A revised stratigraphic framework for later Cenozoic sequences in the northeastern Mediterranean...

Geol Rundsch (1995) 84:794-812 © Springer-Verlag 1995

O R I G I N A L P A P E R

C. Yeti~ • G. Kelling • S. L. GOkgen • F. Baroz

A revised stratigraphic framework for Later Cenozoic sequences in the northeastern Mediterranean region

Received: 7 March 1995 / Accepted: 15 July 1995

Abstract This study describes the lithostratigraphic character of mid-Cenozoic (Oligocene-Pliocene) sequences in different parts of the northeastern Mediter- ranean area and offers a detailed stratigraphic correlation for this region. The sequences concerned are drawn from the Camardi area (south-central Anatolia), the Adana Basin, the Misis Mountains and the Kyrenia Range (northern Cyprus) and the submerged Florence Rise (west of Cyprus). The stratigraphic relationships identified here indicate the following: (a) Following the middle Eocene (Lutetian) regression there was uplift throughout the entire region; (b) Episodes of fluvial and lacustrine deposition in intramontane settings en- sued in most of this region during the late Eocene/early Miocene interval; (c) Following a regionally extensive phase of tectonic compression, major marine transgression commenced in the late Oligocene in northern Cy- prus and in the early Miocene in adjacent southern Turkey, with the exception of the Ecemi~ Fault Zone where continental deposition continued; (d) These O1- igo-Miocene transgressive sequences comprise a broadly diachronous complex of both shallow and deeper marine facies, including reefal carbonates, littoral clastics, basinal shales and fan-turbidites; (e) Deeper marine Miocene facies persisted longer in the Misis area and in northern Cyprus; (f) A regional regression oc-

Prof. Dr. S. L, GOkgen died in 1992.

Cengiz Yeti~ Mut Meslek Ytiksekokulu, Mersin Universitesi, Mut, I~el, Turkey (ori. Jeoloji Mtthendisligi B/31ttmtt, ~ukurova Universitesi, Adana, Turkey)

Gilbert Kelling (1~) Department of Earth Sciences, University of Keele, Staffs., UK

Sungu L. G6k~en Deniz Bilimleri ve Teknolojisi Enst., Dokuz Eyl01 Universitesi, Izmir, Turkey

Francois Baroz Geological Institute, University of Nancy, Nancy, France E-mail: [email protected].

curred throughout most of the area during the late Ser- ravallian to Tortonian interval and is marked by the abrupt, locally discordant appearance of extensive shallow marine, deltaic and fluvial deposits; (g) Continued regression in the Messinian led to the formation of significant evaporite deposits in the western and southern parts of the region, but localized uplift of the Misis area is attested by the initial deformation of the Neogene rocks there and the absence of Messinian sediments from this area; (h) In the Pliocene there was extensive emergence of the northern parts of the region inter- rupted by brief marine incursions. The present-day drainage pattern was established at this time; (i) Ma- rine conditions persisted longer in northern Cyprus, where emergence occurred only in the latest Pliocene.

Introduction

The aim of this study was to describe the lithostratigraphic character of the Mid-Cenozoic (Oligo- cene-Pliocene) sequences of the northeastern Mediter- ranean (in southern Turkey and northern Cyprus), to summarize the biostrat)graphic and other evidence for their correlation and to outline the palaeoenvironmen- tal evolution of this region in the mid-Cenozoic interval. The regional stratigraphic/sedimentological pattern elicited here strongly suggests that the Miocene depositional basin of southern Turkey was elongate, had an original east-west trend, and narrowed eastwards. The studied sequences are derived from: the ~amardi (Ni~de) area of the Ecemi~ Fault Zone, the Adana Ba- sin, the Misis Mountains, the Kyrenia Range of northern Cyprus and the DSDP sites on the Florence Rise, west of Cyprus (Fig. 1).

Southern Turkey is a critical region in evaluating the geological evolution of the Levant and/or eastern Me- diterranean, but as yet the genesis of the Tertiary Medi- terranean basins is poorly understood and there is no consensus on a model for their origin. In most current global geodynamic models the Adana Basin is located

795

TURKEY "e~'" '~;~. ~N~ -~. . b ~

• .. : . ~ . . : . . .~ - . .~. .:~y-x.. . . ~ ~,

~ N ~ Plio-Quaternary

" - q ~ ~ ~ E~ Pre-Miocene

' . , . .' . : . . " ' , . , ; • . • Ankara

~ : ~ . . . - /. . . . . [ % _ _ ~ . . . . ~ ; MEDITERRANEAN SEA I Med~nea~ I:

Fig. 1 General location map for the northeastern Mediterranean region showing the general geological setting and the main localities described herein. DSF=Dead Sea Fault; EAF =East Anatol- ian Fault; EF = Ecemi~ Fault zone; FR = Florence Rise; KB = Kar- santi Basin; KM=Kharamanmara~; KR=Kyrenia (Pentadakty- los) Range; TM=Troodos Massif

above the zone of Cenozoic suturing between the Afro- Arabian and the Euro-Asian plates (Dewey and ~eng6r 1979; Seng6r and Yilmaz 1981; Robertson and Dixon 1985; Gealey 1988). The collision between the Arabian and the Anatolian plates, which was thought to be of Late Cenozoic (Kelling et al. 1987) or Middle Miocene age (~eng6r 1979; ~engOr and Yilmaz 1981; Dewey et al. 1986; G0rtir 1985), has recently been suggested to be Eocene-Miocene in age (Hempton 1985; Yilmaz 1992). This would explain why the Adana Basin records extensional subsidence - possibly post-collisional in origin - as early as the Burdigalian (Williams et al., 1995). Al- ternatively, the Adana Basin may have opened in the Early Miocene as an extensional fore-arc basin north of the Cyprus Trench subduction zone (Jackson and McKenzie 1984). What is clear is that continental collision and suturing result in a strong tectonic control on the sedimentation occurring in and around the suture zone (Miall 1984), with effects that should be manifest in the facies, compositions and thickness variations in the associated basin-fill sequences.

The Camardi (Ni~de) area is situated adjacent to the Ecemi~ Fault Zone, which defines the boundary between the western and eastern Taurus ranges (Fig. 2). The western and eastern parts (blocks) of the ~amardi area present slightly different stratigraphic characters, as does the Ecemi~ Fault Zone itself. Throughout this area the Oligocene-Miocene succession is dominated by continental deposits (Yeti~ 1978, 1984a, 1984b, 1987b; Yeti~ and Demirkol 1984; Fig. 3).

In southern Turkey the most conspicuous of the collision-related troughs are the Antalya and ~ukurova basins (G6k~en et al. 1985, 1986; Kelling et al. 1987; Yeti~ 1988a). The latter now comprises two subsidiary troughs (the Adana and Iskenderun basins) separated

by the Misis complex (Figs. 4 and 5). The Tertiary (O1- igocene-Pliocene) succession of the Adana Basin has been assigned to 12 lithostratigraphic units, which include deposits of pre-transgressive, transgressive and regressive character (Yeti~ 1988a, Fig. 4). On the other hand, the stratigraphic succession within the Misis complex commences with an olistostromic unit (late Oligo- cene to Aquitanian) and the succeeding Miocene units comprise a thick turbiditic formation (Burdigalian-early Tortonian) and regressive, mainly deltaic deposits of late Miocene age (Fig. 6). The deformed sequence in the Misis area is overlain by continental Pliocene-Qua- ternary deposits and Quaternary basalts (Schmidt 1961; Bilgin and Elibol 1984; G6k~en, S. et al. 1985, 1986; Kelling et al. 1987; GOkqen et al. 1991).

The general lithological succession and detailed biostratigraphical analyses of the Cenozoic successions of northern Cyprus, and their stratigraphic correlation with the sequences encountered at DSDP Sites 375-376, have been presented by Baroz and Bizon (1974), Baroz (1979), and Baroz et al. (1978), whereas Robertson and Woodcock (1986) have detailed the Me- sozoic and Cenozoic history of the Kyrenia Range. Fol- lowing late Eocene uplift of the Pentadaktylos (Kyre- nia) Range the Oligo-Miocene sequences to the north and south of this lineament developed in somewhat different settings and display significant differences in stratigraphic expression. In the northern (Kythrea) basin the Oligocene commences with mass-flow conglomerates which are overlain by Oligo-Miocene siliciclastic and calcareous turbidites seen in the Kyrenia Range (Weiler 1969; Baroz et al. 1978, Robertson and Wood- cock 1986). The late Miocene regressive deposits of northern Cyprus are represented by an upwards-shoaling sequence of clastics and Messinian evaporites, followed discordantly by shallow marine to coastal Plio- cene deposits (see Fig. 7).

Stratigraphical details

~amardi (Ni~de) area

The ~amardi Cenozoic basin of Ni~de Province, southern Turkey, (first described by Blumenthal 1941, 1952) straddles the Ecemi~ Fault Zone, an important lineament that separates the western and eastern Taurides. Most of the (mainly sinistral) strike-slip displacement along this fault occurred prior to the Lutetian (Yeti~ 1978, 1984a, b; Yeti~ and Demirkol 1984), resulting in significant differences in the facies and stratigraphic successions seen in the pre-Lutetian rocks on the western and eastern flanks of the fault zone (Fig. 3). The western limit of the ~amardi basin is represented by the Ni~de metamorphics, whereas the basin is delimited to the east by the eastern Tauride nappes.

Three geological sectors can be recognized in the ~amardi region: a western block (west of the Ecemi~ Fault Zone); an eastern block (east of the fault); and an

796

Fig. 2 Simplified geological map of the ~amardi area (Ni~de province, southern Turkey). Note that in this area the 'basement' sequence to the east of the Ecemi~ Fault consists of Tauride nappes, comprising Permo-Triassic massive carbonates (PTrm- Maden Formation); upper Triassic-Jurassic limestones (TrJd-Demirkazik Formation); and the Mazmili ophiolite (Kin-abducted in late Creta- ceous). After Yeti~ 1978

xplanation

~ Alluvium Alluvial Fans & Screes Terrace Deposits

(~a'mlca Conglomerate

liocene ~--~Bur~: Formation lligocene ~-'~'~£ukurba§ Formation &

I ° '~° lK6rpinar Gypsum member iocene ~t~a.'~ goleboynu Formation

~ Karadag Spilite ,k Movras Umestone member l~le~ene ~umardi Formation

Cretaceous [1]~17

I ~^^ l Netumorphlc Basement

q'~, ~ayseri

Ulukisla Kamisli "

,-q/ BLACK SEA X

Ni~de ~ 1t f II f...~ j ~ . . ] km IEDITERRANEAN ...... "1~

elongate basin, confined to the fault zone (Figs. 2 and 3).

Western ~amardi block

The oldest rocks seen here are the Ni~de metamorphics. The depositional age of the Ni~de metamorphics is probably early Late Cretaceous according to foraminiferal evidence obtained from the less metamor- phosed, but lithostratigraphically comparable, sequences of the Kutahya-Bolkardag belt (G6n~tio~lu 1986; Gonqtio~lu et al. 1992, 1993). These metamorphics are disconformably succeeded by the ~amardi formation (Middle-Upper Palaeocene clastic turbidites: 350 m) interbedded with, and overlain in turn by, the Karadag spilite suite (450 m of Upper Palaeocene basaltic flows and sills interleaved with deep marine mudstones, marls and lenticular limestones which form the Mavras Member). On Evliya hill, northwest of Ca- mardi, these Palaeocene rocks are discordantly followed upwards by the basal conglomerates, thick num-

mulitic limestones and shallow marine clastics of the Kaleboynu formation (Lutetian: 200+ m), which also locally oversteps onto the metamorphic basement (Ye- ti~ 1978, Yeti~ and Demirkol 1984; Figs. 2 and 3).

The later Cenozoic sediments of the ~amardi area are represented by the ~ukurba~ and Burg formations (Oligocene-Miocene), best developed within the Ecemi~ Fault Zone (see below). At several localities in the western and eastern blocks these formations rest unconformably on all the older rock units (Yeti~ 1978).

Eastern ~amardi block

In the eastern Camardi block the basement comprises the Tauride thrust slices, which include the Maden formation (massive carbonates of Permian-Early Triassic age), the Demirkazik formation (massive platform carbonates of Late Triassic-Campanian age) and the Maz- mili ophiolitic melange, abducted in the Campan-

Fig 3 Generalized stratigraphic correlation within the Ca- mardi area. (Modified from Yeti~ 1978)

WESTERN BLOCK

ECEMIS FAULT-ZONE

EASTERN BLOCK

797

/ /

/

LEGEND ~ lluvial Gravel --~ Lignite ~ Gypsum ~] Shale/Siltstone ~ Sandstone .~o Pebbly Sandstone o~o Conglomerate -_~ Calc. Siltstone [~] Lirnestone ~ olomite ;~ Serpentinite ~ Spilite

J J

ian-Maastrichtian time interval (Fig. 3). The basement units are overstepped discordantly by the Kaleboynu formation, already mentioned in the sequence of the western block, which marks a phase of rapid marine transgression throughout the region during the Lute- tian. The maximum thickness observed for the Kale- boynu formation in the eastern block is 136 m com- pared with more than 400 m to the west of the Ecemi~ Fault. This difference is attributed to differential era-

sion prior to deposition of the unconformably overlying ~ukurba~ Formation (Oligocene).

The Ecemi¢ Fault Zone

In the ~amardi region this zone, which has an average width of about 3 km, falling between the two main branches of the NNE-SSW trending Ecemi~ Fault, dis-

798

Fig. 4 Generalized stratigraphic successions illustrating the correlation and lithological character of Cenozoic units in the northern part of the Ada- na Basin. (The approximate N-S trending Gildirli-Karais- ali-Kuzgun line is shown in Fig. 5). Modified from Yeti~ 1988a

W

QUATERNARY

PLtOCENE

L )

- - >_ Ld

o 01 Z N Ld

O -- k - O

Z

o Ld

Ld O

OLIOOCENE

PALEOZOIC-MESOZOIC

~ Terroce-Caliche

~ O~kkuyu Gypsum Member Handere Formation

Kuzgun Formation ~ Memisli Member

Solbas Tuff Member Kuzgun Member

Alluvium

. . . . . . . . . - o : T h : . . T U : . = . i T : r ? . : ~ . . . . . . -T.:o] :::..T: z,.._ . . . . . . ,~.-Tkum-:- : . J _~Tkurnz.:-:-:-:_: . . . . . . .

~,,v :..v '..". w. Tkus ".'v.'.', v ..:w. ",'v'.': v :' Tkus~:=:-:Tkur-:- 'Tkuk . . . . ~ .STkuk.

~ ' , ~ o " . ' " " e ~ ' * o : ~ ; 4o ~ " ~ ' d W ~ ; " . " ~ ' ~ "

3d

II II , - . . . . . <;_7' . ' II II A t - - ' - . . . . . f o u - - - ~ S - - - ~ - ~ = ' - =

T k u , ~ - ~ . - - ~ o ' ' o : ° , o :

" . . . . = ,= .~Xkp : : : ~ ~,:~: . , - .~TJk, R-:.~:7:-~. .... k3 . . . . . t - - . z . ~ . . . . . . : - - .~ . . , .~ ~ i , . . . . . . . . . g l . . . . . .

~ GUvenc Formation

~ Cing6z Formation

~ Karaisali Formation

MESSINIAN

TORTONIAN

SERRAVALLIAN

LANGHIAN

BURDIGALiAN

OLIGOCENE 11 Km

PAL./MES.

• Kaplonkoya Formation

~ Gildirti Formation

~ Karsanti Formation

~ Paleozoic-Mesozo]c Basement

plays a thick succession of post-Lutetian sediments. An intramontane basin, recording widespread retreat of the sea, was initiated here in the late Eocene (Yeti~ 1978, 1984a, 1984b, 1987a, 1987b).

Within this zone there are minor outcrops of the late Cretaceous Mazmili ophiolite, unconformably overlain by the ~ukurba~ Formation. This formation (up to 270 m thick) consists of a thick sequence of redbeds, mainly coarse fluvial clastics capped by finer sediments, including both gypsum and thin coal layers. The ~ukurba~ Formation is unfossiliferous but is probably of O1- igocene age both because it occurs between demonstra- bly Eocene and Miocene formations and because of its strong similarity in lithofacies and succession to biostratigraphically dated Oligocene formations in neighbour- ing areas (e.g. the Karsanti basin - ~afak and OnltJgenq 1992; Onltigenq et al. 1992; Yeti~ 1987b). Facies analysis of the conglomerates, coarse-grained sandstones and thin mudstones forming the lower part of the ~ukurba~ Formation demonstrates the existence of both braided and meandering river deposits, whereas the fine- grained sandstones, white marls, and thin gypsum and coal seams in the upper levels of the formation are attributed to playa and lacustrine environments (Yeti~ 1987b; Kerey and Yeti~ 1991).

The overlying Burq Formation, some 270 m thick in the fault zone, consists of an upwards-fining sequence of fine-grained sandstones and red marls with a few thin lignite layers and gypsum bearing mudstones, all indicative of an upwards transition from lacustrine and swamp conditions into lake-delta and meander belt deposits (Yeti~ 1987b; Kerey and Yeti~ 1991). Sparse os-

tracod and gastropod faunas suggest that this formation is Miocene in age (Yeti~ 1978; determinations by N. GOk~en).

Throughout the ~amardi area the youngest deposits are represented by thick sequences of Quaternary alluvial gravels and coarse-grained sands (Figs. 2 and 3).

The Adana Basin

The Adana Basin is situated in the easternmost sector of the Mediterranean coast of Turkey and is bounded by the Ecemi~ Fault to the west, the Taurus Orogenic Belt to the north and the Misis structural high to the east. Southwards it extends to Cyprus (as the Cilicia Basin) beneath the waters of the Mediterranean (Fig. 5). A broader definition of the Adana Basin, adopted here, includes the Tertiary sequence found in the isolated Karsanti Basin, situated within the Taurides immediately north of the main Adana Basin (Fig. 1).

Geological investigations around the Adana Basin began with the studies by Kirk (1935), Maxon (1936) and Foley (1937). Further geological work was under- taken by Blumenthal (1941), and Ternek (1957). Schmidt (1961) was the first to propose a comprehen- sive lithostratigraphic scheme for the Adana Basin and his units have been adopted by many subsequent workers with only minor changes (Ozer et al. 1974; Ilker 1975; G6rttr 1979, 1980, 1982, 1985, 1992; Nazik and Toker 1986; Yalqin 1982; Yal~in and GOrtir et al. 1984; Gtirbtiz et al. 1985; Gtirbtiz and G0kqen 1985; Tanar

Fig. 5 Simplified geological map showing the distribution- of Cenozoic units within the northern Adana Basin, with insets showing the geographical location of the area described and its relationship to adjacent Cenozoic basins. (From Yeti~ 1988a)

. • Black .-U-X, Sea L.. • Ankara ~M "L.

R K E Y 471 a[!.a ........ :,

Mediterranean

r " .7,;,9 ,u°,---b 0 50 1 O0

' km ~ ' [-'---~ Quaternary

Tertiary

Pa{eozeic- Mesozoic

0 ~liocene

°- f O , < Z ~ Miocene

w m o ' "

Oligocene

Mesozeic-Paleezeic~

Terrace - caliche

Handere Formation G6kkuyu Member

Kuzgun Formation Memis]i Member Salbas Member Kuzgun Member G/~venc Formation

Cing6z Formation

Karaisali Formation Kaplankaya Formation

Gildirli Formation

Karsanti Formation

Basement

799

1985). The later Tertiary stratigraphy of the Adana Ba- sin was recently reorganized by Yeti~ (1988a).

The basement of the Tertiary succession of the Ada- na Basin is formed by Paleozoic and Mesozoic units. The Palaeozoic is represented mainly by Middle-Upper Devonian coralline limestones and clastics, and limestones-dolostones of Permo-Carboniferous age. By contrast, the Mesozoic succession comprises platform carbonates (Upper Triassic-Cretaceous) and turbidite sandstones and shales of Late Cretaceous age. Ophiol- itic melange was emplaced in this region after the Late Cretaceous.

The Cenozoic succession is mainly represented by Tertiary units on the northern side of the basin and

Quaternary deposits to the south. The Tertiary rocks unconformably overlie deformed and thrust-emplaced Palaeozoic and Mesozoic rock units, and an upwards succession of pre-transgressive, transgressive and regressive deposits can be distinguished. Recently, the Tertiary succession (Oligocene-Pliocene) of the Adana Basin has been assigned to 12 lithostratigraphic units comprising eight formations and four members (Yeti~ 1987a, 1988a; Fig. 5).

Pre-transgressive deposits

The pre-transgressive deposits of the Adana Basin (sensu lato) include the alluvial-lacustrine-lagoonal

800

Karsanti formation and the alluvial-coastal Gildirli formation, effectively comprising superimposed non-marine depositional sequences that are retrogradationally transgressive onto basement rocks.

The Karsanti formation. The Karsanti formation consists mainly of marls and mudstones, but with substan- tial thickness of redbed conglomerates and pebbly sandstones forming the lowest part of the sequence. Bedding thickness and maximum clast size of the pebbly layers both decrease upwards. Thin lignitic coals, plant-rich shales and gypsum layers occur in the argillaceous upper part of the formation. Small slumps, cross bedding and current ripples are common. The Karsanti formation reaches a maximum thickness of 1500 m in the vicinity of Karsanti town (Schmidt 1961). The Kar- santi formation rests unconformably upon Late Creta- ceous melange and, south of Karsanti town, an angular unconformity can be observed between the Karsanti rocks and the Miocene Kaplankaya and Karaisali formations.

Previous biostratigraphic indications suggested a Miocene age for the Karsanti formation (Abaci 1986; Yurtmen 1986), a date which is anomalous in view of the structural relationships described above. However, recent work by f)nl~gen~ et al. (1993) has clarified the situation. These workers have established the presence of a thin sequence of sandstones and shales, which in- tervenes between the basal Karsanti alluvial member and the upper lacustrine-fluvial units. This sequence has yielded a typical microfauna of planktonic foraminifera and brackish to lagoonal ostracods of Oligocene age including Globigerina tripartitaKoch, Globigerina ciperoensisangulisuturalisBolli, Globigerina ciperoensis augustiumbilicataBolli, Globigerina ouachitaensis oua- chitaensisHowe and Wallace, together with Hemicypri- deis montosaKeen, and Echinocythereissp. Thus, the Karsanti formation represents the fill of an intramontane basin formed in the Oligocene as a precursor half- graben to the main Adana Basin.

The Gildirli formation. Deposits of alluvial fan character crop out around Gildirli village and along the northern flank of the Adana Basin (sensu stricto). They comprise red and brown conglomerates, sandstones and shales arranged in a series of 2- to 20-m-thick fining- upwards cycles (Yeti~ 1988a; GOrUr 1992). Each cycle generally has an erosional base, overlain by conglomerates which pass upwards to cross-bedded pebbly sandstones and sandstones. The sandy layers are followed by alternations of fine-grained sandstone and siltstone. The top of each complete cycle is marked by 20- to 30- cm-thick layers of reddish mudstones which are poorly consolidated and parallel-laminated. Some cycles include calcrete units, generally of nodular type, displaying both small spherical concretions and short, narrow columns arranged perpendicular to the bedding plane. The basement of the Tertiary succession of the main Adana Basin has an irregular topography, the depres-

sions in which are filled by conglomeratic units of the Gildirli formation. This formation attains a maximum thickness of 400 m in the pre-Miocene troughs, but the thickness decreases greatly on the intervening palaeo- highs.

At many localities in the northern and western parts of the Adana Basin the Gildirli alluvial fan redbeds can be observed to rest unconformably upon Palaeozoic and Mesozoic units of the Taurides (Fig. 5). Recently, 15nlt~genq (1993) has described and logged a river section near ~okak village, on the northwestern margin of the Adana Basin, where Gildirli conglomerates rest discordantly on tilted fluvio-lacustrine deposits of the Kar- santi formation. Moreover, the Gildirli formation passes upwards and laterally into the shallow marine Kaplankaya formation (L. Miocene) in the palaeotopographic depressions. By contrast, over the topographic highs the Gildirli formation is concordantly overlain by the reefal Karaisali formation.

Thus, although to date these redbeds have not yielded biostratigraphically useful fossils, the age of the Gildirli formation probably ranges from late Oligocene to basal Miocene (Aquitanian(?)-early Burdigalian).

Transgressive deposits. The transgressive sequence of the Miocene in the Adana Basin comprises the coastal to shallow marine clastics of the Kaplankaya formation, the reefal carbonates of the Karaisali formation, the foraminiferal deep marine shales of the Guvenc formation and the turbidite-fan sandstones of the Cingoz formation (Figs. 4 and 5).

The Kaplankaya formation. This unit was first distinguished by Yeti~ and Demirkol (1986) and is essentially equivalent to the Kabalaktepe member of the Gildirli formation as defined by G6rar (1979; 1992). It consists mainly of fossiliferous sandstones (sometimes pebbly), siltstones, marls and sandy limestones. Alternating sandstones and siltstones dominate the lower part of the formation and the proportion of carbonate in- creases upwards. The marls and clayey/sandy limestones contain abundant remains of echinoids, bivalves, gastropods, small benthic foraminifera, algae, corals, etc. The Kaplankaya formation discordantly overlies Palaeozoic and Mesozoic rock units on the paleotopo- graphic highs (where the Gildirli formation is missing). However, locally there are demonstrable vertical and lateral transitions between the basal Kaplankaya and upper Gildirli units. The upper part of the Kaplankaya formation passes laterally and vertically into both the reefal Karaisali and the basinal Guvenc formations. The thickness is highly variable (from 5 to 100 m) and mainly related to the underlying paleotopography.

The upper levels of the Kaplankaya in the Kutu- koy-Goksun and Kale sections contain a rich foraminiferal association of Burdigalian age (Benda et al. 1977), whereas other localities yield small benthic foraminiferal assemblages that indicate an age range for the Ka- plankaya of Burdigalian to early Langhian (Yeti~

801

1988a). Thus, the coastal to shallow marine Kaplankaya formation appears to represent somewhat diachronous initial deposits resulting from early Miocene marine transgression into the Adana Basin.

The Karaisali formation. This formation is composed mainly of reefal and biostromal carbonate sediments. These are generally white to pale grey, medium to thick bedded and contain algae, corals, foraminifera, echi- noderms, mollusca, bryozoa and worm tubes. Ergene (1972) noted algal, clastic, coralline and micritic sub-facies. GOrt~r (1979, 1980) described coralgal packstones and boundstones, small benthic foraminiferal-algal packstone, coralgal wackestone and packstone, large benthic foraminiferal algal packstone, globigerinid-algal packstone and globigerinid argillaceous wackestone sub-facies for the Karaisali formation. The thickness of the Karaisali formation is variable in the Adana Basin (from 20 to 350 m) and it is thicker to the west of the Gildirli-Karaisali-Kuzgun line (Fig. 5). The age range for the Karaisali, determined from small benthic foraminifera and algae, is Burdigalian to early Serravallian (Yeti~ 1988a), consistent with the observed vertical and lateral passage of basal Karaisali units into upper Ka- plankaya facies and the transitions between upper Ka- raisali beds and the Guvenc and Cingoz clastics.

Thus, the Karaisali formation represents a major reef complex, formed in warm, clear, agitated waters during the acme of early Miocene marine transgression. The first-formed reefs appear to have been confined to palaeotopographic highs, but the Karaisali reef complex later onlapped far to the north across the denuded Tauride units.

The Cingoz formation. Schmidt (196t) designated three members within the Cingoz formation (Ayva, Topalli and Kopekli). Yeti~ (1988a) assigned the first two of these members to his re-defined Cingoz formation (the practice also followed in this paper). The Kopekli shale member was transferred into the Guveng formation by Yeti~ (1988a). The Cingoz formation is developed mainly to the east of the Gildirli-Karaisali-Kuzgun line (Fig. 5). This formation generally starts with poorly bedded, often channelized conglomerates, pebbly sandstones and sandstones which overlie carbonate-rich and carbonate-poor shales that are assigned respectively to the Kaplankaya and Guven~ formations. These shales display large-scale slump-scar features (Gtirbtiz 1993) and several feeder-channel complexes (cut into Kaplan- kaya neritics and filled by Cingoz conglomerates and sandstones) have also been identified on the northern margin of the Adana Basin (Gtirbtiz and Kelling 1992). This basal Cingoz sequence passes upwards into a thick succession of turbidite sandstone-shale alternations. Towards the top of the Cingoz succession shales be- come dominant. The coarse to very fine-grained sandstone layers of the Cingoz display typical turbidite characteristics including a variety of erosional sole marks. The intervening shale layers are fissile, calcareous and,

occasionally, carbonaceous. As defined here, the Cin- goz formation ranges in thickness from 1000 m in the west to 3200 m in the east (Gtirbtiz 1993).

The Cingoz formation is partly coeval with, and partly superimposed upon, the Kaplankaya, Karaisali and Guven~ formations. A recent re-investigation of foraminiferal assemblages indicated that the Cingoz deep marine shales were deposited during the late Burdigal- ian-Serravallian interval (Nazik and Gtirbaz 1992). De- tailed sedimentological investigation has revealed that the Cingoz formation comprises the deposits of two small, coeval submarine fans formed in the northern part of the Adana Basin during the middle Miocene (Gt~rbaz and Kelling 1992, 1993). Each submarine fan starts with channelized bodies of inner fan character, succeeded by thickening upward cycles representative of mid fan lobes and capped by lower fan/basin plain sediments. Overall, these submarine fans are retrograd- ational in style and correspond to turbidite systems types I and II of Mutti (1985).

The Guvenf formation. The Guven~ shale and the Ko- pekli shale members of the Cingoz formation of Schmidt (1961) have been combined recently into the Guven~ formation (Yeti~ 1988a). These predominantly deep marine, foram-rich shales are partly coeval with the Cingoz formation fan turbidites and partly overlie the fan-sequences. The Guven~ formation starts with slope and fore-reef deposits, which pass up into basinal shales. In accordance with the general shallowing of the basin, these basinal shales are transitionally succeeded by storm-dominated outer shelf deposits (Gtirbtiz 1993). Sedimentary and faunal features (the decreasing number and thickness of sandy layers, the decrease in benthic foraminifera and increasing amounts of planktonic foraminifera and Discoaster-type nannofossils) indicate increasing bathymetry from the bottom to the middle section of the Guven~. Prominent pyritization suggests poorly oxygenated environmental conditions in the middle to upper parts of the formation.

On the other hand, an increasing amount of sandy detritus, together with benthic foraminifera and Braa- rudosphaera discula, Braarudosphaera bigelowi-type nannofossils indicate relatively shallow marine conditions in the uppermost part of the succession (Oz~elik 1993; C)z~elik and Yeti~ 1993). The maximum thickness of the Guven~ formation is around 2100 m along the Karaisali-Guven~-Kuzgun line, but the thickness of the unit is significantly reduced to the west. The basal part of the Guven~ formation is partly coeval with, and partly superimposed upon, the Kaplankaya, Karaisali and Cingoz formations, and is disconformably overlain by the terrestrial to shallow marine clastics of the Kuzgun formation. Abundant planktonic foraminifera and nannofossils demonstrate that the Guven~ formation was deposited during the Burdigalian-Serravallian interval (Ozelik 1993).

802

Regressive deposits

The shallow marine and terrestrial clastics of the Kuz- gun formation (Kuzgun, Salba~ Tuffite and Memi~li Members) and the succeeding Handere formation (with the G6kkuyu Gypsum member) form the major regressive cycle of the Adana Basin (Yeti~ 1988a; Figs. 4 and 5).

The Kuzgun formation. The term Kuzgun formation was introduced by Schmidt (1961) for the thick sequence of Late Miocene coarse red and yellow sandstones, conglomerates and shales, with subordinate fossiliferous green sandstones and silty shales and a Wide- spread tuffite layer. Yeti~ and Demirkol (1986) formal- ly sub-divided this formation into three members as described below.

The Kuzgun member. The Kuzgtin member consists mainly of coarse redbeds and shallow marine sediments. The distribution of the main facies within the Kuzgun member has been described by Kerey et al. (1985) and by Yeti~ and Demirkol (1986), who pointed out the regional increase in the proportion of carbonate sediments towards the western part of the Adana Ba- sin. One of the main facies of the Kuzgun member is represented by fluvial cycles of high sinuosity (meandering) type. Each cycle (1-35 m thick) starts with strongly erosional surfaces infilled by trough cross-bedded conglomerates, pebbly sandstones and coarse- grained sandstones. This basal level is succeeded by sandy low-angle cross-bedded point bar deposits, occasionally followed by silty levee deposits.

Reddish-brown mudstones intercalated with the sand bodies represent the floodplain sub-facies. Some of these mudstones contain layers with abundant calcite nodules, probably representing calcrete horizons. Sev- eral horizons (0.5-10 m thick) of greenish-grey sandstones, siltstones and mudstones rich in Ostrea shells represent intercalations of lagoonal or shallow marine deposits within the Kuzgun member. Bedding surfaces of the sandstone layers commonly display burrows, wave- and current ripples. Mudstones and siltstones within these "Ostrea marker units" contain bivalves, gastropods and plant debris.

The Kuzgun member rests erosively upon various levels of the Guven~ formation within the central part of the Adana Basin, but onlaps onto older units (including the Tauride basement rocks) on the western margin, north of Tarsus. In the field the Guven~ and Kuzgun formations appear to be broadly concordant in structural dip, but seismic profiles reveal clear regional discordance at the base of the Kuzgun formations (Wil- liams et al., 1995). The Kuzgun member is transitionally succeeded by the Salba~ Tuffite Member in most parts of the Adana Basin. The maximum thickness of the Kuzgun member is approximately 434 m, and the abundant fauna of benthic foraminifera, bi-valves, gastropods and Hipparionsp. teeth indicates deposition during the late Serravallian to Tortonian (Yeti~ 1988a).

The Salba~ Tuffite member. The Salba~ Tuffite member corresponds to the Upper Tuff member of Schmidt (1961; see Yeti~ 1988a). The Kuzgun member passes upwards into the Salba~ member through a transitional unit comprising light-grey volcaniclastic sandstones, grey siltstones and pebbly sandstones, the sub-rounded pebbles of which are composed mainly of pale-coloured tuffite. This transition zone is covered by a 10-m-thick bed of light-grey/white tuffite which contains variable amounts of clay and silt. On a regional scale this member displays broadly lenticular geometry, the thickness varying from 1 to 15 m (thicker in the west). Vertebrate teeth (Hipparionsp.) and gastropods found in the Sal- ba~ siltstones indicate a Tortonian age for this member (Yeti~ 1988a).

The Memi~li member. The Memi~li member is a sandstone-dominated unit that consists of fluvial, deltaic, lagoonal and shallow marine deposits, lithologically re- sembling the Kuzgun member. The lower part of the member is marked by channelized conglomerates, succeeded by a stacked sequence of fining-upwards fluvial channel-fill cycles (0.5-25m thick), dominated by trough cross-bedded, poorly sorted pebbly sandstones and coarse-grained sandstones, all displaying lenticular geometry. Only a few cycles retain the capping floodplain deposits (fine- to medium-grained, low-angle trough cross-bedded sandstones, reddish siltstones and mudstones). Higher parts of the member display a thickening upwards sequence of siltstones and fine- grained sandstones exhibiting marked bioturbation and parallel or cross lamination.

Shallow marine and lagoonal deposits, with intercalated storm layers, are found mainly in the upper levels of the Memi~li member. They comprise green-grey, sparsely pebbly sandstones, siltstones and mudstones, and yield locally abundant Ostrea and other bivalves, together with gastropods, ostracodes and rare planktonic foraminifera. These faunas indicate a Torton- ian-Messinian(?) age range for this member (Tanar 1985; Gfirb~iz and G6k~en 1985; Inal A., cited in Yeti~ 1988a).

The maximum thickness of the Memi~li member is around 443 m and it passes transitionally upwards into the Handere formation. Throughout most of the Adana Basin the Memi~li rests upon the Salba~ Tuffite. How- ever, locally the basal Memi~li conglomerates are in di- rect contact with upper units of the Kuzgun member, presumably as a result of penecontemporaneous ero- sion of the intervening tuffites.

The Handere formation. Schmidt (1961) gave the name Handere formation to a thick sequence of yellow and beige sandstones, conglomerates, subordinate siltstones and some rnappable gypsiferous layers, of late Miocene to Pliocene age. The name has been successively re- tained, although a recent revision has assigned the gypsiferous sequences to the Gokkuyu Gypsum member (Yeti~ and Demirkol 1986; Yeti~ 1988a). Marked lateral

803

variations in facies are evident, with appreciably coars- er sediments in the western part of the Adana Basin and more thin marine units in the east.

Much of this formation comprises channelized bodies of trough cross-bedded coarse-grained sandstones and conglomerates, formed in moderate-sinuosity fluvial systems. Pale grey, soft sandstones, parallel-laminated siltstones and "flaser-like" sand/silt/mud interbeds represent shallow marine and lagoonal intercalations up to 25 m thick and yield gastropods, ostracods and rare planktonic foraminifera. White and beige, thin-bedded calcareous siltstones, claystones and soft, carbonaceous sandstones form units in the upper part of the Handere succession that are attributed to lagoonal and lacustrine environments. The maximum thickness of this formation is 450 m and the faunal evidence indicates a Messinian-early Pliocene age range (Gtirbfiz and G0k~en 1985; Yeti~ 1988a). Both the base and top of the Handere formation are transitional in character.

The GOkkuyu Gypsum member. This unit, confined to the southwestern part of the Adana Basin, comprises layers of soft grey sandstones and siltstones containing several stringers and lenses (0.25-5 m thick) of gypsum. The gypsum is coarsely to finely crystalline with sac- charine and porphyroblastic textures. Penecontempora- neous re-working is attested by local interbeds of peb- ble- and cobble-grade rudites composed of gypsum clasts. This unit occurs near the top of the Handere formation (Yeti~ 1988a, b) and is generally underlain by the lagoonal/lacustrine siltstones described above.

The maximum thickness of the G6kkuyu Gypsum member observed at outcrop is more than 40 m (Onlti- gent 1993). However, boreholes in the southern part of the onshore Basin reveal 100-900 m of broadly coeval evaporites (gypsum, anhydrite and halite; Yal~in and G6rtir 1984) and even greater thicknesses of Messinian evaporites occur below the seafloor of the Cilicia Basin (Mulder 1973; Mulder et al. 1975; Evans et al. 1978), as elsewhere in the Mediterranean (Hsu 1972; Hsu et al. 1973).

The Misis complex

The southwest-northeast trending Misis Mountain is separated by major faults from the adjacent gently folded sequences of the Adana and Iskenderun basins (Fig. 6). Up to 4300 m thick, the Cenozoic rocks of the Misis Mountains display more complex and intensely deformed structures, dominated by steep northwesterly tectonic dips, local tight lolding and a broadly imbricate arrangement of thrust slices, transected by strike-slip faults (Kelling et al. 1987). The sedimentary succession exposed in the Misis area is of late Oligocene and Mio- cene age, overlain by less deformed Pliocene-Quater- nary sediments and basaltic lavas (Ternek 1957; Eren- toz and Ternek 1962; Ozer et al. 1974; Bilgin and Elibol 1984; G6k~en et al. 1985, 1986, 1988; Kelling et al. 1987; G0kqen et al. 1991). The basinal Tertiary sequences record deposition within a complex trough that under- went rapid subsidence during the Aquitanian-Serraval- ian time interval, followed by gradual shoaling which culminated in shallow water and terrestrial deposition throughout most of this belt in Tortonian-early Messin- ian time. At the base of the succession two main lithostratigraphic units have been recognized, the Isali and Karatas formations (Schmidt 1961; Schiettacatte 1971). The olistostromic origin of the Isali formation was suggested by Schiettecatte (1971), and regional correlation of the Misis sequence with adjacent areas of southern Turkey and Cyprus was proposed by Ùzer et al. (1974) and Baroz et al. (1978). The third unit is the Kizildere formation, present only in the southeastern part of this region. Stratigraphic evidence indicates that deformation of the Misis sequence started in the early- to middle Miocene and continued during the early Pliocene and Quaternary. Detailed micropalaeontological investigations revealed tectonic repetition of stratigraphical- ly distinct successions within the Miocene (G6kgen et al. 1985, 1986, 1988; Kelling et al. 1987; G6k~en et al. 1991).

The Isali formation

Late Pliocene-Quaternary formations. Late Plio- cene-Quaternary formations are represented mainly by the alluvial terrace and caliche deposits associated with the present-day drainage system of the Adana Basin. From bottom to top these comprise Pliocene clays, caliche palaeosols, colluvial materials, and massive and lenticular caliches (Kapur et al. 1986; Yeti~ 1988a; Ka- pur et al. 1993). Rare pebbly sandstones and cross-bedded conglomerates are poorly sorted and grain-supported. A widespread development of caliche is observed above the younger terraces. This unit is up to 30 m thick and southwards it is covered by the recent alluvium of the Adana plain (Seyhan delta).

The Isali formation tectonically overlies the Karatas clastics and represents an olistostromic complex containing large and small blocks of Mesozoic limestones, as well as various mafic and ultramafic igneous rocks, within a matrix of poorly sorted volcanogenic sandstones and shales that have yielded late Oligocene to basal Miocene (Aquitanian) microfaunas, notably Glo- bigerinoidesprimordiusBlow and Banner, Catapsydrax- dissimilisCushman and Bermudaz (GOk~en et al. 1986; GOkqen et al. 1991). Although strongly tectonized, the Isali formation is estimated to be at least 2000 m thick. Volcanogenic material within the Isali formation olis- tostrome includes both basaltic lavas and volcaniclastic sediments. The latter are predominantly mass flow deposits and form most of the olistostromic matrix. On a regional scale the Isali-Karatas contact appears to be a

804

relatively flat-lying thrust in the southwest, but more steeply inclined in the northeast. Genesis of the Isali formation of the Misis complex started no later than the early Miocene or late Oligocene (Bilgin and Elibol 1984; Yalcin and G6rar 1984), with olistostromic accumulation of Mesozoic platform limestones, and ob- ducted ultramaflc and volcanic rocks of back-arc origin, which also provided sandy detritus to form the turbiditic matrix of the olistostromes.

The Karatas formation

This is dominated by deep marine facies, comprising turbidite sandstones alternating with pelagic mudstones and marls, but it also includes rare olistoliths of Maas- trichtian and Palaeocene shelf limestone and volcanic rocks. However, thin sequences of biostromal and co- quinoidal carbonates, marls, cross-stratified sandstones and conglomerates also occur and represent shallow marine and slope settings. The thickness of the Karatas formation varies from 1500 to 3000 m and the age (derived from planktonic foraminiferal assemblages) ranges from Burdigalian to early Tortonian (Kelling et al. 1987; G0k~en et al. 1991).

The Karatas formation deep marine clastics include deposits formed in basin plain, small submarine fan lobe, and marginal slope settings. Palaeocurrent vectors derived from both deep and shallow marine Karatas facies fall into two groups, indicating transport down a southeast-facing slope and southwestwards, along the basin axis. The Karatas clastics include both siliciclastic and carbonate detrital components, and they were derived from source rocks similar to those contributing to the Isali sediments. The junction between the Karatas and the succeeding Kizildere formations is demonstra- bly a thrust or reverse fault in the Yumurtalik sector (Ozer et al. 1974). On the other hand, the Karatas rocks are separated from the fluvial gravels of the Pliocene- age Kuransa formation by a steep reverse fault.

The Kizildere formation

This unit is present only in the southeastern part of the Misis area, and comprises pale grey to pink cross-bedded, coarse, pebbly sandstones, thin siltstones and coa- ly lenses of fluvio-deltaic origin. A few layers of soft dark volcanogenic sandstones and siltstones rich in Os- trea, ostracods and gastropods represent brief marine

Fig. 6 Simplified geological map of the Misis Mountains area with insets showing the geographical location and the relationship to adjacent Ceno- zoic basins. (From G0k~en et al. 1985)

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incursions and this typical fauna is assigned to the Tor- tonian in this region (Schmidt 1961; GOk~en et al. 1991). This Kizildere sequence (up to 500 m of which is preserved) indicates further shoaling of the Adana Ba- sin with development of deltaic and paralic successions characterized by an abundant supply of medium to coarse siliciclastic detritus from the northeast, and subordinate shallow marine re-working and transport in southerly and southwesterly directions.

There is no evidence in the Misis area of deposition during the Messinian. However, a fluvio-lacustrine sequence of reddish, cross-bedded sandstones and conglomerates with buff to white, laminated siltstones (the Kadirli formation) disconformably rests on Misis complex rocks near Osmaniye (Fig. 6) and a late Mio- cene-Pliocene age has been assigned to these deposits (Schmidt 1961). The topmost units in the Misis area are alluvial terrace and caliche deposits (Hamis formation: late Pliocene-Quaternary) and the Delihalli basalt (late Quaternary).

Northern Cyprus

The general lithological succession of the Tertiary (O1- igocene-Pliocene) formations of northern Cyprus and of DSDP Sites 375 and 376 (Florence Rise) has been described and discussed by Baroz and Bizon (1974), Baroz (1979), Baroz et al. (1978). Further sedimentological observations in the Kythrea group are found in the publications of Weiler (1964, 1965, 1969, 1970) and in Robertson and Woodcock (1986). Two different depositional areas are recognized in the Tertiary sediments of Cyprus: 1. To the north and south of the late Eocene Pentadac-

tylos range, the turbiditic and hemipelagic sediments of the Kythrea group (Oligocene-late Miocene) were deposited.

2. In southern Cyprus, around the Troodos Massif, sedimentation of pelagic chalks (which succeed Late Cretaceous ophiolites) continued through the Oligo- cene (Lefkara formation). This was followed in the Miocene by tectonically controlled deposition of the Pakhna formation, comprising a variety of carbonate-dominated facies, from reef bodies to mass-flow rudites (platform to shallow basin). Only the northern area is considered here. In the Kythrea basin the Oligocene is represented by

the Kythrea conglomerate, Bellapais formation and Klepini siltstone from bottom to top (Fig. 7). The transgressive deposits of northern Cyprus continue with the Miocene Flamoudi sandstones, the Panagra marls and the Trapeza marls, whereas the regressive cycle commences with the Davlos sandstone formation followed by the late Miocene evaporitic Lapatza formation. The Pliocene in northern Cyprus is represented by the Me- saoria group, starting with the early Pliocene Myrtou marls and the Nicosia formation, whereas the Athalassa formation is of Mid- to Late Pliocene age. The Quater-

nary deposits of northern Cyprus are exclusively continental. The Miocene and early Pliocene sediments re- covered at Sites 375-376 are broadly comparable with those of the Kythrea group. By contrast, the later Plio- cene and Quaternary lithofacies of Cyprus are funda- mentally different from the coeval pelagic and hemipelagic deposits at site 376.

Kythrea conglomerate

The Kythrea conglomerate, at the base of the Kythrea group, consists mainly of channelized, clast-supported conglomerates and coarse sandstones, often forming fining-upwards cycles a few metres thick. These appear to be of alluvial coastal plain and littoral origin (Ro- bertson and Woodcock 1986) and attain a maximum thickness of about 40 m. This formation rests unconformably on elements of the Pentadactylos (Kyrenia) range complex, a narrow but long-lived structural lineament that suffered intense southwards thrusting in the late Eocene, accompanied by strong uplift to the north (Baroz and Bizon 1974; Robertson and Woodcock 1986).

These coarse clastics have yielded little biostratigraphic evidence of their age, but they rest unconformably on rocks as young as late Eocene (the Priabonian Kalograia-Ardana olistostromes; Baroz 1979) and pass upwards into well-dated marine deposits of early Oligo- cene age (the Bellapais formation). The Kythrea basal conglomerate sequence is therefore most plausibly assigned an Early Oligocene age (Robertson and Wood- cock 1986).

Bellapais formation

The main part of the Bellapais formation (up to 400 m thick) is a sequence of coarse- to medium-grained lithic turbidite sandstones with thin mudstone interbeds which abruptly, but concordantly, succeeds the alluvial-littoral rudites of the Kythrea conglomerate, marking very rapid subsidence of the Kythrea basin. The interbedded mudstones have yielded planktonic foraminifera of Early Oligocene (Stampian) age (Baroz and Bi- zon 1974).

Klepini formation

This concordantly overlying late Oligocene-Aquitanian siltstone-dominated unit consists of alternations of fine- grained distal volcanilithic turbidites (derived from the east), with hemipelagic and pelagic organic-rich marls and planktonic foraminifera-rich limestones. The formation attains a maximum thickness of 160 m and passes transitionally up into the Flamoudi formation turbidites.

806

Fig. 7 Stratigraphic successions from northern Cyprus and DSDP sites on the Flor- ence Rise (west of Cyprus) il- lustrat ing the correlation and lithological character of Ceno- zoic units. Inset: map of Cy- prus shows the locations of the successions depicted. (Aft- er Baroz 1979)

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Flamoudi formation Panagra formation

Traced eastwards along the Pentadaktylos range this unit, comprising about 200 m of bioclastic turbidites and hemipelagic marls (Aquitanian-Burdigalian) passes laterally into a sequence containing a higher proportion of siliciclastic turbidites. Excellent Bouma sequences and abundant sole-marks are characteristic of the Flamoudi turbidites. The abundance of carbonate detritus in these early Miocene clastics has been attributed to a renewed pulse of marine transgression, isolating terrigenous sources and creating wide carbonate-producing marginal shelves (Robertson and Wood- cock 1986; Robertson et al. 1991).

This marl-rich unit (up to 50 m thick) represents a good lithological marker within the Kythrea group of northern Cyprus (Moore 1960). Very similar lithofacies can be traced within this unit from northern Cyprus to the Florence Rise (site 375).

This formation (Langhian in age) is dominated by decimetre-scale sharp-based beds of greenish grey to blue-green, fine-grained limestone grading up into a reddish brown marl with foraminifera-rich laminae. These are interpreted as distal carbonate turbidites and are composed of re-deposited terrigenous and hemipelagic biogenic detritus. Frequent interbeds of light grey to dark red, 1 0 2 0 cm thick, lithic arenite turbidites also occur in the Panagra, together with some thin layers of acid tuffite (Robertson and Woodcock 1986).

807

Trapeza formation

In northern Cyprus the Panagra marls pass transitionally upwards into a sequence of dark marls with some lithic arenites and siltstones that span the Serravallian to Tortonian interval. The thickness of the Trapeza on Cyprus has been estimated as 500 m by Baroz and Bi- zon (1974), whereas the coeval deposits at site 375 are approximately 75 m thick. The typical Trapeza lithofacies comprises centimetre- to decimetre-scale layers grading upwards from dark siltstone (rich in plant debris), through pale marl and pale foram/nannofossil- rich marl and into pelagic limestone (site 375). Occa- sional well-graded, 10- to 20-cm-thick, fine-grained, pale buff lithic arenites also occur in the upper part of the Trapeza.

The regressive sequence of northern Cyprus spans the Tortonian to late Pliocene interval and comprises the Davlos sandstone, the evaporitic Lapatza formation, the Myrtou marls, the Nicosia and Athalassa formations (Fig. 7).

Davlos formation

The Davlos sandstones are restricted to the eastern part of northern Cyprus, on the north and south flanks of the Kyrenia Range. They consist of a 250-m-thick sequence of turbiditic coarse brown to grey, lithic sandstones, siltstones and marlstones which laterally grade westwards into the Trapeza marls lithofacies. These Tortonian sandstones are interpreted as a relatively proximal turbidite complex, derived from the east, and the coeval unit at site 375 (Florence Rise) records the acme of clastic input at this location (Robertson and Woodcock 1986). Some marl-rich intervals in the Day- los formation include beds (up to 3 m thick) of white to pale-yellow acid tuff and hyaloclastites. Similar tuffites have been described by Weiler (1965) from the Kythrea flysch at site 375.

chalks which gradually pass downwards into the Trape- za marls.

Myrtou formation

This 300-m-thick unit of early Pliocene age consists mainly of marls and siltstones with a rich benthic microfauna and overlies the evaporites of the Late Mio- cene Lapatza formation with an angular unconformity marked by basal conglomerates containing gypsum pebbles. The Myrtou marls appear to have been deposited in well-oxygenated shallow marine conditions. A similar unit is observed at site 376 (Florence Rise).

Nicosia and Athalassa formations

In northern Cyprus the middle Pliocene is represented by the Nicosia formation (below) and Athalassa formation (above).

Bioclastic limestones, sandstones and marls of the Nicosia formation (20 m), of shallow marine origin, are overlain by approximately 50 m of caliche-bearing sandy marls, sandstones and conglomerates of the Athalas- sa formation, marking further shoaling and eventual emergence into coastal plains and alluvial fans.

Quaternary

On Cyprus the Quaternary is essentially represented by terrestrial facies with periodic shallow marine intercalations. From the early Quaternary onwards Cyprus was emergent and subjected to pulsed tectonic events. The marine terraces (up to 500 m above present sea level) thus record successive stages of tectonic uplift.

Stratigraphic correlation and discussion

Pre- Miocene framework

Lapatza formation

This unit (up to 150 m thick at outcrop and assigned a Messinian age) consists of marls, pelagic chalks and organic-rich mudstones in its lower part, passing up into thick evaporites. At outcrop the Lapatza evaporites are essentially composed of gypsum and generally occur as chaotic masses, but beds of laminated, saccharoidal and selenitic gypsum are also noted. Site 376 on the Flor- ence Rise penetrated only the uppermost, attenuated part of the Mediterranean evaporite sequence, consist- ing of alternating anhydrite and rock salt layers, overlain by 50 m of gypsum. This is overlain in turn by 92 m of greenish grey, partly dolomitic marl stones with interbedded graded sandstones and siltstones. The main evaporites are underlain by a sequence of marls and

During the middle Eocene a major marine transgression occurred extensively in southern Turkey and adjacent areas, and Lutetian marine deposits are thus encountered throughout the region considered here, from Gamardi to Cyprus (Yeti~ 1978; Baroz et al. 1978). However, late Eocene deformation, manifest throughout the Tauride realm (and most plausibly attributed to final closure of an Inner Tauride Ocean: GOrt~r et al. 1984) induced widespread thrusting and strong uplift. Consequently, there was almost complete retreat of the Neotethyan sea from Asia Minor (Luttig and Steffens 1976; G6k~en et al. 1986; Yeti~ 1988a), resulting in a stratigraphic/structural break of variable magnitude. During the Oligocene a series of intramontane basins developed in southern Anatolia, in which thick bodies of alluvial, lacustrine and marginal marine sediments

808

N ~-

1 C#mardi TURKEY ~ '2 ¢b"

~ / 'y -~

Nd£~)o 7 ~ ilOOkmi

CAMARDI - NIGDE AREA I 20 km

1

MIOCENE ~ A - - 2 ~ ' / ~ Burc Fm.

OLIGOCENE .-~ - ~ ' . I Cukurba~ Fro.

~ Kaleboynu Fro. ~ EOCENE

I

90 km ADANA BASIN

15 km I W 3 E

~ Caliche Terrace

PLIOCENE G ~ Gypsum Mem.

MESSINIAN ~: r:..~.:,- Handere Fm. - -.'~-'2;.'E'.

- - ".L-~-; T.'. Memisli Member : : : : : : ; : : : :

_ ~ TORTONIAN -- ~ - - _ ~ Sa f ba s Tuff M e m b e r . . . . . . . . . -,. '~: -:. ' .~' .

~ Kuzgun Member : 2222 ;22 Z ~2

SERRA~AWAN :.T.:.C~.: ~.:;

~ afivenc Fm, --

LANGHIAN - . . ,

~ Karafsali Fm. !11~ ii ii Kaplankaya Fm. --

B U R D I G A ~ A N , ~ , ~ i :

Gildirli Fm.

OUGQCENE.L

4-- . . . . . . 60 km . . . . . ~ . . . . 120 km . . . .

: : Transgression

. . . . Regression

Caiiche Terrace MISIS COMPLEX

Handere Fm. 5

x x x x x

Kuzgun Fm.

TORTONIAN

Gt3venc Fro. SERRAVALUAN

Cing6z Fro. ~ ~ " " -

. . . . . . . . . . Karatas . . .

~ LANGHIAN

- . - i - . -1 Fro. Karaisali Fro. ' ' ' I

Kaplankaya Fm. . . . . "~ "'~' BURDIGALIAN-

~ ' ~ ~ Isali Fm, -- Gildirli Fm. ~ AQUITAN IAN

ThnJst

Karsantf Fro.

2 5 0 m

Caliche QUATERNARY Basalt

Hamis Fm. PLIOCENE Kadirli Fro.

Kizildere Fro.

-4 S

NORTHERN CYPRUS ~---- 90 km ----~ 6 7

Kyrenia

r:~r:!r:]7

=-I'~l'~t=

N - ,T IT , .--~

!:i.~:i.~-7:--~-- .~?..72.. ~,, - o o o o c

Florence Rise Nicosia Fm, ~ t PLIOCENE P[

"[a'~tza Fro. Ur _ _ MESSIN IAN -- ~: ~ Unit 5

MESSIN IAN Unit 6

T O R T O N I A N "~ -,-,

• -~ "--. Unit 7

Trapeza Maris SERRAVALLIAN

- Unit 8,9

LANGHIAN Unit 10 Pana(lra Maris -- Unit 11

_BU._RDIGALIAN F lamoud i Sandstone AQUITANIAN

t-':i~~-":[:~ Kle'pini Siltstone m:._-.:..-. . : :- T.. : CHA'FFIAN

:':C?':'-i@~ <3 Bel lapa is Fro.

STAMPIAN

Kythrea Conglomerate

Fig. 8 Summary diagram of proposed correlations between the later Cenozoic (Oligocene-Piiocene) successions in the northeastern Mediterranean region illustrating the diachronous nature of the mid-Cenozoic marine transgression in southern Turkey and the near-synchronous character of the Late Miocene regression. Inset shows the geographic location of sequences illustrated. Li- thological symbols are explained in Figs. 3, 5 and 7. (Note: different thickness scales used for sequences in the different basins studied)

were deposited. These include the redbeds at the base of relevant sequences in the Kutukoy-Goksun and Kale-Denizli sections (Benda et al. 1977) and in the ~amardi-Ni~de area considered here (Yeti~ 1978, 1987b), which thus rest unconformably on Lutetian (and older) rocks. The unfossiliferous redbeds of the ~ukurba~ formation in the ~amardi area and the lower, alluvial/coastal deposits of the fossil-bearing Kar- santi formation in the Adana Basin were deposited during the middle Oligocene and pass conformably up into lacustrine deposits of late Oligocene to basal Mio- cene age both in the ~amardi area (Burg formation) and in the upper part of the Karsanti formation of the northern Adana Basin. The alluvial-littoral character of the early Oligocene Kythrea conglomerate in northern Cyprus provides evidence of the same geotectonic episode which here resulted in only brief, local uplift of the Pentadactylos lineament, followed by rapid subsidence in a marine setting that led to accumulation of the Bellapais-Klepini turbidites and hemipelagites.

A further, minor tectonic pulse in the early Miocene (Aquitanian?) is detectable in the northern Adana Ba- sin (post-Karsanti formation and pre-Gildirli) and also to the west in the Mut Basin ($afak et al., in prepara-

tion). The same episode may be responsible for olistostromic deposition in the Misis area (Isali formation) and probably induced the floods of coarse detritus into palaeotopographic hollows that formed the alluvial fan and fan delta deposits of the Gildirli formation (Fig. 8).

Transgressive units

No evidence of Miocene open marine deposition is now found in the ~amardi area, and it seems likely that the early Miocene sea never reached this region. Instead, this area continued to be dominated by terrestrial deposition, although raising of the local base level is suggested by the lake-delta and meander-belt sediments of the Burg formation. Moreover, this marine transgression reached the northern Adana Basin in the Burdi- galian, but invaded the Misis area prior to the Aquitan- ian and left its mark in the deep marine environments of northern Cyprus also in the late Oligocene.

In the Adana Basin the transgressive units com- mence with the shallow marine and littoral clastics of the Kaplankaya formation and the overlying reefal carbonates of the Karaisali formation (Burdigal- ian-Langhian). In the Misis area the faulted base and deep marine nature of the Isali formation (Aquitanian) precludes more accurate dating of the arrival of the mid-Tertiary sea in the southern Adana Basin. In the predominantly deep marine sequences of northern Cy- prus this transgressive phase is weakly reflected in the abrupt appearance of abundant carbonate detritus within the basal Miocene Flamoudi formation turbid-

ites (Fig. 8). Most of this detritus was derived from con- temporaneous sediments formed on the wide, rapidly inundated, carbonate-generating shelf-areas to the north and northeast of northern Cyprus and which are exemplified by the limestones of the Kaplankaya and Karaisali formations in the Adana Basin and by similar early Miocene deposits observed in the Mut Basin (~a- fak and G6kqen 1991) and in the Manavgat/Kopru basins near Antalya (Flecker et al., in press).

In the Adana Basin this early Miocene transgression is succeeded by rapid and widespread subsidence, lead- ing to the accumulation of several kilometres of coarse turbidites and basinal shales (Cingoz and Guvenc formations). Similar subsidence is reflected in the Karatas turbidites of the Misis area and in the deep basin of northern Cyprus and the Florence Rise by the marked lithological change (within the early Langhian) from the turbiditic calcarenites of the Flamoudi to the hemipelagites and distal turbidites of the Panagra and Tra- peza formations (Fig. 8).

The widespread effects of this early Miocene subsidence event are also evident in the Antalya region, where shelf carbonates are again succeeded by thick turbiditic sequences of basinal aspect (Hayward 1984; Flecker et al., in press).

Thus, the early Miocene transgression and rapid subsidence appear to be regionally extensive events, in- fluencing much of the northeastern Mediterranean and probably tectonic in origin. However, the precise geotectonic cause remains problematic. In the Adana Ba- sin and Misis complex these events have been attributed to extension in response to load-induced flexure triggered by Tauride thrusting (Williams and Onlt~genq 1992; Williams et al., 1995; G6k~en et al. 1988). Howev- er, Flecker et al. (in press), citing the evidence of Eaton and Robertson (1993) for the Oligo-Miocene initiation of subduction south of Cyprus, attribute the early Mio- cene regional subsidence to crustal extension generated by back-arc "roll-back" on the Hellenic-Cyprus subduction system, an interpretation also favoured by Postma et al. (1993).

Regressive units

In the Adana Basin a subtle, but regionally important, discordance separates the basinal Cingoz-Guvenc sequence from the regressive deposits represented by the shallow marine, deltaic and fluvial Upper Serraval- lian-Tortonian Kuzgun formation (Fig. 8). The Torton- ian Kizildere formation of the Misis area and the Up- per Tortonian Davlos sandstones of northern Cyprus are lithostratigraphically correlative with the Kuzgun formation of the Adana Basin. However, shallow marine intercalations are more common in the Misis area, whereas the northern Cyprus sequence (Davlos) remains relatively deeper. Widespread explosive volcan- ism of unknown source during the Tortonian is mani- fested by the Salba~ tuffites of the Adana Basin, and by

809

thin volcaniclastic horizons in the lower Kizildere formation of the Misis area, which are laterally equivalent to prominent tuffites encountered in the Davlos formation in northern Cyprus (Figs. 7 and 8). The coarse coastal plain and shallow marine deposits of Messinian to Pliocene age in the Adana Basin (Handere formation) may be correlated with the Lapatza formation of northern Cyprus, whereas the Messinian evaporites of northern Cyprus (and below the Cilicia Basin) are the much-expanded lateral equivalents of the G6kkuyu Gypsum member in the Adana Basin. However, Mes- sinian evaporites have not been detected in the Misis area, which probably was emergent this time. The Myr- tou marls, Athalassa and Nicosia formations of northern Cyprus attest to further shoaling and persistent shallow marine conditions during the Pliocene, which are evident both in northern Cyprus and at sites 375-376 on the Florence Rise.

The dramatic change in lithofacies and palaeoenvi- ronmental character during the Tortonian seen in the Adana Basin is accompanied by significant changes in palaeoflow patterns and sandstone provenance (Yeti~ et al. 1986; Unliigen~ et al. 1992; Unliigenq 1993) marking a major palaeogeographical and geotectonic re-or- ganization. This is most plausibly attributed to incep- tion of the Kharamanmara~ triple-junction (intersec- tion of the East Anatolian Fault and Dead Sea Fault) and the associated uplift of the fully sutured Bitlis sector of southeast Turkey (Fig. 1). These events effectively created the present-day morphological elements of the northeastern Mediterranean.

Deposition during the Quaternary in northern Cy- prus and throughout the late Pliocene and Quaternary in the Cukurova Basin was persistently terrestrial in character. In accordance with the new geomorphologi- cal scenario outlined above these continental lithofacies in the Adana Basin, the Misis area and northern Cy- prus contrast strongly with the pelagic and hemipelagic Quaternary deposits cored at the Florence Rise DSDP sites.

Acknowledgements We wish to pay tribute to our deceased co- worker, Prof. Dr. Sungu G6k~en, who stimulated and contributed significantly to the early stages of this study. We are also indebted to his wife, Prof. Dr. Nuran G0k~en, who undertook many of the microfossil determinations. We express our appreciation to Prof. Dr. T. Ozsayar and Prof. Dr. S. Kapur who critically read the manuscript and Dr. K. Gtirbtiz who assisted in editing the paper. G. Kelling is grateful for financial assistance provided by N.A.T.O. Science Division (grant no. 83/871) and the British Council. Logistic and technical support from the M. T. A. Insti- tute of Turkey (Mediterranean Regional Directorate), ~ukurova University (Adana-Turkey) and the University of Keele (Eng- land) is also warmly acknowledged. We also express our appreciation for technical and secretarial services provided by various members of the support staff in the geology departments of the Universities of Mersin, ~ukurova, Keele and Nancy. We are most grateful for the constructive comments of D. Sciumach and an- other anonymous reviewer.

810

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A revised stratigraphic framework for later Cenozoic sequences in the northeastern Mediterranean...

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