3.a. The Soğucak Formation

The Middle to Upper Eocene Soğucak Formation is exposed in Cape Karaburun where it is represented by different facies of reef limestones (Fig. 4a). It consists of grey to beige biomicrite limestones, massive or thickly to medium bedded and in places containing karstic voids. Intraclasts are angular fragments of corals, bryozoans, gastropods, bivalves, echinoids, algae and micritic or rare sparitic limestones. The massive limestones laterally change into medium- to thick-bedded limestones (Fig. 4a). The rocks are rich in fossils (Fig. 4b, c) and have been interpreted as having been deposited in typical reef environments (Oktay, Eren & Sakınç, Reference Oktay, Eren and Sakınç1992; Sakınç, Reference Sakınç1994).

Figure 4.

Depositional features of the Soğucak Formation. (a) Facial changes of massive reef limestone (6m) (right) into thickly bedded pre-reef limestone; fractured rocks near the contact show no significant displacement. Height of person for scale is 165cm. (b) In situ position of a coral colony in the Soğucak Formation. Hammer for scale is 33cm long. (c) Fossil-rich reef limestones. Coin for scale is 2.7cm diameter. (d) The Soğucak limestone is exposed stratigraphically higher than the alleged unconformity (hard ground) assumed at the base of the Ceylan Formation. Height of person for scale is 165cm. (e) Bedding plane surface coated by iron oxides showing their higher concentration along NW-striking cracks. Coin for scale is 2.7cm diameter.

Only late Priabonian limestones containing Heterostegina gracilis Herb (Shallow Benthic Foraminifera (SBZ) subzone 20) are exposed in Cape Karaburun (Less, Özcan & Okay, Reference Less, Özcan and Okay2011). In the Çatalca region, the similar rocks (Oktay, Eren & Sakınç, Reference Oktay, Eren and Sakınç1992; Sakınç, Reference Sakınç1994) are about 100m thick and contain Lutetian to Priabonian fossils (Yurtsever & Çağlayan, Reference Yurtsever and Çağlayan2002). In the neighbouring regions (south of Karaburun, Çatalca), the Soğucak Formation is underlain by the Lutetian–Priabonian Koyunbaba (or İslambeyli) Formation of sandy limestone with volcaniclastic rocks at the base (2–100m thick).

Some researchers reported unconformable relationships with the overlying Ceylan Formation and point out the existence of a hard ground at the top of the Soğucak Formation (Oktay, Eren & Sakınç, Reference Oktay, Eren and Sakınç1992; Sakınç, Reference Sakınç1994; Less, Özcan & Okay, Reference Less, Özcan and Okay2011). Indeed, this surface is coated with iron oxides but it cannot be considered as the top of the Soğucak Formation for three reasons (Fig. 4d). First, limestones of the Soğucak Formation are also exposed at higher elevation than the hard ground surface (Fig. 4d). Secondly, in places, thin (2–4mm) bedding-parallel surfaces coated with iron oxides may deviate down-dip along fractures (Fig. 4e). Thirdly, signs of bioturbation, incrustation by shells or traces of boring organisms are missing at and below the surface. All of these indicate the absence of a hiatus and unconformity suggested by previous studies between the Soğucak and Ceylan formations.

3.b. The Ceylan Formation

Oktay, Eren & Sakınç (Reference Oktay, Eren and Sakınç1992) and Sakınç (Reference Sakınç1994) described the basal levels of the Ceylan Formation (their Karaburun Formation) as beach and shallow-water coarse-grained clastic sediments that grade into deeper marine shales and marls and then to deltaic sandstone. Our observations do not support this stratigraphic succession because the Soğucak and Ceylan formations are juxtaposed across a SW-dipping normal fault (Figs 2, 5) and the Ceylan Formation leans against the Soğucak Formation at its higher stratigraphic levels. The direct contact between these formations is not exposed but striation, lunate structures (Petit, Reference Petit1987) and the orientations of secondary faults observed in the Soğucak Formation in the closest vicinity to the contact indicate down-dip displacement.

Figure 5.

Structural and stratigraphic relationships between the Soğucak and Ceylan formations. Points a, b and c indicate the same locations in parts (a) and (b). (a) View of the Karaburun Peninsula to the east-southeast. Note gentle dips of the shallow-marine sandstones (diamictites) at point b and the drastic change to vertical dips across the Karaburun Thrust (point c). (b) Geological cross-section of the Karaburun Thrust (see Fig. 2 for location).

Contrary to previous researchers, we distinguish three stratigraphic units in the Ceylan Formation (from bottom to top): (1) deep-marine shales, (2) shallow-marine sandstones and (3) shallow-marine limestones (Fig. 3; A. G. Say, unpub. thesis, Istanbul Technical Univ., 2007). The deep-marine shales show uniform composition and structure across the studied area whereas the shallow-marine sandstones and shallow-marine limestones represent the facies equivalents of each other. The shallow-marine sandstones are distributed in the northeast while the shallow-marine limestones are exposed to the southwest. The older rocks of the Ceylan Formation constitute the core of a broad NW-striking anticline (Fig. 6a) in which deep-marine shales and marls are exposed. The boundary between the overlying shallow-marine units is transitional and roughly coincides with the crest of the broad anticline (Fig. 6a).

Figure 6.

Broad anticline and sedimentary structures of the Ceylan Formation. (a) Facial changes between shallow-marine sandstones and shallow-marine limestone units. N – normal faults. (b) Rhythmic alternation together with graded bedding (see part (c)) in the deep-marine shale unit indicate its turbiditic nature. Height of person for scale is 165cm. (c) Load cast, graded bedding and inverse graded bedding in the deep-marine shale unit. Hammer for scale is 37cm long. (d) Mud diapir exposed on the Terkos Lake shore. It cuts the deep-marine shales and is truncated by a slump of the shallow-marine limestone. (e) Small-scale downlap (short dashed line) and internal convergence structures (long dashed lines) in the shallow-marine sandstone unit. Both indicate that deeper parts of the basin floor were to the west.

The deep-marine shale unit consists of light grey to bluish grey calcareous shale, marl and fine-grained calcareous sandstone revealing a rhythmic alternation (Fig. 6b). It contains various planktonic fossils and coccoliths (Sakınç, Reference Sakınç1994). In places, the rocks reveal graded bedding suggesting a possible turbiditic nature (Fig. 6c). In the upper part of the unit, load structures and inverse graded bedding have been observed (Fig. 6c). Broken pieces of similar composition have also been observed in mud diapirs (see Section 4.b below), which penetrate the deep-marine shales (Fig. 6d).

Clastic dykes (see Section 4.b below) are characteristic features of the deep-marine shales, although their distribution in the unit is irregular. Sandstones in the clastic dykes are calcareous and contain either no or very small amounts of siliciclastic, exotic material. Most likely they formed because of erosion of the Soğucak limestones and compose a stratigraphic level that is not exposed now.

The visible thickness of the deep-marine shales is 56m. We argue that the lower stratigraphic levels of the Ceylan Formation are not exposed in the Karaburun area. Seismic profiling has established a pile of turbidites deposited on the continental slope (upper slope to basin floor environment) to the north of Karaburun. The thickness of the Oligocene rocks increases to the north from 500 to 1500m and shows an onlap structure of turbidite lobes (Sipahioğlu & Menlikli, Reference Sipahioğlu and Menlikli2009; N. Ö. Sipahioğlu, pers. comm.).

In the northeastern part of the Karaburun area, the overlaying shallow-marine sandstone unit consists of coarse-grained sandstones, gravel- and pebble-size (up to 20cm) matrix-supported conglomerates and diamictites (non-genetic term) (Fig. 3). The upper contact of this unit is not exposed; therefore, we estimate its thickness as being more than 54m. Clasts are poorly sorted, rounded and made up of the Soğucak limestone, andesitic porphyries, phaneritic granite, granodiorites and rare gabbro. Nearby exposures of the Late Cretaceous Rhodope–Pontide arc (Şengör & Yılmaz, Reference Şengör and Yılmaz1981) are the most likely source of the igneous material. The bedding planes of the coarse-grained clastic rocks are wavy but parallel. East-orientated internal convergence of layers and small-scale downlaps to the west indicate deepening of the basin floor to the west (Fig. 6e).

The age of this unit is defined by previously found Nummulites vascus (Oktay, Eren & Sakınç, Reference Oktay, Eren and Sakınç1992; Sakınç, Reference Sakınç1994). Additionally, Less, Özcan & Okay (Reference Less, Özcan and Okay2011) reported findings of N. bouillei and other larger foraminifera, which date the rocks as Early Rupelian (SBZ 21).

The second facies of the upper part of the Ceylan Formation consists of shallow-marine limestones that are exposed in the western and southwestern parts of the studied area (Fig. 2). Thick- to medium-bedded limestones of this facies alternate with grey shale and marl (Fig. 3). In terms of allochems, the limestones are variable. Oolites, pellets, grain aggregates, limestone clasts and skeletal grains are present. The total thickness of this unit is more than 20m.

On the Black Sea coast, the shallow-marine coherently bedded limestones overlie the deep-marine shales across a well-developed erosional surface (Fig. 7a) and then are conformably overlain by bluish shales lithologically similar to the deep-marine unit. In Durusu, the shallow-marine limestones reveal submarine slump structures (see below) and are conformably overlain by the bluish shale horizon that wedges out to the east, vertically and laterally passing into the shallow-marine limestone unit (Fig. 7b). These observations are critical for understanding the mechanism of shallowing of the Ceylan Basin.

Figure 7.

(a) A well-developed erosional surface at the top of the deep-marine shale unit. (b) W-vergent drag fold in the deep-marine shales produced by slumping (the Durusu region). Wedging to the east of the shale horizon (marker horizon; see text) indicates general uplift of the region in the hanging wall of the Karaburun Thrust. (c) Imbricated bedding in the slump of the shallow-marine limestones (at Durusu, see Fig. 2) indicating slumping from E to W. (d) Recumbent fold with NW-oriented hinge in the slump horizon. To the east, limestones in the same horizon reveal more coherent bedding and overlie the deep-marine shale across an erosional surface (see part (b)). (e) The head of the submarine slump; the deep-marine shales are also involved in slumping. (f) Clastic dykes that intrude the deep-marine shale unit along conjugated normal faults (Black Sea coast). W-dipping faults have a small magnitude of displacement (arrows) while displacement along the E-dipping fault may be larger than the scale of the photo. In overlying shallow-marine limestones displacement along this fault is 0.5m, which suggests its Miocene–Pleistocene reactivation. Note that the material of the clastic dykes is not foliated.

The shallow-marine limestones reveal a wide spectrum of synsedimentary deformation related to submarine sliding. In Durusu (Fig. 2), an imbricate stack of rotated blocks and fragments of asymmetric synsedimentary folds in the shallow-marine limestones indicate slumping to the west (Fig. 7c). The slump surface cuts clastic dykes in the underlying deep-marine unit. The above-mentioned mud diapirs are also truncated by the slump surface (Fig. 6d). At Cape Kılıç (Fig. 2), submarine slides which are at the boundary between the deep-marine shale and shallow-marine limestone units have many common features with those in Durusu. About 500m to the east of the cape, the above-mentioned erosional surface at the base of the shallow-marine limestone units (Fig. 7a) transforms to a slump surface. Similar to Durusu, deformed limestone layers form recumbent folds indicating a westward transport direction (Fig. 7d). Structures at the slump head thrust the shallow-marine limestone onto the deep-marine shales (Fig. 7e). No slump structures have been observed in the shallow-marine limestone unit to the west of Cape Kılıç and in the western side of Terkos Lake.

We consider these slump structures to be critical events in the tectonic evolution of the area. Displacements along observed normal faults in the deep-marine units do not exceed a few metres and therefore could not drastically change the bathymetry of the basin. However, the facies changes between the upper Ceylan units and general shallowing of the basin suggest an inversion of the basin and change of the extensional regime to a compressional one.

A lens of foliated felsic tuff (0.6–5m) occurs at the base of the shallow-marine limestones (Fig. 3). Several stratigraphic levels of felsic tuffs are mentioned within the Ceylan Formation (Turgut & Eseller, Reference Turgut and Eseller2000; Okay, Yurtsever & Siyako, Reference Okay, Yurtsever and Siyako2006; Varol, Baykal & Ayyıldız, Reference Varol, Baykal and Ayyıldız2009; Less, Özcan & Okay, Reference Less, Özcan and Okay2011) but their exact stratigraphic positions are not reported.

3.c. Mezardere Formation and younger rocks

The stratigraphy of the rocks that are younger than the Lower Oligocene Ceylan Formation is unclear because of the poor outcrop conditions of these rocks in the studied area. At the top of a hill located 3km to the northwest of Balabanburun (Fig. 2), we encountered an outcrop of green marls occurring stratigraphically higher than the shallow-marine limestone unit. These rocks can be interpreted as a change in the depositional environment to a new deepening of the sea. Only the higher structural position of these marls is used for their stratigraphic correlation owing to the absence of fossils. We tentatively correlate these rocks with the Lower Oligocene Mezardere Formation, which consists of shale, marl and subordinate tuff (Okay, Yurtsever & Siyako, Reference Okay, Yurtsever and Siyako2006).

The youngest Çukurçeşme Formation of reddish fluvial clast-supported conglomerates and sands is assigned to the Upper Miocene (Oktay, Eren & Sakınç, Reference Oktay, Eren and Sakınç1992; Sakınç, Reference Sakınç1994). It contains a mammal fauna of the Pannonian (≈Tortonian) Stage, of the Paratethyan Upper Miocene (Sakınç, Yaltırak & Oktay, Reference Sakınç, Yaltırak and Oktay1999). This formation unconformably covers all the above-mentioned Cenozoic rocks in the east of the Karaburun area. Widespread cross-bedding (30–60cm) is planar and has more or less consistent orientations indicating the westward transportation of sediments. The consistency of this orientation and the uniform composition of the rocks suggests their deposition in braided river channels. Rounded pebbles of polymictic composition, including a significant amount of greenschist-facies metamorphic rocks, suggest long-distance transportation.

4.a. Late Eocene – Early Oligocene normal faults

These faults are better developed in the Soğucak Formation. In places, fault planes are striated and show asymmetric structures (R shears, tension gashes, slickensides, etc.) indicating a normal sense of motion. These faults dominantly dip to the southwest at 50–75° but steeper faults have E–W strike and show a dextral sense of motion. Temporal relationships between these fault populations are unknown.

The direct contact between the Soğucak and Ceylan formations is not exposed (see Section 3.b above) but the former abuts against the shallow-marine sandstone, omitting the deep-marine shale unit. A small-scale synsedimentary normal fault with a magnitude of displacement of 2–5cm has been observed in the shallow-marine sandstone unit near the contact. It also dips to the southwest similarly to the normal faults in the Soğucak Formation (Fig. 5).

Further west along the Black Sea coast, NW- and NE-striking normal faults are widespread but unfortunately their age is not always possible to determine. Faults that cut the deep-marine shales seem to be more abundant than those in the shallow-marine limestone. Differences in magnitude of displacement of the stratigraphic contacts in the deep-marine shale and shallow-marine limestone units show that some of the faults were most likely reactivated afterwards in Miocene–Pleistocene time. Usually the normal faults have small magnitudes of displacement (10–50cm) but in places, the correlation of layers in the fault walls is impossible and suggests a larger magnitude of displacement (Fig. 7f). Figure 10a shows the orientation of normal faults that we relate to the extension and the collapse of the carbonate platform.

4.b. Clastic dykes and mud diapirs

Clastic dykes are related to the Early Oligocene extension, the collapse of the carbonate platform, its drowning and the establishment of deep-water turbiditic sedimentation in the studied area. The distribution of the clastic dykes has a clear stratigraphic control: they are observed in the deep-marine shale unit but never in the shallow-marine sandstone and shallow-marine limestone units. Mechanisms for the clastic dyke formation have been discussed by Jolly & Lonergan (Reference Jolly and Lonergan2002) (Fig. 8). The simplest one is a passive filling of pre-existing cracks and voids (Neptunian dykes). In the Çatalca region, Less, Özcan & Okay (Reference Less, Özcan and Okay2011) described a Neptunian dyke in the Soğucak Formation, which fills a crack in the metamorphic basement. Most likely, this dyke indicates the earliest stage of the carbonate platform disruption. But, in the Karaburun area, the majority of the clastic dykes came from below and their emplacement was caused by a rise in the fluid pressure as can be inferred from the dyke compositions (carbonate sandstone similar to the erosional products of the Soğucak limestone that are expected at the base) and the composition of the host rocks. In addition, these dykes often demonstrate transitions to clastic sills, which are characteristic intrusions developing at shallow structural levels (Jolly & Lonergan, Reference Jolly and Lonergan2002).

Figure 8.

Conditions of clastic dyke formation according to Jolly & Lonergan (Reference Jolly and Lonergan2002). (a) In a non-fractured system, a dyke intrudes when pore fluid pressure (P_f) overcomes the principal minimal stress (σ₃) plus the tensile strength (T). (b) In the presence of older faults or fractures, dykes are formed when fluid pressure (P_f) exceeds normal stress (σ_n) across pre-existing faults or fractures. (c) In the case of a downward-propagating fault (typical for extensional settings), it deflects towards a sealed over-pressurized sand body because of local distortion of the stress field (σ_Lith is the lithostatic pressure).

Clastic dykes can be formed in almost all depositional environments but are more frequently seen in the deep-marine turbiditic settings, where sand layers and lenses are entrained in impermeable mud (Jolly & Lonergan, Reference Jolly and Lonergan2002). The critical condition for clastic dyke formation is the excess of fluid pressure over lithostatic pressure (Jolly & Lonergan, Reference Jolly and Lonergan2002). During simple burying, the critical onset of fluid overpressure may be achieved at the depth of 375 to 500m (Jolly & Lonergan, Reference Jolly and Lonergan2002), which is much larger than the visible thickness of the deep-marine unit of the Ceylan Formation. Note that to the north of Karaburun the Oligocene turbidites are thick enough (Sipahioğlu & Menlikli, Reference Sipahioğlu and Menlikli2009) to create overpressure horizons. Taken alone, extension-driven subsidence and the building of lithostatic and fluid pressure are still not favourable for tensile cracking and initiation of clastic dyke formation, because of the sealing effect of impermeable layers, which in our case are marls and calcareous shales (Fig. 8a). Stress gradients should appear to move sand grains together with fluid to zones where hydraulic fracturing is possible (Jolly & Lonergan, Reference Jolly and Lonergan2002; Pluijm & Marshak, Reference Pluijm and Marshak2003). Theoretically, if the fluidization of sand because of the formation of a stress gradient has happened, the dyke intrusion should be vertical, perpendicular to the least stress and should not have a large extent (Fig. 8a). In general, this simple requirement is applicable to the Karaburun geological setting where impermeable marl and shale form seal horizons. However, we also should account for normal faulting. Jolly & Lonergan (Reference Jolly and Lonergan2002) considered two cases: (1) a clastic dyke intrudes the pre-existing fault and (2) a clastic dyke intrudes the propagating fault. Note that boundary conditions for these two cases are the same in terms of the depths at which fluid pressure can overcome the lithostatic pressure. In the first case, the clastic dyke intrusion is controlled by the most favourable orientation of the pre-existing faults. In the Karaburun area, where extension lasted for some period during latest Eocene – Early Oligocene times, clastic dykes should be perpendicular to a certain value of minimal normal stress on planes of pre-existing faults (Fig. 8b). In the second case, the downward-propagating fault plane deflects to a zone of fluid overpressure (e.g. sealed lens of sandstone) and when it intersects its boundary the dyke intrusion starts (Fig. 8c). Application of this theoretical model to Karaburun shows it may explain the emplacement of the majority of clastic dykes in the deep-marine shale unit of the Ceylan Formation.

In the studied area, the clastic dykes have an irregular distribution. Their highest concentration has been observed in the hanging wall of the Karaburun Thrust (Fig. 2). Clastic dykes are absent in the zone of the vertical bedding dips that was produced by the Karaburun Thrust (Fig. 5). Along the Black Sea coast, the dyke density gradually but quickly decreases to the west. Only a few dykes cut the deep-marine shales in the Durusu region but two mud diapir bodies are located there (Fig. 6d).

Clastic dykes consist mainly of light-coloured, well-sorted medium-grained carbonate sandstone. Although erosion products of the Soğucak limestones are most likely their source, clastic rocks of suitable composition are not exposed in the nearby-located regions. In addition to carbonate sandstone, very rare angular clasts of volcanic rocks (up to 7mm) together with rounded limestone pebbles can also be found. In the Durusu region, the composition of the clasts in two neighbouring dykes is different: carbonate sandstone (typical composition) versus polymictic sandstone. It means that the dykes have sampled different overpressure horizons. These two dykes, wedging down, (Fig. 6d) are subvertical. They reflect the state when the tensile strength of the shales and marls (seal rocks) is overcome by brittle failure.

These dykes are located very close to the mud diapirs, one of which (2–3m diameter) has a funnel shape (Fig. 6d), and the shape of the other (20m across) is irregular. The inner structure of these diapirs is chaotic; broken fossils are similar to those of the Soğucak Formation. Mud diapirs can be formed in different settings but mostly they indicate the state of compression (Limonov et al. Reference Limonov, van Weering, Kenyon, Ivanov and Meisner1997; Kopf, Reference Kopf2002; Barreca, Reference Barreca2014).

Along the Black Sea coast, most dykes are located in the hanging wall of the Karaburun Thrust. Away from the thrust, normal faults control dyke intrusions. Faults are often conjugated and the dihedral angle bisected by the σ₃ is ≈ 95–110° (Fig. 7f). Narrow (a few centimetres) zones of foliated cataclasites are restricted to the hanging walls but material of the dyke is not sheared. These relationships indicate a later dyke formation with respect to the faults, which is comparable with the mechanism depicted in Figure 8b. Other dykes in the same area do not have cataclasites and, therefore, for their origin the fault propagation mechanism can be inferred. This mechanism is more obvious if an inclined dyke fits the expected orientation of the maximal shear stresses in the lower part of the outcrop and then changes its orientation to vertical in the upper part. In this case, the dyke initially fills the propagating fault but at a higher structural level it propagates along the tensile crack. The cross-cutting relationships of the clastic dykes along with the variety of mechanisms for their formation indicate the longevity of the dyking process.

An interesting type of clastic dyke is represented by layered dykes (Fig. 9a) where millimetre-scale clastic dyke material alternates with calcite veins. Subhorizontal layering is oriented perpendicular to the dyke's walls. Sometimes calcite veins reveal an asymmetric structure, having a sharp boundary on one side and a defused one on the other, while veins thin out on entrance to the host rocks. These structures require calcite feeding from outside and continuous lengthening of the dyke. We suggest calling them restuffed dykes. Calcite colour variations in restuffed dykes indicate that they were fed from fluid overpressure horizons located at different stratigraphic levels. In places, restuffed dykes are entrained in the inner part of internally homogeneous dykes. These structures suggest dyke feeding from different sources and a long duration of dyke formation.

Figure 9.

(a) Restuffed clastic dyke in the deep-marine shale unit. Thin, parallel calcite veins (light colour) indicate continued elongation of the dyke body. Note their wedging to the right and the sharp left boundary coinciding with the pre-existing fault. Coin for scale is 2.5cm diameter. (b) A curved clastic dyke indicating rotation of the stress field. Pencil for scale (circled) is 17cm long. (c) Various orientations of clastic dykes in the hanging wall of the Karaburun Thrust.

The most complicated orientation of the dykes has been observed in the hanging wall of the Karaburun Thrust (Figs 2, 5, 9c). There, dykes are mainly located to the west of the thrust and are squeezed between two normal faults (Fig. 5) which cut the contact with the shallow-marine sandstone and, therefore, are assigned to the Miocene structural event. We should stress that clastic dykes have never been observed in the shallow-marine limestone or sandstone units and, therefore, we exclude contribution of the Miocene deformation to the dyke formation.

The orientation of clastic dykes in the hanging wall of the Karaburun Thrust is shown in Figure 10b, while Figure 10c shows the orientations of clastic dykes 300m to the west of the thrust. In order to account for tilting caused by the formation of the Karaburun Thrust or tilting caused by the formation of a broad anticline to the west of the thrust (Fig. 6a), the measured orientations of the dykes have been rotated as if the bedding was horizontal. Comparison of the stereonets shows an increase of shallow-dipping dykes towards the hanging wall of the Karaburun Thrust despite the fact that the orientation of the bedding has monoclinal dips at rather shallow angles. Clastic dykes are absent in the zone of the vertical bedding that was produced by the thrust (Fig. 5). These relationships indicate that dyke injections that were initiated during the extension continued during the thrust formation making a bridge between the extensional and shortening events, as a bent clastic dyke shows (Fig. 9b).

Figure 10.

(a) Orientation of latest Eocene – Early Oligocene normal faults in the Soğucak and Ceylan formations (poles, equal-area projection, lower-hemisphere). (b) Clastic dykes in the hanging wall of the Karaburun Thrust (rotated to horizontal bedding). (c) Clastic dykes 250 and more metres to the west of the Karaburun Thrust (rotated to horizontal position of the bedding). (d) Orientation of Miocene normal faults.

4.c. Early Oligocene shortening (Karaburun Thrust)

Oktay, Eren & Sakınç (Reference Oktay, Eren and Sakınç1992) mapped the Karaburun Thrust as a WNW-striking structure of southwestern vergence providing no reasoning for this kinematic interpretation. In the following, we will argue that this structure has a different vergence and has controlled many other aspects of the geological evolution of the Karaburun area and, therefore, must have yet unknown significance for the surrounding regions.

The fault zone (~70m), which is marked by vertical bedding in the shallow-marine sandstone unit, strikes 290–295°. Although steep cliffs prevent direct observations, Google Earth images show the change of bedding strike to 335° towards the southeast. Moreover, these images of adjacent offshore areas indicate another deviation of the strike to 350° to the north of the shoreline (Fig. 2). Thus, we infer a general northwesterly strike for the Karaburun Thrust.

Fault rocks are not exposed but regolith made of sheared clastic rocks (2–4m wide) is exposed to the northeast of the steep attitude of bedding (Fig. 5). Gentle coherent bedding in the footwall shows no evidence of brittle deformation. In the hanging wall, bedding is also coherent except for a few narrow zones (3–5cm) of backthrusts. There, steep bedding planes reveal a gradual transition to a low-angle one (Figs 5, 6a). Normal, northeastward-younging directions are consistent through the asymmetric anticline located in the hanging wall, but in the northeast they change to the opposite direction in the last few metres of the zone of steep bedding. The outline geometry of the bedding necessitates a steep dip for the Karaburun Thrust.

We assume that northeastern motions along the Karaburun Thrust created the broad asymmetric anticline shown in Figure 6a, although the strike of its axis is poorly constrained. Bedding around Terkos Lake is subhorizontal (Fig. 2) but to the east of Durusu, Oktay, Eren & Sakınç (Reference Oktay, Eren and Sakınç1992) and Yurtsever & Çağlayan (Reference Yurtsever and Çağlayan2002) reported northeastern dips of bedding at 20–35°, which suggest a southeastern trend for the broad anticline parallel with the thrust.

The distribution of facies in the Ceylan Formation is controlled by the position of the footwall and the hanging wall of the Karaburun Thrust. The shallow-marine sandstones carrying large clasts of volcanic rocks were deposited in the footwall while deposits of the hanging wall consist mainly of the shallow-marine limestones. Westward-directed submarine slumps, detected along the Black Sea coast and in Durusu, indicate shedding from the uplifted hanging wall of the thrust.

In the hanging wall, the orientations of the clastic dykes change and the amount of gently dipping dykes increases (Fig. 10b, c). As mentioned in Section 4.b above, many dykes have been intruded along planes of pre-existing faults. Variations in sources of the restuffed dykes as well as the evidence indicating a prolonged dyke formation period suggest changes in the controlling factors responsible for the dyke emplacements. The orientation of principal stresses controls the emplacement of both magmatic and clastic dykes (e.g. Jolly & Lonergan, Reference Jolly and Lonergan2002; Pluijm & Marshak, Reference Pluijm and Marshak2003). Figure 9b shows a curved dyke that was found in the hanging wall of the Karaburun Thrust. Its origin is compatible with our inference about the Early Oligocene rotation of the stress field that caused the transition from extension to compression.

4.d. Late Miocene normal faults

Late Miocene normal faults cut the shallow-marine limestone and shallow-marine sandstone units as well as the Çukurçeşme Formation. They have small magnitudes of displacement, which rarely exceed 1–5m. The majority of the observed fault planes are steep (˃60°) and in places they are conjugated. Fault zones (<25cm) are made up of tectonic breccia or incohesive cataclasites; slickenside surfaces are rare indicating very shallow structural levels of fault formation. Fault planes may be coated by iron oxides or bleached along narrow (10cm) dewatering channels. Figure 10d shows the orientation of these faults.

Karstic voids of the Soğucak Formation are filled by stratified and lithified brownish micritic limestones. Lamination of the filling is tilted but at a different angle to the bedding of the Soğucak limestone. This tilting indicates the youngest deformation in the region.