Oligocene Miocene basin evolution in SE Anatolia, Turkey: …forth/publications/Huesing_2009a.pdf · 2009-01-09 · Oligocene Miocene basin evolution in SE Anatolia, Turkey: constraints

Oligocene–Miocene basin evolution in SE Anatolia,

Turkey: constraints on the closure of the eastern Tethys gateway

SILJA K. HUSING1*, WILLEM-JAN ZACHARIASSE2, DOUWE J. J. VAN HINSBERGEN1,

WOUT KRIJGSMAN1, MURAT INCEOZ3, MATHIAS HARZHAUSER4,

OLEG MANDIC4 & ANDREAS KROH4

1Paleomagnetic Laboratory “Fort Hoofddijk”, Department of Earth Sciences, Utrecht

University, Budapestlaan 4, 3584 CD Utrecht, The Netherlands2Stratigraphy and Paleontology Group, Department of Earth Sciences, Utrecht University,

The Netherlands3Department of Geology, Fırat University, Elazıg, Turkey

4Natural History Museum Vienna, Austria

*Corresponding author (e-mail: [email protected])

Abstract: The Oligocene–Miocene was a time characterized by major climate changes as well aschanging plate configurations. The Middle Miocene Climate Transition (17 to 11 Ma) may evenhave been triggered by a plate tectonic event: the closure of the eastern Tethys gateway, themarine connection between the Mediterranean and Indian Ocean. To address this idea, we focuson the evolution of Oligocene and Miocene foreland basins in the southernmost part of Turkey,the most likely candidates to have formed this gateway. In addition, we take the geodynamic evol-ution of the Arabian–Eurasian collision into account.

The Mus and Elazıg basins, located to the north of the Bitlis–Zagros suture zone, were mostlikely connected during the Oligocene. The deepening of both basins is biostratigraphicallydated by us to occur during the Rupelian (Early Oligocene). Deep marine conditions (between350 and 750 m) prevailed until the Chattian (Late Oligocene), when the basins shoaled rapidlyto subtidal/intertidal environment in tropical to subtropical conditions, as indicated by the macro-fossil assemblages. We conclude that the emergence of this basin during the Chattian severelyrestricted the marine connection between an eastern (Indian Ocean) and western (Mediterranean)marine domain. If a connection persisted it was likely located south of the Bitlis–Zagros suturezone. The Kahramanmaras basin, located on the northern Arabian promontory south of theBitlis–Zagros suture zone, was a foreland basin during the Middle and Late Miocene, possiblylinked to the Hatay basin to the west and the Lice basin to the east. Our data indicates that this fore-land basin experienced shallow marine conditions during the Langhian, followed by a rapid dee-pening during Langhian/Serravallian and prevailing deep marine conditions (between 350 and750 m) until the early Tortonian. We have dated the youngest sediments underneath a subduc-tion-related thrust at c. 11 Ma and suggest that this corresponds to the end of underthrusting inthe Kahramanmaras region, i.e. the end of subduction of Arabia. This age coincides in time withthe onset of eastern Anatolian volcanism, uplift of the East Anatolian Accretionary Complex,and the onset of the North and East Anatolian Fault Zones accommodating westward escape tec-tonics of Anatolia. After c. 11 Ma, the foreland basin south of the Bitlis formed not (or no longer) adeep marine connection along the northern margin of Arabia between the Mediterranean Sea andthe Indian Ocean. We finally conclude that a causal link between gateway closure and globalclimate change to a cooler mode, recorded in the Mi3b event (d18O increase) dated at 13.82 Ma,cannot be supported.

Tectonic closure and opening of marine gateways issuggested to have led to substantial reorganizationof surface and deep ocean water currents and mayhave caused important changes in global climate.The closure of the Panama Isthmus between 3.0and 2.5 Ma has influenced the Gulf Stream, trigger-ing major Northern Hemisphere glaciations (Bartoli

et al. 2005; Schneider & Schmittner 2006). Theopening of the Drake Passage allowed the start ofthe Antarctic Circumpolar Current which mighthave initiated the abrupt climate cooling aroundthe Eocene/Oligocene boundary and the extensivegrowth of Antarctic ice sheets (Livermore et al.2005). The restriction of water exchange across

From: VAN HINSBERGEN, D. J. J., EDWARDS, M. A. & GOVERS, R. (eds) Collision and Collapse at theAfrica–Arabia–Eurasia Subduction Zone. The Geological Society, London, Special Publications, 311, 107–132.DOI: 10.1144/SP311.4 0305-8719/09/$15.00 # The Geological Society of London 2009.

12345678910111213141516171819202122232425262728293031323334353637383940414243444546474849505152535455565758

the former straits between Spain and Moroccoresulted in the desiccation of the MediterraneanSea during its Messinian Salinity Crisis (Hsu et al.1973). Likewise, the disconnection of the IndianOcean and the Atlantic/Mediterranean watermasses has been suggested to have caused a majormiddle Miocene climate change, widely recognizedin both the marine (Woodruff & Savin 1989; Flower& Kennett 1994; Zachos et al. 2001; Bicchi et al.2003) and the terrestrial record (Krijgsman et al.1994). It is this disconnection that forms the scopeof this paper.

The middle Miocene is a period characterized bymajor environmental changes during which theEarth’s climate gradually progressed into a coldermode (Zachos et al. 2001). The Miocene ClimateOptimum between 17 to 15 Ma was followed byan interval of global climate variability between15 and 14 Ma, marked by atmospheric and oceaniccooling, East Antarctic Ice Sheet growth, andcarbon cycle variability (Woodruff & Savin 1989;Flower & Kennett 1994; Zachos et al. 2001).Seven major d18O shifts, Mi1 to Mi7, to higher(¼ colder) values documented in marine recordsof the Atlantic reflect brief periods of increasedglaciations (Miller et al. 1991; Wright et al. 1992;Miller et al. 2005). The Mi3a, Mi3b and Mi4events between about 14.5 and 12.5 Ma representthe middle Miocene d18O increase, leading theglobal climate into a colder mode at the same timeas the onset of the Antarctic glaciations (van derZwaan & Gudjonsson 1986; Abels et al. 2005;Miller et al. 2005).

A direct relationship between the MiddleMiocene Climate change, whether recorded inoxygen or carbon isotopes, marine or terrestrialfauna, and the closure of the eastern Tethysgateway has so far never been proven, althoughmany studies suggest a causal link between thetwo events (e.g. Woodruff & Savin 1989; Rogl1999; Flower & Kennett 1993). Part of theproblem is that the sediments that were depositedin the eastern Tethys gateway have on manyoccasions not been recognized or properly dated.In addition, the chronological sequence of tectonicprocesses involved in the convergence of theEurasia and African–Arabian plates is complexand actively debated (see Garfunkel 1998, 2004;Golonka 2004). To assess the timing of gatewayclosure along the northern Arabian promontory,the major geodynamic processes of the Arabia–Eurasia collision and their tectonic responses haveto be taken into account. According to reconstruc-tions of Jolivet & Faccenna (2000) and Bellahsenet al. (2003), Arabia collided first in the easternAnatolian/western Iranian region around 30 Maago. Consequently, it gradually rotated counter-clockwise leading to diachronous collision eastwardfrom Southeastern Anatolia towards the Persian

Gulf (Hessami et al. 2001). Therefore, we decidedto study the southernmost flysch deposits ineastern Anatolia (Fig. 1), these being the mostlikely candidates to represent the youngest sedi-ments deposited just prior to the disconnection ofthe Indian–Arabian gateway.

Geodynamic and geological context

The continental collision of the African–Arabianplate with the Eurasian plate resulted in a tectoniccollage in eastern Anatolia that is generally subdi-vided into: (1) the eastern Rhodope–Pontide Arcin the north; (2) the East Anatolian AccretionaryComplex consisting of an ophiolitic melange over-lain by Paleocene to upper Oligocene sediments;and (3) the Bitlis–Poturge Massif tectonically over-lying the northern part of the Arabian margin(Fig. 1) (Sengor & Yılmaz 1981; Yılmaz 1993;Tuysuz & Erler 1995; Robertson 2000; Sengoret al. 2003; Agard et al. 2005). As north–southshortening continued between the converging Eura-sian and Arabian plate, the relatively soft and irre-sistant East Anatolian Accretionary Complex tookup most of the initial post-collisional convergentstrain by shortening and thickening (Yılmaz et al.1998). Around 13–11 Ma, eastern Anatolia under-went rapid uplift and was confronted with onset ofwidespread volcanism (Dewey et al. 1986; Pearceet al. 1990; Keskin 2003; Sengor et al. 2003),which has been associated with detachment ofnorthward dipping subducted lithosphere (Keskin2003; Faccenna et al. 2006; Hafkenscheid et al.2006). From this moment onward, the ongoingnorthward motion of Arabia (still continuing today)(McClusky et al. 2000; Reilinger et al. 2006;Allmendinger et al. 2007), and the retreat of theHellenic subduction zone to the west (Berckhemer1977; Le Pichon et al. 1982; Jolivet 2001) led towestward tectonic escape of Anatolia along theNorth and East Anatolian Faults (Dewey & Sengor1979; Sengor et al. 1985).

The present-day plate boundary of the Africanand Eurasian plates is determined by the Bitlis–Zagros suture zone (Robertson 2000 and referencestherein; Westaway 2003). On the Arabian plate, tothe south of the suture zone, Eocene and younger(volcano-) sediments are relatively flat lying. Northof the Bitlis–Poturge zone, Tertiary marine sedi-ments crop out rarely and the geology is dominatedby pre-Neogene basement rocks (metamorphicrocks) and Neogene volcanic rocks. The Bitlis–Poturge Massif itself is characterized by a stack ofnappes originated on the Eurasian side of theNeotethys (Robertson 2000; Robertson et al. 2004).

The Bitlis–Poturge Massif runs from southeast-ern Turkey to the eastern Mediterranean basin intothe Cyprus arc, where it meets the East Anatolian

S. K. HUSING ET AL.108

5960616263646566676869707172737475767778798081828384858687888990919293949596979899100101102103104105106107108109110111112113114115116

Fault (EAF). Here the structure becomes morecomplex with several sub-parallel southwestwardsrunning faults and thrusts. The East AnatolianFault is a 2–3 km wide, active left-lateral strike-slip fault extending from Antakya in the west toKarliova in the NE, where it meets the easterntermination of the North Anatolian Fault (NAF)(Figs 1 & 2; EAFZ and NAFZ). The NAF is a right-lateral strike-slip fault extending over a length ofabout 1300 km westward. The relative Africa–Arabia motion is taken up by strike-slip displace-ment along the Dead Sea Fault (Jolivet & Faccenna2000), while the Africa–Anatolia motion is takenup by subduction south of Cyprus. The overall con-vergence between Arabia and Anatolia is taken upalong the North and East Anatolian fault zones(NAFZ and EAFZ) (Fig. 2) (e.g. McClusky et al.2000; Sengor et al. 2005). There is general consen-sus that the NAFZ and EAFZ had the majority oftheir displacement in Plio-Pleistocene times(Barka 1992; Westaway 2003, 2004; Hubert-Ferrariet al. 2008) although incipient motion may havebeen as early as late Serravallian/early Tortonian(c. 12 to 11 Ma) (Dewey et al. 1986; Hubert-Ferrariet al. 2002; Bozkurt 2003; Sengor et al. 2005).Q1

The region that comprises the eastern Tethysgateway has thus been subjected to plate conver-gence and subduction. Sengor et al. (2003)

suggested that this subduction led to southwardmigrating accretion of nappes and overlying deep-marine foreland basin deposits, even though indi-vidual basins that may reflect such evolution havenot been identified in the geological record, whichis, at least in part, due to the young volcanicsequences covering a large part of eastern Turkey.If southward accretion of nappes indeed occurred,one should be able to identify southward youngingflysch deposits. A foredeep likely remains presentuntil continent–continent collision and subsequentslab break-off stalls convergence and the collisionzone is uplifted. Even though small marine basinsmay remain, the long distance between the PersianGulf and the Mediterranean Sea makes foredeepsthe most promising basins to have formed thegateway between these water masses. In the follow-ing paragraphs we will present and discuss the evol-ution of foredeep basins in SE Turkey in the light ofthe closure of the eastern Tethys gateway.

Basin evolution

The Arabian foreland is separated from the EastAnatolian Accretionary Complex (EAAC) by theBitlis–Poturge Massif (Fig. 1). The area of thismassif corresponds to the compression zone located

Fig. 1. Outline of tectonic map of the Middle East region, showing major structures such as the Bitlis–ZagrosSuture zone, the North and East Anatolian Fault Zones (NAFZ and EAFZ), and mountain ranges related to theconvergence of Africa–Arabia and Eurasia (drawn after Geological map of Turkey (Senel 2002)).

OLIGO–MIOCENE SE TURKISH TETHYS FORELAND 109

117118119120121122123124125126127128129130131132133134135136137138139140141142143144145146147148149150151152153154155156157158159160161162163164165166167168169170171172173174

Fig

.2.

Outl

ine

of

schem

atic

geo

logic

alm

apof

SE

Anat

oli

ain

South

east

ern

Turk

eyw

ith

maj

or

tect

onic

stru

cture

s.N

ote

the

thre

eboxes

indic

atin

gth

est

udie

dar

eas:

Mus

inth

eea

ster

nm

ost

par

tan

dE

lazı

g,

both

nort

hof

the

Bit

lis–

Zag

ros

Sutu

rezo

ne

and

Kah

ram

anm

aras

south

of

the

Bit

lis–

Zag

ros

Sutu

reZ

one

(dra

wn

afte

rG

eolo

gic

alm

apof

Turk

ey(S

enel

2002))

.


175176177178179180181182183184185186187188189190191192193194195196197198199200201202203204205206207208209210211212213214215216217218219220221222223224225226227228229230231232

between the two continental crusts, Eurasia andAfrican–Arabian. The massif was stacked to forma nappe complex during the closure of the Neo-Tethys by the middle Miocene (Dewey 1986 andreferences herein).

We have studied the southernmost flysch depos-its in the eastern Anatolian orogenic system. Theseare found in the Mus and Elazıg basins, both north ofthe Bitlis–Poturge Massif, and the Kahramanmarasbasin located south of the Bitlis–Poturge Massif andnear the triple junction of the Arabian, Eurasian andAnatolian plates (Fig. 2).

Geological setting of the Mus basin

The Mus basin is an elongated structure locatednorth of the Bitlis–Poturge Massif and east of theNorth and East Anatolian Fault (Figs 2 & 3).According to previous studies (Saroglu & Yılmaz1986; Sancay et al. 2006) the basin contains upperEocene to lower Miocene limestones, marls and tur-biditic sandstones with marine sedimentation con-tinuous from the Oligocene to Aquitanian. Thesedeposits overlay an upper Cretaceous ophioliticmelange. Saroglu & Yılmaz (1986) suggested thatlower Miocene limestones are widespread in thenorthern part of the Mus area, while middleMiocene strata were not found. These sequencesare unconformably covered by allegedly upper

Miocene and younger continental clastics and vol-canics (Saroglu & Yılmaz 1986; Sancay et al.2006). Detailed biostratigraphy was carried outmainly based on dinoflagellates and palynomorphsyielding a Rupelian (early Oligocene) to Aquitanian(early Miocene) age (Sancay et al. 2006). Theoccurrence of the benthic foraminiferal family ofMiogypsinidae was interpreted as possible indicatorfor a connection with the Indo-Pacific during theOligocene (Sancay et al. 2006).

We sampled two sections in the Mus basin(Fig. 3). The eastern transect comprises allegedlyEocene–Oligocene clastics in the northern part ofthe basin, and Oligocene flysch sediments followedby marine limestones which are covered by volca-nics. The second transect in the western part of thebasin covers the transition from marls to limestones,assuming it is equivalent to the uppermost part of theeastern succession. The entire succession gentlydips towards the NW.

The base of the eastern section (east transect inFig. 3) is determined by a thrust zone emplacingallegedly Eocene clastic sediments onto Pliocenedeposits (see geological map of Turkey, Senel2002) (Fig. 3). The first 20 m of the studiedsection is characterized by an alternation of con-glomerates, clays, sands, and silts (Fig. 4). A layerof limestone (1.5 m) with shell fragments and thepresence of large gastropods clearly indicate shallow

Fig. 3. Simplified tectonic and geological map of the Mus area including the trajectories of the two studied sections:an about 1.4 km long transect in the eastern part of the basin and additionally an about 500 m long transect in thewestern part of the basin equivalent to the uppermost part of the western transect. Refer to Legend for key to lithologyand/or age of outcrops (drawn after Geological map of Turkey (Senel 2002)).


233234235236237238239240241242243244245246247248249250251252253254255256257258259260261262263264265266267268269270271272273274275276277278279280281282283284285286287288289290

Fig. 4. Lithological column of the studied sections in the Mus basin with the biostratigraphic results. The age modelis based on planktonic foraminifer occurrences and the macrofossil assemblage in the uppermost 40 m (mainlylimestones and sands) of the stratigraphy. Planktonic foraminifer occurrences have been correlated to planktonicforaminifer zones, which, in turn, are tied to stages during the Oligocene leading to a correlation to the Geological TimeScale. See legend for key to lithologies, structures and fossils.


291292293294295296297298299300301302303304305306307308309310311312313314315316317318319320321322323324325326327328329330331332333334335336337338339340341342343344345346347348

marine conditions. This sequence is followed by athick (about 1.3 km) succession of alternating clayand sandstone. Occasionally conglomeratic layers,characterized by angular, unsorted material, occurin the succession. These layers have thicknesses ofup to 10 m and are interpreted as debris flows. Thesandstone layers show typical transport character-istics such as fining upwards, Bouma sequence,flute casts and fossil fragments indicating a turbidi-tic origin. These turbidites occur as massive sand-stone layers of thicknesses of up to 15 m or asseveral thinner (up to 50 cm) turbidite layers, prob-ably representing individual events. Only minorslumping, indicating an unstable submarine paleo-slope, and folding occur throughout the succession.The upper part of the section shows shoaling charac-terized by shallow marine limestone, containingechinoderms, bivalves and gastropods, followedby continental clastics.

The western transect (west transect in Fig. 3) isdominated by bluish clay with occasional red sedi-ments. This is followed by a thick, about 100 m,sequence of alternating softer bluish sands, brown-ish sands and indurate bluish sands, probably allof marine origin. These sediments are overlain bycoral limestones, which, in turn, are covered byvolcanic rocks, probably of Miocene age. This suc-cession also clearly indicates shoaling towardsthe top.

Biostratigraphic results of the Mus basin

For biostratigraphy, samples were collected at aboutevery 20 m from both the western and eastern trans-ect (Fig. 3). Not every sample proved to be useful forbiostratigraphy or paleobathymetry. The number offoraminifers is extremely variable and most likelyfluctuate in pace with changes in terrigenousclastic input. Preservation is generally poor withspecimens mostly recrystallized and frequently dis-torted. Samples from the upper 300 m of the westernsection are barren in planktonic foraminifers.

The low diversity in planktonic foraminiferalfauna in both sections is dominated by globoquadri-nids and catapsydracids with occasional occur-rences of Globigerina ciperoensis and Globigerinaangulisuturalis and clearly points to an Oligoceneage for the eastern and lower western section(Fig. 4) (Berger & Miller 1988; Spezzaferri &Premoli Silva 1991).

The basal part of the eastern section correlates toplanktonic foraminiferal biozone P19 of Berger &Miller (1988) on the basis of trace occurrences ofspecimens identical to Turborotalia ampliapertura.This biozone is Rupelian (early Oligocene) in age(Fig. 4) (Berggren et al. 1995). The lowermostoccurrence of Globigerina angulisuturalis isrecorded at 950 m (TR 221) in the eastern section

which together with the highest occurrence ofParagloborotalia opima opima (at 1045 m(TR222)) indicates that the middle part of theMus section correlates to planktonic foraminiferalbiozone P21 of Blow (1969) and Berger & Miller(1988) which is latest Rupelian to early Chattianin age (Berggren et al. 1995). The absence ofParagloborotalia opima opima from sample level1045 m (TR 222) upward in the eastern transectand the occurrence of typical Paragloborotaliapseudokugleri and even of forms transitionalbetween Paragloborotalia pseudokugleri andParagloborotalia kugleri at the top of the section(1360 m (TR 232)) indicate that the upper partextends upwards into the lower part of planktonicforaminiferal biozone P22 of Berger & Miller(1988) being Chattian in age (Berggren et al. 1995).This is confirmed by the presence of Paragloborota-lia siakensis and Globigerinoides primordius in theyoungest samples. Both species make their firstappearance in the lower part of biozone P22 togetherwith Paragloborotalia pseudokugleri (Berger &Miller 1988; Spezzaferri 1994).

In the western section, the co-occurrence ofGlobigerina angulisuturalis and Paragloborotaliaopima opima at 2 m and 195 m (TR 202 andTR 210) indicates that the lower 200 m correlatesto the interval between 900 and 1100 m in theeastern section. Both these intervals belong tobiozone P21. This interval is followed by sedimentsthat are barren in planktonic foraminifers but rela-tively rich in shallow water benthic foraminifers.

The macrofossil assemblage of the uppermost40 m in the eastern transect comprises bivalves, gas-tropods and echinoids. The assemblage is diminishedby complete aragonite leaching. Nevertheless, thefauna is age indicative and allows palaeoecologicalinterpretations. The mollusc fauna comprises typicalOligocene taxa such as the gastropod Ampullinopsiscrassatina (Lamarck 1804) Q2and the bivalves Amussio-pecten labadyei (d’Archiac & Q3Haime 1853) andRingicardium buekkianum Q4(Telegdi-Roth 1914).Some species such as Dilatilabrum sublatissimus(d’Orbigny 1852), Q5;Q6Strombus cf. praecedens Schaffer1912, Cordiopsis incrassatus Q7(Nyst 1836),Amussiopecten subpleuronectes (d’Orbigny 1852),and Hyotissa hyotis Q8(Linnæus 1758) appear duringthe Chattian and persist into the Miocene.

An important biostratigraphic feature is theco-occurrence of the pectinids Amussiopecten laba-dyei and A. subpleuronectes and the occurrence oftransitional morphs. This evolutionary phase isrecorded so far only from the upper Chattian(Mandic 2000). Especially in the Iranian QomBasin, this assemblage co-occurs with the larger for-aminifera Eulepidina dilatata. The last occurrenceof Amussiopecten labadyei precedes the first occur-rence of Miogypsinoides which roughly coincides


349350351352353354355356357358359360361362363364365366367368369370371372373374375376377378379380381382383384385386387388389390391392393394395396397398399400401402403404405406

with the base of the early Miocene. The entiremollusc assemblage is therefore pointing to a lateChattian age. This dating is supported by theechinoid fauna. Parascutella subrotundaeformis(Schauroth 1865),Q9 a sand dollar which occurs mostcommonly in Northern Italy, is restricted to theChattian and Aquitanian.

Comparable assemblages are described from theupper Chattian of the central Iranian Qom For-mation (Mandic 2000; Harzhauser 2004; Reuteret al. 2007) and along the entire northern coast ofthe Western Tethys (Harzhauser et al. 2002). Arelation to the Central Paratethys is indicated bythe occurrence of Ringicardium buekkianum,which is known from the Lower Egerian (UpperChattian) deposits of Hungary (Baldi 1973). Thefaunistic relations towards the east are low. OnlyDilatilabrum sublatissimus (d’Orbigny 1852)reaches to the Zagros Basin and the Arabian shelfduring the Aquitanian (Harzhauser et al. 2007).The echinoderm Clypeaster waageni (Duncan &Sladen 1883),Q10 in contrast, represents ties with theechinoid fauna of the Lower Indus Basin.

Numerical ages for the basin fill are provided bythree planktonic foraminiferal bioevents. However,equating highest and lowest occurrences (ho andlo) with the Last Occurrence (LO) and First Occur-rence (FO) of theses species should be accepted withreservation because the positions are poorly deli-neated due to large sampling distances in combi-nation with scarcity and poor preservation of theage diagnostic species.

The oldest bioevent in the Mus section is thelowest occurrence of Turborotalia ampliaperturasome 300 m above the base of the eastern section(TR 190). The LO of this species is calibrated at30.3 Ma (Berggren et al. 1995) providing aminimum age for the base of the Mus section. Theage for the top of the eastern section should beslightly younger than the age of 25.9 Ma for the FOof Paragloborotalia pseudokugleri (Berggren et al.1995) because of the presence of paragloborotalidsbeing transitional between Paragloborotalia pseu-dokugleri and Paragloborotalia kugleri. The ho ofParagloborotalia opima opima at 1045 m (TR 222)in the eastern section provides an extra age cali-bration point of 27.456 Ma being the calibrated agefor the LO of Paragloborotalia opima opima atODP Site 1218 (Wade et al. 2007). The dating ofthe top of the section is in accordance with the macro-fauna which strongly indicates a late Chattian age forthe upper 40 m the eastern transect.

No numerical ages are provided for the westernsection. However, based on the co-occurrence ofGlobigerina angulisuturalis and Paragloborotaliaopima opima in the lower 200 m, this interval corre-lates to the biozone P21. The upper 300 m lack anyage diagnostic planktonic foraminifer.

Palaeoenvironmental interpretations

for the Mus basin

Benthic foraminifers in the sections were further-more used to estimate the depositional depth. Thecommonly used method of calculating depth bydetermining the ratio between planktonic andbenthic foraminifers (van der Zwaan et al. 1990;van Hinsbergen et al. 2005b) is not reliable heredue to significant downslope transport (seen in pre-sence of notorious epifytes and shallow waterbenthic foraminifers such as Pararotalia andAmphistegina) and poor preservation. Instead, wefocus on the deepest water benthic foraminiferaldepth markers (for list see van Hinsbergen et al.2005b) and the macrofossils. In the easternsection, the depositional environment of the lower20 m is characterized by shallow marine conditions,indicated by shell fragments in the limestone.However, a rapid deepening trend occurs at about50 m indicated by the presence of benthic foramini-feral depth markers (typically Cibicides (pseudo)ungerianus, Gyroidina spp. Uvigerina spp. andoccasionally Oridorsalis spp.), and the absence ofmarkers for deeper water, which points at a deposi-tional depth range of 350 to 750 m (the upper limit isconstraint by the occurrence of Oridorsalis spp.after van Hinsbergen et al. 2005b). Towards thetop of the eastern section rapid shoaling is evidentfrom the presence of macrofossils. Both the mol-luscs and echinoderms of the uppermost 40 m indi-cate a shallow marine, tropical to subtropical,depositional environment with sand bottoms andalgal or sea grass patches. Giant conchs such asDilatilabrum sublatissimus (d’Orbigny 1852) arefound today in sea grass meadows and shelteredlagoons, where they live partly buried in the softsubstrate (Bandel & Wedler 1987). Similarly, theextant representatives of the oyster Hyotissa hyotisprefer shallow subtidal habitats where they areattached to rocks and corals (Slack-Smith 1998).Extant Echinolampas and Clypeaster, too, occurmost commonly on sandy sediments with sea grasspatches (Hendler et al. 1995).

In the western section a shoaling trend in theupper 250 m is observed by the relatively richoccurrence of shallow water benthic foraminifersand occasional red sediments. The differencesbetween west and east suggest that the westernpart of the Mus basin shoaled more rapidly orearlier during the Chattian than the eastern part.

Implications for the Mus basin

Based on the occurrence of turbidites, slumpingand minor folding, this about 1.5 km thick marinesuccession is interpreted as deposits of a deepmarine basin.


407408409410411412413414415416417418419420421422423424425426427428429430431432433434435436437438439440441442443444445446447448449450451452453454455456457458459460461462463464

Shallow marine conditions during the Rupelian(P19) were replaced by rapid deepening of thebasin during biozone P22, late Chattian. The endof the flysch deposition during the Chattian marksthe emergence of the basin which probablyremained shallow marine until the late Chattian.

Considering the biostratigraphic ages, a sedimen-tation rate between 15 and 27 cm/ka is calculated.The constant water depth of 350 to 750 m duringdeposition indicates approximately 2 km of subsi-dence throughout the Oligocene, followed by rapiduplift and exposure of the succession after the lateChattian. Our biostratigraphic dates based on plank-tonic foraminifers in the flysch deposits corroboratethe ages published previously based on dinoflagel-lates and palynomorphs (Sancay et al. 2006).

Geological setting of the Elazıg basin

The studied Gevla section is situated in the eastern-most part of the Elazıg basin, about 40 km NE ofElazıg (Fig. 2). The basin has been studied byseveral workers; however, the literature has beenpublished mostly in Turkish (see Aksoy et al.2005) and no detailed information is available forthe easternmost part of the basin. At present, thebasin fill is exposed in an NE–SW belt in theeastern Taurides of Anatolia. The generalized strati-graphy of the Tertiary sediments has been describedas: lower Paleocene continental deposits at the base,followed by upper Paleocene to lower Miocenemarine deposits and finally Pliocene to Quaternarycontinental deposits. The basement of the Elazıgbasin is formed by Permo-Triassic metamorphicrocks, namely Keban Metamorphics, which wereemplaced over upper Cretaceous magmatic rocksnorth of Elazıg (Perincek 1979; Perincek &Ozkaya 1981; Aktas & Robertson 1984; Bingol1984; Aksoy et al. 2005).

Detailed stratigraphic, sedimentological andtectonic characteristics of the Elazıg area havebeen discussed elsewhere (e.g. Perincek 1979;Perincek & Ozkaya 1981; Aktas & Robertson1984; Bingol 1984; Cronin et al. 2000a, b; Aksoyet al. 2005). From the late Paleocene, shallowmarine carbonates, deposited in an extensionalback-arc setting, were accumulated when the basinfurther subsided until Middle to Late Eocene(Aksoy et al. 2005). During Oligocene to earlyMiocene, after reaching its maximum extendduring the Middle to Late Eocene, deposition wasrestricted to the N–NW and became progressivelyshallower, indicated by Oligocene reefal limestonesuntil the final subaerial exposure at the end ofOligocene. Marine Miocene deposits are restrictedto small areas in the basin and more widespreadnorth of the basin. From Middle Miocene onwardsthe basin was affected by a strong, North–South,

compression. Later, Pliocene to Pleistocene alluvialfan, fluvial and lacustrine sediments were depositedcovering Early Miocene sediments (Cronin et al.2000a, b; Aksoy et al. 2005).

In this setting, we studied a section situated in theeasternmost part of the Elazıg basin. According tothe geological map of Turkey (Senel 2002), in thearea east of the town Basyurt (Fig. 5), Lower toMiddle Eocene continental clastics unconformablyoverly Mesozoic ophiolitic melange. These clasticsediments are, in turn, overlain by either Miocene–Pliocene clastic or volcanic rocks.

The basal part of the studied Gevla succession,about 15 km NE of Basyurt, starts with bluishmarine clay containing bivalves, followed by analternation of clay and sandstone (the sandstonesare up to 50 cm thick or about 5 m thick withcross bedding) (Fig. 6). A distinct layer with abun-dant bivalves and gastropods is located at about50 m stratigraphic position. Three distinct limestonelayers occur between about 100 m and 260 m strati-graphic level. The first one, at about 100 m, is anodular limestone with shell fragments, sponges(up to 30 cm) and corals, followed by two nummu-litic limestone horizons, at 244 m and 255 m. This isfollowed by about 400 m of blue clay grading intoa 600 m thick succession of alternating clay andsandstone, whereby the sand layers show typicaltransport characteristics, such as shell fragments,

Fig. 5. Simplified tectonic and geological map of theeasternmost part of the Elazıg basin. The trajectory of thestudied section, called Gevla, is about 15 km NE of thecity Basyurt and covers the interval between Eocenelimestones and Miocene volcanics to the north. For keyto the lithologies and/or ages refer to Legend (drawnafter Geological map of Turkey (Senel 2002)).


465466467468469470471472473474475476477478479480481482483484485486487488489490491492493494495496497498499500501502503504505506507508509510511512513514515516517518519520521522

displaced nummulites and gastropods (for instanceat 663 m and 1158 m), fining upward sequencesand cross bedding. These layers are interpreted asturbiditic in origin. This succession is followed byabout 300 m of blue clay, and the section endswith a 50 m thick limestone with bivalves (up to5 cm), and clayey intervals with well preservedechinoderms, sponges and corals. These limestones,in turn, are covered by Miocene volcanic rocks. Intotal, the section is about 1.6 km thick.

Slumping at several levels within the successionindicates an unstable submarine slope. Internalfolding is not observed within the succession. Theentire succession gently dips towards the NW.

Biostratigraphic results of the Elazıg basin

Hand samples were collected from about every 20 mthroughout the entire section, but not every samplecontained (diagnostic) planktonic and/or benthicforaminifers. The number of foraminifers is extre-mely variable and most likely fluctuates withchanges in terrigenous clastic input. Preservationis generally poor, with specimens mostly recrystal-lized and frequently distorted. The overall aspectsof the planktonic foraminiferal fauna in thissection is similar to that of the Mus section, whichmeans that the foraminiferal fauna is dominatedby globoquadrinids and catapsydracids withoccasional occurrences of Globigerina ciperoensisand Globigerina angulisuturalis pointing to anOligocene age for this section (Fig. 6) (Berger &Miller 1988; Spezzaferri & Premoli Silva 1991).The presence of Turborotalia ampliapertura up toand including level 287 m (TR 244) providesevidence that the lower part of the section correlateswith planktonic foraminiferal biozone P19 ofBerger & Miller (1988), which is late Rupelian,early Oligocene, in age (Berggren et al. 1995).The lowest occurrence of Globigerina angulisutur-alis is recorded at 477 m (TR 250) which in termsof the zonal scheme of Berger & Miller (1988)would mark the top of biozone P20 although itshould be noted that Globigerina angulisuturalis isneither frequent in this section nor does it displayvery prominent U-shaped sutures. Typical Paraglo-borotalia opima opima is present from level 317 m(TR 245) up to and including level 1445 m (TR 293)indicating that the larger part of the Gevla sectioncorrelates with planktonic foraminiferal biozone

Fig. 6. Lithological column of the studied Gevla sectionin the Elazıg basin with the biostratigraphic results.The age model is based on planktonic foraminifers andmacrofossil assemblage in the 50 m of stratigraphy.Planktonic foraminifer occurrences have been correlated

Fig. 6. (Continued) to planktonic foraminifer zones,which, in turn, are tied to stages during the Oligoceneresulting into a correlation to the Geological Time Scale.Refer to Legend in Figure 4 for key to lithologies,structures and fossils.


523524525526527528529530531532533534535536537538539540541542543544545546547548549550551552553554555556557558559560561562563564565566567568569570571572573574575576577578579580

P21 of Blow (1969) and Berggren et al. (1995)which in terms of chronostratigraphy is latestRupelian to early Chattian in age (Berggren et al.1995). The top of the section post-dates the highestoccurrence of Paragloborotalia opima opima, andcorrelates to the basal part of the late Chattian plank-tonic foraminiferal biozone P22 (Berger & Miller1988), which is evidenced by the joint presence ofParagloborotalia opima nana, Globigerina ciper-oensis, Globigerina angulisuturalis and Globigeri-noides primordius.

The macrofossil assemblage from the upper50 m in this section is similar to the assemblage inthe uppermost 40 m of the eastern transect in theMus basin (Figs 3 & 4). Both assemblages bear atypical late Chattian pectinid fauna with Amussio-pecten labadyei and A. subpleuronectes. PectenarcuatusQ11 (Brocchi 1814), a widespread Oligocenespecies, a typical Western Tethys element, ispresent as well, along with the thin-shelled lucinidbivalveQ12 Anodontia globulosa (Deshayes, 1830).The dominance of such thin shelled species mightindicate a slightly deeper environment than in thecorresponding section of the Mus basin, yet notdeeper than the medium deep sublittoral environ-ment (Mandic and Piller 2001).

The LO of Turborotalia ampliapertura has beencalibrated to 30.3 Ma within Chron 11r (Berggrenet al. 1995). The highest occurrence of this speciesin level 287 m (TR 244) therefore suggests an ageolder than 30.3 Ma for the bottom of the section.The LO of Paragloborotalia opima opima in level1445 m (TR 293) has been recently recalibrated to27.456 Ma within Chron 9n at ODP Site 1218(Wade et al. 2007). This age provides a maximumage for the top of the section since the highest occur-rence of Paragloborotalia opima opima occurs nearthe top of the section (TR 293). A correlation of theupper 50 m of the section to biozone P22 is sup-ported by the mollusc fauna which indicates a lateChattian age.

Paleoenvironmental interpretations

for the Elazıg basin

The depositional environment during the lowerRupelian (biozone P19) was first shallow marineas indicated by the occurrence of limestone withcorals, bivalves and gastropods.

However, the depositional environment rapidlydeepened as indicated by benthic foraminiferaldepth marker species (Cibicides (pseudo-)ungerianus,Gyroidina spp. Uvigerina spp. and occasionallyOridorsalis spp.). Their presence up to the top indi-cates that the basin was 350 to 700 m deep duringmuch of the Oligocene. The benthic foraminifersdo not give any evidence for shoaling, although

the limestone deposits at the top of the section andtheir macrofossils indicates a medium to shallowsubtidal environment for the late Chattian.

Implications for the Elazıg basin

The first 260 m of the studied section was depositedunder shallow marine conditions during Rupelian(biozone P19). This was followed by a rapid deepen-ing during the Rupelian and the deposition of about1.3 km in a relatively deep marine (300 to 750 m)environment. During the late Chattian (biozoneP22), the basin experienced rapid shoaling tomedium deep sublittoral conditions, preferred con-ditions for echinoids and bivalves. The inferredlate Chattian age of the macrofossils in the top ofthe section indicates that the final emergence ofthe basin must have occurred shortly after the Chat-tian followed by widespread Miocene volcanism.

The numerous internal slumping and sandstonelayers, referred to as turbidites, indicate a submar-ine, unstable slope. The entire succession is inter-preted as flysch deposited in a deep marine basin,comparable to the Mus basin. Thus, during theOligocene, rapid (15–30 cm/ka) sedimentation ofclay and turbidites dominated the basin evolution.

These new biostratigraphic ages differ signifi-cantly from the geological map of Turkey (Senel2002) where these sediments are indicated as Lowerto Middle Eocene. Our data suggests instead thatthese sediments were deposited under deep marineconditions during the Oligocene, from the Rupelianuntil the late Chattian, and, additionally, the shallowmarine limestones at the top of the section are lateChattian in age. This data also differs from previousstudies in the area (e.g. see Aksoy et al. 2005 for acompilation of data from the Elazıg basin) wherethe Eocene time has been identified as the mainperiod of deep marine deposition and in the Oligo-cene time shallow marine deposits were restrictedto the NW of the Elazıg basin. Our data howeverindicates that at least the eastern part of the Elazıgbasin was deep marine throughout the Oligoceneand shoaled and emerged only in the late Chattian,latest Oligocene.

Geological setting of the Kahramanmaras

basin

The Kahramanmaras basin is located near the triplejunction of the Arabian, African and Anatolianplates. As a result of the collision of Arabia andEurasia along the Bitlis Suture a trough formed infront of the thrust sheets and was consequentlyfilled by thick alluvial sediments and thick turbiditicflyschsequences (LiceFormation) (Sengor&Yılmaz1981; Perincek & Kozlu 1984; Karig & Kozlu 1990;


581582583584585586587588589590591592593594595596597598599600601602603604605606607608609610611612613614615616617618619620621622623624625626627628629630631632633634635636637638

Yılmaz 1993). According to several studies(Perincek 1979; Perincek & Kozlu 1984), theKahramanmaras basin was part of this elongatedforeland basin extending from Hakkari in southeast-ern Turkey, close to the border to Iran and Iraq, toAdana in southern Turkey (Fig. 2). This basin wasalso called the Lice trough (Dewey et al. 1986;Karig & Kozlu 1990; Derman & Atalik 1993;Derman 1999). Eocene deposits in the Kahraman-maras area are part of the Arabian Platform(Robertson et al. 2004). They indicate a shallowmarine depositional environment with local terres-trial input followed by allegedly lower to middleMiocene reefal limestone and claystone (Gul et al.2005). Oligocene bioclastic limestones are reportedonly from the margin of the Kahramanmaras area(Fig. 7) (Karig & Kozlu 1990). Basal shallow marinered-bed and basalt sequences of the Kalecik

Formation have an inferred age of late Burdigalianto Langhian (Karig & Kozlu 1990). The retreat ofmarine conditions and basin deformation wasassumed to have taken place in the late Miocene,although the age control was not documented(Karig & Kozlu 1990).

Three separate sections (Figs 7, 8a & b), all northof the city of Kahramanmaras, are been studied byus. The lower 200 m were sampled in the hills inthe southern part of the main basin (Hill section),the following about 4.6 km along the road north ofKahramanmaras (Road section) and the upper1.5 km stratigraphic transect near the village ofAvcılar (Avcılar section).

The base of the Hill section consists of nummu-litic limestones according to the Geological Map ofTurkey (Senel 2002) of Eocene age, followed by red,conglomeratic sediments with several basalt layers.

Fig. 7. Simplified tectonic and geological map of the Kahramanmaras area including the trajectories of the threestudied sections north of the city of Kahramanmaras: (1) the lowermost 200 m in the Hill section. (2) about 4.6 km ofsuccession along the road (Road section). (3) the upper 1.6 km up to the thrust studied in the Avcılar section in thenorthernmost part of the region. Refer to Legend for key to lithologies and/or ages (drawn after Geological mapof Turkey (Senel 2002)).


639640641642643644645646647648649650651652653654655656657658659660661662663664665666667668669670671672673674675676677678679680681682683684685686687688689690691692693694695696

Fig. 8. (a) Lithological column of the Hill section and part of the Road section (c. 3 km) in the Kahramanmaras basinwith the biostratigraphic results. The age model is based on planktonic foraminifer occurrences, which delineate thecorrelation of the Hill section to the Langhian and the first c. 3 km of the Road section to the Serravallian.


697698699700701702703704705706707708709710711712713714715716717718719720721722723724725726727728729730731732733734735736737738739740741742743744745746747748749750751752753754

Fig. 8. (Continued ) (b) Lithological column of the upper c. 1.6 km of the Road section and the Avcılar section in theKahramanmaras basin with the biostratigraphic results. The age model is based on planktonic foraminifer occurrences,which delineate the correlation of the upper c. 1.6 km of the Road section to the Serravallian and the lower part of theAvcılar section to the Serravallian, probably overlapping with the Road section, and from c. 700 m to the Tortonianbased on the LCO 0of Globigerinoides subquadratus. Refer to Legend in Figure 4 for key to lithologies, structures andfossils. KM ¼ Kahramanmaras.


755756757758759760761762763764765766767768769770771772773774775776777778779780781782783784785786787788789790791792793794795796797798799800801802803804805806807808809810811812

The studied section begins with a 200 m thick suc-cession of nodular limestone (10–15 m thick) alter-nating with bluish marls. The limestones containmacrofossils such as corals, sponges, echinoderms,bivalves and gastropods, indicating a shallowmarine environment. This succession grades intoan alternation of marl and sandstones layers,which show typical Bouma sequences and flutecasts, which are indicative for a turbiditic origin.

The base of the Road section, however, exposesa strongly different succession, where almost 1 kmof the stratigraphy is dominated by large conglomer-ate lenses. This thick succession of conglomeratescontains sand lenses showing cross bedding, indica-tive of interfingering of braided river channels. Thisthick fluvial succession probably forms, at least inpart, the lateral equivalent of the shallow marinesuccession in the Hill section. This conglomeraticsuccession is followed by a level rich in oysters,indicating a transition into shallow marine con-ditions. This then grades quickly into a very thicksuccession of alternating marl and sandstone (asmentioned above) with occasional conglomeraticlayers, interpreted as debris flows, which cutthrough the stratigraphy. The sandstones showtypical characteristics for turbidites, such as flutecasts and Bouma sequences. Some intervals aredominated by massive sandstone layers and/ordebris flows, while others are characterized bymainly clay. Slumping can be easily differentiatedfrom (minor) internal folding, both occurringthroughout the section. Internal folding, however,does not occur often. The Road section ends at thehighest point in the topography along the roadgoing NNW from Kahramanmaras and since thestratigraphy dips to the NW, a continuation of thestratigraphy was found to the NE, the Avcılarsection (Fig. 7). This section was sampled assumingsufficient overlap with the Road section, until thestratigraphy was cut unconformably by carbonates(Figs 8a & b). The stratigraphy of the Avcılarsection also consists of a thick succession (about1.5 km) of mudstone and sandstone with occasionalconglomeratic layers, interpreted as debris flows.No shoaling trend based on sedimentologicalcharacteristics has been observed towards the topof the section, which ends abruptly with the over-thrusting of pre-Neogene carbonates, which wereemplaced roughly from North. The upper 400 mwere not exposed, except for a few meters justunderneath the thrust (Fig. 8b).

Biostratigraphic results of the

Kahramanmaras basin

About every 20 m hand samples were taken for bios-tratigraphy. Only few samples, however, turned out

to be useful for biostratigraphy and/or paleobathy-metry. Benthic foraminifers in the lower part ofthe Kahramanmaras basin, from the Hill section(Fig. 8a) are dominated by milliolids and represen-tatives of Ammonia, Textularia, Nonion andElphidium indicating shallow marine (inner shelf)conditions although some samples, at 146, 182 and198 m (TR 9, 12 and 14), respectively, containfew planktonic foraminifers such as Globigeri-noides trilobus, Globigerinoides obliquus andOrbulina. Their presence would indicate that thelower part of the Kahramanmaras sequence post-dates the Orbulina datum at 14.74 Ma (Lourenset al. 2004). This age assignment is further con-straint by the presence of the calcareous nannofossilCyclicargolithus abisectus and rare Spenolithusheteromorphus along with the absence of Helico-sphaera ampliaperta. This assemblage is tentativelyassigned to NN5 (Martini 1971) indicating aLanghian age.

Orbulina is common in the samples from theRoad section. The rare occurrences of Globorataliapartimlabiata in the Road section at 165, 175, 2725,2788 and 2875 m (TR 20, 21, 71, 73 and 77),respectively, are remarkable because they representthe first recording of this species in Turkey. It hasbeen first described from the middle Miocene ofSicily (Ruggieri & Sprovieri 1970) and since thenreported from the Mediterranean (o.a. Foresi et al.

Q131998 and references herein; Hilgen et al. 2000;Turco et al. 2001; Foresi et al. 2002a, b; Hilgenet al. 2003; Abels et al. 2005) and adjacent NorthAtlantic (Chamley et al. 1986) and even from theIndian Ocean off NW Australia (Zachariasse1992). Ages of FO and LO of Globoratalia partim-labiata in the Mediterranean have recently beenrecalibrated at 12.771 and 9.934 Ma (Husing et al.2007). Its presence in the basal part of the Roadsection along with Globigerinoides subquadratusat 4224 m (TR 114) indicates that the larger partof the Road section, up to 4200 m, is Serravallian,Middle Miocene, in age, since the base of the Torto-nian has been defined at a level close to the LastCommon Occurrence (LCO) of Globigerinoidessubquadratus (Hilgen et al. 2000, 2005) with anew astronomical age of 11.625 Ma (Husing et al.2007). It cannot be excluded that the Road sectionterminates into the lowermost Tortonian sinceParagloborotalia siakensis at 4532 m (TR 120) isthe only biostratigraphic marker species presentabove 4200 m.

In the Avcılar section, the occasional occur-rences of Paragloborotalia siakensis up to 710 m(TR151) along with Globorotalia partimlabiata at590 m and Globigerinoides subquadratus at 710and 730 m, respectively, (TR 151 and 152) indicatethat the lower 700 m of this section is also Serraval-lian in age. The absence of Globigerinoides


813814815816817818819820821822823824825826827828829830831832833834835836837838839840841842843844845846847848849850851852853854855856857858859860861862863864865866867868869870

subquadratus and Paragloborotalia siakensis in theupper part of the Avcılar section along with the pre-sence of Globorotalia partimlabiata near the top ofthe section (TR 169 and 172) suggests that thesection extends up into the Tortonian.

The Avcılar section has been sampled assuminga significant overlap with the Road section and if theuppermost part of the Road section indeed extendsinto the Tortonian, we might assume an overlapof up to 1 km between these two sections. Themaximum age range of the road section andAvcılar section is indicated by the age range ofGloboratalia partimlabiata of 12.771–9.934 Ma(Husing et al. 2007).

Paleoenvironmental interpretations

for the Kahramanmaras basin

Deposition of the lower 200 m occurred in shallowmarine conditions, as indicated by the occurrenceof benthic foraminifers, calcareous nannofossils,poorly preserved echinoderms, gastropods and thelarge estuarine oyster Crassostrea gryphoides(Schlotheim, 1820).Q14 This large-sized Oligocene toMiocene species is restricted to brackish waterenvironments with a high nutrient input andprefers building colonies on mud flats of outer estu-aries (Slack-Smith 1998). Benthic foraminiferalspecies of the flysch succession from the road andAvcılar, such as Cibicides (pseudo-)ungerianus,Gyroidina spp., Uvigerina spp., Oridorsalis spp.and occasionally Siphonina reticulata, suggestwater depths between 350 and 750 m during depo-sition of this section without evidence for shoalingtowards the top, which is, in turn, cut by the thrustin the Avcılar section.

Implications for the Kahramanmaras basin

During the Langhian – early Serravallian, shallowmarine conditions prevailed in the Kahramanmarasbasin. The basin deepened during late Langhian/early Serravallian as indicated by the change fromlimestones and/or conglomerates to an alternationof marl and turbidites. Since neither in the lithology,nor in the biostratigraphic data, a shoaling trendtowards the top of the section is observed, deepmarine conditions (350–750 m) prevailed in thebasin until the early Tortonian.

We interpret the whole section as a characteristicforeland basin flysch succession (as Dewey 1986;Karig & Kozlu 1990; Derman & Atalik 1993;Derman 1999). Assuming the Road and Avcılar sec-tions were sampled with no overlap, the maximumthickness is about 6.1 km, but assuming an overlapof up to 1 km, the maximum thickness is about5.1 km. It is very difficult to estimate a

sedimentation rate for this basin, because three sec-tions were sampled with an unknowing overlap.Secondly the accuracy of the age indicative biostra-tigraphic events is uncertain due to poor preser-vation and poor sampling resolution. Furthermore,the age indicative biostratigraphic events, LCO ofGlobigerinoides subquadratus and LO of Globoro-talia partimlabiata are recorded in different sec-tions, which makes the determination of thesedimentation rate between these two calibrationpoints nearly impossible. The sedimentation ratesthus vary much, between 50 and 450 cm/ka, butincluding slumps, debris flows and turbidites depos-ited in front of and during the activity of the thrustthat now covers the top of the sequence. Taking aconservative estimate of 100 to 200 cm/ka, alsobecause the LO of Globigerinoides subquadratusmight not correspond to the true LCO, dated at11.625 Ma (Husing et al. 2007), but might behigher in the stratigraphy, the age of the youngestflysch is about 11 Ma.

This age range, from Langhian to early Torto-nian, differs significantly from the assigned Oligo-cene age of the open marine flysch and limestonedeposits in the Mus and Elazıg basins. The continu-ous marine sedimentation in the Kahramanmarasbasin from Langhian to early Tortonian at a constantdepth indicates that tectonic subsidence, possibly upto the order of 5 km, dominated the evolution ofthe basin.

Discussion

Evolution of the east Anatolian basins

The stratigraphic results from the east Anatolianbasins are summarized in Figure 9, and are corre-lated to the Geological Time Scale (Gradsteinet al. 2005). This figure schematically illustratesthat the individual basins belong to two different,major basins: (1) a basin north of the Bitlis–Poturge Massif, encompassing the Elazıg and Musbasins, which was filled with clastic mass flowdeposits during the Rupelian and Chattian (Oligo-cene); (2) a basin south of the Bitlis–PoturgeMassif, a foreland basin which was filled withclastic sediments during the Langhian, Serravallianand early Tortonian (Middle and early LateMiocene).

The Mus and Elazıg basin, both north of theBitlis–Poturge Massif, show similar stratigraphicevolution during the Oligocene: Deepening of thebasin occurred during the Rupelian and deepmarine conditions (350–750 m) prevailed until theChattian. Both basins evidence a shoaling trendduring the Chattian. The macrofossil assemblagein the sandy limestones, such as molluscs and


871872873874875876877878879880881882883884885886887888889890891892893894895896897898899900901902903904905906907908909910911912913914915916917918919920921922923924925926927928

echinoderm, indicates shallow marine, tropical tosubtropical deposition, similar to a sheltered lagoonenvironment, with species preferring subtidal andintertidal environments. In addition, the macrofossilassemblage is comparable to assemblages found in

Central Iran, the entire northern coast of theWestern Tethys, the Central Paratethys and thelower Indus Ocean, indicating an open marine con-nection between these marine realms prior to theemergence during the late Chattian. We suggest

Fig. 9. Chronology of Paleogene–Neogene foreland basin development in Southeastern Turkey. The evolutionof the three basins from this study, Mus, Elazıg and Kahramanmaras, is compared to the Hatay and Lice regions. Forpurpose of comparison the stratigraphy of all areas has been simplified. Refer to Legend for key to lithologiesand structures. Dashed lines of the lithological columns indicate uncertain ages for the section. All ages are given in Ma.


929930931932933934935936937938939940941942943944945946947948949950951952953954955956957958959960961962963964965966967968969970971972973974975976977978979980981982983984985986

that the Mus and Elazıg basins were connectedforming a large East–West elongated deep marinebasin during Rupelian and Chattian.

The rapid deepening of the basin north of theBitlis Massif may be related either to onset offlexural subsidence associated with (northward)underthrusting within the prevailing overall com-pressional regime (e.g. for the Elazıg basinQ15 Cronin(2000)) or to the late stages of an extensional defor-mation period that persisted in the Paleocene andEocene (e.g. in the Malatya basin (Kaymakci et al.2006) and in the Mus area (Sengor et al. 1985)).These two scenarios are controversial and our dataprovide age and depth constraints on the Mus andElazıg basins, which do not allow to eliminate orprefer either of these scenarios.

In the Kahramanmaras basin, south of theBitlis–Poturge Massif, shallow marine sedimentswere deposited during the Langhian. A rapid dee-pening during the Langhian to Serravallian indi-cated by the rapid transition to deep marine (350to 750 m) flysch deposits, was followed by depo-sition of continuously deep marine sediments untilthe early Tortonian. Since no shoaling trend isobserved we suggest that the age of the youngestflysch underneath the thrust, biostratigraphicallydated as early Tortonian, at about 11 Ma, coincideswith the end of underthrusting.

The rapid deepening of the foreland basin southof the Bitlis–Poturge Massif during the Langhian toSerravallian, followed by the deposition of a thickdeep marine flysch succession, can be interpretedas northern Arabia and more specifically the areaof the Kahramanmaras basin, entering into the sub-duction zone underneath Anatolia. The end of flyschdeposition and thus the youngest flysch underneaththe thrust of the overriding Bitlis–Poturge Massifcould be envisaged as the end of subduction, thusunderthrusting, at about 11 Ma (Tortonian), whichis likely followed by rapid uplift in the region.Such episodes of very rapid uplift and folding offoreland basins associated with the stalling of under-thrusting is, for example, also well documented inthe western Aegean region (Richter et al. 1978;van Hinsbergen et al. 2005c, d).

Our new results of the Kahramanmaras basin canbe compared to previously published data from theHatay (around Antakya) and Lice regions (seeFigs 2 & 9), which have been interpreted as forelandbasins related to southward thrusting of the Taurusallochton over the Arabian continental marginbelonging to an East–West elongated forelandbasin overlying the Arabian promontory (e.gPerincek 1979; Karig & Kozlu 1990; Derman andAtalik 1993; Derman 1999; Robertson 2004).Q16

The stratigraphy and chronology of the Hatayarea is very similar to the evolution of the Kahra-manmaras basin. The chronology in the Hatay area

has recently been redefined based on micropaleon-tological dating (Boulton et al. 2007) and we cantherefore correlate the evolution of the Kahraman-maras basin to the Hatay area. The stratigraphy inthe Hatay area is characterized by a pronouncedangular unconformity between middle Eocene andoverlying lower Miocene sediments, with a hiatusin the Oligocene (Boulton & Robertson 2007).Sedimentation resumed during the Aquitanian toBurdigalian (Early Miocene) with deposition ofconglomerates and mudstones. In the Kahraman-maras area, Derman & Atalik (1993) and Derman(1999) assigned a lower Miocene age to the about1 km thick series of fluvial deposits, which precedethe thick flysch deposits. We, however, have no ageconstraints on the fluvial deposits and can thereforenot confirm an Early Miocene age. During theLanghian both basins experienced shallow marinelimestone deposition and the basin progressivelydeepened during Serravallian to Tortonian (Boultonet al. 2007; Boulton & Robertson 2007). The depo-sition of shallow marine limestones in the Hatayarea have been interpreted to be related to furtherloading of the lithosphere in response to flexuralsubsidence and the progressive deepening to flex-ural control (Boulton & Robertson 2007). Whereflexural subsidence exceeded the build up of a car-bonate platform hemipelagic sediments were depos-ited (Boulton et al. 2007; Boulton & Robertson2007). A similar scenario can be envisaged for theKahramanmaras basin indicated by coarseningupwards in the thick flysch deposition. In the Hatayarea, by the end of the Miocene, the tectonic regimechanged and the Pliocene–Quaternary HatayGraben structure was formed in a transtensionalsetting related to the EAF (Perincek & Cemen1990; Boulton et al. 2007; Boulton & Robertson2007), while deep marine sediments in the Kahra-manmaras basin were overthrusted already duringthe early Tortonian. This comparison might indicatea diachronous evolution of these two basins with theKahramanmaras basin emerging during the Torto-nian and the Hatay area remaining open marineuntil the deposition of Messinian evaporites, ordifferent basin evolutions due to the relativelywestern position of the Hatay area thus closer tothe present-day extent of the eastern Mediterranean.

A comparison to the Lice basin, which is situatedto the east of the Kahramanmaras basin, would evi-dently give constraints on the syn- or diachronousevolution of the southernmost, Arabian forelandbasin. However, the chronology of sediments inthe Lice basin is scarcely documented in the litera-ture (e.g. Perincek 1979; Dewey 1986; Karig &Kozlu 1990; Robertson et al. 2004). On the geologi-cal map of Turkey (Senel 2002) shallow marineclastic and carbonatic sediments have been indi-cated as Early Miocene in age and continental


987988989990991992993994995996997998999100010011002100310041005100610071008100910101011101210131014101510161017101810191020102110221023102410251026102710281029103010311032103310341035103610371038103910401041104210431044

clastic rocks as Middle to Late Miocene in age. Thissuccession would pre-date the flysch deposition inthe Kahramanmaras and Hatay area and would indi-cate diachronous evolution of the elongated Arabianforeland basin. Other studies assigned, howeverwithout documenting an age control, a Tortonianage to the Lice flysch (Dewey 1986). If the flyschdeposits in the Lice, Kahramanmaras and Hatayarea are indeed synchronous, we would assume asynchronous evolution of the Arabian forelandbasin which emerged during the Tortonian.However since the chronology of the Lice basin isnot well documented, firm correlation to the Kahra-manmaras basin and Hatay area remains impossible(see question marks in Fig. 9).

The basin south of the Bitlis–Poturge Massifincluding the Kahramanmaras, Hatay and Licebasins, is interpreted as the southernmost and

youngest foreland basin in the east Anatolian fold-and thrustbelt, which formed as a large East–Westtrending foreland basin on the subducting Arabianplate. The end of underthrusting in the Kahraman-maras basin is dated at about 11 Ma, but mighthave been diachronous relative to the emergenceof the Hatay and Lice basin.

Tectonic closure of the eastern Tethys gateway

Based on the presented data herein, we envisagethe following scenario for the Oligocene toMiocene evolution of the basins north and south ofthe Bitlis Massif in SE Turkey (Fig. 10). Duringthe early Oligocene, marine sediments were depos-ited in a large basin to the north of the Bitlis–Poturge Massif (Fig. 10a). However, our data doesnot allow us to constrain whether the deepening of

c .

c . c .

c .

c . c . c .(a1) (a2)

(b)

(c)

Fig. 10. The evolution of the Oligocene–Miocene basins in SE Turkey are illustrated schematically in three majorphases: (a) during the Oligocene from c. 30 to c. 23 Ma: a marine basin was situated north of the Bitlis–Poturge (BP)until the end of latest Oligocene, Chattian, when this basin emerged; (a1) related to extension, (a2) related to thrusting.(b) during the Langhian to early Tortonian (c. 13 to c. 11 Ma): areas of present-day northern Arabia enter the position ofthe foreland basin South of the BP and the northern Arabian promontory was subducted underneath the BP from theLanghian until the early Tortonian, c. 11 Ma, and finally. (c) since the early Tortonian to Recent: the end of large-scaleunderthrusting at c. 11 Ma in east Anatolia coincides with the onset of collision-related volcanism, uplift of the EastAnatolian Accretionary Complex (EAAC), the onset of shearing along the North and East Anatolian Faults (NAF andEAF). Refer to text for further discussion. BP ¼ Bitlis-Poturge; BPM ¼ Bitlis-Poturge Massif; NAF ¼ North AnatolianFault; EAF ¼ East Anatolian Fault; EAAC ¼ East Anatolian Accretionary Complex; KM ¼ Kahramanmaras.


1045104610471048104910501051105210531054105510561057105810591060106110621063106410651066106710681069107010711072107310741075107610771078107910801081108210831084108510861087108810891090109110921093109410951096109710981099110011011102

the basin during the Rupelian was related to eitherlarge scale extension (Fig. 10a1) or thrusting(Fig. 10a2), with the Bitlis–Poturge Massif situatedon the overriding plate. Nevertheless, our datasuggest that until the Chattian, flysch was depositedin a deep marine environment, as recorded in thearea of Mus and Elazıg. The emergence of thebasin north of the Bitlis–Poturge Massif duringthe late Chattian (see also Fig. 9) probably coincideswith the accretion of the Bitlis–Poturge Massif tothe Anatolian plate.

On the southern side of the Bitlis–PoturgeMassif, oceanic subduction was probably ongoingdue to Africa/Arabian’s relative distal position(e.g. Besse & Courtillot 2002).

During the Langhian to Serravallian the basinsouth of the Bitlis–Poturge Massif deepened rapidly,which might be related to the northern Arabian pro-montory, the present-day northern margin of theArabian plate (Kahramanmaras, Hatay and Licebasins), entering into the subduction zone (Fig. 10b)below the Bitlis–Poturge Massif. During the Serra-vallian and early Tortonian, the Kahramanmarasbasins remained deep marine indicated by thickflysch deposition until, at least, the early Tortonian.The youngest flysch underneath the thrust in the Kah-ramanmaras area, biostratigraphically dated at about11 Ma, might be linked to the end of the large-scaleunderthrusting (subduction) in eastern Anatolia(Fig. 10c). In models proposed by Keskin (2003)and Sengor (2003), it is assumed that the Bitlis–Poturge Massif was accreted with Arabia during lateEocene, while Robertson et al. (2004) suggestedLate Oligocene–earliest Miocene time. Our data, onthe other hand, indicate the presence of a deepmarine realm between the Bitlis–Poturge Massifand Arabia during Serravallian and early Tortonian,which we suggest is associated with the continuoussubduction of Arabia underneath the Anatolian plate.

The timing of the end of thrusting agrees with theonset of the collision-related volcanism at about11 Ma north of the present-day suture line (Keskin2003), the uplift of the East Anatolian AccretionaryComplex inferred to start around 11 Ma onwards(Sengor et al. 2003) and the onset of the Northand East Anatolian Fault (Dewey 1986; Hubert-Ferrari et al. 2002; Sengor et al. 2005) (Fig. 10c).Collision-related volcanism and uplift of the EastAnatolian High Plateau despite the relatively thincrust (45 km) has been related to an anomalouslyhot mantle underneath the eastern Anatolia(Keskin 2003; Sengor et al. 2003). This, as well asthe onset of westward extrusion of Anatolia andthe onset of formation of the North and East Anato-lian faults, have been explained by slab detachmentat about 11 Ma in eastern Anatolia (Keskin 2003;Sengor et al. 2003; Sengor et al. 2005; Faccennaet al. 2006), which is in line with a recent tomogra-phy study of Hafkenscheid et al. (2006). Our new

results from the Kahramanmaras area can thus beconsidered in line with the previously suggestedscenarios, the end of underthrusting and the onsetof extrusion of Anatolia in the Late Miocene (atabout 11 Ma) (Fig. 10c).

Constraints on the closure of the eastern

Tethys gateway

The continuous northward migration of the African–Arabian plate led to the disruption of the Tethysseaway and the final closure related to continentalcollision of Arabia and Eurasia. The paleogeographicextent of the Tethys during the Paleogene andNeogene thus underwent significant changes untilthe connection was finally closed. Several authorssuggested that the final closure of the easternTethys gateway may have resulted in significantchanges in the paleoceanographic circulation andconsequently in a major change in global climate(e.g. Woodruff & Savin 1989, 1991; Jacobs et al.1996; Flower & Kennett 1993; Yılmaz 1993).

Our data from eastern Anatolia indicate that adeep marine connection was present north of theBitlis–Poturge Massif from Rupelian to late Chat-tian. The shoaling of this northern basin during thelate Chattian led to severe disruption between aneastern (Indian Ocean) and western (Mediterranean)marine domain; particularly the deep-water circula-tion was disrupted during the Chattian. The emer-gence of this basin after the late Chattian resultedin the closure of at least this branch of the southernNeotethys and coincides with the late Oligocenewarming, reducing the extent of the Antarctic ice,which was punctuated by the Mi-1 glaciationsaround the Oligocene–Miocene boundary (Zachoset al. 2001).

Other studies suggest that the Tethys seaway wasopen until the early Miocene and became severelyrestricted during the Burdigalian (c. 19 Ma), whenmammal fauna and shallow marine macrofaunalrecords from the eastern Mediterranean region indi-cate the existence a landbridge (Gomphotheriumlandbridge) connecting Africa/Arabia and Eurasia(Popov 1993; Rogl 1998, 1999; Harzhauser et al.2002, 2007). These authors claim that, sincec. 19 Ma, biogeographic separation between theMediterranean-Atlantic and Indo-Pacific regionspersisted; despite some short-lived periodic marineconnections between the two domains until themiddle Miocene (Rogl 1999; Meulenkamp &Sissingh 2003; Golonka 2004; Harzhauser et al.2007). If a causal link between the closure of theeastern Tethys gateway and global climate coolingexists, a major change in global, or at least local,climate must be expected during the Burdigaliantime. The most significant climatic change duringthe Burdigalian, as evidenced in both the d18O and


1103110411051106110711081109111011111112111311141115111611171118111911201121112211231124112511261127112811291130113111321133113411351136113711381139114011411142114311441145114611471148114911501151115211531154115511561157115811591160

d13C record, indicates a change from a cooling to awarming trend which led into the Mid-Miocene Cli-matic Optimum (Zachos et al. 2001).

Our data suggest that if a deep marine connectionbetween the eastern and western marine realm per-sisted after the late Chattian, it was probablylocated south of the Bitlis–Poturge Massif. Thestudied basins along the south Bitlis suture zone ineastern Anatolia, however, do not comprise thestratigraphic interval between late Chattian andLanghian (25–15 Ma). Consequently, it is not poss-ible to constrain the tectonic evolution and thepalaeogeographic extent of the Tethys seawayduring this time interval from the stratigraphicrecord of the east Anatolian basins.

In the context of global climate change, the mainoxygen and carbon isotope shift corresponds to thesecond and major step (Mi3b) of the middleMiocene global cooling, and has recently beenastronomically dated at 13.82 Ma, close to theLanghian–Serravallian boundary, in a section onMalta (Abels et al. 2005). The middle Miocenedecrease in d18O values was previously attributed,amongst other hypotheses, to a possible localexpression of the isolation of the MediterraneanSea from the Indo-Pacific Ocean (van der Zwaan& Gudjonsson 1986; Jacobs et al. 1996). Abelset al. (2005), however, show that this eventcoincides with a period of minimum amplitudesobliquity related to the 1.2-Ma cycle andminimum values of eccentricity as part of both the400– and 100-ka eccentricity cycle, thus suggestingastronomical forcing (see Abels et al. 2005).

If a link between gateway closure and middleMiocene climate change exists, the south Bitlisgateway must have re-opened to finally close inthe middle Miocene, which is very unlikely in anoverall converging setting. Moreover, our datadoes not show evidences for a final closure of theseaway in the middle Miocene. In contrast, thedata from Kahramanmaras indicates rapid deepen-ing during the Langhian to Serravallian and prevail-ing deep marine conditions until the earlyTortonian. This has been interpreted as related tocontinuous northward subduction underneath theBitlis–Poturge Massif and finally continental col-lision during the Serravallian to early Tortonian.Our data suggest that a deep marine connectionlocated between the Bitlis–Poturge Massif andArabia, whether periodic or not, was disrupted atlatest during the early Tortonian, giving an upperlimit of c. 11 Ma to the final closure between theIndian Ocean and the Mediterranean along thenorthern Arabian. The above analysis shows thatthe end of foreland basin existence in SE Turkey –and therefore the closure of the southern Tethyangateway – can not straightforwardly be linked tothe middle Miocene climate change. Future assess-ment of the timing of the Tethys gateway closure

should focus on detailed stratigraphy of the young-est foreland basins in SE Turkey, NW Iran, Syriaand N Iraq, the region of the Bitlis–Zagrossuture zone.

Conclusions

The marine basin north of the Bitlis–Poturge Massifencompassing the Elazıg and Mus basins emergedduring Chattian, which was followed by shallowmarine limestone deposition during the late Chattianand finally closed after the late Chattian. This marksthe disruption of the Tethys gateway north of theBitlis–Poturge Massif connecting an eastern(Indian Ocean) and western (Mediterranean Sea)domain.

The Kahramanmaras basin south of the Bitlis–Poturge Massif, probably linked to the Hatay andLice basins, experienced shallow marine conditionsduring Langhian, rapidly deepened during Langhianto Serravallian and remained deep marine during theSerravallian and early Tortonian. No shoaling trendhas been observed in the Kahramanmaras basin andthe age of the youngest flysch underneath thesubduction-related thrust has been biostratigraphi-cally dated at early Tortonian, at about 11 Ma. Theend of flysch deposition in the Kahramanmarasarea is probably related to the end of subduction,thus the end of underthrusting. The age coincideswith the onset of collision-related volcanism, upliftof the East Anatolian Accretionary Complex, andthe timing of shearing along the NAF and EAF.Our new results suggest a strong link between theprocesses outlined above, which have beenexplained by slab detachment at about 11 Ma ineastern Anatolia.

This age, early Tortonian, about 11 Ma, is theyoungest possible age for a deep marine connectionbetween the Mediterranean-Atlantic and Indo-Pacific regions. We can thus constrain the timingof the final closure of a deep marine Tethysgateway to an upper limit of about 11 Ma. The emer-gence of the basin north of the Bitlis–Poturge Massifduring the late Chattian thus provides a lower limit ofthe closure of the eastern Tethys gateway.

In the southern basins marine foreland depo-sition was continuous during Serravallian andearly Tortonian and our data does not support alink between the Middle Miocene climate coolingdated at 13.82 Ma and the closure of the easternTethys gateway. In contrast, the age of the youngestflysch deposits, thus the youngest foreland basin inSE Turkey is early Tortonian, about 11 Ma.

We thank C. Langereis, N. Kaymakci and E. Yılmaz fortheir help in the field and M. Triantaphyllou for the deter-mination of nannofossils in the Kahramanmaras section.We are also grateful for the constructive comments ofE. Turco and N. Kaymakci, which led to a significantly


1161116211631164116511661167116811691170117111721173117411751176117711781179118011811182118311841185118611871188118911901191119211931194119511961197119811991200120112021203120412051206120712081209121012111212121312141215121612171218

improved manuscript. We acknowledge support by theNetherlands Research Centre for Integrated Solid EarthSciences (ISES) and by the Netherlands GeosciencesFoundations (ALW) with financial aid from theNetherlands Organization for Scientific Research (NWO).This work was carried out under the program of theVening Meinesz Research School of Geodynamics(VMSG) and the Netherlands Research School of Sedi-mentary Geology (NSG).

References

ABELS, H. A., HILGEN, F. J., KRIJGSMAN, W., KRUK,R. W., RAFFI, I., TURCO, E. & ZACHARIASSE, W. J.2005. Long-period orbital control on middle Mioceneglobal cooling: Integrated stratigraphy and astronomi-cal tuning of the Blue Clay Formation on Malta. Paleo-ceanography, 20, 1–17.

AGARD, P., OMRANI, J., JOLIVET, L.Q17 & MOUTHEREAU.2005. Convergence history across Zagros (Iran):constraints from collisional and earlier deformation.International Journal of Earth Sciences, 94, 401–419.Q7

AKSOY, E., TURKMEN, I. & TURAN, M. 2005. Tectonicsand sedimentation in convergent margin basins: anexample from the Tertiary Elazig basin, EasternTurkey. Journal of Asian Earth Science, 25, 459–492.

AKTAS, G. & ROBERTSON, A. H. F. 1984. The MadenComplex, SE Turkey: evolution of a Neotethyanactive margin. In: DIXON, J. E. & ROBERTSON,A. H. F. (eds) The Geological Evolution of theEastern Mediterranean. Blackwell Scientific Publi-cations, Oxford, 375–402.

ALLMENDINGER, R. W., REILINGER, R. & LOVELESS, J.2007. Strain and rotation rate from GPS in Tibet,Anatolia and the Altiplano. Tectonics, 26, TC3013,doi: 10.1029/2006TC002030.

BALDI, T. 1973. Mollusc fauna of the Hungarian igocene(Egerian). Budapest, Akademiai Kiado.

BANDEL, K. & WEDLER, E. 1987. Hydroid, amphineuranand gastropod zonation in the littoral of the CaribbeanSea, Colombia. Senckenbergiana maritima, 19,1–130.

BARKA, A. 1992. The North Anatolian fault zone. Annalsof Tectonics, 6, 164–195.Q18

BARTOLI, G., SARNTHEIN, M., WEINELT, M.,ERLENKEUSER, H., GARBE-SCHONBERG, D. &LEA, D. W. 2005. Final closure of Panama and theonset of northern hemisphere glaciation. Earth andPlanetary Science Letters, 237, 33–44.

BELLAHSEN, N., FACCENNA, C., FUNICIELLO, F.,DANIEL, J.-M. & JOLIVET, L. 2003. Why did Arabiaseparate from Africa? Insights from 3-D laboratoryexperiments. Earth and Planetary Science Letters,216, 365–381.

BERCKHEMER, H. 1977. Some aspects of the evolution ofmarginal seas deduced from observations in theAegean region. In: MONTADERT, L. (ed.) StructuralHistory of the Mediterranean Basins. Technip. Paris,Split, Yugoslavia, 303–313.

BERGER, W. H. & MILLER, K. G. 1988. Paleogenetropical planktonic foraminiferal biostratigraphy andmagnetobiostratigraphy. Micropaleontology, 34,362–380.

BERGGREN, W. A., KENT, D. V., SWISHER III, C. C. &AUBRY, M. P. 1995. A revised Cenozoic Geochro-nology and Chronostratigraphy. SEPM Special Publi-cation, 54, 129–212.

BESSE, J. & COURTILLOT, V. 2002. Apparent and truepolar wander and the geometry of the geomagnetic fieldover the last 200 Ma. Journal Geophysical Research,107 (B11), 2300, doi:10.1029/2000JB000050.

BICCHI, E., FERRERO, E. & GONERA, M. 2003. Paleocli-matic interpretation based on Middle Miocene plank-tonic foraminifera: the Silesia Basin (Paratethys) andMonferrato (Tethys) records. Paleogeography, Paleo-climatology, Paleoecology, 196, 265–303.

BINGOL, E. 1984. Geology of the Elazig area in theEastern Taurus region. In: Geology of the TaurusBelt. MTA (Turkish Geological Survey), Ankara,209–216.

BLOW, W. H. 1969. Late middle Eocene to Recent plank-tonic foraminiferal biostratigraphy. In: BRONNIMANN,P. & RENZ, H. H. (eds) Proceedings of the First Inter-national Conference on Planktonic Microfossils(Geneva, 1967), Volume 1: E. J. Brill, Leiden,199–421.

BOULTON, S. J. & ROBERTSON, A. H. F. 2007. TheMiocene of the Hatay area, S Turkey: Transitionform the Arabian passive margin to an underfilledforeland basin related to the closure of the SouthernNeotethys Ocean. Sedimentary Geology, 198, 93–124.

BOULTON, S. J., ROBERTS, A. H. F., ELLAM, R. M.,SAFAK, U. & UNLUGENC, U. C. 2007. Strontiumisotopic and micropaleontological dating used to helpredefine the stratigraphy of the neotectonic HatayGraben, Southern Turkey. Turkish Journal of EarthSciences, 16, 141–179.

BOZKURT, E. 2003. Origin of NE-trending basins inwestern Turkey. Geodinamica Acta, 16, 61–81.

CHAMLEY, H., MEULENKAMP, J. E., ZACHARIASSE,W. J. & VAN DER ZWAAN, G. J. 1986. Middle tolate Miocene marine ecostratigraphy: clay mineral,planktonic foraminifera and stable isotopes fromSicily. Oceanological Acta, 9, 227–238.

CRONIN, B. T., HARTLEY, A. J., CELIK, H., HURST, A.,TURKMEN, I. & KEREY, E. 2000a. Equilibriumprofile development in graded deep-water slopes:Eocene, Eastern Turkey. Journal of the GeologicalSociety, London, 157, 946–955.

CRONIN, B. T., HURST, A., CELIK, H. & TURKMEN, I.2000b. Superb exposure of a channel, levee and over-bank complex in an ancient deep-water slope environ-ment. Sedimentary Geology, 132, 205–216.

DERMAN, A. S. 1999. Braided river deposits related toprogressive Miocene surface uplift in KahramanMaras area, SE Turkey. Geological Journal, 34,159–174.

DERMAN, A. S. & ATALIK, E. 1993. Sequence Strati-graphic Analysis of Miocene Sediments in MarasMiocene Basin and Effect of Tectonism in the Devel-opment of Sequences. Special Publications SequenceStratigraphy, Sedimentology Study Group, 1, 43–52[in Turkish].

DEWEY, J. F. & SENGOR, A. M. C. 1979. Aegean and sur-rounding regions: Complex multiplate and continuumtectonics in a convergent zone. Geological Society ofAmerica Bulletin, 90, 84–92.


1219122012211222122312241225122612271228122912301231123212331234123512361237123812391240124112421243124412451246124712481249125012511252125312541255125612571258125912601261126212631264126512661267126812691270127112721273127412751276

DEWEY, F. J., HEMPTON, M. R., KIDD, W. S. F.,SAROGLU, F. & SENGOR, A. M. C. 1986. Shorteningof continental lithosphere: the neotectonics of easternAnatolia – a young collision zone, In: COWARD,M. P. & RIES, A. C. (eds) Collision Tectonics,Geological Society, London, Special Publication,19, 3–36.

FACCENNA, C., BELLIER, O., MARTINOD, J.,PIROMALLO, C. & REGARD, V. 2006. Slab detach-ment beneath eastern Anatolia: A possible cause forthe formation of the North Anatolian Fault. Earthand Planetary Science Letters, 242, 85–97.

FLOWER, B. P. & KENNETT, J. 1993. Middle MioceneOcean-Climate Transition: High-Resolution Oxygenand Carbon Isotopic Records from Deep Sea DrillingProject Site 558A, Southwest Pacific. Paleoceanogra-phy, 8, 811–843.

FLOWER, B. P. & KENNETT, J. P. 1994. The middleMiocene climate transition, East Antarctic ice sheetdevelopment, deep ocean circulation and globalcarbon cycle. Paleogeography, Paleoclimatology,Paleoecology, 108, 537–555.

FORESI, L. M., IACCARINO, S., MAZZEI, R. &SALVATORINI, G. 1998. New data on calcareousplankton biostratigraphy of the middle-late Miocene(Serravallian/Tortonian) of the Mediterranean area.Reviews of Italian Paleontology and Stratigraphy,104, 95–114.Q18

FORESI, L. M., BONOMO, S., CARUSO, A., DI STEFANO,E., SALVATORINI, G. & SPROVIERI, R. 2002a.Calcareous plankton high resolution biostratigraphy(foraminifera and nannofossils) of the uppermostLanghian-lower Serravallian ‘Ras-il-Pellegrin’section (Malta). Reviews of Italian Paleontology andStratigraphy, 337–353.Q18

FORESI, L. M., CARUSO, A., DI STEFANO, E., LIRER, F.,IACCARINO, S., SALVATORINI, G. & SPROVIERI, R.2002b. High-resolution biostratigraphy of the upperSerravallian/lower Tortonian sequence of the TremitiIslands. Reviews of Italian Paleontology and Stratigra-phy, 108,Q18 257–253.

GARFUNKEL, Z. 1998. Constrains on the origin andhistory of the Eastern Mediterranean basin. Tectono-physics, 298, 5–35.

GARFUNKEL, Z. 2004. Origin of the Eastern Mediterra-nean basin: a re-evaluation. Tectonophysics, 391,11–34.

GOLONKA, J. 2004. Plate tectonic evolution of thesouthern margin of Eurasia in the Mesozoic and Ceno-zoic. Tectonophysics, 381, 235–273.

GRADSTEIN, F. M., OGG, J. G. & SMITH, A. G. (eds)2005. A geological time scale 2004. CambridgeUniversity Press.

GUL, M., DARBAS, G. & GURBUS, K. 2005.Tectono-Stratigraphic position of Alacik Formation(Latest Middle Eocene – Early Miocene) in theKahraman Maras Basin. In Turkish with Englishabstract and summary. Istanbul University Muh. Fak.Yerbilimleri Dergisi, 18, 183–197.Q18

HAFKENSCHEID, E., WORTEL, M. J. R. & SPAKMAN, W.2006. Subduction history of the Tethyan region derivedfrom seismic tomography and tectonic reconstructions.Journal Geophysical Research, 111, doi: 10.1029/2005JB003791.

HARZHAUSER, M. 2004. Oligocene Gastropod Faunas ofthe Eastern Mediterranean (Mesohellinic Trough/Greece and Esfahan-Sirjan Basin/Central Iran).Courier Forschungsinstitut Senckenberg, 248, 93–181.

HARZHAUSER, M., KROH, A., MANDIC, O., PILLER,W. E., GOHLICH, U., REUTER, M. & BERNING, B.2007. Biogeographic responses to geodynamics: Akey study all around the Oligo-Miocene TethyanSeaway. Zoologischer Anzeiger, doi: 10.1016/j.jcz.2007.05.001.

HARZHAUSER, M., PILLER, W. E. & STEININGER,F. F. 2002. Circum-Mediterranean Oligo-Miocenebiostratigraphic evolution – the gastropods’ point ofview. Paleogeography, Paleoclimatology, Paleoecol-ogy, 183, 103–133.

HENDLER, G., MILLER, J. E., PAWSON, D. L. & KIER,P. M. 1995. Sea Stars, Sea Urchins and Allies –Echinoderms of Florida and the Carribbean. Smitho-nian Institute Press, Washington, London, 1–390.

HESSAMI, K., KOYI, H. A., TALBOT, C. J., TABASI, H. &SHABANIAN, E. 2001. Progressive unconformitieswithin an evolving foreland fold-thrust belt, ZagrosMountains. Journal of the Geological Society,London, 158, 969–981.

HILGEN, F. J., KRIJGSMAN, W., RAFFI, I., TURCO, E. &ZACHARIASSE, W. J. 2000. Integrated stratigraphyand astronomical calibration of the Serravallian/Tortonian boundary section at Monte Gibliscemi(Sicily, Italy). Marine Micropaleontology, 38, 181–211.

HILGEN, F. J., ABDUL AZIZ, H., KRIJGSMAN, W., RAFFI,I. & TURCO, E. 2003. Integrated stratigraphy andastronomical tuning of the Serravallian and lowerTortonian at Monte dei Corvi (Middle-UpperMiocene, northern Italy). Paleogeography, Paleocli-matology, Paleoecology, 199, 229–264.

HILGEN, F. J., ABDUL AZIZ, H., BICE, D., IACCARINO, S.,KRIJGSMAN, W., KUIPER, K. F., MONTANARI, A.,RAFFI, I., TURCO, E. & ZACHARIASSE, W. J. 2005.The Global Boundary Stratotype Section and Point(GSSP) of the Tortonian Stage (Upper Miocene) atMonte dei Corvi. Episodes, 28, 6–17.

HSU, K. J., RYAN, W. B. F. & CITA, M. B. 1973. LateMiocene desiccation of the Mediterranean. Nature,242, 240–244.

HUBERT-FERRARI, A., ARMIJO, R., KING, G. C. P.,MEYER, B. & BARKA, A. 2002. Morphology, displa-cement and slip rates along the North Anatolian Fault(Turkey). Journal Geophysical Research, 107,101029–101059. Q1

HUSING, S. K., HILGEN, F. J., ABDUL AZIZ, H. &KRIJGSMAN, W. 2007. Completing the Neogene geo-logical time scale between 8.5 and 12.5 Ma. Earthand Planetary Science Letters, 253, 340–358.

IGRS-IFP, 1966. Etude Geologique de l’Epire (Grecenord-occidentale). Paris, Institut Francais du Petrol.

JACOBS, E., WEISSERT, H. & SHIELDS, G. 1996. TheMonterey event in the Mediterranean: A record fromshelf sediments of Malta. Paleoceanography, 11,717–728.

JACOBSHAGEN, V. 1986. Geologie von Griechenland.Berlin-Stuttgart, Borntraeger. Q21

JOLIVET, L. 2001. A comparison of geodetic and finitestrain pattern in the Aegean, geodynamic implications.Earth and Planetary Science Letters, 187, 95–104.


1277127812791280128112821283128412851286128712881289129012911292129312941295129612971298129913001301130213031304130513061307130813091310131113121313131413151316131713181319132013211322132313241325132613271328132913301331133213331334

JOLIVET, L. & FACCENNA, C. 2000. Mediterranean exten-sion and the Africa-Eurasia collision. Tectonics, 19,1094–1106.

KARIG, D. E. & KOZLU, H. 1990. Late Paleogene–Neogene evolution of the triple junction region nearMaras, south-central Turkey. Journal of the GeologicalSociety, London, 147, 1023–1034.

KAYMAKCI, N., INCEOZ, M. & ERTEPINAR, P. 2006.3D-Architecture and Neogene Evolution of theMalatya Basin: Inferences from the Kinematics of theMalatya and Ovacik Fault Zones. Turkish Journal ofEarth Sciences, 15, 123–154.

KESKIN, M. 2003. Magma generation by slab steepeningand breakoff beneath a subduction-accretion complex:An alternative model for collision-related volcanismin Eastern Anatolia, Turkey. Geophysical ResearchLetters, 30, doi:10.1029/2003GL018019, 2003.

KRIJGSMAN, W., LANGEREIS, C. G., DAAMS, R. &VAN DER MEULEN, A. J. 1994. Magnetostratigraphicdating of the middle Miocene change in the continentaldeposits of the Aragonian type area in the Calayud-Terueal basin (Central Spain). Earth and PlanetaryScience Letters, 128, 513–526.

LE PICHON, X., ANGELIER, J. & SIBUET, J.-C. 1982.Plate boundaries and extensional tectonics. Tectono-physics, 81, 239–256.

LIVERMORE, R., NANKIVELL, A., EAGLES, G. &MORRIS, P. 2005. Paleogene opening of DrakePassage. Earth and Planetary Science Letters, 236,459–470.

LOURENS, L. J., HILGEN, F. J., SHACKLETON, N. J.,LASKAR, J. & WILSON, D. 2004. The NeogenePeriod. In: GRADSTEIN, F. M., OGG, J. G. & SMITH,A. G. (eds) A geological time scale 2004. CambridgeUniversity Press, 409–440.

MANDIC, O. 2000. Oligocene to Early Miocene pectinidbivalves of Western Tethys (N-Greece, S-Turkey,Central Iran and NE-Egypt) – taxonomy and paleobio-geography. Unpubl. PhD thesis, University of Vienna.

MANDIC, O. & PILLER, W. E. 2001. Pectinid coquinas andtheir paleoenvironmental implications – examplesfrom the early Miocene of northeastern Egypt. Paleo-geography, Paleoclimatology, Paleoecology, 172,171–191.

MARTINI, E. 1971. Standard Tertiary and Quaternary cal-careous nannoplankton zonation. In: FARINACCI, A.(ed.) Proceedings of the Second Planktonic ConferenceRoma, Edizioni Technoscienza, Roma. 739–785.Q19

MCCLUSKY, S., BALASSANIAN, S., BARKA, A., DEMIR,A., ERGINTAV, S., GEORGIEV, I. ET AL. 2000. GlobalPositioning System constrains on plate kinematics anddynamics in the eastern Mediterranean and Caucasus.Journal Geophysical Research, 105, 5695–5719.

MEULENKAMP, J. E. & SISSINGH, W. 2003. Tertiarypaleogeography and tectonostratigraphic evolution ofthe Northern and Southern Peri-Tethys platforms andthe intermediate domains of the African-Eurasian con-vergent plate boundary zone. Paleogeography, Paleo-climatology, Paleoecology, 196, 209–228.

MILLER, K. G., WRIGHT, J. D. & FAIRBANKS, R. G.1991. Unlocking the Ice House: Oligocene–Mioceneoxygen isotopes, eustacy, and margin erosion.Journal Geophysical Research, 96, 6829–6848.

MILLER, K. G., WRIGHT, J. D. & BROWNING, J. V. 2005.Visions of ice sheets in a greenhouse world. MarineGeology, 217, 215–231.

PEARCE, J. A., BENDER, J. F., DE LONG, S. E., KIDD,W. S. F., LOW, P. J., GUNER, Y., SAROGLU, F.,YILMAZ, Y., MOORBATH, S. & MITCHELL, J. G.1990. Genesis of collision volcanism in Eastern Anato-lia, Turkey. Journal of Volcanology and GeothermalResearch, 44, 189–229.

PERINCEK, D. 1979. The geology of Hazro–Korudag–Cungus–Maden–Ergani–Hazar–Elazig–MalatyaArea. The Geological Society of Turkey.

PERINCEK, D. & CEMEN, I. 1990. The Structural relation-ship between the Eastern Anatolian and Dead Sea faultzones in southeastern Turkey. Tectonophysics, 172,331–340.

PERINCEK, D. & KOZLU, H. 1984. Stratigraphy andstructural relations of the units in the Afsin-Elbistan-Dogansehir region (Eastern Taurus), In: Proceedingsof the international symposium, Geology of TaurusBelt, MTA, Ankara-Turkey, Q20181–198.

PERINCEK, D. & OZKAYA, I. 1981. Arabistan LevhasiKuyzey Kenari Tektonik evrimi. Tectonic evolutionof the northern margin of Arabian plate. Yerbilimleri,Hacettepe Universitesi, Bulletin of Institute of EarthSciences of Hacettepe University Beytepe Ankara,8, 91–101.

POPOV, S. V. 1993. Zoogeography of the Late EoceneBasins of Western Eurasia Based on Bivalve Mollusks.Stratigraphy and Geological Correlation, 2, 103–118.

REILINGER, R., MCCLUSKY, S., VERNANT, P.,LAWRENCE, S., ERGINTAV, S., CAKMAK, R. ET AL.2006. GPS constraints on continental deformation inthe Africa–Arabia–Eurasia continental collision zoneand implications for the dynamics of plate interactions.Journal of Geophysical Research, 111, V05411,doi:10.1029/2005JB004051.

REUTER, M., PILLER, W. E., HARZHAUSER, M.,MANDIC, O., BERNING, B., ROGEL, F. ET AL. 2007.The Oligo-/Miocene Qom Formation (Iran): evidencefor an early Burdigalian restriction of the TethyanSeaway and closure of its Iranian gateways. Inter-national Journal of Earth Sciences, doi 10.1007/s00531-007-0269-9.

RICHTER, D., MARIOLAKOS, I. & RISCH, H. 1978. Themain flysch stages of the Hellenides. In: CLOSS, H.,ROEDER, D. & SCHMIDT, K. (eds) Alps, Apennines,Hellenides. Inter-Union Commission on GeodynamicsScientific Report, 38, 434–438.

ROBERTSON, A. H. F. 2000. Mesozoic-Tertiary tectonic-sedimentary evolution of a south Tethyan oceanicbasin and its margins in southern Turkey, In:BOZKURT, E., WINCHESTER, J. A. & PIPER,J. D. A. (eds) Tectonics and magmatism in Turkeyand the Surrounding Area. Geological Society,London, Special Publication, 173, xx–xx. Q24

ROBERTSON, A. H. F., UNLUGENC, U. C., INAN, N. &TASLI, K. 2004. The Misis-Andirin Complex: a Mid-tertiary melange related to late-stage subduction ofthe Southern Neotethys in S Turkey. Journal of AsianEarth Science, 22, 413–453.

ROGL, F. 1998. Paleogeographic considerations forMediterranean and Paratethys Seaways (Oligocene to


1335133613371338133913401341134213431344134513461347134813491350135113521353135413551356135713581359136013611362136313641365136613671368136913701371137213731374137513761377137813791380138113821383138413851386138713881389139013911392

Miocene). Annalen des Naturhistorischen Museums inWien, 99, 279–310.

ROGL, F. 1999. Circum-Mediterranean Miocene Paleo-geography. In: ROSSNER, G. & HEISSIG, K. (eds)The Miocene Land Mammals of Europe. Dr. FritzPfeil Verlag, Munchen, 39–48.

RUGGIERI, G. & SPROVIERI, R. 1970. I microforaminiferidelle ‘marne di S. Cipirello’ Translated Title: Forami-nifera of the San Cipirello marls. Palermo University,Institute of Geology; (Lavori.10;1970).Q18

SANCAY, R. H., BATI, Z., ISIK, U., KIRICI, S. & AKCA,N. 2006. Palynomorph, Foraminifera, and CalcareousNannoplankton Biostratigraphy of Oligo-MioceneSediments in the Mus Basin, Eastern Anatolia,Turkey. Turkish Journal of Earth Sciences, 15,259–319.

SAROGLU, F. & YILMAZ, Y. 1986. Geological evolutionand basin models during neotectonic episode in theEastern Anatolia. Bulletin Mineral Research andExploration Institute of Turkey, 107, 61–83.Q18

SCHNEIDER, B. & SCHMITTNER, A. 2006. Simulatingthe impact of the Panamanian seaway closure onocean circulation, marine productivity and nutrientcycling. Earth and Planetary Science Letters, 246,367–380.

SENEL, M. 2002. Geological Map of Turkey, Ankara,MTA.

SENGOR, A. M. C. & YILMAZ, Y. 1981. Tethyan evol-ution of Turkey: a plate tectonic approach. Tectono-physics, 75, 181–241.

SENGOR, A. M. C., GORUR, N. & SAROGLU, F. 1985.Strike-slip faulting and related basin formation inzones of tectonic escape: Turkey as a case study. In:CHRISTIE-BLICK, N. (ed.) Basin Formation andSedimentation. Society of Economic Paleontologistsand Mineralogists Special Publications, 37, 227–264.

SENGOR, A. M. C., OZEREN, S., GENC, T. & ZOR, E.2003. East Anatolian high plateau as a mantle-support, North–South shortened domal structure.Geophysical Research Letters, 30, doi: 10.1029/2003GL017858.

SENGOR, A. M. C., TUYSUZ, O., IMREN, C., SAKINC, M.,EYIDOGAN, H., GORUR, N., LE PICHON, X. &RANGIN, C. 2005. The North Anatolian Fault: a newlook. Annual Reviews in Earth and Planetary Sciences,33, 37–112.

SLACK-SMITH, S. M. 1998. Ostreoidea. In: BEESLEY,P. L., ROSS, G. J. B. & WELLS, A. (eds) Mollusca:The southern synthesis. Fauna of Australia, CSIROPublishing, Melbourne, 5, Part A, CSIRO Publishing,Melbourne, 268–274.

SPEZZAFERRI, S. 1994. Planktonic foraminiferal biostrati-graphy of the Oligocene and lower Miocene in theoceanic record. An overview. PaleontographiaItalica, 81, 1–187.

SPEZZAFERRI, S. & PREMOLI SILVA, I. 1991. Oligoceneplanktonic foraminiferal biostratigraphy and paleocli-matic interpretation from Hole 538A, DSDP Leg 77,Gulf of Mexico. Paleogeography, Paleoclimatology,Paleoecology, 83, 217–263.

TURCO, E., HILGEN, F. J., LOURENS, L. J., SHACKLE-

TON, N. J. & ZACHARIASSE, W. J. 2001, Punctuatedevolution of global climate cooling during the late

Middle to early Late Miocene: High-resolutionplanktonic foraminiferal, oxygen isotope recordsfrom the Mediterranean. Paleoceanography, 16,405–423.

TUYSUZ, N. & ERLER, A. 1995. Geology and geotectonicimplications of Kazikkaya area, Kagizman-Kars(Turkey). In: ORCEN, S. (ed.) Geology of the BlackSea Region. Proceedings of the Internationalsymposium on the Geology of the Black sea Region,7–11 September 1995, Ankara, Turkey. Generaldirectorate of mineral research and exploration andchamber of geological engineering, Ankara, 76–81.

VAN DER ZWAAN, G. J. & GUDJONSSON, L. 1986. MiddleMiocene–Pliocene stable isotope stratigraphy andpaleoceanography of the Mediterranean. MarineMicropaleontology, 10, 71–90.

VAN DER ZWAAN, G. J., JORISSEN, F. J. & STIGTER,H. C. 1990. The depth dependency of planktonic/benthic foraminiferal ratios: Constrains and appli-cations. Marine Geology, 95, 1–16.

VAN HINSBERGEN, D. J. J., HAFKENSCHEID, E.,SPAKMAN, W., MEULENKAMP, J. E. & WORTEL,M. J. R. 2005a. Nappe stacking resulting from subduc-tion of oceanic and continental lithosphere belowGreece. Geology, 33, Q22325–328.

VAN HINSBERGEN, D. J. J., KOUWENHOVEN, T. J. & VAN

DER ZWAAN, G. J. 2005b. Paleobathymetry in thebackstripping procedure: Correction for oxygenationeffects on depth estimates. Paleogeography, Paleocli-matology, Paleoecology, 221, 245–265.

VAN HINSBERGEN, D. J. J., LANGEREIS, C. G. &MEULENKAMP, J. E. 2005c. Revision of the timing,magnitude and distribution of Neogene rotations inthe western Aegean region. Tectonophysics, 369,1–34.

VAN HINSBERGEN, D. J. J., ZACHARIASSE, W. J.,WORTEL, M. J. R. & MEULENKAMP, J. E. 2005d.Underthrusting and exhumation: A comparisonbetween the External Hellenides and the ‘hot’ Cycladicand ‘cold’ South Aegean core complexes (Greece).Tectonics, 24, doi: 10.1029/2004TC001692.

WADE, B. S., BERGGREN, W. A. & OLSSON, R. K. 2007.The biostratigraphy and paleobiology of Oligoceneplanktonic foraminifera from the equatorial PacificOcean (ODP Site 1218). Marine Micropaleontology,62, 167–179.

WESTAWAY, R. 2003. Kinematics of the Middle East andeastern Mediterranean updated. Turkish Journal ofEarth Sciences, 12, 5–46.

WESTAWAY, R. 2004. Kinematic consistency between theDead Sea Fault Zone and the Neogene and Quaternaryleft-lateral faulting in SE Turkey. Tectonophysics, 391,203–237.

WOODRUFF, F. & SAVIN, S. M. 1989. Miocene deepwateroceanography. Paleoceanography, 4, 87–140.

WOODRUFF, F. & SAVIN, S. M. 1991. Mid-Mioceneisotope stratigraphy in the deep sea: high resolutioncorrelations, paleoclimatic cycles, and sediment pres-ervation. Paleoceanography, 6, 755–806.

WRIGHT, J. D., MILLER, K. G. & FAIRBANKS, R. G.1992. Early and Middle Miocene stable isotopes:implications for deepwater circulation and climate.Paleoceanography, 7, 357–389.


1393139413951396139713981399140014011402140314041405140614071408140914101411141214131414141514161417141814191420142114221423142414251426142714281429143014311432143314341435143614371438143914401441144214431444144514461447144814491450

YILMAZ, Y. 1993. New evidences and model on the evol-ution of the southeast Anatolian orogen. GeologicalSociety of America Bulletin, 105, 251–271.

YILMAZ, Y., GUNER, Y. & SAROGLU, F. 1998. Geologyof the quaternary volcanic centers of east Anatolia.Journal of volcanology and geothermal research, 85,173–210.Q23

ZACHARIASSE, W. J. 1992. Neogene planktonic foramini-fers from Sites 761 and 762 off NorthwesternAustralia. Proceedings of the Ocean DrillingProgram, Scientific Results, 122, 665–675.

ZACHOS, J., PAGANI, M., SLOAN, L., THOMAS, E. &BILLUPS, K. 2001. Trends, rhythms, and aberrations inglobal climate 65 Ma to present. Science, 292, 686–693.


Oligocene Miocene basin evolution in SE Anatolia, Turkey: …forth/publications/Huesing_2009a.pdf · 2009-01-09 · Oligocene Miocene basin evolution in SE Anatolia, Turkey: constraints

Documents