Tectonic implications of transtensional supradetachment basin development in an extension-parallel transfer zone: the Kocacay Basin, western Anatolia,Turkey

Tectonic implications of transtensionalsupradetachment basin development in anextension-parallel transfer zone: the Kocac� ay Basin,western Anatolia,TurkeyHasan Sozbilir, Bilal Sarı, Bora Uzel, Okmen Sumer and Serkan Akkiraz

Dokuz Eylˇl �niversitesi, Mˇhendislik Fakultesi, JeolojiMˇhendislig� i B˛lˇm , 35100 Bornova-izmir,Turkey

ABSTRACT

TheKocac� ay Basin (KC� B) is a key area in western Anatolia ^ a well-known extended terrane whereregional segmentation has received limited attention ^ for investigating strike-slip faultskinematically linked to detachment faults. In this paper, we present results of an integratedsedimentologic, stratigraphic, and structural study ofMiocene alluvial fan/fan-delta/lacustrinedeposits that accumulated in the KC� B, a NE-trending basin with connections to theMenderesMetamorphic Core Complex (MCC).We mapped and evaluated most of the key faults in the KC� B,many for the ¢rst time, and recognised di¡erent deformation events in the study area near the Emargin of theMCC.We also present ¢eld evidence for kinematic connections between low-anglenormal and strike-slip faults whichwere developed in an intermittently active basement-involvedtransfer zone in western Anatolia.We ¢nd that the KC� B contains a detailed record ofMiocenetranstensional sedimentation and volcanism that accompanied exhumation of theMCC. Structuraldata reveal that the basin was initially formed by transtension (D1phase) and subsequently upliftedand deformed, probably as a result of early Pliocene wrench- to extension-dominated deformation(D2 phase) overprinted by Plio-Quaternary extensional tectonics (D3 phase).These results areconsistent with progressive deformationwherein the axis of maximum extension remained in thehorizontal plane but the intermediate andmaximum shortening axes switched position in the verticalplane. Combining our results with published studies, we propose a newworking hypothesis that theKC� B was a transtensional supradetachment basin during theMiocene.The hypothesis could providenew insights into intermittently active extension-parallel zone ofweakness inwestern Anatolia.Theseresults also suggest that the termination of low-angle normal fault systems within an extensionparallel transfer zone may have resulted in a transtensional depressions which are di¡erent fromclassical supradetachment basins with respect to the sedimentation and deformational pattern of thebasin in¢lls.

INTRODUCTION

One of the best known extended terranes in the world iswestern Anatolia, where regional segmentation hasreceived limited attention because most workers have fo-cused on low-angle normal faults and related meta-morphic core complexes (MCC) and supradetachmentbasins (Hetzel et al., 1995; Emre, 1996; Emre & S˛zbilir,1997; Koc� yig� it et al., 1999; Yılmaz et al., 2000; S˛zbilir,2001, 2002; Seyitog� lu et al., 2002; Bozkurt & S˛zbilir,2004).However,many basins formed in transtensional set-tings involve an association of strike-slip and low-anglenormal faults that are responsible for the exhumation of

MCCs (e.g. Christie-Blick & Biddle, 1985; Sylvester, 1988;Burch¢el et al., 1995; Friedmann & Burbank, 1995; Dorsey&Barajas,1999;Murphy etal., 2002). In these systems, nor-mal faults have developed orthogonal to the extensionaldirection, whereas strike-slip dominated transfer zonesthat caused the segmentation of extended terranes are or-iented parallel or slightly oblique to the extension direc-tion (Aydın & Nur, 1982; Gibbs, 1984; Bosworth, 1985;Faulds & Varga, 1998; Fodor, 2007; Umhoefer et al., 2007).

Segmentation ofN^S extension inwesternAnatoliawasaccommodated by the Izmir-Balıkesir Transfer Zone(IBTZ), a deep crustal transform fault in the Late Cretac-eous that later acted as a transtensional transfer fault zoneduring theNeogene (Okay & Siyako,1993; Okay et al., 1996;Ring etal.,1999;�zkaymak&S˛zbilir, 2008;Uzel &S˛zbi-lir, 2008; Uzel et al., 2010). The IBTZ separates two dis-tinctly extended terrains ^ western Anatolia to the E and

EAGE

Correspondence: Hasan S˛zbilir, Dokuz Eylˇl �niversitesi, Mˇ-hendislik Fakultesi, JeolojiMˇhendislig� i B˛lˇmˇ, 35100Bornova-izmir,Turkey. E-mail: [email protected]

BasinResearch (2011) 23, 423–448, doi: 10.1111/j.1365-2117.2010.00496.x

r 2010 The AuthorsBasin Researchr 2010 Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists 423

mailto:[email protected]

the Aegean region to theW (Ring et al., 1999).The IBTZ ischaracterised by a NE^SW structural trend, which is con-siderably di¡erent from the principal structural frame-work of western Anatolia. In the E part of the zone, akinematically linked crustal- scale MCC ^ the MenderesMassif (MCC) ^ and approximatelyE^WandNE^SWdis-

secting basins are the most prominent features of westernAnatolia (Fig. 1). Detachment fault systems in this pro-vince are associated with domal uplift of the MCC in thelower plate and the formation of asymmetric supradetach-ment basins in the upper plate. E-striking normal faultsbounding the hanging-wall blocks appear to be rooted in

Fig.1. (a) Regional map showing main tectonic structures inTurkey and surrounding regions and (b) simpli¢ed fault map showing themain fault system ofwestern Anatolia, with locations of structures described in this study (compiled from Reng˛r et al., 1985; Bozkurt &S˛zbilir, 2004, 2006;�zkaymak&S˛zbilir, 2008;Uzel&S˛zbilir, 2008; and the presentwork).Double arrows indicate reactivation of theIBTZ as a sinistral and later a dextral zone. GD, Gediz Detachment Fault; BMD, BˇyˇkMenderes Detachment Fault; AT, AegeanTrench; CT, CyprusTrench, IAZ, Izmir-Ankara Suture Zone.

r 2010 The AuthorsBasin Researchr 2010 Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists424

H. S˛zbiliret al.

Seyitog� lu et al. (2002) or to cut the present low-angle nor-mal faults (Koc� yig� it etal.,1999;Yılmaz etal., 2000; S˛zbilir,2001, 2002), which are commonly exposed in culminationsof theMCC (Bozkurt & Park, 1994). Recent structural stu-dies and seismicity in the IBTZ (Akyol etal., 2006; Bozkurt& S˛zbilir, 2006; Zhu et al., 2006; �zkaymak & S˛zbilir,2008; Uzel & S˛zbilir, 2008; Uzel et al., 2010) have revealedthat N^S extension is accommodated (1) by dip-slipdisplacements on E-striking normal faults and (2) byslip on conjugate arrays of NW-striking sinistral andNE-striking dextral strike-slip faults (Fig. 2). Several linesof evidence, including geologic relationships, GPS mea-surements, and historic seismicity indicate that the seis-mically active IBTZ is undergoing E^W shortening in

addition toN^S extension (Aktug� &Kılıc� og� lu, 2006; Aktaret al., 2007).

To provide more detailed information on regional seg-mentation in western Anatolia, we present the results ofan integrated sedimentologic, stratigraphic, and structuralstudy ofMiocene alluvial fan/fan-delta deposits that accu-mulated in a lacustrine-dominated basinwith connectionsto the Aegean Sea. In this study, we mapped and evaluatedmost of the key faults in the KC� B, many of which were es-sentially unmapped, and recognised di¡erent deformationevents in the study area near the E margin of the MCC.Herein, we also document ¢eld evidence for the kinematicconnections between low-angle normal faults and strike-slip faults in a zone of transtensional tectonic setting.

Fig. 2. Simpli¢ed geologic map and cross section of NE-trending Cumaovası Basin andKocac� ay Basin.The basins were separated bybasement uplifts (e.g. theNifdag� ıHigh) during deposition of basin- ¢ll units.MFZ,Mahmutdag� ıFaultZone; SFZ,Spildag� ıFaultZone;KF, KemalpaSa Fault; KTF, Kalkancatepe Fault; OFZ, Orhanlı Fault Zone.


Tectonic implications of transtensional supradetachment basin

METHODS

Field-based studies consist of (1) mapping of geologicalstructures at a scale of1/25 000, (2) measurement of strati-graphic/sedimentologic sections of basin- ¢ll units, and(3) documentation of outcrop-scale faults and their kine-matic relationships. Some of these faults have beenmapped previously (Inci, 1991; Seyitog� lu et al., 2002; Kayaet al., 2007), but their kinematic features have not yet beendocumented in detail. We identi¢ed numerous locationswhere the sense of slip along brittle faults can be deter-mined using well-documented structural criteria (Han-cock, 1985; Means, 1987; Petit, 1987; Stewart & Hancock,1991).Virtually all faults have slickensides indicative of slipdirection, and in some cases, shear sense. Grooves andstriations from abrasion during slip, or elongate calcitecrystals present in dilational fault-surface jogs, also de-lineate slip direction. In this study, shear sense was deter-mined fromRiedel shears and from carbonate in dilationaljogs along slip surfaces. Relative ages of the di¡erent setsof faults were established by cross-cutting and o¡set rela-tionships. The observed faults were grouped into threesets that represent distinct styles or periods of faulting,herein referred to as D1 through D3. Some faults partici-pated in two periods of movement, as indicated by di¡er-ent sets of lineations on the same fault plane. A computerprogram (Angelier, 1984) was used to compute the charac-teristics of the palaeostress tensor associatedwith the per-iods of faulting.

THE KOCAC� AY BASIN (KC� B)

The basin formerly known as ‘KemalpaSa-Torbalı Basin’(Seyitog� lu, 1992) is herein referred to as the KC� B becausethe town ofKemalpaSa lies at the edge andwithin thewell-de¢ned Gediz Graben, not in the transtensional basin ofthe present study. Likewise, the town of Torbalı more cor-rectly lies in theKˇc� ˇk-Menderes Graben.‘Kocac� ay’ is thename of a river £owing along the centre of the elongate ba-sin (Figs 2 and 4).

The KC� B is a NE-trending asymmetric basin420 kmlong at the W termination of the MCC at the S margin ofthe Gediz Graben, continuing between two structuralhighs (Nifdag� ı High to the W and Mahmutdag� ı High tothe E) and terminating near theTorbalı district (Figs 1 and2). Exhumation of the core complex proceeded along thewell-known E-striking Gediz detachment fault (Hetzelet al., 1995; Emre, 1996; Emre & S˛zbilir, 1997; Koc� yig� itet al., 1999; Yılmaz et al., 2000; Lips et al., 2001; S˛zbilir,2001, 2002; ISık et al., 2003; Bozkurt & S˛zbilir, 2004).

Pre-basin-fill units

Basement rocks in the study area belong to theTauride-Anatolide block, which consists of many tectonostrati-graphic units, three of which are exposed: the MenderesMCC, the Bornova Flysch Zone, and the Cycladic Massif(Figs1 and 2).

The menderesMCC

TheMCC, a¡ected by Eocene Alpine metamorphism, is apart of the Alpine mountain chain in the easternMediter-ranean (Reng˛r & Yılmaz, 1981; Reng˛r et al., 1984). TheMCC was formed as a NE-oriented tectonic window; it isbordered to the N and W by rocks of the Izmir-Ankarazone and to the S by the Lycian Nappes.

The MCC is composed mostly of greenschist-faciesmetamorphic rocks intruded byTriassic andMiocene plu-tonic rocks.These rocks are tectonically overlain by med-ium-pressure amphibolite-facies migmatites andorthogneisseswith granitic to granodioritic compositions,consisting of para- and ortho-derived rocks with eclogiteboudins and metaophiolite lenses (Candan et al., 2001;�zer et al., 2001; �zer & S˛zbilir, 2003).These two tecto-nostratigraphic packages are separated by the well-knownGediz detachment fault, which is a crustal- scale low-angle normal fault forming the S boundary of the Gedizsupradetachment basin (Fig. 1). The detachment placesMiocene clastic units on chlorite breccia of myloniticgneiss, with brecciation a¡ecting a structural thickness ofabout 100m. Footwall rocks in the study area are repre-sented by a Permo-Carboniferous metasedimentary se-quence, i.e. the Bayındır Nappe ^ the structurally lowestunit in the MCC, composed of quartzites, biotite-richmica schists, garnet mica schists, and phyllites with inter-calations of thinly bedded grey and black marble lenses(Gessner et al., 2001; Okay, 2001).

The CycladicMassif

The Cycladic Massif is composed mainly of mica schistswith intercalations of calc-schists, marbles, thinly beddedcherts, and lenses of ma¢c metavolcanic rocks (now in theform of amphibolite and chlorite schist exhibiting Eocenehigh-pressure metamorphism) overlain by platform car-bonates (Okay, 2001).The contact between the MCC andthe overlying CycladicMassifwas inferred to be an uncon-formity by Erdog� an & Gˇng˛r (1992). However, structuralstudies suggest that the contact is a shear zone with a top-to-the-NE sense of shear (Okay, 2001).

The Bornova Flysch Zone

The Bornova Flysch Zone (‘Bornova melange’of Erdog� an,1990) is composed of various-sized blocks of Mesozoiclimestones, cherts, submarine volcanics, and serpentinitesembedded in a £ysch-type sedimentary matrix (Okay etal.,1996). It is a NE-trending tectonic zone lying between theMCC and the Izmir-Ankara Suture (Fig. 1). The BornovaFlysch Zone, Late Cretaceous to Palaeocene in age, hasundergone signi¢cant Alpine deformation, with a verylow metamorphic grade (Erdog� an, 1990; Okay & Siyako,1993; Okay & Altıner, 2007). Details of the stratigraphy ofthe pre-basin- ¢ll units are beyond the scope of this paperbut are summarised in recent literature (Erdog� an, 1990;Okay & Siyako, 1993; Okay et al., 1996; Bozkurt &Oberh�n-slı, 2001; Okay & Altıner, 2007; Okay, 2008).


H. S˛zbiliret al.

Basin-fill stratigraphy and depositional facies

The Neogene stratigraphic record of the KC� B is repre-sented by two unconformity-bounded sequences: the Ke-malpaSa Group and the Kızılca Formation (Figs 2 and 4).Strata of the KemalpaSa Group were initially describedbyVerdier (1963) as a Neogene detrital sequence and sub-sequently divided into lower and upper units by Inci(1991). Akdeniz et al. (1986) renamed the lower unit as theKesmedag� ı Formation and the upper unit as the ViSneliFormation. Seyitog� lu (1992) divided the basin ¢ll into twoconformable units: the Kesmedag� ı Formation, consistingof alternating reddish conglomerate, sandstone, and mud-stone, and the C� altaSı Formation, comprising alternatingyellow^brown^gray conglomerate, sandstone, and mud-stone with lenses of limestone and coal. However, Seyito-g� lu (1992) did not map the distribution of these units.Recently, Kaya et al. (2007) mapped the N part of the KC� Band divided the basin ¢ll into two formations ^ theKemal-paSa Formation and the unconformably overlying AlaSehirFormation. However, none of these studies mapped theentire basin or included measured sections.

In this study, all basin- ¢ll units were mapped, the best-exposed sections were measured, and a number of ¢eldcross sections were constructed in order to establish thestratigraphic relationships. Figures 3 and 4 summarisethe geology and stratigraphy of theKC� B ¢ll.The main de-positional units were combined as the ‘KemalpaSa Group’in this study. This group comprises the Derek˛y and ViSneli formations. The middle to upper Miocene KızılcaFormation lies unconformably above the ancient basin- ¢llunits.The youngest sediments are the Pleistocene Sˇtc� ˇ-ler Formation and Holocene alluvium. Two stratigraphicsections, ranging in thickness from 540 to 1300m, were

measured from di¡erent parts of the basin in order to de-termine lithologic changes and facies characteristics of thebasin- ¢ll units.Twelve sedimentary facies based on lithol-ogy, sedimentary structures, grain size, grading, and ma-trix content were recognised within the sections of theDerek˛y and ViSneli formations (Table 1).These facies arenamed in accord with the schemes of Miall (1977, 1978,1985), Lowe (1982), and Smith (1986). A complete and de-tailed description of the facies is beyond the scope of thepaper; however, an outline of the types of facies is pre-sented inTable 1.These facies are grouped into four faciesassociations, each representing di¡erent depositional en-vironments: alluvial fan, subaerial fan-delta, subaqueousfan-delta, and lacustrine. Detailed descriptions of themeasured sections are given below.

Derek˛y formation

The name ‘Derek˛y Formation’ is here given to a coarsen-ing-then- ¢ning-upward clastic sequence, 41300m inthickness, with lenses of limestone and coal. The lowerpart of the unit, herein termed the ‘Kesmedag� ı Member’,includes blocks of re-crystallised limestone-dolomite invarious sizes (Fig. 5a). The succession is made up ofalluvial-fan and £uvial clastics that grade into lacustrinesediments that are thickest towards the E margin of thebasin. Near the Mahmutdag� ı High and adjacent to theMahmutdag� ı fault zone (MFZ), the unit consists mainlyof coarse-grained clastic facies (Fig. 5b). Clasts are mostlyderived from marble and schist of the MCC/CycladicMassif and limestone, sandstone, and ophiolitic rock ofthe Bornova Flysch Zone. Upsection, ¢ning-upwardconglomerates display erosive bases and occur with

Fig. 3. Lithostratigraphic section of study area and correlationwith Cumaovası Basin.



cross-bedded sandstone. Some erosive surfaces are ob-served in underlying dark-brown sandstones. Conglomer-ates at these levels also include clasts derived from algallimestone lenses (Fig. 5c and d). Lacustrine micritic, algallimestone interlayers and organic-rich limestone lensesare also observed (Fig. 5e and f). In general, grain size de-creases westwards, indicating westward transport from atopographic high (Mahmutdag� ı High) E of theMFZ.Thissuggests that theMFZmarks the E margin of the Derek˛yFormation.

Derek˛y stratigraphic section

This section is located along the Mahmutdag� ı-Derek˛yroad (Fig. 4).The lower 50m is characterised by the relative

abundance of facies Gh, Gmc, and Sm (Fig. 7a and b;see Table 1 for detailed facies descriptions, and Fig. 6for explanation). Coarse-grained red to dark-red sand-stone with conglomerate lenses lies at the base of the se-quence. Conglomerate clasts include schist, marble, andquartzite.Red to dark-red colours are dominant at the low-er levels, whereas yellowish-orange and light-green col-ours are dominant at the upper levels, consistent with atransition from a subaerial to subaqueous fan-delta set-ting. The subaerial part of the fan-delta complex wasdominated by hyperconcentrated stream £ow and lesscommon debris £ow processes. Such a succession alongwith the presence of rounded clasts suggests an alluvialfan environmentwith high-energybraided streams (Miall,1996).

Fig.4. Detailed geologic map ofKocac� ay Basin (location in Fig. 2).MFZ,Mahmutdag� ı Fault Zone;KoF,Karaot Fault; KeF,Kesmedag� ıFault; Kıf, Kızılkaya Fault; AkF,Akkayatepe Fault; ReF, RekerolukFault; DeF, Derek˛y Fault; DaF,Dag� kızılca Fault; KaF, KarabeldereFault.


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Massive-to-planar-bedded clast- supported conglom-erate (Gm, Gh) is dominant at the 51^136m level of thesuccession. Conglomerate beds in this part exceed 10m.Clasts include algal limestone, schist, quartzite, and mar-ble. Coarse-grained sediments were transported bybraided streams across alluvial fan complexes into the lake.Such sediment in£uxes are evidenced by channelised con-glomerate containing intraclasts of underlying stromatoli-tic limestones. Some conglomerates grade into cream-coloured, thin-bedded, cross-laminated sandstone (Sr)and mudstone (Fm). The lowest limestone level (Lms) is1.8m thick and observed at the 150m level. Limestonesare grey, cream-coloured, laminated, and stylolitic as wellas bituminous at the base. Overlying beds are dominatedby cream and grey thin-to-medium-bedded sandstones(Sh, Sr, and Sm), which pass upwards into greenish-creamlaminated mudstones (Fm) up to the 230m level. Faciescharacteristics suggest a stream £ow-dominated alluvialfan to a subaerial fan-delta. This interval is overlain by a30-m-thick massive conglomerate (Fig.7). Clast composi-tions include quartzite, white marble, cream and grey algallimestone, dark-grey marble, dark-red chert, green ser-pentinite, and light-brown sandstone. Upsection andwestwards, the succession is represented by intercalationsof conglomerate and sandstone-mudstone packages up tothe 552m level.A thin coal lens is observedwithin pinkish-creammudstone at 309m.Fining- and coarsening-upwardbeds in the conglomerates are common. These deposits

are overlain by a 2-m-thick dark-grey algal limestone bed(Lms), which contains algal laminites and oncoids. Thelower and upper parts of the limestone bed are rich in bi-tuminous material. Clastic channel- ¢ll deposits are over-lain by stromatolitic and cryptalgal laminites intercalatedwith thin coal lenses. These relationships imply periodsof ponding on the channel £anks (Marzo & Anadon,1988). Overall, the facies characteristics suggest a subaerialto subaqueous fan-delta together with a wave-in£uencednearshore lacustrine environment (Horton & Schmitt,1996).

Meandering stream deposits are abundant in the mid-dle part of the Derek˛y Formation.They consist of ¢ning-upward sequences that are about 5^20m thick, consistingof a basal channelised conglomerate and cross-beddedsandstone overlain by mudstone. Some sequences arecapped by algal limestones and thin coal seams.The con-glomeratic units rapidly thin basinwards and merge intochannel- ¢ll conglomerates intercalated with sandstone-mudstone couplets. These include cross-bedded sand-stone (Sr) and horizontally laminated conglomerate (Gh)deposited by braided streams and thinly inter-beddedsandstone (Sh) and mudstone (Fm) formed as overbankdeposits with coal seams. This facies association re£ectslow-sinuousity braided stream deposition in a delta plainenvironment.

Above a 15-m-thick covered interval, a mudstone-dominated package caps the underlying clastic interval

Table1. Description and interpretation of lithofacies of KemalpaSa Group. Facies codes and interpretations are modi¢ed fromMiall(1977, 1978, 1985), Lowe (1982), and Smith (1986).

Facies codes Description Interpretation

Gm Poorly sorted massive, matrix-supported conglomerate,disorganized conglomerate with no appaerant locallyinversely graded

Debris £ow representing sedimentation in an alluvial fanenvironment or subaqueous debris £ow

Gmc Poorly sorted, massive to normally graded, clast-supported conglomerate; massive, poymictic, sandymatrix

Subaerial hyperconcentrated £ow, high-gradient braidedstream or subaqueous high-density turbidity currents

Gp Poorly sorted moderately to well rounded, cross-beddedconglomerate, coarse sand matrix, ¢ning upwardwitherosive base

Channel ¢ll to channel bar representing sedimentation ina stream-dominated alluvial fan environment

Gh Poorly sorted, clast supported congomerate, horizontalstrati¢cation, erosional to nonerosional boundary

Subaerial sheet£ood or wave-driven bedload traction

Sm Coarse grained,massive sandstone, pebbly, non-erosive toerosive base

Sand dominated hyperconcentrated £ow or subaqueoushigh-density turbidity current

Sh Medium to thin bedded, medium grained, horizontallystrati¢ed sandstone with erosive base

Subareal sheet£ood, wave-driven bedload traction orhigh/low density turbidity current

Sr Medium bedded, medium to coarse ¢ne grainedsandstone with ripple-cross lamination

Subaqueous deposits at lower £ow regime, current ripples

Fm Massive, thin bedded to laminated mudstone, locallybitumous with plant remnants

Low energy, overbank deposit, suspension fallout

Lms Thick, medium bedded to laminated limestone, algallaminated oncoids and pisoids, stromatolitic, localy richin bituminous material

Low energy, carbonate precipitation and binding ofcalcareous sediment by algal mats, ephemeral lake

C Brownish-black coloured, thin bedded to laminatedorganic material

Low energy, coastal plain, overbank deposit, swamp

Vft Massive to thick bedded, locally graded felsic tu¡ Pyroclastic £ow, locally reworked



(Fig. 7). This package continues upwards to 1023m andincludes conglomerate, sandstone, and limestone inter-beds. Cream, greenish-grey, and dark-red coal interlayersare locally present, but sandstone (Sm) interbeds witherosive and gradational bases are rare. The medium-to-thick-bedded conglomerates have erosive bases andexhibit local £aser-like bedding. Coarsening- and ¢ning-upward beds are common in the conglomerates. Algallimestone beds are grey to dark-grey and includeabundant tube-like and spherical (oncoidal) algal struc-tures. The sequence at 1023^1310m is represented byintercalations of conglomerates and sandstone-mudstonecouplets with erosive bases. The facies characteristics

suggest a fan-delta plain and a nearshore lacustrineenvironment.

Visneli formation

The ‘ViSneli Formation’ (Akdeniz et al., 1986) here refers toa ¢ning-upward succession of alluvial fan and sheet£ood/lacustrine deposits and interbedded tu¡s. At the base, theformation comprises an interval of reddish-grey andlight-green^grey conglomerate, laminated sandstone,and mudrock. The succession grades upward into papershales that contain sporadic thin intercalations of sand-stone and felsic tu¡, and alternate upwards with volcani-

Fig. 5. Field photographsrepresenting basin- ¢ll facies: (a)Limestone olistolith embedded inred coarse sandstone alternatingwith poorly sorted conglomerate.(b)Massive blocky conglomeratefacies at the base of unconformity-bounded packages of Derek˛yFormation. (c) Cross-beddedpebbly sandstone facies betweenclast-supported conglomerate bedswith a reddish-brown matrix ofsand and ¢ne gravel. Hammeris �34 cm long. (d) Reddish-wine-coloured sandstone and channel-¢ll conglomerate with erosive lowercontact. (e)Whitish-beige-coloured laminated to thin-beddedlimestone, conformably overlain byreddish-brown sandstone andconglomerate. (f) Thin- tomedium-bedded lacustrinelimestone facies, includingsynsedimentary folds. (g) Creamand light-grey planar-laminatedwell-bedded mudstones. (h) Tu¡bedwith sharp lower boundary andgradational upper boundary withsandstone beds of ViSneliFormation.


H. S˛zbiliret al.

clastic sandstone, lithic conglomerate, and thicker felsictu¡ (Fig. 5g and h).The distinctive characteristics includethe laminated mudstone and interlayers of volcanic tu¡(Fig. 3). One stratigraphic section was measured betweenthe villages of ViSneli andDag� kızılca in order to determinelithological characteristics and depositional facies. Detailsof this section are as follows.

Visneli stratigraphic section

Typical outcrops of the Visneli Formation are observedalong roadcuts between ViSneli and Dag� kızılca (Fig. 4).The lower 50m of the formation comprise matrix-sup-ported conglomerates (Gm) with mudstone (Fm) inter-layers. The conglomerates are thick-bedded and grey,dark-grey, and yellowish-grey (Fig. 8, see Fig. 6 for expla-nation). Clasts are mostly derived from the Bornova £yschzone and include limestone, sandstone, quartzite, chert,and serpentinite.The alternation of debris £ow (Gm) andsuspension fallout (Fm) clastics are suggestive of deposi-tion in a subaqueous fan-delta environment.

Mudstones become dominant after the 53m level andcontinue to 129m. The mudstone sequence comprises,from bottom to top, cream and beige well-bedded mud-stones and light-grey planar laminated mudstones. A 50-cm-thick grey and dark-grey recrystallised limestone ap-pears at the 93m level. Three 20^50-cm-thick sandstonebeds are present within the mudstone sequence at 100^110m. The light-grey sandstone beds are planar-bedded

and have gradational boundaries with the mudstones.These composite assemblages may suggest a nearshore la-custrine environment. Upsection, a sequence representedby intercalations of mouth-bar conglomerates and sand-stones occurs at 129^157m. Polymodal matrix-supportedconglomerates with slightly erosive to non-erosive basalcontacts are observed in the middle of the succession. Asubaqueous mass- £ow origin is suggested by the lack ofinternal strati¢cation and non-erosive basal contacts.Sandstone interbeds within the conglomerate interval aregrey and coarse-grained and include lens geometries.Theuppermost part of the succession is represented by ma-trix-supported cross-bedded conglomerates.These faciessuggest a transition from Gm-dominated deposits of asubaqueous fan-delta environment to Gp-dominated fa-cies of a subaerial fan-delta environment.

The mixed conglomerate and sandstone interval isoverlain by amudstone-dominated succession (principallyfacies Fm) with intercalations of sandstone (Sm), tu¡ (Vft),and conglomerate (Gm).Tu¡ levels are massive-to-planar-laminated and have both sharp and gradational bound-aries.The clastic-dominated Visneli Formation is cut by afault at the 540m level. Suspension sedimentation (Fm)and turbidity-current deposition (Sm) together with thepresence of synsedimentary deformation and debris- £owconglomerates (Gm) may suggest an alternation of a near-shore/o¡shore lacustrine environment and subaqueousfan-delta environment for the upper ViSneli Formation.The observed Gm, Gp, Sm, Sh, Sr, Sm, Fm, Lms, and

Fig. 6. Explanatory chart oflithofacies and sedimentarystructures used in measuredstratigraphic sections.



Fig.7.

(a,b)M

easuredstratigraphicsections

ofDerek˛y

Form

ation(lo

cation

inFig.4),showingthickn

ess,litho

facies,and

sedimentary

structures.See

Fig.6forexplanation

andTable1

ford

escription

andinterpretation

offacies

codes.


H. S˛zbiliret al.

Fig. 8. Measured stratigraphic section of theViSneli Formation (location in Fig. 4. See Fig. 6 for explanation andTable1 for descriptionand interpretation of facies codes.



Vft facies of the ViSneli Formation can be classi¢ed intothree facies associations: (1) subaerial fan-delta, (2) suba-queous fan-delta, and (3) lacustrine.

Kızılca formation

Unconformably overlying the ancient basin- ¢ll succes-sion is the Kızılca Formation. Well exposed between thevillage of Kızılca and Nifdag� ı High, it is a package of non-marine clastic rocks and lacustrine carbonate exposed atthe SW end of the Gediz Graben. The formation is con-¢ned not only to the interior of the basin; but it also occurswell away from the basin centre (Figs 2 and 4).

The formation is made up of brown poorly sorted allu-vial conglomerate at the base, passing upwards into grey-ish-brown sandstone and mudstone.The succession endswith light-grey and yellowish-white clayey limestone in-tercalations at the top. Carbonate-dominated strata repre-senting shallow lacustrine facies associations are well

exposed at Kalkanca Hill, where the lower part of the for-mation is composed of yellowish-grey wavy laminatedmudstones intercalated with brecciated algal limestones.Upsection, the unit consists of whitish-beige thick-bedded limestones alternating with yellowish-brownmudstones, including iron oxide coatings of algal carbo-nate nodules.Towards the W, the lateral equivalent of theformation is theYenik˛yFormation,which is middle to lateMiocene in age, which is known on the basis of radio-metric ages for the interbedded Cumaovası Volcanics(Uzel & S˛zbilir, 2008). The formation has been upliftedand dissected into several small- scale horsts and grabensby a series of strike-slip and oblique slip normal faults.

Sˇtc� ˇler formation

The Sˇtc� ˇler Formation comprises thick-bedded alluvial-fan conglomerates derived from the uplifted Spildag� ıHigh. These basin-margin fans are exposed along the N

Table 2. Detailed results of 40Ar/39Ar analysis carried out on tu¡ sample KT-01; (a) analytical and (b) graphical data.

T ( 1C) cum.39 Atmos. (%) 36Ar/39Ar 37Ar/39Ar 38Ar/39Ar 40Ar/39Ar 40Ar/36Ar Date (Ma)

(a)680 .017 92.211 0.7581E� 01 0.1995E100 0.2511E� 01 0.2430E102 0.3206E103 20.5 � .8750 .069 88.307 0.3944E� 01 0.1373E100 0.1845E� 01 0.1321E102 0.3350E103 16.8 � .5820 .163 78.204 0.1685E� 01 0.1149E100 0.1408E� 01 0.6384E101 0.3788E103 15.1 � .2900 .254 23.477 0.1327E� 02 0.8016E� 01 0.1154E� 01 0.1671E101 0.1259E104 13.7 � .1950 .373 25.335 0.1456E� 02 0.7459E� 01 0.1184E� 01 0.1702E101 0.1169E104 13.6 � .11000 .485 29.905 0.1806E� 02 0.7152E� 01 0.1193E� 01 0.1793E101 0.9926E103 13.5 � .11050 .582 36.569 0.2443E� 02 0.6956E� 01 0.1211E� 01 0.1986E101 0.8129E103 13.5 � .21100 .677 44.000 0.3285E� 02 0.5630E� 01 0.1233E� 01 0.2223E101 0.6766E103 13.4 � .11150 .770 49.410 0.3893E� 02 0.4566E� 01 0.1241E� 01 0.2348E101 0.6030E103 12.8 � .21200 .855 53.591 0.4691E� 02 0.3808E� 01 0.1259E� 01 0.2607E101 0.5559E103 13.0 � .21250 .923 52.387 0.4600E� 02 0.3359E� 01 0.1256E� 01 0.2616E101 0.5687E103 13.4 � .21300 .961 63.985 0.7549E� 02 0.3406E� 01 0.1308E� 01 0.3509E101 0.4648E103 13.6 � .41350 .984 80.486 0.1683E� 01 0.3893E� 01 0.1489E� 01 0.6201E101 0.3685E103 13.1 � 1.01500 1.000 83.573 0.1987E� 01 0.4235E� 01 0.1600E� 01 0.7048E101 0.3547E103 12.5 � .8

Sample mass5 29.4mg.J-value5 .00604875 .00004160.Integrated date513.8 .1Ma.

(b)


H. S˛zbiliret al.

and S margins of the modern Gediz graben (Fig. 2). Thematrix-supported, poorly sorted conglomerates formingthe main constituent of the alluvial fan deposits consist ofsub-rounded clasts with poor sphericity. The matrix ismoderately to poorly consolidated.The thick-bedded con-glomerates dip towards the KemalpaSa Plain.This forma-tion is unconformably covered by Quaternary alluviumforming the modern KemalpaSa Basin.

Alluvium

The alluvium unit is the ¢nal product of the KemalpaSaPlain, which is bounded by normal faults to the S and N(Figs 2 and 4).This unit is made up of coarse-grained allu-vial-fan and ¢ne-grained £uvial deposits.The alluvial-fandeposits are controlled by theKemalpaSa fault to the S andthe SpilMountain fault zone to theN.The £uvial depositsare axial river deposits of the KemalpaSa River. The fansand axial river deposits inter- ¢nger laterally and verticallyand represent typical graben ¢lls. The middle part of thestudy area is also covered by alluvium consisting of allu-vial-fan and river deposits. The alluvium observed S ofthe village of Derek˛y to the Torbalı district was trans-ported by the Kocac� ay River. These deposits continue asa long corridor in the NE^SW direction and terminate asa S-widening fan towards theTorbalı district.

Age and stratigraphic relationships ofbasin-fill units

The age of the KC� B sedimentary succession has been as-sessed by previous studies on the basis of regional strati-graphic correlations. Inci (1991) assigned a late Miocene-early Pliocene age on the basis of data on the tectono-sedi-mentary evolution of the western Anatolian graben sys-tems (Arpat & Bing˛l, 1969; Kaya, 1979). Seyitog� lu (1992)reported that coal levels in the basin did not yield any ageinformation. In the present study, however, the best ageconstraints on the basin ¢ll came from mammalian fossilsin mudrock and 40Ar/39Ar dating of tu¡s.

Within the upper Derek˛y Formation, a greenish-greymudrock zone contains a fossil assemblage (Aliveria sp.,Anomalomys cf. aliveriensis, Cricetodon sp., Democricetodonsp.,Eumyarion sp.,Galerix sp.,Megacricetodon cf. collongensis,Microdyromys sp., Erinaceidae, and Lagomorpha) indicat-ing a late Early Miocene age (Kaya et al., 2007, see Fig. 4for location of the MN4 mammalian zone, UTM coordi-nate: 053985E/424405N).

In addition to existing age data, we obtained a new40Ar/39Ar biotite age for a tu¡ (sample KT-01) from theupper ViSneli Formation (Table 2).The sample of biotite-rich felsic volcanic tu¡ yielded a 13.4� 0.1Ma plateau ageand 13.8� 0.1Ma isochron age (Figs 4 and 8; Table 2).40Ar/39Ar step-heating experiments were carried out onKT-01 at the Chinese Academy of Science. Aliquots ofmineral separates were wrapped in aluminium foil packetsand stacked in a quartz canister along with the neutron£ux monitor GA1500 biotite, which has a K-Ar age of

97.9 � 0.7Ma (Baksi et al., 1996). Samples were irradiatedat the H8 position in the 49-2 reactor at Beijing, China for28.5 h. After irradiation, samples were removed from thealuminium foil and degassed for 30min at 350 1C. Twosampleswere heated stepwise with a double-vacuum resis-tance furnace following the step schedule listed inTable 2.The released gas was puri¢ed with a titanium sponge andZr^Al getters. Isotopic compositions were measured usinganMM5499mass spectrometer at the Institute ofGeologyand Geophysics, Chinese Academy of Sciences, China.After corrections for mass discrimination, system blanks,isotope interferences, and radiometric decay, 40Ar/39Arages were calculated according to isotopic ratios and the Jvalue obtained by analyses of monitor standards.The iso-tope interference correction factors were (36Ar/37Ar)Ca5

2.609� 10� 4� 1.148� 10� 5, (39Ar/37Ar)Ca57.236� 10� 4

� 2.814� 10� 5, (40Ar/39Ar)K52.648� 10� 2� 2.254�10� 4. Plateau ages were calculated from adjacent ages thatagreedwithin a range of 2s.

In addition, the following fossil assemblages (foramini-fers and ostracods) were obtained from the ¢ne-grainedinterval of the Derek˛y Formation:Globorotalia obesa,Glo-bigerinoides trilobus, Globoquadrina dehiscens, Globorotalia cf.menardii, Globigerinoides obliquus, Globorotalia scitula scitula,Orbulina universa, and Indet. Candonas. Although this as-sociation does not indicate a precise age, the co-occur-rence of planktonic foraminifers and ostracods indicate amarine connection with the lacustrine basin.

Frompublished data, the age and stratigraphic relation-ships of deposits in the neighbouring Cumaovası Basincan be outlined as follows.The oldest ¢ll in the Cumaovasıbasin is represented by the C� atalca Formation (Fig. 3),which is composed of laminated siltstones, sandstones,and shale alternations including lignite lenses and thin-bedded conglomerate levels (Genc� etal., 2001).TheC� atalcaFormation has been interpreted as lacustrine fan-delta se-dimentation and dated as early to middle Miocene in agefrom palaeontological and palynological studies (Akartu-na,1962;Kaya,1979,1981;Genc� etal., 2001;Uzel &S˛zbilir,2008). The conformably overlying Yamanlar volcanics arecomposed of several lava £ows, pyroclastic rocks, dykes,and domes of dacitic, andesitic, rhyolitic, and basalticcompositions. Ages of 19.2^14.7Ma have been reportedfor the Yamanlar Volcanics (Borsi et al., 1972; SavaSc� ın,1978; Ercan et al., 1996), indicating an early Middle Mio-cene age.The unconformably overlying volcano-sedimen-tary sequence beginswith the�rkmezFormation,which isdominated by red conglomerate and sandstone alternatingwith lacustrine limestone lenses (ESder & RimSek, 1975;Genc� etal., 2001).Upsection, a ¢ning-upward pro¢le is de-¢ned by conglomerates that are replaced gradually with¢ne-grained conglomerates, sandstones, and mudstones.In the upper levels, these clastic rocks are interbeddedwith lacustrine limestones of the Yenik˛y Formation (ESder&RimSek,1975;Genc� etal., 2001).These units are inter-bedded with pyroclastic rocks and rhyolitic lava £ows ofthe Cumaovası Volcanics, which yielded K/Ar ages of 11.5^9Ma (Borsi etal.,1972; ESder&RimSek,1975;�zgenc� ,1978;



Genc� et al., 2001). On the basis of the stratigraphic relation-ships of the basin- ¢ll units with the Cumaovası Basin ¢ll, itis proposed that the lower-middleMiocene C� atalca Forma-tion observed at the base of the Cumaovası Basin is the lat-eral equivalent of the Derek˛y Formation and that theunconformably overlying middle-to-upper Miocene Yeni-k˛yFormation correlateswith theKızılcaFormation (Fig. 3).

STRUCTURES

We recognised two main types of structures: (1) those thatshaped theKC� B andplayed key roles during sedimentationwithin the basin (D1phase) and (2) those that deformed ba-

sin- ¢ll and basin-bounding growth faults (D2 and D3phases). Syn- to post-sedimentary structural features havebeen observed in the ¢eld as indicated in Figs 4, 9 and10.

Synsedimentary structures (D1phase)

D1structures thatwere developed during the formation ofthe KC� B fall into three categories: (1) theMFZ, (2) synse-dimentary folds, and (3) intraformational unconformities.

TheMFZ

The breakaway faults and ¢ll of the supradetachment ba-sins in the Gediz Graben end abruptly near the boundary

Fig.9. Geologic cross sections showing stratigraphic and structural relationships of basin ¢llwith older basement units. SectionsA-A0,B-B0, andC-C0 are taken from theKesmedag� ımember of theDerek˛y Formation. SectionsD-D0, E-E0 , and F-F0 represent syntectonicintraformational unconformities in the lower Derek˛y formation.


H. S˛zbiliret al.

between the basin ¢ll of the KC� B and the MCC at theNE-strikingMFZ, which forms right-stepping structuralcon¢gurations (Fig. 4). The MFZ is a N201E-strikingrange-front fault zone that is approximately 2 km wideand 22 km long. It is located in the area betweenTorbalı intheS andAkkayatepe in theN,where it meets at theSmar-gin-bounding faults of the Gediz Graben.

The MFZ consists of two types of faults: NE-strikingstrike- slip and E-striking low-angle normal faults (Figs9d, f, 10a, and c). The NE-striking strike- slip segmentsare en echelon and right stepping, and they mark thestructural contact between the MCC/Cycladic Massifand the Derek˛y Formation. Several low-angle fault sur-

faces have been established on the structural steps withinthe fault zone, where a sharp striated contact de¢nes thedetachment surface.These low-angle fault surfaces strikeE^Wanddip northwards, but southward tilting can also beobserved owing to post-Miocene tectonics. Where ex-posed, the fault zone is de¢ned by a1^2-m-thick cataclas-tic zone containing angular mylonitic clasts of footwallrocks. Immediately below the cataclastic zone is a se-quence of mylonitised biotite schist and marble intercala-tions. Well-developed mylonitic texture is preserved inmylonitic gneiss along the W £ank of Mahmutdag� ı High,yielding a top-to-the-NE sense of shear. The averagelineation of the metamorphic rocks is N201E, and fold

Fig.10. Field photographs of (a^c)synsedimentary and (d^h)postsedimentary structures. (a)Low-angle normal faults alongwhich metamorphic rocks and algallimestone of the Derek˛yFormation are tectonicallyjuxtaposed. (b) Close-up view oflow-angle fault surface withoblique-slip slickensides. (c) Close-up view of brecciated marble withwell-de¢ned slip surfaces.Slickensides indicate dextral strike-slip motion. (d) Superimposedslickensides with older strike-slipmotion and younger dip-slipnormal motion. (e) Field view of N-striking strike-slip fault. (f) Detailof slip surface within Kızılkayafault, showing dextral strike-slipslickensides. (g)Well-developedfault plane of Derek˛y fault with anearly dip-slip slickenline. (h)Close-up view of slickensidedplane of NW-striking sinistralstrike-slip fault.



axes developed parallel to that direction are observed.Synkinematic indicators within the lower plate of theGediz detachment also indicate top-to-the-NE motionbetween 19.5 � 1.4 and 7 � 0.11Ma on the basis of an40Ar/39Ar amphibole cooling age and biotite fabric agefrom a ‘synextensional’ granodiorite, respectively (Hetzelet al., 1995; Lips et al., 2001).

The NE-striking fault segments are dextral strike-slipfaults with strikes averaging N201E, with a 101NE averagerake of slip lines (Fig.10c and Table 3) while the low-anglefault surfaces are those of an oblique-slip normal fault,dipping an average of 201NW with a minor componentof dextral strike slip (Fig. 10a and b). These data indicatekinematic linkages between normal and transverse

Table 3. Detailed information on faults and stress tensors where slickenlines were measured in the Kocac� ay Basin.

Name of fault Nature of fault Number of slip data s1 s1 s1 f ANG

D1Phase Mahmutdag� ı Normal-slip 05 080/72 316/11 223/15 0.418 06Mahmutdag� ı Strike-slip 06 096/18 300/71 188/08 0.388 06

D2 Phase Derek˛y footwall Strike-slip 12 152/06 036/77 243/12 0.139 13Kesmedag� ı Strike-slip 05 222/04 116/76 313/13 0.279 12Kalkancatepe Normal-slip 08 291/74 186/04 095/16 0.469 12Karabeldere Normal-slip 07 057/72 156/03 247/18 0.168 06Derek˛y Normal-slip 11 027/78 134/04 225/12 0.363 06Karaot Normal-slip 07 094/22 336/50 199/32 0.652 14

D3 Phase Spildag� ı and KemalpaSa Normal-slip 11 222/82 332/03 062/08 0.702 10

Fig.11. Lower-hemisphere, equal-area projection of fault planes,slickenlines, and stress orientationsfor tectonic phases that operatedduring (D1) and aftersedimentation (D2 andD3) inKocac� ay basin development.


H. S˛zbiliret al.

strike-slip faults in accommodating regional extensionand that the KC� B is a transtensional supradetachmentbasin.

In order to determine the stress tensor responsible forthe kinematic character of these faults, we used the pa-laeostress method of Angelier (1984).The results de¢ne ahorizontal s3, approximately trending N^S (Fig. 11 andTable 3). The obtained stress axes are similar to those ofthe low-angle normal faulting determined in the Gedizdetachment (S˛zbilir, 2001, 2002; Bozkurt & S˛zbilir,2004) and are considered to characterise a state of regionalstress developed during theMiocene.

Synsedimentary folds

At many outcrops, lacustrine deposits exhibit strong soft-sediment deformation such as convolute laminations,water-escape structures, and disrupted beds, with someof them associated with folds (centimetre-scale); only twosites show m-scale synsedimentary folds trending NE^SW (Fig. 5f). Given their structural association, the originof these minor structures appears to be connected to syn-sedimentary activity on strike-slip faults that a¡ected thelacustrine deposits. The observed synsedimentary foldsare overturned to the W, suggesting mass- £ow progressfrom the E margin of the basin to theW.

Intraformational unconformities

Four step-like structures covered by di¡erent levels of thesedimentary succession were established along the MFZ(Figs 4 and 9). Each unit of the succession close to the faultis unconformity-bounded.These unconformable contactspass downdip and along the strike into conformable rela-tionships. A sourceward stepping of unconformity-bounded sedimentary packages indicates that the faultprogressively stepped back towards the source area. Upliftwithin the source area and subsidence in the basin wereapparently discontinuous, with periods of active faultmovement accompanied by development of an alluvialfan/fan-delta followed by tectonically quiescent periodswhen subaerial erosion created intraformational angularunconformities. The most convincing evidence is givenby di¡erences in the degrees of stratal dip between uncon-formity-bounded sequences.

The syntectonic unconformities (similar to Riba, 1976)observed in the Derek˛y Formation are restricted to themarginal areas of the E basin margin, and they die outacross the basin centre towards the W.The best examplesof these are observed on the roadcut around the village ofKaraot (Fig. 4). A representative ¢eld cross section wasconstructed to show the stratigraphic position of intrafor-mation unconformities within the Derek˛y succession(Fig. 9). Each package begins with blocky conglomerate atthe base (Fig. 5b), passing upwards into ¢ne-grained clas-tics intercalated with thin carbonate lenses.This suggestsmotion along the E-striking low-angle normal faults, gen-erating extension subparallel to the NE-trending strike-

slip margin of the basin throughout the depositional his-tory. The intraformational unconformities indicate thatthe MFZ acted as a NE-striking growth fault controllingearlyMiddleMiocene sedimentation within the KC� B.

Structures that postdate the basin fill

Structures that deform basin- ¢ll units formed during awrench- to extension-dominated deformational phaseswith local compressional strain (D2 phase), whichwas fol-lowed by an extensional phase (D3 phase). Evidence fordistinct deformational phases comes from (i) two types offold-to-fault relationships: folding parallel to the E-strik-ing normal-fault traces and folds oblique to theNE-strik-ing strike-slip fault traces and (ii) reactivation of olderfaults, as evidenced by local superposition of steeply plun-ging slickenlines with a normal sense of shear over oldersubhorizontal slickenlines with strike-slip kinematics.

D2 phase

TheD2phase is characterised by the coexistence of conju-gate strike-slip faults, E-striking reverse faults, variouslyoriented folds, and N- to E-striking normal faults. Theobserved strike-slip faults are hard-linked by moderatelyto gently dipping oblique-normal connecting faults.Theydisplay geometries and kinematics consistent withchanges in ¢nite strain axes from wrench- to extension-dominated deformation, with the maximum shorteningaxis oriented in the horizontal and vertical positions, re-spectively.These structures are outlined as follows.

Karabeldere fault zone

This fault zone is mapped as three discrete fault traces ex-tending for about 20 km between Kızılca in the N and theTorbalı district in the S (Fig. 4).TheN part consists of a 3-km-long NE-striking fault zone. Clastic and carbonaterocks of the Kesmedag� ı member are intensively shearedalong the fault zone, which includes several dextral faultsegments showing an en echelon fault pattern in a mainlyNE^SW direction. Several well-exposed fault planesstrike approximately NE^SW, with an average 72^881SEdip and rakes of slip lines averaging 2^221NE.

In the axis of the fault zone,wemeasured two di¡erentlyoriented striation sets on the same slip surface of the Kar-abeldere fault zone (Fig. 10d), which strike N651E and dip821SW. The younger set is represented by slip lines withrakes of 80^891SW that overprint the older striation setwith an average rake of 51NE.The two observed slicken-side lineations with di¡erent plunges and slip senses onthe same fault plane suggest that the strike-slip surfacewas overprinted by dip-slip movement.

To the S, along the Kocac� ay River, the moderate to stee-ply dipping Derek˛y Formation is cut by well-exposedslickensided fault surfaces.Here, fault planes show normalfaulting with a minor dextral- slip component striking ap-proximately N^S and dipping 49^701W (with rakes of 78^851N).



Karadere fault

TheN-striking right-lateral Karadere fault cuts theDere-k˛y Formation where it is bounded by the E-striking Ak-kayatepe fault to the N (Fig. 4). The Karadere fault is anapproximately 1-km-long N-striking dextral strike-slipfault that displays a steeply dipping high-relief (up to 5m)fault scarp (Fig.10e).

Kizilkaya fault

The Kızılkaya fault de¢nes a linear E-trending topo-graphic valley. To the E, the fault appears to terminate atthe Karabeldere fault (Fig. 4). Several fault-sense criteria,such as Riedel fractures and displaced clasts at the end ofthe grooves, indicate dextral slip. Several well-exposedfault planes strike approximately E, with average dips of64^891N in conglomerates of the Derek˛y Formation.Well-preserved slickenlines on these planes show dextralstrike-slip o¡set with a minor reverse-slip componenthaving rakeso251E (Fig.10f).

Reverse faults

No mappable reverse faults were observed in the KC� B.However, in the W block of the Karabeldere fault (alongthe KemalpaSa-Torbalı road), the red clastic part of theKesmedag� ı Member is cut and deformed by low-anglethrust planes dipping S with well-developed slickensides.In the outcrop, the features occur as a brittle zone com-posed of curvilinear and closely spaced reverse-faultplanes.Detailed studies of fault structures such as slicken-sides, steps on the slip surfaces, and cross-cut markers in-dicate reverse motion on the E-striking fault.The fault iscut and downthrown by a high-angle N-dipping normalfault. This indicates that the reverse faults postdate sedi-mentation of the Kesmedag� ı Formation but are older thanthe E-striking normal faults associated with the NE^SWextension.

Kalkancatepe fault

An obvious N-striking scarp marks the Kalkancatepefault, where limestones of the Bornova £ysch zone are jux-taposed against lacustrine carbonates and clastics of theKızılca Formation. The fault forms the NE boundary ofthe Nifdag� ı High (Fig. 4). East of Kalkanca Hill, the faultsegment displays a well-exposed fault plane with a 5^10mrelief andwell-preserved slickenlines. Stereographic plotsshow nearly dip-slip normal o¡set with attitudes aver-aging N-S/821E and N401W/381NE, where the rakes ofslip lines are 801N and 841SE, respectively.The Kalkanca-tepe fault is systematically cut and displaced by NE-strik-ing strike-slip faults (Figs 2 and 4;Table 1).

Kesmedag� i fault

The Nifdag� ı block is bounded to the E by a NE-strikingconcave-shaped strike-slip fault ^ the Kesmedag� ı fault

(Fig. 4). The fault zone is characterised by a highly brec-ciated zonewith limestone olistoliths.There are numeroussteep-to-subvertical slip planes with nearly horizontalslickensides. Several subparallel faults at the S end of theKesmedag� ı fault are de¢ned by NE-trending topographiclineaments (Figs 2 and 4).The principal motion along theslip surfaces is dextral strike-slip.There is some evidence,however, for earlier sinistral strike-slip motion along thefault zone where the calculated rakes for the older sinistralmotion are 2^101E.

Akkayatepe fault

The Akkayatepe fault forms the N structural boundary ofthe KC� B (Fig. 4). It juxtaposes the Kızılca Formation witheither metamorphic basement or the Derek˛y Formation.It displays a graben-facing step-like pattern connectedwith the S margin of the Gediz Graben E of the village ofKızılca, and the Kesmedag� ı fault to the W. It strikes E be-tween the Nifdag� ı High and Akkaya Hill and strikes NEnear �renk˛y.

The N side of Akkaya Hill is shaped by closely spacedlistric-fault surfaces showing a step-like con¢guration.There, the fault is an E-striking normal fault dipping atan average of 401N without preserved slickenlines. At itsWend,however, the fault has several slickensided fault sur-faces that indicate oblique and strike slips. Although nocross-cutting relationships exist between these two vari-able slickenlines, the strike-slip may be related to themovement on the Kesmedag� ı fault, whereas the oblique-slip may be related to Quaternary graben formation.

Derek˛y fault

The Derek˛y fault is mapped between the Karabel gateand the village of Derek˛y (Fig. 4). It is a 50-m-wide and2-km-long NW-striking normal fault along which thefolded and strike-slip-faulted Derek˛y Formation and theSE-dippingViSneliFormation are tectonically juxtaposed.The fault zone includes fault planes dipping at an averageof 651SWwith rakes of 70^881NW (Fig.10g).

Some well-preserved fault planes are exposed in thefootwall of the Derek˛y fault in a roadcut and are markedby a zone of slickensided calcite mineralisation (Fig. 10h).Two orientations of slickensided fault planes are domi-nant: N-S/801Wand N551W/701NE. Subhorizontal slick-enlines, together with accretionary calcite steps, indicate adominantly sinistral and dextral sense of motion, respec-tively (Table 3).These two fault segments represent conju-gate strike-slip faults.

Karaot fault

The Karaot fault is characterised by a curvilinear range-front fault along which W-dipping strata of the Derek˛yFormation and E-dipping strata of the ViSneli Formationare tectonically juxtaposed (Fig. 4). It is a 10-m-wide and4-km-long NW-striking and SW-dipping normal-faultzone, located between the Kocac� ay stream and the village


H. S˛zbiliret al.

of Karakızlar. The Karaot fault displays several well-ex-posed fault planes and slickenlines. Kinematic data showoblique-slip normal o¡set with a minor sinistral strike-slip component.The slickenlines show rakes of 48^801SE.

Folds

Folds occur as a series of anticlines and synclines with ap-proximately parallel axes (Fig. 4).They mainly trend E^Wand NE^SW. E-trending folds are well exposed betweenthe villages of ViSneli and Dag� kızılca, where they are cutby the Kesmedag� ı fault. They are large, open, gently W-plunging anticlines with gentle to moderately dippinglimbs.The fold axes are parallel to associated normal faultsand perpendicular to theN^S extension.The origin of themapped folds can be explained by (1) N^S compressionthat postdates deposition of the ViSneli Formation and (2)extensional tectonics with (i) buried normal faults under aseries of open folds or (ii) ramp- £at bends in detachmentsurfaces at depth (Schlıshe, 1995).They may be similar tofault-bend longitudinal folds mapped in the basin ¢ll ofthe Gediz Graben (S˛zbilir, 2001, 2002; C� iftc� i & Bozkurt,2008).

The origin of E^W trending folding that deforms Neo-gene units in western Anatolia has also been debated.Var-ious models fall into two groups: short-lived N^Scompression or synextensional folding. Folding may be re-lated to local fault-related folding (e.g. S˛zbilir, 2002) or ashort-lived event of contraction (e.g. Koc� yig� it et al., 1999).Punctuated N^S compression corresponding to the un-conformity between Miocene and upper Pliocene sedi-ments has been documented in other western Anatolianbasins (Koc� yig� it et al., 1999; Kaya et al., 2004; Bozkurt &Rojay, 2005; Bozkurt & S˛zbilir, 2006; Emre & S˛zbilir,2007; C� iftc� i & Bozkurt, 2008). However, in the absence ofsubsurface data, it may be di⁄cult to determine the originof these structures (Dorsey & Barajas, 1999).

Several mapped mesoscale folds are associated withstrike-slip faults and typically arranged in en echelon pat-terns oblique to the principal direction of shear. Miocenerocks are folded into en echelon NE- and NW-trendinganticlines and synclines that die out within approximately1km.Typically, en echelon folds are distributed in a rela-tively narrow zone adjacent to strike-slip faults.They mayhave formed in a zone between two strike-slip faults astheydo in theNpart of theKC� B,whereNE-trending foldslie at the E fault block of theKarabeldere fault (Fig. 4).Thefold trend suggests NW^SE compression, possibly linkedto growth of the Karabeldere fault. Localised domains ofextension and shortening may also re£ect the bending orstepping geometry of strike-slip faults, thus explainingthe presence of en echelon folds oblique to the zone of de-formation (e.g. Sylvester, 1988).

D3 phase

TheD3phase is characterised byNW^SE-striking range-front fault zones developed in an extensional tectonic re-

gime. The normal faults are dip- to oblique-slip and arewell exposed as short (0.5^10 km) and long (410 km) faultsegments.These structures are outlined as follows.

Kemalpasa fault

Miocene strata and strike-slip faults are cut by theKemal-paSa fault,which is the bounding structure of the southernKemalpaSa Basin and forms the SW end of the modernGediz Graben. It is a 1-km-wide and 15-km-long WNW-striking normal fault zone between the town ofKemalpaSato theWand the �renk˛y district to the E, where the othermajor fault forms the S margin of the Gediz Graben andhas a cross-cutting relationship with the KemalpaSa fault.There, the two overstepping major fault segments join analong-strike bend to form a breached relay ramp. Similarrelay ramps are reported along the S margin of the GedizGraben SE of Turgutlu (C� iftc� i & Bozkurt, 2007) and alongthe Manisa fault zone (Bozkurt & S˛zbilir, 2006; �zkay-mak & S˛zbilir, 2008). The KemalpaSa fault juxtaposesfootwall rocks of the Bornova Fysch Zone, basin- ¢ll unitsof the KC� B, and the Kızılca and Sˇtc� ˇler formationsagainst modern basin- ¢ll sediments in theNdownthrownhanging wall.The major fault segments display triangularfacets and locally preserved slickensided fault planes thatshow the KemalpaSa fault zone as an oblique-slip normalfault dipping at an average of 701N, with minor dextralstrike-slip o¡set.

Spildag� i fault zone

This is a 2-km-wide and 12-km-long NW-striking faultzone, which separates the Spildag� ı High from the Kemal-paSa plain (Fig. 2). Older basement rocks constituting anupthrown S footwall and Quaternary graben- ¢ll sedi-ments of the N downthrown hanging wall are tectonicallyjuxtaposed along major fault segments. The fault zoneconsists of four fault sets bounding the graben and dis-playing a step-like normal-fault pattern. Some of thesefault sets display well-preserved slickensides, with stereo-graphic plots showing oblique-slip normal o¡sets dippingat an average of 481SW with a minor sinistral strike-slipcomponent.

PALAEOSTRESS ANALYSIS

Syn- to post-sedimentary structures of the KC� B preserveevidence for various deformation phases (Fig.11and Table3). Major episodes include (1) an early to late Miocenetranstensional phase forming the basin; (2) an early Plio-cene wrench- to extension-dominated phase causing up-lift, deformation, and erosion of the basin ¢ll units; and (3)a Plio-Quaternary extensional phase forming the modernNW^SE trending depressions.

At the Emargin of the basin, the earliest faulting (D1) isrepresented by two families of faults, including NE-strik-ing dextral strike-slip faults and low-angle normal faults.D1faults were documented along theMFZ. Inverse analy-



sis results of fault-slip measurements for strike-slip fault-ing de¢ne steeply plunging s2 axes (711) but gently plun-ging s1 and s3 axes (18 and 081). The results suggest thatstrike-slip faulting developed under the approximatelyN^S extension is associated with E^Wcompression (Fig.11). Fault-slip measurements for low-angle normal fault-ing de¢ne a near-horizontal s3, trending approximately2231, and a151 plunge, whereas thes1 ands2 axes have at-titudes of approximately 080/721 and 316/111, respectively.The results suggest NE-SWextension (Table 3). D1 faultsare kinematically congruent with the NE^SW extensionassociated with the opening of supradetachment basinsduring the early Miocene, an interpretation supported bythes3 orientation calculated for D1strike-slip faults.

Later faulting (D2) is characterised by E-striking dip-slip and NE-striking dextral strike-slip faults, consistentwith the approximately NE^SWextension associatedwithNW^SE compression.These tensors were collected fromstations along the Derek˛y and Karaot faults. Fault- slipmeasurements along the Derek˛y fault de¢ne a relativelysteeply plunging s1 axis (781) but a horizontal s2 axis(041).The orientation of thes3 axis is similar for both lo-calities, with attitudes of 225/121 and199/321, respectively.Data from strike-slip faults in the footwall of the Derek˛yfault de¢ne an approximately vertical s2, plunging 771,whereas thes1ands3 axes are almost horizontal, plunging06 and 121, and trending 152 and 2431, respectively. Theresults suggest NW^SE compression and associatedNE^SWextension.During this stage, basin- ¢ll units weredeformed into mesoscale anticlines and synclines.

D2 also produced approximately N^S-striking normalfaults. Some D2 strike-slip faults were also reactivated, asshown by strike-slip slickensides overprinted by dip-slipslickenlines (Fig. 10d). In the outcrop, two stress sites arecharacterised by tensors with a subvertical s1 consistentwith the approximately ENE^WSW extension, and sub-horizontal s2 and s3 axes (Fig. 11 and Table 3).These ten-sors were collected from stations along the Kalkancatepeand Kocac� aydere faults. Fault-slip measurements alongthe Kalkancatepe fault de¢ne steeply plunging s1 axes(741) but horizontal s2 axes (041). The attitude of the s3

axis is 095/161. Along the strike of the Kocac� aydere fault,the computed results show remarkable similarities and de-¢ne s1 plunging 721. The s2 and s3 axes are plunging 03and 181, and trending 156 and 2471, respectively. Fault- slipdata collected from theKesmedag� ı fault de¢ne an approxi-mately verticals2, plunging761,whereas thes1ands3 axesare almost horizontal, plunging 04 and 131, and trending222 and 3131, respectively.

Normal faults belonging to the youngest generation ofstructures (D3) formed the modern KemalpaSa Basin tothe N and the Kˇc� ˇk Menderes Graben to the S. Com-puted results for bounding faults of the KemalpaSa Basinde¢ne an approximately vertical s1 axis (139/751), whereass2 (281/121) ands3 (013/091) are almost horizontal (Fig.11and Table 3).

We note that the extension direction based on structuraldata (including the paleostress analysis) does not change,

but is almost horizontal and NE^SW in all three stages(Fig.11). However, the axes of intermediate and maximumshortening switched position in the vertical plane duringD2 stage.To explain the complex deformational pattern ofbasin ¢ll units, we prepared a synoptic diagram of defor-mational structures developed during early Pliocenewrench- to extension-dominated tectonics with the maxi-mum shortening axis oriented in horizontal (1) and verti-cal (2) position, respectively (Fig. 12f). The mappedconjugate strike-slip faults and some folds may result fromthe horizontal shortening component of wrench-dominated deformation. However, some folds initiallyformed perpendicular to the axis of shortening may bemodi¢ed due to vertical-axis rotation. Field evidence alsosuggests that older strike-slip faults may be reactivated asoblique-slip normal faults due to a switch fromwrench- toextension-dominated transtension. Thus, some earlystructures may be modi¢ed during subsequent shearingalong the main fault zone.

GEOLOGIC EVOLUTION OF KC� B

A regional tectonic model is discussedwhich suggests thatdeformation was controlled by a combination of right- toleft-lateral shear at theW termination of the core complexformation, a change in local strain axes, and an inheritedzone of weakness.We also report new structural data thatshow reactivation of a pre-existing zone of weakness, theIBTZ in which the KC� B is located. The tectonic historyof the IBTZ can be traced back to the Late Cretaceous,when a transform fault between the Neotethys Ocean andthe Anatolide-Tauride carbonate platform was active (Fig.12a). The transform fault led to the deposition of £ysch-like sediments (the Bornova Flysch Zone) along the faultzone (Okay et al., 1996) where a thick sequence of earlyMiddle Eocene nummulitic limestone stratigraphicallyabove the Bornova Flysch Zone suggests that a relativelytectonically stable period existed during the Eocene (S˛z-bilir, 2005).

Pre-late-Eocene rockswithin theAegean region experi-enced a lateMesozoic-larlyTertiary Alpine compressionalhistory.This has been related to closure of the Neotethys,which is believed to have been sutured along the Izmir-Ankara and Vardar ophiolite belts in Turkey and Greece,respectively. Compression caused folding, thrusting, andcrustal thickening, which was accompanied by the forma-tion of subduction-related metamorphism and caused S-verging thrusting and nappe stacking in the Anatolide-Tauride block. Evidence for the earliest Tertiary event inthe region corresponds to the top-to-the-NE ductileshearing within the gneissic nappe pile of the MenderesMassif, interpreted as having re£ected late Eocene crustalthickening and nappe stacking (Lips etal., 2001). Intracon-tinental N^S convergence associatedwith the Palaeocene-Eocene collision along the Izmir-Ankara suture zone con-tinued until the Oligocene.

At the beginning of the Oligocene, a NE^SW-trendingLycian molasse basinwas formed at the tectonic boundary


H. S˛zbiliret al.

Fig.12. Schematic representation of tectonic evolution of theKocac� ayBasin. (a)Campanian: formation of BornovaFlyschZone along aNE-striking crustal- scale transform fault that formed during N^S contraction. (b) Early to middleMiocene: kinematically coupledmovement alongMahmutdag� ı fault zone andGediz detachment fault system that led to formation of E^W-trending supradetachmentbasin (Gediz Graben) andNE-trending transtensional supradetachment basin (Kocac� ay Basin) within a margin-parallel, right-lateraltransfer fault zone. (c)Middle to lateMiocene: widening of basin and deposition of volcanosedimentary basin ¢ll under transtensionalconditions, (d) Early Pliocene: deformation of basin ¢ll by variously oriented folds, and normal and reverse faulting associatedwithstrike-slip faults under wrench- to extension-dominated conditions, (e) Plio -Quaternary: formation ofNW-striking normal faults thatcut anddisplaceMiocene units and related structures. (f) Synoptic diagram of deformational structures developed during earlyPliocenewrench- to extension-dominated tectonics with the maximum shortening axis oriented in the horizontal (1) and vertical (2) positions,respectively (secondary structures relative to principal strain axes after Sanderson&Marchini,1984;Sylvester,1988;DePaola etal., 2008).



between theMCC and theLycianNappes during theMen-deres orogenic collapse, thus initiating the ¢rst exten-sional phase (S˛zbilir, 2005; ten Veen et al., 2009).Evidence for the presence ofOligocene-lowerMiocene ex-tensional basins includes (1) the top-to-the-NE shearsense indicators in the Menderes Massif and (2) a seriesof NE-trending Oligocene extensional basins containinghigh-grade metamorphic rock fragments of the MCC(Seyitog� lu et al., 2004; S˛zbilir, 2005; tenVeen et al., 2009).In response to NW^SE compression, the Lycian molassebasin subsequently underwent deformation at the begin-ning of the Miocene, while western Anatolia was charac-terised by a core complex mode of extension duringwhich supradetachment basins were formed and exten-sion-related deformation caused pervasive mylonitisationof the footwall, with top-to-the-NNE ductile and thenbrittle shear (Fig. 12b).The main episode of extension oc-curred from about 20 to 7Ma (Lips etal., 2001); it has beenviewed in terms of volcanism, sedimentation, detachmentfaulting, and strike-slip faulting accommodating N^S ex-tension.

From the early to middle Miocene, crustal thinning inthe centralMenderesMassifwas associatedwith the denu-dation of MCCs in the footwalls of the Gediz and Bˇyˇk-Menderes detachment faults (Emre & S˛zbilir, 1997; Lipset al., 2001).These include mylonitised, metamorphic, andgranitic rocks lying below a low-angle detachment fault,with associated chlorite brecciation and a supradetach-ment basin containing a thick succession of nonmarinestrata (Hetzel et al., 1995; Emre & S˛zbilir, 1997; Koc� yig� itet al., 1999; S˛zbilir, 2001, 2002; Seyitog� lu et al., 2002; ISıket al., 2003, 2004; Bozkurt & S˛zbilir, 2004). Terrigenousand coal-bearing rocks of early to middle Miocene lie inlow-angle fault contact on older units ranging in age fromPalaeozoic through Cretaceous. The footwall meta-morphic rocks were progressively mylonitised, exhumed,and intruded by syndeformational granitoids (Turgutluand Salihli granodiorites: the Salihli granodiorite yielded39Ar-40Ar amphibole isochron and biotite plateau coolingages of19.5^1.4 and12.2^0.4Ma, respectively) (e.g.Hetzel etal., 1995; Emre, 1996; Koc� yig� it et al., 1999; Gessner et al.,2001; Lips et al., 2001; S˛zbilir, 2001, 2002; Seyitog� luet al., 2002; ISık et al., 2003, 2004; Bozkurt & S˛zbilir,2004). A sense of brittle shear, deduced from fault-planestructures, mesoscopic o¡sets, and brittle and semi-brittle shear bands, consistently has attitudes of top-to-the-NNE. This suggests kinematic continuity fromductile to brittle conditions during exhumation of footwallrocks of the detachment fault system (Lips et al., 2001;S˛zbilir, 2001, 2002; ISık et al., 2003).

At the beginning of theMiocene, theKC� B,which formsthe E margin of the IBTZ, underwent NE^SW transten-sion that was characterised by kinematically linked low-angle normal faults and strike-slip faults. Maximum sub-sidence occurred along the E margin of the basin, wheresynsedimentary folding and syntectonic unconformitieshave been observed, forming an asymmetric depression.From early to middleMiocene times, aNE-trending base-

ment high separated the basin of deposition into two long-itudinal depressions ^ the KC� B to the E and theCumaovası basin to theW (Fig.12c).This type of basementhigh may be related to an extension-parallel footwall cor-rugation of the detachment fault (for details of origins ofextensional folds, see Janecke et al., 1998). During theNeo-gene, the basin of deposition mayhave served as a tempor-ary marine connection with the Aegean Sea.

Following the tectonic activity in the late Miocene, theregion was deformed under the di¡erent tectonic forcesthat arose after sedimentation of the Kızılca Formation,thus postdating the region-wide late Miocene lacustrinesedimentation. During the early Pliocene, the transten-sional supradetachment basin was uplifted and deformedby strike-slip faults, reverse faults, and folds (Fig. 12d).The early Pliocene phase was associated with a shift fromwrench- to extension-dominated deformation.The early-middle Miocene basin- ¢ll deposits are limited by NE-striking the Kesmedag� ı fault zone (Kef) to theW. Some ofthe E-striking normal faults end before or terminateagainst the Kef. Connecting normal faults have strikes of50^801 from the Kef.They may be explained by a soft- orhard-linked transtensional relay ramp that is breachedthrough a set of normal faults striking at a high angle tothe main strike-slip fault (Fodor, 2007).This suggests thatthe NE-striking strike-slip fault and E-striking normalfaults were active at the same time during the early Plio-cene faulting phase.

The ¢nal tectonic phase (Fig. 12e) produced pull-apartbasins in the late Pliocene, which resulted from the ap-proximately NE^SW extension (Uzel & S˛zbilir, 2008).This phase occurred as a response to initial westward tec-tonic escape of Anatolia along the North Anatolian Faultand East Anatolian Fault (Dewey & Reng˛r, 1979; Reng˛ret al., 1985; Reng˛r, 1987). During this stage, theKemalpaSafault and Spildag� ı fault zone acted together to form themodern Gediz Graben (Fig. 12e). On the basis of recentseismotectonic studies (Akyol et al., 2006; Zhu et al., 2006),the tension direction is also determined to be NE^SW.

DISCUSSION AND CONCLUSION

The KC� B exhibits many of the typical features of a trans-tensional supradetachment basin as described in NWMexico (Dorsey & Barajas, 1999). These features include(1) rapid basinward facies changes, (2) low-angle normalfaults, (3) strike-slip faults, (4) intraformational unconfor-mities, and (5) an asymmetric depocentre. The stratalthickness of the KC� B is asymmetric, being thickest(1300m) along the E margin of the basin, as observed inmany transtensional and extensional basins with a half-graben geometry inwhich the throw of one bounding faultis much greater than that of the other (Gawthorpe&Colel-la, 1990; Janecke et al., 1998; Umhoefer et al., 2007).

The studied basin is bounded to theW by the BornovaFlysch Zone and to the E by metamorphic rocks of theMenderes/Cycladic Massif. One of the most distinctive


H. S˛zbiliret al.

characteristics of the basin is a sedimentation that is syn-chronous with tectonism. Evidence includes the presenceof intraformational angular unconformities and syndepo-sitional structures.

Sedimentologic, stratigraphic, and structural evidencefrom the sedimentary succession suggests that basin de-velopment was kinematically related to movement alongNE-striking strike-slip and E-striking low-angle normalfaults. Major movement along the nearby Gediz detach-ment also occurred at this time.We suggest that spatiallyand temporally related basin growth and sedimentationre£ect kinematically coupled movement along a combinedMFZ and Gediz detachment fault system. The impor-tance of the termination of low-angle normal fault systemswithin a transfer zone has already been recognised in thewestern Anatolia extensional province (Ring et al., 1999;S˛zbilir et al., 2003).

The IBTZ represents a major transverse strike-slip-dominated zone that accommodated the lateral termina-tion of E-striking graben-faults and linked spatiallydiscrete loci of extension.This zone might be classi¢ed asa transverse or replacement structure, de¢ned by Reng˛ret al. (1985) as a structure coincident in space with palaeo-tectonic structures but that which does not perform thesame function during a later tectonic episode. As statedby Duebendorfer & Black (1992),‘transverse structure’ is anongenetic term that applies to any zone of discontinuitythat is oriented nearly parallel to the extensional direction.Faults within the IBTZ may have had their earliest recordin the Late Cretaceous, related to an existing transformzone perpendicular to the Neotethys.

Two di¡erent views address the late Cenozoic stress-re-gime history ofwestern Anatolia: (1) uninterrupted exten-sional tectonics with an approximately N-trending s3

(Dewey & Reng˛r, 1979; Seyitog� lu & Scott, 1991, 1992; Ar-mijo et al., 1996; Seyitog� lu et al., 2002), and (2) N^S exten-sional tectonics interrupted by punctuated or recurrentcompressional stresses (Angelier et al., 1981; Dumont et al.,1981; Mercier, 1981; Angelier, 1984; Koc� yig� it et al., 1999;Bozkurt & S˛zbilir, 2004, 2006; Kaya et al., 2004, 2007;S˛zbilir, 2005; van Hinsbergen & Meulenkamp, 2006). Inthis study, we consider the possibility of a strain historythat changedwith time. A further hypothesis suggests thatan inherited zone of weakness played an important role inde¢ning fault geometry.

The KC� B underwent several phases of Neogene toQuaternary faulting, resulting in a dense network of faultsofvariable orientation.The basin is interpreted to have de-veloped in response to transtension (D1) within a margin-parallel zone of right-lateral transfer faulting. During theearly Miocene, this zone may have provided a temporarymarine connection with the Aegean Sea. The next phaseof deformation involved early Pliocenewrench-dominatedtectonism (D2) associatedwith a shift fromwrench- to ex-tension-dominated deformation. At the same time, theGediz detachment ceased activity coevally with develop-ment of the North Anatolian Fault Zone.The terminationof detachment faulting and rise of the middle crust may

have changed extensional and transtensional regions towrench-dominated zones partly because of the presenceof a basement block that caused lateral rheological changesin the upper crust.This phase resulted in deformation anderosion of the uplifted basin- ¢ll units.

At least two ages of deformational structures can be re-cognised, particularly in the N part of the basin. Earlierformed faults comprise strike-slip faults with slip vectorsraking 0^251. Later deformational structures comprisingnormal faults are commonly observed to reactivate the ear-lier formed faults. This can be explained by progressivedeformation due to a localised switch fromwrench- to ex-tension-dominated tectonics. Even though the regionalstress ¢eld has been characterised by a subhorizontalminimum principal stress axis, the IBTZ probably in-duced di¡erences in the local stress ¢eld.Within the faultzone, palaeostress analyses point to a local stress ¢eldwithNW^SE subhorizontal compression andNE^SWsubhor-izontal tension. With increasing strain intensity, the for-mer wrench-dominated con¢guration changed toextension-dominated deformation where the axis of max-imum extension remained horizontal, but the axis of max-imum shortening changedwith time (Fig.12f).Di¡erencesin the strain history have resulted in kinematically hetero-geneous and structurally complex fault associations ofstrike-slip faults and normal faults (Figs 4 and12). Subse-quently, the deformed basin ¢ll was unconformably over-lain by up to 100m of post-detachment deposits of Plio-Quaternary age.The base of the post-detachment succes-sion is an erosional unconformity that truncates thefolded and strike-slip-faulted basin ¢ll. The ¢nal post-detachment deformation included Plio-QuaternaryNE^SWextension.

ACKNOWLEDGEMENTS

This research was supported byTubitak National ScienceFoundationGrant Number102Y065.We thank Fuat Erkˇl,Yalc� ınErsoy, FundaAkgˇn, Sebla Ercan, andMedihaKılıc�for assistance in the ¢eld.The authors are greatly indebtedto P. Umhoefer, whose suggestions greatly improved themanuscript. We also are grateful to Associate Editor B.Horton for his comments and substantial improvementsto the manuscript. The paper was edited by two editingservices the International ScienceEditing, and theEditageo¡ered byWiley-Blackwell.

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10.1029/2006GL025842

10.1029/2006GL025842

Tectonic implications of transtensional supradetachment basin development in an extension-parallel transfer zone: the Kocacay Basin, western Anatolia,Turkey

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