Gondwana Research 24 (2013) 243–274

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GR Focus Review

What caused the denudation of the Menderes Massif: Review of crustal evolution,lithosphere structure, and dynamic topography in southwest Turkey

Klaus Gessner a,⁎, Luis A. Gallardo b, Vanessa Markwitz c, Uwe Ring d, Stuart N. Thomson e

a Western Australian Geothermal Centre of Excellence, and Centre for Exploration Targeting, The University of Western Australia, M006, 35 Stirling Highway, Crawley 6009, Australiab Earth Science Division, CICESE, Carretera Ensenada-Tijuana No. 3918, CP 22860, Ensenada, Mexicoc Centre for Exploration Targeting, The University of Western Australia, M006, 35 Stirling Highway, Crawley 6009, Australiad Department of Geological Sciences, Stockholm University, SE-106 91 Stockholm, Swedene Department of Geosciences, University of Arizona, Gould-Simpson Building, 1040 E. 4th St., Tucson, AZ 85721-0077, USA

⁎ Corresponding author at: Geological Survey of WesE-mail addresses: klaus.gessner@uwa.edu.au (K. Ges

thomson@email.arizona.edu (S.N. Thomson).

1342-937X/$ – see front matter © 2013 International Ahttp://dx.doi.org/10.1016/j.gr.2013.01.005

a b s t r a c t

a r t i c l e i n f o

Article history:Received 31 March 2012Received in revised form 28 January 2013Accepted 31 January 2013Available online 16 February 2013

Handling Editor: M. Santosh

Keywords:Metamorphic core complexContinental extensionTurkeyAegean SeaMenderes MassifLithosphere delaminationDynamic topography

The deformation of Earth's lithosphere in orogenic belts is largely forced externally by the sinking slab, butcan also be driven by internal delamination processes caused by mechanical instabilities. Here we presentan integrated analysis of geophysical and geological data to show how these processes can act contempora-neously and in close proximity to each other, along a lithosphere scale discontinuity that defines the lateralboundary between the Hellenide and Anatolide segments of the Tethyan orogen in western Turkey. TheHellenides and Anatolides have experienced similar rates of convergence, but display remarkable differencesin the structure of Earth's crust and lithospheric mantle across the Aegean coast of the Anatolian peninsula.We review the tectonics of southwest Turkey in the light of new and published data on crustal structure,cooling history, topography evolution, gravity, Moho topography, earthquake distribution and seismic to-mography. Geological data constrain that one of Earth's largest metamorphic core complexes, the MenderesMassif, experienced early Miocene tectonic denudation and surface uplift in the footwall of a north-directedextensional detachment system, followed by late Miocene to recent fragmentation by E–W and NW–SEtrending graben systems. Gravity data, earthquake locations and seismic velocity anomalies highlight anorth–south oriented boundary in the upper mantle between a fast slab below the Aegean and a slow as-thenospheric region below western Turkey. Based on the interpretation of geological and geophysical datawe propose that the tectonic denudation of the Menderes Massif and the delamination of its subcontinentallithospheric mantle reflect the late Oligocene/early Miocene onset of transtension along a lithosphere scaleshear zone, the West Anatolia Transfer Zone (WATZ). We argue that the WATZ localised along the boundaryof the Adriatic and Anatolian lithospheric domains in the Miocene, when southward rollback of the Aegeanslab started to affect the central Aegean–Menderes portion of the Tethyan orogen. Transtension across theWest Anatolia Transfer Zone affected the entire Menderes Massif in the Early Miocene. The current crustal ex-pression of this boundary is a NNE-trending, distributed brittle deformation zone that localised at the west-ern margin of the denuded massif. Here, sinistral transtension accommodates the continuing velocitydifference between relatively slow removal of lithospheric mantle below western Anatolia and trench retreatin the rapidly extending Aegean Sea region. Our review highlights the significance of lateral variations of thelower plate in subduction–collision systems for evolving structure and surface processes in orogenic belts,particularly in relation to the formation of continental plateaux and metamorphic core complexes.

© 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2442. Regional tectonic overview . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 245

2.1. Structure of the Hellenides in the Aegean Sea region . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2462.2. Structure of the Anatolides in western Turkey . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2462.3. Controversies on Alpine tectonics of the Menderes Massif . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249

tern Australia, Department of Mines and Petroleum, 100 Plain Street, East Perth, WA 6004, Australia.sner), lgallard@cicese.mx (L.A. Gallardo), vanessa.markwitz@uwa.edu.au (V. Markwitz), uwe.ring@geo.su.se (U. Ring),

ssociation for Gondwana Research. Published by Elsevier B.V. All rights reserved.

244 K. Gessner et al. / Gondwana Research 24 (2013) 243–274

2.3.1. Alpine crustal shortening and the age of deformation fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2492.3.2. Significance of the Selimiye shear zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2492.3.3. Stratigraphic position of low grade metasediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249

2.4. Miocene to recent extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2502.4.1. Extension of the Anatolide belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2502.4.2. Magmatic record of crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252

2.5. Controversies on crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2552.5.1. Fabric overprinting — extension or contraction? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2552.5.2. Exhumation of the Gördes submassif and the role of the Simav detachment . . . . . . . . . . . . . . . . . . . . . . . . 2552.5.3. Block rotation versus diffuse extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 255

3. Topographic response to crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2563.1. Methods and materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2563.2. Topographic profiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2563.3. Drainage channels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2563.4. Interpretation of topography and river channel data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257

4. Upper mantle structure and active deformation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2594.1. Geophysical evidence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 261

4.1.1. Gravity anomaly and Moho depth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2614.1.2. Earthquake hypocentres . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2614.1.3. 3D model of seismic tomography and earthquake hypocenters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 262

4.2. Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2635. Tectonic synthesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 264

5.1. Lateral differences in lithospheric structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2655.2. Sinistral transtension across West Anatolian Transfer Zone as a driver for Menderes extension . . . . . . . . . . . . . . . . . . . . . 2665.3. Continuous versus punctuated crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2685.4. Open questions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 269

6. Summary points . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 269Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 270Appendix A. Supplementary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 270References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 270

1. Introduction

Lithosphere architecture and strain distribution can vary substan-tially in orogenic belts, both across and along strike. Along-strike var-iations and structural complexity are common features of mountainbelts such as the European Alps (e.g. Schmid et al., 2004), the Andes(e.g. Allmendinger et al., 1997), the Himalayas (e.g. An, 2006) andthe Hellenide–Anatolide orogen in southeastern Europe (Ring et al.,1999a; Gessner et al., 2001c; Gessner et al., 2011; van Hinsbergenand Schmid, 2012). The causes for along-strike variations are likelyto differ in individual orogenic belts, but will generally be a conse-quence of compositional and architectural variations in the accretingor colliding continental lithosphere fragments. Along-strike varia-tions, however, depend not only on the composition and architectureof these fragments, but also on the differential dynamics generated bythe sinking slab, and by mechanical instabilities that affect the accre-tion of continental arcs even at distances far from the actual tectonicmargin, and shape the geology, topography and the lithosphere struc-ture sensed by geophysical data. It has been recognised that through-out the Earth's history tectonic and magmatic accretion of continentalarcs not only have played an important role in the growth of conti-nents (Rudnick, 1995), but also as regions of long-lived thermallyweakened mobile belts (Hyndman et al., 2005). Conceptual and nu-merical models of generic and regionally specific continental arcssuggest that deformation is not only mainly driven by external forc-ing by the sinking slab (Royden, 1993; Collins, 2002; Schellart et al.,2007; Spakman and Hall, 2010), but also internally, by gravitationalinstabilities within thermally weakened lithosphere (Houseman etal., 1981; England and Houseman, 1989; Molnar et al., 1993; Plattand England, 1993; Houseman and Molnar, 1997; Stern et al., 2006),with mechanical and thermal coupling across the subduction zonedetermining how these processes interact (Faccenda et al., 2009).The significance of considering ‘internal drivers’ such as gravitational

instabilities in addition to ‘external drivers’ such as sinking slabs, isthat synchronous contraction and extension can be accommodatedin the Earth's crust over relatively short across-strike distances(Gögüs and Pysklywec, 2008; Faccenda et al., 2009). Such internaldriving processes pose a challenge to the existence of regional orfar-field force continua across orogens, an assumption that is oftenmade a priori when deformation fabrics are linked with geodynamicprocesses in ancient orogenic belts. The partition of deformationalong active continental collision zones such as the Tethyan orogenin the Eastern Mediterranean provide a natural laboratory wherethe recent and current evolution of geological structures can be stud-ied and interpreted in the context of surface processes, gravity anom-alies, seismicity, geodetic measurements, and mantle tomography. Inthe Eastern Mediterranean the southward rollback of the Hellenicsubduction zone and the westward motion of Anatolia dominate thekinematics of continental plate fragments as they occur at muchhigher rates than the convergence between Africa and Eurasia(e.g. Reilinger et al., 2006; Pérouse et al., 2012) (Fig. 1). This study fo-cuses on southwest Turkey, where the westward movement of Ana-tolia changes to the southward movement of the Aegean, where theAnatolian plateau gives way to the Aegean Sea, and where theHellenide and Anatolide segments of the Tethyan orogen meet. Wedescribe the regional structure across the Hellenide–Anatolide transi-tion and, in the light of new and published apatite fission track data,discuss the tectonic models put forward for the Menderes Massif, par-ticularly with regard to key structures like the Simav detachment andthe Selimiye shear zone. We then use the structure of the Alpinenappe stack as a marker to track the deformation imposed on westernAnatolia by the late Miocene to recent extension, as evidenced by to-pography and drainage channel morphology. Using geophysical datasuch as gravity, seismic velocity anomaly, and earthquake hypocentrelocations we show how the geological along-strike-differences be-tween the Hellenic and the Anatolide crustal domains relate to the

Fig. 1. Kinematic configuration and geodetic measurements of continental fragmentsin the Eastern Mediterranean, Arabia and the Caucasus relative to a fixed Eurasia; no-tice the relatively high velocities of Anatolia and the Aegean — driven by suction of theAegean slab — relative to the convergence between Nubia/Arabia and Eurasia — drivenby slab pull, and the uncertainty of how the difference in movement direction betweenAnatolia and the Aegean is accommodated.Modified from Reilinger et al. (2006).

245K. Gessner et al. / Gondwana Research 24 (2013) 243–274

upper mantle structure below western Turkey. Finally, we synthesiseour evidence to discuss the lateral differences in lithosphere structureas the driver of the Menderes extension and in the geodynamics of

Fig. 2. Simplified tectonic overview of the Aegean Sea region, with Adriatic plate units in blueoceanic units in green. The Pindos unit, including widespread high-pressure metamorphicBlueschist unit overlies the Menderes Massif, which lacks Alpine high-pressure metamorph

the Eastern Mediterranean. We propose that a lithospheric scaletransfer zone, the West Anatolia Transfer Zone (WATZ) defines thelateral boundary between the Hellenide and Anatolide orogens,where slab rollback in the Aegean and delamination of the lithospher-ic mantle in western Anatolia have operated contemporaneously andin close proximity to each other; causing tectonic denudation of thelower crust in the Aegean and in the Menderes Massif.

2. Regional tectonic overview

The Hellenide orogen of Greece and the Anatolide belt of westernTurkey form an arcuate orogenic belt north of the Hellenic subductionzone (Fig. 2). Both the Hellenides and the Anatolides consist ofstacked tectonic units that are overlain by the Late Cretaceous toPaleogene Vardar–İzmir–Ankara suture zone to the north. TheAdriatic plate has played a key role in the tectonic development ofthe Mediterranean region. It has rifted from the northern margin ofGondwana in the Cretaceous and still moves independently of theEurasian Plate. In the eastern Mediterranean little is known aboutthe Mesozoic to early Tertiary paleogeography of the Adriatic plate,which appears to pinch out eastwards. In the Mesozoic, continentalcrust of the Adriatic plate as exposed today on the Attic Peninsulaand in the Aegean, varied between normal thickness, highly stretchedand thinned; and locally may have been oceanic (Jacobshagen, 1986;Robertson et al., 1991). The continental fragment directly east of theAdriatic plate was termed Anatolide–Tauride platform (Sengör andYilmaz, 1981), or — following Gessner et al. (2001c) — also as Anato-lia. Tectonic units within the Hellenide–Anatolide orogen are aligned

, Anatolian plate units in pink, Eurasian plate units in brown, and Vardar–İzmir–Ankararocks, overlies the External Hellenides in the west. In the east the equivalent Cycladicism. The box shows the extent of Figs. 4, 15, 16, and 17.

246 K. Gessner et al. / Gondwana Research 24 (2013) 243–274

parallel to the present-day Hellenic subduction zone. Once believed tobe a ‘Median Crystalline Belt’ (Dürr et al., 1978), the metamorphicrocks of the Pelagonian zone (Aubouin, 1959), the Cycladic zone, andthe Menderes Massif (Paréjas, 1940; Brinkmann, 1971) are nowknown to be different tectonic units (Robertson and Dixon, 1984;Erdogan and Güngör, 1992; Robertson et al., 1996; Ring et al., 1999b;Gessner et al., 2001c; Jolivet and Brun, 2010; Ring et al., 2010; vanHinsbergen and Schmid, 2012). Rather than representing an eastern ex-tension of the Carboniferous basement and Permo-Mesozoic cover ofthe Adriatic plate, the Anatolide belt is made up of two different unitsthat do not share the same Alpine tectono-metamorphic history, theCycladic Blueschist unit and the underlying Menderes nappes (Ringet al., 1999a; Gessner et al., 2001a; Gessner et al., 2001c; Regnier et al.,2003). In theMenderes nappes, pronouncedmagmatic activity occurredat the Proterozoic/Cambrian boundary (Hetzel and Reischmann, 1996;Dannat and Reischmann, 1999; Gessner et al., 2001a; Reischmann andLoos, 2001; Zlatkin et al., 2012), the mid-Triassic (Dannat, 1997;Koralay et al., 2001) and the Miocene (Hetzel et al., 1995b; Seyitogluand Scott, 1996; Isik and Tekeli, 2001; Ring and Collins, 2005; Glodnyand Hetzel, 2007; Ersoy et al., 2008; Akay, 2009; Dilek andAltunkaynak, 2009; Hasozbek et al., 2010; Prelevic et al., 2010a;Prelevic et al., 2010b; Hasozbek et al., 2011; Öner and Dilek, 2011;Altunkaynak et al., 2012a, 2012b; Catlos et al., 2012; Hasozbek et al.,2012). In the Cycladic zone, the granitic basement is of Carboniferousage (Reischmann, 1997; Engel and Reischmann, 1998). In addition,there are Triassic intrusions (Reischmann, 1997; Ring et al., 1999b) andprominent Miocene to recent magmatic activity (Altherr et al., 1982).

2.1. Structure of the Hellenides in the Aegean Sea region

The Hellenides consist of five tectonic units. These are from top(north) to bottom (south), (1) the Eurasian plate units, such as for ex-ample the Serbo-Macedonian Block, (2) the Vardar–İzmir–AnkaraOceanic units, (3) the Pelagonian Zone, (4) the Pindos Unit (includingthe Cycladic Blueschist Unit), (5) the External Hellenides, includingthe Gavrovo–Tripolitza Block and the underlying Ionian Block, and(6) the Mediterranean Ridge Accretionary Complex (Fig. 2). Ofthese, only the top three units can be correlated across from theHellenides to the Anatolides in western Turkey (Ring et al., 1999a).The Pindos unit is a subduction complex that formed between ca.55 Ma and 30 Ma (Ring and Layer, 2003; Jolivet and Brun, 2010;Ring et al., 2010) and comprises normal-thickness continentalbasement–cover sequences, as well as thick radiolarite sequences in-dicating that locally it was underlain by oceanic crust, or by thinnedcontinental crust (Pe-Piper and Piper, 1984; Robertson et al., 1991).In the Cyclades, the uppermost unit of the Pindos Unit is the highly at-tenuated ophiolitic Selçuk Mélange (Okrusch and Bröcker, 1990; Ringet al., 1999b; Katzir et al., 2000), which forms the upper part of theCycladic Blueschist Unit. The lower part of the Cycladic BlueschistUnit comprises Carboniferous schist and orthogneiss, and a late- topost-Carboniferous passive-margin sequence of marble, metapeliteand volcanics (Dürr et al., 1978).

The Gavrovo–Tripolitza Block is a continental platform unit of Trias-sic to Eocene age, partly overlain by late Eocene to early Oligocene turbi-dites (Jacobshagen, 1986). Subduction of the Gavrovo–Tripolitza Blockcommenced at ca. 35–30 Ma (Thomson et al., 1998; Sotiropoulos andKamberis, 2003). In the Cyclades, high-pressure rocks of the Gavrovo–Tripolitza Block that are locally exposed in tectonic windows belowthe Cycladic Blueschist Unit are usually referred to as the Basal Unit(Godfriaux, 1968; Shaked et al., 2000; Ring et al., 2001a). In the Pelopon-nese and in Crete, the rocks of the Gavrovo–Tripolitza Block and thePindos Unit are only weakly metamorphosed. The Ionian block com-prises late Carboniferous to possibly Triassic rocks overlain by limestoneand late Eocene to Miocene turbidites (Jacobshagen, 1986). Rocks ofboth the Gavrovo–Tripolitza and Ionian blocks do not crop out in west-ern Turkey. Tectonic units in the footwall of the PelagonianUnit lack any

Cretaceous orogenic history, and were metamorphosed to high pres-sures at least 20 Ma later than the Pelagonian Unit and the Vardar–İzmir–Ankara Oceanic Units (Ring et al., 2010). The most outboard tec-tonic domain of the Hellenides is theMediterranean Ridge AccretionaryComplex (Fig. 2) (Kopf et al., 2003). Along the central MediterraneanRidge, East Mediterranean oceanic crust has been subducted and theleading edge of the African passive continental margin is currently en-tering the subduction zone.

The subduction of the Vardar–İzmir–Ankara Ocean that fringedAdria and Anatolia on its northern sides caused high-pressure meta-morphism in these oceanic units in the Late Cretaceous (Sherlock etal., 1999). Beginning in the Early Tertiary, the northern edge of thePindos Unit was underthrust causing high-pressure metamorphism inlarge parts of the Cycladic Blueschist Unit in the central Aegean Sea re-gion (Cyclades islands) and westernmost Anatolia. Well-constrainedages for high-pressure metamorphism range from ca. 53 Ma to 30 Ma(Ring and Layer, 2003; Tomaschek et al., 2003; Putlitz et al., 2005;Ring et al., 2007b). High-pressure metamorphism took place in the Ex-ternal Hellenides in Crete and the Peloponnesus in the latest Oligoceneto Miocene, at ca. 25–20 Ma (Seidel et al., 1982; Jolivet et al., 1996).

Recent reviews (Jolivet and Brun, 2010; Ring et al., 2010; Jolivet etal., in press) demonstrated the progression of high-pressure thatmetamorphism gets younger towards the south. Major along-strikevariations in the Hellenide–Anatolide orogen therefore should be re-lated to the arrival of Anatolia in the eastern Mediterranean subduc-tion systems in the Eocene (Gessner et al., 2001c). During incipientunderthrusting of the leading edge of Anatolia the high-pressuremetamorphosed Cycladic Blueschist Unit was thrust onto the Mende-res nappes between 42 Ma and 32 Ma (Ring et al., 2007a).

2.2. Structure of the Anatolides in western Turkey

In the Anatolide belt ofwestern Turkey the Pindos Unit (representedby the Cycladic Blueschist unit) overlies theMenderes nappes (Figs. 3, 4,and 5)—which are part of Anatolia—whereas in the Aegean region thePindos Unit overlies the Basal Unit — which is part of the ExternalHellenides (Gavrovo–Tripolitza) (Dürr, 1975; Robertson et al., 1991;van Hinsbergen et al., 2005). The Vardar–İzmir–Ankara Oceanic unitscontain Triassic to Eocene remnants of Neothethys which were sub-ducted below Sakarya since the Cretaceous (Okay and Tüysüz, 1999;Okay, 2011). In western Turkey, Cretaceous to Palaeogene subduc-tion–accretion complexes constitute the footwall of the Vardar–İzmir–Ankara suture, including the Tavşanlı zone, and the Bornova Flyschzone (Okay and Tüysüz, 1999; Okay, 2011). The Ören/Afyon zone andthe Lycian nappes (Okay and Tüysüz, 1999; Pourteau et al., 2010) occurstructurally below the ophiolitic parts of the Vardar–İzmir–Ankara Oce-anic units, parts of which may constitute remains of a separate, continu-ous Anatolian ophiolite nappe (Okay, 2010). The Tavşanlı zone, and theAfyon/Ören units were metamorphosed under blueschist-facies condi-tions in the Late Cretaceous and Palaeocene (Sherlock et al., 1999;Rimmelé et al., 2003; Pourteau et al., 2010), and overlie the Pindos unit,represented by the Cycladic Blueschist Unit. In western Turkey, the Cy-cladic Blueschist occurs above the Menderes Nappes, separated by theCyclades–Menderes Thrust (Fig. 4) (Gessner et al., 2001c).

We follow the tectonic division of the Menderes Massif as an Al-pine nappe stack (Gessner et al., 1998; Partzsch et al., 1998; Ring etal., 1999a; Gessner et al., 2001c; Regnier et al., 2003) consistingof, from top to bottom (1) the Selimiye Nappe, (2) the Cine Nappe,(3) the Bozdağ Nappe, and (4) the Bayındır Nappe (Fig. 4). The Çineand Bozdağ nappes have a polyorogenic history, which extends backinto the Neoproterozoic/Cambrian (Kröner and Sengör, 1990; Hetzeland Reischmann, 1996; Candan et al., 2001; Gessner et al., 2001a;Gessner et al., 2004; Ring et al., 2004; Catlos and Çemen, 2005; Ringand Collins, 2005; Oberhänsli et al., 2010; Candan et al., 2011;Zlatkin et al., 2012).

Fig. 3. Schematic architecture of tectonic units in the Aegean Sea region and western Anatolia.

247K. Gessner et al. / Gondwana Research 24 (2013) 243–274

According to this subdivision the structurally lowest unit of theMenderes nappes, the Bayındır nappe, has only been affected by onemajor Alpine tectonometamorphic event, whereas in the overlyingBozdağ, Çine and Selimiye nappes pre-Alpine and Alpine events aredocumented. The Cyclades–Menderes thrust cuts through severalnappes of the underlying Menderes nappe stack. Deformation/meta-morphism relations across the Cyclades–Menderes thrust indicate thatthe breakdown of garnet and biotite to chlorite in the Bozdağ nappeat temperatures below ca. 400 °C occurred during mylonitization. Ac-cordingly, the Cyclades–Menderes thrust has been interpreted as alate Alpine out-of-sequence thrust (Gessner et al., 2001c).

The structurally highest tectonic unit, the Selimiye Nappe, con-tains Palaeozoic metapelite, metabasite and marble (Schuiling, 1962;Çaglayan et al., 1980; Loos and Reischmann, 1999; Regnier et al.,2003; Gessner et al., 2004). The Eocene Selimiye Shear Zone separatesthe Selimiye Nappe from the underlying Çine Nappe (Fig. 4) (Bozkurtand Park, 1994; Bozkurt and Park, 1997; Gessner et al., 2004). Most ofthe Çine nappe consists of deformed orthogneiss, largely undeformedmetagranite andminor pelitic gneiss, eclogite and amphibolite. Protoliths

Fig. 4. Interpretative thrust sequence during formation of Anatolide belt. Notice that the Cycthe Menderes nappes.After Gessner et al. (2011). Age data refer to Lips (1998)†, Loos and Reischmann (1999)††,

of much of the orthogneiss/metagranite intruded at ca. 560–530 Ma(Hetzel and Reischmann, 1996; Hetzel et al., 1998; Loos andReischmann, 1999; Gessner et al., 2001a, 2004; Zlatkin et al., 2012). Theunderlying Bozdağ Nappe is made up of metapelite containing amphibo-lite, eclogite andmarble lenses. Protolith ages of the BozdağNappemeta-morphics are unknown, but geological constraints (Candan et al., 2001;Gessner et al., 2001a, 2004) suggest a Precambrian age. Like the ÇineNappe, the Bozdağ Nappe was intruded by granitoids at ca. 240–230 Ma (Dannat and Reischmann, 1999; Koralay et al., 2001). TheBayındır Nappe contains phyllite, quartzite, marble and greenschist of in-ferred Permo-Carboniferous to Mesozoic age (Özer and Sozbilir, 2003)that were affected by a single Eocene greenschist-facies metamorphism(Lips et al., 2001; Catlos and Çemen, 2005; Cemen et al., 2006). TheBayındır nappe was deformed by the first common deformation eventrecorded in the Menderes Massif and the Cycladic blueschist unit(Gessner et al., 2001c). The corresponding foliation is associated with afine-grained N-trending stretching lineation associated with ductileshear bands and sigma-type objects indicating a top-to-the-S shearsense (Gessner et al., 2001c).

lades–Menders Thrust emplaces units with a high-pressure accretion history on top of

and Gessner et al. (2001c)†††.

Fig. 5. Exposed below the Vardar–İzmir–Ankara Ocean suture and overlying high-pressure metamorphic units, the Menderes Massif is the structurally lowest part of the Tethyanorogen in western Anatolia. Early Miocene extensional detachments at the massif's northern boundary constitute Stage 1 of northeast stretching and tectonic denudation. DuringStage 2, the Central Menderes Metamorphic Core Complex (CMCC) has formed within already exhumed Stage 1 basement.

248 K. Gessner et al. / Gondwana Research 24 (2013) 243–274

The analysis of regional structures and metamorphism shows thatthe tectonic units below the Cycladic Blueschist Unit are different fromthe Aegean area from those in western Turkey. The oldest known base-ment rocks in the Phyllite–Quartzite Unit in Crete are about 510 Ma old(Romano et al., 2004), whereas there is widespread evidence for aPan-African orogenic cycle in parts of theMenderes Nappes. The late Tri-assic to Eocene platform sequence of the Gavrovo–Tripolitza Block hasno equivalent in theMenderes Nappes. The orogenic history of both tec-tonic unitswas also different: theGavrovo–Tripolitza Block did not enterthe subduction zone until about 35–30 Ma, whereas the MenderesNappes were already underthrust by that time.

In contrast to the Aegean Sea region, high-pressure metamor-phism in the Anatolide belt is absent in the structural deeper units.Quantitative data from the Menderes Nappes so far have producedno evidence for Tertiary high-pressure metamorphism (Candan etal., 2001; Ring et al., 2001b; Whitney and Bozkurt, 2002; Regnier etal., 2003; Ring et al., 2004; Catlos and Çemen, 2005; Baker et al.,2008; Oberhänsli et al., 2010). Tertiary metamorphism in the BayındırNappe, which is the structurally deepest nappe in the pile (Gessner

et al., 1998; Ring et al., 1999a; Gessner et al., 2001c, 2010), reached4–6 kbar at a maximum of 400–450 °C (Ring et al., 2007b). Availableage data indicate ages of 42–37 Ma for greenschist-facies metamor-phism in the Menderes Nappes (Hetzel and Reischmann, 1996;Catlos and Çemen, 2005; Baker et al., 2008). The Menderes Nappes,together with the overlying Cycladic Blueschist Unit, the Afyon–Ören Unit and the Lycian nappes formed a southward propagatingthrust stack in the Late Eocene and Oligocene (Fig. 4) (Collins andRobertson, 1997, 1998; Gessner et al., 2001c; Rimmelé et al., 2003;Pourteau et al., 2010). While the underthrusting of Anatolia caused agreenschist-facies metamorphic belt in western Turkey, ongoing deepsubduction in the Aegean caused an orogenic wedge characterised bysustained high-pressure metamorphism (Ring et al., 2007b). The struc-tural data constrain two important aspects: firstly, there is no evidencefor Alpine high-pressure metamorphism in the Menderes nappes, andsecondly, the available data are consistent with the proposal thatmaximum temperature and age of metamorphism associated withAlpine shortening decrease structurally downward. Temperatures inthe Selimiye nappe were >450 °C and occurred before 43–37 Ma

249K. Gessner et al. / Gondwana Research 24 (2013) 243–274

(Hetzel and Reischmann, 1996), whereas in the Bayındır nappe temper-atures barely reached 400 °C and occurred later at ca. 37 Ma (Lips et al.,2001).

2.3. Controversies on Alpine tectonics of the Menderes Massif

The Menderes Massif is a complex geological terrain that stillyields unresolved issues regarding its tectonic and metamorphichistory.

2.3.1. Alpine crustal shortening and the age of deformation fabricsAlpine shortening of the Menderes Massif has been interpreted in

terms of large-scale recumbent fold (Okay, 2001; Gessner et al., 2002),a series of nappes stacked during south directed thrusting (Ring et al.,1999a; Gessner et al., 2001c), and a series of north-directed thruststhat collapsed either in a bivergent fashion (Hetzel et al., 1995a) orthrough top-to-south extension (Bozkurt and Park, 1994; Bozkurt,2007). The key controversies are focused onwhich structures are relatedto the kinematics of early Tertiary Alpine crustal shortening, which onesare related to late Tertiary crustal extension, and how this fits with theobserved large scale architecture of the Massif. While the role of Mio-cene to Pliocene normal fault systems bounding the Gediz and BüyükMenderes grabens (see section ‘Miocene to recent extension’) is lesscontroversial, the age of kinematic indicators in the metamorphicrocks of the Massif, and in some cases the age of the protoliths arecontroversial. Top-N–N/NE shear sense indicators are common inamphibolite facies metamorphic rocks in the Menderes Massif. Outsidethe contact aureoles of Miocene intrusions, these fabrics predateMiocene extension, and have also been interpreted as Alpine nappestacking (Bozkurt and Park, 1994; Bozkurt, 1995; Hetzel et al., 1995a;Hetzel et al., 1998; Lips et al., 2001; Bozkurt, 2007). Based on detailedregional fabric mapping and cross-cutting relationships, a number ofstudies have shown that the Menderes nappe stack was assembled bysouth-directed shearing under greenschist facies conditions and thatnorth-directed fabrics often are relics of earlier deformation events inindividual tectonic units (Gessner et al., 2001a, 2001c, 2004). Regionallysignificant north-directed kinematics alsowould be difficult to reconcilewith regional tectonic models that have shown that Tertiary conver-gence encompasses south directed shearing and thrusting (e.g. Sengörand Yilmaz, 1981; Sengör et al., 1984; Collins and Robertson, 1998;Gessner et al., 2001c; van Hinsbergen et al., 2010b; Gessner et al.,2011; van Hinsbergen and Schmid, 2012). High-grade metamorphismin the Çine and Bozdağ nappes (Candan et al., 2001; Ring et al., 2001b;Ring et al., 2004; Oberhänsli et al., 2010) occurred before the intrusionof Neoproterozoic to Cambrian granites, and reliable P–T estimates forthe Tertiary tectonometamorphic evolution only exist for the uppermostnappe of the Menderes nappe pile, the Selimiye Nappe (Fig. 4) (Whitneyand Bozkurt, 2002; Regnier et al., 2003). Metasediments in the SelimiyeNappe reached pressures of ca. 6 kbar and temperatures of ca. 500 °Cnear the base of the nappe, decreasing up section (Regnier et al., 2003).Themineral isograds in the SelimiyeNappe runparallel to the regional fo-liation and parallel to the Selimiye Shear Zone and suggest that theSelimiye Shear Zone formed during this prograde greenschist to loweramphibolite-facies metamorphic event. No reliable P–T estimates existfor chlorite-stable mylonitic rocks within the Cyclades–MenderesThrust. However, biotite is destroyed in the mylonite, and pressures of4–6 kbar in the rocks of the Selimiye Nappe below the thrust suggestP–T conditions of b4–6 kbar and b400 °C in themylonite. These P–T es-timates are largely similar to those from mylonitic metagabbroswithin the Cycladic Blueschist unit (Ring et al., 2007b).

2.3.2. Significance of the Selimiye shear zoneThe tectonic significance of the greenschist facies deformation fab-

rics in the Selimiye shear zone (Fig. 4 and 6) remains controversial.Interpretations include (i) Alpine shortening (Gessner et al., 2001c;Gessner et al., 2004), (ii) Precambrian and Alpine polymetamorphic

deformation (Regnier et al., 2003), (iii) post-Precambrian, pre-Alpinemonometamorphic deformation (Regnier et al., 2006), (iv) folding dur-ing Alpine shortening (Erdogan and Güngör, 2004), and (v) late Alpineextension (Bozkurt and Park, 1994; Bozkurt, 2007). While the currentdown-dip, south-directed sense of shear suggests an apparent exten-sional deformation, the orientation of the Selimiye Shear Zone relativeto Earth's surface may well have been different when the deformationfabrics formed. Also there is inconsistent evidence for a telescopedmetamorphic field gradient, or for a change in cooling history acrossthe Selimiye Shear Zone (Gessner et al., 2001c, 2004). Another conten-tious issue is that a number of authors claim that the granitic rocks in-trude lithologies that can be correlated with Mesozoic sediments andare therefore ‘Alpine’ in age (Sengör et al., 1984; Bozkurt et al., 1993;Bozkurt and Oberhänsli, 2001; Bozkurt et al., 2001; Erdogan andGüngör, 2004). Radiometric ages of the intrusions, however, consistent-ly have given Late Proterozoic to Cambrian ages (Reischmann et al.,1991; Hetzel and Reischmann, 1996; Gessner et al., 2001c, 2004), andwe regard stratigraphy based on lithological correlations in the highlydeformed metasediments of the Selimiye Nappe as problematic.

A further problem is that the Selimiye shear zone appears to bewrapped around the granites and orthogneisses towards the westernoutcrop limit of these lithologies, which has lead to contradicting inter-pretations (Gessner et al., 2001a; Regnier et al., 2003; Gessner et al.,2004; Regnier et al., 2006). Based on lithological and metamorphic sim-ilarities the schists andmarbles overlying the Selimiye nappe can be cor-related with Cycladic blueschists, and that these, as well as the Afyon–Ören Unit preserve high pressure metamorphic relics for which thereis no evidence in the Menderes nappes (Oberhänsli et al., 1998a,1998b; Ring et al., 1999b; Oberhänsli et al., 2001; Rimmelé et al., 2003;Pourteau et al., 2010).

2.3.3. Stratigraphic position of low grade metasedimentsOne of the most contentious geological issues of the Menderes

Massif has been the tectonic position of Carboniferous to Mesozoicmetasedimentary rocks. Along the southern margin of the Çinesubmassif fossilferous Palaeozoic metasediments in what we classifyas the Selimiye nappe, have been known as the Göktepe Formation(Kaaden and Metz, 1954; Schuiling, 1962; Dürr, 1975). Thesemetasediments — our Selimiye nappe — overlie amphibolite faciesmetamorphic rocks that occur above orthogneisses and granitoids thatwe would classify as Çine nappe. In the Aydın Mountains and theBozdağ range in the central massif greenschist facies metasedimentaryunits that have been correlated with the Göktepe Formation occurbelow the amphibolite facies metapelites (our Bozdağ nappe), which,in turn are overlain by what we would consider Çine nappe ortho-gneisses (Dora et al., 1995; Hetzel et al., 1998; Candan et al., 2001;Gessner et al., 2007). This situation has been explained in two differentways: as a recumbent fold, or as a south-directed thrust stack. The recum-bent fold hypothesis (Okay, 2001) is based on the assumption that thecontact between orthogneisses above and below the Palaeozoic to Meso-zoic metasediments represents an equivalent tectonostratographic posi-tion. The thrust hypothesis is based on the analysis of deformationfabric elements (Hetzel et al., 1995a, 1995b, 1998; Gessner et al., 2001a,2001b, 2001c, 2002). These studies suggest that tectonic contacts in theAydın Mountains and the Bozdağ range have formed in tectonic eventsthat include Neoproterozoic to Cambrian shortening, Eocene contraction,andMiocene to recent extension of the crust. According to this hypothesisthe juxtaposition of amphibolite facies rocks with Paleozoic to Mesozoicschists has occurred by south-directed thrusting during Alpine contrac-tion and by bivergent tectonic denudation during Neogene extension(Gessner et al., 2001b).Whilewe consider the case for recumbent foldingon the 100 km scale unlikely for the Menderes Massif (Gessner et al.,2002), reports of non-cylindrical folding, particularly at the southern,and the lateral margins of the Çine submassif (Rimmelé et al., 2003;Erdogan and Güngör, 2004; Regnier et al., 2006; Candan et al., 2011)present a challenge to existing tectonic models of the Menderes Massif.

Fig. 6.Map showing tectonic units of the Alpine nappe stack and the key structures within the Menderes Massif. Due to their unresolved stratigraphic position, the Karaburun pen-insula rocks have been left separate. For order of stacking refer to Fig. 3.

250 K. Gessner et al. / Gondwana Research 24 (2013) 243–274

The thrust hypothesis implies that the Palaeozoic to Mesozoicmetasediments in the Aydın Mountains and the Bozdağ range repre-sents the lowest tectonic unit of the Menderes Massif, and could thusbe a parautochtonous unit that correlates with the Bey Dağları unit(Figs. 4, 5 and 6) (Gessner et al., 2001c; van Hinsbergen et al., 2010b).

2.4. Miocene to recent extension

After earlier extension in the northern Aegean — e.g. Eocene in theRhodope (Dinter, 1998; Burg, 2011) — the onset of north–south exten-sion in the central Aegean Sea region and in the Anatolide Belt of west-ern Turkey has been placed around the Oligocene–Miocene boundary(Schermer et al., 1990; Hetzel et al., 1995b; Seyitoglu and Scott, 1996;Dinter, 1998; Gessner et al., 2001b; Keay et al., 2001; Ring et al.,2003a; Kumerics et al., 2005; Ring and Collins, 2005; Cemen et al.,2006; Thomson and Ring, 2006; Glodny and Hetzel, 2007; Thomson etal., 2009; Öner and Dilek, 2011; Catlos et al., 2012), but the overall mag-nitude of extension differs significantly in both regions. Extension in theAegean has been estimated at ca. 350 km (Gautier et al., 1999), and atca. 150 kmacross theMenderesMassif (vanHinsbergen, 2011). The dif-ference in the amount of extension is also apparent in the topography ofboth regions. The Aegean is largely submerged with the Cycladic archi-pelago representing a horst structure between themore highly extendednorthern Aegean Sea and the Cretan Sea (Tirel et al., 2004). Western

Turkey is characterised by thicker crust than the Aegean (Makris andStobbe, 1984; Saunders et al., 1998; Tirel et al., 2004; Zhu et al., 2006b;Özeren and Holt, 2010; Mutlu and Karabulut, 2011) and this alsoreflected in peak elevation exceeding 2 km. The E–W-oriented grabensin western Turkey bend to the south and curve into a NE orientation inthe vicinity of the Aegean Sea (Fig. 7). An early Miocene or older bound-ary between the Aegean and Anatolian domains has been proposed by anumber of studies, based on the differences in extension geometry andmetamorphic history between Samos and western Anatolia (e.g. Ringet al., 1999b, 2010; Gessner et al., 2011), the tectonic controls on the for-mation of the Late Cretaceous to Palaeocene Bornova Flysch Zone (Okay,2011), and also on the occurrence of NNE-trending active fault systemsand Cenozoic to recent basins in western Anatolia (Sözbilir et al., 2003;Özkaymak and Sozbilir, 2008; Uzel and Sozbilir, 2008, and referencestherein; Erkül, 2010). Uzel and Sozbilir (2008) and Sözbilir et al.(2011) have proposed that this seismically active NNE-trending corridorof crustal deformation represents the transfer zone between the Aegeanand Anatolia and named it the İzmir–Balıkesir Transfer Zone (Fig. 7).

2.4.1. Extension of the Anatolide beltSince the early Miocene the Anatolide belt underwent extensional

deformation (Dewey and Sengör, 1979). Miocene extension is expressedby normal-fault systems of Miocene to recent age (Hancock and Barka,1987; Cohen et al., 1995; Hetzel et al., 1995a, 1995b; Gessner et al.,

251K. Gessner et al. / Gondwana Research 24 (2013) 243–274

2001b; Isik and Tekeli, 2001; Ring and Collins, 2005; Emre and Sözbilir,2007; Glodny and Hetzel, 2007; Erkül, 2010).

The Menderes Massif has experienced a two-stage cooling history(Table 1). Three crustal segments differing in structure and cooling his-tory are identified. The Central Menderes Metamorphic Core Complex(CMCC) represents an ‘inner’ axial segment of the Anatolide Belt and ex-poses its lowest structural levels, whereas the two ‘outer’ submassifs,the Gördes submassif to the north and the Çine submassif to the south,represent higher levels of the nappe stack (Figs. 6 and 7). Rocks inÇine submassif and the Gördes submassif, as well as in the upper struc-tural levels of the CMCC recorded significant coolingduring the latest Ol-igocene and early Miocene (Fig. 8). In the northern part of the Gördessubmassif, cooling most likely occurred as a consequence of rapid tec-tonic denudation during N to NNE-directed movement on the Simavand Alacamdağ detachment systems (Isik and Tekeli, 2001; Ring andCollins, 2005; Erkül, 2010; Bozkurt et al., 2011; Catlos et al., 2012). Inthis area, apatite-fission-track ages show a northward younging trendin the direction of hanging wall movement of the detachments (Fig. 8)(Thomson and Ring, 2006). This view, however has been questionedby some recent studies (Akay, 2009; Hasozbek et al., 2010; Hasozbeket al., 2011; Hasozbek et al., 2012) and we refer to Section 2.5.2, wherewe discuss this controversy in more detail.

There is also strong evidence for relatively rapid cooling in the lateOligocene and early Miocene in the Çine submassif. However, fieldevidence for a well-developed extensional detachment system islacking (Ring et al., 2003a). The apatite fission track data in Fig. 8show a gradient towards older ages across the boundary between

Fig. 7.Map highlighting the extent of the Menderes Nappes and of Tertiary sediments andmNE oriented long axis. Tectonic units overlying the Menderes nappes (Fig. 5) are shown in

the Cycladic Blueschist Unit and the Ören unit. This pattern could beexplained by either a top-S extensional reactivation of the basal thrustof the Ören unit, the tilting of the crustal section (Fitzgerald et al., 1991;Foster and John, 1999) in the footwall of a— now eroded— detachmentsystem, or a combination of both.

The second phase of cooling in the Anatolide belt is related to the for-mation of the CMCC. Since the lateMiocene/Pliocene, twoopposite-facingcontemporaneous normal-fault systems, the Kuzey detachment in thenorth (also known as the Karadut fault, the Alaşehir detachment, or theGediz detachment) and the Güney detachment (also known as theBüyük Menderes detachment) in the south (Hetzel et al., 1995b; Emreand Sözbilir, 1997; Gessner et al., 2001b) have caused symmetrical foot-wall uplift, thus forcing a synform structure on the relatively flat lying Al-pine age structures (Gessner et al., 2001b; van Hinsbergen et al., 2010a)(Fig. 8). Within the CMCC Eocene foliation and the boundaries of the tec-tonic units define an east-trending synform with a wavelength of ca.45 km and an amplitude of ca. 10 km. Across this synform fission-trackcooling ages become younger in the hangingwall displacement direction(Fig. 8) (Gessner et al., 2001b; Ring and Layer, 2003; Thomson and Ring,2006). Miocene sediments only occur in fault-bounded blocks in thehanging wall of the detachment faults (Seyitoglu and Scott, 1996; Emreand Sözbilir, 1997; Çiftçi and Bozkurt, 2009a; Çiftçi and Bozkurt, 2010;Öner and Dilek, 2011). Defined by structure and cooling history, theCMCC extends ca. 100 km east–west and 50 km north–south in the cen-tral part of the Anatolide belt. The detachment systems cut the upperlevels of the Alpine nappe stack for a lateral distance of ca. 80 km,displacing the hanging wall regions to the north above the Kuzey

agmatic rocks. Notice that overall the outcrop of the Menderes nappes is elliptical with agrey.

Table 1Apatite fission-track data.

Sample

no.

Location(in degrees and decimal minutes)

Unit/nappe No. of

crystals

Track density (×106 tr cm-2) Age dispersion

(Pc2)

Central age (Ma)

(± 2s)

Apatite mean track

length (µm ± 1 s.e.)

(no. of tracks)

Standard

deviation

(µm)

THT28 37°38.28’N; 28°18.50’E Çine 20 0.8569

(355)

9.305

(3855)

1.240

(8561)

<0.01%

(99.8%)

20.5 ± 2.7 14.36 ± 0.11

(100)

1.06

THT29 37°31.01’N; 28°21.02’E Çine 20 0.0196

(13)

0.2394

(159)

1.234

(8519)

0.70%

(91.4%)

18.1 ± 10.5 −−

THT31 37°23.31’N; 27°48.01’E Çine 20 0.0453

(35)

0.3595

(278)

1.228

(8477)

<0.01%

(95.0%)

27.7 ± 10.1 −−

THT33 37°26.53’N; 27°42.24’E Selimiye 20 0.644

(44)

0.5544

(379)

1.221

(8434)

<0.02%

(89.1%)

25.4 ± 8.3 14.96 ± 0.16

(37)

0.94

THT34 37°27.11’N; 27°43.04’E Çine 14 0.1247

(62)

1.024

(509)

1.215

(8392)

<0.01%

(93.8%)

26.5 ± 7.4 14.93 ± 0.17

(27)

0.87

THT35 37°23.36’N; 27°45.04’E Selimiye 20 0.0797

(78)

0.6199

(607)

1.209

(8350)

<0.01%

(91.2%)

27.8 ± 7.0 13.81 ± 0.29

(7)

0.71

THT37 37°28.35’N; 27°35.05’E Selimiye 20 0.0924

(90)

0.8350

(813)

1.203

(8307)

0.01%

(97.2%)

23.9 ± 5.6 −−

G1 37°12.87’N; 27°34.87’E 13 0.1557

(34)

1.099

(240)

1.301

(4060)

<0.01%

(95%)

31.5 ± 5.8† −−

G2 37°12.92’N; 27°34.93’E 20 0.1785

(41)

0.9361

(215)

1.292

(4031)

0.13%

(96%)

42.1 ± 7.2† −−

G4 37°12.66’N; 27°34.89’E 10 0.0931

(29)

0.5555

(173)

1.283

(4003)

0.01%

(82%)

36.7 ± 7.4† −−

rs

(Ns)

ri

(Ni)

rd

(Nd)

Notes:

(i). Analyses by external detector method using 0.5 for the 4p/2p geometry correction factor

(ii). Ages calculated using dosimeter glass: CN5 with zCN5 = 358.8 ± 12.7; CN2 with zCN2 = 130.7 ± 2.8

(†) CN5 with ζCN5 = 342.5±3.8

(iii). Pc2 is the probability of obtaining a 2 value for v degrees of freedom where v = no. of crystals - 1

252 K. Gessner et al. / Gondwana Research 24 (2013) 243–274

detachment, and to the south in the Güney detachment. The Kuzey de-tachment dips 15°–20°N and its hanging wall consists of south-dippingMiocene continental basin sequences, locally underlain by small volumesof amphibolite-grade orthogneiss. The footwall exposes a greenschist fa-cies mylonitic shear zone of middle Miocene age (Hetzel et al., 1995a,1995b; Emre andSözbilir, 1997;Glodny andHetzel, 2007). TheGüneyde-tachment is exposed along the northern shoulder of the BüyükMenderesgraben as a 0°–15°S dipping ductile to cataclastic shear zone that consti-tutes the basal cut-off to Neogene basins (Fig. 7) (Gessner et al., 2011).While the Küçük Menderes graben in the centre of the CMCC also datesback to the Miocene, it mainly developed in the Plio-Quaternary andhas not experienced nearly as much extension as the Gediz and BüyükMenderes graben systems (Rojay et al., 2005, and references therein).

Detailed work on the Gediz–Alaşehir graben system at the northernmargin of the CMCC (Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi andBozkurt, 2010) has confirmed the hypothesis of Gessner et al. (2001b)that displacement originated along a high-angle normal fault systemand became shallower in orientation due to footwall uplift. It was alsoshown that the Gediz–Alaşehir graben system has grown from a seriesof smaller normal fault segments that controlled the subsidence inearly Miocene sub-basins, to a larger structure during its later activity(Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt, 2010). An alterna-tive hypothesis, where the Miocene to Pliocene basinal strata in Gediz–Alaşehir graben system are interpreted as having formed in a supra-detachment basin above an initially shallow-dipping detachment(Öner andDilek,2011) isdifficult to reconcilewith theobserved footwalluplift (Gessner et al., 2001b) and with seismic reflection data that sug-gest that sediments accumulated much closer to its southern than itsnorthern margin (Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt,2010). A distinct garnet-bearing orthogneiss, that occurs in the internalpart of the Central Menderes Metamorphic Core Complex as well as in

the hanging wall of the Kuzey detachment suggests a minimumdown-dip displacement of ca. 12 km (Gessner et al., 2001b); thisorder of magnitude of displacement has been supported by numericalmodels of core complex formation (Wijns et al., 2005). Assuming thatthe overall structural symmetry between the two detachment systemsalso applies to displacement-to-length relationships, displacementsalong Güney detachment are likely tomirror those of the Kuzey detach-ment. The Kuzey and the Güney detachments root in theMiocene to re-cent Gediz graben and the Büyük Menderes graben, which continue tobe active (Schaffer, 1900; Eyidogan and Jackson, 1985). The Gediz andBüyük Menderes grabens are associated with a number of geothermalfields (Simsek, 1985; Gökgöz, 1998; Faulds et al., 2009; Gessner et al.,2010), andMiocene to recent volcanic activity north of theGediz grabenhas been associated with ongoing lithospheric extension (Seyitogluet al., 1997; Ersoy et al., 2008; Prelevic et al., 2010a).

The Gediz graben and the BüyükMenderes graben separate the Cen-tral Menderes Metamorphic Core Complex from adjacent plateau-likeareas: the Gördes massif to the north and the Çine massif to the south(Figs. 6 and 7). In both the Gördes and Çine massifs flat-lying Miocenesediments overlie rocks of the Menderes nappe stack. When viewedparallel to the Miocene extension direction the Eocene foliation, thebedding of the Miocene sediments and the remnants of a late Mioceneerosion surface are flat-lying and parallel to each other, althoughthere are pronounced changes along strike, that will be addressed inmore detail in a subsequent section on dynamic topography.

2.4.2. Magmatic record of crustal extensionMagmatic activity related to Alpine convergence in western Turkey

ranges fromEocene toHolocene in agewith the largest volumes of igne-ous rocks produced during the Miocene (e.g. Ersoy et al., 2008). Ingeneral there is a trend from older, subduction-related sub-alkaline

Fig. 8. Map of apatite fission-track age locations. White data points with white numbers represent new data (Table 1); others are taken from published sources (Gessner et al.,2001b; Ring et al., 2003b; Thomson and Ring, 2006). Colours generated by spline interpolation in ESRI ARCGIS10.2, using faults and tectonic regions (grey. purple) as barriers. Noticethe pronounced age gap between the Tavşanlı zone and northern Menderes Massif, compared to a smaller jump in ages between the Menderes and the Ören Unit in the south. Thedecrease of cooling ages towards fault zones in the centre outline the second denudation stage (Central Menderes Metamorphic Core Complex, CMCC).

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magmatic compositions that intruded into the Izmir–Ankara zone in thenorth, to younger, alkaline compositions in the south (Seyitoglu et al.1996; Dilek and Altunkaynak, 2009; Ersoy et al. 2010). Although thesource regions and geodynamic setting of the magmatic rocks havebeen discussed controversially, the emerging consensus appears to bethat the Oligocene–Miocene igneous activity took place in a post-collisional crustal extension setting, and documents thermal melting ofa previously metasomatized subcontinental lithospheric mantle (SCLM).Ersoy et al. (2010) have pointed out that Miocene high-Mg volcanicsalong the NNE-trending Izmir–Balıkesir Transfer Zone—which coincideswith the edge of the Aegean slab (cf. Section 4 ‘Upper mantle structureand active deformation’) — tend to be K-rich, whereas ultrapotassicand shoshonotic suites are common the eastern parts on the volcanicprovince. Dilek and Altunkaynak (2009) have proposed that volcaniccentres along the eastern margin of the magmatic province in theAfyon–Isparta region are related to the western edge of the Cyprus slab.

While Miocene to recent magmatic activity has recorded increasingtemperatures and shallower depths of melting that are consistent withremoval of large portions of the lithosphericmantle below theMenderesMassif, it is unclear whether the removed lithospheric mantle has beenautochthonous or not. A related question is if the metasomatic eventthat produced the subduction signature within the Oligocene to

Miocene igneous compositions relates to Alpine convergence or recordsan older event. Based on the composition of Oligocene to Miocene igne-ous rocks Dilek and Altunkaynak (2009) and Altunkaynak et al. (2012a,2012b) argue that the subcontinental lithosphericmantle of theMende-res Massif was metasomatised in the Miocene by flat subduction of acontinuous African slab that comprised the now separated Cyprus andAegean slabs. A further argument for a flat slab has been made basedon ultra-depleted harzburgitic xenoliths within Miocene lamproiticrocks in western Turkey. Prelevic et al. (2010b) argued that these xeno-liths originated from an intraoceanic subduction system within a flatslab, because they considered the other possible source, Archaean litho-sphere, unlikely. There is, however, increasing evidence for Archaeanmodel ages of crust formation, documented in detritic andmagmatic zir-cons within metamorphic units of the Menderes Massif (Kröner andSengör, 1990; Ring and Collins, 2005; Candan et al., 2011; Zlatkin et al.,2012). In contrast to a Cenozoic metasomatism, Pe-Piper and Piper(2007) consider a large component of mantle metasomatism to be ofNeoproterozoic age. As will be discussed in our section on geophysicalimaging of upper mantle structure, the proposition that a continuousslab of oceanic lithosphere has replaced autochthonous lithosphere be-neath the Menderes Massif during Alpine convergence cannot be easilyreconciled with the existing geophysical data.

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2.5. Controversies on crustal extension

2.5.1. Fabric overprinting — extension or contraction?The discovery that down-dip greenschist facies deformation fabrics

overprint north-directed shearing at higher metamorphic grades inthe southern Çine submassif and in the area between Aydın Mountainsand Bozdağ range led to the proposal that this overprinting relation rep-resents crustal extension (Bozkurt and Park, 1994; Hetzel et al., 1995b;Bozkurt and Park, 1997; Hetzel et al., 1998). In the case of the AydınMountains and Bozdağ range, the greenschist facies fabrics, however,were folded into the large scale synform thatwas forced by the symmet-ric uplift of the detachment footwall areas. For both areas it is also ques-tionable whether or not there are sufficiently ‘telescoped’metamorphicfield gradients that would support a crustal thinning scenario (Gessneret al., 2001a, 2001c). In the case of the Selimiye shear zone, new apatitefission track data presented as part of our data compilation (Fig. 8) shownear uniform ages of between 25 and 30 Ma on both sides of the struc-ture, requiring that it was sealed by late Oligocene times. This shear zonetherefore could did not undergo anyMiocene extension— at least not attemperatures above ca. 120 °C.

2.5.2. Exhumation of the Gördes submassif and the role of the Simavdetachment

Most studies agree that the Gördes submassif was exhumed in theMiocene as a consequence of tectonic denudation. An intriguing map-scale feature of the Gördes submassif is the corrugation-like alternationbetween northeast trending Miocene sedimentary basins, and basementhighs that typically expose Çine nappe orthogneisses (Fig. 7). A numberof hypotheses were put forward to explain this pattern, including theformation of an array of cross-faults that accommodated differentialstretching of the Kuzey detachment hangingwall (Sengör, 1987), a com-ponent of ESE–WNW shortening that accompanied Miocene north–south extension (Yilmaz, 1981; Bozkurt and Park, 1997; Bozkurt, 2003;Cemen et al., 2006), oblique slip faulting in N–S extension (Yilmaz etal., 2000), early strike-slip movement of NE–SW trending faults thatwere later reactivated as normal faults (Özkaymak and Sozbilir, 2008;Özkaymak and Sözbilir, 2012), and the existence of spoon-shaped de-tachment faults at the base of Miocene basins (Purvis and Robertson,2004; Purvis and Robertson, 2005). Recent three-dimensional modelsof Miocene to Pliocene basins in the Alaşehir–Gediz graben system(Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt, 2010) haveshown that the corrugation pattern in the Gördes massif exists as base-ment topography in an axial direction of the Alaşehir–Gediz graben sys-tem, i.e. along strike of the Kuzey detachment. This may support thehypothesis that ESE–WNW shortening was involved in controllingbasin topography, which is also consistent with the observation thatMiocene sediments often show onlap relations towards foldedorthogneiss (Purvis and Robertson, 2005; Cemen et al., 2006). Foldingparallel to extension has been described from other extensional prov-inces, such as the Basin and Range province in North America (e.g. Yin,1991; Fletcher and Bartley, 1994; Fletcher et al., 1995) and the AegeanSea region (Avigad et al., 2001; Jolivet et al., 2004). Scaled physical exper-iments suggest that folding parallel to extension is an unstable deforma-tion mode, where elastic folding of the upper crust gets imposed onviscous mid- to lower crustal layers (Venkat-Ramani and Tikoff, 2002;Lévy and Jaupart, 2011). Lévy and Jaupart (2011) point out that in theirmodel shortening induced by these folds takes place as an elastic re-sponse perpendicular to extension without the need of externally im-posed far-field shortening, and suggest that this process may becommon in extensional provinces. While the relationship between thisfolding pattern and the Simav and Alacamdağ detachment systems hasnot been investigated, the Kuzey detachment appears to postdate the

Fig. 9. Movement of western Turkey relative to Eurasia based on published GPS measuremvalues are in millimetre per year; (a) shows all components (black), (b) westward compofor reference. Notice the pronounced increase in southward component and the decrease in

ESE–WNW shortening. Temporal overlap of folding and normal faultingis, however, very likely. Zhu et al. (2006a, 2006b) presentmoment tensordeterminations that provide evidence for ongoing NNE–SSW extensioncontemporaneous with ESE–WNW shortening in the central Menderesarea. A recent study also reviews and synthesises reports of N–S andNE–SW shortening inMiocene basin rocks of the Alaşehir graben, and in-terprets these in the context of progressive simple shear in an overall ex-tensional situation rather than as episodes of Cenozoic shortening thatwere suggested by earlier workers (cf. Şengör and Bozkurt, 2013; andreferences therein). The tectonic denudation of the Gördes submassifin the footwall of north-directed detachments has been questioned bystudies that favourMiocene cooling of theMenderesmassif due to upliftrelated lithosphere scale cooling after a change from steep to flat slabsubduction below the Menderes Massif (Westaway, 2006; Prelevic etal., 2010b). Furthermore, recent studies on the Simav, Koyunoba andAlaçam plutons (Akay, 2009; Hasozbek et al., 2010; Hasozbek et al.,2011; Hasozbek et al., 2012) have suggested that these magmatic com-plexes stitch tectonic units related to Alpine crustal shortening, andquestioned whether they were emplaced in the footwall of a synchro-nously operating Miocene extensional detachment system. While weacknowledge that the intrusion–deformation relationships in the areamay be more complex than previously suggested, we are very scepticalabout these hypotheses. A strong argument for tectonic denudation ofthe Gördes submassif is the existence of ophiolitic klippen thatdirectly overlie Çine nappe orthogneiss across the Gördes submassif(Fig. 6). Together with the large jump in apatite fission track ages acrossthe contact between the Menderes Massif and the overlying Tavşanlızone (Fig. 8), this suggests to us that large parts of the Alpine nappestack have been cut out by an early Miocene Simav–Alaçamdağ detach-ment system (Isik and Tekeli, 2001; Ring and Collins, 2005; Thomsonand Ring, 2006; Erkül, 2010), which can be traced to ophiolitic klippenthat occur as far south as the southern Gördes massif (e.g. Fig. 6).

The age of normal faulting across the Menderes Massif has beenconstrained by K–Ar dating of brittle fault rocks from the high-angleSimav fault, and the low-angle Kuzey and Güney detachment systems(Hetzel et al., 2013). Hetzel et al. (2013) suggest that the onset of brittlefaulting in the CMCC was diachronous, with cataclasite formation in thehanging wall units dating back to ca. 22 Ma in the Güney detachmentand to ca, 9 Ma in the Kuzey detachment. Both faults, however recordedgouge formation as late as 4–3 Ma. According to Hetzel et al. (2013)brittle faulting in the Simav fault dates back to ca. 17–16 Ma. Bozkurtet al. (2011) produced Rb–Sr ages as young as 12–10 Ma from grainsof a late biotite generation in a mylonitic detachment near Gördes, butthe synkinematic growth of the dated grains is questionable (Hetzel etal., 2013).

Overall we favour a two-stage denudation model that is consistentwith the symmetric structure of the CMCC, and with the age gap be-tween the exhumation of the Simav detachment footwalls betweenca. 23 Ma (Isik and Tekeli, 2001; Ring and Collins, 2005; Erkül,2010) and 16–17 Ma (Hetzel et al., 2013), and the onset of major de-nudation of the CMCC as constrained by the cooling below the Kuzeydetachment in the Late Miocene–Pliocene (Ring et al., 2003a). Wenote that while these two stages are based on structures that can bedetected and mapped across the northern part of the MenderesMassif, it is also plausible that the sequence of detachments, foldsand faults reflects the changing style of strain localisation in a pro-gressively exhuming metamorphic terrain.

2.5.3. Block rotation versus diffuse extensionThe normal fault systems bordering the Gediz and BüyükMenderes

grabens are prominent geological and topographical features with aprolonged normal faulting history that have played an important role

ents (Aktug et al., 2009; McClusky et al., 2000). Stations are located at end of arrow;nent (blue), and (c) southward component (red); outline of Menderes Massif is givenwestward component towards the southwest.

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in the tectonic denudation of the CMCC. There are, however, conflictinginterpretations as to what extent crustal extension across western Tur-key has been localised by these graben systems. Based onpalaeomagnetic studies, it has beenproposed that theGediz graben rep-resents a ‘breakaway’ structure, with fragmented blocks to its south andsouthwest— including the CMCC having rotated around a pole near De-nizli in a counter-clockwise direction since the Early Miocene, whilecrust to the north has not experienced such rotation (van Hinsbergenet al., 2010a). While consistent with topographic features across south-west Turkey, block rotation on the proposed scale is not supported bythe strain field calculated from geodetic GPS measurements (Aktug etal., 2009; Özeren and Holt, 2010; Pérouse et al., 2012). Instead, thesestudies have shown that the area to the west and southwest of the An-atolian plateau is currently deforming in a relatively homogeneous dis-placement field that does neither show distinct block rotations, norsharp strain gradients across structures like the NNE-trending İzmir–Balıkesir Transfer Zone (Erkül, 2010); or the West Anatolian ShearZone proposed by Papanikolaou and Royden (2007).

The apparent discrepancy between the fragmented blocks and thedistributed strain could be that the rotation is a component of thewestward increasing sinistral shearing of western Turkey. Pérouseet al. (2012) have shown that in an absolute plate motion frame thevelocity pattern is toroidal relative to the edges of the Hellenic sub-duction zone, with displacement directions defining rotation polesin northwest Greece and some 200 km west of Cyprus in the EasternMediterranean Sea. Hence, geodetic displacement measurements(Fig. 9) would capture the quasi instantaneous strain component,while the palaeomagnetic data have recorded the finite rotationalstrain after many million years of this shearing.

The block rotation scenario is also incompatible with the observa-tion by Çiftçi and Bozkurt (2010) that the oldest sub-basins in Gedizgraben occur in the east, and the graben system propagated from eastto west.). In a block rotation scenario one would expect propagationof the graben from west to east, i.e. in the direction of decreasing tan-gential displacement towards the proposed rotation pole near Denizli.Both the Gediz and Büyük Menderes grabens, however, are narrowand deep in the east, and become wider and shallower towards thewest. Maximum depths to basement are on the order of 4 km in theeastern Büyük Menderes graben and 2 km in the eastern Gediz graben(Sari and Şalk, 2006; Çiftçi and Bozkurt, 2009a; Isik and Senel, 2009).The eastern (Alaşehir) segment of the Gediz graben also contains theoldest Miocene sedimentary infill (Çiftçi and Bozkurt, 2009a; Gürer etal., 2009; Çiftçi and Bozkurt, 2010;Öner andDilek, 2011). To us this sug-gests an east–west propagation of the grabens, and we speculate thatthey may have originated from the area near Denizli.

3. Topographic response to crustal extension

Miocene crustal extension in western Turkey has been accompa-nied by surface uplift that has exposed and eroded the MenderesMassif around the time of the first stage of tectonic denudation.Thermochronological (Gessner et al., 2001a, 2001b, 2001c; Ring et al.,2003a, 2003b; Thomson and Ring, 2006) and sedimentological studies(Yilmaz et al., 2000, and references therein) suggest that erosion tonear base level produced an extensive and relatively flat land surfacecovered by shallow continental basins from the Gördes submassif inthe north across what is now the CMCC and including most of theÇine submassif. Currently, the centralwest coast of the Anatolian penin-sula is characterised by the transition from the Anatolian plateau to theAegean Sea. The landforms of the area are mainly controlled by a seriesof E–W and ESE–WNW oriented horsts and grabens that bound moun-tain ranges and highlands. This basin and range type topography, whichis particularly conspicuous in the central Menderes Massif, is a conse-quence of Neogene to recent normal and strike-slip faulting combinedwith uplift and erosion. Constraints on the ‘background uplift’ of thearea exist for the Gediz submassif, for which Westaway et al. (2004)

estimated ca. 400 m of regional surface uplift since the Middle Pleisto-cene.We have produced a digital terranemodel to generate topograph-ic swath models, and to model drainage channels across the MenderesMassif. The surface models we present here highlight the short wave-length effect of faulting on the topography of the central Menderes re-gion as well as the long wavelength effect, which we tentatively linkto lower crustal flow.

3.1. Methods and materials

We have used SRTM data (Farr et al., 2007) to generate a terrainmodel from which we have extracted two east–west and two north–south oriented topography swaths, modelled the catchments in theMenderes region, and extracted a representative selection of drainagechannel profiles. The area that defines the swaths is shown in Fig. 10.Swath 1 extends from 27.32°E to 27.53°E, and from 37.02°N to 39.39°N(ca. 19 km×236 km) and swath 2 from 28.12°E to 28.38°E, and from37.02°N to 39.39°N (ca. 23 km×263 km). Swath 3 extends from 26.5°Eto 29.43°E, and from 38.12°N to 38.23°N (256 km×12 km) and swath4 from 26.5°E to 29.43°E, and from 38.7°N to 39°N (253 km×33 km).The data have a spatial resolution of approximately 92 m in N–S, andrange between 72 mand74 m in E–Wdirection. To calculate the profilesshown in Fig. 11, we determined the minima, maxima and mean valuesperpendicular to the orientation of the profiles. No smoothing has beenapplied parallel to the orientation of the profiles.

3.2. Topographic profiles

The higher amplitude and shorter wavelength topography of thenorth–south profiles reflect the dominant E–Worientation of the gra-ben systems (Fig. 11). Profiles 3 and 4 show a long wavelength nega-tive elevation gradient from their eastern end to around 27.5°E. Westof this latitude, roughly corresponding to the western limit of theMenderes Massif and the location of the İzmir–Balıkesir TransferZone (Uzel and Sozbilir, 2008) (Fig. 7), the character of elevationchanges, with the ranges closer the coast showing no systematiclong wavelength elevation trend. This observation is corroborated byProfile 1, which runs along the western margin of the Menderes Massif(Figs. 10 and 11) and also lacks a long wavelength trend. Profile 2 runsacross the Menderes Massif at a high angle to the graben systems andthus highlights the difference between the more plateau-like characterof the Gördes and Çine submassifs, and the much higher Bozdağ andAydın Mountains that were uplifted in the footwall of the detachmentsystems of the CMCC (Fig. 10). The drainage elevation of both the Çinesubmassif and the Gördes submassifs shows a negative gradient to-wards the CMCC, where the base level is higher.

3.3. Drainage channels

We discuss general forms of the channel data using concavity datafor inferring zones where uplift rates change along profiles and to testif there aremajor differences between the CMCC and adjacent plateaux.The stream channel profiles from the CMCC and the adjacent plateauxare distinctly different. The Gördes submassif is drained to the south-west by a number of shallow gradient channels (Fig. 12); the Çinesubmassif by similarly shallow channels towards the northwest. Mostof the channels originating on the plateaux north and south of theCMCC feed in to the axial drainages of the Gediz and Büyük Menderesrivers (channels 28 and 5), in general, the plateau channels get steepertowards the west. Most of the drainage channels originating on the pla-teaux (Fig. 12) have smooth concave profiles with knick points separat-ing upstream and downstream channel segments with different slopes.In general, profiles located on the eastern plateau regions (profiles 4–7)are less concave. However, some of these profiles (especially 7 but also5 and 6) have steep segments in their downstream parts of the profilebefore they enter the valleys around the CMCC uplift area. The outward

Fig. 10. Map showing location of swath profiles shown in the figure, and topographic features such as relevant river valleys, mountain ranges, and coastal embayments.

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draining channels of the CMCC are distinctlymore concave than the pla-teaux channels (Fig. 13). Most of the profiles resemble well-developedriver channel profiles (Wobus et al., 2006) with minor knick points;however, concavity is particularly prominent in profile 30, which tran-sects a well-exposed portion of the Kuzey detachment, and in profiles22 and 23. The inward draining channels of the CMCC (Fig. 14) also rep-resent well-developed river channel profiles, but they are steeper,smoother, and lack major knick points. The N–S oriented channels,which run close to perpendicular to the Kuzey and Güney detachmentsand drain the slopes of the Küçük Menderes valley, are the steepestchannels in the study area and have nearly identical profiles. The axialdrainage channel, the Küçük Menders river (channel 13 in Fig. 14),has amuch shallower gradient than theN–S oriented channels, but nev-ertheless is much steeper than axial channels in the Büyük Menderesand Gediz graben systems (channels 28 and 5 in Fig. 12).

3.4. Interpretation of topography and river channel data

The pronounced mountain ranges and steep drainage channels inthe CMCC (Figs. 13 and 14) are consequences of high uplift rates inthe footwall areas of the Kuzey and Güney detachments. The distinctknick point in profile 30 in the footwall of the Kuzey detachment coin-cides spatially with the youngest apatite fission track data (Fig. 8)(Gessner et al., 2001b; Ring et al., 2003a) and is therefore likely to re-flect an active uplift pulse. Other than the distinct knick point near theKuzey detachment footwall, the inward and outward draining channelprofiles on either side of the CMCC are similar, suggesting similar upliftrates in the footwalls of the opposite facing normal fault systems.

The topographic profiles and the drainage models support the hy-pothesis that the CMCC is a symmetrical uplift that was superimposedon a Miocene peneplain (Yilmaz et al., 2000; Gessner et al., 2001b;Ring et al., 2003a; Thomson and Ring, 2006). The area of this Miocenepeneplain is nearly identical with the outcrop area of the MenderesMassif (Figs. 6 and 8). The channel profiles for the plateaux reflectrelatively slow uplift, which is likely to be consistent with regionalscale long wavelength uplift, as proposed by Westaway et al. (2004).We note that these authors considered the role of footwall uplift andE–W graben formation in the central Menderes to be insignificantfor a lower crustal flow system. Based on our data we find this inter-pretation difficult to support. While data from Miocene fluviatile de-posits in the Gördes submassif indicate mostly north or east directedflow directions (Yilmaz, 1979; Purvis and Robertson, 2005) the chan-nels now are reversed, and presently drain to the southwest. Compa-rable data for the Çine massif are not known, but we note that thenorthward slope of the drainage level is consistent with a symmetricsurface tilting of both plateaux towards the CMCC (Fig. 11). To accom-modate the drastic footwall uplift associated with the Kuzey and theGüney detachments, and the formation of the Miocene to Holocenebasins in the Büyük Menderes and Gediz grabens, redistribution ofmid- to lower crustal material may have played a major role in thecentral Menderes area. Flow of the lower crust as a response to meta-morphic core complex formation (Wdowinski and Axen, 1992; Wijnset al., 2005; Gessner et al., 2007; Rey et al., 2009; Schenker et al., inpress) occurs when lateral pressure gradients drive low viscositylower crust from regions of high overburden into the thinned crustal‘gap’ generated by tectonic denudation. We would argue that the

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Fig. 12. Profiles of channels draining the plateau areas to the north, east and south of the Central Menderes Metamorphic Core Complex. Notice that overall channel slopes are rel-atively small, and increase towards the west.

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overall higher values in the maximum, mean and minimum elevationprofiles, as well as the steeper drainage channels in the CMCC are con-sistent with a lower crustal flow system driven by the Kuzey andGüney detachment systems. Estimates for the viscosity of the lowercrust are similar to values suggested for the Basin and Range provincein North America. For the Menderes, values are likely to be on theorder of between 1019 and 1020 Pa s (Westaway et al., 2004; Sodoudiet al., 2006). We would argue that in the central Menderes regionthe weak lower crust has been driven toward the Gediz and BüyükMenderes grabens from below the southern Gördes and the northernÇine submassifs, and from below the CMCC. The oldest cooling ages inthe CMCC are similar in age to those found near the Miocene peneplain

Fig. 11. Topography profiles across the Menderes Massif calculated from swaths across a terrminimum elevation, which also corresponds to local drainage elevation. Locations of geomorwest of the MM), and Profile 2 (within MM), expressed in Profile 2's higher amplitudes, andderes valleys (Profile 2); a feature not present in Profile 1. Profiles 3 and 4 show a long wacaptures parts of the eastern Aydın Mountains and Bozdağ range area.

on the plateaux to the north and the south, but the rocks occur in core ofthe synform atmuch lower elevations. Even though the CMCC supportshigh mountain ranges in the footwall of the opposite-facing detach-ment systems — probably by elastic flexure — its overall loss of lowercrustal thickness may have caused the central CMCC to ‘sink’ to alower elevation relative to the neighbouring plateaux (Fig. 11).

4. Upper mantle structure and active deformation

Studies based on P-wave and S-wave tomography have shownthat parts of the upper mantle below Anatolia are asthenosphericrather than lithospheric (Spakman et al., 1988; Spakman et al.,

ain model. Red is the maximum elevation, black the lateral mean elevation, and blue thephological features are also shown in Fig. 9. Notice difference between Profile 1 (mostlyan overall trend of decreasing drainage elevation towards the Gediz and Büyük Men-velength decrease in elevation form E to W, which is less clear in Profile 3, because it

Fig. 13. Profiles of channels draining the Central Menderes Metamorphic Core Complex ‘outward’ into the Gediz and Büyük Menderes systems. Notice that slopes are much steepercompared to plateau draining channels (Fig. 11).

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1993; Spakman, 1999; Wortel and Spakman, 2000; Şengör et al.,2003; Faccenna et al., 2006; Berk Biryol et al., 2011; Mutlu andKarabulut, 2011; Paul et al., 2011; Zhu et al., 2012; Jolivet et al., inpress). Due to the inherently low resolution of mantle tomographycompared to geological data, and the variety of the methods used(e.g. body wave tomography in older studies versus adjoint tomogra-phy in Zhu et al., 2012) the shape of the anomaly is not consistentacross the models, and cannot always be related to crustal structurewith confidence. In general a slow wavespeed region of variableshape below western Turkey is a common and robust feature in thetopography models, with the Berk Biryol et al. (2011) model probablybeing the most detailed.

A number of studies have proposed that the asthenospheric windowbelow western Turkey (Fig. 15) originated from a tear in the high-velocity material that separates the oceanic lithosphere domains of theAfrican plate into an Aegean section and a Cyprus section (Dilek andSandvol, 2009; van Hinsbergen et al., 2010b; Berk Biryol et al., 2011;Mutlu and Karabulut, 2011). This scenario would imply that prior to

the rupture, the current Aegean slab extended across all of westernAnatolia, including across the Hellenide–Anatolide transition, and wasconnected to the Cyprus slab (Fig. 15). Dilek and Altunkaynak (2009)on the basis of geochemical arguments, proposed that the Miocene toPliocene trend of southward younging alkaline volcanics between Eski-şehir and Isparta record the formation of a tear, which they interpret tohave separated the Cyprus slab and the Aegean slab, that prior to tearingwould have constituted a continuous slab from central Anatolia to theIonian Sea. While the alkaline volcanic trend is likely to record a signifi-cant lithospheric feature, this model does not account for the easterntermination of the Aegean slab below western Turkey as imaged bymost tomography models. An alternative to tearing a continuous slabcould be that the window reflects a primary lithospheric feature, forexample thick buoyant sub-lithospheric mantle domains within theAnatolides, for which Proterozoic and Archaean ages are known(Kröner and Sengör, 1990; Ring and Collins, 2005; Candan et al., 2011;Zlatkin et al., 2012). To better understand the uppermantle architecture,and to assess how a mantle-scale discontinuity across the Aegean

Fig. 14. Profiles of channels draining the Central Menderes Metamorphic Core Complex ‘inward’ into the Küçük Menderes river. Notice that slopes are much steeper compared toplateau draining channels (Fig. 11).

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coastline of Anatolia relates to crustal structure, we have produced grav-ity anomaly and crustal thicknessmodels across this area, and have pro-duced a three-dimensional representation of earthquake hypocentresand the MIT-P08 seismic tomography model (Li et al., 2008).

4.1. Geophysical evidence

4.1.1. Gravity anomaly and Moho depthWe analysed Sandwell and Smith's 1′×1′ free gravity anomaly grid

(Sandwell and Smith, 2009) and transformed it into a conventionalBouguer anomaly (Fig. 16a) using bathymetry and topography datafrom Smith and Sandwell (1997). We concentrated our modelling effortson Bouguer anomalies with wavelengths longer than 10 km, sinceshorter wavelength anomalies are masked by the lateral resolution ofthe original data and do not capture the gravity signatures of mantle-depth heterogeneities. We used these Bouguer anomaly data to estimateMoho depths assuming lateral continuity of the interface and initial

average depths of 30 km as reported from teleseismic receiver functions(Saunders et al., 1998; Zhu et al., 2006b; Özeren and Holt, 2010; Mutluand Karabulut, 2011). We inverted the gravity data for the relief of theMoho interface using a layered inversion algorithm (Gallardo et al.,2005) assuming a density contrast of 400 kg m−3 (Tirel et al., 2004)and a tension factor for layer continuity of 1×10−4 km−1 for a regularmesh of 5 km-wide cells. These parameters provided a model of depthsthat range between 18.6 and 40.5 km. The gravity response of this inter-face shows regional gravity anomaly (Fig. 16) and fits the data at a stan-dard deviation of 15.26 mGal, which is well above the reported precisionof the satellite gravity data ranging 1.8 to 3.6 mGal (Sandwell and Smith,2009). Residual gravity anomalies not justifiedbyourmodel (Fig. 16c) arelikely to be caused by crustal-depth heterogeneities.

4.1.2. Earthquake hypocentresEarthquake data include magnitude and hypocentre location of

13,935 events listed in the USGS PDE catalogue (USGS, 2011) between

Fig. 15. Location and depth of the upper limit of the fast p-wave speed anomaly of the Hellenic and Cyprus slab fragments, as interpreted by Berk Biryol et al. (2011). The MenderesMassif is located above the western margin of a ca. 300 km wide ‘asthenospheric window’; a slow wave speed anomaly that is commonly interpreted as a tear in the African plate(see text for details).

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19 April 1973 and 14 June 2011, and also including data from eventsthat occurred before 1973 but were not systematically recorded. Datathat lacked either magnitude or depth information were not consid-ered. We subdivided the earthquake locations into different domains,according to the spatial distribution within the study area.

Fig. 16. Bouguer gravity anomaly of western Anatolia (a) modelled from satellite gravity a(c) high frequency residual. The data show a long wavelength positive anomaly (b) in the sgradient can be seen towards central Anatolia, where a slight ridge below the Central Mend(c) show north to northeast oriented linear trends broken up by negative anomalies in the gthe Stage 1 core complex.

4.1.3. 3D model of seismic tomography and earthquake hypocentersFor the three-dimensionalmodel, we projected earthquake data and

seismic velocity anomalies in Cartesian coordinates in UTM Zone 35 Nusing PARADIGM™ SKUA© v.2009.3. Seismic tomography data aresubsampled fromMIT global depth-corrected p-wave velocity anomaly

nd topography data (Sandwell and Smith, 2009); (b) filtered for low frequencies; andouthern Aegean Sea region that becomes weaker to the north. A pronounced negativeeres Metamorphic Core Complex separates negative anomalies. Short wavelength dataraben systems surrounding the Stage 2 core complex and along the northern margin of

Fig. 17. Locations and magnitudes (M) of 11,995 earthquakes (a subset within the region of interest of the total 13,935 shown in Fig. 17 and in the supplementary material)recorded in western Anatolia from 19 April 1973 to 22 April 2011 displayed with (a) active faults and p-wave anomaly data (dVp) at 22.6 km depth and (b) depth of Mohomodelled after satellite gravity data (cf. Fig. 15). Data show relative few earthquakes in the Stage 1 core complex (outlined), except along its northern boundary, and along thegrabens defining the Stage 2 core complex (Fig. 7), especially at an intersection of two graben systems in the east. Earthquake locations data are also presented in 3D PDF format(cf. Supplementary material 1).

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dataset MIT-P08 (Li et al., 2008) and interpolated with SKUA's DiscreteSmooth Interpolator algorithm from1568data points inUTMzone 35 ata resolution of ca. 78 km inN–S (seven rows per depth layer), ca. 61 kmin E–W (eight columns per depth layer); and 45.2 km depth resolution(28 layers). Earthquake data include all events listed in the previoussection.

4.2. Results

A long wavelength, gravity low is located in the Aegean Sea re-gion; it has a steep gradient parallel to the Aegean coastline, and amuch shallower gradient across strike to the north. Crustal thicknessincreases from below 20 km in the Aegean to ca. 40 km at the easternmargin of the Menderes Massif (Figs. 17 and 18). Short wavelengthgravity anomalies reveal linear corrugation trends that are consistentwith folded basement ridges and elongate Miocene basins coveringthe Stage 1 core complex footwall, particularly in the northern partof the Menderes Massif. These linear trends are offset by negativeanomalies of the grabens that formed above the Stage 2 detachmentfaults that eventually fragmented the Stage 1 core complex.

Bouguer gravity (Fig. 17), seismic velocities and crustal thickness(Fig. 18) are consistent with published data (Tirel et al., 2004; Sodoudiet al., 2006; Zhu et al., 2006b) and show how extensional tectonicscaused exhumation of denser and faster material to shallower crustallevels, predominantly in the eastern Aegean but also in western Turkey.

The pattern of earthquake distribution can be used to define dis-tinct domains of seismic activity (i) in the north around Simav and(ii) within the Menderes Massif south of Simav, west of the MenderesMassif, and in the southern part of the study area (Figs. 17, 18, 19,

Table 2). Within the Menderes Massif as well as to its north andwest, earthquakes occur in the shallow crust, with the mean depthbeing shallower in the Simav domain (9.7 km) compared to the west-ern domain (11.9 km) and the Menderes (11.2 km) domain. Our datashow the Menderes Massif as a distinct, seismically relatively quietarea with a crustal thickness of about 30 km (Fig. 17). Relativelyfew seismic events have occurred across the Menderes Massif, exceptfor activity along east to southeast trending graben systems (Figs. 17,18, and 19). Overall, seismic activity shows strong spatial correlationwith the western and northernmargin of the Stage 1 core complex andwith Anatolia's western and southern coastline. The spatial distributionchanges markedly across a NE-trend along the western margin of theMenderes Massif (Figs. 17 and 18). There are also a number of shallowand deep crustal earthquake clusters. In a significant deviation from theoverall depth pattern earthquakes occur at great depths below the Gulfof Gökova separated by a gap from an even deeper cluster below theDodekanes (down to 176 km), in a pronounced, sheet-like cluster thatgets deeper towards the west (Fig. 19, and supplementary material).

Surface rendering of the MIT-P08 tomography model (Fig. 20)shows the outline of the north-dipping Aegean slab. In the Aegean,a low velocity layer in the upper mantle defines a hot mantle wedgethat below the Menderes Massif connects to asthenospheric mantlein the southeast of the model (Fig. 20). The shape of the slab is a ro-bust feature that has been consistently imaged by virtually all tomog-raphy studies and has been interpreted by van Hinsbergen et al.(2005) to record continuous subduction of the African plate sincethe Mesozoic. The location and depth of subducting African litho-sphere is also consistent with seismic receiver function imaging(Sodoudi et al., 2006). The slab has a marked edge in the upper

Fig. 18. Map of earthquakes in the Simav (yellow), the Menderes Massif (blue), the Western (green) and Southern (orange) earthquake domains. The Menderes Massif is outlinedin red; subdivision is based on the overall geographic earthquake distribution pattern. Notice the change of event density in the central and northern areas of the map, where do-mains are defined by a NE-trend at the western margin of the Menderes Massif, and continuing north from there. See Fig. 18 for cross section views; coordinates refer to WGS1984UTM zone 35 N (in thousands of metres).

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mantle below western Turkey but continues eastward for 0.5%, 0.6%,and 0.7% p-wave velocity anomaly contours. The surface representa-tion of the slab edge as based on the MIT-P08 model, differs some-what from Berk Biryol et al. (2011), who see the slab continuingslightly further to the east (Fig. 15).

Previous studies have suggested that the Benioff zone of the Aegeansubduction zone may extend to the Turkish coastline (Papazachos et al.,2000; Sodoudi et al., 2006). While this may be the case for theDodekanes area, the deep earthquake cluster below the Gulf of Gökovadoes not correlate with any positive seismic velocity anomaly in theMIT-P08 model that would indicate the presence of a slab (Fig. 18),and the deeper earthquakes occur in a region for whichmost other seis-mic tomography models do not show a steep E–W oriented slab thatcould be interpreted as a Benioff zone. A possible exception is the Zhuet al. (2012) model where anomalies in the vertically polarised compo-nent of the shear wave speed outline a gradient parallel to the Gulf ofGökova in the 275 km depth slice.

The Gökova deep earthquake cluster bears strong resemblance inshape, size and depth extent, to a cluster in the south-easternCarpathians, which has been interpreted as a strain pattern of continen-tal delamination (Fillerup et al., 2010), or of a Rayleigh–Taylor type in-stabilitywithin continental lithosphere (Lorinczi andHouseman, 2009).We would argue that the location and shape of the Gökova earthquake

cluster suggests a similar scenario where deep earthquakes occur in asteep sheet of delaminated or otherwise detached continental crust.

5. Tectonic synthesis

We have reviewed the structure of the Menderes Massif in the lightof new and published geological and geophysical data. The picture thatis emerging is a snapshot of an extending orogenic system, situated in alaterally inhomogeneous convergent geodynamic setting. The challengein understanding the tectonic evolution of southwest Turkey lies inconstraining how closely the dynamics inferred from mantle structurecan be related to the evolution of geological structure. The datasets wehave produced in this study provide new detail on western Anatolia'slithospheric structure, highlight pronounced differences to the Aegean,and support our claim of the existence of the West Anatolia TransferZone (WATZ), a transtensional structure that has originated from awide lithosphere scale transtension zone that denuded the MenderesMassif in the Miocene. We argue that the İzmir–Balıkesir TransferZone, the structural corridor that limits the western exposure of theMenderes Massif, is the current upper crustal expression of the litho-spheric mantle-driven transtension across theWestern Anatolia Trans-fer zone.

Fig. 19. Cross-sections and summary statistics of earthquake depth data shown in Figs. 16 and 17. The features that stand out are the deep earthquakes in the southern domain thatrange from ca. 60 km in the east to 176 km in the west with a notable gap between ca. 520,000 and 550,000 mE, and vertical clusters in the crust at a northing of ca. 41,000 in thesouthern domain along the Gulf of Gökova (see Fig. 17 for location), and a second cluster southeast of Izmir at a northing of ca. 42,220. There is also an increase in mean depth fromnorth to south, and a much greater depth range in the southern domain. Clusters to the northwest and north of the Menderes Massif have different depths, with the Simav domainhaving the shallowest mean depth and also a lower mean magnitude than the Western domain (also see Figs. 16 and 17).

Table 2Summary statistics of earthquake depth and magnitude.

Depth

Domain n Min [m] Max [m] Mean [m] Median [m] Standarddeviation

Simav 2581 1000 60,000 9747 10,000 3894South 2086 2000 170,000 28,044 13,000 35,342Menderes(MM)

670 1000 58,000 11,176 10,000 6957

West 6656 1000 176,000 11,912 10,000 7734

Magnitude

Domain n Min Max Mean Median Standarddeviation

Simav 2581 1.70 7.30 2.86 2.70 0.44South 2086 2.40 7.10 3.55 3.40 0.50Menderes(MM)

670 2.50 6.50 3.40 3.20 0.56

West 6656 1.90 7.50 3.11 3.00 0.48

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5.1. Lateral differences in lithospheric structure

Below the Aegean Sea, the north-subducting slab is imaged by boththe gravity data (Fig. 16) andby the P-wave velocitymodel of themantle(Figs. 17 and 20). We envisage that in the late Oligocene/early Miocenethe Hellenide–Anatolide boundary was expressed as the lateral transi-tion from dense oceanic lithosphere in the Aegean to the thickened con-tinental root below the Anatolide belt. It is difficult to assess the tectonicrecord of this transition, but it is likely that it represents a former domainboundary (e.g. a passivemargin) between the Adriatic and Anatolian do-mains, and was already tectonically active during late Mesozoic conver-gence, when the NNE-trending Bornova Flysch Zone (Fig. 5) formed in atranstensional setting (Okay, 2011). We argue that the continuingsouthwest-ward rollback of the Aegean slab by the late Oligocene/earlyMiocene initiated theWest Anatolia Transfer Zone (WATZ) either by lo-calization across a lithosphere scalematerial boundary, or by reactivationof an earlier feature. We would argue that rollback turned the north–south oriented trailing edge of the Aegean slab into a rigid transtensional

Fig. 20. Three-dimensional model of P-wave velocity (dVp) anomaly contours andhypocentres in the eastern Aegean and western Anatolia from the surface to a depthof ca. 1250 km. Land surface (white, transparent) is shown for reference. Thenorth-dipping slab (a) is discontinuous in the upper mantle (b), but continues east-ward for 0.5%, 0.6%, and 0.7% dVp contours (white arrows). Slow, hot mantle abovethe slab is represented as a negative dVp anomaly in the Aegean (a) that is connectedto a vertical anomaly southeast. In the south, a westward deepening cluster of earth-quake hypocentres in normal velocity lithosphere reaches depths of 176 km. Thesedata are also presented in 3D PDF format (cf. supplementary materials 2 and 3).

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boundary relative towestern Anatolia's sub-continental lithosphere. Thiskinematic framework caused tearing at the western and eastern bound-aries of the continental lithosphere and delamination of large parts of thesubcontinental lithosphericmantle accompanied by uplift, crustal exten-sion, and alkaline magmatism (Figs. 20, 21, and 22).

There is no evidence for a presently subducting slab below theMenderes Massif in the MIT-P08 tomography model. Tomographicdata presented in other studies are variable in orientation, shape,and size of the slow wave speed anomaly (de Boorder et al., 1998;Govers and Wortel, 2005; van Hinsbergen et al., 2010b; Berk Biryolet al., 2011; Paul et al., 2011; Zhu et al., 2012; Jolivet et al., in press).Following the STEP fault concept (Govers and Wortel, 2005), moststudies have proposed that asthenospheric material was emplacedalong a tear fault that ruptured a previously continuous slab to ac-commodate the difference between the fast rollback of the Aegeanslab section versus the slow rollback of the Cyprus slab section(Govers and Wortel, 2005; Dilek and Altunkaynak, 2009). Thiswould require that the lithospheric mantle below the Menderes Mas-sif would have been either no different from the Aegean one, or— if itwas different — would have been sheared off during convergence, assuggested by proponents of flat slab hypotheses (Westaway, 2006;Prelevic et al., 2010b; van Hinsbergen et al., 2010b). In the lattercase it is unclear where the sheared-off autochthonous continentalmantle material would be now, and how it could be imaged. If theca. 300 km wide asthenospheric window formed within a previouslycontinuous flat slab in the Miocene (van Hinsbergen et al., 2010b),the western and eastern edges of the slab (Fig. 15), now separatedby ca. 250 km, would have needed to experience an E–W componentof separation at rates between ca. 7 mm/a (in case of onset in Eocene)and ca. 17 mm/a (onset in the Miocene). Not only is there a lack of ev-idence for an E–W difference within this range in geodetic measure-ments (Fig. 9), but there is also no record of any crustal deformationthat could relate to these kinematics. Furthermore, a gradually wid-ening tear in a continuous slab should have produced a pattern ofsymmetrically outward migrating alkalic magmatic rocks in the Mio-cene, rather than the observed southward progression focused alongthe eastern and western edges of the asthenospheric window(Fig. 15).

There are considerable differences between the Aegean Sea regionand the Menderes Massif, particularly regarding lithospheric evolu-tion. In the light of Archaean zircon ages produced from magmaticand metamorphic rocks (Kröner and Sengör, 1990; Candan et al.,2011; Zlatkin et al., 2012), we regard it possible that the lithospherebelow the Menderes Massif may have been considerably older, andthicker than Aegean lithosphere prior to its Miocene delamination.We propose a scenario where the Menderes arrived into the conver-gence zone containing thick buoyant continental lithosphere with apassive margin sequence towards the north, and possibly to thewest. Upon the collision, the continental lithosphere thickened, andat some stage the leading, oceanic lithospheric domain to the northdecoupled (van Hinsbergen et al., 2010b). The thick lithospherebelow the Menderes then became mechanically unstable either byan increase in temperatures (Houseman et al., 1981; England andHouseman, 1989; Molnar et al., 1993; Platt and England, 1993;Houseman and Molnar, 1997; Stern et al., 2006), due to advection ofasthenospheric mantle where the Aegean slab decoupled, or as a re-sult of heterogeneity in plastic strength (Gorczyk et al., 2012) causingthe thick lithosphere below the Menderes to delaminate (Fig. 22).Such a model would be similar to scenarios proposed for theCarpathians (Lorinczi and Houseman, 2009), the North AmericanGreat Basin (West et al., 2009), New Zealand's North Island (Sternet al., 2006), Eastern Anatolia (Gögüs and Pysklywec, 2008), andintraplate orogenic systems in general (Gorczyk et al., 2013). Asmany studies have pointed out before, southward progressing remov-al of lithospheric mantle below western Turkey is consistent with thechange in age and composition of Miocene to recent volcanic rocks

(Seyitoglu et al., 1997; Pe-Piper and Piper, 2007; Ersoy et al., 2008;Dilek and Altunkaynak, 2009; Ersoy et al. 2010; Prelevic et al.,2010a; Altunkaynak and Dilek, 2009). We would argue that theMiocene southward progression of alkaline volcanic rocks bothalong the Anatolian west coast and along the Afyon–Isparta linetrack the edges of the delaminated Anatolian crustal domain.

5.2. Sinistral transtension across West Anatolian Transfer Zone as a driverfor Menderes extension

The current strainfield shows a clearwestward gradient to higher ve-locities, both in overall movement, and in the southward component(Fig. 9) (Reilinger et al., 2006; Aktug et al., 2009; Pérouse et al., 2012),but it is difficult to assess how fast the two driving processes, slab roll-back and removal of the lithospheric mantle, could have operatedthrough time. In general the geodeticmeasurements of theHellenic sub-duction zone suggest amovement of ca. 33 mm/a (Reilinger et al., 2006),which is at the lower range ofwhat Stegman et al. (2006) and Schellart et

Fig. 21. Summary figure showing the two denudation stages of the Menderes Massif in the context of regional structures, including the area approximate area currently affected bythe West Anatolian Transfer Zone (hatched). Elevations of more than 1000 m are shown with white transparent overlay. Box indicates extent of model shown in Fig. 21.

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al. (2007) consider characteristic for narrow oceanic slabs, but very closeto the ca. 30 mm/a for decoupled continental collision systems inFaccenda et al. (2009). In the Aegean, fast displacement rates havebeen proposed for individual normal fault systems: ca. 20 mm/a alongthe Cretan detachment (Ring and Reischmann, 2002) and 6.5 mm/aalong the Vari detachment on Syros and Tinos. Gessner et al. (2001a,2001b, 2001c) and Wijns et al. (2005) estimated displacement alongthe Kuzey detachment at 2 mm/a. The Aegean displacement rates aretypical for fast displacement rates in metamorphic core complexes, theKuzey detachment would be typical for slower displacement rates(Gessner et al., 2007). The velocity of southward delamination of litho-spheric mantle below western Turkey is difficult to constrain directly,but can be approximated by the progression of alkaline magmatism inwestern Anatolia. Data compiled by Dilek and Altunkaynak (2009) andErsoy et al. (2010) imply that Miocene volcanic activity is limited byboth the eastern edge of the Aegean slab and the western edge of theCyprus slab, and that this activity progressed southward at a rate of ca.10–15 mm/a. Even though these are only rough estimates, these ratessuggest that the southward retreat of the Hellenic subduction zonetook place at least twice as fast as delamination progressed in westernAnatolia. Therefore, we would argue that the formation of the West An-atolian Transfer Zone caused very differentMiocene developments on ei-ther side: rapid rollback of the dense lithosphere of the Adriatic plate inthe Aegean Sea region triggered a surge of lithospheric extension in themid Miocene. In western Anatolia at the same time a plateau with asso-ciated (Yilmaz et al., 2000), and increasingly alkaline volcanic activityformed in the northern Menderes (Seyitoglu et al., 1997; Pe-Piper and

Piper, 2007; Ersoy et al., 2008; Dilek and Altunkaynak, 2009; Prelevic etal., 2010a; Altunkaynak et al., 2012a, 2012b), while the Lycian nappeswere thrust onto the Bey Dağları foreland (Fig. 5) to the south (Collinsand Robertson, 1998). Such close proximity of crustal shortening and ex-tension in conjunction with alkali magmatism is also known from east-ern Anatolia (Şengör et al., 2003; Gögüs and Pysklywec, 2008), and theNorth Island of New Zealand (Stern et al., 2006); areas in convergent set-tings forwhich removal of lithosphere has been proposed. Transtensionalkinematics could provide an explanation for folding about NE–SW toNNE–SSW axes that is most prominent in the northern Menderes(Bozkurt, 2003; Cemen et al., 2006), but has also been described fromother parts of the Menderes (Schuiling, 1962; Rimmelé et al., 2003;Regnier et al., 2006) and may be the reason for the sinusoidal basementtopography in the Alaşehir–Gediz graben system (Çiftçi and Bozkurt,2009a, 2010). Folding parallel to extension at the observed NNE–SSWfold axis orientations and wavelengths of ca. 20–40 km may haverecorded the initial elastic response during Stage 1 crustal extension.Thinning of the crust parallel to stretching could therefore partially havebeen countered by shortening in a perpendicular direction, hence crustalthickness in this casewould be no adequatemeasure of crustal stretching,as for example implied by Zhu et al. (2006b). Folding caused uplift ofbasement in the anticlines, while providing accommodation space forthe Miocene basins in the synclines. The relation between basementtopography and fault segment length in the Alaşehir–Gediz graben sys-tem suggests that transtensional folding may have overlapped with theearly stages of tectonic denudation of the CMCC. This means that theMenderesMassif has experienced northeast–southwest stretching before

Fig. 22. Conceptual model of the present slab dynamics in southwest Turkey, where the southwest retreat of the Aegean slab with its vertical edge maintains a transtensional sit-uation that controls diffuse brittle deformation along the coast and inboard of the Aegean. The two stages of Menderes denudation exposed one of Earth's largest metamorphic corecomplexes displaying frozen in mid-crustal levels of transtensional deformation.

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north–south fragmentation by the Kuzey and Güney detachment sys-tems, and that overall extension in western Anatolia may be greaterthan estimated by the displacement of rigid blocks. The previous estimateof 150 km extension (van Hinsbergen, 2011) is mainly based on the as-sumption that the Simav detachment footwall experienced little or no in-ternal deformation, before being fragmented by the Kuzey and Güneydetachments.

While earthquake distribution patterns support the fragmentationof western Anatolia into different domains (Figs. 18 and 19) geodeticmeasurement shows that the overall strain is very homogeneous(Aktug et al., 2009; Özeren and Holt, 2010). This discrepancy may beexplained to some extent by the variation of rock composition in thestudy area, where heterogeneous rock units such as the tectonic me-langes of the Bornova Flysch and the Tavşanlı Zonemay respond to duc-tile flow of the lower crust by fracturing more often and in a moredistributed manner than the metamorphic rocks of the Menderes Mas-sif. It is also clear that, compared to the current situation, transtensionacross theWATZwould have played amuch larger role during theMio-cene, when transtension affected the entire area of the thermally weak-ened Menderes Massif lithosphere. Structures like the İzmir–BalıkesirTransfer Zone (Uzel and Sozbilir, 2008) — which we argue would bethe present day upper crustal expression of the West Anatolia TransferZone—may at present only constitute a rheological boundary in a widezone of bottom-driven crustal deformation between the North Anato-lian Fault Zone and the Hellenic Trench. That the WATZ does not cutthe surface as one distinct fault zone is likely to be the consequence ofthe overall rheological stratification of the continental lithosphere inthe Aegean and western Turkey. The weak lower crust constitutes athick viscous layer that mechanically decouples the brittle-elasticupper crust from the bottom-driven kinematics, such that the onlything the upper crust ‘feels’ is a westward increase in crustal extension(Fig. 9) resulting in a thinner crust and shallower Moho in the Aegeanrelative to the Menderes area.

5.3. Continuous versus punctuated crustal extension

The geological evidence points to a two-stage denudation of theMenderes Massif, with movement on the Simav detachment slowingdown or stopping around ca. 19–16 Ma (Ring and Collins, 2005;Thomson and Ring, 2006; Hetzel et al., 2013). Synkinematic granitesthat were exhumed in the footwall of the Simav detachment were prob-ably intruded in conjunction with asthenospheric flow caused bydecoupling of the leading part of the African slab (van Hinsbergen et al,2010b). We envisage a scenario where footwall up-doming at the scaleof the Menderes Massif caused secondary late-stage top-S extension atthe base of the Afyon/Ören Unit at the southern margin of the dome,explaining the break in fission-track ages across the base of this unit(Fig. 8).

There is also strong evidence for relatively rapid cooling in the lateOligocene and early Miocene in the Çine submassif. However, field evi-dence for a well-developed extensional detachment system is lacking(Ring et al., 2003a). The steep gradient in apatite fission track ages be-tween the Cycladic Blueschist Unit and the Afyon/Ören unit in Fig. 8could be explained by either a top-S extensional reactivation of thebasal thrust of the Afyon/Ören unit, but also by footwall uplift belowthe original Simav–Alaçam detachment system (Fig. 23), if indeed it ex-tended this far south.

We propose that the development of a wide, distributed WestAnatolia Transfer Zone caused the lithospheric extension in theMenderes Massif and the dome-shape of the evolving core complex(Fig. 23). Extension waned by 19–16 Ma in the northern part of themassif. In the central Menderes Massif granites intruded at 16–15 Ma (Glodny and Hetzel, 2007) and the basin fill suggests ongoingextension in the mid Miocene (Çiftçi and Bozkurt, 2010). Modestfootwall cooling by this time (Ring et al., 2003a, 2003b) suggests lim-ited extensional activity by this time. Continuing uplift of the evolvingStage 1 Menderes core complex led to the formation of a plateau with

Fig. 23. Conceptual model of the two stage tectonic denudation of the Menderes Massif from the Early Miocene to the present. According to our model the monocline in the south-ern Menderes formed as a footwall uplift after Miocene detachment faulting. Miocene crustal melts get exhumed either soon after intrusion in the north, of by the Kuzey detach-ment in the Late Miocene. Corrugation occurring due to E–W shortening of the basement in the footwall of the Early Miocene detachments still shape the drainage of the Çine andGördes submassifs (Fig. 9); the CMCC footwall ‘inherited’ corrugations as topographic features that control the orientation of drainage in the Aydın and Bozdağ mountains (Figs. 12and 13).

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an associated peneplain (Yilmaz et al., 2000). Our surface topographyanalysis supports the hypothesis that this peneplain has been signifi-cantly modified by the movement along the Kuzey and Güney detach-ments and the high-angle normal fault systems in the Gediz andBüyük Menderes grabens since the Pliocene (Gessner et al., 2001b).Whether or not the early Miocene extensional pulse was clearlyseparated in time from the distinct Pliocene to Recent extensionalactivity is difficult to assess. It is conceivable that the two eventsrepresent a continuum of lithospheric extension that commenced inthe early Miocene and then slowed down. As a result of the denuda-tion and cooling of the Menderes Massif, transtension across theWATZ changed character from the wide transtensional deformationof thermally weakened crust to a more focussed structural corridorat the western margin of the massif.

5.4. Open questions

We propose that sinistral movement along the boundary of theAegean and Anatolian domains played amuch larger role in the tectonicevolution of the Anatolian peninsula than previously proposed, particu-larly with regard to the extensional deformation in the Menderes Mas-sif, but also with respect to the distribution of seismic hazard, and thestructural control on hydrothermal metallic resources and geothermalreservoirs (Yigit, 2006; Faulds et al., 2009; Yigit, 2009; Gessner et al.,2010). TheWest Anatolia Transfer Zone is awide anddiffuse lithospher-ic deformation zone that has localised at the western margin of oneof its earlier products, the Menderes Massif, and constrained theshape and location of Anatolia's Aegean coastline. Despite large ad-vances in understanding this rapidly deforming region, there is still a

considerable lack of detailed knowledge, for example on how deforma-tion has partitioned across theWATZ over time, how extensional strainhas been accommodated at the flanks of theMenderes Massif, and howthe crustal structure changes toward the east. There is a need to betterresolve crustal and mantle structure below western Turkey to increaseour knowledge of three-dimensional architecture and to further con-strain the processes that have lead to the low velocity anomaly belowthe Menderes Massif. Our review shows the significance and benefitof integrating geological and geophysical data in three dimensions to ar-rive at a better understanding of lithospheric structure and tectonicevolution.

6. Summary points

• The Hellenide and Anatolide domains of the Tethyan orogen can bedefined on the grounds of their geological history that encompassesdifferences in the age of pre-Alpine basement rocks, as well as instructure, metamorphic and magmatic history related to continentalsubduction and crustal extension.

• The lithospheric mantle across the Hellenide and Anatolide domainsis heterogeneous. Seismic velocity anomalies show a sharp verticalboundary between the fast, cold and dense slab below the Aegeanand a slow, hot and buoyant asthenospheric region below westernTurkey. Gravity data show a north–south oriented boundary betweena high in the Aegean and lower gravity values below the MenderesMassif, and towards the Anatolian plateau to the east.

• We propose that geological differences between the Hellenide andthe Anatolide domains are closely related to the discontinuity in thelithospheric mantle. We interpret this discontinuity as a lithosphere

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scale shear zone, theWest Anatolia Transfer Zone (WATZ), which hasaccommodated the difference between fast roll back of the slab in theAegean and slow delamination of the Anatolian continental litho-sphere since the Miocene, and links the North Anatolian Fault zoneto the Hellenic trench.

• The lack of oceanic lithosphere belowwestern Anatolia's upper man-tle together with the lack of high-pressure metamorphism supportsthe hypothesis that thick buoyant continental lithosphere of the An-atolian microplate brought about an end to continental subduction inwestern Turkey in the Eocene.

• We link the Late Oligocene/Early Miocene to recent crustal extensionin the Anatolide belt in western Turkey to sinistral transtensionacross the WATZ. Within this kinematic framework the MenderesMassif, of one of Earth's largest metamorphic core complexes, has ex-perienced NNE–SSW extension, including extensional detachmentfaulting, folding parallel to extension, doming, and footwall uplift.

• The Menderes Massif appears to be coincident with a Miocene pene-plain, but since the Late Miocene has been fragmented by E–W andWNW–ESE trending graben systems. We propose that fragmentationof the plateau has driven flow in the weak lower crust towards theCentral Menderes area, causing a dynamic topographic responseacross the fracture system in the plateau and in the footwall of Mio-cene to Pliocene graben bounding detachment faults.

• Earthquake locations in western Anatolia strongly correlate with thespatial distribution of tectonic units. Seismic activity in the MenderesMassif is lower than in adjacent regions, but it strongly localises alonggraben systems.

• Most earthquakes occur at shallow to mid-crustal depths, with thenotable exception of a narrow E–Woriented zone of very deep earth-quakes below the Gulf of Gökova. We tentatively link this steep seis-mic zone to delamination of continental lithosphere.

• Our findings highlight the significance of lateral variations in evolv-ing continental arcs for the structure of orogenic belts, particularlywith respect to the formation of metamorphic core complexes.

Acknowledgements

K. Gessner wishes to acknowledge funding by a 2010 University ofWestern Australia Professional Development Award, the AustralianResearch Council (LP100200785), and support by Ariana ResourcesPty. Ltd. and the ARC Centre of Excellence for Core to Crust FluidSystems (Publication 279). We thank J.-P. Burg, P.A. Cawood, B. Çiftçi,M. Fiorentini, T. Güngör, R. Hetzel, A.I.S. Kemp, E. Koralay, Y. Lu, D.J.J.van Hinsbergen, F. Wedin, G. Duclaux, and K.-H. Wyrwoll for discus-sions and comments on previous versions. Y. Dilek and an anonymousreviewer are thanked for formal reviews that helped to improve themanuscript; M. Santosh and T. Horscroft are thanked for outstandingeditorial support, and patience.

Appendix A. Supplementary data

The supplementarymaterial consists of three 3DPDFpages in onefile.Notice that the 3D PDF models, which contain earthquake locations andp-wave anomaly contours — similar to the content shown in Fig. 20 —

allow a range of interactions, including the selection of objects in themodel tree. Explicit instructions are given in Supplementary Fig. 1.Supplementary data associated with this article can be found, in the on-line version, at http://dx.doi.org/10.1016/j.gr.2013.01.005.

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Klaus Gessner is the 3D GeoscienceManager at the Geolog-ical Survey of Western Australia, and Adjunct Research Fel-low at The University of Western Australia. He received a‘Diplom’ in Geology–Palaeontology from Johann WolfgangGoethe University in Frankfurt, Germany (1996), and a PhDfrom Johannes Gutenberg University in Mainz, Germany(2000). Klaus has worked as a structural geologist at CSIROExploration and Mining, and has taught at The Universityof Western Australia. His research focus is on the structuralevolution of Phanerozoic, Proterozoic andArchaean orogenicbelts in Turkey, Australia, and New Zealand, and on the pro-cesses involved in hydrothermal mineral deposits and geo-thermal systems.

Luis A. Gallardo is a titular researcher in the Departmentof Applied Geophysics at CICESE, Mexico; a National Scien-tist from the National Council of Science and Technology inMexico; and an Adjunct Scientist of the Centre for Explora-tion Targeting at the University of Western Australia. Heobtained an MSc in Applied Geophysics from CICESE(1997), a PhD in Environmental Science from LancasterUniversity, UK (2005) and held the Goodeve Lectureshipat The University of Western Australia from 2009 to2011. Luis' research focuses on geophysical inverse theoryand the joint inversion of gravity, electromagnetic andseismic data. He has worked on geophysical imaging for

mineral and petroleum exploration in Western Australia,Western Turkey, Southeast Brazil as well as East and West

Africa. He has also worked on near surface imaging projects around the world for en-vironmental and geotechnical applications.

Vanessa Markwitz is a Research Fellow at the Centre forExploration Targeting at The University of Western Austra-lia in Perth. She graduated from Johann Wolfgang GoetheUniversity in Frankfurt, Germany with a ‘Diplom’ in Geology–Palaeontology (1996). Vanessa has carried out structuralresearch on the RhenishMassif in Germany as a contract geol-ogist for the Geological Survey of Rhineland Palatinate inMainz, Germany. Vanessa is a GIS specialist at the Centre forExploration Targeting and has worked on uranium, nickeland gold prospectivity in Australia, Zimbabwe, West Africa,and Turkey.

Uwe Ring is a Professor in the Department of GeologicalSciences at Stockholm University in Sweden. He obtaineda ‘Diplom’ in Geology from the Technische HochschuleDarmstadt, Germany (1985), and a PhD in Geology fromEberhard-Karls-Universität in Tübingen, Germany (1988).Uwe was a Humboldt Fellow at Yale University in theUSA, has taught Geology at Johannes Gutenberg Universityin Mainz, Germany, and at the University of Canterbury inChristchurch, New Zealand. His research interests includecontinental extensional tectonics in Greece, Turkey, andthe East African Rift. Uwe also has worked on the exhuma-tion of metamorphic rocks from great depths and on the

interactions between climate and tectonics.

Stuart N. Thomson is a Research Scientist at the Universi-ty of Arizona. He obtained a BSc in Geology from DurhamUniversity, UK (1988) and a PhD in Geology from Univer-sity College London (1993). His research focuses primarilyon the application of geochronology and low temperaturethermochronology to problems in geology, tectonics,structural geology, and tectonic geomorphology. He hasworked on projects around the world including on Ceno-zoic glacial–tectonic interactions in the South AmericanAndes, various projects on Mediterranean tectonics in Italy,Greece, and Turkey, to methodological work on apatiteU–Pb dating.