Messinian forced regressions in the Adana Basin: a near ... - Tübitak

Turkish Journal of Earth Sciences 

http://journals.tubitak.gov.tr/earth/ 

Research Article 

Turkish J Earth Sci 

(2013) 22: 

© TÜBİTAK 

doi:10.3906/yer-1208-3 

Messinian forced regressions in the Adana Basin: a near-coincidence of tectonic and 

eustatic forcing 

Ayhan ILGAR 1, *, Wojciech NEMEC 2 , Aynur HAKYEMEZ 1 , Erhan KARAKUŞ 1 

1 Department of Geological Research, General Directorate of Mineral Research and Exploration (MTA), 06520 Ankara, Turkey 

2 Department of Earth Science, Faculty of Mathematics and Natural Sciences, University of Bergen, 5007 Bergen, Norway 

Received: 10.08.2012 Accepted: 19.12.2012 Published Online: 00.00.2013 Printed: 00.00.2013 

Abstract: This sedimentological and sequence-stratigraphic study focuses on the late Miocene deposits in one of the largest peri- 

Mediterranean basins of southern Turkey, the Adana Basin, which formed as a Tauride foreland depression accumulating molasse 

deposits. The Tortonian–Messinian shallow-marine Handere Formation, previously interpreted as a regressive succession, appears to 

have recorded several relative sea-level changes. The formation base recorded a forced regression attributed to the end-Serravalian (Tor- 

1) eustatic fall in sea level. The lower to middle part of the formation is transgressive, culminating in offshore mudstones. The upper part 

is regressive and its 3 isolated conglomeratic members represent sharp-based Gilbert-type deltas with incised fluvial valley-fill deposits, 

recording a forced regression followed by marine reflooding. The time of this regression is biostratigraphically constrained to ~7.8 to 

6.4 Ma B.P. on the basis of planktonic foraminifera in delta bottomset deposits. The regression is attributed to the tectonic conversion 

of the Adana foreland shelf into a piggyback basin, as indicated by seismic sections and compressional basin-margin deformation. The 

reflooding of the basin ~6.4 Ma B.P. is ascribed to a postthrusting flexural subsidence of the foreland under increased crustal load. The 

marine transgression brought an almost immediate evaporitic sedimentation, which suggests invasion of hypersaline Mediterranean 

water. The basin was subsequently emerged and its gypsiferous deposits were extensively eroded due to a second Messinian forced 

regression, attributed to the early evaporative drawdown in the Mediterranean Sea (~6 Ma B.P.). The postorogenic isostatic uplift of 

the Taurides had meanwhile elevated the basin enough to prevent its reflooding by the Zanclean regional transgression. Stratigraphic 

comparison with coeval peri-Mediterranean basins to the west demonstrates that interbasinal correlations are difficult, and that a 

superficial linking of comparable events may be quite misleading. The local timing of the late Miocene relative sea-level changes and the 

landward extent of the Zanclean flooding were apparently determined by the combination of eustasy, tectonics, basin topography, and 

sediment supply, whereby the eustatic signal was modulated and often obscured by local conditions. However, the individual basin-fill 

successions bear a high-resolution record of local events and give unique insights into the local role of tectonics, sediment yield, and 

sea-level changes. 

Key Words: Sedimentology, sequence stratigraphy, Taurides, piggyback basin, Gilbert-type delta, incised valley-fill, Messinian salinity 

crisis, stratigraphic correlation 

1. Introduction 

The late Miocene Mediterranean event known as the 

Messinian salinity crisis was triggered by a glacioeustatic 

sea-level fall combined with the region’s tectonic 

separation from the Atlantic at the latest stages of the 

Alpine orogeny (Hsü et al. 1972; Ryan & Cita 1978; Cita & 

McKenzie 1986; Ryan 2009; Hüsing et al. 2010). The event 

culminated in the evaporitic Lago Mare phase of partial or 

nearly complete desiccation and ended with the Zanclean 

marine flooding, when the Atlantic waters reclaimed the 

Mediterranean Basin. A consentient 3-phase scenario 

for the Messinian event postulates (Roveri & Manzi 

2006): (1) a preevaporitic phase 7.25–5.96 Ma B.P., when 

organic-rich euxinic deposits recorded a significantly 

* Correspondence: ayhan_ilgar@yahoo.com 

reduced circulation of Mediterranean deep waters and 

when microbial stromatolitic limestones formed in some 

peripheral basins; (2) the deposition phase of Lower 

Evaporites 5.96–5.60 Ma B.P., when the precipitation of 

gypsum occurred in shallow-water peripheral basins; and 

(3) the deposition phase of Upper Evaporites 5.60–5.33 

Ma B.P., when the nonmarine Lago Mare environment 

formed in the lowest parts of a desiccating Mediterranean 

Basin. The bulk amplitude of relative sea-level fall is 

estimated at 2000 to 3000 m (Ryan 2009). 

It may thus seem surprising that a regional event 

of such a great magnitude, originally recognised from 

thick evaporites in the centre of the Mediterranean 

Basin, is much less conspicuous at the basin margins, 

1

where stratigraphic correlations of relative sea-level 

changes are difficult and controversial (Ryan 2009). 

One of the most contentious issues is the timing of the 

onsets of hypersalinity and evaporative drawdown in 

the Mediterranean Sea, with direct implications for the 

negative imbalance between the rate of water influx from 

the Atlantic and the regional rate of evaporation. Regional 

studies suggest that the first precipitates at the deep 

bottom of the Mediterranean Basin were preceded by a 

long stepwise advance towards hypersalinity, with gypsum 

in peripheral basins precipitated well before the nominal 

onset of the regional salinity crisis (see review by Ryan 

2009). Most researchers also suggest that the salinity crisis 

was preceded by a considerable early drawdown, with the 

isolation of peripheral basins as evaporating lagoons and 

their eventual emergence (Rouchy 1982; Rouchy & Saint 

Martin 1992; Clauzon et al. 1996; Riding et al. 1999; Soria 

et al. 2005; Maillard & Mauffret 2006; Rouchy & Caruso 

2006; Roveri & Manzi 2006). The early drawdown might 

not exceed 200 m (Dronkert 1985; Krijgsman et al. 1999), 

but would mark a negative water budget and would 

expectedly have a major impact on the peripheral basins 

and their stratigraphy. However, the peri-Mediterranean 

late Miocene stratigraphic record is fuzzy, combining 

relative sea-level changes caused by eustatic and local 

tectonic forcing. 

The diversified tectono-geomorphic conditions 

of peripheral basins resulted in intricate stratigraphic 

successions that are difficult to correlate and also difficult 

to relate to the evaporitic successions in offshore wells. 

Regional correlations are complicated by the fact that 

evaporites are found in only some of the peripheral 

basins, where they may either predate or postdate the 

Mediterranean desiccation (Riding et al. 1999). The 

key indicator of the early evaporative drawdown in the 

peripheral basins might thus be not evaporites, but a 

regional surface of erosion (Ryan 2009). However, neither 

feature can easily be recognised and correlated in the basins 

(Riding et al. 1991, 1998; Roep et al. 1998; Soria et al. 2005; 

Roveri & Manzi 2006). The Messinian surface of subaerial 

erosion is highly irregular due to the varied rates of local 

denudation and it is not marked by any significant climatic 

change. It has been elevated by tectonics and overtaken 

by Plio–Pleistocene erosion in many basins (Glover et al. 

1998; Dilek et al. 1999; Deynoux et al. 2005; Monod et al. 

2006), and it splits into 2 or more erosion surfaces towards 

the deep part of the Mediterranean Basin (Ryan 2009) and 

commonly also landwards in the peripheral basins (Butler 

et al. 1995; Clauzon et al. 1996; Riding et al. 1998; Soria et 

al. 2003). 

The evidence of late Miocene regressions, commonly 

multiple, has been recognised in virtually all peri- 

Mediterranean basins at both active and passive margins 

2 

ILGAR et al. / Turkish J Earth Sci 

(Ryan 2009), but these events are difficult to correlate and 

have been variously attributed to eustasy, high sediment 

supply, or local tectonic uplift (e.g., Clauzon et al. 1996; 

Riding et al. 1999; Karabıyıkoğlu et al. 2000; Larsen 2003; 

Soria et al. 2003, 2005; Deynoux et al. 2005; Flecker et al. 

2005; Roveri & Manzi 2006; Çiner et al. 2008). Regional 

studies have pointed to the importance of local tectonics 

in controlling the late Miocene palaeogeography (Butler 

et al. 1995; Roveri & Manzi 2006). Many areas of the 

Mediterranean were still subject to the final stages of the 

Alpine orogeny at that time, whereas it is generally difficult 

to distinguish between eustatically forced and tectonically 

forced regressions, particularly if both factors were 

potentially involved. The local timing of the late Miocene 

relative sea-level changes and the landward extent of 

the Zanclean marine reflooding were probably both 

determined by the combination of eustasy, local tectonics, 

basin topography, and sediment supply. 

This regional issue is addressed by the present study 

from the Adana Basin at the north-eastern corner of the 

Mediterranean (Figure 1a), where a regression due to 

tectonic inversion of the basin nearly coincided with the 

Messinian evaporative drawdown. The principal aims of 

the study are to: (1) give a palaeontologically constrained 

revised sequence stratigraphy of the Adana Basin, with a 

focus on the late Miocene part of the basin-fill succession; 

(2) assess the role of tectonics and eustasy in forcing the 

Messinian relative sea-level changes in the basin; and (3) 

compare the late Miocene stratigraphy of the Adana Basin 

with that of the adjacent peri-Mediterranean basins in 

order to draw regional implications. 

2. Terminology 

The term “regression” denotes seaward displacement 

of shoreline, resulting in a relative increase of land area 

(Posamentier & Vail 1988; Posamentier et al. 1992). 

Regression reflects the interplay between the relative 

sea-level change (i.e. the available accommodation) 

and the supply of sediment to the shoreline (i.e. the 

accommodation infilling). Their interplay may result in 

a normal or a forced regression (Posamentier et al. 1992; 

Posamentier & Morris 2000). A normal regression signifies 

relative sea-level stillstand or slow rise, with the high 

sediment supply causing seaward shoreline displacement. 

A forced regression signifies a relative sea-level fall, with 

the latter causing seaward shoreline displacement, even if 

the sediment supply to the shoreline is negligible. A forced 

regression may be caused by a eustatic sea-level fall, a 

tectonic uplift, or a coincidental combination of these 2 

factors. 

The basic sequence-stratigraphic terminology used here 

is according to Catuneanu (2006). Stratigraphic sequence 

is a sedimentary succession deposited during a full cycle of

A 

AGE FORMATION 

Holocene 

alluvium 

unconformity 

Pliocene- 

Pleistocene 

alluvial terraces 

unconformity 

Gökkuyu evaporitic mb 

Late Miocene 

Late Serravallianearly 

Tortonian 

Middle Miocene 

Mesozoic 

Handere Fm. 

Karaisalı 

Fm. 

Güvenç 

Fm. 

deltaic members 

Handere Fm. 

conglomerates, sandstones 

mudstones & gypsum 

Kuzgun Fm. 

cong.,sst.& mudst 

unconformity 

Karaisalı Fm. 

reefal limestones, marls 

Güvenç Fm. 

mudstones, sandstones 

unconformity 

Carbonate bedrock 


5 

Muratlı 

Mb. 

7 

Middle 

Miocene 

Middle 

Miocene 

Lower 

Miocene 

Oligocene 

Palaeozoic 

& Mesozoic 

bedrock 

2 

6 

Tepeçaylak 

Mb. 

Taurus Orogenic Belt 

map C 

EXPLANATIONS 

deep-marine clastics Holocene 

fluvial-lagoonal clastics 

Pliocene- 

Pleistocene 

fluvial-lacustrine clastics 

Upper 

Miocene 

lacustrine carbonates 

Middle-Upper 

Miocene 

ophiolitic mélange 

limestones & clastics 

Middle-Upper 

Miocene 

Söğütlü 

Mb. 

alluvial terraces 

fluvial-shallow 

marine clastics 

marine 

marl-limestones 

reefal limestones 

Figure 1. (a) Topographic image of Anatolia (90-m resolution SRTM from Jarvis et al. 2008), showing the location of the Adana Basin 

and other main peri-Mediterranean Miocene basins and major tectonic lineaments referred to in the text. (b) Simplified geological 

map of the southern part of the Adana Basin and the adjacent Mut Basin (based on Şenel 2002 and Ulu 2002); note the study area in 

the former basin (frame) and the location of a late Tortonian delta in the latter basin. (c) Detailed geological map of the study area in 

the Adana Basin. Note the 3 isolated deltaic members of the uppermost Handere Formation. The points 1–7 in maps B and C indicate 

outcrop localities to which the paper’s other figures refer. 

1 

Çakıt river 

Kuzgun 

Salbaş 

3 

5 km 

4 

alluvium 

B 

C 

3

change in accommodation (i.e. its decrease and subsequent 

increase), coupled with sediment supply. The term 

“sedimentary system” denotes a sedimentary environment 

and refers to its specific facies assemblage, whereas a 

systems tract is a succession of such palaeoenvironmental 

facies assemblages. A sequence is considered to be a 

vertical succession of relatively conformable systems 

tracts, bounded by unconformities (erosional surfaces 

of sediment bypass) that grade seawards into correlative 

conformities. A parasequence is a succession of relatively 

conformable deposits bounded by flooding surfaces and 

lacking evidence of a relative base-level fall. 

Following Helland-Hansen (2009), we distinguish 

3 basic types of systems tracts as the building blocks 

of stratigraphic sequences: a forced-regressive systems 

tract (FRST) formed during a relative sea-level fall; a 

transgressive systems tract (TST) formed during a relative 

sea-level rise; and a normal-regressive systems tract 

formed during either highstand (HST) or lowstand (LST) 

and recording sea-level stability or minor relative rise. 

The basis for distinguishing systems tracts is the vertical 

stacking of sedimentary facies assemblages and the 

stratigraphic palaeoshoreline trajectory (Helland-Hansen 

& Martinsen 1996). A FRST has a falling trajectory, but 

the regressive shoreline shift may involve deposition or be 

fully erosional, depending on the sediment supply rate and 

the rate and magnitude of relative sea-level fall (Plint 1988; 

Helland-Hansen & Gjelberg 1994). With the sea-level fall 

compensated by tectonic subsidence, some sequences, 

referred to as the sequences of type 2 (Jervey 1988), may 

show no recognisable shoreline fall and masquerade as 

parasequences (e.g., Ghinassi 2007; Messina et al. 2007). 

Parasequences consist of a TST overlain by a HST. 

The descriptive sedimentological terminology used 

in this study is according to Harms et al. (1975, 1982) 

and Collinson and Thompson (1982). In biostratigraphic 

analysis, the Mediterranean planktonic foraminifer zones 

of Iaccarino et al. (2007) are followed, and the definition 

of species is based mainly on Kennett and Srinivasan 

(1973), Iaccarino (1985), and Bolli and Saunders (1985). 

Biostratigraphic age estimates refer to the astronomically 

tuned ATNTS2004 scale (Lourens et al. 2004). 

3. Regional geological setting 

The Tauride orogen of southern Turkey is the youngest of 

the eastern peri-Mediterranean Alpine mountain chains. 

It is arbitrarily divided into 3 segments (Figure 1a): the 

Western Taurides, west of the Isparta Angle, passing 

westwards into the Hellenides and sometimes referred to 

as the Eastern Hellenides due to their tectonic link with 

the Hellenic subduction arc; the Central Taurides between 

the Isparta Angle and the Ecemiş Fault to the east; and 

the Eastern Taurides that pass eastwards into the Zagros 

Mountains. The orogeny culminated at the end of the 

4 


Eocene (Şengör 1987; Clark & Robertson 2002), but lowrate 

plate convergence in the Central Taurides persisted 

until the mid-Oligocene (Kelling et al. 1987; Andrew 

& Robertson 2002), when the Cyprian subduction arc 

eventually stepped back to the south of Cyprus (Figure 1a). 

Orogenic deformation proceeded until the late Miocene 

in the Eastern Taurides, where the Misis Structural High 

popped up by folding and thrusting (Michard et al. 1984; 

Aktaş & Robertson 1990; Dilek & Moores 1990; Yılmaz 

1993; Yılmaz et al. 1993; Robertson 2000; Sunal & Tüysüz 

2002), and also at the transition of the Western and 

Central Taurides, where the Lycian and Hoyran-Hadım 

nappe fronts collided north of the Isparta Angle (Collins 

& Robertson 1998, 2000; Poisson et al. 2003; Sagular & 

Görmüş 2006). The Miocene thus saw the last stages of 

localised compressional deformation, while the Taurides 

in general had already become subject to postorogenic 

isostatic uplift and crustal extension with the development 

of orogen-collapse basins (Seyitoğlu & Scott 1991, 1996; 

Jaffey & Robertson 2005; Bartol et al. 2011; Koç et al. 2011; 

Cosentino et al. 2012). 

The peri-Mediterranean basins in southern Turkey 

formed nonsynchronously during the early Miocene and 

ranged from relatively simple intramontane grabens or 

half-grabens (Alçiçek et al. 2005; Alçiçek 2010) to more 

complex extensional depressions (Flecker et al. 1995, 

2005; Larsen 2003; Şafak et al. 2005; Çiner et al. 2008), 

strike-slip pull-apart features (Ilgar & Nemec 2005), and 

compressional foreland troughs (Hayward 1984a, 1984b; 

Burton-Ferguson et al. 2005; Alçiçek & Ten Veen 2008). 

The basin-fill successions of these isolated molasse basins 

are highly diversified in terms of sedimentary facies and 

sequence stratigraphy, and are difficult to correlate (Tekeli 

& Göncüoğlu 1984; Yetiş et al. 1995; Durand et al. 1999; 

Bozkurt et al. 2000; Kelling et al. 2005). However, they 

provide crucial information on an early postorogenic 

tectono-geomorphic evolution of the Tauride belt and its 

interaction with the Mediterranean Sea. As pointed out 

by Kelling et al. (2005, pp. 1–13), the palaeogeographical 

and chronostratigraphical resolution of the local basin-fill 

successions far exceeds that of geophysical lithospheric 

models and gives unique regional insights into the relative 

role of tectonics, climate, sediment yield, and sea-level 

changes. Detailed palaeogeographical reconstructions and 

the recognition of major sediment-transfer fairways to 

the offshore zone (Satur et al. 2005) are vital to regional 

hydrocarbon prospecting (Görür & Tüysüz 2001). 

The Adana Basin is one of the largest Miocene 

peripheral basins in southern Turkey, located between 

the Taurus orogenic front to the north-west and the Misis 

Structural High to the south-east (Figure 1a). The SWtrending 

basin passes offshore into the Cilicia Basin north 

of Cyprus. The Adana Basin and its smaller counterpart, 

İskenderun Basin on the other side of the Misis High,

form the Çukurova Basin Complex at the Kahramanmaraş 

junction of the Afro-Arabian, Anatolian, and Eurasian 

plates (Ünlügenç et al. 1990). The structural development 

in this region involved 3 major tectonic lineaments 

(Figure 1a): the Bitlis-Zagros Suture Zone separating the 

Arabian and Anatolian-Eurasian plates; the eastern arm 

of the Cyprian arc of intra-Tethyan plate subduction; and 

the sinistral strike-slip Dead Sea Fault between Africa and 

Arabia, passing to the north-east into the East Anatolian 

Fault (Kelling et al. 1987; Ünlügenç et al. 1990; Williams et 

al. 1995; Robertson 2000). The system of the East Anatolian 

and North Anatolian faults lead the neotectonic westward 

“expulsion” of the compound Anatolian craton (Dewey & 

Şengör 1979; Şengör & Yılmaz 1981). Derivatives of this 

neotectonic strike-slip system include the Burdur-Fethiye- 

Pliny Fault to the west and the Ecemiş Fault separating the 

Adana Basin from the coeval Mut Basin (Figure 1a). 

The Adana Basin formed in the early Miocene on 

a wedge-shaped sliver of the Tethyan shelf that was 

structurally entrapped between the Anatolian and Arabian 

plates and converted into the local Tauride foreland. Seismic 

interpretation by Burton-Ferguson et al. (2005) suggests 

that the Adana foreland developed by flexural subsidence 

under the load of a SE-advancing orogenic front and then 

turned into a piggyback basin in the Tortonian, with the 

Misis High pop-up ridge separating it from the İskenderun 

foredeep to the south-east (Figures 1a and 1b). The foreland 

model explains the Miocene strong subsidence and great 

thickness of sediments accumulated in the basin as well as 

the basin’s late Miocene compressional tectonic inversion. 

4. Dynamic stratigraphy of the Adana Basin 

The stratigraphy of the Adana Basin was established by 

Schmidt (1961) and refined by subsequent studies (Yalçın 

& Görür 1984; Kelling et al. 1987; Yetiş 1988; Ünlügenç 

et al. 1990; Görür 1992; Yetiş et al. 1995; Nazik 2004; 

Satur et al. 2005). The present study contributes further 

to this topic. The basin-fill succession comprises up to 6 

km of Miocene to Quaternary siliciclastic and calcareous 

deposits. Bedrock consists of Palaeozoic and Mesozoic 

sedimentary rocks, which include Devonian coralline 

limestones and sandstones, Permo-Carbonifereous 

limestones, a Late Triassic to Cretaceous thick carbonate 

platform, and Late Cretaceous turbidites. These rocks were 

postdated by the tectonic emplacement of a nappe of Late 

Cretaceous ophiolitic mélange (Figure 1b). 

Sedimentation in the basin commenced in the early 

Miocene with deposition of the alluvial fan redbeds of the 

Gildirli Formation (Figure 2), including conglomerates, 

sandstones, and mudstones. The Burdigalian to early 

Langhian Kaplankaya Formation recorded the first 

episode of marine sedimentation in the basin, with 

sandstones, siltstones, marlstones, and sandy limestones. 


This transgressive formation has a broader lateral extent, 

particularly northwards, and unconformably overlies 

bedrock palaeotopography. Reefal limestones formed 

in the marginal zone of the basin, while open-marine 

deep neritic conditions prevailed in the basin interior. 

The Kaplankaya Formation thus passes laterally into and 

is partly overlain by the late Burdigalian–Serravalian 

reefal Karaisalı Formation, whose basinal equivalents 

are sublittoral tempestitic sandstones of the Cingöz 

Formation and offshore mudstones of the Güvenç 

Formation (Figure 2). There is also evidence of stormgenerated 

erosive turbidity currents transferring abundant 

sand across the shelf edge to the deep-water realm of the 

adjoining Cilicia Basin (Satur et al. 2005). The reefal and 

coeval nearshore to offshore deposits show an overall 

shallowing-upwards trend, with the upper part of the 

Güvenç Formation increasingly richer in sandstones 

(Figure 2). 

The marine sedimentation was interrupted when 

the basin emerged due to a relative sea-level fall at the 

end of Serravalian (Figure 2). River valleys were incised 

and then filled with the fluvial deposits of the Kuzgun 

Formation, as the basin was subsequently reflooded due 

to an early Tortonian relative sea-level rise. A transgressive 

ravinement surface with a lag of wave-worked oysterbearing 

gravel marks the marine reflooding. The 

transgression initiated shallow-marine sedimentation 

with a second generation of reefal limestones along 

the basin margin, the Tırtar Formation, superimposed 

directly on the older limestones of the Karaisalı Formation 

(Figure 2). The coeval Handere Formation in the basin 

interior consists of shoreface sandstones that pass upward 

into finer-grained sandstones, siltstones, and mudstones of 

an offshore-transition environment and further into thick 

offshore mudstones (Figure 2). The deposition of offshore 

mudstones in the upper part of the Handere Formation 

marked the maximum marine flooding in the basin, 

reached in the late Tortonian. 

These transgressive deposits are sharply overlain by 

the latest Tortonian–Messinian regressive deposits of 

the uppermost Handere Formation (Figure 2), which 

comprise shallow-marine sandstones and siltstones 

and include 3 isolated conglomeratic members (see the 

Muratlı, Tepeçaylak, and Sögütlü members in Figure 1c). 

These conglomeratic deposits, previously interpreted 

as fluvial, are the main topic of the present study, which 

documents them as sharp-based deltas with associated 

incised fluvial valley-fills. There are also erosional relics 

of the uppermost gypsiferous Gökkuyu Member of the 

Handere Formation preserved in the southern part of 

the basin (Figures 1c and 2). The evaporites overlie both 

the deltaic conglomeratic members and offshore clastic 

deposits of the Handere Formation. However, there is no 

evidence of the Zanclean regional marine transgression in 

5

6 

0 

1 

2 

3 

4 

5 

6 

7 

8 

9 

10 

11 

12 

13 

14 

15 

16 

17 

18 

19 

20 

21 

22 

23 

Chronostratigraphy (Ma) 

Epoch Stage 

Holocene Tyrrhenian 

Ionian 

Pleistocene 

Pliocene 

Early Late 

Miocene 

Early Middle 

Late 

Calabrian 

1.81 

Gelasian 

2.59 

Piacenzian 

3.60 

Zanclean 

5.33 

Messinian 

7.25 

Tortonian 

11.61 

Serravallian 

13.65 

Langhian 

15.97 

Burdigalian 

20.43 

Aquitanian 

Plankt. 

foram 

zones 

Mple2 

Mple1 

Mpl6 

Mpl5 

Mpl4 

Mpl3 

Mpl2 

Mpl1 

MMi13 

MMi12 

MMi11 


Nannofossil 

zones 

MNN21 

MNN20 

MNN19 

MNN18 

MNN17 

MNN16 

14-15 

MNN13 

MNN12 

Nondistinctive Zone 

MMi10 

MMi9 

MMi8 

c 

MMi7 b 

a 

MMi6 

MMi5 

MMi4 

MMi3 

MMi2 

MMi1 

c 

b 

a 

b 

a 

MNN11 

MNN10 

MNN9 

MNN8 

MNN7 

MNN6 

MNN5 

MNN4 

MNN3 

MNN2 

MNN1 

c 

b 

a 

b 

a 

b 

a 

c 

b 

a 

Tırtar 

Fm. 

Muratlı 

deltaic mb 

Karaisalı 

Fm. 

ADANA BASIN 

Gökkuyu 

evaporitic mb 

Sequence 

stratigraphy 

?HST + FRST 

+ LST 

Figure 2. Revised stratigraphy of the Adana Basin and its interpretation in terms of systems tracts. The letter code is as used in the text 

(terminology after Helland-Hansen 2009): LST – normal-regressive lowstand systems tract; TST – transgressive systems tract; HST – 

normal-regressive highstand systems tract; and FRST – forced-regressive systems tract. 

Cingöz 

Fm. 

Kaplankaya Fm. 

MESOZOIC 

BEDROCK 

v v v v v v v 

v v v v v v v 

v v v v v 

v v v 

Handere Fm. 

Kuzgun Fm. 

Güvenç Fm. 

Gildirli Fm. 

Post - Miocene 

alluvium 

(fluvial terraces) 

FRST 

HST 

TST 

TST 

FRST + LST 

HST 

TST 

LST

the basin. The post-Miocene deposits are fluvial terraces of 

coarse-grained alluvium with caliche. 

The Messinian gypsum evaporites in the basin are up to 

a few metres thick (Figure 3a) and their main varieties range 

from massive to laminated and enterolithic (Figure 3b), 

micro- to coarse-crystalline (Figures 3c and 3d), nodular 

(Figure 3e), and wave-worked gypsarenitic (Figure 3f). 

X-ray diffraction data show that selenitic gypsum is the 

sole evaporitic mineral (Karakuş 2011). X-ray fluorescence 

analyses of major oxide composition indicate that the 

evaporitic precipitation occurred in homogeneous 

hydrochemical conditions, with a Sr-signature similar to 

the Messinian evaporites in adjacent peri-Mediterranean 

basins (Karakuş 2011). 

5. Facies architecture of deltaic members 

The study focuses on the 3 isolated conglomeratic members 

of the Handere Formation: the Muratlı, Tepeçaylak, and 

Sögütlü members (Figure 1c). They have been exhumed by 

Quaternary erosion and are laterally surrounded by sparsely 

preserved shallow-marine sandstones of the Handere 

Formation, with which they sharply overlie the formation’s 

offshore mudstones (Figure 2). These conglomeratic 

members consist of 2 main facies associations, Gilberttype 

deltaic deposits and fluvial incised valley-fill deposits, 

which are described and interpreted in the present section. 

5.1. Gilbert-type delta deposits 

These conglomeratic deposits are generally well-bedded, 

forming clinoformal wedges that are stacked basinwards 

in a downstepping pattern and are up to 15–40 m thick. 

The clinoformal architecture consists of inclined foreset 

beds overlain by horizontal topset beds (Figure 4a), as 

is generally characteristic of Gilbert-type deltas (Barrell 

1912; Colella 1988; Postma 1990). 

5.1.1. Foreset facies 

Foreset deposits are conglomerate and subordinate 

sandstone beds inclined basinwards at up to 25° 

(Figure 4b). They show an overall coarsening-upwards 

trend and comprise facies commonly reported from 

Gilbert-type deltas (Postma 1984; Nemec et al. 1999; 

Lønne & Nemec 2004; Breda et al. 2007). Some outcrops 

show the foreset beds passing tangentially downdip into 

the gently inclined and finer-grained beds of delta toeset 

(Figure 4a). Conglomerate beds are 10–75 cm thick, but 

mainly 15–35 cm, and are tabular to mound-shaped. They 

consist of granule to coarse-pebble gravel and occasionally 

contain scattered cobbles of up to 15 cm in size. The gravel 

is subrounded to rounded and has mainly a clast-supported 

texture (Figure 4c). Clasts are derived from the bedrock 

Mesozoic limestones and serpentinite, and from the basinmargin 

Miocene reefal limestones. The matrix is moderately 

well-sorted sand with granules. Many conglomerate 

beds and the majority of associated sandstone beds 


show planar-parallel stratification (Figure 4b) indicating 

tractional deposition from fully turbulent hyperpycnal 

flows (Bornhold & Prior 1990; Nemec 1990), which means 

river-generated low-density turbidity currents (sensu 

Lowe 1982). Massive conglomerate beds are nongraded or 

inversely graded (Figure 4d), tabular in dip section, and 

mound-shaped lenticular in strike section, interpreted to 

be deposits of cohesionless debris flows (Nemec & Steel 

1984; Nemec 1990). 

Sandstones predominate in the down-dip part of the 

foreset and at its toe, forming tabular or wedge-shaped 

beds that are 5–30 cm thick, composed of very coarse- to 

fine-grained sand and alternating with thin beds of granule 

conglomerate (Figure 4a). Common are scattered pebbles 

of up to 5 cm in size. The sandstone beds show planarparallel 

stratification with or without normal grading 

and are often capped by current-ripple cross-lamination 

with down-dip transport direction. Some of the foreset 

beds, up to 40 cm thick, are isolated backsets of up-slope 

dipping cross-strata composed of coarse sand and/or finepebble 

gravel (Figure 4e). They occur on the stoss side of 

mound-shaped massive conglomerate bodies (debris-flow 

deposits) or as the infill of trough-shaped scours (deltaslope 

chutes). The backsets indicate tractional deposition 

by low-density turbidity currents subject to hydraulic jump 

(Nemec 1990). There are also sporadic slump deposits of 

variable scale and thickness (Figure 4f). 

5.1.2. Topset facies 

The delta topset deposits (Figures 4a and 5) are pebble 

conglomerates and coarse-grained sandstones. Their 

fining-upwards bedsets, 60–140 cm thick, have erosional 

bases and are commonly stacked on top of one another, 

apparently representing multistorey palaeochannels of 

braided streams a few metres wide (Collinson 1996; Miall 

1996). Their laterally discontinuous basal layers of coarse 

clast-supported conglomerate are thought to be channelfloor 

lag deposits (Miall 1985; Nemec & Postma 1993). 

Planar parallel-stratified and cross-stratified beds, 10–45 

cm thick, are interpreted, respectively, to be deposits of 

longitudinal and transverse or oblique midchannel bars 

(Miall 1985; Nemec & Postma 1993). 

The deltas suffered erosion and their topset deposits 

are inconsistently preserved, generally better in the 

upstream part (Figure 4a). The sparser downstream 

preservation of delta topset may be due to a negative 

subaerial accommodation (Bhattacharya & Willis 2001), 

or to removal by post-Miocene erosion (Figure 2). The 

relationship of the delta topset to the foreset is invariably 

oblique (erosional), which supports the notion of a falling 

delta-shoreline trajectory (Breda et al. 2007, 2009) based 

on evidence that the horizontal topset was incrementally 

stepping down in the basinward direction, as discussed in 

the next section. 

7

5.1.3. Bottomset facies 

The basal deltaic facies are only locally exposed, but are 

generally similar. The gently inclined delta-toe deposits 

(Figures 4a and 6) are thinly bedded, fine-grained 

sandstones and siltstones with plane-parallel stratification 

and minor ripple cross-lamination. Delta bottomset 

consists of thin siltstone and sandstone beds intercalated 

with mudstones (Figure 6). The microfauna content of 

these deposits is described in a subsequent section. 

5.2. Incised valley-fill deposits 

This facies assemblage is exposed in a chance crosscut 

section in the proximal part of the Muratlı Member 

(Figures 1c and 5), but similar unexposed deposits 

presumably also occur in the 2 other coeval deltaic members 

8 


A B 

C 

1 m 

E F 

Figure 3. Messinian evaporites in the Adana Basin. (a) An example outcrop of the evaporites at 

the top of Handere Formation. (b) Enterolithic gypsum. (c) Crystalline gypsum with a chevron 

growth structure. (d) Crystalline gypsum with a grassy growth structure. (e) Nodular gypsum. 

(f) Gypsarenite with wave-ripple cross-lamination. The coin (scale) is 2 cm. 

D 

of the Handere Formation, where no similar transverse 

sections are available. These deposits are recognisably 

coarser-grained, comprising pebble conglomerates and 

subordinate coarse-grained sandstones with scattered 

cobbles and boulders (Figure 5). Conglomerates are clastsupported, 

with a matrix of medium to coarse sand and 

granules. Gravel clasts are moderately sorted, mainly 

subangular to subrounded, and of the same provenance 

as the delta gravel. Scattered boulders are derived from 

the basin-margin Miocene reefal limestones. The deposits 

form erosionally based, vertically stacked fining-upwards 

bedsets (Figure 5) interpreted to be multistorey braidedstream 

palaeochannels (Collinson 1996; Miall 1996), 

mainly 0.8–1.6 m thick and 10–25 m wide. Coarse gravelly 

channel-floor lags are poorly developed, but solitary

delta topset 

delta foreset 


B C 

E F 

foreset 

backset 

Figure 4. Outcrop details of the Muratlı delta, Adana Basin. (a) Longitudinal outcrop section showing an 

erosional angular contact between the delta’s foreset and topset deposits. (b) Conglomeratic foreset deposits 

comprising planar parallel-stratified and massive beds. Note the rapid upward increase in the bedding 

inclination. (c) Close-up detail of the delta topset, showing submature gravel composed of bedrock limestone 

and serpentinite clasts mixed with large fragments of Miocene corals and reefal limestones. (d) Delta foreset 

deposits including massive, inversely graded, and nongraded conglomerate beds. (e) Delta foreset detail showing 

a backset of upslope-dipping gravelly cross-strata. (f) Slump deposits within the delta foreset. Picture A is from 

locality 1, pictures B–E from locality 2, and picture F from locality 3 in Figure 1c. 

D 

coral fragment 

A 

9

10 

° 

or multiple cross-strata sets, 25–60 cm thick, indicate 

transverse or oblique midchannel bars (Miall 1985). 

This contrasting facies assemblage in the Muratlı 

Member forms the infill of an axial valley that was deeply 

incised in the delta. The top part of the valley-fill and the 

surrounding host delta are not preserved in the outcrop 

section, but the measured depth of incision is at least 

15 m and the valley width is up to 60 m. The incised 

valley is clearly not an integral part of the prograding 

delta’s topset (see discussion by Hampson et al. 1997) and 

instead indicates erosional cannibalisation of the delta by 

a deep incision of its fluvial feeder system in response to 

pronounced base-level fall (see Mellere et al. 2002; Ilgar 

& Nemec 2005; Breda et al. 2009). The high depth/width 

ratio of the valley and the scattered boulders suggest 

relatively rapid incision, with a minimal lateral shifting 

of the fluvial system (see Yoxall 1969; Wood et al. 1993) 

and with the stream competence significantly increased by 

the topographic confinement (Schumm 1993). The valley 

incision seems to have occurred concurrently with the late 

stages of the delta progradation, when the entrenching 

fluvial system acted as a feeder for the youngest telescoping 

parts of the delta (Figure 7a). The down-stepping pattern 

of delta topset (Figure 7b) strongly supports the notion of 

an incremental fall of the delta shoreline trajectory. 

The relatively narrow valley filled with fluvial deposits 

indicates that the infilling of the valley was relatively rapid, 

under a high rate of sediment supply, with little lateral 

wandering of the river and no significant valley-side 

collapses at the sea-level lowstand stage. 


Figure 5. An oblique transverse section through the Muratlı delta, Adana Basin, showing a gravel-filled axial fluvial valley deeply 

incised in the delta deposits. Palaeotransport direction is away from the viewer, obliquely to the right. Picture from locality 2 in 

Figure 1c. 

6. Sequence-stratigraphic interpretation 

The late Tortonian shoreline of the basin is represented 

by the reefal limestones of the Tırtar Formation, which 

were superimposed directly on the earlier basin-margin 

limestones of the Karaisalı Formation (Figure 2) and were 

later extensively eroded (Figure 1c). The 3 deltas in their 

location appear to have been offset basinwards by ~25 km 

with respect to the late Tortonian shoreline and emplaced 

directly onto the offshore mudstones of the Handere 

Formation, which indicates a forced-regressive erosional 

shift of the shoreline. 

A forced regression is indicated by the downstepping 

geometry of the delta topset (Figure 7b), the clinoformal 

foreset wedges that sharply downlapped the basin floor 

(Figures 4a and 6), and further by the incision of fluvial 

valley along the delta axis (Figures 5 and 7a). The sharp, 

erosional basal surface of the delta marks an abrupt 

facies change and passes basinwards into a correlative 

depositional conformity (Figure 7a). The aggradational 

infilling of the incised valley documents a subsequent rise 

of relative sea level, and the overlying gypsiferous deposits 

indicate a brief drowning of the deltas. The evaporites 

occur as erosional relics of the latest marine deposits 

(HST) in the basin, which implies yet another subsequent 

forced regression (see the late Messinian erosional FRST 

in Figure 2). 

The Tortonian–Messinian deposits of the Kuzgun 

and Handere formations (Figure 2) have previously been 

interpreted as a simple regressive succession (Yetiş 1988; 

Yetiş et al. 1995), which would imply a normal-regressive

Figure 6. Simplified vertical profile of the bottomset part of the 

one of the Muratlı delta wedges, showing the pattern of sediment 

sampling for biostratigraphic analysis. The dots numbered 1–24 

indicate the location of samples. Log from locality 2 in Figure 1c. 

HST. The present study indicates that this sedimentary 

succession, in reality, bears a high-resolution record of 

several major sea-level changes in the basin and comprises 

2 stratigraphic sequences bounded by the erosional 

surfaces of forced regression. The following interpretive 

stratigraphic scenario is suggested (Figure 2): 

• The Kuzgun Formation represents a forced regression 

that is recognisable around the Mediterranean and 

attributed to the end-Serravalian (Tor-1) eustatic fall in sea 

level (Haq et al. 1988; Haq 1991); the formation’s erosional 

basal part is an erosional FRST, whereas the bulk of the 

formation comprises a LST and possibly the earliest TST. 

• The main Tortonian part of the overlying Handere 

Formation and the coeval Tırtar Formation (basin-margin 

reefal platform) constitute a TST, recording the subsequent 

eustatic sea-level rise (Haq et al. 1988; Haq 1991). Basin 

subsidence may have enhanced this marine transgression. 

• The upper part of the Handere Formation, with 

its isolated deltaic members, might have commenced 

its deposition as a normal-regressive HST, but there is 


no facies evidence to support this notion. The lack of a 

recognisable HST suggests that the transgression was 

interrupted by a relative sea-level fall, which would mean 

a TST overlain directly by a depositional FRST. This 

stratigraphic configuration of systems tracts may indicate 

a eustatic transgression terminated by a tectonically forced 

stepwise regression. 

• The sharp base of the deltaic members and coeval 

littoral deposits of the Handere Formation (Figure 7a) is 

a regressive surface of marine erosion (Plint 1988; Plint 

& Nummedal 2000), expectedly passing basinwards into 

a correlative conformity (MacEachern et al. 1999). This 

surface was developing incrementally during the entire 

time of the stepwise relative sea-level fall and hence is 

probably diachronous (Embry 2002). 

• The sharp-based gravelly deltas and coeval shallowmarine 

deposits would thus represent a depositional forced 

regression (Plint 1988; Helland-Hansen & Gjelberg 1994; 

Plint & Nummedal 2000). The basinward advance of the 

deltas (FRST and LST) was followed by a relative sea-level 

rise (TST), when the incised valleys were filled with fluvial 

deposits and the deltas were shallowly drowned with an 

abrupt landward shift of the shoreline and river outlets. 

• The marine transgression brought an almost 

immediate deposition of evaporites (HST), which suggests 

flooding by hypersaline sea water (Figure 7b). The Adana 

Basin was subsequently emerged and its gypsiferous 

deposits were extensively eroded due to another forced 

regression (the late Messinian FRST, Figure 2). The 

evaporitic Gökkuyu Member of the uppermost Handere 

Formation (Figure 2) is sparsely preserved in the northern 

part of the basin, but its thickness reaches a few hundred 

metres in the southern part and exceeds 1 km in the 

adjoining inner Cilicia Basin (Aksu et al. 2005; Burton- 

Ferguson et al. 2005). It can thus be precluded that these 

evaporates are local deposits, formed in an isolated coastal 

lagoon or sabkha. 

It would then appear that 2 consecutive forced 

regressions occurred in the Adana Basin in the Messinian 

time, the first depositional and possibly forced by tectonics 

and the next erosional and probably eustatic. The key issue 

in this hypothetical scenario is the exact timing and actual 

causes of the 2 regressions, which apparently followed 

each other closely. 

7. Biostratigraphic dating 

The heterolithic, fine-grained bottomset deposits of the 

best-exposed Muratlı delta have been systematically 

sampled (Figure 6) to estimate the time of the delta 

progradation. The samples of mudstone and silty mudstone 

layers were disaggregated in a 30% hydrogen peroxide 

solution and washed over 63-, 125-, and 425-µm sieves. 

In total, 24 samples have been analysed (Figure 6), with a 

focus on planktonic foraminifera. 

11

Planktonic foraminifera occur throughout the studied 

section, except for 3 samples from its uppermost part 

(samples 18, 21, and 22 in Figure 6). The abundance of 

foraminifer assemblages varies from medium to low, 

whereas their diversity and degree of preservation are 

generally moderate to high. The lowermost part of the 

section (samples 1–5) lacks age-diagnostic species and 

12 

A 

B 

reefal limestone 

Surface of forced 

regression 

Relic evaporitic cover 

on terminal delta slope 

Messinian evaporites 

topset 

delta wedge 


Delta progradation direction 

step-down 

topset 

25 km 

step-down 

delta wedge 

sea-level fall 

topset 

delta wedge 

offshore 

mudstones 

Figure 7. (a) Schematic model for the downstepping forced-regressive progradation of the Muratlı delta with concurrent 

incision of an axial fluvial valley; diagram not to scale. (b) Panoramic view of the Muratlı delta longitudinal outcrop showing 

the downstepping pattern of clinoformal foreset wedges towards the left (SSE direction); picture from locality 2 in Figure 1c. 

200 m 

shows an assemblage of benthic forams (with Ammonia sp. 

and Elphidium spp.), echinid spines, gastropods, and rare 

specimens of Globoturborotalita, Globigerina, Orbulina, 

Globigerinella, and Globigerinoides. 

Planktonic foraminifer assemblages are most 

diversified in the middle part of the section (samples 

6–16, Figure 6), including Globorotalia suterae Catalano

a1 

d2 

i1 

a2 


i2 

l1 l2 

r 

o 

a3 

d3 e 

s 

m2 

Figure 8. Selected planktonic foraminifera identified in the latest Miocene delta bottomset deposits in the Adana and Mut basins: (a1) Turborotalita multiloba (Romeo) in 

spiral, (a2) umbilical, and (a3) side view, sample 20 from Adana Basin; (b) Turborotalita quinqueloba (Natland) in umbilical view, sample 9 from Adana Basin; (c) Tenuilellinata 

angustiumbilicata (Bolli) in umbilical view, sample 5 from Mut Basin; (d1) Neogloboquadrina acostaensis (Blow) in umbilical view and (d3) in spiral view, sample 2 from Mut 

Basin; (d2) Neogloboquadrina acostaensis (Blow) in umbilical view, sample 6 from Adana Basin; (e) Globigerinita glutinata (Egger) in umbilical view, sample 4 from Mut Basin; 

(f) Globigerinita uvula (Ehrenberg) in side view, sample 2 from Mut Basin; (g) Catapsydrax parvulus Bolli in umbilical view, sample 4 from Mut Basin; (h) Globigerinoides 

bollii Blow in umbilical view, sample 2 from Mut Basin; (i1) Neogloboquadrina continuosa (Blow) in spiral and (i2) umbilical view, samples 4 and 2 from Mut Basin; (j1) 

Globoturborotalita woodi (Jenkins) in spiral and (j2) umbilical view, sample 13 from Adana Basin; (k) Globoturborotalita apertura (Cushman) in umbilical view, sample 13 from 

Adana Basin; (l1) Neogloboquadrina humerosa (Takayanagi & Saito) in spiral and (l2) umbilical view, samples 11 and 9 from Adana Basin; (m1) Globorotalia suterae Catalano 

and Sprovieri in oblique and (m2) spiral view, sample 24 from Adana Basin; (n) Globoturborotalita decoraperta (Takayanagi & Saito) in spiral view, sample 24 from Adana 

Basin; (o) Orbulina universa d’Orbigny, sample 6 from Adana Basin; (p) Globigerinoides bulloideus Crescenti in spiral view, sample 1 from Mut Basin; (q) Globigerina bulloides 

d’Orbigny in spiral view, sample 6 from Adana Basin; (r) Globigerinella obesa (Bolli) in umbilical view, sample 13 from Adana Basin; (s) Orbulina suturalis Brönnimann, 

sample 9 from Adana Basin; (t) Globigerinella siphonifera (d’Orbigny) in side view, sample 6 from Adana Basin; and (u) Globigerinoides seigliei Bermudez and Bolli in spiral 

view, sample 3 from Mut Basin. The scale bar is 75 µm in pictures a–d and 100 µm in pictures e–u. 

j1 

m1 

b 

t 

f 

j2 

p 

n 

c 

g 

u 

k 

q 

d1 

h 

13

14 

Age (Ma) 

6 6.033 

7 

8 

9 

10 

11 

ATNTS2004 

C3A 

C3B 

and Sprovieri, Globoturborotalita apertura (Cushman), G. 

nepenthes (Todd), G. decoraperta (Takayanagi & Saito), G. 

woodi (Jenkins), Globigerinelloides bulloideus Crescenti, 

G. obliquus Bolli, G. trilobus (Reuss), G. quadrilobatus 

(d’Orbigny), Globigerina bulloides d’Orbigny, G. 

falconensis Blow, Globigerinella siphonifera (d’Orbigny), 

G. obesa (Bolli), Orbulina suturalis Brönnimann, O. 

universa d’Orbigny, Neogloboquadrina acostaensis Blow, N. 

humerosa Takayanagi & Saito, Dentoglobigerina altispira 

altispira (Cushman & Jarvis), Turborotalita quinqueloba 

(Natland), Tenuitellinata angustiumbilicata (Bolli), and 

Globigerinita uvula (Ehrenberg) (Figure 8). The presence 

of Globorotalia suterae Catalano & Sprovieri is particularly 

important, because its first stratigraphic occurrence 

serves to identify the base of subzone MMi12b (i.e. the 

Globorotalia suterae subzone of Iaccarino 1985) and is 

astronomically dated to 7.81 Ma B.P. (Figure 9; Sprovieri 

et al. 1999; Lourens et al. 2004; Iaccarino et al. 2007). 

The coeval occurrence of Globorotalia suterae Catalano 

& Sprovieri, Globigerinoides bulloideus Crescenti, and 

Globoturborotalita woodi (Jenkins) allows this part of 

the section to be assigned to undifferentiated biozone 

MMi12b–MMi13a (Figure 9; D’Onofrio et al. 1975; 

Iaccarino 1985; Iaccarino et al. 2007). The lack of the 

Globorotalia miotumida group (particularly Globorotalia 

Polarity 

Magneto 

zones 

C3 

7.140 

7.528 

C4 

8.699 

C4A 

9.779 

C5 

Period 


Chronostratigraphy 

N E O G E N E 

M I O C E N E 

Epoch 

L A T E 

Stage 

Messinian 

7.246 

Tortonian 

Mediterranean 

planktonic 

foraminiferal 

biostratigraphy 

Nondistinctive 

Zone 

MMi11 

MMi9 

Figure 9. Mediterranean planktonic foraminiferal biozones plotted against the 

ATNTS2004 magnetic chronostratigraphy, with the stratigraphic distribution of 

diagnostic species in the outcrop section sampled in the Adana Basin (Figure 6). 

MMi12 MMi13 

MMi10 

c 

b 

a 

b 

a 

Globoturborotalita woodi 

Globigerinoides bulloideus 

Globorotalia suterae 

Neogloboquadrina humerosa 

Globigerinella siphonifera 

conomiozea Kennett, G. sahelina Catalano & Sprovieri, 

G. mediterranea Catalano & Sprovieri, and G. miotumida 

Jenkins) in the samples does not permit a more precise 

assignment. Other fauna in this part of the section includes 

scarce Bulimina echinata d’Orbigny, representatives 

of Uvigerina and Bulimina, miliolids, echinid spines, 

gastropods, and bivalves. 

Sample 18 (Figure 6) bears no planktonic foraminifers, 

but contains some benthic forms, gastropods, bivalves, 

and other shell fragments. The slightly higher sample 20 

(Figure 6) shows a planktonic foraminifer assemblage 

that consists almost entirely of Turborotalita multiloba 

(Romeo), T. quinqueloba (Natland), and Tenuitellinata 

angustiumbilicata (Bolli), accompanied by Bulimina 

echinata d’Orbigny, miliolids, and echinid spines. These 

foram species are known to have been most common prior 

to the deposition of the Messinian Lower Evaporites in the 

Mediterranean Sea (Kouwenhoven et al. 2006; Manzi et 

al. 2007; Morigi et al. 2007; Iaccarino et al. 2008; Orszag- 

Sperber et al. 2009; Di Stefano et al. 2010). Furthermore, 

Turborotalita multiloba (Romeo) is an important marker 

species in the Mediterranean, because its narrow 

stratigraphic range shortly predates the deposition of the 

Lower Evaporites (D’Onofrio et al. 1975; Iaccarino 1985). 

The first brief influx of this taxon is dated to 6.42 Ma B.P., 

Turborotalita multiloba

A 

5 

4 

Middle Messinian 


NW SE 

6 Late Messinian 

3 

2 

1 

N 

Inner 

Cilicia Basin 

Gökkuyu Mb. evaporites 

Latest Tortonian–early Messinian 

Early Messinian erosional unconformity (forced regression) 

Gravelly deltaic member of Handere F. 

(cross-cut by incised fluvial valley) 

Handere Fm. Misis High rises 

as a pop-up ridge 

Tortonian 

Late Burdigalian–Serravalian 

Karaisalı Fm. 

2.0 

0.5 

1.0 

2.5 

Late Aquitanian–early Burdigalian 

1.5 

3.0 

3.5 

4.0 

Adana 

Basin 

Cross-section line 

Cingöz Fm. - Güvenç Fm. 

Kaplankaya Fm. 

Base-Miocene erosional unconformity Gildirli Fm. 

0 km 50 

Present-day erosional 

margin of Miocene basin 

(outcrop limit) 

37 00'N 

Modern shoreline 

İskenderun 

Basin 

35 00'E 36 00'E 

B 

Late Messinian erosional unconformity (forced regression) 

Early Tortonian erosional unconformity 

Tırtar Fm. Handere Fm. 

Kuzgun Fm. 

Proto-Misis High 

Figure 10. (a) The relationship of Adana Basin to adjacent structural units, with a seismic map of the base-Miocene unconformity in the basin (depths below sea level 

in seconds of 2-way travel time); modified from Burton-Ferguson et al. (2005, Fig. 8). Note the trough-shaped, SW-plunging basin-floor palaeotopography and the high 

palaeotopographic relief along the basin’s eastern margin (Misis High). The NW–SE cross-section line pertains to the cartoon below. (b) Schematic cartoon showing the 

interpreted Miocene tectono-stratigraphic development in the Adana Basin (based on key features revealed by the seismic sections in Burton-Ferguson et al. 2005; not 

to scale). The successive stages are: 1– The Mesozoic bedrock of the Tauride foreland, affected by thrusting and elevated in the Eocene, is gradually denudated by erosion 

and eventually covered with alluvial deposits in the late Aquitanian to early Burdigalian; 2– The area becomes a foreland shelf zone as a result of mid-Burdigalian marine 

transgression, which drowns the basin and is followed by a late Burdigalian to Serravalian normal regression; 3– The basin is emerged by early Tortonian eustatic sea-level fall, 

and the resulting erosional unconformity is covered with alluvial deposits and then drowned by marine transgression; 4– Orogenic thrusting in the latest Tortonian to early 

Messinian converts the basin into a piggyback feature, with the thrust-induced uplift causing a forced regression and progradation of Gilbert-type deltas incised by fluvial 

valleys; 5– The foreland subsides under the increased load of the orogen thrust-sheets, which invites a marine transgression that brings in hypersaline water; and 6– The basin 

is emerged and subject to erosion due to the late Messinian evaporative drawdown of the Mediterranean Sea. Post-Miocene extensional deformation not considered. 

37 30'N 

36 30'N 

15

within subzone MMi13b, before its common occurrence 

in subzone MMi13c (Sierro et al. 2001; Drinia et al. 2004; 

Lourens et al. 2004; Kouwenhoven et al. 2006; Manzi et 

al. 2007). The higher samples show no occurrence of 

Turborotalita multiloba (Romeo) and the stratigraphic 

level of sample 20 is thus tentatively assigned to subzone 

MMi13b (Figure 9). 

Samples 21 and 22 (Figure 6) bear no planktonic 

foraminifera and contain miliolids, gastropods, ostracods, 

and echinid spines. Sample 24 from the top of the section 

is rich in planktonic foraminifera (Globorotalia suterae 

Catalano & Sprovieri, Globoturborotalita apertura 

(Cushman), G. nepenthes (Todd), G. decoraperta 

(Takayanagi & Saito), Neogloboquadrina acostaensis 

(Blow), N. humerosa (Takayanagi & Saito), Globigerina 

bulloides d’Orbigny, Orbulina universa d’Orbigny, 

16 


Figure 11. Schematic interpretation of the latest Miocene relative sea-level changes in the Adana Basin (cf. Figure 2); the 

suggested time intervals are based on biostratigraphic dating. (a) The late Tortonian tectonic thrusting lifts up the basin and 

forces a gradual regression with the deposition of downstepping Gilbert-type deltas in FRST. (b) The postthrusting foreland 

subsidence due to crustal load results in shallow reflooding of the basin by hypersaline water, with evaporitic TST and HST. (c) 

The Mediterranean early evaporative drawdown causes the second eustatically forced and nondepositional regression (FRST) 

in the basin. 

O. suturalis Brönnimann, O. bilobata d’Orbigny, and 

Globigerinella siphonifera (d’Orbigny); Figure 8) and also 

contains benthic foraminifers (Ammonia sp., Bulimina 

echinata d’Orbigny, and some miliolids and nodosarids), 

gastropods, bivalves, and echinid spines. However, it 

lacks Turborotalita multiloba (Romeo), Globigerinoides 

bulloideus Crescenti, and Globoturborotalita woodi 

(Jenkins), which allows this stratigraphic level to be also 

assigned to subzone MMi13b (Figure 9). 

8. Discussion 

8.1. Interpretation of Miocene events in the Adana Basin 

In this section, the chronology and regional causes of 

the late Miocene relative sea-level changes in the Adana 

Basin are discussed with reference to the basin’s tectonic 

development interpreted from multichannel seismic

A 

B 

1 m 

sections by Burton-Ferguson et al. (2005). However, the 

chronostratigraphy of tectonic events suggested by these 

latter authors needs to be rectified in the light of outcropderived 

biostratigraphic data. The main corrections are as 

follows: 

• The Güvenç Formation was considered by Burton- 

Ferguson et al. (2005, Fig. 3) to be of a Langhian to end- 

Tortonian age, whereas in reality it is no younger than 

Serravalian, truncated by the base-Tortonian erosional 

unconformity (Figure 2). The time-span of the Güvenç 

Formation is the same as that of the Cingöz Formation, 

which is its nearshore time-equivalent. 

• The Kuzgun Formation was considered by Burton- 

Ferguson et al. (2005, Fig. 3) to be Tortonian to earliest 

Messinian in age, whereas in reality this terrestrial 

formation is of an earliest Tortonian age, covered 

transgressively by the marine Handere Formation whose 

basal part bears planktonic forams of zone MMi10 and 

nannofossils of the uppermost zone MNN7c (Figure 2). 

• The Handere Formation was considered by Burton- 

Ferguson et al. (2005, Fig. 3) to be of a Pliocene age and 

overlying Messinian evaporites, whereas its actual age is 

Tortonian to Messinian and the Messinian evaporites 

occur at its top (Figure 2). 


C 

1 m 1 m 

Figure 12. (a) Late Miocene shallow-marine deposits folded by compressional tectonic deformation at the inner 

margin of the Adana Basin; picture from locality 4 in Figure 1b. (b, c) The exhumed planes of extensional normal 

faults in Miocene limestones at the inner margin of the Mut Basin; picture from locality 5 and in Figure 1b. 

SE 

Accordingly, the timing of the Miocene events in the 

basin in our interpretation, summarised in Figure 10, 

differs slightly from the chronology suggested by Burton- 

Ferguson et al. (2005). 

The seismic data indicate that the Miocene basin was 

initially a trough plunging towards the SW and passing 

into the Cilicia Basin (Figure 10a). The incipient Misis 

Structural High was probably a low-relief horst (Figure 10b; 

see Aksu et al. 2005, Fig. 8), perhaps shallowly submerged. 

Progressive unconformities on its flanks (Burton-Ferguson 

et al. 2005, Fig. 9) indicate that it became gradually elevated 

by differential subsidence during the Langhian–Tortonian 

and eventually popped up as a northward extension of 

the Kyrenia fold-and-thrust ridge in the latest Tortonian 

to early Messinian time (see Aksu et al. 2005, Fig. 12), 

when the Miocene succession also evidently underwent 

thrusting (Figure 10b; see Burton-Ferguson et al. 2005, 

Figs. 14 & 15). This event is thought to have marked the 

tectonic conversion of the Miocene Adana foredeep into a 

thrust-soled piggyback basin (Figure 10b). 

The biostratigraphic data indicate that the forced 

regression which formed the Gilbert-type deltas and 

caused fluvial valley incision (Figure 11a) occurred in 

the latest Tortonian to early Messinian, ~7.8 to 6.4 Ma 

17

B.P. (Figure 9), predating the Messinian early evaporative 

drawdown of the Mediterranean Sea (dated to 5.96 ± 

0.02 by Krijgsman et al. 1999). The cause of this stepwise 

and depositional forced regression is thought to have 

been the late Tortonian tectonic conversion of the Adana 

foredeep shelf into a thrust wedge-top (piggyback) basin, 

as is also evidenced by the compressional deformation 

of late Miocene basin-margin deposits (Figure 12a). 

The tectonic thrusting would likely cause a stepwise 

basin-floor uplift and relative sea-level fall (Figure 7b). 

The biostratigraphically constrained time frame for 

18 

A 

15 

10 

5 

thickness (metres) 

0 

delta topset 

Delta foreset 

B 

30 

25 

20 

mudstone samples for 

planktonic foraminifera 

Tırtar Fm. 

clay sand gravel 

silt 


delta foreset 

Delta topset 

coral and Miocene 

limestone blocks 

large coral 

fragments 

upslope-dipping 

backset deposits with 

large coral fragments 

Delta foreset 

bioturbated sandstones 

and oyster shells 

foreset direction 

C 

coral and reefal 

limestone fragments 

D 

backset 

coral 

fragments 

Figure 13. (a) Longitudinal outcrop section of the latest Tortonian Gilbert-type delta in the Mut Basin. (b) Vertical sedimentological 

log of the delta deposits, showing a coarsening-upwards coarse-grained succession overlying directly the Tırtar limestone. (c) Closeup 

view of the delta topset conglomerates, rich in pebble- to cobble-sized fragments of Miocene reefal limestones and corals. (d) A 

backset of upslope-dipping cross-strata within the delta conglomeratic foreset. Pictures and log from locality 7 in Figure 1b. 

the regression thus gives a more accurate timing of this 

tectonic event in the Adana Basin (cf. Burton-Ferguson et 

al. 2005). 

The formation of a piggyback basin would signify 

the climax of the structural contraction and thickening 

of the orogen, which would normally be followed by a 

regulating flexural subsidence of the foreland under the 

increased crustal load (DeCelles & Giles 1996). Such an 

episode of postthrusting isostatic subsidence is thought 

to have terminated the basinward advance of the deltas 

(LST) and caused their shallow marine drowning (TST)

in an estimated period between ~6.4 and ~6.0 Ma B.P. 

(Figure 11b). The marine highstand sedimentation (HST) 

was evaporitic, which indicates flooding by hypersaline 

water and thus suggests that the hypersalinity in the 

eastern Mediterranean Sea was reached at least ~6.5 

Ma B.P. The deposition of gypsiferous HST was terminated 

by the regional onset of the early evaporative drawdown in 

the Mediterranean Sea ~6 Ma B.P. (Krijgsman et al. 1999), 

which caused the second, erosional forced Messinian 

regression in the Adana Basin (Figure 11c). 

The early evaporative drawdown in the Mediterranean 

Sea might not exceed 200 m (Dronkert 1985; Krijgsman 

et al. 1999), but it was more than sufficient to emerge the 

shallow-marine peripheral basin. During the ensuing 

desiccation of the Mediterranean Sea, the postorogenic 

regional isostatic uplift (Jaffey & Robertson 2005; 

Cosentino et al. 2012) had apparently elevated the Adana 

Basin sufficiently high to prevent its reflooding by the 

Zanclean regional transgression dated to 5.3 Ma B.P. 

The basin thus remained terrestrial and accumulated 

a succession of Pliocene–Quaternary fluvial terraces 

(Figure 2) in response to the continuing uplift of the 

Taurides combined with concurrent eustatic sea-level 

changes (Haq et al. 1988). 

8.2. Comparison with the Mut Basin 

The end-Serravalian fall and rise of relative sea level 

recorded in the Adana Basin (Figure 2) are also well 

recognisable in the other Mediterranean peripheral basins 

of southern Turkey (Kelling et al. 2005), such as the 

adjacent Mut-Ermenek Basin (Atabey et al. 2000; Ilgar & 

Nemec 2005) and the Antalya Basin farther to the west 

(Figure 1a; Flecker et al. 1995; Karabıyıkoğlu et al. 2000; 

Deynoux et al. 2005; Monod et al. 2006; Çiner et al. 2008). 

An array of fan deltas prograded and became drowned 

in the Antalya Basin at that time (Larsen 2003), whereas 

an incised fluvial valley filled with a Gilbert-type delta, 

stratigraphically equivalent to the Kuzgun Formation in 

the Adana Basin (Figure 2), separates the amalgamated 

Burdigalian–Serravalian and Tortonian reefal platforms in 

the Ermenek Basin (Ilgar et al. unpublished data). The great 

amplitude of this early Tortonian eustatic cycle (see Haq et 

al. 1988) rendered it widely recognisable and correlative. 

However, the subsequent relative sea-level changes that 

occurred in these peripheral basins in Tortonian over a 

period of nearly 4 Ma are by no means correlative. 

In contrast to the Adana Basin, the other peripheral 

basins formed as postorogenic intramontane collapse 

depressions and their individual sedimentation history 

had recorded the interplay between lower-amplitude 

eustatic cycles and the gradual isostatic uplift of the 

orogen, local extensional tectonics and variable sediment 

supply. The gravelly deltas that formed in these basins in 

the Tortonian are mainly noncorrelative local sequences 


and parasequences. As an instructive example, we discuss 

here the development of a late Tortonian Gilbert-type 

delta in the nearby Mut Basin (Figure 13a; see locality in 

Figure 1b). 

The late Tortonian deltaic deposits in the Mut 

Basin belong to the shallow-marine Ballı Formation 

(stratigraphic equivalent of the Handere Formation in the 

Adana Basin, Figure 2) and overlie sharply an erosional 

ramp of the basin-margin reefal limestones of the Tırtar 

Formation. The Gilbert-type delta prograded from the 

basin’s northern margin towards the SSW and formed a 

coarsening-upwards clastic succession up to 30 m thick 

(Figure 13b). The delta’s fluvial topset has an erosional 

base and consists of coarse conglomerates and subordinate 

sandstones, rich in pebble- to cobble-sized coral fragments 

and Miocene limestone debris (Figure 13c). The gravel is 

mainly angular to subangular and moderately sorted, with 

a clast-supported texture and a matrix of poorly sorted 

coarse sand and granules. Fining-upwards bedsets of planar 

parallel-stratified gravel, 40–90 cm thick, are thought to 

be multistorey palaeochannels of braided streams filled 

mainly by the deposition of longitudinal bars (Nemec 

& Postma 1993; Miall 1996). Nonstratified gravel beds 

relatively richer in sand matrix, with or without normal 

grading, are probably deposits of hyperconcentrated flows 

and debris flows generated by the sediment-sweeping 

action of stream floods (Wasson 1977, 1979; Nemec & 

Muszyński 1982; Ridgway & DeCelles 1993). 

The erosional, oblique contact between the topset 

and foreset deposits (Figures 13a and 13b) indicates a 

nonrising trajectory of the delta shoreline (Breda et al. 

2007), which may suggest a normal or a forced regression. 

Foreset beds are mainly tabular, 10–45 cm thick, and 

inclined at up to 25°, composed of fine- to very coarsegrained 

sandstones rich in granules and scattered pebbles 

of up to 4 cm in size. Scattered oyster shells and isolated 

burrows are also common. The beds have erosional bases 

and show plane-parallel stratification, with or without 

normal grading. Their tractional deposition is attributed to 

fully turbulent hyperpycnal flows, or river-generated lowdensity 

turbidity currents (sensu Lowe 1982). Among the 

foreset beds are backsets of upslope-dipping cross-strata 

(Figures 13b and 13d), which occur as the infill of deltaslope 

chutes, up to 250 cm deep. The chute-fill deposits are 

rich in large coral fragments and reefal limestone debris, 

up to 100 cm in size, derived from the basin-margin 

Miocene carbonate platform. Delta bottomset consists of 

thin siltstone and fine-grained sandstone beds intercalated 

with mudstones (Figure 13b). 

The delta, thus built out directly on a wave-cut 

carbonate ramp, shows an erosional topset/foreset contact 

and contains debris derived from denudation of the basinmargin 

reefal limestones. Furthermore, the delta in its distal 

19

part is covered with thinly bedded reefal limestones, which 

indicates a shallow drowning, cessation of fluvial supply, 

and thus a considerable landward shift of the shoreline. 

There are no younger marine deposits in the basin, which 

implies its subsequent emergence. Taken together, the 

evidence indicates a forced regression followed by limited 

marine reflooding, and the delta’s association with the 

upper part of the Tırtar Formation might thus render it 

correlative with the deltas in the Adana Basin. In reality, 

the delta in the Mut Basin is significantly older than the 

deltas in the Adana Basin, as evidenced by biostratigraphic 

data. 

The age of the deltaic deposits in the Mut Basin has 

been determined on the basis of planktonic foraminifera 

in 10 mudstone samples from a vertical outcrop section 

of the delta bottomset deposits. Planktonic foraminifera 

are abundant and well preserved, but show a relatively low 

to moderate diversity and are accompanied by benthic 

foraminifers (representatives of Ammonia, Elphidium, 

Bulimina, Bolivina, Uvigerina, Nonion, miliolids, and 

nodosarids), gastropods, bivalves, ostracods, and 

echinid spines. The dominant planktonic species are 

Globigerina and Globigerinoides, accompanied by 

abundant globoturborotalids, neogloboquadrinids, and 

orbulinids. All samples consistently contain Globigerina 

bulloides d’Orbigny, Globigerinoides trilobus (Reuss), G. 

quadrilobatus (d’Orbigny), G. obliquus Bolli, G. bulloideus 

Crescenti, Globigerinella obesa (Bolli), Neogloboquadrina 

acostaensis (Blow), N. continuosa (Blow), and 

Globoturborotalita woodi (Jenkins). Species that are less 

consistently present or rare include Globoturborotalita 

decoraperta (Takayanagi & Saito), Tenuitellinata 

angustiumbilicata (Bolli), Catapsydrax parvulus Bolli, 

Loeblich & Tappan, Globorotalia scitula (Brady), 

Globigerinita uvula (Ehrenberg), Globigerinita incrusta 

Akers, G. glutinata (Egger), and Globigerinoides seigliei 

Bermudez & Bolli. The majority are long-range species, 

but the cooccurrence of Neogloboquadrina continuosa 

(Blow), N. acostaensis (Blow), Globigerinoides bollii Blow, 

and G. seigliei Bermudez & Bolli indicates a Tortonian age 

of the deposits. The cooccurrence of these last 3 species 

suggests an age probably not older than biozone MMi11 

(Figure 9; D’Onofrio et al. 1975; Iaccarino 1985; Iaccarino 

et al. 2007). The age can be constrained further by the 

presence of Neogloboquadrina continuosa Blow, whose last 

stratigraphic occurrence in the Mediterranean is in the 

subzone Globigerinoides obliquus extremus/G. bulloideus 

(Iaccarino 1985) or subzone MMi12a of Iaccarino et 

al. (2007). The deltaic bottomset deposits, coeval with 

the delta built-out, can thus be assigned to the biozone 

interval MMi11–MMi12a (Figure 9; D’Onofrio et al. 

1975; Iaccarino 1985; Iaccarino et al. 2007). The delta 

thus appears to have prograded in an approximate time of 

20 


8.5–7.8 Ma B.P., which corresponds to the regressive part 

(8.3–7.8 Ma B.P.) of the third-order eustatic sea-level cycle 

TB3.2 of Haq et al. (1988). 

In summary, the isolated late Tortonian delta in the Mut 

Basin, although seemingly correlative with similar deltas in 

the Adana Basin, is significantly older. Its development and 

drowning were apparently related to a third-order eustatic 

sea-level cycle, rather than local tectonics. There was no 

compressional Neogene deformation in the Mut Basin, 

whose margin instead shows normal faults indicating 

postorogenic tectonic extension (Figures 12b and 12c). 

The amplitude of eustatic cycle TB3.2 did not exceed 30 m 

(Haq et al. 1988), but the sea-level change was recognised 

in several other peri-Mediterranean basins (e.g., Rouchy 

& Saint Martin 1992; Larsen 2003; Roveri & Manzi 2006). 

However, this eustatic cycle was also modulated by 

local tectonics and became recorded differently in different 

basins. The marine reflooding of the Mut Basin was 

probably outpaced by the regional isostatic uplift of the 

Central Taurides, whereby the basin became emerged well 

before the hypersalinity state and evaporative drawdown 

of the Mediterranean Sea. In the Adana Basin, this 

eustatic cycle went unrecognisable (Figure 2), probably 

because the low-amplitude eustatic sea-level fall was fully 

compensated by the postthrusting isostatic subsidence of 

the foreland basin. In the Antalya Basin, in contrast, the 

Tortonian succession as a whole is regressive (Flecker et 

al. 1995; Deynoux et al. 2005; Çiner et al. 2008), yet shows 

no less than 4 cycles of a marine drowning and renewed 

progradation of the basin-margin fan deltas (Larsen 

2003). Their exact timing is uncertain and none of these 

cyclothems is recognisably forced-regressive, but at least 

some of them may be third-order eustatic sequences of 

type 2. 

These regional comparisons illustrate the difficulty 

with interbasinal stratigraphic correlations and serve as a 

warning against a superficial linking of seemingly similar 

events in the late Miocene peri-Mediterranean basins (see 

also discussions by Clauzon et al. 1996; Soria et al. 2003; 

Roveri & Manzi 2006). Reliable stratigraphic correlations 

require biostratigraphic constraints, a good understanding 

of the basin’s bathymetric conditions and sedimentary 

systems, and a careful account of the basin’s tectonic 

development. False hypothetical correlations of regional 

events lead to mistaken regional interpretations. 

9. Conclusions 

This combined sedimentological, sequence-stratigraphic, 

and biostratigraphic study of the late Miocene deposits 

in the Adana Basin has shown that the shallow-marine 

Handere Formation of Tortonian–Messinian age, 

previously interpreted as a simple regressive succession, 

bears a high-resolution record of several major relative

sea-level changes in the basin. The base of this formation 

recorded a forced regression corresponding to the end- 

Serravalian (Tor-1) eustatic fall in sea level. The lower to 

middle part of the formation is transgressive, culminating 

in offshore mudstones. The upper part is regressive and 

its 3 isolated conglomeratic members represent sharpbased 

Gilbert-type deltas with axially incised fluvial 

valleys that recorded a depositional forced regression, 

biostratigraphically dated to ~7.8 to 6.4 Ma B.P. on the basis 

of planktonic foraminifera in delta bottomset deposits. 

This latest Tortonian–early Messinian regression is 

attributed to the tectonic conversion of the Adana foredeep 

shelf into a piggyback basin, as is independently indicated 

by seismic profiles and compressional basin-margin 

deformation. The marine reflooding of the basin, estimated 

at ~6.4 to 6 Ma B.P., is ascribed to a postthrusting flexural 

subsidence of the foreland under increased crustal load. 

The transgression brought in evaporitic sedimentation, 

which suggests invasion of hypersaline Mediterranean 

water. The Mediterranean Sea in its eastern part would thus 

appear to have reached hypersalinity around 6.5 Ma B.P. 

The basin was subsequently emerged and its 

gypsiferous deposits were extensively eroded due to 

another Messinian forced regression, which is attributed 

to the early evaporative drawdown of the Mediterranean 

Sea, regionally dated to ~6 Ma B.P. The Mediterranean 

Sea then desiccated. The postorogenic regional isostatic 

uplift of the Taurides had meanwhile elevated the Adana 

Basin sufficiently to prevent its reflooding by the Zanclean 

regional transgression at 5.3 Ma B.P. The basin remained 

terrestrial in the post-Miocene time, accumulating a 

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