Gondwana Research 24 (2013) 243–274
Contents lists available at SciVerse ScienceDirect
Gondwana Research
journal homepage: www.elsevier.com/locate/gr
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, Australia
Earth Science Division, CICESE, Carretera Ensenada-Tijuana No. 3918, CP 22860, Ensenada, Mexico
Centre for Exploration Targeting, The University of Western Australia, M006, 35 Stirling Highway, Crawley 6009, Australia
d
Department of Geological Sciences, Stockholm University, SE-106 91 Stockholm, Sweden
e
Department of Geosciences, University of Arizona, Gould-Simpson Building, 1040 E. 4th St., Tucson, AZ 85721-0077, USA
b
c
a r t i c l e
i n f o
Article history:
Received 31 March 2012
Received in revised form 28 January 2013
Accepted 31 January 2013
Available online 16 February 2013
Handling Editor: M. Santosh
Keywords:
Metamorphic core complex
Continental extension
Turkey
Aegean Sea
Menderes Massif
Lithosphere delamination
Dynamic topography
a b s t r a c t
The deformation of Earth's lithosphere in orogenic belts is largely forced externally by the sinking slab, but
can also be driven by internal delamination processes caused by mechanical instabilities. Here we present
an integrated analysis of geophysical and geological data to show how these processes can act contemporaneously and in close proximity to each other, along a lithosphere scale discontinuity that defines the lateral
boundary between the Hellenide and Anatolide segments of the Tethyan orogen in western Turkey. The
Hellenides and Anatolides have experienced similar rates of convergence, but display remarkable differences
in 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 tomography. Geological data constrain that one of Earth's largest metamorphic core complexes, the Menderes
Massif, experienced early Miocene tectonic denudation and surface uplift in the footwall of a north-directed
extensional detachment system, followed by late Miocene to recent fragmentation by E–W and NW–SE
trending graben systems. Gravity data, earthquake locations and seismic velocity anomalies highlight a
north–south oriented boundary in the upper mantle between a fast slab below the Aegean and a slow asthenospheric region below western Turkey. Based on the interpretation of geological and geophysical data
we propose that the tectonic denudation of the Menderes Massif and the delamination of its subcontinental
lithospheric mantle reflect the late Oligocene/early Miocene onset of transtension along a lithosphere scale
shear zone, the West Anatolia Transfer Zone (WATZ). We argue that the WATZ localised along the boundary
of the Adriatic and Anatolian lithospheric domains in the Miocene, when southward rollback of the Aegean
slab started to affect the central Aegean–Menderes portion of the Tethyan orogen. Transtension across the
West Anatolia Transfer Zone affected the entire Menderes Massif in the Early Miocene. The current crustal expression of this boundary is a NNE-trending, distributed brittle deformation zone that localised at the western margin of the denuded massif. Here, sinistral transtension accommodates the continuing velocity
difference between relatively slow removal of lithospheric mantle below western Anatolia and trench retreat
in the rapidly extending Aegean Sea region. Our review highlights the significance of lateral variations of the
lower 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.
2.
Introduction . . . . . . . . . . . . . . . . . . . . . . . .
Regional tectonic overview . . . . . . . . . . . . . . . . . .
2.1.
Structure of the Hellenides in the Aegean Sea region . .
2.2.
Structure of the Anatolides in western Turkey . . . . .
2.3.
Controversies on Alpine tectonics of the Menderes Massif
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⁎ Corresponding author at: Geological Survey of Western Australia, Department of Mines and Petroleum, 100 Plain Street, East Perth, WA 6004, Australia.
E-mail addresses: [email protected] (K. Gessner), [email protected] (L.A. Gallardo), [email protected] (V. Markwitz), [email protected] (U. Ring),
[email protected] (S.N. Thomson).
1342-937X/$ – see front matter © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
http://dx.doi.org/10.1016/j.gr.2013.01.005
244
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
2.3.1.
Alpine crustal shortening and the age of deformation fabrics . . . . . . . . . .
2.3.2.
Significance of the Selimiye shear zone . . . . . . . . . . . . . . . . . . . . .
2.3.3.
Stratigraphic position of low grade metasediments . . . . . . . . . . . . . . .
2.4.
Miocene to recent extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.4.1.
Extension of the Anatolide belt . . . . . . . . . . . . . . . . . . . . . . . .
2.4.2.
Magmatic record of crustal extension . . . . . . . . . . . . . . . . . . . . .
2.5.
Controversies on crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.5.1.
Fabric overprinting — extension or contraction? . . . . . . . . . . . . . . . .
2.5.2.
Exhumation of the Gördes submassif and the role of the Simav detachment . . .
2.5.3.
Block rotation versus diffuse extension . . . . . . . . . . . . . . . . . . . . .
3.
Topographic response to crustal extension . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.1.
Methods and materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.
Topographic profiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.
Drainage channels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.4.
Interpretation of topography and river channel data . . . . . . . . . . . . . . . . . . .
4.
Upper mantle structure and active deformation . . . . . . . . . . . . . . . . . . . . . . . .
4.1.
Geophysical evidence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1.1.
Gravity anomaly and Moho depth . . . . . . . . . . . . . . . . . . . . . . .
4.1.2.
Earthquake hypocentres . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1.3.
3D model of seismic tomography and earthquake hypocenters . . . . . . . . .
4.2.
Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.
Tectonic synthesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.1.
Lateral differences in lithospheric structure . . . . . . . . . . . . . . . . . . . . . . .
5.2.
Sinistral transtension across West Anatolian Transfer Zone as a driver for Menderes extension
5.3.
Continuous versus punctuated crustal extension . . . . . . . . . . . . . . . . . . . . .
5.4.
Open questions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.
Summary points . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Appendix A.
Supplementary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
1. Introduction
Lithosphere architecture and strain distribution can vary substantially in orogenic belts, both across and along strike. Along-strike variations and structural complexity are common features of mountain
belts 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) and
the Hellenide–Anatolide orogen in southeastern Europe (Ring et al.,
1999a; Gessner et al., 2001c; Gessner et al., 2011; van Hinsbergen
and Schmid, 2012). The causes for along-strike variations are likely
to differ in individual orogenic belts, but will generally be a consequence of compositional and architectural variations in the accreting
or colliding continental lithosphere fragments. Along-strike variations, however, depend not only on the composition and architecture
of these fragments, but also on the differential dynamics generated by
the sinking slab, and by mechanical instabilities that affect the accretion of continental arcs even at distances far from the actual tectonic
margin, and shape the geology, topography and the lithosphere structure sensed by geophysical data. It has been recognised that throughout the Earth's history tectonic and magmatic accretion of continental
arcs not only have played an important role in the growth of continents (Rudnick, 1995), but also as regions of long-lived thermally
weakened mobile belts (Hyndman et al., 2005). Conceptual and numerical models of generic and regionally specific continental arcs
suggest that deformation is not only mainly driven by external forcing by the sinking slab (Royden, 1993; Collins, 2002; Schellart et al.,
2007; Spakman and Hall, 2010), but also internally, by gravitational
instabilities within thermally weakened lithosphere (Houseman et
al., 1981; England and Houseman, 1989; Molnar et al., 1993; Platt
and England, 1993; Houseman and Molnar, 1997; Stern et al., 2006),
with mechanical and thermal coupling across the subduction zone
determining how these processes interact (Faccenda et al., 2009).
The significance of considering ‘internal drivers’ such as gravitational
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instabilities in addition to ‘external drivers’ such as sinking slabs, is
that synchronous contraction and extension can be accommodated
in the Earth's crust over relatively short across-strike distances
(Gögüs and Pysklywec, 2008; Faccenda et al., 2009). Such internal
driving processes pose a challenge to the existence of regional or
far-field force continua across orogens, an assumption that is often
made a priori when deformation fabrics are linked with geodynamic
processes in ancient orogenic belts. The partition of deformation
along active continental collision zones such as the Tethyan orogen
in the Eastern Mediterranean provide a natural laboratory where
the recent and current evolution of geological structures can be studied and interpreted in the context of surface processes, gravity anomalies, seismicity, geodetic measurements, and mantle tomography. In
the Eastern Mediterranean the southward rollback of the Hellenic
subduction zone and the westward motion of Anatolia dominate the
kinematics of continental plate fragments as they occur at much
higher rates than the convergence between Africa and Eurasia
(e.g. Reilinger et al., 2006; Pérouse et al., 2012) (Fig. 1). This study focuses on southwest Turkey, where the westward movement of Anatolia changes to the southward movement of the Aegean, where the
Anatolian plateau gives way to the Aegean Sea, and where the
Hellenide and Anatolide segments of the Tethyan orogen meet. We
describe the regional structure across the Hellenide–Anatolide transition and, in the light of new and published apatite fission track data,
discuss the tectonic models put forward for the Menderes Massif, particularly with regard to key structures like the Simav detachment and
the Selimiye shear zone. We then use the structure of the Alpine
nappe stack as a marker to track the deformation imposed on western
Anatolia by the late Miocene to recent extension, as evidenced by topography and drainage channel morphology. Using geophysical data
such as gravity, seismic velocity anomaly, and earthquake hypocentre
locations we show how the geological along-strike-differences between the Hellenic and the Anatolide crustal domains relate to the
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
245
the Eastern Mediterranean. We propose that a lithospheric scale
transfer zone, the West Anatolia Transfer Zone (WATZ) defines the
lateral boundary between the Hellenide and Anatolide orogens,
where slab rollback in the Aegean and delamination of the lithospheric mantle in western Anatolia have operated contemporaneously and
in close proximity to each other; causing tectonic denudation of the
lower crust in the Aegean and in the Menderes Massif.
2. Regional tectonic overview
Fig. 1. Kinematic configuration and geodetic measurements of continental fragments
in the Eastern Mediterranean, Arabia and the Caucasus relative to a fixed Eurasia; notice the relatively high velocities of Anatolia and the Aegean — driven by suction of the
Aegean slab — relative to the convergence between Nubia/Arabia and Eurasia — driven
by slab pull, and the uncertainty of how the difference in movement direction between
Anatolia and the Aegean is accommodated.
Modified from Reilinger et al. (2006).
upper mantle structure below western Turkey. Finally, we synthesise
our evidence to discuss the lateral differences in lithosphere structure
as the driver of the Menderes extension and in the geodynamics of
The Hellenide orogen of Greece and the Anatolide belt of western
Turkey form an arcuate orogenic belt north of the Hellenic subduction
zone (Fig. 2). Both the Hellenides and the Anatolides consist of
stacked tectonic units that are overlain by the Late Cretaceous to
Paleogene Vardar–İzmir–Ankara suture zone to the north. The
Adriatic plate has played a key role in the tectonic development of
the Mediterranean region. It has rifted from the northern margin of
Gondwana in the Cretaceous and still moves independently of the
Eurasian Plate. In the eastern Mediterranean little is known about
the Mesozoic to early Tertiary paleogeography of the Adriatic plate,
which appears to pinch out eastwards. In the Mesozoic, continental
crust of the Adriatic plate as exposed today on the Attic Peninsula
and in the Aegean, varied between normal thickness, highly stretched
and thinned; and locally may have been oceanic (Jacobshagen, 1986;
Robertson et al., 1991). The continental fragment directly east of the
Adriatic plate was termed Anatolide–Tauride platform (Sengör and
Yilmaz, 1981), or — following Gessner et al. (2001c) — also as Anatolia. Tectonic units within the Hellenide–Anatolide orogen are aligned
Fig. 2. Simplified tectonic overview of the Aegean Sea region, with Adriatic plate units in blue, Anatolian plate units in pink, Eurasian plate units in brown, and Vardar–İzmir–Ankara
oceanic units in green. The Pindos unit, including widespread high-pressure metamorphic rocks, overlies the External Hellenides in the west. In the east the equivalent Cycladic
Blueschist unit overlies the Menderes Massif, which lacks Alpine high-pressure metamorphism. 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 to
be a ‘Median Crystalline Belt’ (Dürr et al., 1978), the metamorphic
rocks of the Pelagonian zone (Aubouin, 1959), the Cycladic zone, and
the Menderes Massif (Paréjas, 1940; Brinkmann, 1971) are now
known 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; van
Hinsbergen and Schmid, 2012). Rather than representing an eastern extension of the Carboniferous basement and Permo-Mesozoic cover of
the Adriatic plate, the Anatolide belt is made up of two different units
that do not share the same Alpine tectono-metamorphic history, the
Cycladic Blueschist unit and the underlying Menderes nappes (Ring
et al., 1999a; Gessner et al., 2001a; Gessner et al., 2001c; Regnier et al.,
2003). In the Menderes nappes, pronounced magmatic activity occurred
at the Proterozoic/Cambrian boundary (Hetzel and Reischmann, 1996;
Dannat and Reischmann, 1999; Gessner et al., 2001a; Reischmann and
Loos, 2001; Zlatkin et al., 2012), the mid-Triassic (Dannat, 1997;
Koralay et al., 2001) and the Miocene (Hetzel et al., 1995b; Seyitoglu
and Scott, 1996; Isik and Tekeli, 2001; Ring and Collins, 2005; Glodny
and Hetzel, 2007; Ersoy et al., 2008; Akay, 2009; Dilek and
Altunkaynak, 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 Carboniferous
age (Reischmann, 1997; Engel and Reischmann, 1998). In addition,
there are Triassic intrusions (Reischmann, 1997; Ring et al., 1999b) and
prominent 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 example the Serbo-Macedonian Block, (2) the Vardar–İzmir–Ankara
Oceanic units, (3) the Pelagonian Zone, (4) the Pindos Unit (including
the Cycladic Blueschist Unit), (5) the External Hellenides, including
the Gavrovo–Tripolitza Block and the underlying Ionian Block, and
(6) the Mediterranean Ridge Accretionary Complex (Fig. 2). Of
these, only the top three units can be correlated across from the
Hellenides 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 continental
basement–cover sequences, as well as thick radiolarite sequences indicating that locally it was underlain by oceanic crust, or by thinned
continental crust (Pe-Piper and Piper, 1984; Robertson et al., 1991).
In the Cyclades, the uppermost unit of the Pindos Unit is the highly attenuated ophiolitic Selçuk Mélange (Okrusch and Bröcker, 1990; Ring
et al., 1999b; Katzir et al., 2000), which forms the upper part of the
Cycladic Blueschist Unit. The lower part of the Cycladic Blueschist
Unit comprises Carboniferous schist and orthogneiss, and a late- to
post-Carboniferous passive-margin sequence of marble, metapelite
and volcanics (Dürr et al., 1978).
The Gavrovo–Tripolitza Block is a continental platform unit of Triassic to Eocene age, partly overlain by late Eocene to early Oligocene turbidites (Jacobshagen, 1986). Subduction of the Gavrovo–Tripolitza Block
commenced at ca. 35–30 Ma (Thomson et al., 1998; Sotiropoulos and
Kamberis, 2003). In the Cyclades, high-pressure rocks of the Gavrovo–
Tripolitza Block that are locally exposed in tectonic windows below
the Cycladic Blueschist Unit are usually referred to as the Basal Unit
(Godfriaux, 1968; Shaked et al., 2000; Ring et al., 2001a). In the Peloponnese and in Crete, the rocks of the Gavrovo–Tripolitza Block and the
Pindos Unit are only weakly metamorphosed. The Ionian block comprises late Carboniferous to possibly Triassic rocks overlain by limestone
and late Eocene to Miocene turbidites (Jacobshagen, 1986). Rocks of
both the Gavrovo–Tripolitza and Ionian blocks do not crop out in western Turkey. Tectonic units in the footwall of the Pelagonian Unit lack any
Cretaceous orogenic history, and were metamorphosed to high pressures at least 20 Ma later than the Pelagonian Unit and the Vardar–
İzmir–Ankara Oceanic Units (Ring et al., 2010). The most outboard tectonic domain of the Hellenides is the Mediterranean Ridge Accretionary
Complex (Fig. 2) (Kopf et al., 2003). Along the central Mediterranean
Ridge, East Mediterranean oceanic crust has been subducted and the
leading edge of the African passive continental margin is currently entering the subduction zone.
The subduction of the Vardar–İzmir–Ankara Ocean that fringed
Adria and Anatolia on its northern sides caused high-pressure metamorphism in these oceanic units in the Late Cretaceous (Sherlock et
al., 1999). Beginning in the Early Tertiary, the northern edge of the
Pindos Unit was underthrust causing high-pressure metamorphism in
large parts of the Cycladic Blueschist Unit in the central Aegean Sea region (Cyclades islands) and westernmost Anatolia. Well-constrained
ages 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 External Hellenides in Crete and the Peloponnesus in the latest Oligocene
to 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 et
al., in press) demonstrated the progression of high-pressure that
metamorphism gets younger towards the south. Major along-strike
variations in the Hellenide–Anatolide orogen therefore should be related to the arrival of Anatolia in the eastern Mediterranean subduction systems in the Eocene (Gessner et al., 2001c). During incipient
underthrusting of the leading edge of Anatolia the high-pressure
metamorphosed Cycladic Blueschist Unit was thrust onto the Menderes nappes between 42 Ma and 32 Ma (Ring et al., 2007a).
2.2. Structure of the Anatolides in western Turkey
In the Anatolide belt of western Turkey the Pindos Unit (represented
by the Cycladic Blueschist unit) overlies the Menderes nappes (Figs. 3, 4,
and 5) — which are part of Anatolia — whereas in the Aegean region the
Pindos Unit overlies the Basal Unit — which is part of the External
Hellenides (Gavrovo–Tripolitza) (Dürr, 1975; Robertson et al., 1991;
van Hinsbergen et al., 2005). The Vardar–İzmir–Ankara Oceanic units
contain Triassic to Eocene remnants of Neothethys which were subducted below Sakarya since the Cretaceous (Okay and Tüysüz, 1999;
Okay, 2011). In western Turkey, Cretaceous to Palaeogene subduction–accretion complexes constitute the footwall of the Vardar–İzmir–
Ankara suture, including the Tavşanlı zone, and the Bornova Flysch
zone (Okay and Tüysüz, 1999; Okay, 2011). The Ören/Afyon zone and
the Lycian nappes (Okay and Tüysüz, 1999; Pourteau et al., 2010) occur
structurally below the ophiolitic parts of the Vardar–İzmir–Ankara Oceanic units, parts of which may constitute remains of a separate, continuous Anatolian ophiolite nappe (Okay, 2010). The Tavşanlı zone, and the
Afyon/Ören units were metamorphosed under blueschist-facies conditions 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 Cycladic Blueschist occurs above the Menderes Nappes, separated by the
Cyclades–Menderes Thrust (Fig. 4) (Gessner et al., 2001c).
We follow the tectonic division of the Menderes Massif as an Alpine nappe stack (Gessner et al., 1998; Partzsch et al., 1998; Ring et
al., 1999a; Gessner et al., 2001c; Regnier et al., 2003) consisting
of, from top to bottom (1) the Selimiye Nappe, (2) the Cine Nappe,
(3) the Bozdağ Nappe, and (4) the Bayındır Nappe (Fig. 4). The Çine
and Bozdağ nappes have a polyorogenic history, which extends back
into the Neoproterozoic/Cambrian (Kröner and Sengör, 1990; Hetzel
and Reischmann, 1996; Candan et al., 2001; Gessner et al., 2001a;
Gessner et al., 2004; Ring et al., 2004; Catlos and Çemen, 2005; Ring
and Collins, 2005; Oberhänsli et al., 2010; Candan et al., 2011;
Zlatkin et al., 2012).
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
247
Fig. 3. Schematic architecture of tectonic units in the Aegean Sea region and western Anatolia.
According to this subdivision the structurally lowest unit of the
Menderes nappes, the Bayındır nappe, has only been affected by one
major Alpine tectonometamorphic event, whereas in the overlying
Bozdağ, Çine and Selimiye nappes pre-Alpine and Alpine events are
documented. The Cyclades–Menderes thrust cuts through several
nappes of the underlying Menderes nappe stack. Deformation/metamorphism relations across the Cyclades–Menderes thrust indicate that
the breakdown of garnet and biotite to chlorite in the Bozdağ nappe
at temperatures below ca. 400 °C occurred during mylonitization. Accordingly, the Cyclades–Menderes thrust has been interpreted as a
late Alpine out-of-sequence thrust (Gessner et al., 2001c).
The structurally highest tectonic unit, the Selimiye Nappe, contains 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 separates
the Selimiye Nappe from the underlying Çine Nappe (Fig. 4) (Bozkurt
and Park, 1994; Bozkurt and Park, 1997; Gessner et al., 2004). Most of
the Çine nappe consists of deformed orthogneiss, largely undeformed
metagranite and minor pelitic gneiss, eclogite and amphibolite. Protoliths
of much of the orthogneiss/metagranite intruded at ca. 560–530 Ma
(Hetzel and Reischmann, 1996; Hetzel et al., 1998; Loos and
Reischmann, 1999; Gessner et al., 2001a, 2004; Zlatkin et al., 2012). The
underlying Bozdağ Nappe is made up of metapelite containing amphibolite, eclogite and marble lenses. Protolith ages of the Bozdağ Nappe metamorphics are unknown, but geological constraints (Candan et al., 2001;
Gessner et al., 2001a, 2004) suggest a Precambrian age. Like the Çine
Nappe, the Bozdağ Nappe was intruded by granitoids at ca. 240–
230 Ma (Dannat and Reischmann, 1999; Koralay et al., 2001). The
Bayındır Nappe contains phyllite, quartzite, marble and greenschist of inferred 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). The
Bayındır nappe was deformed by the first common deformation event
recorded in the Menderes Massif and the Cycladic blueschist unit
(Gessner et al., 2001c). The corresponding foliation is associated with a
fine-grained N-trending stretching lineation associated with ductile
shear bands and sigma-type objects indicating a top-to-the-S shear
sense (Gessner et al., 2001c).
Fig. 4. Interpretative thrust sequence during formation of Anatolide belt. Notice that the Cyclades–Menders Thrust emplaces units with a high-pressure accretion history on top of
the Menderes nappes.
After Gessner et al. (2011). Age data refer to Lips (1998)†, Loos and Reischmann (1999)††, and Gessner et al. (2001c)†††.
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K. Gessner et al. / Gondwana Research 24 (2013) 243–274
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 Tethyan
orogen in western Anatolia. Early Miocene extensional detachments at the massif's northern boundary constitute Stage 1 of northeast stretching and tectonic denudation. During
Stage 2, the Central Menderes Metamorphic Core Complex (CMCC) has formed within already exhumed Stage 1 basement.
The analysis of regional structures and metamorphism shows that
the tectonic units below the Cycladic Blueschist Unit are different from
the Aegean area from those in western Turkey. The oldest known basement rocks in the Phyllite–Quartzite Unit in Crete are about 510 Ma old
(Romano et al., 2004), whereas there is widespread evidence for a
Pan-African orogenic cycle in parts of the Menderes Nappes. The late Triassic to Eocene platform sequence of the Gavrovo–Tripolitza Block has
no equivalent in the Menderes Nappes. The orogenic history of both tectonic units was also different: the Gavrovo–Tripolitza Block did not enter
the subduction zone until about 35–30 Ma, whereas the Menderes
Nappes were already underthrust by that time.
In contrast to the Aegean Sea region, high-pressure metamorphism in the Anatolide belt is absent in the structural deeper units.
Quantitative data from the Menderes Nappes so far have produced
no evidence for Tertiary high-pressure metamorphism (Candan et
al., 2001; Ring et al., 2001b; Whitney and Bozkurt, 2002; Regnier et
al., 2003; Ring et al., 2004; Catlos and Çemen, 2005; Baker et al.,
2008; Oberhänsli et al., 2010). Tertiary metamorphism in the Bayındır
Nappe, which is the structurally deepest nappe in the pile (Gessner
et al., 1998; Ring et al., 1999a; Gessner et al., 2001c, 2010), reached
4–6 kbar at a maximum of 400–450 °C (Ring et al., 2007b). Available
age data indicate ages of 42–37 Ma for greenschist-facies metamorphism 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 propagating
thrust stack in the Late Eocene and Oligocene (Fig. 4) (Collins and
Robertson, 1997, 1998; Gessner et al., 2001c; Rimmelé et al., 2003;
Pourteau et al., 2010). While the underthrusting of Anatolia caused a
greenschist-facies metamorphic belt in western Turkey, ongoing deep
subduction in the Aegean caused an orogenic wedge characterised by
sustained high-pressure metamorphism (Ring et al., 2007b). The structural data constrain two important aspects: firstly, there is no evidence
for Alpine high-pressure metamorphism in the Menderes nappes, and
secondly, the available data are consistent with the proposal that
maximum temperature and age of metamorphism associated with
Alpine shortening decrease structurally downward. Temperatures in
the Selimiye nappe were >450 °C and occurred before 43–37 Ma
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
(Hetzel and Reischmann, 1996), whereas in the Bayındır nappe temperatures 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 still
yields unresolved issues regarding its tectonic and metamorphic
history.
2.3.1. Alpine crustal shortening and the age of deformation fabrics
Alpine 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 thrusts
that collapsed either in a bivergent fashion (Hetzel et al., 1995a) or
through top-to-south extension (Bozkurt and Park, 1994; Bozkurt,
2007). The key controversies are focused on which structures are related
to the kinematics of early Tertiary Alpine crustal shortening, which ones
are related to late Tertiary crustal extension, and how this fits with the
observed large scale architecture of the Massif. While the role of Miocene to Pliocene normal fault systems bounding the Gediz and Büyük
Menderes grabens (see section ‘Miocene to recent extension’) is less
controversial, the age of kinematic indicators in the metamorphic
rocks of the Massif, and in some cases the age of the protoliths are
controversial. Top-N–N/NE shear sense indicators are common in
amphibolite facies metamorphic rocks in the Menderes Massif. Outside
the contact aureoles of Miocene intrusions, these fabrics predate
Miocene extension, and have also been interpreted as Alpine nappe
stacking (Bozkurt and Park, 1994; Bozkurt, 1995; Hetzel et al., 1995a;
Hetzel et al., 1998; Lips et al., 2001; Bozkurt, 2007). Based on detailed
regional fabric mapping and cross-cutting relationships, a number of
studies have shown that the Menderes nappe stack was assembled by
south-directed shearing under greenschist facies conditions and that
north-directed fabrics often are relics of earlier deformation events in
individual tectonic units (Gessner et al., 2001a, 2001c, 2004). Regionally
significant north-directed kinematics also would be difficult to reconcile
with regional tectonic models that have shown that Tertiary convergence encompasses south directed shearing and thrusting (e.g. Sengör
and 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 metamorphism
in the Çine and Bozdağ nappes (Candan et al., 2001; Ring et al., 2001b;
Ring et al., 2004; Oberhänsli et al., 2010) occurred before the intrusion
of Neoproterozoic to Cambrian granites, and reliable P–T estimates for
the Tertiary tectonometamorphic evolution only exist for the uppermost
nappe of the Menderes nappe pile, the Selimiye Nappe (Fig. 4) (Whitney
and Bozkurt, 2002; Regnier et al., 2003). Metasediments in the Selimiye
Nappe reached pressures of ca. 6 kbar and temperatures of ca. 500 °C
near the base of the nappe, decreasing up section (Regnier et al., 2003).
The mineral isograds in the Selimiye Nappe run parallel to the regional foliation and parallel to the Selimiye Shear Zone and suggest that the
Selimiye Shear Zone formed during this prograde greenschist to lower
amphibolite-facies metamorphic event. No reliable P–T estimates exist
for chlorite-stable mylonitic rocks within the Cyclades–Menderes
Thrust. However, biotite is destroyed in the mylonite, and pressures of
4–6 kbar in the rocks of the Selimiye Nappe below the thrust suggest
P–T conditions of b4–6 kbar and b400 °C in the mylonite. These P–T estimates are largely similar to those from mylonitic metagabbros
within the Cycladic Blueschist unit (Ring et al., 2007b).
2.3.2. Significance of the Selimiye shear zone
The tectonic significance of the greenschist facies deformation fabrics 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
249
deformation (Regnier et al., 2003), (iii) post-Precambrian, pre-Alpine
monometamorphic deformation (Regnier et al., 2006), (iv) folding during Alpine shortening (Erdogan and Güngör, 2004), and (v) late Alpine
extension (Bozkurt and Park, 1994; Bozkurt, 2007). While the current
down-dip, south-directed sense of shear suggests an apparent extensional deformation, the orientation of the Selimiye Shear Zone relative
to Earth's surface may well have been different when the deformation
fabrics formed. Also there is inconsistent evidence for a telescoped
metamorphic field gradient, or for a change in cooling history across
the Selimiye Shear Zone (Gessner et al., 2001c, 2004). Another contentious issue is that a number of authors claim that the granitic rocks intrude lithologies that can be correlated with Mesozoic sediments and
are therefore ‘Alpine’ in age (Sengör et al., 1984; Bozkurt et al., 1993;
Bozkurt and Oberhänsli, 2001; Bozkurt et al., 2001; Erdogan and
Güngör, 2004). Radiometric ages of the intrusions, however, consistently have given Late Proterozoic to Cambrian ages (Reischmann et al.,
1991; Hetzel and Reischmann, 1996; Gessner et al., 2001c, 2004), and
we regard stratigraphy based on lithological correlations in the highly
deformed metasediments of the Selimiye Nappe as problematic.
A further problem is that the Selimiye shear zone appears to be
wrapped around the granites and orthogneisses towards the western
outcrop limit of these lithologies, which has lead to contradicting interpretations (Gessner et al., 2001a; Regnier et al., 2003; Gessner et al.,
2004; Regnier et al., 2006). Based on lithological and metamorphic similarities the schists and marbles overlying the Selimiye nappe can be correlated with Cycladic blueschists, and that these, as well as the Afyon–
Ören Unit preserve high pressure metamorphic relics for which there
is 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 metasediments
One of the most contentious geological issues of the Menderes
Massif has been the tectonic position of Carboniferous to Mesozoic
metasedimentary rocks. Along the southern margin of the Çine
submassif fossilferous Palaeozoic metasediments in what we classify
as the Selimiye nappe, have been known as the Göktepe Formation
(Kaaden and Metz, 1954; Schuiling, 1962; Dürr, 1975). These
metasediments — our Selimiye nappe — overlie amphibolite facies
metamorphic rocks that occur above orthogneisses and granitoids that
we would classify as Çine nappe. In the Aydın Mountains and the
Bozdağ range in the central massif greenschist facies metasedimentary
units that have been correlated with the Göktepe Formation occur
below the amphibolite facies metapelites (our Bozdağ nappe), which,
in turn are overlain by what we would consider Çine nappe orthogneisses (Dora et al., 1995; Hetzel et al., 1998; Candan et al., 2001;
Gessner et al., 2007). This situation has been explained in two different
ways: as a recumbent fold, or as a south-directed thrust stack. The recumbent fold hypothesis (Okay, 2001) is based on the assumption that the
contact between orthogneisses above and below the Palaeozoic to Mesozoic metasediments represents an equivalent tectonostratographic position. The thrust hypothesis is based on the analysis of deformation
fabric elements (Hetzel et al., 1995a, 1995b, 1998; Gessner et al., 2001a,
2001b, 2001c, 2002). These studies suggest that tectonic contacts in the
Aydın Mountains and the Bozdağ range have formed in tectonic events
that include Neoproterozoic to Cambrian shortening, Eocene contraction,
and Miocene to recent extension of the crust. According to this hypothesis
the juxtaposition of amphibolite facies rocks with Paleozoic to Mesozoic
schists has occurred by south-directed thrusting during Alpine contraction and by bivergent tectonic denudation during Neogene extension
(Gessner et al., 2001b). While we consider the case for recumbent folding
on 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.
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K. Gessner et al. / Gondwana Research 24 (2013) 243–274
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 peninsula rocks have been left separate. For order of stacking refer to Fig. 3.
The thrust hypothesis implies that the Palaeozoic to Mesozoic
metasediments in the Aydın Mountains and the Bozdağ range represents the lowest tectonic unit of the Menderes Massif, and could thus
be 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 the
Rhodope (Dinter, 1998; Burg, 2011) — the onset of north–south extension in the central Aegean Sea region and in the Anatolide Belt of western 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 et
al., 2009; Öner and Dilek, 2011; Catlos et al., 2012), but the overall magnitude of extension differs significantly in both regions. Extension in the
Aegean has been estimated at ca. 350 km (Gautier et al., 1999), and at
ca. 150 km across the Menderes Massif (van Hinsbergen, 2011). The difference in the amount of extension is also apparent in the topography of
both regions. The Aegean is largely submerged with the Cycladic archipelago representing a horst structure between the more highly extended
northern Aegean Sea and the Cretan Sea (Tirel et al., 2004). Western
Turkey is characterised by thicker crust than the Aegean (Makris and
Stobbe, 1984; Saunders et al., 1998; Tirel et al., 2004; Zhu et al., 2006b;
Özeren and Holt, 2010; Mutlu and Karabulut, 2011) and this also
reflected in peak elevation exceeding 2 km. The E–W-oriented grabens
in western Turkey bend to the south and curve into a NE orientation in
the vicinity of the Aegean Sea (Fig. 7). An early Miocene or older boundary between the Aegean and Anatolian domains has been proposed by a
number of studies, based on the differences in extension geometry and
metamorphic history between Samos and western Anatolia (e.g. Ring
et al., 1999b, 2010; Gessner et al., 2011), the tectonic controls on the formation of the Late Cretaceous to Palaeocene Bornova Flysch Zone (Okay,
2011), and also on the occurrence of NNE-trending active fault systems
and Cenozoic to recent basins in western Anatolia (Sözbilir et al., 2003;
Özkaymak and Sozbilir, 2008; Uzel and Sozbilir, 2008, and references
therein; Erkül, 2010). Uzel and Sozbilir (2008) and Sözbilir et al.
(2011) have proposed that this seismically active NNE-trending corridor
of crustal deformation represents the transfer zone between the Aegean
and Anatolia and named it the İzmir–Balıkesir Transfer Zone (Fig. 7).
2.4.1. Extension of the Anatolide belt
Since the early Miocene the Anatolide belt underwent extensional
deformation (Dewey and Sengör, 1979). Miocene extension is expressed
by normal-fault systems of Miocene to recent age (Hancock and Barka,
1987; Cohen et al., 1995; Hetzel et al., 1995a, 1995b; Gessner et al.,
K. 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 history are identified. The Central Menderes Metamorphic Core Complex
(CMCC) represents an ‘inner’ axial segment of the Anatolide Belt and exposes 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 in
Çine submassif and the Gördes submassif, as well as in the upper structural levels of the CMCC recorded significant cooling during the latest Oligocene and early Miocene (Fig. 8). In the northern part of the Gördes
submassif, cooling most likely occurred as a consequence of rapid tectonic denudation during N to NNE-directed movement on the Simav
and Alacamdağ detachment systems (Isik and Tekeli, 2001; Ring and
Collins, 2005; Erkül, 2010; Bozkurt et al., 2011; Catlos et al., 2012). In
this area, apatite-fission-track ages show a northward younging trend
in the direction of hanging wall movement of the detachments (Fig. 8)
(Thomson and Ring, 2006). This view, however has been questioned
by some recent studies (Akay, 2009; Hasozbek et al., 2010; Hasozbek
et al., 2011; Hasozbek et al., 2012) and we refer to Section 2.5.2, where
we discuss this controversy in more detail.
There is also strong evidence for relatively rapid cooling in the late
Oligocene and early Miocene in the Çine submassif. However, field
evidence for a well-developed extensional detachment system is
lacking (Ring et al., 2003a). The apatite fission track data in Fig. 8
show a gradient towards older ages across the boundary between
251
the Cycladic Blueschist Unit and the Ören unit. This pattern could be
explained by either a top-S extensional reactivation of the basal thrust
of the Ören unit, the tilting of the crustal section (Fitzgerald et al., 1991;
Foster and John, 1999) in the footwall of a — now eroded — detachment
system, or a combination of both.
The second phase of cooling in the Anatolide belt is related to the formation of the CMCC. Since the late Miocene/Pliocene, two opposite-facing
contemporaneous normal-fault systems, the Kuzey detachment in the
north (also known as the Karadut fault, the Alaşehir detachment, or the
Gediz detachment) and the Güney detachment (also known as the
Büyük Menderes detachment) in the south (Hetzel et al., 1995b; Emre
and Sözbilir, 1997; Gessner et al., 2001b) have caused symmetrical footwall uplift, thus forcing a synform structure on the relatively flat lying Alpine age structures (Gessner et al., 2001b; van Hinsbergen et al., 2010a)
(Fig. 8). Within the CMCC Eocene foliation and the boundaries of the tectonic units define an east-trending synform with a wavelength of ca.
45 km and an amplitude of ca. 10 km. Across this synform fission-track
cooling ages become younger in the hanging wall 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 the
hanging wall of the detachment faults (Seyitoglu and Scott, 1996; Emre
and Sözbilir, 1997; Çiftçi and Bozkurt, 2009a; Çiftçi and Bozkurt, 2010;
Öner and Dilek, 2011). Defined by structure and cooling history, the
CMCC extends ca. 100 km east–west and 50 km north–south in the central part of the Anatolide belt. The detachment systems cut the upper
levels of the Alpine nappe stack for a lateral distance of ca. 80 km,
displacing the hanging wall regions to the north above the Kuzey
Fig. 7. Map highlighting the extent of the Menderes Nappes and of Tertiary sediments and magmatic rocks. Notice that overall the outcrop of the Menderes nappes is elliptical with a
NE oriented long axis. Tectonic units overlying the Menderes nappes (Fig. 5) are shown in grey.
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K. Gessner et al. / Gondwana Research 24 (2013) 243–274
Table 1
Apatite fission-track data.
Track density (×106 tr cm-2)
rs
ri
rd
(Ns)
(Ni)
(Nd)
Age dispersion
(Pc2)
Central age (Ma)
(± 2s)
Apatite mean track
length (µm ± 1 s.e.)
(no. of tracks)
Standard
deviation
(µm)
1.240
(8561)
<0.01%
(99.8%)
20.5 ± 2.7
14.36 ± 0.11
(100)
1.06
0.2394
(159)
1.234
(8519)
0.70%
(91.4%)
18.1 ± 10.5
−
−
0.0453
(35)
0.3595
(278)
1.228
(8477)
<0.01%
(95.0%)
27.7 ± 10.1
−
−
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
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
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†
−
−
Sample
no.
Location
(in degrees and decimal minutes)
Unit/nappe
No. of
crystals
THT28
37°38.28’N; 28°18.50’E
Çine
20
0.8569
(355)
9.305
(3855)
THT29
37°31.01’N; 28°21.02’E
Çine
20
0.0196
(13)
THT31
37°23.31’N; 27°48.01’E
Çine
20
THT33
37°26.53’N; 27°42.24’E
Selimiye
THT34
37°27.11’N; 27°43.04’E
THT35
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
detachment, and to the south in the Güney detachment. The Kuzey detachment dips 15°–20°N and its hanging wall consists of south-dipping
Miocene continental basin sequences, locally underlain by small volumes
of amphibolite-grade orthogneiss. The footwall exposes a greenschist facies mylonitic shear zone of middle Miocene age (Hetzel et al., 1995a,
1995b; Emre and Sözbilir, 1997; Glodny and Hetzel, 2007). The Güney detachment is exposed along the northern shoulder of the Büyük Menderes
graben as a 0°–15°S dipping ductile to cataclastic shear zone that constitutes 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 dates
back to the Miocene, it mainly developed in the Plio-Quaternary and
has not experienced nearly as much extension as the Gediz and Büyük
Menderes graben systems (Rojay et al., 2005, and references therein).
Detailed work on the Gediz–Alaşehir graben system at the northern
margin of the CMCC (Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and
Bozkurt, 2010) has confirmed the hypothesis of Gessner et al. (2001b)
that displacement originated along a high-angle normal fault system
and became shallower in orientation due to footwall uplift. It was also
shown that the Gediz–Alaşehir graben system has grown from a series
of smaller normal fault segments that controlled the subsidence in
early Miocene sub-basins, to a larger structure during its later activity
(Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt, 2010). An alternative hypothesis, where the Miocene to Pliocene basinal strata in Gediz–
Alaşehir graben system are interpreted as having formed in a supradetachment basin above an initially shallow-dipping detachment
(Öner and Dilek, 2011) is difficult to reconcile with the observed footwall
uplift (Gessner et al., 2001b) and with seismic reflection data that suggest that sediments accumulated much closer to its southern than its
northern margin (Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt,
2010). A distinct garnet-bearing orthogneiss, that occurs in the internal
part of the Central Menderes Metamorphic Core Complex as well as in
the hanging wall of the Kuzey detachment suggests a minimum
down-dip displacement of ca. 12 km (Gessner et al., 2001b); this
order of magnitude of displacement has been supported by numerical
models of core complex formation (Wijns et al., 2005). Assuming that
the overall structural symmetry between the two detachment systems
also applies to displacement-to-length relationships, displacements
along Güney detachment are likely to mirror those of the Kuzey detachment. The Kuzey and the Güney detachments root in the Miocene to recent Gediz graben and the Büyük Menderes graben, which continue to
be active (Schaffer, 1900; Eyidogan and Jackson, 1985). The Gediz and
Büyük Menderes grabens are associated with a number of geothermal
fields (Simsek, 1985; Gökgöz, 1998; Faulds et al., 2009; Gessner et al.,
2010), and Miocene to recent volcanic activity north of the Gediz graben
has been associated with ongoing lithospheric extension (Seyitoglu
et al., 1997; Ersoy et al., 2008; Prelevic et al., 2010a).
The Gediz graben and the Büyük Menderes graben separate the Central Menderes Metamorphic Core Complex from adjacent plateau-like
areas: 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 Miocene
sediments overlie rocks of the Menderes nappe stack. When viewed
parallel to the Miocene extension direction the Eocene foliation, the
bedding of the Miocene sediments and the remnants of a late Miocene
erosion surface are flat-lying and parallel to each other, although
there are pronounced changes along strike, that will be addressed in
more detail in a subsequent section on dynamic topography.
2.4.2. Magmatic record of crustal extension
Magmatic activity related to Alpine convergence in western Turkey
ranges from Eocene to Holocene in age with the largest volumes of igneous rocks produced during the Miocene (e.g. Ersoy et al., 2008). In
general there is a trend from older, subduction-related sub-alkaline
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
253
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. Notice
the 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. The
decrease of cooling ages towards fault zones in the centre outline the second denudation stage (Central Menderes Metamorphic Core Complex, CMCC).
magmatic compositions that intruded into the Izmir–Ankara zone in the
north, to younger, alkaline compositions in the south (Seyitoglu et al.
1996; Dilek and Altunkaynak, 2009; Ersoy et al. 2010). Although the
source regions and geodynamic setting of the magmatic rocks have
been discussed controversially, the emerging consensus appears to be
that the Oligocene–Miocene igneous activity took place in a postcollisional crustal extension setting, and documents thermal melting of
a previously metasomatized subcontinental lithospheric mantle (SCLM).
Ersoy et al. (2010) have pointed out that Miocene high-Mg volcanics
along the NNE-trending Izmir–Balıkesir Transfer Zone — which coincides
with the edge of the Aegean slab (cf. Section 4 ‘Upper mantle structure
and active deformation’) — tend to be K-rich, whereas ultrapotassic
and shoshonotic suites are common the eastern parts on the volcanic
province. Dilek and Altunkaynak (2009) have proposed that volcanic
centres along the eastern margin of the magmatic province in the
Afyon–Isparta region are related to the western edge of the Cyprus slab.
While Miocene to recent magmatic activity has recorded increasing
temperatures and shallower depths of melting that are consistent with
removal of large portions of the lithospheric mantle below the Menderes
Massif, it is unclear whether the removed lithospheric mantle has been
autochthonous or not. A related question is if the metasomatic event
that produced the subduction signature within the Oligocene to
Miocene igneous compositions relates to Alpine convergence or records
an older event. Based on the composition of Oligocene to Miocene igneous rocks Dilek and Altunkaynak (2009) and Altunkaynak et al. (2012a,
2012b) argue that the subcontinental lithospheric mantle of the Menderes Massif was metasomatised in the Miocene by flat subduction of a
continuous African slab that comprised the now separated Cyprus and
Aegean slabs. A further argument for a flat slab has been made based
on ultra-depleted harzburgitic xenoliths within Miocene lamproitic
rocks in western Turkey. Prelevic et al. (2010b) argued that these xenoliths originated from an intraoceanic subduction system within a flat
slab, because they considered the other possible source, Archaean lithosphere, unlikely. There is, however, increasing evidence for Archaean
model ages of crust formation, documented in detritic and magmatic zircons within metamorphic units of the Menderes Massif (Kröner and
Sengör, 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 of
Neoproterozoic age. As will be discussed in our section on geophysical
imaging of upper mantle structure, the proposition that a continuous
slab of oceanic lithosphere has replaced autochthonous lithosphere beneath the Menderes Massif during Alpine convergence cannot be easily
reconciled with the existing geophysical data.
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K. Gessner et al. / Gondwana Research 24 (2013) 243–274
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 in
the southern Çine submassif and in the area between Aydın Mountains
and Bozdağ range led to the proposal that this overprinting relation represents crustal extension (Bozkurt and Park, 1994; Hetzel et al., 1995b;
Bozkurt and Park, 1997; Hetzel et al., 1998). In the case of the Aydın
Mountains and Bozdağ range, the greenschist facies fabrics, however,
were folded into the large scale synform that was forced by the symmetric uplift of the detachment footwall areas. For both areas it is also questionable whether or not there are sufficiently ‘telescoped’ metamorphic
field gradients that would support a crustal thinning scenario (Gessner
et al., 2001a, 2001c). In the case of the Selimiye shear zone, new apatite
fission track data presented as part of our data compilation (Fig. 8) show
near uniform ages of between 25 and 30 Ma on both sides of the structure, requiring that it was sealed by late Oligocene times. This shear zone
therefore could did not undergo any Miocene extension — at least not at
temperatures above ca. 120 °C.
2.5.2. Exhumation of the Gördes submassif and the role of the Simav
detachment
Most studies agree that the Gördes submassif was exhumed in the
Miocene as a consequence of tectonic denudation. An intriguing mapscale feature of the Gördes submassif is the corrugation-like alternation
between northeast trending Miocene sedimentary basins, and basement
highs that typically expose Çine nappe orthogneisses (Fig. 7). A number
of hypotheses were put forward to explain this pattern, including the
formation of an array of cross-faults that accommodated differential
stretching of the Kuzey detachment hanging wall (Sengör, 1987), a component 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 et
al., 2000), early strike-slip movement of NE–SW trending faults that
were later reactivated as normal faults (Özkaymak and Sozbilir, 2008;
Özkaymak and Sözbilir, 2012), and the existence of spoon-shaped detachment faults at the base of Miocene basins (Purvis and Robertson,
2004; Purvis and Robertson, 2005). Recent three-dimensional models
of Miocene to Pliocene basins in the Alaşehir–Gediz graben system
(Çiftçi and Bozkurt, 2009a, 2009b; Çiftçi and Bozkurt, 2010) have
shown that the corrugation pattern in the Gördes massif exists as basement topography in an axial direction of the Alaşehir–Gediz graben system, i.e. along strike of the Kuzey detachment. This may support the
hypothesis that ESE–WNW shortening was involved in controlling
basin topography, which is also consistent with the observation that
Miocene sediments often show onlap relations towards folded
orthogneiss (Purvis and Robertson, 2005; Cemen et al., 2006). Folding
parallel to extension has been described from other extensional provinces, such as the Basin and Range province in North America (e.g. Yin,
1991; Fletcher and Bartley, 1994; Fletcher et al., 1995) and the Aegean
Sea region (Avigad et al., 2001; Jolivet et al., 2004). Scaled physical experiments suggest that folding parallel to extension is an unstable deformation mode, where elastic folding of the upper crust gets imposed on
viscous mid- to lower crustal layers (Venkat-Ramani and Tikoff, 2002;
Lévy and Jaupart, 2011). Lévy and Jaupart (2011) point out that in their
model shortening induced by these folds takes place as an elastic response perpendicular to extension without the need of externally imposed far-field shortening, and suggest that this process may be
common in extensional provinces. While the relationship between this
folding pattern and the Simav and Alacamdağ detachment systems has
not been investigated, the Kuzey detachment appears to postdate the
255
ESE–WNW shortening. Temporal overlap of folding and normal faulting
is, however, very likely. Zhu et al. (2006a, 2006b) present moment tensor
determinations that provide evidence for ongoing NNE–SSW extension
contemporaneous with ESE–WNW shortening in the central Menderes
area. A recent study also reviews and synthesises reports of N–S and
NE–SW shortening in Miocene basin rocks of the Alaşehir graben, and interprets these in the context of progressive simple shear in an overall extensional situation rather than as episodes of Cenozoic shortening that
were suggested by earlier workers (cf. Şengör and Bozkurt, 2013; and
references therein). The tectonic denudation of the Gördes submassif
in the footwall of north-directed detachments has been questioned by
studies that favour Miocene cooling of the Menderes massif due to uplift
related lithosphere scale cooling after a change from steep to flat slab
subduction below the Menderes Massif (Westaway, 2006; Prelevic et
al., 2010b). Furthermore, recent studies on the Simav, Koyunoba and
Alaçam plutons (Akay, 2009; Hasozbek et al., 2010; Hasozbek et al.,
2011; Hasozbek et al., 2012) have suggested that these magmatic complexes stitch tectonic units related to Alpine crustal shortening, and
questioned whether they were emplaced in the footwall of a synchronously operating Miocene extensional detachment system. While we
acknowledge that the intrusion–deformation relationships in the area
may be more complex than previously suggested, we are very sceptical
about these hypotheses. A strong argument for tectonic denudation of
the Gördes submassif is the existence of ophiolitic klippen that
directly overlie Çine nappe orthogneiss across the Gördes submassif
(Fig. 6). Together with the large jump in apatite fission track ages across
the contact between the Menderes Massif and the overlying Tavşanlı
zone (Fig. 8), this suggests to us that large parts of the Alpine nappe
stack have been cut out by an early Miocene Simav–Alaçamdağ detachment system (Isik and Tekeli, 2001; Ring and Collins, 2005; Thomson
and Ring, 2006; Erkül, 2010), which can be traced to ophiolitic klippen
that occur as far south as the southern Gördes massif (e.g. Fig. 6).
The age of normal faulting across the Menderes Massif has been
constrained by K–Ar dating of brittle fault rocks from the high-angle
Simav fault, and the low-angle Kuzey and Güney detachment systems
(Hetzel et al., 2013). Hetzel et al. (2013) suggest that the onset of brittle
faulting in the CMCC was diachronous, with cataclasite formation in the
hanging wall units dating back to ca. 22 Ma in the Güney detachment
and to ca, 9 Ma in the Kuzey detachment. Both faults, however recorded
gouge 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. Bozkurt
et al. (2011) produced Rb–Sr ages as young as 12–10 Ma from grains
of a late biotite generation in a mylonitic detachment near Gördes, but
the synkinematic growth of the dated grains is questionable (Hetzel et
al., 2013).
Overall we favour a two-stage denudation model that is consistent
with the symmetric structure of the CMCC, and with the age gap between the exhumation of the Simav detachment footwalls between
ca. 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 denudation of the CMCC as constrained by the cooling below the Kuzey
detachment in the Late Miocene–Pliocene (Ring et al., 2003a). We
note that while these two stages are based on structures that can be
detected and mapped across the northern part of the Menderes
Massif, it is also plausible that the sequence of detachments, folds
and faults reflects the changing style of strain localisation in a progressively exhuming metamorphic terrain.
2.5.3. Block rotation versus diffuse extension
The normal fault systems bordering the Gediz and Büyük Menderes
grabens are prominent geological and topographical features with a
prolonged normal faulting history that have played an important role
Fig. 9. Movement of western Turkey relative to Eurasia based on published GPS measurements (Aktug et al., 2009; McClusky et al., 2000). Stations are located at end of arrow;
values are in millimetre per year; (a) shows all components (black), (b) westward component (blue), and (c) southward component (red); outline of Menderes Massif is given
for reference. Notice the pronounced increase in southward component and the decrease in westward component towards the southwest.
256
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
in the tectonic denudation of the CMCC. There are, however, conflicting
interpretations as to what extent crustal extension across western Turkey has been localised by these graben systems. Based on
palaeomagnetic studies, it has been proposed that the Gediz graben represents a ‘breakaway’ structure, with fragmented blocks to its south and
southwest — including the CMCC having rotated around a pole near Denizli in a counter-clockwise direction since the Early Miocene, while
crust to the north has not experienced such rotation (van Hinsbergen
et al., 2010a). While consistent with topographic features across southwest Turkey, block rotation on the proposed scale is not supported by
the strain field calculated from geodetic GPS measurements (Aktug et
al., 2009; Özeren and Holt, 2010; Pérouse et al., 2012). Instead, these
studies have shown that the area to the west and southwest of the Anatolian plateau is currently deforming in a relatively homogeneous displacement field that does neither show distinct block rotations, nor
sharp strain gradients across structures like the NNE-trending İzmir–
Balıkesir Transfer Zone (Erkül, 2010); or the West Anatolian Shear
Zone proposed by Papanikolaou and Royden (2007).
The apparent discrepancy between the fragmented blocks and the
distributed strain could be that the rotation is a component of the
westward increasing sinistral shearing of western Turkey. Pérouse
et al. (2012) have shown that in an absolute plate motion frame the
velocity pattern is toroidal relative to the edges of the Hellenic subduction zone, with displacement directions defining rotation poles
in northwest Greece and some 200 km west of Cyprus in the Eastern
Mediterranean Sea. Hence, geodetic displacement measurements
(Fig. 9) would capture the quasi instantaneous strain component,
while the palaeomagnetic data have recorded the finite rotational
strain after many million years of this shearing.
The block rotation scenario is also incompatible with the observation by Çiftçi and Bozkurt (2010) that the oldest sub-basins in Gediz
graben occur in the east, and the graben system propagated from east
to west.). In a block rotation scenario one would expect propagation
of the graben from west to east, i.e. in the direction of decreasing tangential displacement towards the proposed rotation pole near Denizli.
Both the Gediz and Büyük Menderes grabens, however, are narrow
and deep in the east, and become wider and shallower towards the
west. Maximum depths to basement are on the order of 4 km in the
eastern 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 the
oldest Miocene sedimentary infill (Çiftçi and Bozkurt, 2009a; Gürer et
al., 2009; Çiftçi and Bozkurt, 2010; Öner and Dilek, 2011). To us this suggests an east–west propagation of the grabens, and we speculate that
they may have originated from the area near Denizli.
estimated ca. 400 m of regional surface uplift since the Middle Pleistocene. We have produced a digital terrane model to generate topographic swath models, and to model drainage channels across the Menderes
Massif. The surface models we present here highlight the short wavelength effect of faulting on the topography of the central Menderes region as well as the long wavelength effect, which we tentatively link
to lower crustal flow.
3.1. Methods and materials
We have used SRTM data (Farr et al., 2007) to generate a terrain
model from which we have extracted two east–west and two north–
south oriented topography swaths, modelled the catchments in the
Menderes region, and extracted a representative selection of drainage
channel 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 from
37.02°N to 39.39°N (ca. 23 km×263 km). Swath 3 extends from 26.5°E
to 29.43°E, and from 38.12°N to 38.23°N (256 km×12 km) and swath
4 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, and
range between 72 m and 74 m in E–W direction. To calculate the profiles
shown in Fig. 11, we determined the minima, maxima and mean values
perpendicular to the orientation of the profiles. No smoothing has been
applied parallel to the orientation of the profiles.
3.2. Topographic profiles
The higher amplitude and shorter wavelength topography of the
north–south profiles reflect the dominant E–W orientation of the graben systems (Fig. 11). Profiles 3 and 4 show a long wavelength negative elevation gradient from their eastern end to around 27.5°E. West
of this latitude, roughly corresponding to the western limit of the
Menderes Massif and the location of the İzmir–Balıkesir Transfer
Zone (Uzel and Sozbilir, 2008) (Fig. 7), the character of elevation
changes, with the ranges closer the coast showing no systematic
long wavelength elevation trend. This observation is corroborated by
Profile 1, which runs along the western margin of the Menderes Massif
(Figs. 10 and 11) and also lacks a long wavelength trend. Profile 2 runs
across the Menderes Massif at a high angle to the graben systems and
thus highlights the difference between the more plateau-like character
of the Gördes and Çine submassifs, and the much higher Bozdağ and
Aydın Mountains that were uplifted in the footwall of the detachment
systems of the CMCC (Fig. 10). The drainage elevation of both the Çine
submassif and the Gördes submassifs shows a negative gradient towards the CMCC, where the base level is higher.
3. Topographic response to crustal extension
3.3. Drainage channels
Miocene crustal extension in western Turkey has been accompanied by surface uplift that has exposed and eroded the Menderes
Massif 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 to
near base level produced an extensive and relatively flat land surface
covered by shallow continental basins from the Gördes submassif in
the north across what is now the CMCC and including most of the
Çine submassif. Currently, the central west coast of the Anatolian peninsula is characterised by the transition from the Anatolian plateau to the
Aegean Sea. The landforms of the area are mainly controlled by a series
of E–W and ESE–WNW oriented horsts and grabens that bound mountain ranges and highlands. This basin and range type topography, which
is particularly conspicuous in the central Menderes Massif, is a consequence of Neogene to recent normal and strike-slip faulting combined
with uplift and erosion. Constraints on the ‘background uplift’ of the
area exist for the Gediz submassif, for which Westaway et al. (2004)
We discuss general forms of the channel data using concavity data
for inferring zones where uplift rates change along profiles and to test
if there are major differences between the CMCC and adjacent plateaux.
The stream channel profiles from the CMCC and the adjacent plateaux
are distinctly different. The Gördes submassif is drained to the southwest by a number of shallow gradient channels (Fig. 12); the Çine
submassif by similarly shallow channels towards the northwest. Most
of the channels originating on the plateaux north and south of the
CMCC feed in to the axial drainages of the Gediz and Büyük Menderes
rivers (channels 28 and 5), in general, the plateau channels get steeper
towards the west. Most of the drainage channels originating on the plateaux (Fig. 12) have smooth concave profiles with knick points separating 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 also
5 and 6) have steep segments in their downstream parts of the profile
before they enter the valleys around the CMCC uplift area. The outward
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
257
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.
draining channels of the CMCC are distinctly more concave than the plateaux channels (Fig. 13). Most of the profiles resemble well-developed
river channel profiles (Wobus et al., 2006) with minor knick points;
however, concavity is particularly prominent in profile 30, which transects a well-exposed portion of the Kuzey detachment, and in profiles
22 and 23. The inward draining channels of the CMCC (Fig. 14) also represent 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 detachments
and drain the slopes of the Küçük Menderes valley, are the steepest
channels in the study area and have nearly identical profiles. The axial
drainage channel, the Küçük Menders river (channel 13 in Fig. 14),
has a much shallower gradient than the N–S oriented channels, but nevertheless is much steeper than axial channels in the Büyük Menderes
and 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 in
the CMCC (Figs. 13 and 14) are consequences of high uplift rates in
the footwall areas of the Kuzey and Güney detachments. The distinct
knick point in profile 30 in the footwall of the Kuzey detachment coincides spatially with the youngest apatite fission track data (Fig. 8)
(Gessner et al., 2001b; Ring et al., 2003a) and is therefore likely to reflect an active uplift pulse. Other than the distinct knick point near the
Kuzey detachment footwall, the inward and outward draining channel
profiles on either side of the CMCC are similar, suggesting similar uplift
rates in the footwalls of the opposite facing normal fault systems.
The topographic profiles and the drainage models support the hypothesis that the CMCC is a symmetrical uplift that was superimposed
on a Miocene peneplain (Yilmaz et al., 2000; Gessner et al., 2001b;
Ring et al., 2003a; Thomson and Ring, 2006). The area of this Miocene
peneplain is nearly identical with the outcrop area of the Menderes
Massif (Figs. 6 and 8). The channel profiles for the plateaux reflect
relatively slow uplift, which is likely to be consistent with regional
scale long wavelength uplift, as proposed by Westaway et al. (2004).
We note that these authors considered the role of footwall uplift and
E–W graben formation in the central Menderes to be insignificant
for a lower crustal flow system. Based on our data we find this interpretation difficult to support. While data from Miocene fluviatile deposits in the Gördes submassif indicate mostly north or east directed
flow directions (Yilmaz, 1979; Purvis and Robertson, 2005) the channels now are reversed, and presently drain to the southwest. Comparable data for the Çine massif are not known, but we note that the
northward slope of the drainage level is consistent with a symmetric
surface tilting of both plateaux towards the CMCC (Fig. 11). To accommodate the drastic footwall uplift associated with the Kuzey and the
Güney detachments, and the formation of the Miocene to Holocene
basins in the Büyük Menderes and Gediz grabens, redistribution of
mid- to lower crustal material may have played a major role in the
central Menderes area. Flow of the lower crust as a response to metamorphic core complex formation (Wdowinski and Axen, 1992; Wijns
et al., 2005; Gessner et al., 2007; Rey et al., 2009; Schenker et al., in
press) occurs when lateral pressure gradients drive low viscosity
lower crust from regions of high overburden into the thinned crustal
‘gap’ generated by tectonic denudation. We would argue that the
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259
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 relatively small, and increase towards the west.
overall higher values in the maximum, mean and minimum elevation
profiles, as well as the steeper drainage channels in the CMCC are consistent with a lower crustal flow system driven by the Kuzey and
Güney detachment systems. Estimates for the viscosity of the lower
crust are similar to values suggested for the Basin and Range province
in North America. For the Menderes, values are likely to be on the
order of between 1019 and 1020 Pa s (Westaway et al., 2004; Sodoudi
et al., 2006). We would argue that in the central Menderes region
the weak lower crust has been driven toward the Gediz and Büyük
Menderes grabens from below the southern Gördes and the northern
Çine submassifs, and from below the CMCC. The oldest cooling ages in
the CMCC are similar in age to those found near the Miocene peneplain
on the plateaux to the north and the south, but the rocks occur in core of
the synform at much lower elevations. Even though the CMCC supports
high mountain ranges in the footwall of the opposite-facing detachment systems — probably by elastic flexure — its overall loss of lower
crustal thickness may have caused the central CMCC to ‘sink’ to a
lower 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 shown
that parts of the upper mantle below Anatolia are asthenospheric
rather than lithospheric (Spakman et al., 1988; Spakman et al.,
Fig. 11. Topography profiles across the Menderes Massif calculated from swaths across a terrain model. Red is the maximum elevation, black the lateral mean elevation, and blue the
minimum elevation, which also corresponds to local drainage elevation. Locations of geomorphological features are also shown in Fig. 9. Notice difference between Profile 1 (mostly
west of the MM), and Profile 2 (within MM), expressed in Profile 2's higher amplitudes, and an overall trend of decreasing drainage elevation towards the Gediz and Büyük Menderes valleys (Profile 2); a feature not present in Profile 1. Profiles 3 and 4 show a long wavelength decrease in elevation form E to W, which is less clear in Profile 3, because it
captures parts of the eastern Aydın Mountains and Bozdağ range area.
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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 steeper
compared to plateau draining channels (Fig. 11).
1993; Spakman, 1999; Wortel and Spakman, 2000; Şengör et al.,
2003; Faccenna et al., 2006; Berk Biryol et al., 2011; Mutlu and
Karabulut, 2011; Paul et al., 2011; Zhu et al., 2012; Jolivet et al., in
press). Due to the inherently low resolution of mantle tomography
compared to geological data, and the variety of the methods used
(e.g. body wave tomography in older studies versus adjoint tomography in Zhu et al., 2012) the shape of the anomaly is not consistent
across the models, and cannot always be related to crustal structure
with confidence. In general a slow wavespeed region of variable
shape below western Turkey is a common and robust feature in the
topography models, with the Berk Biryol et al. (2011) model probably
being the most detailed.
A number of studies have proposed that the asthenospheric window
below western Turkey (Fig. 15) originated from a tear in the highvelocity material that separates the oceanic lithosphere domains of the
African plate into an Aegean section and a Cyprus section (Dilek and
Sandvol, 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 western
Anatolia, including across the Hellenide–Anatolide transition, and was
connected to the Cyprus slab (Fig. 15). Dilek and Altunkaynak (2009)
on the basis of geochemical arguments, proposed that the Miocene to
Pliocene trend of southward younging alkaline volcanics between Eskişehir and Isparta record the formation of a tear, which they interpret to
have separated the Cyprus slab and the Aegean slab, that prior to tearing
would have constituted a continuous slab from central Anatolia to the
Ionian Sea. While the alkaline volcanic trend is likely to record a significant lithospheric feature, this model does not account for the eastern
termination of the Aegean slab below western Turkey as imaged by
most tomography models. An alternative to tearing a continuous slab
could be that the window reflects a primary lithospheric feature, for
example thick buoyant sub-lithospheric mantle domains within the
Anatolides, 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 upper mantle architecture,
and to assess how a mantle-scale discontinuity across the Aegean
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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 to
plateau draining channels (Fig. 11).
coastline of Anatolia relates to crustal structure, we have produced gravity anomaly and crustal thickness models across this area, and have produced a three-dimensional representation of earthquake hypocentres
and the MIT-P08 seismic tomography model (Li et al., 2008).
4.1. Geophysical evidence
4.1.1. Gravity anomaly and Moho depth
We analysed Sandwell and Smith's 1′×1′ free gravity anomaly grid
(Sandwell and Smith, 2009) and transformed it into a conventional
Bouguer anomaly (Fig. 16a) using bathymetry and topography data
from Smith and Sandwell (1997). We concentrated our modelling efforts
on Bouguer anomalies with wavelengths longer than 10 km, since
shorter wavelength anomalies are masked by the lateral resolution of
the original data and do not capture the gravity signatures of mantledepth heterogeneities. We used these Bouguer anomaly data to estimate
Moho 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; Mutlu
and Karabulut, 2011). We inverted the gravity data for the relief of the
Moho 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 regular
mesh of 5 km-wide cells. These parameters provided a model of depths
that range between 18.6 and 40.5 km. The gravity response of this interface shows regional gravity anomaly (Fig. 16) and fits the data at a standard deviation of 15.26 mGal, which is well above the reported precision
of the satellite gravity data ranging 1.8 to 3.6 mGal (Sandwell and Smith,
2009). Residual gravity anomalies not justified by our model (Fig. 16c) are
likely to be caused by crustal-depth heterogeneities.
4.1.2. Earthquake hypocentres
Earthquake data include magnitude and hypocentre location of
13,935 events listed in the USGS PDE catalogue (USGS, 2011) between
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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 Menderes
Massif 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).
19 April 1973 and 14 June 2011, and also including data from events
that occurred before 1973 but were not systematically recorded. Data
that lacked either magnitude or depth information were not considered. We subdivided the earthquake locations into different domains,
according to the spatial distribution within the study area.
4.1.3. 3D model of seismic tomography and earthquake hypocenters
For the three-dimensional model, we projected earthquake data and
seismic velocity anomalies in Cartesian coordinates in UTM Zone 35 N
using PARADIGM™ SKUA© v.2009.3. Seismic tomography data are
subsampled from MIT global depth-corrected p-wave velocity anomaly
Fig. 16. Bouguer gravity anomaly of western Anatolia (a) modelled from satellite gravity and topography data (Sandwell and Smith, 2009); (b) filtered for low frequencies; and
(c) high frequency residual. The data show a long wavelength positive anomaly (b) in the southern Aegean Sea region that becomes weaker to the north. A pronounced negative
gradient can be seen towards central Anatolia, where a slight ridge below the Central Menderes Metamorphic Core Complex separates negative anomalies. Short wavelength data
(c) show north to northeast oriented linear trends broken up by negative anomalies in the graben systems surrounding the Stage 2 core complex and along the northern margin of
the Stage 1 core complex.
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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 Moho
modelled 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 the
grabens 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).
dataset MIT-P08 (Li et al., 2008) and interpolated with SKUA's Discrete
Smooth Interpolator algorithm from 1568 data points in UTM zone 35 at
a resolution of ca. 78 km in N–S (seven rows per depth layer), ca. 61 km
in E–W (eight columns per depth layer); and 45.2 km depth resolution
(28 layers). Earthquake data include all events listed in the previous
section.
4.2. Results
A long wavelength, gravity low is located in the Aegean Sea region; it has a steep gradient parallel to the Aegean coastline, and a
much shallower gradient across strike to the north. Crustal thickness
increases from below 20 km in the Aegean to ca. 40 km at the eastern
margin of the Menderes Massif (Figs. 17 and 18). Short wavelength
gravity anomalies reveal linear corrugation trends that are consistent
with folded basement ridges and elongate Miocene basins covering
the Stage 1 core complex footwall, particularly in the northern part
of the Menderes Massif. These linear trends are offset by negative
anomalies of the grabens that formed above the Stage 2 detachment
faults 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; Sodoudi
et al., 2006; Zhu et al., 2006b) and show how extensional tectonics
caused exhumation of denser and faster material to shallower crustal
levels, predominantly in the eastern Aegean but also in western Turkey.
The pattern of earthquake distribution can be used to define distinct domains of seismic activity (i) in the north around Simav and
(ii) within the Menderes Massif south of Simav, west of the Menderes
Massif, 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 and
west, earthquakes occur in the shallow crust, with the mean depth
being shallower in the Simav domain (9.7 km) compared to the western domain (11.9 km) and the Menderes (11.2 km) domain. Our data
show the Menderes Massif as a distinct, seismically relatively quiet
area with a crustal thickness of about 30 km (Fig. 17). Relatively
few seismic events have occurred across the Menderes Massif, except
for activity along east to southeast trending graben systems (Figs. 17,
18, and 19). Overall, seismic activity shows strong spatial correlation
with the western and northern margin of the Stage 1 core complex and
with Anatolia's western and southern coastline. The spatial distribution
changes markedly across a NE-trend along the western margin of the
Menderes Massif (Figs. 17 and 18). There are also a number of shallow
and deep crustal earthquake clusters. In a significant deviation from the
overall depth pattern earthquakes occur at great depths below the Gulf
of Gökova separated by a gap from an even deeper cluster below the
Dodekanes (down to 176 km), in a pronounced, sheet-like cluster that
gets 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 wedge
that below the Menderes Massif connects to asthenospheric mantle
in the southeast of the model (Fig. 20). The shape of the slab is a robust feature that has been consistently imaged by virtually all tomography studies and has been interpreted by van Hinsbergen et al.
(2005) to record continuous subduction of the African plate since
the Mesozoic. The location and depth of subducting African lithosphere is also consistent with seismic receiver function imaging
(Sodoudi et al., 2006). The slab has a marked edge in the upper
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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 outlined
in 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 domains 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 WGS1984
UTM zone 35 N (in thousands of metres).
mantle below western Turkey but continues eastward for 0.5%, 0.6%,
and 0.7% p-wave velocity anomaly contours. The surface representation of the slab edge as based on the MIT-P08 model, differs somewhat from Berk Biryol et al. (2011), who see the slab continuing
slightly further to the east (Fig. 15).
Previous studies have suggested that the Benioff zone of the Aegean
subduction zone may extend to the Turkish coastline (Papazachos et al.,
2000; Sodoudi et al., 2006). While this may be the case for the
Dodekanes area, the deep earthquake cluster below the Gulf of Gökova
does not correlate with any positive seismic velocity anomaly in the
MIT-P08 model that would indicate the presence of a slab (Fig. 18),
and the deeper earthquakes occur in a region for which most other seismic tomography models do not show a steep E–W oriented slab that
could be interpreted as a Benioff zone. A possible exception is the Zhu
et al. (2012) model where anomalies in the vertically polarised component of the shear wave speed outline a gradient parallel to the Gulf of
Gökova in the 275 km depth slice.
The Gökova deep earthquake cluster bears strong resemblance in
shape, size and depth extent, to a cluster in the south-eastern
Carpathians, which has been interpreted as a strain pattern of continental delamination (Fillerup et al., 2010), or of a Rayleigh–Taylor type instability within continental lithosphere (Lorinczi and Houseman, 2009).
We would argue that the location and shape of the Gökova earthquake
cluster suggests a similar scenario where deep earthquakes occur in a
steep sheet of delaminated or otherwise detached continental crust.
5. Tectonic synthesis
We have reviewed the structure of the Menderes Massif in the light
of new and published geological and geophysical data. The picture that
is emerging is a snapshot of an extending orogenic system, situated in a
laterally inhomogeneous convergent geodynamic setting. The challenge
in understanding the tectonic evolution of southwest Turkey lies in
constraining how closely the dynamics inferred from mantle structure
can be related to the evolution of geological structure. The datasets we
have produced in this study provide new detail on western Anatolia's
lithospheric structure, highlight pronounced differences to the Aegean,
and support our claim of the existence of the West Anatolia Transfer
Zone (WATZ), a transtensional structure that has originated from a
wide lithosphere scale transtension zone that denuded the Menderes
Massif in the Miocene. We argue that the İzmir–Balıkesir Transfer
Zone, the structural corridor that limits the western exposure of the
Menderes Massif, is the current upper crustal expression of the lithospheric mantle-driven transtension across the Western Anatolia Transfer zone.
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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 that
range 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 the
southern 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 from
north 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 domain
having the shallowest mean depth and also a lower mean magnitude than the Western domain (also see Figs. 16 and 17).
Table 2
Summary statistics of earthquake depth and magnitude.
5.1. Lateral differences in lithospheric structure
Depth
Domain
n
Min [m]
Max [m]
Mean [m]
Median [m]
Standard
deviation
Simav
South
Menderes
(MM)
West
2581
2086
670
1000
2000
1000
60,000
170,000
58,000
9747
28,044
11,176
10,000
13,000
10,000
3894
35,342
6957
6656
1000
176,000
11,912
10,000
7734
Domain
n
Min
Max
Mean
Median
Standard
deviation
Simav
South
Menderes
(MM)
West
2581
2086
670
1.70
2.40
2.50
7.30
7.10
6.50
2.86
3.55
3.40
2.70
3.40
3.20
0.44
0.50
0.56
6656
1.90
7.50
3.11
3.00
0.48
Magnitude
Below the Aegean Sea, the north-subducting slab is imaged by both
the gravity data (Fig. 16) and by the P-wave velocity model of the mantle
(Figs. 17 and 20). We envisage that in the late Oligocene/early Miocene
the Hellenide–Anatolide boundary was expressed as the lateral transition from dense oceanic lithosphere in the Aegean to the thickened continental root below the Anatolide belt. It is difficult to assess the tectonic
record of this transition, but it is likely that it represents a former domain
boundary (e.g. a passive margin) between the Adriatic and Anatolian domains, and was already tectonically active during late Mesozoic convergence, when the NNE-trending Bornova Flysch Zone (Fig. 5) formed in a
transtensional setting (Okay, 2011). We argue that the continuing
southwest-ward rollback of the Aegean slab by the late Oligocene/early
Miocene initiated the West Anatolia Transfer Zone (WATZ) either by localization across a lithosphere scale material boundary, or by reactivation
of an earlier feature. We would argue that rollback turned the north–
south oriented trailing edge of the Aegean slab into a rigid transtensional
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boundary relative to western Anatolia's sub-continental lithosphere. This
kinematic framework caused tearing at the western and eastern boundaries of the continental lithosphere and delamination of large parts of the
subcontinental lithospheric mantle accompanied by uplift, crustal extension, and alkaline magmatism (Figs. 20, 21, and 22).
There is no evidence for a presently subducting slab below the
Menderes Massif in the MIT-P08 tomography model. Tomographic
data 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 Biryol
et al., 2011; Paul et al., 2011; Zhu et al., 2012; Jolivet et al., in press).
Following the STEP fault concept (Govers and Wortel, 2005), most
studies have proposed that asthenospheric material was emplaced
along a tear fault that ruptured a previously continuous slab to accommodate the difference between the fast rollback of the Aegean
slab section versus the slow rollback of the Cyprus slab section
(Govers and Wortel, 2005; Dilek and Altunkaynak, 2009). This
would require that the lithospheric mantle below the Menderes Massif would have been either no different from the Aegean one, or — if it
was different — would have been sheared off during convergence, as
suggested by proponents of flat slab hypotheses (Westaway, 2006;
Prelevic et al., 2010b; van Hinsbergen et al., 2010b). In the latter
case it is unclear where the sheared-off autochthonous continental
mantle material would be now, and how it could be imaged. If the
ca. 300 km wide asthenospheric window formed within a previously
continuous flat slab in the Miocene (van Hinsbergen et al., 2010b),
the western and eastern edges of the slab (Fig. 15), now separated
by ca. 250 km, would have needed to experience an E–W component
of 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 evidence for an E–W difference within this range in geodetic measurements (Fig. 9), but there is also no record of any crustal deformation
that could relate to these kinematics. Furthermore, a gradually widening tear in a continuous slab should have produced a pattern of
symmetrically outward migrating alkalic magmatic rocks in the Miocene, rather than the observed southward progression focused along
the eastern and western edges of the asthenospheric window
(Fig. 15).
There are considerable differences between the Aegean Sea region
and the Menderes Massif, particularly regarding lithospheric evolution. In the light of Archaean zircon ages produced from magmatic
and metamorphic rocks (Kröner and Sengör, 1990; Candan et al.,
2011; Zlatkin et al., 2012), we regard it possible that the lithosphere
below the Menderes Massif may have been considerably older, and
thicker than Aegean lithosphere prior to its Miocene delamination.
We propose a scenario where the Menderes arrived into the convergence zone containing thick buoyant continental lithosphere with a
passive margin sequence towards the north, and possibly to the
west. Upon the collision, the continental lithosphere thickened, and
at some stage the leading, oceanic lithospheric domain to the north
decoupled (van Hinsbergen et al., 2010b). The thick lithosphere
below the Menderes then became mechanically unstable either by
an increase in temperatures (Houseman et al., 1981; England and
Houseman, 1989; Molnar et al., 1993; Platt and England, 1993;
Houseman and Molnar, 1997; Stern et al., 2006), due to advection of
asthenospheric mantle where the Aegean slab decoupled, or as a result of heterogeneity in plastic strength (Gorczyk et al., 2012) causing
the thick lithosphere below the Menderes to delaminate (Fig. 22).
Such a model would be similar to scenarios proposed for the
Carpathians (Lorinczi and Houseman, 2009), the North American
Great Basin (West et al., 2009), New Zealand's North Island (Stern
et al., 2006), Eastern Anatolia (Gögüs and Pysklywec, 2008), and
intraplate orogenic systems in general (Gorczyk et al., 2013). As
many studies have pointed out before, southward progressing removal of lithospheric mantle below western Turkey is consistent with the
change 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 the
Miocene southward progression of alkaline volcanic rocks both
along the Anatolian west coast and along the Afyon–Isparta line
track the edges of the delaminated Anatolian crustal domain.
5.2. Sinistral transtension across West Anatolian Transfer Zone as a driver
for Menderes extension
The current strain field shows a clear westward gradient to higher velocities, 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 rollback and removal of the lithospheric mantle, could have operated
through time. In general the geodetic measurements of the Hellenic subduction zone suggest a movement of ca. 33 mm/a (Reilinger et al., 2006),
which is at the lower range of what Stegman et al. (2006) and Schellart et
Fig. 20. Three-dimensional model of P-wave velocity (dVp) anomaly contours and
hypocentres in the eastern Aegean and western Anatolia from the surface to a depth
of ca. 1250 km. Land surface (white, transparent) is shown for reference. The
north-dipping slab (a) is discontinuous in the upper mantle (b), but continues eastward for 0.5%, 0.6%, and 0.7% dVp contours (white arrows). Slow, hot mantle above
the slab is represented as a negative dVp anomaly in the Aegean (a) that is connected
to a vertical anomaly southeast. In the south, a westward deepening cluster of earthquake hypocentres in normal velocity lithosphere reaches depths of 176 km. These
data are also presented in 3D PDF format (cf. supplementary materials 2 and 3).
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
267
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 by
the 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.
al. (2007) consider characteristic for narrow oceanic slabs, but very close
to the ca. 30 mm/a for decoupled continental collision systems in
Faccenda et al. (2009). In the Aegean, fast displacement rates have
been proposed for individual normal fault systems: ca. 20 mm/a along
the Cretan detachment (Ring and Reischmann, 2002) and 6.5 mm/a
along the Vari detachment on Syros and Tinos. Gessner et al. (2001a,
2001b, 2001c) and Wijns et al. (2005) estimated displacement along
the Kuzey detachment at 2 mm/a. The Aegean displacement rates are
typical for fast displacement rates in metamorphic core complexes, the
Kuzey detachment would be typical for slower displacement rates
(Gessner et al., 2007). The velocity of southward delamination of lithospheric mantle below western Turkey is difficult to constrain directly,
but can be approximated by the progression of alkaline magmatism in
western Anatolia. Data compiled by Dilek and Altunkaynak (2009) and
Ersoy et al. (2010) imply that Miocene volcanic activity is limited by
both the eastern edge of the Aegean slab and the western edge of the
Cyprus slab, and that this activity progressed southward at a rate of ca.
10–15 mm/a. Even though these are only rough estimates, these rates
suggest that the southward retreat of the Hellenic subduction zone
took place at least twice as fast as delamination progressed in western
Anatolia. Therefore, we would argue that the formation of the West Anatolian Transfer Zone caused very different Miocene developments on either side: rapid rollback of the dense lithosphere of the Adriatic plate in
the Aegean Sea region triggered a surge of lithospheric extension in the
mid Miocene. In western Anatolia at the same time a plateau with associated (Yilmaz et al., 2000), and increasingly alkaline volcanic activity
formed in the northern Menderes (Seyitoglu et al., 1997; Pe-Piper and
Piper, 2007; Ersoy et al., 2008; Dilek and Altunkaynak, 2009; Prelevic et
al., 2010a; Altunkaynak et al., 2012a, 2012b), while the Lycian nappes
were thrust onto the Bey Dağları foreland (Fig. 5) to the south (Collins
and Robertson, 1998). Such close proximity of crustal shortening and extension in conjunction with alkali magmatism is also known from eastern Anatolia (Şengör et al., 2003; Gögüs and Pysklywec, 2008), and the
North Island of New Zealand (Stern et al., 2006); areas in convergent settings for which removal of lithosphere has been proposed. Transtensional
kinematics could provide an explanation for folding about NE–SW to
NNE–SSW axes that is most prominent in the northern Menderes
(Bozkurt, 2003; Cemen et al., 2006), but has also been described from
other parts of the Menderes (Schuiling, 1962; Rimmelé et al., 2003;
Regnier et al., 2006) and may be the reason for the sinusoidal basement
topography in the Alaşehir–Gediz graben system (Çiftçi and Bozkurt,
2009a, 2010). Folding parallel to extension at the observed NNE–SSW
fold axis orientations and wavelengths of ca. 20–40 km may have
recorded the initial elastic response during Stage 1 crustal extension.
Thinning of the crust parallel to stretching could therefore partially have
been countered by shortening in a perpendicular direction, hence crustal
thickness in this case would be no adequate measure of crustal stretching,
as for example implied by Zhu et al. (2006b). Folding caused uplift of
basement in the anticlines, while providing accommodation space for
the Miocene basins in the synclines. The relation between basement
topography and fault segment length in the Alaşehir–Gediz graben system suggests that transtensional folding may have overlapped with the
early stages of tectonic denudation of the CMCC. This means that the
Menderes Massif has experienced northeast–southwest stretching before
268
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
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 situation 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 core
complexes displaying frozen in mid-crustal levels of transtensional deformation.
north–south fragmentation by the Kuzey and Güney detachment systems, and that overall extension in western Anatolia may be greater
than estimated by the displacement of rigid blocks. The previous estimate
of 150 km extension (van Hinsbergen, 2011) is mainly based on the assumption that the Simav detachment footwall experienced little or no internal deformation, before being fragmented by the Kuzey and Güney
detachments.
While earthquake distribution patterns support the fragmentation
of western Anatolia into different domains (Figs. 18 and 19) geodetic
measurement shows that the overall strain is very homogeneous
(Aktug et al., 2009; Özeren and Holt, 2010). This discrepancy may be
explained to some extent by the variation of rock composition in the
study area, where heterogeneous rock units such as the tectonic melanges of the Bornova Flysch and the Tavşanlı Zone may respond to ductile flow of the lower crust by fracturing more often and in a more
distributed manner than the metamorphic rocks of the Menderes Massif. It is also clear that, compared to the current situation, transtension
across the WATZ would have played a much larger role during the Miocene, when transtension affected the entire area of the thermally weakened Menderes Massif lithosphere. Structures like the İzmir–Balıkesir
Transfer Zone (Uzel and Sozbilir, 2008) — which we argue would be
the present day upper crustal expression of the West Anatolia Transfer
Zone — may at present only constitute a rheological boundary in a wide
zone of bottom-driven crustal deformation between the North Anatolian Fault Zone and the Hellenic Trench. That the WATZ does not cut
the surface as one distinct fault zone is likely to be the consequence of
the overall rheological stratification of the continental lithosphere in
the Aegean and western Turkey. The weak lower crust constitutes a
thick viscous layer that mechanically decouples the brittle-elastic
upper crust from the bottom-driven kinematics, such that the only
thing the upper crust ‘feels’ is a westward increase in crustal extension
(Fig. 9) resulting in a thinner crust and shallower Moho in the Aegean
relative to the Menderes area.
5.3. Continuous versus punctuated crustal extension
The geological evidence points to a two-stage denudation of the
Menderes Massif, with movement on the Simav detachment slowing
down or stopping around ca. 19–16 Ma (Ring and Collins, 2005;
Thomson and Ring, 2006; Hetzel et al., 2013). Synkinematic granites
that were exhumed in the footwall of the Simav detachment were probably intruded in conjunction with asthenospheric flow caused by
decoupling of the leading part of the African slab (van Hinsbergen et al,
2010b). We envisage a scenario where footwall up-doming at the scale
of the Menderes Massif caused secondary late-stage top-S extension at
the 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 late
Oligocene and early Miocene in the Çine submassif. However, field evidence for a well-developed extensional detachment system is lacking
(Ring et al., 2003a). The steep gradient in apatite fission track ages between the Cycladic Blueschist Unit and the Afyon/Ören unit in Fig. 8
could be explained by either a top-S extensional reactivation of the
basal thrust of the Afyon/Ören unit, but also by footwall uplift below
the original Simav–Alaçam detachment system (Fig. 23), if indeed it extended this far south.
We propose that the development of a wide, distributed West
Anatolia Transfer Zone caused the lithospheric extension in the
Menderes Massif and the dome-shape of the evolving core complex
(Fig. 23). Extension waned by 19–16 Ma in the northern part of the
massif. In the central Menderes Massif granites intruded at 16–
15 Ma (Glodny and Hetzel, 2007) and the basin fill suggests ongoing
extension in the mid Miocene (Çiftçi and Bozkurt, 2010). Modest
footwall cooling by this time (Ring et al., 2003a, 2003b) suggests limited extensional activity by this time. Continuing uplift of the evolving
Stage 1 Menderes core complex led to the formation of a plateau with
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
269
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 southern 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 detachment 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 and
Gördes submassifs (Fig. 9); the CMCC footwall ‘inherited’ corrugations as topographic features that control the orientation of drainage in the Aydın and Bozdağ mountains (Figs. 12
and 13).
an associated peneplain (Yilmaz et al., 2000). Our surface topography
analysis supports the hypothesis that this peneplain has been significantly modified by the movement along the Kuzey and Güney detachments and the high-angle normal fault systems in the Gediz and
Büyük Menderes grabens since the Pliocene (Gessner et al., 2001b).
Whether or not the early Miocene extensional pulse was clearly
separated in time from the distinct Pliocene to Recent extensional
activity is difficult to assess. It is conceivable that the two events
represent a continuum of lithospheric extension that commenced in
the early Miocene and then slowed down. As a result of the denudation and cooling of the Menderes Massif, transtension across the
WATZ changed character from the wide transtensional deformation
of thermally weakened crust to a more focussed structural corridor
at the western margin of the massif.
5.4. Open questions
We propose that sinistral movement along the boundary of the
Aegean and Anatolian domains played a much larger role in the tectonic
evolution of the Anatolian peninsula than previously proposed, particularly with regard to the extensional deformation in the Menderes Massif, but also with respect to the distribution of seismic hazard, and the
structural control on hydrothermal metallic resources and geothermal
reservoirs (Yigit, 2006; Faulds et al., 2009; Yigit, 2009; Gessner et al.,
2010). The West Anatolia Transfer Zone is a wide and diffuse lithospheric deformation zone that has localised at the western margin of one
of its earlier products, the Menderes Massif, and constrained the
shape and location of Anatolia's Aegean coastline. Despite large advances in understanding this rapidly deforming region, there is still a
considerable lack of detailed knowledge, for example on how deformation has partitioned across the WATZ over time, how extensional strain
has been accommodated at the flanks of the Menderes Massif, and how
the crustal structure changes toward the east. There is a need to better
resolve crustal and mantle structure below western Turkey to increase
our knowledge of three-dimensional architecture and to further constrain the processes that have lead to the low velocity anomaly below
the Menderes Massif. Our review shows the significance and benefit
of integrating geological and geophysical data in three dimensions to arrive at a better understanding of lithospheric structure and tectonic
evolution.
6. Summary points
• The Hellenide and Anatolide domains of the Tethyan orogen can be
defined on the grounds of their geological history that encompasses
differences in the age of pre-Alpine basement rocks, as well as in
structure, metamorphic and magmatic history related to continental
subduction and crustal extension.
• The lithospheric mantle across the Hellenide and Anatolide domains
is heterogeneous. Seismic velocity anomalies show a sharp vertical
boundary between the fast, cold and dense slab below the Aegean
and a slow, hot and buoyant asthenospheric region below western
Turkey. Gravity data show a north–south oriented boundary between
a high in the Aegean and lower gravity values below the Menderes
Massif, and towards the Anatolian plateau to the east.
• We propose that geological differences between the Hellenide and
the Anatolide domains are closely related to the discontinuity in the
lithospheric mantle. We interpret this discontinuity as a lithosphere
270
•
•
•
•
•
•
K. Gessner et al. / Gondwana Research 24 (2013) 243–274
scale shear zone, the West Anatolia Transfer Zone (WATZ), which has
accommodated the difference between fast roll back of the slab in the
Aegean and slow delamination of the Anatolian continental lithosphere since the Miocene, and links the North Anatolian Fault zone
to the Hellenic trench.
The lack of oceanic lithosphere below western Anatolia's upper mantle together with the lack of high-pressure metamorphism supports
the hypothesis that thick buoyant continental lithosphere of the Anatolian microplate brought about an end to continental subduction in
western Turkey in the Eocene.
We link the Late Oligocene/Early Miocene to recent crustal extension
in the Anatolide belt in western Turkey to sinistral transtension
across the WATZ. Within this kinematic framework the Menderes
Massif, of one of Earth's largest metamorphic core complexes, has experienced NNE–SSW extension, including extensional detachment
faulting, folding parallel to extension, doming, and footwall uplift.
The Menderes Massif appears to be coincident with a Miocene peneplain, but since the Late Miocene has been fragmented by E–W and
WNW–ESE trending graben systems. We propose that fragmentation
of the plateau has driven flow in the weak lower crust towards the
Central Menderes area, causing a dynamic topographic response
across the fracture system in the plateau and in the footwall of Miocene to Pliocene graben bounding detachment faults.
Earthquake locations in western Anatolia strongly correlate with the
spatial distribution of tectonic units. Seismic activity in the Menderes
Massif is lower than in adjacent regions, but it strongly localises along
graben systems.
Most earthquakes occur at shallow to mid-crustal depths, with the
notable exception of a narrow E–W oriented zone of very deep earthquakes below the Gulf of Gökova. We tentatively link this steep seismic zone to delamination of continental lithosphere.
Our findings highlight the significance of lateral variations in evolving continental arcs for the structure of orogenic belts, particularly
with respect to the formation of metamorphic core complexes.
Acknowledgements
K. Gessner wishes to acknowledge funding by a 2010 University of
Western Australia Professional Development Award, the Australian
Research Council (LP100200785), and support by Ariana Resources
Pty. Ltd. and the ARC Centre of Excellence for Core to Crust Fluid
Systems (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 discussions and comments on previous versions. Y. Dilek and an anonymous
reviewer are thanked for formal reviews that helped to improve the
manuscript; M. Santosh and T. Horscroft are thanked for outstanding
editorial support, and patience.
Appendix A. Supplementary data
The supplementary material consists of three 3D PDF pages in one file.
Notice that the 3D PDF models, which contain earthquake locations and
p-wave anomaly contours — similar to the content shown in Fig. 20 —
allow a range of interactions, including the selection of objects in the
model tree. Explicit instructions are given in Supplementary Fig. 1.
Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.gr.2013.01.005.
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Klaus Gessner is the 3D Geoscience Manager at the Geological Survey of Western Australia, and Adjunct Research Fellow at The University of Western Australia. He received a
‘Diplom’ in Geology–Palaeontology from Johann Wolfgang
Goethe University in Frankfurt, Germany (1996), and a PhD
from Johannes Gutenberg University in Mainz, Germany
(2000). Klaus has worked as a structural geologist at CSIRO
Exploration and Mining, and has taught at The University
of Western Australia. His research focus is on the structural
evolution of Phanerozoic, Proterozoic and Archaean orogenic
belts in Turkey, Australia, and New Zealand, and on the processes involved in hydrothermal mineral deposits and geothermal systems.
Luis A. Gallardo is a titular researcher in the Department
of Applied Geophysics at CICESE, Mexico; a National Scientist from the National Council of Science and Technology in
Mexico; and an Adjunct Scientist of the Centre for Exploration Targeting at the University of Western Australia. He
obtained an MSc in Applied Geophysics from CICESE
(1997), a PhD in Environmental Science from Lancaster
University, UK (2005) and held the Goodeve Lectureship
at The University of Western Australia from 2009 to
2011. Luis' research focuses on geophysical inverse theory
and the joint inversion of gravity, electromagnetic and
seismic 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 environmental and geotechnical applications.
Vanessa Markwitz is a Research Fellow at the Centre for
Exploration Targeting at The University of Western Australia in Perth. She graduated from Johann Wolfgang Goethe
University in Frankfurt, Germany with a ‘Diplom’ in Geology–
Palaeontology (1996). Vanessa has carried out structural
research on the Rhenish Massif in Germany as a contract geologist for the Geological Survey of Rhineland Palatinate in
Mainz, Germany. Vanessa is a GIS specialist at the Centre for
Exploration Targeting and has worked on uranium, nickel
and gold prospectivity in Australia, Zimbabwe, West Africa,
and Turkey.
Uwe Ring is a Professor in the Department of Geological
Sciences at Stockholm University in Sweden. He obtained
a ‘Diplom’ in Geology from the Technische Hochschule
Darmstadt, Germany (1985), and a PhD in Geology from
Eberhard-Karls-Universität in Tübingen, Germany (1988).
Uwe was a Humboldt Fellow at Yale University in the
USA, has taught Geology at Johannes Gutenberg University
in Mainz, Germany, and at the University of Canterbury in
Christchurch, New Zealand. His research interests include
continental extensional tectonics in Greece, Turkey, and
the East African Rift. Uwe also has worked on the exhumation of metamorphic rocks from great depths and on the
interactions between climate and tectonics.
Stuart N. Thomson is a Research Scientist at the University of Arizona. He obtained a BSc in Geology from Durham
University, UK (1988) and a PhD in Geology from University College London (1993). His research focuses primarily
on the application of geochronology and low temperature
thermochronology to problems in geology, tectonics,
structural geology, and tectonic geomorphology. He has
worked on projects around the world including on Cenozoic glacial–tectonic interactions in the South American
Andes, various projects on Mediterranean tectonics in Italy,
Greece, and Turkey, to methodological work on apatite
U–Pb dating.