts suggest that the basal red conglomerates, exposed near the Asi River, were derived, eroded and then deposited as an alluvial fan. The overlying thin-bedded sediments show evidence of shallow-water deposition with periodical sub-aerial exposure resulting in desiccation cracks and immature palaeosols. The close association of these two facies suggests the existence of a fan delta (Nemec and Steel, 1988). The sedimentary structures in the overlying thinbedded limestones point to a shallow-marine, possibly lagoonal or peritidal setting. Any lagoonal-type facies could be laterally equivalent to the peritidal carbonate sequence low in the type section at Kozkalesi. However, such very shallow-marine sediments are not present further west in the coastal exposures, near Çevlik. This could be due to deeper-water conditions in this area at the time of deposition, to erosion, or simply to a lack of exposure.

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Fig. 11. Sedimentary log of the upper part of the Sofular Formation along the Asi River (Map reference Antakya-P36-d4 0235097/ 3998128); log 4 on Fig. 2.

A gradual decrease in stratal dip, as observed northeastwards along the Asi River, has been interpreted to reflect tilting related to syn-sedimentary normal faulting (Boulton et al., 2006). As a result, a local lowangle discordance (described above) could reflect such a tectonic disturbance. The presence of borings on the upper surface of the fine-grained limestone suggests a period of very low net sediment accumulation or nondeposition within the sequence. The diversity of biological material, mostly fragmentary and reworked, in the limestone above this discordance suggests a return to more normal rates of marine accumulation. In addition, the interbedded conglomerates in the upper part of the succession represent channel-fill deposits, possibly equivalent to the slope facies in the type section at Kozkalesi.

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5.3.5. Patch reefs Moving north-eastwards away from the coast, the thickness of the Sofular Formation decreases to b10 m at many localities. Specifically, the Sofular Formation decreases from ∼ 150 m at Çevlik, to b20 m near the village of Kesecik (Fig. 2). Where present, a thin limestone (i.e. wackestone–rudstone/wackestone– packstone) contains articulated bivalves, suggesting that significant transport of this bioclastic material has not taken place. The base of the formation, near Kesecik (Figs. 1 and 6C), is conformable on the underlying Lower Miocene succession and is composed of reefal wackestone (Table 2). Stratigraphically above this is a reef rudstone (Table 2), overlain by reefal packstone. Along strike (∼ 250 m) to the southwest (at 0238686/ 4015181), reddish-brown, pebbly sandstone and conglomerate beds overlie the Lower Miocene succession.. The top of this exposure is reef rudstone, ∼ 1.5 m thick, dominantly composed of Porites coral fragments. The coral, Porites, is known for tolerance to finegrained clastic input (Hubbard and Pocock, 1972), as seen in the Miocene of SW Turkey (Hayward et al., 1996). The disappearance of coral upwards suggests that the environment became unsuitable for coral growth. Coral reefs with low species diversity, where Porites dominates, occur elsewhere around the Eastern Mediterranean basin, including the Late Miocene Koronia Member, of the Pakhna Formation Cyprus (Follows, 1992; Follows et al., 1996). Associated sandstones can be compared with the ‘Marginal Terrigenous Complex’ (Esteban, 1996), an association of coral reef and siliciclastic facies present in a number of Mediterranean areas, including Sicily (Esteban, 1996) and Tuscany (Bossio et al., 1978). Only one intact patch-reef was observed in the field; however, coralline rubble and coral fragments probably representing fore-reef talus, were observed at several locations. This suggests that coral build-ups were once more commonly developed, but underwent penecontemporaneous erosion and redeposition in a relatively high-energy, shallow-marine setting. Similar, reef talus is common in the Miocene Pakhna Formation of southern Cyprus, where in situ reefs are not preserved now either (Eaton and Robertson, 1993). Coral reefs were common in some parts of the Mediterranean during the Langhian (Esteban, 1996), apparently related to 2nd order eustatic sea-level highstands (Haq et al., 1987). Other major times of coral reef formation occurred during the Aquitanian and the Late Tortonian/Messinian, as in Cyprus (Follows et al., 1996) and the Isparta Angle, SW Turkey (Flecker

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Fig. 12. A) Photograph showing a clastic horizon observed directly above the bored surface along the Asi River section; Middle Miocene in age; B); Thick conglomerate horizon in the Nurzeytin Formation; clasts up to 2 m; C) Thick sandstone interbeds in the Nurzeytin Formation; D) Blocks of alabastrine gypsum in a gypsiferous marl, near Nurzeytin.

et al., 2005); these reefs also correlate with 2nd order highstands. Elsewhere, Langhian carbonate platforms with numerous coral reefs, generally composed of Porites, Tarbellastrae and Stylophora, are found in the Red Sea (Coniglio et al., 1996; Purser et al., 1996), Catalonia, Spain (Permanyer and Esteban, 1973; Alverez et al., 1977) and Israel (Buchbinder, 1996). Small fringing reefs with a variety of morphologies predominate, whereas true barrier reef systems were uncommon (Esteban, 1996). 5.4. Mid–Late Miocene — Serravallian to Tortonian 5.4.1. Facies description The succession passes transitionally upwards into a thick (∼ 300 m) unit of foraminiferal marl (Fig. 13; Tables 3 and 6), with interbedded calcarenites (Fig. 12B, C; Tables 5 and 6), limestone and limestone conglomerates (Table 4), known as the Nurzeytin Formation. The upper levels of this formation are formed of the Messinian, gypsiferous Vakıflı Member (Fig. 2, 12D).

5.4.2. Interpretation The predominant marl facies is interpreted as a relatively deep-marine hemipelagic facies that accumulated from suspension. The ratio of benthic to planktic foraminifera suggests a water depth of up to 700 m (Boulton et al., 2006). A single conglomerate bed is exposed at several locations suggesting lateral continuity (Fig. 12C). Its matrix-supported fabric suggests that the conglomerate formed by a mass-flow process. The characteristics of the calcarenite facies (Tables 5 and 6), including parallel and cross-lamination, suggest deposition from turbidity currents, although complete Bouma sequences (Bouma, 1962) are absent. The second calcarenite facies (Table 5) also has sedimentary structures indicative of turbidites. Beds are locally associated with small-scale slumped horizons, indicating sediment instability. Litharenite beds may either represent gravity-driven grain flows, or the TA intervals of turbidites. Thick sandstone turbidites, of similar composition to those described above, occur as up to metre-scale beds and

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Fig. 13. Sedimentary log of the type section of the Nurzeytin Formation; B is a more detailed log of the lower 120 m of log A, which represents the whole formation. — log 5 on Fig. 2.

series of beds, in which the base of the beds exhibit flute and groove casts. Clear palaeocurrent indicators are rare (Fig. 14). However, measurements of mainly flutes and groove casts at four locations indicate variable current directions (Boulton et al., 2006). Northerly exposures exhibit palaeocurrents flowing to the south and the west, whereas further south an easterly flow is indicated. In summary, the lithologically variable interbeds within the background hemipelagic marl facies are indicative of sediment being reworked down-slope to the basin-floor, predominantly by turbidites of variable grainsize, composition, density and sediment volume. 5.5. Messinian At five localities, the Mid–Upper Miocene facies exhibit an upward transition to evaporites of inferred Messinian age, known as the Vakıflı Member. The type

section is composed of disturbed laminated alabastrine gypsum set in a gypsiferous marl matrix (Fig. 15, location a; Fig. 12D). The fine-grained gypsum is diagenetically altered to selenitic gypsum in places. At the other locations the gypsum is composed of selenite. At one locality (Fig. 15; location e), a primary layering of small selenite crystals was observed that could be equivalent to the banded-stacked selenite described from Cyprus (Robertson et al., 1995). However, generally there is no such organisation to the selenite crystals and it is unclear whether they represent primary (reworked) gypsum, or secondary (diagenetic) gypsum. 5.5.1. Interpretation Gypsum formed when the basin became semiisolated from the Mediterranean Sea due to a fall in sea-level during the Messinian (Hsü, 1972; Hsü et al., 1978). The Hatay area evolved into a small semiisolated marginal basin. The fine-grained albastrine

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Fig. 14. Map of the field area with the Nurzeytin Formation highlighted in grey showing palaeocurrent data measured at four locations.

gypsum partially formed by precipitation at the sediment–water or air–water interface during the initial stages of desiccation (e.g. Schreiber et al., 1976). This material was then reworked towards basin depocentres, or was deposited directly in these areas (Fig. 15). Minor amounts of carbonate gypsiferous marl and evaporite accumulated later in the salinity crisis. In the Hatay area the alabastrine gypsum that directly overlies the relatively deep-water Upper Miocene marls could have formed in a relatively deep-water setting (tens of metres), as inferred for the lower evaporites in the Polemi and Maroni basins in Cyprus (Robertson et al., 1995) and of the Veno del Gesso Basin in the Apennines (Manzi et al., 2005). By contrast, the overlying selenitic gypsum (where primary; e.g. Fig. 15, locality e) formed in a very shallow-water, lagoonaltype setting. Broken selenite crystals at the top of the succession may represent gypsum debris flows, possibly triggered by tectonic activity, as described from Cyprus (Robertson et al., 1995).

6. Discussion 6.1. Summary of basin development Key facts necessary for an assessment of the tectonic setting and evolution of the Miocene basin in the Hatay basin are summarised below. The oldest sediments in the area date from the uppermost Cretaceous (Kaleboğazı Formation), and were deposited after the southward emplacement of the Hatay Ophiolite, onto the Arabian platform. These sediments, east of Antakya (Fig. 2), are composed of shallowmarine (intertidal) carbonates, overlain by palaeosols and occasional conglomerates this indicates that this area was emergent in the latest Cretaceous/Palaeocene. A regression then occurred due to sea-level fall or tectonic uplift. Marine conditions returned during the Eocene (Okçular Formation). A transgressive sequence of shallowmarine, intertidal carbonates, similar to the Upper

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Fig. 15. Outline map of the Hatay area adapted from Fig. 2, showing the locations of Messinian evaporites in the area. The shaded area represents the possible extent of the evaporite basin during the Messinian.

Cretaceous limestones, passes upwards into deeper marine Nummulites-rich shelf limestones. The lack of detrital material in the limestones shows that the underlying ophiolite lacked any significant topographic relief during this time, and as a result shelf limestones accumulated over the whole area (Fig. 16). The Eocene succession is folded and is marked by an angular unconformity at the top of the succession. The upper surface of the limestone is often karstified. No Oligocene sediments were identified in the field area. This, combined with the fact that the overlying Early Miocene sediments are unfolded, implies that the tectonic folding took place during the Late Eocene/ Oligocene. The area was probably emergent and eroding during much of Oligocene time. Sedimentation recommenced during the Aquitanian– Burdigalian (Early Miocene Balyatağı Formation) with a thin, localised shallow-marine pebbly sandstone

horizon exposed near Harbiye. This represents a brief marine incursion followed by a considerable thickness (∼ 175 m) of braided-river deposits that are only present in the north of the area (Fig. 16). These sediments are rich in ophiolitic material and exhibit palaeocurrent evidence of sediment transport towards the north, from an uplifted area to the south that probably included the Baer–Bassit Massif in southernmost Turkey/northernmost Syria. The Middle Miocene sediments (Sofular Formation) are varied. Shallow-marine limestones dominate, indicating an overall relative sea-level rise since the Early Miocene. The sediment thickness is greatest in the south/southwest, suggesting that more accommodation space existed in this area. In the south the limestones directly overlie ophiolitic rocks, whereas in the northeast the limestones overlie Early Miocene and older sediments. The southwest was mainly an area of non-

114 S.J. Boulton, A.H.F. Robertson / Sedimentary Geology 198 (2007) 93–124 Fig. 16. Simplified block-diagrams illustrating the inferred sedimentary-tectonic evolution of the Hatay area in its regional context; from A, Latest Cretaceous to D, Late Miocene. See the text for explanation.

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deposition and erosion during Oligocene–Early Miocene times, whereas sedimentation was ongoing in the northeast (Fig. 16). Rising sea-levels then submerged the entire area during most of Middle Miocene time. Similar deepening upwards trend can be seen elsewhere in the Mediterranean, for example in Italy (Ricci-Lucci, 1986; Boccaletti et al., 1990) and southern Turkey (Flecker et al., 2005; Bassant et al., 2005). In the southwest, where the sediments are the thickest, there was a progression from peritidal facies, to lagoonal, then to shelf limestones deposited along a carbonate ramp during the Mid-Miocene. Progressive deepening appears to have taken place towards the northwest. This suggests that the overall basin topography was similar to that of the Eocene. The peritidal carbonates are composed of repeated transgressive cycles that probably reflect short-periodicity eustatic sea-level fluctuations superimposed on a slow, but continuous, subsidence (Miller et al., 2005). The presence of locally developed hardgrounds indicates that sediment accumulation rates were low, at times. Reefal material and rare intact patch reefs are more common in the northeast where the environment was better suited to coral growth (Fig. 16). By contrast, the southeast of the area was characterised by greater water depths, lower energy environments and less substrate topography and was unsuited to coral growth. The Middle Miocene sediments contain some ophiolitic material as well as limestone clasts, showing that both of the underlying limestone formations (Kaleboğazı and Oçkular formations) and the ophiolite were being eroded. Most of the Hatay basin area was submerged during the Mid-Miocene and so the clastic material was probably eroded from the hinterland to the south. The Upper Miocene hemipelagic marls (Nurzeytin Formation) are concordant with the underlying limestone but, in places, unconformably overlie older facies. Planktic/benthic foraminiferal ratios suggest that the sea deepened to the north (Boulton et al., 2006). Eustatic sea-level was falling throughout the Late Miocene (Haq et al., 1987; Miller et al., 2005); therefore, it is likely that subsidence was tectonically controlled. The large thickness of marls (∼300 m) indicates that the Hatay area experienced relatively deep-water conditions (N200 m) from ∼ 9.5 Ma to the Messinian (7.25– 5.33 Ma; Gradstein et al., 2004). Gravity reworking of sediments reflects basin-slope topography during this time (Fig. 16). The marl and reworked sediments are mixed, suggesting continuing erosion of both ophiolitic and sedimentary rocks. Both subaerial and shallowmarine areas must have existed to supply the clastic material to the basin.

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White mica first appears in the Upper Miocene marls of the Nurzeytin Formation (Boulton et al., 2006). There are no micaceous rocks in the exposed ophiolitic basement of the Hatay area; therefore, this material must have been sourced from outside the local basin. Muscovite appears earliest within Lower Miocene sediments in the Kırıkhan area to the northeast to the area discussed in this paper (Fig. 1), along with a high concentration of quartz, suggesting that this area was close to the source. A potential source area was Palaeozoic sandstones within the Arabian foreland succession which are today exposed, for example, in the Amanos Mountains to the north (Parlak et al., 1998). An alternative source is the Tauride thrust front further north. However, sandstones derived from the regional Tauride allochthon are more lithoclastic and commonly contain ophiolite-derived material and radiolarian chert (Gökçen et al., 1988; Robertson et al., 2004). The evaporites are inferred to have accumulated near the depocentre of the basin in the Hatay area. These evaporite outcrops are now at quite different elevations. This is consistent with a marked changed in tectonic regime, associated with differential uplift, following the end of the Miocene (Boulton et al., 2006). 6.2. Peripheral foreland basins We now compare the Miocene sediments of the Hatay region with well-documented foreland basins elsewhere, specifically “underfilled” peripheral foreland basins. 6.2.1. Underfilled foreland basin trinity Peripheral forelands can be broadly categorised as ‘underfilled’ or ‘filled’ (Tankard, 1986). Underfilled forelands are composed of dominantly deep-marine sediments, as opposed to filled foreland basins that are dominated by shallow-marine, to continental sediments. Underfilling is characteristic of the early stages of foreland basin evolution when high rates of subsidence, due to tectonic loading, create accommodation space that for a time remains unfilled (Allen et al., 1991). The sedimentary fill of such underfilled basins forms a classic sequence that was termed the “underfilled trinity” by Sinclair (1997). These three units are diachronous and are superimposed on one another due to the migration of facies away from the orogenic front. These units generally overlie the basement and passive margin sediments above an unconformity (e.g. Allen et al., 1991; Coakley and Watts, 1991; Sinclair, 1997). The “trinity” represents sedimentation in three different areas of an “underfilled” peripheral foreland basin. The lower unit reflects shallow-marine sedimentation on

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the inboard (cratonic) margin of the basin, unconformably above a passive margin succession. The basal deposits record the initiation of subsidence at any given location. In the Alpine foreland basin, for example, the basal foreland sediments were deposited above a regionally extensive basal unconformity. This unconformity resulted from a combination of flexural uplift and normal faulting related to the formation of the Rhine Graben (Sinclair et al., 1998). The middle unit records the period when relative sea-level rise, linked to flexural subsidence, outpaces the growth of shallow-marine carbonates. The upper part of this unit represents the deepest part of the basin and records the position of the basin axis at that time. The upper unit is typically dominated by deep-water siliciclastic sediment (“flysch”) derived from the over-riding thrust wedge. In the Alpine foreland, continued advance of the orogenic wedge eventually led to the drowning of the carbonate ramp and deepening of the basin. This in turn, led to the deposition of pelagic marls and eventually turbidites — the classic flysch facies (Allen et al., 1991; Gupta, 1997; Sinclair et al., 1998). By constrast, in the Appenines Oligocene to Miocene turbidites (Ricci-Lucci, 1986, 1990) are overlain by Tortonian/Messinian euxinic shales, turbiditic sandstones and evaporites, both primary shallowwater facies and deep-water resedimented deposits (Roveri et al., 2001). Pliocene–recent sediments in the Adriatic Sea, the present foreland basin to the Apennines, are represented by Pliocene fan deltas along the coast and large thicknesses of turbidites and interbedded hemipelagic mudstones within the basin. The Pleistocene is characterised more by the presence of deltaic sandstone and mudstone (Ori et al., 1986). Previous studies have shown that the majority of “underfilled” foreland basins evolve into filled, or overfilled, basins with the deposition of significant thickness of continental sediments (Tankard, 1986; Coakley and Watts, 1991; Allen et al., 1991; Sinclair, 1997). For example, the Oligocene to Miocene of the Alpine system was dominated by ‘molasse’ derived from the overriding thrust wedge. Initially there was a transition from the Eocene deep-marine turbidites to shallow-marine sand deposition, but by the Early Oligocene terrestrial conglomerates were deposited in alluvial fans (Kuhleman and Kempf, 2002). This change in facies marked the transition from an ‘underfilled’ to a ‘filled’ foreland basin. 6.2.2. Application of the “underfilled foreland basin model” A simple application of the “underfilled foreland basin trinity” to the Miocene Hatay sediments results in

the following interpretation. The passive margin sediments are represented by the Eocene and older platform carbonates of the Arabian Platform and the Early Miocene fluvial conglomerates and sandstones (Balyatağı Formation). The top of the platform carbonates (Eocene) is a regional unconformity and local unconformities are present at the top of the Balyatağı Formation. The Early Miocene represents a river system that drained the inboard (“cratonic”) margin of the emerging foreland. In the foreland basin model the clastics represent material eroded from the southern passive margin and a coincident flexural forebulge (the isostatic response to the lithospheric load). Therefore, in the model, the top of the formation marks the boundary between passivemargin and convergent-margin sedimentation. The lower unit of the “underfilled trinity” is then represented by the Middle Miocene Sofular Formation, composed of peritidal, lagoonal, reef and ramp carbonates, together with local palaeosols. Similar limestones to those of the Hatay basin are, for example, reported from the Alps (Sinclair, 1997), Papua New Guinea (Pigram et al., 1989) and the Persian Gulf (Murris, 1980). Lateral facies variation similar to the Hatay area is known from several of these basins, including the Alps. The thickness of these units (30 m–2500 m) is comparable with that of the Sofular Formation (250– 300 m). The middle unit of the “underfilled trinity” is represented by the Upper Miocene Nurzeytin Formation, foraminiferal-rich marl, this is very similar to carbonate pelagic–hemipelagic sediments (e.g. mudstones) of other foreland basins, including the Alps (Sinclair, 1997), the Italian Apennines (Ricci-Lucci, 1986) and Papua New Guinea (Pigram et al., 1989); these sediments range from 50–400 m in thickness. The upper unit of the “underfilled trinity” is not present in the Hatay area. This is surprising as there is a progression to turbiditic sediments in most documented foreland basins (Ricci-Lucci, 1986, 1990; Sinclair, 1997; Sinclair et al., 1998). 6.2.3. Potential anomalies with the “underfilled foreland basin model” There are several problems (local and regional) with applying the simple “underfilled foreland basin model” to the Miocene Hatay sediments, which we now consider. Firstly, in the model (Sinclair, 1997) the Lower Miocene fluvial clastic sediments (Baltatağı Formation) would be attributed to the top of the pre-existing passive margin sequence. However, a major structural and stratigraphic break exists in the area between these Miocene alluvial sediments and the underlying, folded,

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Mid-Eocene carbonates. Only minor discontinuities exist between the top of the Baltatağı Formation and the overlying Mid-Miocene Sofular Formation. This suggests that the fluvial clastics (Baltatağı Formation) should instead be interpreted as the base of the foreland basin succession. If so, the presence of thick clastic sediments differs from the basal facies of most other foreland basins. The unconformity beneath the Lower Miocene clastics (Baltatağı Formation) would then be attributed to the southward passage of a flexural forebulge from the over-riding plate located to the north. In the simple model, the passive margin should not be deformed until flexural uplift takes place. By contrast, the Cretaceous and Mid-Eocene successions of the Arabian foreland in the Hatay area were locally deformed (folded) prior to the Early Miocene, as noted above. This deformation was probably related to the initial collision of the Tauride micro-continent to the north with the Arabian continental margin to the south as the Southern Neotethys closed (Yılmaz, 1993; Robertson, 1998). Rather than being simply related to the passage of a flexural forebulge, the uplift of the foreland may also relate to basement reactivation. Specifically, the uplift to the south of the Hatay area that created the source for the Lower Miocene coarse fluvial sediments may relate to the Syrian Arc. The uplift of the Baer–Bassit Massif, the probable source area to the south, has been attributed to movements associated with the Syrian Arc (Al-Riyami et al., 2002; Hardenberg and Robertson, submitted for publication). The Syrian Arc structures extend throughout the Levant, from southern Turkey to Sinai (Egypt). The Syrian Arc was initiated in the Late Cretaceous and, evolved during the Early Cenozoic, associated with stratigraphic inversion and uplift of pre-existing basement structures (e.g. rift faults) (Chaimov et al., 1990; Alsdorf et al., 1995; Walley, 2001). A second problem with the ideal model, is that the upper turbidite unit of the “underfilled foreland basin trinity” is absent from the Hatay area. An upper turbidite unit is to be expected since southward thrusting is known to have continued along the Tauride thrust front to the north during Mid–Late Miocene time (Gökçen et al., 1988; Robertson et al., 2004) and should have supplied sediment to the basin. The influx of quartz and mica, initially to the northeastern of the Hatay area (Kırıkhan area, Fig. 1) during the Early Miocene may record erosion of the uplifting Palaeozoic basement of the Arabian foreland, rather than derivation from the over-riding Tauride allochthon, as noted earlier. This would again be related to reactivation of structures within the Arabian foreland.

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In summary, the sedimentary record of the Hatay area differs from the ideal model mainly because the foreland was not passive, but rather was structurally reactivated, before, during and after the development of the Miocene foreland basin. 6.2.4. Comparison with more northerly part of the foreland basin The Miocene Hatay sediments can be usefully compared to the more proximal part of the same foreland basin system, located c. 75 km further north, near the Tauride thrust front (Fig. 1). This proximal basin area extends eastwards from the Misis Mountains, through the Andırın Range and then through SE Turkey to Iran (Sharland et al., 2004). Early Miocene terrigenous deepwater turbidites, interpreted as a proximal (inboard) peripheral foreland basin deposits, are overthrust by the Tauride allochthon throughout this area (Aktaş and Robertson, 1984; Yılmaz, 1993; Gökçen et al., 1988; Robertson et al., 2004). The Tauride allochthon is structurally underlain by deep-water terrigenous turbidites of Early to Mid-Miocene age (Gül, 2006). In the west, in the Misis area, Middle–Upper Miocene shallow-water clastic sediments (Kızıldere Formation) are structurally underlain (southwards) by Middle Miocene deep-water turbidites (Karataş Formation; Gökçen et al., 1988), and then by Middle–Late Miocene shallow-marine to nonmarine clastic sediments, based on outcrop and well data (Gökçen et al., 1988). These Miocene clastic sediments are interpreted as the thrust-imbricated proximal (northerly) part of a peripheral foreland basin, close to the Tauride thrust front. More similar to the Miocene Hatay area sediments are Miocene sequences that directly overlie the Arabian foreland basin succession further south. These sediments are exposed south of the exposed thrust front, extending eastwards from the Iskenderun (Andırın) area in the west (Fig. 1) (Robertson et al., 2004). In general, Eocene– Oligocene white hemipelagic carbonates, interpreted as the uppermost part of the Arabian shelf succession, pass upwards into shallow-water carbonates rich in bioclastic material (e.g. Nummulites, bivalves and echinoids). These sediments are unconformably overlain by nonmarine sediments (2–500 m thick), including red siltstones and mudstones, abundant nodular caliche and channelised pebbly conglomerates of inferred Early Miocene age (Kalecik Formation). These clastic sediments are interpreted as a northward-flowing braided fluvial system (Robertson et al., 2004). The clastic sediments pass transitionally upwards into Middle Miocene shallow-water limestones (Upper Langhian– Serravallian Horu Formation), together with calcareous

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mudstones and localised reef and lagoonal facies (10– 100 m thick Menzelet Formation; Önalan, 1988; Derman et al., 1996). During the Late Miocene thick deltaic clastic sediments (sandstones and conglomerates) prograded over the area, ranging from non-marine in the east (Kahramanmaraş area), to marginal margin and shallowmarine in the west (Andırın–Misis area). The above, more northerly Miocene sediments that overlie the Arabian platform succession show many similarities with those of the Hatay region further south, strengthening our interpretation of this as the distal (southerly) part of a larger “underfilled” foreland basin. In both areas a regional unconformity is present above the Eocene, or older, Arabian platform succession and the unconformity surface is overlain by Lower Miocene fluvial sediments (up to several hundred metres thick), derived from the south. Another similarity is that the clastic sediments in the north pass into Mid-Miocene shallow-marine carbonates including localised coral build-ups, as in the Hatay area. During the Mid-Miocene a shallow-water carbonate-depositing basin, therefore existed for up to ∼ 75 km from north to south, extending east–west for hundreds of kilometres. However, a major difference between the proximal (northerly) and distal (southerly) areas is that coarse clastic sediments occur only in the north. These northerly clastic sediments may correspond to the upper unit of the “underfilled trinity”, although these facies are mainly non-marine rather than turbidites as in the simple model. There are several possible explanations for the absence of the upper deep-marine unit of the “trinity” in the Hatay area further south. The first is that the area was simply further from the over-riding allochthon (∼75 km). The second is that structural barriers existed within the foreland. The presence of clastic sediment within the Miocene succession, apparently eroded from the Palaeozoic succession of the Arabian foreland in the Amanos Mountains (Fig. 1), implies that uplift of this area took place and thus that a barrier to southward sediment dispersal probably existed. The third explanation is that a westward non-marine, to marine, transition existed during the Late Miocene in the north of the region, close to the thrust front (Derman et al., 1996; Robertson et al., 2004). This implies that a palaeoslope existed in this direction i.e. towards the deep Eastern Mediterranean basin. Clastic sediment was, thus, transported axially (westwards) rather than prograding southwards across the foreland. It is likely that all three of these factors (i.e. distance from source, structural barriers and axial sediment transport) contributed the absence of the upper unit of the “underfilled trinity” in the Hatay area. Thus existing ideal foreland basin

models have to be modified to take account of the palaeogeography and tectonics of the Hatay area. 7. Proposed model: distal foreland basin in a diachronous collisional setting We envisage that the main stages of the sedimentary development of the distal (southerly) part of the Miocene underfilled peripheral foreland basin correspond to stages of overthrusting of the Tauride allochthon from the north. Stage 1: Early Miocene (Fig. 17). This period includes uplift and deformation of the Upper Cretaceous and Eocene Arabian carbonate platform and the emplaced ophiolite, resulting in a regional unconformity. Unlike some foreland basins (i.e. Alps) this is not attributable to simple forebulge uplift in view of the amount of observed compressional deformation (e.g. folding) and the presence of focused uplift well to the south of the Tauride thrust front (N 100 km). Instead, the passive margin was already affected by some form of initial collision related to the closure of the Southern Neotethys Ocean to the north, although the exact cause of this deformation remains unknown. As suggested above, this regional collision may have re-activated the Syrian Arc structures, which extend for hundreds of kilometres south of the collisional front. This could explain why the Lower Miocene fluvial facies (Balyataği Formation) are thick and extensive. A greater topographic relief was created than in other foreland basins at this stage of development where similar reactivation of basement structures did not take place. Stage 2: Mid-Miocene (Fig. 17). The development of the basin during the deposition of the Sofular Formation is compatible with an ideal foreland basin model. The basin subsided under flexural control, as the Tauride thrust load advanced; accommodation space was created but remained underfilled, resulting in basin deepening. A carbonate ramp deepened generally northwards towards the thrust front. Offshore from the marginal carbonate ramp hemipelagic muds were deposited closer to the collisional front, representing the area where flexural subsidence exceeded the upbuilding of the marginal carbonate platform. Similar deeper-water sediments are exposed further north along the edge of the Iskenderun Bay (Boulton, 2006). Short and longer periodicity glacio-eustatic sea-level change also influenced sedimentation. Within the Miocene of the Hatay area, localised evidence of syn-sedimentary faulting (Boulton et al., 2006) suggests that carbonate deposition (Sofular Formation) was at least partly controlled by normal

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Fig. 17. Schematic cross-sections showing the regional development from the Taurides in the north to the Arabian Platform in the south; the box on each diagram indicates the approximate position of the field area for each time-slice. First, deformation and uplift of the Arabian passive margin related to final closure of the Southern Neotethys ocean and continent–continent, collision leading to erosion of the flexural bulge and fluvial sedimentation; secondly, further southward thrusting of the over-riding Tauride allochthon associated with flexurally controlled subsidence of the foreland, resulting in migration of the foreland basin and the deposition of shallow-marine carbonates; thirdly, during the Late Miocene further southward overthrusting of the Tauride allochthon and flexural subsidence initiates deeper-water hemipelagic marl deposition, with evaporite deposition near the depocentre during the Messinian. Finally, in the Plio-Quaternary collision ceases and the Hatay area becomes non-marine.

faulting. Normal faulting has also been documented in the lower/middle unit in other forelands basins (e.g. Timor, Veevers et al., 1978; Taconic foreland basin, Bradley and Kidd, 1991; Pelagian Shelf, Argani and Torelli, 2001; the Alps, Sinclair, 1997). Such faulting in the Hatay area is attributed to the continued reactivation of faults within the continental margin due to stresses generated during loading of the lithosphere. Stage 3: Mid–Late Miocene (Fig. 17). Deeper-water facies were established over the shallow-marine limestones in response to further loading of the lithosphere and resulted in additional flexural subsidence. Local faulting was still active and possibly influenced the deposition of the various gravity-driven sediments. The eventual onset of the Messinian salinity crisis resulted in evaporite deposition in local depocentres. The absence of the upper unit of the theoretical “underfilled trinity”

can be explained by a combination of a southerly position (∼ 75 km south of the thrust front), structural barriers within the foreland, or the diversion of clastic sediment, axially into the deep Mediterranean basin to the west. The tectonic regime changed after the end of the Miocene and the Plio-Quaternary Hatay Graben then formed in a transtensional setting related to the westward escape of Anatolia (Boulton et al., 2006). 8. Conclusions During latest Cretaceous and Eocene times the Hatay area near the Mediterranean coast in south-central Turkey was dominated by a regionally extensive carbonate shelf on which deposition ranged from shallowmarine, to intertidal, to non-marine with palaeosol

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development. This sedimentation took place along the northern margin of the Arabian (African) plate adjacent to the Southern Neotethys, which remained partly open to the north. Folding after the Middle Eocene (pre-Early Miocene) resulted from an early phase of continental collision, which probably reactivated basement structures. The pronounced angular unconformity between the Middle Eocene and overlying Early Miocene succession probably reflects the development of a flexural forebulge related to overthrusting of the Tauride allochthon over the Arabian foreland during the final closure of the Southern Neotethys. However, the uplift may also reflect the re-activation of deep-seated structures (e.g. Triassic rift faults) as part of the regionally extensive Syrian Arc that extends from southern Turkey/ northern Syria southwards through the Levant. Sedimentation resumed in the Early Miocene after a hiatus in the Oligocene, with the deposition of a thick sequence of continental conglomerates and palaeosols, which is observed only in the north of the area. Sediments were mainly deposited from braided rivers flowing northwards from an uplifted source area, probably the Baer–Bassit region of northern Syria. During the Middle Miocene, tectonic subsidence during a time of sea-level rise favoured the deposition of shallow-marine carbonates on an overall northwarddeepening carbonate ramp, with lagoons and reefs developed locally. Sediment deposition was influenced by coeval normal faulting. Similar shallow-marine sediments are exposed north of the Hatay area, near the overridding thrust front, suggesting that the Mid-Miocene shallow-marine basin was up to ∼ 75 km across. Tectonic subsidence and deepening continued during the Late Miocene, resulting in widespread hemipelagic marl deposition with interbeds of material transported from marginal environments, probably to the south. Marine sedimentation ceased in the Messinian due to isolation of the Mediterranean from global eustatic sealevel, resulting in gypsum precipitation and reworking near the basin depocentre. The Miocene stratigraphy of the Hatay region can be modelled as the distal part of an “underfilled” peripheral foreland basin related to flexural collapse of the lithosphere during southward thrusting of the Tauride allochthon. However, only the middle unit of Sinclair's (1997) tripartite “trinity” is ideally represented. The presence of thick, coarse Lower Miocene fluvial sediments above the regional unconformity differs from the ideal lower unit of the “trinity”, and probably reflects important collision-related basement reactivation in the Hatay area. The middle unit of the “trinity” records an increase in accommodation space due to southward

migration of the thrust load and is represented by the accumulation of deeper-water marls in the Hatay area. The upper unit of the “trinity” is absent altogether; the most likely reasons being that the Hatay area was far from the supply of terrigenous sediment from the overthrust load; that, structural barriers existed within the intervening foreland, or the clastic sediment from overriding Tauride allochthon was shed laterally towards the deep Eastern Mediterranean Sea during the Late Miocene rather than southwards over the Arabian platform. By the end of the Miocene, convergence along the Tauride thrust front halted and was replaced by strike– slip, transtension and westward ‘tectonic escape’. As a result, the Miocene sediments of the Hatay area were faultdissected to form the Plio-Quaternary neotectonic Hatay Graben. Marginal areas of the basin were uplifted to form the flanks of the graben, while the depocentre gradually filled and regressed during Plio-Quaternary time. Acknowledgements SJB acknowledges receipt of an NERC PhD Studentship (NER/S/A/2002/10361) while at the University of Edinburgh, and additional funding by the American Association of Petroleum Geologists for financial support. AHFR thanks the Carnegie Trust for financial assistance with fieldwork. We thank Prof. U. Ünlügenç for logistical and scientific assistance with this work. We also thank T. Mistik; N. Temizkan; A. Kop and Y. Dokumacı for access to their unpublished M.Sc. theses on Hatay Region. We would also like to thank two anonymous reviewers for their helpful comments. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j. sedgeo.2006.12.001. References Aktaş, G., Robertson, A.H.F., 1984. The Maden Complex, S E Turkey: evolution of a Neotethyan continental margin. In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, vol. 17, pp. 375–402. Allen, P.A., Homewood, P.W., Williams, G.D., 1986. Foreland basins: an introduction. In: Allen, P.A., Howewood, P.W. (Eds.), Foreland Basins: International Association of Sedimentologists Special Publication, vol. 8, pp. 3–12. Allen, P.A., Crampton, S.L., Sinclair, H.D., 1991. The inception and early evolution of the North Alpine foreland basin, Switzerland. Basin Research 3, 143–163.

S.J. Boulton, A.H.F. Robertson / Sedimentary Geology 198 (2007) 93–124 Allen, P.A., Burgess, P.M., Galewsky, J., Sinclair, H.D., 2001. Flexural-eustatic numerical model for drowning of the Eocene periapline carbonate ramp and implications for Alpine geodynamics. Geological Society of America Bulletin 113 (8), 1052–1066. Al-Riyami, K., Robertson, A.H.F., Dixon, J.E., Xenophontos, C., 2002. Origin and emplacement of the Late Cretaceous Baer–Bassit ophiolite and its metamorphic sole in NW Syria. Lithos 65, 225–260. Alsdorf, D., Barazangi, M., Litak, R., Seber, D., Sawaf, T., Al-Saad, D., 1995. The intraplate Euphrates fault system—Palmyrides mountain belt junction and relationship to Arabian plate boundary tectonics. Annali Di Geofisica 38 (3–4), 385–397. Alverez, G., Busquets, P., Permanyer, A., Vilaplana, M., 1977. Growth dynamic and stratigraphy of Sant Pau d'Ordal Miocene patch reef. Province of Barcelona, Catalonia. Second International Symposium of Corals and Fossil Coral Reefs, Paris, 1975. Bureau de Récherches Géologiques et Minières, Mémoires 89, 367–377. Argani, A., Torelli, L., 2001. The Pelagian Shelf and its graben system (Italy/Tunisia). In: Ziegler, P.A., Cavazza, W., Robertson, A.H.F., Crasquin-Soleau, S. (Eds.), Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basins and Passive Margins. Mémoires du Muséum Nationale d'Histoire Naturelle, vol. 186, pp. 529–544. Arpat, E., Şaroğlu, F., 1972. The East Anatolian Fault System; thoughts on its development. Bulletin of the Mineral Research and Exploration (MTA) Institute of Turkey 78, 33–39. Atan, O.R., 1969. Eğrıbucak Kırıkhan bölgesindeki ofıyolilerin jeolojisi ve petrografisi [Geology and Petrography of the Kırıkhan District Ophiolite]. MTA Bulletin no 150, pp.71, Ankara. Barka, A.A., Kadinsky-Cade, C., 1988. Strike–slip geometry in Turkey and its influence on earthquake activity. Tectonics 7, 663–684. Bassant, P., Van Buchem, E.S.P., Strasser, A., Görür, N., 2005. The stratigraphic architecture and evolution of the Burdigalian carbonate–siliciclastic sedimentary systems in the Mut Basin, Turkey. Sedimentary Geology 173, 151–186. Beaumont, C., 1981. Foreland Basins. Royal Astronomical Society Geophysical Journal 137, 291–329. Bentham, P.A., Talling, P.J., Burbank, D.W., 1993. Braided stream and flood-plain deposition in a rapidly aggrading basin: the Escanilla formation, Spanish Pyrenees. In: Best, J.L., Bristow, C.S. (Eds.), Braided Rivers, Geological Society, London. Special Publications, vol. 75, pp. 177–194. Boccaletti, M., Calamite, F., Deina, G., Gelati, L., Massari, F., Moratti, G., Ricci-Lucchi, F., 1990. Migrating foredeep-thrust belt systems in the Northern Apennines and Southern Alps. Palaeogeography, Palaeoclimatology, Palaeoecology 77 (1), 3–14. Bossio, A., Bradley, F., Esteban, M., Giannelli, L., Landini, W., Mazzanti, R., Mazzei, R., Ricci Lucci, F., Slavatorini, G., 1978. Some aspects of the Upper Miocene in Tuscany. Pisa, Messinian Seminar 4. International Geological Correlation Project, vol. 96. 88 pp. Boulton, S.J., 2006. Tectonic-sedimentary evolution of the Cenozoic Hatay Graben, South Central Turkey. University of Edinburgh, unpublished PhD thesis, 414 pp. Boulton, S.J., Robertson, A.H.F., Unlügenç, Ü.C., 2006. Tectonic and sedimentary evolution of the Cenozoic Hatay Graben, Southern Turkey: a two-phase, foreland basin then transtensional basin model. In: Robertson, A.H.F., Mountrakis, D. (Eds.), Tectonic Development of the Eastern Mediterranean Region. Geological Society, London, Special Publications, vol. 260, pp. 613–634. Boulton, S.J., Robertson, A.H.F., Ellam, R.M., Şafak, Ü., Ünlügenç., U.C., in press. Strontium isotopic and micropalaeontological dating

121

used to help redefine the stratigraphy of the Neotectonic Hatay Graben, southern Turkey –Turkish Journal of Earth Sciences. Bouma, A.H., 1962. Sedimentology of Some Flysch Deposits: a Graphi Approach to Facies Interpretation. Elsevier, Amsterdam. 168 pp. Bradley, D.C., Kidd, W.S.F., 1991. Flexural extension of the upper continental crust in collisional foredeeps. Geological Society of America Bulletin 103, 1416–1438. Buchbinder, B., 1996. Miocene carbonates of the Eastern Mediterranean, the Red Sea and the Mesopotamian basin: geodynamic and eustatic controls. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean regions. Society for Sedimentary Geology, Concepts in Sedimentology and Palaeontology, vol. 5, pp. 89–97. SEPM, Tulsa, Oklahoma, USA. Chaimov, T., Barazangi, D., Al-Saad, D., Sawaf, T., Gebren, A., 1990. Crustal shortening in the Palymaride Fold Belt, Syria, and implications for movement along the Dead Sea Fault System. Tectonics 9 (6), 1369–1386. Clevis, Q., De Boer, P.L., Nijman, W., 2004. Differentiating the effect of episodic tectonism and eustatic sea-level fluctuations in foreland basins filled by alluvial fans and axial deltaic systems: insights from a three-dimensional stratigraphic forward model. Sedimentology 51 (4), 809–835. Coakley, B.J., Watts, A.B., 1991. Tectonic controls of the development of unconformities: the North Slope, Alaska. Tectonics 10, 101–130. Çoğulu, E., 1974. Ultrabasic tectonites and layered periodites of the Hatay area (Turkey). Bulletin of the Mineral Research and Exploration (MTA) 83, 139–147. Coliacicchi, R., Passeri, L., Pialli, G., 1975. Evidence of tidal environment deposition in the Calcave Massiccio Formation Central Appeninnes, Lower Lias. In: Ginsburg, R.N. (Ed.), Tidal Deposits. Springer-Verlag, Berlin, pp. 345–353. Collinson, J.D., 1996. Alluvial sediments, In: Reading, H.G. (Ed.), Sedimentary Environments: Processes, Facies and Stratigraphy, 3rd edition. Blackwell Science, pp. 37–82. Coniglio, M., James, N.P., Aissaoui, D.M., 1996. Abu Shaar Complex (Miocene) Gulf of Suez, Egypt: deposition and diagenesis in an active rift setting. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean Regions. Society for Sedimentary Geology, Concepts in Sedimentology and Palaeontology, vol. 5, pp. 367–386. SEPM, Tulsa, Oklahoma, USA. Delaloye, M., Wagner, J.-J., 1984. Ophiolites and volcanic activity near the western edge of the Arabian plate. In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, vol. 17, pp. 225–233. Delaloye, M., Pişkin, Ö., Selçuk, H., Vuagnat, M., Wagner, J.-J., 1980. Geological section through the Hatay ophiolite along the Mediterranean coast, southern Turkey. Ofioliti 52 (3), 205–216. Derman, A.S., Akdağ, K., Gül, M.A., Yeniay, G., 1996. Relationship between sedimentation and tectonics in the Maraş Miocene basin. Turkish Association of Petroleum Geologists. 11th Petroleum Congress, Turkey, pp. 91–102. Dubertret, L., 1939. Sur la genèse et l’ages des roches vertes syriennes. Comptes Rendus de l'Academie des Sciences, Paris 209, 763. Dubertret, L., 1953. Géologie des roches vertes du nord-ouest de la Syrie et du Hatay (Turquie). Notes et Mémoires sur le Moyen-Orient. Muséum National d'Histoire Naturelle 5, 5–179 (Paris).

122

S.J. Boulton, A.H.F. Robertson / Sedimentary Geology 198 (2007) 93–124

Eaton, S., Robertson, A.H.F., 1993. The Miocene Pakhna Formation, Cyprus, and its relationship to the Neogene tectonic evolution of the Eastern Mediterranean. Sedimentary Geology 86, 273–296. Erol, O., Pirazzoli, P.A., 1992. Seleucia Pieria: an ancient harbour submitted to two successive uplifts. The International Journal of Nautical Archeaology 21 (4), 317–327. Esteban, M., 1996. An overview of Miocene reefs from Mediterranean areas. Tectonic controls of Miocene reefs and related sediments in Cyprus. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean Regions. Society for Sedimentary Geology. Concepts in Sedimentology and Palaeontology, vol. 5, pp. 3–54. SEPM, Tulsa, Oklahoma, USA. Flecker, R., Poisson, A., Robertson, A.H.F., 2005. Facies and palaeogeographic evidence for the Miocene evolution of the Isparta Angle in its regional eastern Mediterranean context. Sedimentary Geology 173, 233–276. Follows, E.J., 1992. Patterns of reef sedimentation and diagenesis in the Miocene of Cyprus. Sedimentary Geology 79, 225–253. Follows, E.J., Robertson, A.H.F., Scoffin, T.P., 1996. Tectonic controls of Miocene reefs and related sediments in Cyprus. Tectonic controls of Miocene reefs and related carbonate facies in Cyprus. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean regions. Society for Sedimentary Geology, Concepts in Sedimentology and Palaeontology, vol. 5, pp. 295–316. SEPM, Tulsa, Oklahoma, USA. Ginsburg, R.N., 1975. Tidal Deposits: a Casebook of Recent Examples and Fossil Counterparts. Springer-Verlag, New York. 428 pp. Ginsburg, R.N., Hardie, L.A., Bricker, O.P., Garret, P., Wanless, H.R., 1977. Exposure index: a quantitative approach to defining position within the intertidal zone. In: Hardie, L.A. (Ed.), Sedimentation on the Modern Carbonate Tidal Flats of Northwest Andros Island, Bahamas. John Hopkins University Press, Baltimore, pp. 7–11. Gökçen, S.L., Kelling, G., Gökçen, N., Floyd, P., 1988. Sedimentology of the Late Cenozoic collisional sequence: the Misis Complex, Adana, Southern Turkey. Sedimentary Geology 59, 205–235. Gradstein, F.M., Ogg, J.G., Smith, A.G., Agterberg, F.P., Bleeker, W., Cooper, R.A., Davydov, V., Gibbard, P., Hinnov, L.A., House, M.R., Lourens, L., Luterbacher, H.P., McArthur, J., Melchin, M.J., Robb, L.J., Shergold, J., Villeneuve, M., Wardlaw, B.R., Ali, J., Brinkhuis, H., Hilgen, F.J., Hooker, J., Howarth, R.J., Knoll, A.H., Laskar, J., Monechi, S., Plumb, K.A., Powell, J., Raffi, I., Röhl, U., Sadler, P., Sanfilippo, A., Schmitz, B., Shackleton, N.J., Shields, G.A., Strauss, H., Van Dam, J., van Kolfschoten, T., Veizer, J., Wilson, D., 2004. A Geologic Time Scale, 2004. Cambridge University Press. 589 pp. Gupta, S., 1997. Tectonic control on paleovalley incision at the distal margin of the Early Tertiary Alpine Foreland Basin, Southeastern France. Journal of Sedimentary Research 67 (6), 1030–1043. Gül, M., 2006. Evolution of the turbidite system in the Kahraman Maraş Basin. Cükürova University, Unpublished PhD thesis. Gvirtzman, G., Honigstein, A., Reiss, Z., Almogi-Labin, A., Moshkovitz, S., Lewy, Z., 1989. Upper Cretaceous high-resolution multiple stratigraphy, northern margin of the Arabian platform, central Israel. Cretaceous Research 10 (2), 107–135. Haq, B.U., Hardenbol, J., Vail, P.R., 1987. Chronology of fluctuating sea-levels since the Triassic. Science 4793, 1156–1167. Hardenberg, M., Robertson, A.H.F., submitted for publication. Sedimentary evolution of the NW margin of the Arabian plate and the SW–NE-trending Nahr El-Kabir half-graben in northern Syria during latest Cretaceous–Cenozoic time. Sedimentary Geology.

Hayward, A.B., Robertson, A.H.F., Scoffin, T.P., 1996. Miocene patch reefs from a Mediterranean terrigenous setting in southwest Turkey. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean Regions. Society for Sedimentary Geology, Concepts in Sedimentology and Palaeontology, vol. 5. SEPM, Tulsa, Oklahoma, USA. 391 pp. Hempton, M.R., 1987. Constraints on Arabian Plate motion and extensional history of the Red Sea. Tectonics 6, 687–705. Hiscott, R.N., Pickering, K.T., Beeden, D.R., 1986. In: Allen, P.A., Homewood, P. (Eds.), Foreland Basins: International Association of Sedimentologists Special Publication, vol. 8, pp. 309–326. Hsü, K.J., 1972. The desiccated deep-basin model for the Messinian events. In: Drooger, C.W. (Ed.), Messinian Events in the Mediterranean. North-Holland Pub. Co, Amsterdam, pp. 60–67. Hsü, K.J., Montadert, L., Bernoulli, D., Cita, M.B., Garrison, R.E., Kidd, R.B., Melières, F., Müller, C., Wright, R., 1978. History of the Mediterranean salinity crisis. Initial Reports Deep Sea Drilling Project 42, 1053–1078. Hubbard, J.A.E.B., Pocock, Y.P., 1972. Sediment rejection by Scleractinian corals: a key to palaeoenvironmental reconstruction. Geologisches Rundschau 61, 598–626. Kelling, G., Gökçen, S.L., Floyd, P.A., Gökçen, N., 1987. Neogene tectonics and plate convergence in the Eastern Mediterranean: new data from southern Turkey. Geology 15, 425–429. Kelling, G., Robertson, A.H.F., Van Buchem, F., 2005. Cenozoic sedimentary basins of southern Turkey: an introduction. Sedimentary Geology 173, 1–13. Kozlu, H., 1997. Tectono-stratigraphic units of the Neogene basins (İskenderun, Misis–Andırın) and their tectonic evolution in the eastern Mediterranean region. Çükurova Univ. Natural Science Institute, Adana—Turkey, Unpublished Ph.D. Thesis, 189 p. (in Turkish). Kraus, M.J., Middleton, L.T., 1987. Contrasting architecture of two alluvial suites in different structural settings. In: Etheridge, F.G., Flores, R.M., Harvey, M.D. (Eds.), Recent Developments in Fluvial Sedimentology. Society of Economic Palaeontologists and Mineralogists, Special Publications, vol. 39, pp. 253–262. Kuhleman, J., Kempf, O., 2002. Post-Eocene evolution of the North Alpine foreland basin and its response to Alpine tectonics. Sedimentary Geology 152, 45–78. Lyberis, N., 1988. Tectonic evolution of the Gulf of Suez and the Gult of Aqaba. Tectonophysics 153 (1–4), 209–220. Manzi, V., Lugli, S., Ricci Lucchi, F., Roveri, M., 2005. Deep-water clastic evaporites deposition in the Messinian Adriatic foredeep northern Apennines, Italy: did the Mediterranean ever dry out? Sedimentology 52 (4), 875–902. Mart, Y., Rabinowitz, P.D., 1986. The northern Red Sea and the Dead Sea Rift. Tectonophysics 124, 85–113. Meschede, M., Weiss, R., Schmiedl, G., Hemleben, C., 2002. Benthic foraminiferal distribution and sedimentary structures suggest tectonic erosion at the Costa Rica convergent plate margin. Terra Nova 14, 388–396. Miall, A.D., 1978. Fluvial sedimentology. Canadian Society of Petroleum Geologists Memoir 5 (859 pp.). Miall, A.D., 1985. Architectural element analysis: a new method of facies analysis applied to fluvial deposits. Earth-Science Reviews 22, 261–308. Miall, A.D., 1996. The Geology of Fluvial Deposits: Sedimentary Facies, Basin Analysis and Petroleum Geology. Springer-Verlag, Berlin. 582 pp.

S.J. Boulton, A.H.F. Robertson / Sedimentary Geology 198 (2007) 93–124 Miller, K.G., Kominz, M.A., Browning, J.V., Wright, J.D., Mountain, G.S., Katz, M.E., Sugarman, P.J., Cramer, B.S., Christie-Blick, N., Pekar, S.F., 2005. The Phanerozoic record of global sea-level change. Science 310 (5752), 1293–1298. Murris, R.J., 1980. Middle East: stratigraphic evolution and oil habitat. American Association of Petroleum Geologists Bulletin 64, 597–618. Nemec, W., Steel, R.J., 1988. Alluvial and coastal conglomerates: flow deposits. In: Koster, E.H., Steel, R.J. (Eds.), Sedimentolgy of Gravels and Conglomerates. Memoirs of the Canadian Society of Petroleum Geology, vol. 10, pp. 1–31. Önalan, M., 1988. Kahramanmarş Tersiyer kenar havzasının jeolojik evrimi. Geological evolution of the Kahramanmaraş Tertiary peripheral basin. Geological Bulletin of Turkey 31, 1–10 (in Turkish). Ori, G.G., Roveri, M., Vannoni, F., 1986. In: Allen, P.A., Homewood, P. (Eds.), Foreland Basins: International Associaiton of Sedimentologists Special Publication, vol. 8, pp. 183–198. Parlak, O., Kop, A., Ünlügenç, U.C., Demirkol, C., 1998. Geochronology and geochemistry of basaltic rocks in the Karasu graben around Kırıkhan Hatay, S. Turkey. Turkish Journal of Earth Sciences 7, 53–61. Permanyer, A., Esteban, M., 1973. El arrecife mioceno de Sant Pau d'Ordal provinca de Barcelona, Barcelona. Institut d'Investigacions Geologiques 28, 45–72. Pfinner, O.A., 1986. Evolution of the north Alpine foreland basin in the Central Alps. In: Allen, P.A., Homewood, P. (Eds.), Foreland Basins: International Associaiton of Sedimentologists Special Publication, vol. 8, pp. 219–228. Pigram, C.J., Davies, P.J., Feary, D.A., Symonds, P.A., 1989. Tectonic controls on carbonate platform evolution in southern Papua New Guina: passive margin to foreland basin. Geology 17, 199–202. Pirazzoli, P.A., Laborel, J., Saliège, J.F., Erol, O., Kayan, I., Person, A., 1991. Holocene raised shorelines on the Hatay coasts Turkey: palaeoecological and tectonic implications. Marine Geology 96, 295–311. Pişkin, O., Delaloye, M., Selçuk, H., Wagner, J., 1986. Guide to Hatay geology SE Turkey. Oflioliti 11, 87–104. Pratt, B.R., James, N.P., Cowan, C.A., 1992. Peritidal carbonates. In: Walker, R.J., James, N.P. (Eds.), Facies, Models and Response to Sea-level Change, Geological Association of Canada, St John's, Newfoundland, pp. 303–322. Purser, B.H., Plaziat, J.-C., Rosen, B.R., 1996. Miocene reefs of the northwest Red Sea. In: Franseen, E.K., Esteban, M., Ward, W.C., Rouchy, J.-M. (Eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of Mediterranean Regions. Society for Sedimentary Geology, Concepts in Sedimentology and Palaeontology, vol. 5, pp. 347–366. SEPM, Tulsa, Oklahoma, USA. Rafini, S., Mercier, E., 2002. Forward modelling of foreland basins progressive unconformities. Sedimentary Geology 146 (1–2), 75–89. Reilinger, R.E., McClusky, S.C., Oral, M.B., King, R.W., Toksoz, M.N., 1997. Global positioning system measurements of present-day crustal movements in the Arabia–Africa–Eurasia plate collision zone. Journal of Geophysical Research 102, 9983–9999. Ricci-Lucci, F., 1986. The Oligocene to Recent foreland basins of the northern Appenines. In: Allen, P.A., Homewood, P. (Eds.), Foreland Basins: International Association of Sedimentologists Special Publication, vol. 8, pp. 105–139. Ricci-Lucci, F., 1990. Turbidites in foreland and on-thrust basins of the northern Apennines. Palaeogeography, Palaeoclimatology, Palaeoecology 77 (1), 51–66.

123

Robertson, A.H.F., 1998. Mesozoic–Tertiary Tectonic evolution of the easternmost Mediterranean area: integration of marine and land evidence. In: Robertson, A.H.F., Emeis, K.-C., Richter, C., Camerlenghi, A. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results, vol. 160, pp. 723–782. Robertson, A.H.F., 2000. Mesozoic–Tertiary tectonic-sedimentary evolution of a south Tethyan oceanic basin and its margins in southern Turkey. In: Bozkurt, E., Winchester, J.A., Piper, J.D. (Eds.), Tectonics and Magmatism in Turkey and the Surrounding Area. Geological Society, London, Special Publications, vol. 173, pp. 43–82. Robertson, A.H.F., 2002. Overview of the genesis and emplacement of Mesozoic ophiolites in the Eastern Mediterranean Tethyan region. Lithos 65, 1–67. Robertson, A.H.F., Dixon, J.E., 1984. Introduction: aspects of the geological evolution of the Eastern Mediterranean. In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution of the Eastern Mediterranean. Geological Society, London, Special Publications, 17, pp. 1–74. Robertson, A.H.F., Eaton, S., Follows, E.J., Payne, A.S., 1995. Depositional processes and basin analysis of Messinian evaporites in Cyprus. Terra Nova 7, 233–253. Robertson, A.H.F., Ünlügenç, U.C., İnan, N., Taşli, K., 2004. The Misis–Andırın Complex: a Mid-Tertiary melange related to latestage subduction of the Southern Neotethys in S Turkey. Journal of Asian Earth Sciences 22, 413–453. Rojay, B., Heimann, A., Toprak, V., 2001. Neotectonic and volcanic characteristics of the Karasu fault zone Anatolia, Turkey: the transition zone between the Dead Sea transform and the East Anatolian fault zone. Geodinamica Acta 14, 197–212. Rosenfeld, A., Hirsch, F., 2005. The Paleocene–Eocene of Israel. In: Hall, J.K., Krasheninnikov, V.A., Hirsch, F., Benjamini, C., Flexer, A. (Eds.), Geological Framework of the Levant, Volume II: The Levantine Basin and Israel. Historical Productions — Hall, Jerusalem, pp. 437–459. Roveri, M., Bassetti, M.A., Ricci Lucchi, F., 2001. The Mediterranean Messinian salinity crisis: an Apennine foredeep perspective. Sedimentary Geology 140, 201–214. Şafak, Ü., 1993. Antakya Havzası ostrakod foraminifer biyostratigrafisi (The ostracod biostratigraphy of the Antakya Basin) in Turkish, with English abstract. Geological Bulletin of Turkey 36, 115–137. Saller, A., Armin, R., Ichram, L.O., Sulliven, C., 1993. Sequence stratigraphy of aggrading and backstepping carbonate shelves, Oligocene, Central Kalmantan, Indonesia. In: Loucks, R.G., Sarg, J.F. (Eds.), Carbonate Sequence Stratigraphy: Recent Developments and Applications. American Association of Petroleum Geologists Memoir, vol. 57, pp. 267–290. Sartorio, D., Venturini, S., 1988. Southern Tethys Biofacies. AGIP, Italy. 235 pp. Schreiber, B.C., Friedman, G.M., Decima, A., Schreiber, E., 1976. Depositional environments of the Upper Miocene Messinian evaporate deposits of the Silician Basin. Sedimentology 23, 729–760. Şengör, A.M.C., Yılmaz, Y., 1981. Tethyan evolution of Turkey: a plate tectonic approach. Tectonphysics 75, 181–241. Şengör, A.M.C., Görür, N., Şaroğlu, F., 1985. Strike–slip faulting and related basin formation in zones of tectonic escape: Turkey as a case study. Society of Economic Palaeontology and Mineralogy Special Publication, vol. 37, pp. 227–264. Selçuk, H., 1981. Etude géologique de la partie meridionale du Hatay (Turquie): Doctoral thesis, University of Geneva. p.116.

124

S.J. Boulton, A.H.F. Robertson / Sedimentary Geology 198 (2007) 93–124

Sharland, P.R., Casey, D.M., Davies, R.B., Simmons, M.D., Sutcliffe, O.E., 2004. Arabian Plate stratigraphy—revisions to SP2. GeoArabia 9, 199–214. Sinclair, H.D., 1997. Tectonstratigraphic model for underfilled peripheral foreland basins: an Alpine perspective. Geological Society of America Bulletin 109, 324–346. Sinclair, H.D., Sayer, Z.R., Tucker, M.E., 1998. Carbonate sedimentation during early foreland basin subsidence: the Eocene succession of the French Alps. In: Wright, V.P., Burchette, T.P. (Eds.), Carbonate Ramps. Geological Society, London, Special Publications, vol. 149, pp. 205–227. Stockmal, G.S., Beaumont, C., 1987. Geodynamic models of convergent margin tectonics: the Southern Canadian cordillera and the Swiss Alps. In: Beaumont, C., Tankard, A.J. (Eds.), Sedimentary Basins and Basin Forming Mechanisms: Canadian Society of Petroleum Geologists Memoir, vol. 12, pp. 93–411. Tankard, A.J., 1986. On the depositional response to thrusting and lithospheric flexure: examples from the Appalachian and Rocky Mountain basins. In: Allen, P.A., Homewood, P. (Eds.), Foreland Basins. International Association of Sedimentologists Special Publications, vol. 8, pp. 369–392. Tinkler, C., Wagner, J.J., Delaloye, M., Selçuk, H., 1981. Tectonic history of the Hatay Ophiolites South Turkey and their relationship with the Dead Sea rift. Tectonophysics 72, 23–41.

Tucker, M.E., 1991. Sedimentary Petrology: an Introduction to the Origin of Sedimentary Rocks, 2nd edition. Blackwell Science, Oxford. 253 pp. Veevers, J.J., Falvey, D.A., Robins, S., 1978. Timor Trough and Australia: facies show topographic wave migrated 80 km during the past 3 myr. Tectonophysics 45, 217–227. Walley, C.D., 2001. The Lebanon passive margin and the evolution of the Levanine Neo-Tethys. In: Ziegler, P., Cavazza, W., Robertson, A.H.F., Crasquin-Soleau, S. (Eds.), Peri-Tethys Memoir. 5 PeriTethyan Rift/Wrench Basins and Passive Margins. Memoirs du Museum National D'Histoire Naturelle, pp. 407–439. Watts, K.F., Garrison, R.E., 1986. Sumeini Group, Oman — evolution of a Mesozoic carbonate slope on a south Tethyan continental margin. Sedimentary Geology 48, 107–168. Wright, V.P., Burchette, T.P., 1996. Shallow water carbonate environments, In: Reading, H.G. (Ed.), Sedimentary Environments: Processes, Facies and Stratigraphy, 3rd edition. Blackwell Science, pp. 325–394. Yılmaz, Y., 1993. New evidence and model on the evolution of the southeast Anatolian orogen. Geological Society of America Bulletin 105, 251–271.