Miocene–Recent evolution of the western Antalya Basin and its linkage with the Isparta Angle, eastern Mediterranean
Interpretation of ~ 9500 km of multichannel seismic reflection profiles showed the presence of two major tectonic histories in western Antalya Basin, spanning from the Miocene (or older) to the Pliocene–Quaternary. A prominent fold–thrust belt affects the Miocene succession in the offshore. The thrusts swing from a NW–SE strike, with SW vergence, in the northeast of the mapped area to a more northerly strike, with westerly vergence at the western shelf edge of the deep basin. The Miocene deformation appears to continue offshore from the westerly-directed thrusting seen onshore that characterizes the eastern edge of the Isparta Angle. The contraction is consistent with the counterclockwise Miocene rotation of the western side of the Isparta Angle determined from paleomagnetic studies. The thrust belt forms the western extremity of the wider regional Aksu–Kyrenia–Misis oroclinal culmination. The tectonic activity experienced a period of relative quiescence across the western Antalya Basin during the Messinian. A major kinematic change occurred in the Pliocene, when the regional strain was partitioned into three spatially localized tectonic domains: an extensional domain confined to the Pliocene–Quaternary Unit 1, occupied the northeastern portion of the study area, a predominantly extensional domain with a few re-activated pre-existing Miocene contractional structures occupied the southern and central portion of the study area, and an extensional and/or transtensional domain occupied the continental shelf and slope in the westernmost Antalya Basin. These complexities relate to strike-slip motions as the crustal block within the Isparta Angle moved northwards relative to the blocks to the north.
Miocene–Recent evolution of the western Antalya Basin and its linkage
with the Isparta Angle, eastern Mediterranean
⁎, A.E. Aksu
, H. King
, A. Gogacz
, G. Çifçi
Department of Earth Sciences, Centre for Earth Resources Research, Memorial University of Newfoundland, St. John's, Newfoundland A1B 3X5, Canada
Department of Geological Engineering, Faculty of Mines Istanbul Technical University, Ayazağa, Istanbul 34426, Turkey
Institute of Marine Sciences and Technology, Dokuz Eylül University, Haydar Aliyev Caddesi No: 10, İnciraltı,İzmir 35340, Turkey
Received 26 July 2013
Received in revised form 11 December 2013
Accepted 16 December 2013
Available online 31 December 2013
Communicated by: D.J.W. Piper
western Antalya Basin
Interpretation of ~9500 km of multichannel seismic reﬂection proﬁles showed the presence of twomajor tectonic
histories in western Antalya Basin, spanning from the Miocene (or older) to the Pliocene–Quaternary. A prominent
fold–thrust belt affects the Miocene succession in the offshore. The thrusts swing from a NW–SE strike, with SW
vergence, in the northeastof the mapped area to a more northerly strike, with westerly vergence at the western
shelf edge of the deep basin. The Miocene deformation appears to continue offshore from the westerly-directed
thrusting seen onshore that characterizes the eastern edge of the Isparta Angle. The contraction is consistent
with the counterclockwise Miocene rotation of the western side of the IspartaAngle determined from paleomag-
netic studies. The thrust belt forms the western extremity of the wider regional Aksu–Kyrenia–Misis oroclinal
culmination. The tectonic activity experienced a period of relative quiescence across the western Antalya Basin
during the Messinian. A major kinematic change occurred in the Pliocene, when the regional strain was partitioned
into three spatiallylocalizedtectonic domains: an extensional domain conﬁned to the Pliocene–Quaternary Unit 1,
occupied the northeastern portion of the study area, a predominantly extensional domain with a few re-activated
pre-existing Miocene contractional structures occupied the southern and central portion of the study area, and
an extensional and/or transtensional domain occupied the continental shelf and slope in the westernmost Antalya
Basin. These complexities relate to strike-slip motions as the crustal block within the Isparta Angle moved north-
wards relative to the blocks to the north.
© 2013 Elsevier B.V. All rights reserved.
Orogenesis is one of themost fundamental of Earth processesand is
responsible for most of the relief that we see in the World today, includ-
ing the Alps, the Rocky Mountains and the Appalachian Mountains
(Moores and Twiss, 1995). Similarly, deep arcuate oceanic trenches
observed today adjacent to island arcs and continents are also products
of orogenesis, where one oceanic lithospheric plate is forced to plunge
beneath the continental or oceanic lithosphere of another plate,
depressing or uplifting the overriding plate edge. This study focuses
on the geologically recent evolution of an orogen, being caused by the
collision between the African and the Eurasian continental plates and
the squeezing and shufﬂing of the smaller microplates and continental
fragments in the eastern Mediterranean. Speciﬁcally, it is focused on
the Miocene to Recent tectonic and sedimentary evolution of the
western Antalya Basin, which in a larger plate tectonic context, is a
forearc basin north of the African Plate–Aegean–Anatolian microplate
boundary (Şengör et al., 1985; Dewey et al., 1986). Here the subduction
has possibly ceased. Papazachos and Papaioannou (1999) showed the
presence of a Benioff zone below the western Antalya Basin, but only
diffuse seismicity farther east as evidence for the tearing of the
subducting African Plate. Biryol et al. (2011) suggest that slab break
off has also occurred here, detaching the remnants from the still-
subducting slab below the Hellenic Arc along a Subduction Transform
Edge Propagator (‘STEP’)fault(Govers and Wortel, 2005). As such,
the study area is an excellent modern laboratory for the understanding
of the processes that govern the deformation during the early stages of
slab break off and the ultimate transition to continent to continent
collision, which is largely hidden in ancient orogenic belts, such as the
Appalachian Mountains of eastern North America. During the last
~20–25 Ma, the forearc experienced profound tectonic changes when
former marine basins, such as the Aksu, Köprüçayand Manavgat basins,
were uplifted to become nestled on the foothills of the evolving Central
and Western Taurus Mountains, while the deep Antalya Basin experi-
enced complementary subsidence and marine sedimentation.
The main focus of this paper is the interpretation of high-resolution
multichannel seismic reﬂection proﬁles collected during four Memorial
University of Newfoundland–Dokuz Eylül University research cruises in
1992, 2001, 2008 and 2010 from the Antalya Basin and environs in the
eastern Mediterranean, complemented by industry seismic reﬂection
proﬁles. The primary scientiﬁc objectives of this paper are: (i) to establish
Marine Geology 349 (2014) 1–23
⁎Corresponding author.Fax: +1 709 864 2589.
E-mail address: firstname.lastname@example.org (J. Hall).
0025-3227/$ –see front matter © 2013 Elsevier B.V. All rights reserved.
Contents lists available at ScienceDirect
journal homepage: www.elsevier.com/locate/margeo
a seismic stratigraphic framework for the Miocene to Recent successions
observed in the seismic reﬂection proﬁles, and a chronostratigraphy
for these successions using correlations with the litho- and/or bio-
stratigraphic data from an exploration well from the onland Manavgat
Basin; (ii) to delineate and map the structural elements affecting the seis-
mic stratigraphic units and to determine the age of the deformation using
growth stratal architecture and progressive syn-tectonic unconformities
observed in the seismic reﬂection proﬁles; (iii) to relate the large-scale
tectonic elements mapped within the marine Antalya Basin with their
counterparts in the Isparta Angle and Beydağlarıand Antalya Complex
regions of southwestern Turkey and the Kyrenia Mountains of northern
Cyprus and (iv) to develop a tectonic and kinematic model for the
Miocene to Recent structures of the western Antalya Basin that explains
the evolution of the region within the context of the greater eastern
1.1. Tectonic framework of the eastern Mediterranean
The Cyprus Arc is a large south convex structure in the eastern
Mediterranean (Fig. 1). The western segment of the arc terminates
against a broad transform fault zone which includes three prominent
sinistral strike-slip faults, known as the Ptolemy, Pliny and Strabo
trenches (Mascle et al., 1986). This transform fault zone (a Subduction
Transform Edge Propagator fault, Govers and Wortel, 2005) links the
Cyprus Arc to the Hellenic Arc, below which subduction continues as
the arc rolls back ( e.g., Jolivet and Brun,2010). The Anaximander Moun-
tains (Zitter et al., 2003; Aksu et al., 2009) are enigmatic underwater
highs that are situated at or near the junction between the Hellenic
Arc and the Cyprus Arc (Fig. 1). The Florence Rise–Cyprus Arc–Tartus
Ridge deﬁnes the eastern segment of the oblique convergent boundary
between the African Plate and the Aegean–Anatolian microplate
(Fig. 1). Subduction along this boundarywas initiated in the Late Creta-
ceous as evidenced today by an ophiolitic suture. In the west, the suture
includes the ophiolites of the Antalya Complex and Isparta Angle,
whereas in the east it includes the ophiolites of the Hatay, Kızıldağ,
Baër–Bassit, and Troodos complexes (Fig. 1;Biju-Duval et al., 1978).
Subsequent events in the Eocene and late Miocene shaped an arcuate
fold–thrust belt with major culminations centered on re-imbricated
elements of the ophiolitic suture (Yılmaz, 1993; Hall et al., 2005a,b;
Faccenna et al., 2006; van Hinsbergen et al., 2010). At present, the
east-trending southern segment of the margin is characterized by
contraction in a forearc setting, related to northward subduction of
the African Plate with ensuing collision of the Eratosthenes Seamount
(Fig. 1;Ben Avraham et al., 1995; Robertson, 1998). The northeast-
trending eastern segment is in sinistral transtension along strands of
the East Anatolian Transform Fault, facilitating the westward escape of
the Anatolian microplate (Şengör et al., 1985; Kempler and Garfunkel,
1994). The Neogene marine Antalya and onland Aksu, Köprüçay and
Manavgat basins are situated inboard of the Cyprus Arc (Fig. 1). During
the Miocene, these basins developed in a broad foredeep south and east
of the evolving Central and Western Taurus Mountains. In the late
Miocene, the Aksu, Köprüçay and Manavgat basins experienced a
protracted uplift, while the marine Antalya Basin experienced consider-
able subsidence. A large crustal-scale culmination developed during the
Pliocene–Quaternary, extending from the Aksu thrust system onland
toward the southeast into the Kyrenia Range of northern Cyprus
(Fig. 1;Işler et al., 2005).
Recent studies showed that subduction has ceased along the Florence
Rise–Cyprus Arc, but is continuing along the Hellenic Arc in the west
(Woodside et al., 2002; Govers and Wortel, 2005). Several studies
suggested that a dextral wrench developed along the Florence Rise,
associated with the cessation of subduction along the Cyprus Arc, but
that the contractional deformation continued throughout the Pliocene–
Quaternary (Zitter et al., 2003). In this region the relative motion be-
tweentheAfricanPlateandtheAegean–Anatolian microplate has nearly
come to a halt and the subduction of the northern fringes of the African
Plate along the Hellenic Arc is accompanied by slab roll-back (Govers
and Wortel, 2005). In such land-locked basins the overriding plate
shows back-arc extension in response to the movement of the trench,
such as the north–south extension seen in the western segment of the
Aegean–Anatolian microplate (Robertson, 1998). Another consequence
of subduction along the Hellenic Arc and no subduction along the
Cyprus Arc is the tearing of the lithosphere along transform-parallel
zones. The tearing transform segment along the present-day Ptolemy–
Pliny–Strabo trenches (van Hinsbergen and Schmid, 2012) is referred
to as a Subduction-Transform Edge Propagator, or STEP fault (Govers
and Wortel, 2005).
Thus, the Antalya Basin is situated north of this broad conver-
gence zone delineating the boundary between the African Plate and
the Aegean–Anatolian microplate (Fig. 1). In fact, the zone of deforma-
tion associated with the convergence is very wide, extending in the
marine areas at least, from the Florence Rise–Cyprus Arc–Tartus Ridge,
Fig. 1. Simpliﬁed plate tectonic map of the eastern Mediterranean Sea and surrounding regions, showing majorplate/microplate boundaries, ophioliticrocks (greenﬁll: ac = Antalya com-
plex, bb = Baër-Bassit complex, hk = Hatay and Kızıldağcomplexes, tc = Troodos complex)and major tectonicelements. Ab = Antalya Basin, AKMB = Aksu, Köprüçay, Manavgat ba-
sins, IB = Iskenderun Basin. Half arrowsindicate transform/strike-slip faults.(For interpretation of the references to color inthis ﬁgure legend, thereader is referred to the webversion of
2J. Hall et al. / Marine Geology 349 (2014) 1–23
approximately 300 km toward the north (e.g., Aksu et al., 2005a,b; Calon
et al., 2005a,b; Hall et al., 2005a,b; Işler et al., 2005). This broad deforma-
tion zone is characterized by three prominent south-convex arcuate
zones, which parallel the trend of the Florence Rise–Cyprus Arc–Tartus
Ridge: the Amanos–Larnaka and Misis–Kyrenia–Aksu zones and the
Central Taurus Mountains (Fig. 1). Within this backdrop, the Antalya
Basin emerges as an arcuate forearc basin that appears to curve towards
(i.e., strike into) the onland Isparta Angle. The latter has an extensive his-
tory of nappe emplacement associated with ocean basin closure during
late Cretaceous to early Tertiary time (Robertson et al., 2003). Tightening
of the angle between western and eastern limbs is established from pa-
leomagnetic studies (Kissel and Poisson, 1986; van Hinsbergen et al.,
2007) suggesting continued intermittent convergence up to Miocene
time. Neogene deformation observed in the offshore Antalya Basin
(Işler et al., 2005) is characterized by the Middle to Late Miocene fold–
thrust belt and regionally-partitioned Pliocene to Recent extension/
transtension and transpression. Thus a critical question addressed in
Basin and its onland extension into the area of the Isparta Angle.
1.2. Bathymetry of the eastern Mediterranean Sea
In the east Mediterranean region, the morphology and topography
are controlled by large-scale tectonic features, such as theAnaximander
Mountains, the Florence Rise, the Misis–Kyrenia–Aksu zone, and the
Cyprus Arc in the marine areas and the Isparta Angle, the Taurus Moun-
tains, Kyrenia Range onland (Fig. 2). The Antalya Basin is an embayment
in the eastern Mediterranean (Fig. 2). The continental shelf around the
Antalya Basin is very narrow, ranging between 2 and 6 km. The shelf-
slope break occurs at ~100–150 m depth, and steep slopes lead to the
continental rise and abyssal plain. There is no multibeam data from
the Antalya Basin, but the available bathymetry maps with 200 m
isobaths show that the slope face is dissected by numerous submarine
canyons, presumably feeding submarine fans, similar to those seen in
continental slopes around the western Mediterranean (e.g., Droz et al.,
2001; Lastras et al., 2002). The continental rise occurs between 1800
and 2000 m water depth, where the slope gradient decreases consider-
ably (Fig. 2). The abyssal plain occurs at ~ 2400 m water depth: the
maximum depth is ~ 2600 m, observed as a near-circular depression
in a central location within the Antalya Basin (Fig. 2).
1.3. Marine Miocene basins in the northeastern Mediterranean
The geology of the Isparta Angle and its surroundings is dominated
by the predominantly carbonate successions of Mesozoic age and the
ophiolitic remnants of the last vestiges of the Neo-Tethys Ocean, such
as the Lycian nappes and the Antalya Complex in the west and the
Hadim, Bolkar, Bozkır and Beyşehir nappes in the east (Fig. 3). Between
Late Cretaceous (Campanian–Maastrichtian) and Early Eocene, several
ophiolitic units, including mélanges, were emplaced onto the margins
of the Tauride carbonate platform (Mackintosh and Robertson, 2012).
In southwestern Anatolia, thereare several predominantly marine Mio-
cene basins, which are presently perched on the Central and Western
Taurus Mountains, such as the Kasaba, Aksu, Köprüçay, Manavgat,
Mut, Ecemişand Adana basins (Figs. 1, 3). In these basins, the Early–
Late Miocene deposits unconformably overlie Cretaceous to Eocene
basement rocks of a thin-skinned fold–thrust belt (Burton-Ferguson
et al., 2005;Monod et al., 2006). This depositional architecture suggests
that the basement was exhumed and eroding prior to the Miocene
transgression (Erişet al., 2005). In the Mut, Ecemişand Adana basins,
the ﬁrst marine inundation is dated as Early Miocene (Bassant et al.,
2005; Erişet al., 2005; Ilgar and Nemec, 2005). Similarly, sediments
occupying the paleo-river valleys also date as Early Miocene in the
Mut Basin (Erişet al., 2005), Adana Basin (Ocakoğlu, 2002) and the
Aksu, Köprüçay, and Manavgat basins (Deynoux et al., 2005;
Karabıyıkoğlu et al., 2005). In the offshore, immediately south of
the Central and Western Taurus Mountains, there are several deep ba-
sins, including the Rhodes, Finike, Antalya, Cilicia and Iskenderun ba-
sins, which contain signiﬁcant thicknesses of Miocene deposits, in
addition to near-complete Pliocene–Quaternary successions (Aksu
et al., 2005a,b, 2009; Bridge et al., 2005; Hall et al., 2005a,b, 2009; Işler
et al., 2005). These offshore basins often are directly paired with an
onshore basin, only separated by the narrow continental shelf and the
adjacent steep continental slope, such as the onland Kasaba Basin and
its offshore continuation into Finike Basin, the onland Mut and Adana
basins and their offshore continuations into the Cilicia and Iskenderun
basins and the onland Aksu, Köprüçay, and Manavgat basins and their
offshore continuations into the Antalya Basin (Fig. 1).
The evolution of the Miocene basins in the eastern Mediterranean is
controlled by the development of a large, nearly east–west-trending
foredeep in front of the Tauride fold–thrust belt (Williams et al.,
1995). The Tauride culmination was characterized by an arcuate thrust
Fig. 2. Physiographyof the eastern Mediterranean showingthe Antalya Basin and itsrelationship withthe major tectonic elements in the region.The topography andbathymetry are com-
piled fromGeoMapApp (Ryan et al.,2009), the coastline andthe selected isobathscontours are from the International OceanographicCommission (1981).Inset is the study area shown in
3J. Hall et al. / Marine Geology 349 (2014) 1–23
front that delineated a broadsyntaxis, comprising several smaller thrust
culminations which developed in the foredeep itself. There are remark-
ably similar marine Aquitanian–Tortonian successions in the now-
onland Mut and Adana basins (Erişet al., 2005; Şafak et al., 2005),
Aksu, Köprüçay and Manavgat basins (Poisson et al., 2003a,b; Deynoux
et al., 2005; Karabıyıkoğlu et al., 2005) and the Mesaoria Basin of central
Cyprus (Robertson and Woodcock, 1986). The depositional similarities
further continue into the fold–thrust panels of the Misis Mountains
(Gökçen et al., 1988) and the Kyrenia Range (Calon et al., 2005a,b),
the Aksu Thrust (Poisson et al., 2003a,b), as well as the marine Cilicia,
Iskenderun, Antalya and Finike basins (Uffenorde et al., 1990; Aksu
et al., 2005a,b, 2009; Işler et al., 2005). These strong regional depositional
similarities suggest the presence of a single large basin in the Early
Miocene which encompassed what are now seemingly isolated basins
in the eastern Mediterranean (Fig. 1). This large ancestral basin probably
extended into the Karsantıand Maraşbasins in the east (Calon et al.,
2005a; Hall et al., 2005a; Ilgar and Nemec, 2005; Satur et al., 2005;
Hüsing et al., 2009) and the Antalya and Kasaba basins in the west
(Işler et al., 2005; Çiner et al., 2008). The development of crustal-scale
thrust culminations (e.g., the Misis–Kyrenia–Anamur lineament, the
Amanos–Larnaka fault zone and the Tartus ridge), perhaps associated
with the onset of escape tectonics associated with the ﬁnal collision of
the Arabian and Aegean–Anatolian microplates in the latest Miocene
and Pliocene–Quaternary (Şengör et al., 1985), essentially split the
foredeep into several large piggy-back basins: the Mut–Adana–Cilicia
basin complex, the Iskenderun–Latakia–Mesaoria basin complex, and
the Cyprus, Antalya, Finike and Rhodes basins (e.g., Calon et al., 2005a;
Hall et al., 2005a, 2009; Aksu et al., 2009). While the origins of these
basins lie in Miocene contraction, extensional structures are overprinted
on them in many places, especially during the Pliocene to Recent,
reﬂecting regionally-variable transtension.
An examination of the elevation of various Miocene successions
within the onland-offshore linked basins shows that correlative
shallow-marine units are routinely vertically separated by 3000–
5000 m across short distances of 5–10 km. Clearly, there must have
been primary seabed gradients within these basins, so that some of
the observed vertical stratigraphic offset can be attributed to variations
in the water depth in the ancestral Miocene basin. However, a signiﬁ-
cant proportion of this large offset is the result of rapid subsidence in
the offshore basins coupled withdramatic tectonic uplift of the onshore
basins associated with the rise of the Taurus Mountains (Erişet al.,
2005; Karabıyıkoğlu et al., 2005; Satur et al., 2005; Schildgen et al.,
2011; Cosentino et al., 2012; Koç et al.,2012).
2. Data acquisition and methods
The principal data usedin this paper consist of (a) ~4500 km of mul-
tichannel seismic reﬂection proﬁles collected in 1992 using the Memo-
rial University of Newfoundland (MUN) systems on RV Koca Piri Reis of
the Institute of Marine Sciences and Technology (IMST), Dokuz Eylül
University, (b) ~ 3000 km of multichannel seismic reﬂection proﬁles
using the MUN source and the IMST streamer on RV Koca Piri Reis,
(c) ~ 2000 km of multichannel seismic reﬂection proﬁles provided
by the Turkish Petroleum Corporation and (d) biostratigraphic and
lithostratigraphic data from two onshore exploration wells, provided
by the Turkish Petroleum Corporation (Fig. 4). The source for the
MUN multichannel data consisted of a Halliburton sleeve gun array,
employing gun sizes of 40, 20 and 10 in.
(656, 328 and 164 cm
with the total volume varying during maintenance cycling of the guns,
but typically 200 in.
) in 2008 and 2010 and 90–120 in.
) in 1992. Shots were ﬁred every 25 m, and reﬂections
were detected by the full 48 channels in 1991 (group interval =
Fig. 3. Geological map of the Western Taurus Mountains (simpliﬁed and redrawn from Blumenthal, 1963). AKMB = Aksu, Köprüçay, Manavgat basins, ANT = Antalya complex, BEY =
Beyşehir nappes, BOL = Bolkar nappes, BOZ = Bozkır nappes, LYC = Lycian nappes.
4J. Hall et al. / Marine Geology 349 (2014) 1–23
12.5 m), the nearest 12 channels of the same 48 × 12.5 m analog
multichannel streamer in 1992, 96-channel digital streamer (group
interval = 6.25 m) in 2008 and 216-channel digital streamer
(group interval = 6.25 m) in 2010. The resultant 12-fold (1991 and
2008), 3-fold (1992) and 27 fold (2010) data were recorded digitally
for 3–7 s (with delay dependent on water depth) at 1 millisecond
Fig. 4. Location map showing the position of seismic reﬂection proﬁles used in this study. Solid red lines = high-resolution multichannel seismic reﬂection proﬁles, dashed purple
lines = industry multichannel seismicreﬂection proﬁles.Seismic proﬁles shown as thick lines A–N are illustrated in text ﬁgures.Also shown are the locations of exploration wells drilled
in the onland Aksu, Köprüçay and Manavgat basins. Selected isobaths contours(in meters) are from the International Oceanographic Commission (1981). (For interpretation of the ref-
erences to color in this ﬁgure legend, the reader is referred to the web version of this article.)
Fig. 5. Stratigraphy of the AntalyaBasin showing the correlations betweenseismic stratigraphic units and the sedimentary successions on land, compiled from: (i) Adana Basin = Yalçın
and Görür(1984),Kozlu (1987),Yılmaz et al. (1988) and Gökçen et al. (1988), (ii)Mesaoria Basinand Kyrenia Range = Weiler (1969),Cleintaur et al. (1977)and, Robertsonet al. (1995),
(iii) Kasaba Basin = Hayward (1984),Şenel (1997a,b),andŞenel and Bölükbaşı (1997); (iv) Aksu, Köprüçayand Manavgat basins = Akay and Uysal (1985),Akay et al. (1985),Flecker
et al. (1998),andKarabıyıkoğluet al. (2000, 2005). Stratigraphy of the Manavgat-1and Manavgat-2wells is from the Turkish PetroleumCorporation(unpublished data). Units 1 through 4
are discussed in text. M and N are reﬂectors delineating the top and base of the Messinian successions, discussed in text.
5J. Hall et al. / Marine Geology 349 (2014) 1–23
sample rate, using a DFS V instrument in 1992 and a NTRS2 seismograph
in 2008 and 2010. The multichannel data were processed at Memorial
University of Newfoundland, with automatic gain control, short-gap
deconvolution, velocity analysis, normal move-out correction, stack,
ﬁlter (typically 50–200 Hz bandpass), Kirchhoff time migration, and
adjacent trace sum.
The sonic logs in the exploration wells show that the velocities in
the Pliocene–Quaternary sediments increase from ~1500 m s
sediment–water interface to ~2100–2300 m s
at the base of the
succession. Similarly, borehole data reveal that the Miocene siliciclastic
successions have interval velocities of 3000–3500 m s
. Interval veloc-
ities calculated during seismic data processing reveal that the Messinian
evaporites of Unit 2 in the marine Antalya Basin often exhibit values
ranging between 4200 and 5000 m s
. Note that thedip indicators in
ﬁgures illustrating the seismic reﬂection proﬁles are calculated using
1500 m s
sound velocity: dip estimates below the seabed should be
multiplied by the ampliﬁcation factor (i.e., interval velocity/water
3. Seismic stratigraphy and chronology
On the basis of acoustic character, stratigraphic position and age,
four distinct seismic stratigraphic units are identiﬁed in the Antalya
Basin and environs (Fig. 5): Unit 1: Pliocene–Quaternary siliciclastic
successions; Unit 2: Messinian evaporites and interbedded siliciclastic
successions, Unit 3: pre-Messinian Miocene siliciclastic and carbonate
successions and Unit 4: undifferentiated pre-Miocene sedimentary,
igneous and possibly metamorphic rocks. Two prominent and laterally
continuous markers are identiﬁed in the seismic reﬂection proﬁles,
delineating the Units 2–3 and 1–2 boundaries, respectively. Based on
their stratigraphic position and age, these markers are correlated with
the well-known M- and N-reﬂectors in the eastern Mediterranean.
Where possible, the lithostratigraphic makeup and chronology of
these units are determined by correlations withthe onshore exploration
wells. These four units are further correlated with lithostratigraphic
units identiﬁed in the adjacent Kasaba Basin in the west and the
Adana, Cilicia and Mesaoria basins in the east (Fig. 5).
There are four exploration wells drilled in the onland Aksu,
Köprüçay and Manavgat basins: Manavgat-1, Manavgat-2, Aksu-1 and
Ismail-1 (Fig. 4). The chronology of the Manavgat-1 and Manavgat-2
wells is most critical for this study because the successions encountered
in the well can be readily correlated with the seismic stratigraphic units
identiﬁed in the seismic reﬂection proﬁles. The Manavgat-2 well was
drilled to a total depth of 2565 m (Fig. 6). The well recovered ~204 m
of loosely consolidated to unconsolidated claystonewith few sandstone
interbeds. These sediments are assigned to the Pliocene–Quaternary
Yenimahalle Formation (Fig. 6; Turkish Petroleum Corporation, unpub-
lished data). Below this upper veneer, there is a 290 m-thick siliciclastic
succession composed of sandstones and shales with several volcanictuff
horizons. On the basis of biostratigraphic data this succession is corre-
lated with theLate Miocene Taşlık Formation, which is the lateral equiv-
alent of the evaporitic deposits of the Gebiz Formation deposited
associated with the Messinian Salinity Crisis (e.g., Garrison et al.,
1978). The Taşlık Formation is conformably underlain by a 445 m
thick siliciclastic succession consisting of sandstone, siltstone and
claystone interbeds, which is correlated with the Tortonian Karpuzçay
Formation (Fig. 6; Turkish Petroleum Corporation, unpublished data).
Below the Karpuzçay Formation the well recovered an approximately
436 m thick siliciclastic succession with several well-deﬁned limestone
beds (Fig. 6). This succession is correlated with the Aquitanian–
Serravallian Geceleme Formation, which is underlain by a 171 m-thick
prominent limestone unit, which is well known in the Aksu, Köprüçay
and Manavgat basins as the Oymapınar Formation (Akay and Uysal,
1985; Akay et al., 1985). At the base of the Oymapınar Formation the
well encountered 575-m thick Geceleme Formation clearly indicating
a repetition of stratigraphy. A northeast–southwest trending industry
seismic reﬂection proﬁle explains this age reversal: the Manavgat-2
well drilled through a broadly northeast-verging thrust at ~1.1 second
depth where a strongly reﬂective seismic package is clearly duplicated
(Fig. 6). The Manavgat-1 well a few km to the northwest was drilled
away from the thrust, and so does not include the duplication seen in
the Manavgat-2 well. The Manavgat-2 well recovered an additional
128 m of siliciclastic successions with carbonate interbeds, which are
correlated with the Aquitanian–Burdigalian Aksu Formation (Fig. 6).
3.1. Unit 1: Pliocene–Quaternary
The youngest succession identiﬁed in the Antalya Basin is char-
acterized by a strongly reﬂective, laterally continuous package of
high-frequency reﬂections which extends from the seabed to the
M-reﬂector (Figs.7,8). This unit is imaged on all seismic reﬂection
proﬁles, but shows dramatic thickness variations across the Antalya
Basin as further explained below. Unit 1 is also identiﬁed in the onland
seismic reﬂection proﬁle, and corresponds to the predominantly
Fig. 6. Industry seismic reﬂection proﬁle (A)showing the projectedlocations of the Manavgat-1 and Manavgat-2 exploration wells. Notethat there is a major NE-verging thrust thatpro-
duced a duplication of the lower Miocene successions in the Manavgat-2 well. Proﬁleis kindly provided by the Turkish Petroleum Corporation. See Fig. 5 for the details of theformations.
Location is shown in Fig. 4.
6J. Hall et al. / Marine Geology 349 (2014) 1–23
Fig. 7. High-resolution multichannel seismic reﬂection proﬁle (B) showing the architectures of seismic stratigraphic units described in text. The prominent M-reﬂector and N-reﬂectors
deﬁne thetop and base of the evaporite successions of Unit2. Note the co-occurrence of thrustsand extensionalfaults, discussedin text. Domains 1Band 3B are explainedin text. Location
is shown in Fig. 4.
Fig. 8. High-resolution multichannel seismic reﬂection proﬁle (C) showing the architectures of seismic stratigraphic units described in text. Note thatthe present-day continental slope is
delineated by a huge thrust culmination, and thatthe M-reﬂector deﬁnes a major erosionalunconformity acrossthe shallower regionof the Antalya Basin where theMessinian evaporites
of Unit 2 are absent. Domains 1A/B and 3A/B are explained in text. Location is shown in Fig. 4.
7J. Hall et al. / Marine Geology 349 (2014) 1–23
siliciclastic successions of the Yenimahalle Formation (Fig. 6). In the ma-
rine Antalya Basin, the base of Unit 1 is marked by a strong and distinc-
tive reﬂector, identiﬁed in the eastern Mediterranean as the M-reﬂector,
which is dated at 5.3 Ma. In a regional context, this unit is correlated
with: Kuranşa and Handere Formations of the Adana and Cilicia basins;
the Athalassa and Nicosia formations of the Mesaoria Basin; and the
Mirtou Formation of the Kyrenia Mountains of northern Cyprus (Fig. 5).
The thickness of the Pliocene–Quaternary succession varies across
the study area from b100 ms in the shallow nearshore region to
N1600 ms in elongated basins in deep western Antalya Basin (Fig. 9).
In general, Unit 1 is thinnest along the continental shelf and slope, but
thickens considerably toward the deeper water regions. The thickest
packages of Pliocene–Quaternary sediments, exceeding 1500 ms of
sediments are found along broadly east–west trending elongated de-
pressions situated across the southern portion of the western Antalya
Basin (Fig. 9). This trend is also illustrated in Işler et al. (2005). Toward
the east there are several northwest–southeast trending tear-drop
shaped basins which also contain 1500–1600 ms thick Pliocene–
Quaternary sediments. A broadly north–south trending elongated
lobe of Pliocene–Quaternary sediments occurs beneath the axis of
the present-day bathymetric channel in the western Antalya Basin
(Fig. 9). The distribution and thickness variations of the Pliocene–
Quaternary Unit 1 are predominantly controlled by the prevailing tec-
tonic regime in western Antalya Basin, where thick Pliocene–Quaternary
packages are found either (i) as growth strata wedges developed in
piggy back basins (Işler et al., 2005), or (ii) in mini basins between elon-
gated salt rollers and walls. The relationship between the distribution
and thickness variations of the Pliocene–Quaternary and ongoing tecto-
nism is further discussed later in this paper.
3.2. Unit 2: Miocene (Messinian)
Unit 2 is characterized by a weakly reﬂective package displaying
complex internal architecture with weak, discontinuous and often cha-
otic reﬂections that is bounded at their top by the M-reﬂector and base
by the N-reﬂector (Figs. 7, 8). Reﬂectors within this unit have generally
lower frequencies than those observed within the overlying Pliocene–
Quaternary succession. This unit also occursacross the southern Antalya
Basin and Florence Rise, wh ere Deep Sea Drilling Proj ect Leg 42, Site 37 5
and 376 results show that it is primarily composed of halite with alter-
nating smaller layers of anhydrite and limestone (Baroz et al., 1978).
Based on its acoustic character, the correlations with the Manavgat-1
and Managvat-2 wells (Turkish Petroleum unpublished data) and the
existing studies onland (Akay et al., 1985; Karabıyıkoğlu et al., 2000)
Unit 2 is correlated with the Messinian deposits (Fig. 5). In the onland
Aksu, Köprüçay and Manavgat basins the equivalent successions of
Unit 2 include siliciclastic series interbedded with anhydrite-bearing
and/or gypsiferous andcarbonaceous sediments of the Gebiz Formation,
Fig. 9. Isopach maps of the Pliocene–Quaternary succession of Unit 1 (top) and Messinian evaporite successionof Unit 2 (bottom) in the marine Antalya Basin. Unit1 thickness data from
the easternand western Antalya Basin are fromIşler et al. (2005) and King (2013),respectively.Note that there isa good correlation between the northernand western edgesof Unit 2 and
the 2000 m isobath (from International Oceanographic Commission, 1981), except inthe deep Finike Basin where the evaporitesare absent.
8J. Hall et al. / Marine Geology 349 (2014) 1–23
as well as the terrestrial coarser siliciclastic sediments of the Eskiköy
Formation (Akay et al., 1985; Karabıyıkoğlu et al., 2000). In a regional
context, this unit is further correlated with the Haymanseki Formation
of the Adana and Cilicia basins, the Kalavasos Formation of the Mesaoria
Basin and the Lapatza Formation of the Kyrenia Mountains (Fig. 5).
The thickness of the Messinian evaporite succession of Unit 2 varies
across thestudy area from 0 ms to ~ 1700 ms in the deep westernAntal-
ya Basin (Fig. 9). The edge of Unit 2 broadly follows the 2000 m
isobaths, except in the deep Finike Basin where there are no Messinian
evaporite deposits (Fig. 9;Aksu et al., 2009). The Messinian isopach map
clearly shows that the present-day post-halokinetic thickness of Unit 2
in western Antalya Basin is characterized by several broadly north-
west–southeast elongated tear-drop shaped lobes where Messinian
deposits range in thicknesses between 700 ms and 1600 ms (Fig. 9).
These thick lobes are separated from one another by similarly trending
zones where the Messinian evaporites are notably thin, ranging from
100 ms to 400 ms. A comparison betweenthe Pliocene–Quaternary iso-
pach map and the Messinian isopach map reveals that the axes of the
thickest sediment deposits between these two maps are offset, and
that the thickest regions in Unit 2 correspondwith the thinnest regions
in Unit 1. This observation indicates that the growth of the Pliocene–
Quaternary depocenters is controlled by halokinesis andthe associated
developments of salt diapirs and walls in the region.
The 0 ms isopach contour can be interpreted as denoting the depo-
sitional edge of the Messinian evaporites in western Antalya Basin.
However two lines of evidence argue that the depositional edge of
Unit 2 may have been much higher along the slope: (1) the wide-
spread occurrence of evaporite deposits in the Cilicia and Adana basins
in the easternMediterranean suggests that the sea level must have risen
considerably sometime during the Messinian to allow evaporite
deposition within these basins which are situated 1500 m–2000 m
above the present-day ﬂoor of the Antalya Basin, and (2) onland studies
clearly show that the Gebiz Formation in the Aksu, Köprüçay and
Manavgat basins include anhydritic and gypsiferous intervals, suggest-
ing that these now onland basins must have been inundated and were
receiving evaporitic deposition during the Messinian (Akay and Uysal,
1985; Akay et al., 1985). The occurrence of evaporites in the Cilicia
and Adana basins as well as in the Aksu, Köprüçay and Manavgat basins
argues that evaporite deposits must have developed higher along the
slope and/or the surrounding continental shelves of the Antalya Basin
during the Messinian. These possibly thinner evaporite deposits may
have migrated downslope during the Pliocene–Quaternary mobilization
of Unit 2 (Bridge et al., 2005; Işler et al., 2005), thus creating weld sur-
face along much of the present-day continental slope. Thus, the 0 ms
isopach contour not only shows the present-day edge of the evaporites,
but also suggests that bulk of the evaporite deposition has taken place in
the deeper parts of the Antalya Basin.
3.3. Unit 3: Miocene (pre-Messinian)
Unit 3 is situated below the N-reﬂector where Messinian evaporites
of Unit 2 are present and below the M-reﬂector where the Messinian
evaporites are absent (Figs. 7, 8). Unit 3 is characterized by strongly
reverberatory, high reﬂective and low amplitude reﬂections with signif-
icant lateral continuity (Fig. 8). Correlations with the onland Manavgat-1
and Manavgat-2 wells show that this unit is composed of siliciclastic and
carbonate successions of the Aquitanian–Tortonian age. In the broader
regional context, Unit 3 is further correlated with the Pakhna Formation,
including the Koronia and Terra members of the Mesaoria Basin, the
Kythrea Group of the Kyrenia Mountains in Northern Cyprus (Follows
Fig. 10. Pre-Messinian Miocene tectonic map of the western Antalya Basin showing major thrust faults (with ﬁlled triangular ticks on the hanging walls),the crestal hinge lines of prom-
inent ridges (shown with diamond ticks) and trough lines of major piggy-back basins developed on the backlimbs of major thrusts. Thick lines with half arrows depict strike-slip faults.
Onland thrusts are compiled from the geological map of the Antalya region (Blumenthal, 1963).
9J. Hall et al. / Marine Geology 349 (2014) 1–23
and Robertson, 1990); and also the Elekdağ, Kasaba and Sinekli forma-
tions of the Kasaba Basin (Şenel, 1997a,b; Şenel and Bölükbaşı, 1997;
3.4. Unit 4: Cretaceous–Eocene
Unit 4 constitutes the acoustic basement in the study area and con-
sists of a diverse collection of regional lithostratigraphic units ranging
from the Paleozoic to possibly Eocene (Fig. 5). Along the westernmost
Antalya Basin immediately east of the Antalya Complex there is an
acoustically dark-transparent unit which occurs directly below a very
thin veneer of Unit 1, separated by the M-reﬂector. The unit exhibits
only few coherent reﬂections emanating from seemingly chaotic and
disordered surfaces. This acoustic character is often observed associated
with massive igneous or metamorphic rock units. In this region Unit 4
probably includes the ophiolitic Antalya Complex (Bağcıand Parlak,
2009) and the Alanya Complex of high-pressure metamorphic rocks
(Okay and Özgül, 1984).
4. Structural interpretation
The structures and their associations observed in the seismic reﬂec-
tion proﬁles are described using three temporal phases: Phase 1 = pre-
Messinian Miocene; Phase 2 = Messinian and Phase 3 = Pliocene–
Quaternary. Phases 1 and 3 are each further dividend into three spatial
domains, as explained below.
4.1. Phase 1: pre-Messinian Miocene
Phase 1 affected the entire western Antalya Basin west of the
Beydağlarıand Antalya complexes (Fig. 10). The deformation associated
with Phase 1 also extends eastward toward the Kyrenia Range of north-
ern Cyprus and southward toward the Florence Rise. It is characterized
by structures developed during aperiod of protracted contractional de-
formation. On the basis of the predominant morpho-tectonic elements
and their trends and to a lesser extent the style of deformation, Phase
1 can further be subdivided into threespatial domains: the easternarcu-
ate mainly NW–SE trending domain 1A, the western primarily N–S
trending domain 1B and the triangular-shaped south-central domain
1C that straddles between domains 1A and 1B (Fig. 10).
4.1.1. Domain 1A
Domain 1A is a NW–SE trending zone which occupies the continen-
tal shelf and slope in theeastern portion of the study area (Fig. 10). It is
characterized by an arcuate SW- and W-verging, NW–SE and N–S
trending fold–thrust belt (Fig. 10). The belt is composed of 9–12 prom-
inent thrust panels, best imaged in the industry seismic reﬂection
proﬁles (e.g., Figs. 11, 12). Within domain 1A, the M-reﬂector is a dis-
tinctive marker and deﬁnes a prominent erosional unconformity
which separates the Pliocene–Quaternary successions of Unit 1 from
the underlying pre-Messinian Miocene successions of Unit 3 (Figs. 11,
12). In this area the gently northerly-dipping pre-Messinian Miocene
reﬂectors are clearly erosionally-truncated at the level of the M-
reﬂector. The leading thrust of the fold thrust belt is located at the
base of slope in the western Antalya Basin (T3 in Figs. 11, 12). The foot-
wall–hanging wall cutoffs of this fault are clearly visible in the seismic
reﬂection proﬁles. The thrust trajectory can be conﬁdently traced from
the seabed to ~5000 ms depth, deﬁning a prominent listric thrust
surface (e.g., Figs. 11, 12). At depth, this surface further deﬁnes a
200–300 ms thick distinctive reﬂector bundle that gently plunges
northward to depths exceeding 6000 ms. Thus, the slope face is the
forelimb of a huge thrust culmination carried by the T3 thrust
(Fig. 11). A secondary thrust (i.e., T3a) splays from T3 and creates a
wedge-shaped package bounded on its top and base by two thrusts:
this package occurs prominently at the base of slope and is readily
seen in most seismic reﬂection proﬁles (Figs. 11, 12).
Trajectories of the individual thrusts in the fold thrust belt can be
delineated in the seismic reﬂection proﬁles. For example, all thrusts
near the M-reﬂector exhibit relatively high angles (15–20°), but pro-
gressively become ﬂattened with depth (e.g., Figs. 11, 12). Here the
thrust trajectories imaged range from ~6° and can be conﬁdently
traced down to 6000 ms depth in the seismic reﬂection proﬁle,
where they merge into a near-horizontal detachment surface.
These depths, when converted using the interval velocities in the
seismic proﬁles, correspond to ~ 10–15 km, suggesting that the pre-
Messinian Miocene fold–thrust belt deﬁnes a crustal-scale feature
in front of the Western Taurus Mountains of south-central Turkey
(Figs. 1, 2). The folds within the fold–thrust belt are clearly asym-
metric having short, S- and SW-dipping forelimbs and longer, N-
and NE-dipping backlimbs (Figs. 11, 12).Thedistancebetweenthe
hinge lines of a given syncline to the adjacent anticline is notably
shorter than the distance between the hinge lines of the anticline
to the adjacent syncline, conﬁrming the asymmetry of the fold–
Fig. 11. Industry multichannel seismicreﬂection proﬁle (D) showing the Miocene structuralarchitecture of the western Antalya Basin. Note that theM-reﬂector is a prominent erosional
unconformity separating the Pliocene–Quaternary successions of Unit 1 from the pre-Messinian Miocene successions of Unit 3. Further note that the leading thrust of the foldthrust belt
delineates the base of slope in the western Antalya Basin, and that the slope face is the forelimb of a huge thrust culmination. Proﬁle is kindly provided by the Turkish Petroleum Corpo-
ration. Location is shown in Fig. 4.
10 J. Hall et al. / Marine Geology 349 (2014) 1–23
thrust belt. These fold geometries indicate a SW and W vergence di-
rection for the fold thrust belt (Fig. 11). A prominent back-thrust is
identiﬁed in the frontal portion of the huge thrust culmination that
deﬁnes the morphology of the present-day continental slope (BT in
In the trailing portion of the fold thrust belt, there are several distinc-
tive thrust panels where large piggy-back basins developed on the back
of the thrusts (Figs. 12, 13). Although the majority of the pre-Messinian
Miocene successionsof Unit 3 show no discernible sedimentary growth,
it is clearly observed within the uppermost portion of Unit 3, where the
upper Tortonian succession is mildly thicker within the trough of the
piggy-back basin and thins toward the ramp anticline (Fig. 13). This
sedimentary architecture suggests that thrusting occurred during the
Late Miocene. The tip points of the thrusts in domain 1A in the eastern
segment of the study area are invariably situated at or below the M-
reﬂector (Figs. 11, 12). By comparison, in the western portion of the
Fig. 12. Industry multichannel seismic reﬂection proﬁle(E) showing the Miocene structural architecture of the western Antalya Basin. Note that the M-reﬂector is a prominent erosional
unconformity separatingthe Pliocene–Quaternary successions of Unit 1 fromthe pre-Messinian Miocene successions of Unit 3 and that theprominent wedge-shaped package at the base
of slope delineated by two large thrusts. Also note the Pliocene–Quaternary reactivation of the Miocene thrusts in the westernmost portion of the proﬁle in domain 1B. Proﬁle is kindly
provided by the Turkish Petroleum Corporation. Location is shown in Fig. 4.
Fig. 13. High-resolution multichannelseismic reﬂection proﬁle(F) showing the detailed structural architecture of domains1A and 1B. Note thatthe M-reﬂectoris deﬁnes a major erosional
unconformity across the shelf region of the Antalya Basin, and thatthe Messinian evaporites of Unit 2 is absent in thisarea. Also note that the Pliocene–Quaternary successions are cut by
numerous normal faults, forming prominent horst and graben structures. Location is shown in Fig. 4.
11J. Hall et al. / Marine Geology 349 (2014) 1–23
study area in domains 1B and 1C (see below), thrusts extend into the
Pliocene–Quaternary, creating notable structures within these succes-
sions, as well as controlling the present-day morphology of the sea
ﬂoor (Figs. 12, 13).
The fold–thrust belt in the southeastern segment of domain 1A is
oriented in a NW–SE trend. Traced toward the northwest, these thrusts
progressively swing clockwise to assume a broadly N–S orientation in
the northern portion of the domain. An E–W trending industry proﬁle
located ~8 km south of the present-day shoreline conﬁrms that
the SW-convex arcuate fold thrust belt continues toward the present-
day coast (Fig. 14), and suggests that the prominent thrusts in this
belt must link with the prominent thrust panels mapped onland
(e.g., Poisson et al., 2003a,b). These onland thrust panels further deﬁne
the southeastern structures of the Isparta Angle (Poisson et al., 2003a,b).
Fig. 14. Industry multichannel seismic reﬂection proﬁle(G) showing the Miocene structural architecture of the western AntalyaBasin. Note that the M-reﬂector deﬁnes a prominent ero-
sionalunconformity separating the Pliocene–Quaternary successions of Unit1 from the pre-MessinianMiocene successions of Unit 3. Also notethe presence of a prominentramp anticline
delineated by the M-reﬂector. Proﬁle is kindly provided by the Turkish Petroleum Corporation. Location is shown in Fig. 4.
Fig. 15. High-resolution multichannelseismic reﬂectionproﬁle (H) showing the detailed structural architecture of domains 1A and 1B. Notethat the M-reﬂector deﬁnes a major erosional
unconformity acrossthe shelf region of theAntalya Basin, andthat the Messinianevaporites of Unit2 are absent in this area.Also note that thePliocene–Quaternary successions of domain
1C are cut by numerous high-angle planarnormal faults. Location is shown in Fig. 4.
12 J. Hall et al. / Marine Geology 349 (2014) 1–23
4.1.2. Domain 1B
Domain 1B is a narrow zone situated between domains 1A in the
east and 1C in the west (Fig. 10). The structural architecture of domain
1B is characterized bya large west-verging broadly N–S trending thrust
panel, which is developed as a secondary splay from the major thrust T3
of domain 1A (Fig. 12). The zone occupies the narrow bathymetric
channel in western Antalya Basin, and widens considerably toward
the abyssal plain in the south (Figs.2,10). The eastern boundary of do-
main 1B is delineated by the prominent thrust T3 and its westerly splay
T3a (Figs. 10, 12). The western boundary of domain 1B is marked by a
distinctive thrust which rises from at least 5000 ms depth in the east,
immediately west of domain 1A, westward to 2000–3000 ms depth in
the seismic reﬂection proﬁles (T4 in Figs. 10, 12). A prominent lower-
frequency reﬂector bundle clearly highlights the thrust trajectory,
deﬁning a nearly ﬂat and gently listric fault surface. The apparent thrust
angle is notably gentle (i.e., 6–9°) in the southern portion of the domain,
but gradually steepens to 15–20° in the northern portion of the domain.
Footwall and hanging wall cutoffs on the M-reﬂector suggest consider-
able contractional separation along this thrust (Fig. 12). A notable ramp
anticline developed on thehanging wall of thrust T4 (Fig. 12). This ramp
anticline can be readily mapped across the domain, deﬁning a promi-
nent structure along the western margin of the domain (Figs. 10, 15).
Seismic reﬂection proﬁles show that tectonic activity continued into
the Pliocene–Quaternary, as indicated by growth strata development in
Unit 1 (Figs. 12, 15). However, several lines of arguments can be made
for the initiation of the tectonic activity in the Miocene. For example,
there are several ramp anticlines which show steeply dipping shorter
forelimb segments, and gently dipping longer backlimb segments
(Figs. 11, 12, 15). This architecture suggests a preferred deformation
direction for the Unit 3 successions. The fact that these structures are
discordant with respect to the structures observed in Unit 1 above the
M-reﬂector suggests that they are developed somewhat independent
of the Pliocene–Quaternary tectonic phase.
4.1.3. Domain 1C
Domain 1C is situated along the continental shelf and slope of the
westernmost portion of the Antalya Basin, immediately west of the
Beydağlarıand Antalya complexes (Fig. 10). It is characterized by a
broadly N–S trending wet-verging fold–thrust belt which consists of
anels(Figs. 10, 15, 16). In the northern portion of the domain
the thrusts have well-developed ramp anticlines delineated by the
lower frequency reﬂections in Unit 3, as well as the M-reﬂector
(Figs. 12, 16). They cut the M-reﬂector, tipping within the lowermost
portion of the Pliocene–Quaternary successions of the unit. The eastern-
most thrust of this belt (T6 in Figs. 10, 15, 16) displays a listric thrust
trajectory and soles deep into Unit 3. Immediately west of thrust T6,
the seismic architecture of the pre-Messinian Miocene Unit 3 suggests
the presence of another east-verging thrust; however, this structure is
Fig. 16. High-resolution multichannel seismic reﬂection proﬁle(J) showing the detailed structural architecture of domains 1B/3B and1C/3C. Note that the Messinian evaporites of Unit 2
are very thin across thelower slope and absentin middle and upperslopes. Also note thatthe Pliocene–Quaternarysuccessions of domain1 are cut by numerous high-angle planarnormal
faults. The map distribution of the faults bracketed by the X symbol is illustrated in Fig. 17. Location is shown in Fig. 4.
13J. Hall et al. / Marine Geology 349 (2014) 1–23
not well imaged in the seismic reﬂection proﬁles. Traced toward
the south, the architecture of domain 1C becomes progressively more
complicated as structures become buried beneath the steep western
continental slope of the Antalya Basin.However, the structural elements
observed in the northernmost portion of the domain can still be mapped
in this region. For example, the thrust that demarks the boundary
between domains 1B and 1C is well imaged in the seismic reﬂection
proﬁles (T6 in Figs. 15, 16).
4.2. Phase 2: Messinian
The Messinian was a period of signiﬁcant changes in the eastern
Mediterranean Sea, both in the tectonic style as well as the morphology
of the basins and the surrounding landmass. From the latest Tortonian
into the Messinian, the progressive evaporation of the western exten-
sion of the Tethys Ocean (the evolving Mediterranean Sea) exposed
the continental shelves and slopes to subaerial processes (Hsü et al.,
1978; Rouchy and Caruso, 2006; Garcia-Castellanos and Villaseñor,
2011). During the Messinian, the entire Antalya Basin became a deep
largely subaerially exposed basin, with a very shallow water depth.
Periodic inundation of this deep but shallow basin throughout the
Messinian allowed the sedimentation of an approximately 2000 m-
thick evaporite successions within the deeper portions of the erosional
The tip points of the pre-Messinian Miocene thrusts occur at o r below
the M-reﬂector. Most of the thrusts do not affect the M-reﬂector and the
structuring that is observed on this marker appears predominantly
erosional in nature. These observations together with the absence of
growth strata development within the evaporite succession collectively
suggest that most of the Miocene thrust activity ceased by the time
when the M-reﬂector was formed, and that Messinian was a period of
relative tectonic quiescence. The Messinian evaporites have been subse-
quently mobilized, but this event has clearly taken place during the
4.3. Phase 3: Pliocene–Quaternary
The Pliocene–Quaternary structures in the Antalya Basin are largely
overprinted on the older structures. Some of these young structures
appear to have developed completely independent of older structures,
yet others are clearly evolved over the pre-existing structures by
the re-activation of the faults bounding these structures. Pliocene–
Quaternary structures can be described in three spatial domains,
where each domain displays distinctive sets of characteristics, including
structural trends and style and depth of deformation (Fig. 17). These do-
mains overlap nearly perfectly with the domains identiﬁed in Phase 1.
4.3.1. Domain 3A
The Pliocene–Quaternary domain 3A is a zone occupied by numer-
ous superﬁcial extensional faults which are developed over the pre-
Messinian Miocene fold–thrust belt (i.e., domain 1A; Fig. 10). The struc-
tural architecture of domain 3A is characterized by an arcuate belt
which has a predominantly NW–SE trend in the southeastern portion
of the study area, but progressively swings clockwise to assume N–S
Fig. 17. Pliocene–Quaternary tectonic map of the western AntalyaBasin showing majorthrust faults (withﬁlled triangular ticks on the hangingwalls), the crestalhinge lines of prominent
ridges (shownwith diamond ticks) and trough lines of major piggy-back basins developed on the backlimbs of major thrusts. Thick lines with half arrows depict strike-slip faults. The
architecture of the region bracketed by X is illustrated in Fig. 16. Map is compiled from data from the Department of Earthquake Research (Deprem Araştırma Dairesi), Ankara, Turkey
14 J. Hall et al. / Marine Geology 349 (2014) 1–23
trend in the northwest segment of domain (Fig. 17). This morpho-
tectonic character appears to overprint the pre-existing structures of
the pre-Messinian Miocene domain 1A, but displays a dramatically
different style of deformation, dominated by extensional faults (Fig. 17).
Domain 3A is characterized by numerous NW–SE trending and SW-
dipping extensional faults that deﬁne a series of internally parallel
fault blocks (Figs. 17, 18). Most of these faults cut the entire Pliocene–
Quaternary successions of Unit 1, extending to the depositional surface
where they create distinctive steps on the seaﬂoor. Some of these faults
show NE dips and create horst and graben structures along the slope
(Fig. 18). Along the upper slope, the extensional faults are steeper and
cut the M-reﬂector, where they create 200–300 ms vertical stratigraph-
ic separation. Along the lower slope the fault trajectories are notably
listric, and faults sole immediately below the M-reﬂector, near or onto
the tips of the underlying pre-Messinian Miocene thrusts described in
domain 1A. A few faults in this region only cut the middle portion of
Unit 1. These faults sole in bedding parallel detachments. Prominent
reﬂectors across the footwalls and hanging walls of the extensional faults
show very little sedimentary growth across these faults, suggesting that
the fault development post-dates most of the deposition of the Pliocene–
In the central portion of domain 3A and along the steeper lower
slope immediately above the abyssal plain the slope face is carved by
several superﬁcial listric extensional faults which developed over the
forelimb of a huge thrust culmination (e.g., Fig. 13). Landward of this
frontal margin, the morphology of the slope is very similar to that de-
scribed in the southeastern portion of the domain. This region is charac-
terized by numerous NW–SE trending and NE- and SW-dipping planar
extensional faults that deﬁne a series of internally-parallel horst and
graben structures (Figs. 10, 19). The majority of these faults show tip
points at the depositional surface where they create steps on the sea-
ﬂoor, giving the seaﬂoor an irregularly corrugated 2D appearance.
Most of these faults terminate near or at the M-reﬂector, but some cut
the M-reﬂector where they create 10–100 ms vertical stratigraphic sep-
aration (Fig. 19). Seismic proﬁles show little-to-no sedimentary growth
across these faults.
Farther toward the northwest the domain 3A is characterized by
numerous broadly N–S trending and E- and W-dipping high angle ex-
tensional faults (Figs. 8, 10, 16). The horst and graben structures prom-
inently imaged in the southeastern and central portion of the domain
become progressively less deﬁned toward the north. The high angle
extensional faults are largely conﬁned to the Pliocene–Quaternary suc-
cession of Unit 1 (Figs. 8, 16).
4.3.2. Domain 3B
Domain 3B is an arrow-head shaped zone situated between the
arcuate extensional faults of domain 3A in the east and the NE–SW
trending high-angle extensional faults of domain 3C in the west
(Fig. 17). The region exhibits complex structures involving re-activation
of the pre-Messinian Miocene thrusts, development of prominent fans
of superﬁcial extensional faults and numerous halokinetic structures as-
sociated with withdrawal and/or migration of the Messinian evaporites
of Unit 2. These structures are developed within an internally coherent
spatial framework with extensional, contractional and halokinetic struc-
tures having similar orientations and trends. The structural architecture
of the domain is described under three headings: south, central and
220.127.116.11. Southern extensional area. The southernmost portion of domain
3B is delineated by a N-convex arcuate region that extends from the
eastern fringes of domain 3C to the southwestern edge of domain 3A
(Fig. 18). A major south-dipping north convex arcuate listric normal
fault (referred to as the master fault) marks the northern boundary of
this region (N1 in Figs. 7, 16, 17, 20). Several smaller synthetic and
Fig. 18. High-resolution multichannel seismic reﬂection proﬁle(K) showing the detailed structural architecture of domains 1A/3A. Note that the edge of the evaporite Unit 2 is located at
the base of continental slope. Further note that the Pliocene–Quaternary successions are cut by numerous faults with normal-sense stratigraphic separations. Location is shown in Fig. 4.
15J. Hall et al. / Marine Geology 349 (2014) 1–23
antithetic faults are developed seaward and landward of the master
fault (Figs. 7, 16, 20). The master fault and its subsidiary splays deﬁne
the northern margin of a large Pliocene–Quaternary basin situated in
the southern portion of the western Antalya Basin (Fig. 17). They
show clear listric trajectories, cutting the M-reﬂector and the Messinian
evaporites of Unit 2 when present, extending into the pre-Messinian
Miocene Unit 3 (Figs. 7, 16, 20). The faults further extend to the seaﬂoor,
where they create 50–300 ms steps on the seabed. Prominent growth
strata wedges are developed within the Pliocene–Quaternary succes-
sion of Unit 1 immediately seaward of the master fault and its subsidiary
splays (Figs. 7, 16, 20). These growth strata wedges deﬁne a prominent
roll-over structure suggesting that the master fault was active during
the entire Pliocene–Quaternary. The structural architecture of the re-
gion south of the master fault and its subsidiary splays is characterized
by a distinctive fan of superﬁcial faults that are developed entirely
within the Pliocene–Quaternary succession of Unit 1. Most of these su-
perﬁcial faults are developed in the upper portion of Unit 1 and extend
up-section to the seaﬂoor where they create corrugated seabed
morphology. The trajectories of thesesuperﬁcial faults are notably listric
and in many instances the fault surfaces progressively curve to assume
bedding-parallel detachment surfaces (Figs. 7, 16, 20). The preponder-
ance of the shallow near-bedding-parallel detachment surfaces ob-
served within the Pliocene–Quaternary succession and the complexity
of their faults traces imply that the region experienced a transtensional
stress regime during the Quaternary (further discussed later).
18.104.22.168. Central extensional area. The region west and northwest of the
arcuate listric master fault is characterized by a prominent fault fan
which is entirely conﬁned to the Pliocene–Quaternary successions of
Unit 1 (X in Fig. 17). The fan consists of numerous broadly NE–SW
trending and southeast-dipping planar extensional faults, with a
smaller number of similarly trending, but northwest-dipping antithetic
faults (X in Figs. 16, 20, 21). In some instances, these planar faults
appear to deﬁne a series of domino faults, such as seen in Fig. 21,
and in other instances the faults are developed over the halokinetic
structures, such as seen in Fig. 16. The tip points of the faults often
extend to the depositional surface where they create steps on the sea-
ﬂoor (Figs. 20, 21). There is little-to-no sedimentary growth within
the Pliocene–Quaternary succession associated with these faults, sug-
gesting that the faulting largely postdates the deposition of most of
the Pliocene–Quaternary successions.
22.214.171.124. Northern area with extensional and contractional structures. The
northernmost portion of domain 3B is a narrow belt situated between
domains 3A in the east and 3C in the west (Fig. 18). The tectonic archi-
tecture of the region is deﬁned by two distinctly different structural
styles that overprint one another (e.g., Figs. 7, 12, 13, 15, 16, 21). The re-
gion is characterized by two large broadly N–S trending and oppositely-
verging thrusts (west-verging T3A, Figs. 12, 13, and east-verging T6,
dipping extensional faults (Fig. 17). The prominent thrust T6 is clearly
mapped displaying a large ramp anticline developed within the pre-
Messinian Miocene Unit 3 (Figs. 12, 13, 15, 17). Seismic proﬁle shows
that the lower portion of the Pliocene–Quaternary succession is also
incorporated into the ramp structure, which shows clear growth strata
wedge on its backlimb, suggesting that thrust T6 was active during the
early Pliocene–Quaternary (e.g., Fig. 13). In the southern portion of the
northern area, there are two additional thrusts that were active during
the Pliocene–Quaternary (T5 in Fig. 17). These thrusts exhibit well-
developed ramp anticlines (e.g., Figs. 7, 21). These thrusts sole deep
into the pre-Messinian Miocene Unit 3.
Fig. 19. High-resolution multichannel seismic reﬂection proﬁle (L) showing the detailed structural architecture of domains 1A/3A across the Antalya Basin continental slope. The
architecture of the pre-MessinianMiocene successions is characterized bylarge thrusts. Notethat evaporites(Unit 2) are conﬁnedto local pockets. Also notethat the Pliocene–Quaternary
successions are cut by numerous faults with normal-sense stratigraphic separations, creating horst and graben structures. Location is shown in Fig. 4.
16 J. Hall et al. / Marine Geology 349 (2014) 1–23
The northeastern and/or southwestern margins of the large
Pliocene–Quaternary ramp anticlines are cut by several high-angle
planar extensional faults (Figs. 13, 16). These faults invariably sole ei-
ther on the M-reﬂector or near the tip of the underlying thrusts. The ob-
served geometric relationship between the ramp anticlines and the
superﬁcial extensional faults strongly suggests that the extensional
faults are developed as a response to the thrust activity. The extension-
al/transtensional character of the structure in the southern and possibly
central segments of the domain and the predominantly contractional
character of the structure in the northern segment of the domain
strongly suggest that the strain is spatially partitioned in this area.
4.3.3. Domain 3C
Domain 3C is situated in the westernmost portion of the studyarea.
It occupies the shelf and slope regions of the western Antalya Basin. The
domain is characterized by 7–9NE–SW- and NNE–SSW-trending high-
angle normal faults (Figs. 15–17, 21). The shallow occurrence of the ﬁrst
multiple, coupled with the steep continental slope with numerous large
slide and slump masses (discussed below) renders poorer temporal and
lateral resolutions of seismic markers below the M-reﬂector. However,
primary reﬂectors can still be identiﬁed in the pre-Messinian Miocene
Unit 3, as well as the M-reﬂector, and other Pliocene–Quaternary
features, including the seabed morphology, so that the locations of
these high-angle extensional faults can be delineated.
Some of the faults show tip points at or below the M-reﬂector, yet
others show tip points extending into the lower portion of the
Pliocene–Quaternary successions of Unit 1. In all cases, the faults sole
deep into Unit 3. The slope face in this area is marked by several large
lenticular units that rest over the steeply (10–20°) southeast-dipping
M-reﬂector. Seismic reﬂection proﬁles show that these lenticular units
are bounded at their upslope ends by numerous superﬁcial listric de-
tachment surfaces (Fig. 15). These shallow structures are very similar
in their morphology and internal seismic architecture to the submarine
slide and slump masses described elsewhere (e.g., Hiscott and Aksu,
1994). In the southern portion of the study area, the domain is notice-
ably broader and is deﬁned by several high-angle extensional faults
The structural data described above reveal the presence of a complex
tectonic history in the western Antalya Basin, spanning from the
Miocene (or older) to the Pliocene–Quaternary. The Miocene succes-
sions deﬁne a southwest to west verging imbricate fold–thrust belt
(Fig. 10) aligned with structures mapped onland that deﬁne the eastern
limb of the Isparta Angle (e.g., Poisson et al., 2003a,2003b). Tortonian
and older successions are involved in the fold–thrust panels, suggesting
that the Isparta Angle continued to evolve at least into the late Miocene.
The tectonic activity experienced a period of relative quiescence across
the western Antalya Basin during the Messinian. In a broad regional
sense, this period follows a major tectonic re-organization in the eastern
Mediterranean Sea: the collision of the Arabian Microplate with the
Eurasian Plate and its ﬁnal suturing along the Bitlis–Zagros belt (around
11 Ma, e.g., Şengör et al., 1980; Robertson, 1998) initiated the west-
escape of the Aegean–Anatolian microplate (Dewey and Şengör, 1979;
Şengör et al., 1985; Dewey et al., 1986). This westward escape is accom-
modated along a number of major crustal-scale transform faults, includ-
ing the North and East Anatolian Transform faults, the Kozan Fault, the
Tuzgölü Fault, and the prominent Misis–Kyrenia–Aksu fault zone
(Aksu et al., 2005a; Işler et al., 2005). During the Pliocene–Quaternary
Fig. 20.High-resolutionmultichannelseismic reﬂectionproﬁle (M) showingthe detailed structural architecture of the southwestern AntalyaBasin. Note thatthe edge of the evaporiteUnit
2 is locatedat the base of continental slope andthat Unit 2 is absentacross the continental slope andshelf. Further notethat a prominent listric extensional faultdeﬁnes the framework of
the deep basin in this region. Several shallow faults deﬁne near-bedding-parallel detachments. Location is shown in Fig. 4.
17J. Hall et al. / Marine Geology 349 (2014) 1–23
the tectonic evolution in the eastern Mediterranean is verystrongly spa-
tially partitioned into contractional, extensional and strikeslip domains.
5.1. Morpho-tectonic elements of the Miocene fold–thrust belt
A prominent crustal-scale imbricate fold–thrust belt formed in the
offshore western Antalya Basin during the Miocene. The thrusts display
broadly arcuate map traces which trend NW–SE in the central portion of
the Antalya Basin, but toward the northwest these thrusts progressively
assume a NNE–SSW trend. During the Middle–Late Miocene, the entire
Antalya Basin, including its present day onshore sedimentary basins,
was situated between the evolving Tauride culminations in the north
and the subduction zone at the leading edge of the Neotethys Ocean
to the south. The Late Miocene (mainly Serravallian to Tortonian)
successions of the Karpuzçay and Aksu formations onland and their
correlative successions imaged in the marine seismic stratigraphic Unit
3 were deposited within a large elongated, broadly east–west trending
foredeep extending from the Bitlis Ocean in the east (e.g., Şengör et al.,
1985), across the present-day Iskenderun, Adana, Cilicia basins (Aksu
et al., 2005a,b; Burton-Ferguson et al., 2005), and the Kyrenia Range
of northern Cyprus (Calon et al., 2005a,b) into the eastern Antalya
Basin (Işler et al., 2005) and the western Antalya Basin in the west. The
ﬁnal collision of the Arabian Microplate with the eastern portion of the
Aegean–Anatolian microplate in the Late Miocene splits the broadly
east–west trending foredeep with arcuate deformation fronts, such as
the Misis–Kyrenia fold–thrust belt, the Amanos–Larnaka fold–thrust
belt and the zone which links the Tartus Ridge with the Cyprus Arc
(Fig. 1), and to the west, the Kyrenia–Aksu fold thrust belt (Fig. 1;Işler
et al., 2005). The fold–thrust structures associated with this Late Miocene
compression are clearly seen in the western Antalya Basin, where there
is protracted contractional deformation during at least the deposition of
the upper portion of Miocene Unit 3 (Işler et al., 2005, and this study).
Thus the Antalya Basin shows a similar late Miocene history to that
described in the Cilicia and Iskenderun basins (Aksu et al., 2005a,b), in
the Kyrenia Range (Calon et al., 2005a,b), as well as the Latakia Basin
(Hall et al., 2005a,b).
At the end of the Tortonian, the Mediterranean Sea was situated
at approximately the same subtropical latitude as today and was
completely isolated both from the Atlantic Ocean and the Indian
Ocean (Rouchy and Caruso, 2006; Garcia-Castellanos and Villaseñor,
2011). This conﬁguration led to the Messinian Salinity Crisis (Hsü
et al., 1978).
Fig. 21. High-resolution multichannel seismic reﬂection proﬁle (N) showing the detailed structural architecture of domains 1B/3B in western Antalya Basin. Note that the Pliocene–
Quaternarysuccessions are cut by numerous planar faults which form tilted domino-pattern. Location is shown in Fig. 4.
18 J. Hall et al. / Marine Geology 349 (2014) 1–23
During the Messinian, the Mediterranean Sea became desiccated
and the ensuing lowering of the base level and subsequent subaerial
exposure led to profound erosion of all the Mediterranean basins. This
erosional event is represented by the N-reﬂector where the Messinian
evaporites are present and by the M-reﬂector where they are absent.
The observed thickness of the Messinian evaporites range from
3000 m in the Herodotus Basin to ~2500 m in the vicinity of Florence
Rise (Biju-Duval et al., 1978) and to ~1000 m in the Cilicia and Latakia
basins (Aksuet al., 2005a; Hall et al., 2005a). The ﬁnal phase of desicca-
tion of the Mediterranean Sea at the end of the Messinian (Hsü et al.,
1978) and the associated subaerial exposure of the sea-ﬂoor resulted
in the development of the well-known unconformity represented in
the seismic reﬂection proﬁles by the M-reﬂector. The truncation of the
folded strata of Unit 3 by the M-reﬂector (e.g., domain 1A) implies
that the initial thicknesses of the Miocene sedimentary ﬁll of the
piggy-back basins were greater than what is now observed on the
seismic reﬂection proﬁles. The progressionof contractional deformation
during the early Messinian is difﬁcult to establish, because the architec-
ture of the evaporite unit was considerably changed by both contrac-
tional deformation and halokinesis that took place after the early
5.2. Morpho-tectonic elements of the Pliocene–Quaternary
A major kinematic change occurred during the transition from the
Miocene to the Pliocene, when the regional strain was partitioned into
three spatially localized tectonic domains, as described above: (i) an
extensional domain (3A) conﬁned to the Pliocene–Quaternary Unit 1,
occupying the northeastern portion of the study area, which co-exists
with a few re-activated Miocene thrusts, (ii) a predominantly exten-
sional domain (3B) with few re-activated pre-existing Miocene contrac-
tional structures in the southern and central portion of the study area,
and (iii) an extensional and/or transtensional domain (3C) occupying
the continental shelf and slope in the westernmost Antalya Basin.
5.2.1. Pliocene–Quaternary extensional fault zone over large
Normal faults were observed affecting the Pliocene–Quaternary suc-
cessionsalong the Aksu–Kyrenia deformation zone by Işler et al. (2005),
who suggested that the faults also accommodate signiﬁcant strike slip
displacements. Within the central portion of the Antalya Basin, small
amounts of growth strata observed in the hanging walls of these faults
suggest that the faulting may have locally initiated during the lower
The presence of this zone dominated by extension and transtension
immediately co-occurring with a zone dominated by transpression is
enigmatic. Işler et al. (2005) proposed that this lineament developed
through the partitioning of displacements created by the ensuing
westward escape of the Aegean–Anatolian microplate in the north
(i.e., Dewey et al., 1986), and accommodated by an arcuate splay of
the East Anatolian Transform Fault, extending from the Misis Mountains
of southern Turkey to the Kyrenia Range of northern Cyprus and farther
west to the Antalya Basin.
Işler et al. (2005) used GPS data (e.g., McClusky et al., 2000, 2003)to
argue that the sense of movement along the Misis–Kyrenia–Aksu fault
zone must be sinistral. Sinistral strike slip is also suggested by Meijers
et al. (2011). However, several other studies suggested that the
Aksu fault zone is a dextral strike-slip system overprinting large re-
Fig. 22. Physiographyof the eastern Mediterranean Sea showinga selection of GPS vectors, relative to a ﬁxed Anatolia, redrawn from McClusky et al. (2000). Thetopography and bathym-
etry are compiled from GeoMapApp (Ryan et al., 2009), the coastline and the selected isobaths contours are from the International Oceanographic Commission (1981). Major structural
lineaments are drawn in white;coastlines in black; GPSvectors and site names are shownin red with white outline.Study area of the western Antalya basin is shown as orange rectangle.
PST = Pliny–Strabo trenches, BDL = BeydağlarıLineament. Strike-slip motions acr oss major lineaments indicated by GPS vectors are shown in red, withpossible extensions indicated with
dashed arrows in the triangular area south of the Isparta Angle. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)
19J. Hall et al. / Marine Geology 349 (2014) 1–23
activated thrusts (Barka and Reilinger, 1997; Yağmurlu et al., 1997;
Poisson et al., 2003a,b, 2011; Piper et al., 2006; Çiner et al., 2008;
Toprak et al., 2009). These studies imply that the offshore continuations
of these re-activated thrusts must also have a dextral sense of slip.
A distinctly spatially localized contractional zone is situated in
the eastern and central portion of the western Antalya Basin. Here
the prominent pre-existing Miocene thrusts are re-activated in the
Pliocene–Quaternary, as indicated by the growth strata architecture
that developed in the associated piggy-back basins. Comparison of the
structures in the greater onland Aksu, Köprü, Manavgat and Antalya
basins revealed that the offshore re-activated fold–thrust belt can be
readily linked with the 4–7 large thrust panels identiﬁed onland, associ-
ated with the Aksu phase of compression (Poisson et al., 2003a,b, 2011).
Poisson et al. (2003a) documented that the compressional deformation
continued into the lower–middle Pliocene in the onland Aksu, Köprü
and Manavgat basins. However, our offshore data also reveal that the
thrust activity locally continued into the upper Pliocene–Quaternary,
particularly in the southern and southeastern segment of the study
area. Industry seismic reﬂection proﬁles showed that this thrust belt
had very deep roots, extending well into 6–8s(or~10–15 km), strong-
ly suggesting that this fold–thrust belt was a crustal-scale feature, and
the associated tectonic activity was thick-skinned.
The correlation of the offshore thrust panels with those mapped
onland suggests that the fold–thrust belt associated with the Aksu
phase of thrusting formed a 350 km long and 30–50 km wide deforma-
tion front, extending through the Antalya Basin to the thrust panels of
theKyreniarangeofnorthernCyprus(Calon et al., 2005a,b; Işler et al.,
2005), and then farther east. Aksu et al. (2005a) showed that the fold–
thrust panels mapped on the Kyrenia Mountains extend across the
Cilicia Basin and link with the fold–thrust panels mapped in the Misis
Mountains of southern Turkey, and then toward the Kahramanmaraş
triple junction in southeastern Turkey (Fig. 1). Thus, we identify a
750 km long south-convex deformation front referred to as the Misis–
Kyrenia–Aksu fault zone (c.f., Aksu et al., 2005a,b; Calon et al., 2005a,
b; Işler et al., 2005). We regard this as a major orocline in eastern
Mediterranean Neogene evolution, related in part to the Isparta Angle
but with much greater regional ramiﬁcations. Onland in the Isparta
Angle, the individual faults (i.e., the Aksu Fault, Kırkkavak Fault;
Figs. 10, 17) of the northwestern limb of this arcuate deformation
front exhibit notable dextral Pliocene–Quaternary strike slip displace-
ments (Barka and Reilinger, 1997; Yağmurlu et al., 1997; Çiner et al.,
2008). The individual faults on the northeastern limb of the oroclinal
deformation zone (i.e., the Misis–Kyrenia fault zone, Misis Thrust,
Kyrenia Thrust, AslantaşThrust) all are known to have sinistral strike
slip movements (Kelling et al., 1987; Kozlu, 1987; Gökçen et al.,
1988). The fact that the northwestern and northeastern limbs of the
arcuate deformation front exhibit oppositely-directed slip directions
suggests that the southernmost apex of the deformation front must be
a zone of intense contractional deformation. The thrusting observed in
the Kyrenia Range of northern Cyprus during the Quaternary may be
partially explained by the convergence of these two slip vectors in the
region of Kyrenia Range.
5.2.2. Extensional/transtensional zone in western Antalya Basin
A broad northeast–southwest trending zone of invariably southeast-
dipping extensional structures cuts the Pliocene–Quaternary strata of
Domain 3B at its boundary with domain 3C (Fig. 17). The steeply-
dipping faults cut the Pliocene–Quaternary succession of Unit 1 and
become deeply rooted in Miocene Unit 3. This fault system appears to
control the morphology of the present-day continentalmargin. Indeed,
there are several prominent scarps along the shoreline where the strike
of the scarp face is nearly identical to the strike of the individual faults in
this system. The fact that these faults are steeply deeply cutting into the
Miocene (or older) Unit 3, and that they delineate a series of sharp
escarpments, both onland and across the shelf break suggests that
they form part of a large crustal-scale structure which shapes the
westernmost Antalya Basin. Recent mapping of the onland Beydağları
and Antalya complexes also delineated numerous Quaternary and
younger faults which appear to have developed over the similarly
trending Miocene and older thrust surfaces (Fig. 10)anddisplay
north-northeast–south-southwest trends and extend along the entire
length of the Kemer Peninsula (Yaltırak, unpublished data). This fault
system appears to link with a major north–south trending dextral
strike-slip fault zone that extends ~200 km from the towns of Kırka to
Afyon, and then to Isparta (Savaşçın et al., 1995), deﬁning the Antalya
fault zone (Fig. 17). Glover and Robertson (1998a,b) also described a
dextral strike-slip fault zone along the Beşadalar–Kemer zone immedi-
ately west of the marine Antalya Basin. However, Barka and Reilinger
(1997) use modern GPS data to suggest that this is a sinistral strike-
slip zone. Along the apex of the Isparta Angle, this strike-slip fault
zone is associated with a prominentnorth–south trending potassicalka-
line volcanic belt situated between the Menderes and Kirşehir Massifs
(Savaşçın et al., 1995). These authors used radiometric ages and geo-
chemical data on the volcanic successions to show that the volcanism
occurred along this north–south zone paralleling the bisector of the Is-
parta Angle. They showed that the age of the volcanic rocks becomes
progressively younger from the north (i.e., the region of Kırka and
Afyon dated at 21–17 Ma) toward the south (i.e., Isparta at 4 Ma, and
Antalya 1.5–3.0 Ma; Besang et al., 1977; Sunder, 1982; Lefèvre et al.,
1983). They further argued that the fault zone exhibits dextral strike
While Glover and Robertson (1998a,b) and Zitter et al. (2003) sug-
gest that the Pliocene–Quaternary tectonism in western Antalya Basin
is dominated by extension, Işler et al. (2005) argued that except for
very superﬁcial extensional faulting associated with gravity sliding,
the region is dominated by contractional deformation. New data and
mapping showed that indeed Glover and Robertson (1998a,b) and
Zitter et al. (2003) were correct: there is a prominent Late Pliocene–
Quaternary tectonic phase of extensional faulting. However, there was
also a distinct phase of Early–Middle Pliocene–Quaternary phase
of contractional deformation, which locally continued until the Late
Pliocene–Quaternary. The co-occurrence of extensional and contrac-
tional deformation within roughly the same interval is suggestive of
strain partitioning in a strike-slip regime.
5.3. Linkage with the Isparta Angle
The Isparta Angle is a north-pointing triangular-shaped tectonic
province in southern Turkey. It constitutes one of the most important
structures in southern Turkey which can be correlated with other
major lineaments in the eastern Mediterranean such as the Kyrenia
Range of northern Cyprus. The Isparta Angle has an autochthonous
core in the Beydağları, overthrust from the west by the Lycian nappes
and from the east by nappes of the Antalya Complex and the Hoyran–
Beyşehir–Hadım(Monod, 1977; Robertson et al., 2003). Nappe
emplacement occurred during early Tertiary closure of the Pamphylian
Basin, which originally separated the Beydağlarıand the Western Taurus
platforms during the Mesozoic (Poisson et al., 2003a,b), but nappe devel-
opment continued intermittently until Miocene time. The Isparta Angle
is transected by several younger structural lineaments. The Lycian
nappes are cut by the active NE–SW-Burdur–Fethiye fault zone, charac-
terized by sinistral strike-slip faults with considerable normal dip-slip
component (Şaroğlu et al., 1987; Price and Scott, 1994; Barka et al.,
1997) and which may be a northeasterly extension of the Pliny–Strabo
Trenches STEP transform fault zone linking Cyprus and Hellenic arcs.
Miocene basins occupy part of the onland eastern margins of the Isparta
angle and are cut by the transpressional, N–S, Kırkkavak fault and the
NNW–SSE-trending, westward-verging Aksu Thrust, which may link to
the southeast through the offshore Antalya Basin to the Kyrenia Range
of northern Cyprus (Işler et al., 2005). Normal faults also cut parts of
these Miocene basins (e.g., Çiner et al., 2008). The evolution of the
nappe systems and their relationship with closure of strands of the
20 J. Hall et al. / Marine Geology 349 (2014) 1–23
Tethyan Ocean are still widely debated (Robertson et al., 2003; Güngör,
2013). However of particular interest here is the later, late Miocene to
Recent history, which is characterized by rotations about vertical axes
documented from paleomagnetic studies. During the Miocene, the
western limb of the Isparta angle, including the Beydağlarıcarbonate
massif experienced a 30° counterclockwise (CCW) rotation, according
to Kissel and Poisson (1986),andMorris and Robertson (1993). Recent
re-evaluation of this rotation (van Hinsbergen et al., 2010) indicates a
20° CCW rotation during 16–5 Ma, i.e., Middle to Upper Miocene time.
wise (CW) rotation since the Eocene (Kissel et al., 1993), though the
highly variable estimates (from 7° to 56°) of such rotations suggest
that they may be localized to individual thrust sheets and not character-
istic of the region (Meijers et al., 2011).
The northernmost seismic reﬂection proﬁle where thrusts are
observed in the marine data is only 5 km south of the present-day
shoreline (Fig. 14), allowing a correlation of the marine structures
with the similarly-trending and similarly-verging structures mapped
onland. In fact, the marine fold–thrust belt can be readily correlated
with the eastern limb of the Isparta Angle: the structures mapped in
Unit 3 of the western Antalya Basin are the seaward continuation of
the structures mapped and described onland. Seismic reﬂection proﬁles
and the borehole data further document that the Tortonian and older
successions are involved in the fold–thrust panels, suggesting that the
tightening of the Isparta Angle continued to evolve at least into the
The thrusts mapped in the marine area invariably have south-
westerly vergence. This is consistent with the trend and vergence
of structures mapped onland across the eastern limb of the Isparta
Angle (e.g., Waldron, 1984; Poisson et al., 2003a,b, 2011). Furthermore,
continued thrusting during the Pliocene and Quaternary observed
offshore matches the continuation of thrusting observed onland in, for
example, the Aksu Basin (Çiner et al., 2008) at least in the Pliocene.
In summary, the Miocene westerly-directed thrusting occurring off-
shore matches that observed onland on the eastern limb of the Isparta
Angle. In the broadest regional sense, this is interpreted as part of a
much longer orocline, extending to the East Anatolian Fault, but locally
its development may also be related to the counter-clockwise rotation
of the western and central parts of the Isparta Angle, which indicates
contraction across Antalya bay (van Hinsbergen et al., 2010).
5.4. Relationship to current deformation determined from GPS studies
How can the possibly contrary directions of strike-slip across the
western Antalya Basin and adjacent Beydağlarıbe related? Fig. 22
shows a simpliﬁed version of GPS motion vectors from McClusky et al.
(2000) relative to a ﬁxed Aegean–Anatolian microplate. The Aegean
Sea is generalized by vectors showing strong southerly motion. Vectors
from stations MATR and HELW show that the African Plate is moving
northward about a rotation pole not far off the eastern edge of the
map. Convergence at the Florence Rise increases westward. Vectors
from stations ANTG, KASO and SIRA show that the triangular block be-
tween the Burdur–Fethiye fault zone and the Aksu–Kyrenia fault zone
is moving northward, more slowly than the African Plate, indicating
contraction across the Florence Rise, at least in the west. The northward
motion of the triangular block contrasts with the southerly motion of
the Aegean segment of the Aegean–Anatolian microplate, conﬁrming
the sense of sinistral strike-slip across the Burdur–Fethiye fault zone.
The southwesterly vector at BURD contrasts with the northerly vector
at SIRA, suggesting that sinistral transpression extends some distance
into the triangular block from the Burdur Fethiye fault zone. Northeast
of the Aksu–Kyrenia fault zone, the vector at SEKI indicates not only
continuing contraction across the thrust zone, but also dextral strike-
slip relative to vectors at ANTG, KASO and SIRA. Such dextral strike
slip may well extend farther into the triangular zone, and thus conﬁrm
oblique slip across many of the steep faults at the western margin of the
western Antalya basin. The BeydağlarıLineament (BDL, Fig. 22) may
mark the boundary of the dextral strike slip fault zone from the sinistral
strike-slip faulting closer to the Burdur–Fethiye fault zone. There is no
clear evidence of present-day extension in the triangular zone from
the GPS vectors. In essence, it appears that continuing northward mo-
tion of the African Plate, increasing to the west, results in contraction
across theFlorence Rise but that northward motion ispartly transmitted
to the triangular block, resulting in sinistral strike slip towards its
western margin and dextral strike slip towards its eastern margin.
How these general motions are expressed locally in transtension or
transpression will depend on local orientations of older structures
being reactivated, thus presenting a diversity of apparently co ntradictory
strains in adjacent blocks.
Offshore seismic mapping of Miocene to Recent structures in the
offshore western Antalya Basin leads to recognition of two contrasted
phases of structural development. A west- to southwesterly-verging
late Miocene fold–thrust belt links similar structures on the eastern
limb of the Isparta Angle to the Aksu–Kyrenia–Misis oroclinal contrac-
tional zone, which developed in an ancestral foredeep basin subsequent
to the Arabia–Eurasia collision and the initiation of the westward tec-
tonic escape of the Aegean–Anatolian microplate. Locally this contraction-
al deformation may also relate to the counterclockwise rotation of the
western limb of the Isparta Angle. Pliocene–Quaternary deformation is
spatially partitioned and characterized by transtension and transpression,
partly reactivating older structures, but consistent with present day
differential motions recorded by GPS, indicating northward motion of
the block within the Isparta Angle and extending offshore, relative to
the adjacent blocks to the north of the Angle.
We thank the Ofﬁcers, crew and scientiﬁc personnel of the RV Koca
Piri Reis for their assistance in data acquisition. Special thanks are
extended to the Turkish Petroleum Corporation for kindly providing
paper copies of their multichannel seismic proﬁles and well informa-
tion. We acknowledge research funds from the Natural Sciences and
Engineering Research Council of Canada (NSERC) to Aksu and Hall,
and contributions from the Ofﬁce of the Vice-President Research at
Memorial University of Newfoundland. Seismic data were processed
at Memorial University of Newfoundland, using the ProMAX© software
donated by Landmark Graphics. Assistance with data processing was
provided by Sharon Deemer. We thank D.J.J. van Hinsbergen and
A. Çiner for their valuable reviews of the originalversion of this paper.
Akay, E., Uysal, S., 1985. Orta Toroslarınbatısındaki (Antalya) Neojen çökellerinin
stratigraﬁsi, sedimantolojisi ve yapısal jeolojisi. Mineral Research Exploration
Institute (MTA), unpublished report, 276 pp.
Akay, E., Uysal, S., Poisson, A., Cravette, J., Müller, C., 1985. Antalya Neojen havzasının
stratigraﬁsi (The stratigraphy of the Neogene Antalya Basin). Bulletin of the Geolog-
ical Society of Turkey 28, 105–119.
Aksu,A.E.,Calon,T.J.,Hall,J.,Mansﬁeld, S., Yaşar, D., 2005a. The Cilicia–Adana Basin
complex, Eastern Mediterranean: Neogene evolution of an active fore-arc basin in
an obliquely convergent margin. Marine Geology 221, 121–159.
Aksu,A.E.,Calon,T.J.,Hall,J.,Yaşar, D., 2005b. Origin and evolution of the Neogene
Iskenderun Basin, northeastern Mediterranean Sea. Marine Geology 221, 161–187.
Aksu, A.E., Hall, J., Yaltırak, C., 2009. Neogene evolution of the Anaximander Mountains
and Finike Basin at the junction of Hellenic and Cyprus Arcs, Eastern Mediterranean.
Marine Geology 258, 24–47.
Bağcı, U., Parlak, O., 2009. Petrology of the Tekirova (Antalya) ophiolite (Southern
Turkey): evidence for diverse magma generations and their tectonic implications
during Neotethyan-subduction. International Journal of Earth Sciences 98, 387–405.
Barka, A., Reilinger, R., 1997. Active tectonics of the eastern Mediterranean region:
deduced from GPS, neotectonic and seismicity data. Annali di Geoﬁsica 40, 587–610.
Barka, A., Reilinger, R.,Şaroğlu, F., Şengör, A.M.C., 1997. The Isparta Angle: its importance
in the neotectonics of the Eastern Mediterranean region. International Earth Sciences
21J. Hall et al. / Marine Geology 349 (2014) 1–23
Colloquium on the Aegean Region (IESCA-1995). Proceedings of the National Acade-
my of Sciences of the United States of America 1, 3–17.
Baroz, F., Bernoulli, D., Biju-Duval, B., Bizon, G., Bizon, J.-J.,Letouzey, J., 1978. Correlations
of the Neogene formations of the Florence Rise and of northern Cyprus: paleogeo-
graphic and structural implications. In: Hsü, K., Montadert, L., et al. (Eds.), Initial
Reports of the Deep Sea Drilling Project 42 (1). U.S. Government Printing Ofﬁce,
Washington, pp. 903–926.
Bassant, P., Van Buchem, F.S.P., Strasser, A., Görür,N., 2005. The stratigraphic architecture
and evolution of the Burdigalian carbonate—siliciclastic sedimentary systems of the
Mut Basin, Turkey. Sedimentary Geology 173, 187–232.
Ben Avraham, Z., Tibor, G., Limonov, A.F., Leybov, M.B., Ivanov, M.K., Torkarev, M.Y.,
Woodside, J.M., 1995. Structure and tectonics of the eastern Cyprean Arc. Marine
and Petroleum Geology 12, 236–271.
Besang, C., Eckhart, F.J., Harre, W., Kreuzer, G., Muller, P., 1977. Radiometrische
Alterbestimmung am neogenen Eruptivgesteinen der Türkei. Geologisches Jahrbuch
Biju-Duval, B., Letouzey, J., Montadert, L., 1978. Structure and evolution of the Mediterra-
nean Basins. In: Hsü, K., Montadert, L., et al. (Eds.), Initial Reports of the Deep Sea
Drilling Project 42, Part 1. U.S. Government Printing Ofﬁce, Washington, pp. 951–984.
Biryol, C.B., Beck, S.L.,Zandt, G., Özacar, A.A., 2011.Segmented Africanlithosphere beneath
the Anatolian region inferred from teleseismic P-wave tomography. Geophysical
Journal International 184, 1037–1057.
Blumenthal, M., 1963, (Compiler). Geological Map of Turkey, Konya-sheet, 1:500,000,
Institute of Mineral Resources and Exploration (MTA), Ankara, Turkey.
Bridge, C., Calon, T.J., Hall, J., Aksu,A.E., 2005. Salt tectonics in two convergent margin ba-
sins of the Cyprus Arc, northeastern Mediterranean. Marine Geology 221, 223–259.
Burton-Ferguson, R., Aksu, A.E., Calon, T.J., Hall, J., 2005. Seismic stratigraphy and
structural evolution of the Adana Basin, Eastern Mediterranean. Marine Geology
Calon, T.J.,Aksu, A.E., Hall, J., 2005a. The Neogene evolution of the Outer Latakia Basin and
its extension into the Eastern Mesaoria Basin (Cyprus), Eastern Mediterranean.
Marine Geology 221, 61–94.
Calon, T.J., Hall, J., Aksu, A.E., 2005b. The Oligocene–Recent evolution of the Mesaoria
Basin (Cyprus) and its western marine extension, Eastern Mediterranean. Marine
Geology 221, 95–120.
Çiner, A., Karabıyıkoğlu, M., Monod, O., Deynoux, M., Tuzcu, S., 2008. Late Cenozoic sedi-
mentary evolution of the Anyalya Basin, southern Turkey. Turkish Journal of Earth
Sciences 17, 1–41.
Cleintaur, M.R., Knox, G.J., Ealey, P.J., 1977. The geology of Cyprus and its place in the
eastern Mediterranean framework. Geologie em Mijnbouw 56 (1), 66–82.
Cosentino, D., Schildgen, T., Cipollari, P., Faranda, C., Gliozzi, E., Hudáčková, N., Lucifora, S.,
Strecker, M., 2012. Late Miocene surface uplift of the southern margin of the Central
Anatolian Plateau, Central Taurides, Turkey. Geological Society of America Bulletin
124 (1/2), 133–145 (January/February 2012).
Dewey, J.F., Hempton, M.R., Kidd, W.S.F., Şaroğlu, F., Şengör, A.M.C., 1986. Shortening of
continental lithosphere: the neotectonics of eastern Anatolia —a young collision
zone. In: Coward, M.P., Ries, A.C.(Eds.), Collision Tectonics. Geological Society Special
Publication, 19, pp. 3–36.
Dewey, J.F., Şengör, A.M.C., 1979. Aegean and surrounding regions: complex multiplate
and continuum tectonics in a convergent zone. Bulletin of the Geological Society of
America 90, 84–92.
Deynoux, M., Çiner, A., Monod, O., Karabıyıkoğlu, M., Manatschal, G., Tuzcu, S., 2005.
Facies architecture and depositional evolution of alluvial fan to fan-delta complexes
in the tectonically active Miocene Köprüçay Basin, Isparta Angle, Turkey. Sedimentary
Geology 173, 315–343.
Droz, L., Kergoat, R., Cochonat, P., Berné, S., 2001. Recent sedimentary events in the
western Gulf of Lions (Western Mediterranean). Marine Geology 176, 23–37.
Eriş, K.K., Bassant, P., Ülgen, U.B., 2005. Tectono-stratigraphic evolution of an Early
Miocene incised valley-ﬁll (Derinçay Formation) in the Mut Basin, Southern Turkey.
Sedimentary Geology 173, 151–185.
Faccenna, C., Bellier, O., Martinod, J., Primallo, C., Regard, V., 2006. Slab detachment be-
neath eastern Anatolia; a possible cause for the formation of the North Anatolian
Fault. Earth and Planetary Science Letters 242, 85–97.
Flecker, R., Ellam, R.M., Müller, C., Poisson, A., Robertson, A.H.F., Turner, J., 1998. Applica-
tion of Sr isotope stratigraphy and sedimentary analysis to the origin and evolution of
the Neogene basins in the Isparta Angle, southern Turkey. Tectonophysics 298,
Follows, E.J., Robertson, A.H.F., 1990. Sedimentology and structural setting of reefal lime-
stones in Cyprus. In: Malpas, J., Moores, E.M., Panayiotou, A., Xenophontos, C. (Eds.),
Ophiolites, Oceanic Crustal Analogues. Proceedings of the Symposium “Troodos
1987”. Geological Survey Department. Ministryof Agriculture and Natural Resources,
Nicosia, Cyprus, pp. 207–215 (733 pp.).
Garcia-Castellanos, D., Villaseñor, A., 2011. Messinian salinity crisis regulated by compet-
ing tectonics and erosion at the Gibraltar arc. Nature 480, 359–363. http://dx.doi.org/
Garrison, R.E., Schreiber, B.C., Bernoilli, D., Fabricius, F.H., Kidd, R.B., Mélière, F., 1978.
Sedimentary petrology and structures of Messinian evaporite sediments in the
Mediterranean Sea, Leg 42A, Deep Sea Drilling Project. In: Hsü, K., Montadert, L., et
al. (Eds.), Initial Reports of the Deep Sea Drilling Project 42 (1). U.S. Government
Printing Ofﬁce, Washington, pp. 571–612.
Glover, C.P., Robertson, A.H.F., 1998a. Role of regional extension and uplift in the Plio-
Pleistocene evolution of the AksuBasin, SW Turkey. Journal of the Geological Society
of London 155, 364–387.
Glover, C.P., Robertson, A.H.F., 1998b. Neotectonic intersection of the Aegean and Cyprus
tectonic arcs: extensional and strike-slip faulting in the Isparta Angle, SW Turkey.
Tectonophysics 298, 103–132.
Gökçen, S.L., Kelling, G., Gökçen, N., Floyd, P.A., 1988. Sedimentology of a late Cenozoic
collisional sequence: the Misis Complex, Adana, southern Turkey. Sedimentary
Geology 59, 205–235.
Govers, R., Wortel, M.J.R., 2005. Lithosphere tearing at STEP faults: response to edges of
subduction zones. Earth and Planetary Science Letters 236, 505–523.
Güngör, T., 2013. Kinematics of the Central Taurides during Neotethys closure and
collision, the nappes in the Sultan Mountains, turkey. International Journal of Earth
Sciences 102, 1381–1402.
Hall, J., Aksu, A.E., Calon, T.J., Yaşar, D., 2005a. Varying tectonic control on basin develop-
ment at an active microplate margin: the Iskenderun–Latakia Basin complex, Eastern
Mediterranean. Marine Geology 221, 15–60.
Hall, J., Calon, T.J., Aksu, A.E., Meade, S.R., 2005b. Structural evolution of the Latakia Ridge
and Cyprus Basin at the front of the Cyprus Arc, Eastern Mediterranean Sea. Marine
Geology 221, 261–297.
Hall, J., Aksu, A.E., Yaltırak, C., Winsor, J.D., 2009. Structural architecture of the Rhodes
Basin: a deep depocentre that evolved since the Pliocene at the junction of Hellenic
and Cyprus Arcs, eastern Mediterranean. Marine Geology 258, 1–23.
Hayward, A.B., 1984. Sedimentation and basin formation related to ophiolite nappe
emplacement; Miocene, SW Turkey. Sedimentary Geology 40, 105–129.
Hiscott, R.N., Aksu, A.E., 1994. Submarine debris ﬂows and continental slope evolution in
front of Quaternary ice sheets, Bafﬁn Bay, Canadian Arctic. American Association of
Petroleum Geologists 78 (3), 445–460.
Hsü, K.J., Montadert, L., Bernouilli, D., Cita, M.B., Erikson, A., Garrison, R.G., Kidd, R.B.,
Mélières, F., Müller, C., Wright, R., 1978. History of the Mediterranean salinity crisis.
In: Hsü, K.J., Montadert, L., et al. (Eds.), Initial Rep. Deep Sea Drill. Proj., 42A. US
Government Printing Ofﬁce, Washington, DC, pp. 1053–1078.
Hüsing, S.K., Zachariasse, W.J., van Hinsbergen, D.J.J., Krijgsman, W., Inceöz, M.,
Harzhauser, M., Mandic, O., Kroh, A., 2009. Oligo-Miocene foreland basin evolution
in SE Anatolia: constraints on the closure of the eastern Tethys gateway. In: van
Hinsbergen, D.J.J., Edwards, M.A., Govers, R. (Eds.), Geodynamics of Collision and
Collapse at the Africa–Arabia–Eurasia Subduction Zone. Geol. Soc. Lond. Spec. Publ.,
311, pp. 1–7(132).
Ilgar, A., Nemec, W., 2005. Early Miocene lacustrine deposits and sequence stratigraphy of
the Ermenek Basin, Central Taurides, Turkey. Sedimentary Geology 173, 233–275.
International Oceanographic Commission, 1981. International Bathymetric Chart of the
Mediterranean (1:1,000,000 scale). Head Department of Navigation and Oceanogra-
phy, Leningrad, former USSR.
Işler, F.I., Aksu, A.E., Hall, J., Calon, T.J.,Yaşar, D., 2005. Neogene dev elopment of the Antalya
Basin, Eastern Mediterranean: an active fore-arc basin adjacent to an arc junction.
Marine Geology 221, 299–330.
Jolivet, L., Brun, J.-P., 2010. Cenozoic geodynamic evolution of the Aegean. International
Journal of Earth Sciences 99, 109–138.
Karabıyıkoğlu, M., Çiner, A., Monod, O., Deynoux, M., Tuzcu, S., Örçen, S., 2000.
Tectonosedimentary evolution of the Miocene Manavgat Basin, western Taurides,
Magmatism in Turkey and the Surrounding Area. Geol. Soc. Lond. Spec. Publ.,
173, pp. 271–294.
Karabıyıkoğlu, M., Tuzcu, S., Çiner, A., Deynoux, M., Örçen, S., Hakyemez, A., 2005. Facies
and environmental setting of the Miocene coral reefs in the late-orogenic ﬁll of the
Antalya Basin, western Taurides, Turkey: implications for tectonic control and sea-
level changes. Sedimentary Geology 173, 345–371.
Kelling,G., Gökçen, S.L., Floyd, P.A.,Gökçen, N., 1987. Neogenetectonics and plate conver-
gence in the eastern Mediterranean: new data from southern Turkey. Geology 15,
Kempler, D., Garfunkel, Z., 1994. Structure and kinematics in the northeastern
Mediterranean: a study of irregular plate boundary. Tectonophysics 234, 19–32.
King, H., 2013. Structural and Stratigraphic Evolution of the Western Antalya Basin,
Eastern Mediterranean Sea. (Unpublished MSc thesis) Memorial University of
Kissel, C., Poisson, A., 1986. Étude paleomagnetique prélininaire des formations
Cenozoique des Bey Dağları(Taurides occidentales - Turquie). Comptes Rendue,
Academy of Science. Paris 302 Ser. 11 (8), 343–348.
Kissel, C., Averbuch, O., Frizon de Lamotte, D., Manod, O., Allerton, S., 1993. First paleo-
magnetic evidence for a post-Eocene clockwise rotation of the western Taurus thrust
belt, eastof the Isparta reentrant (southwestern Turkey). Earth and Planetary Science
Letters 117, 1–14.
Koç, A., Kaymakcı, N., van Hinsbergen, D.J.J., Kuiper, K.F., Vissers, R.L.M., 2012. Tectono-
sedimentary evolution and geochronology of the Middle Miocene Altınapa Basin,
and implications for the Late Cenozoic uplift history of the Taurides, southern
Turkey. Tectonophysics 532–535, 134–155.
Kozlu, H., 1987. Structural development and stratigraphy of Misis–Andirin region.
Proceedings of the 7th Petroleum Congress of Turkey. Turkish Association of Petroleum
Geologists, pp. 104–116.
Lastras, G., Canals, M., Hughes-Clarke, J.E., Moreno, A., De Batist, M., Masson, D.G.,
Cochonat, P., 2002. Seaﬂoor imagery from the BIG′95 debris ﬂow, western Mediterra-
nean. Geology 30, 871–874.
Lefèvre, C., Bellon, H., Poisson, A., 1983. Présences de leucitites dans le volcanisme
pliocène de la région d'Isparta (Taurides occidentales, Turquie). Comptes Rendus de
l'Académie des Sciences 297, 367–372.
Mackintosh, P.W., Robertson, A.H.F., 2012. Sedimentary and structural evidence for two-
phase Upper Cretaceous and Eocene emplacement of the Tauride thrust sheets in
central southern Turkey. In: Robertson, A.H.F., Parlak, O., Ünlügenç, U.C. (Eds.),
Geological Development of Anatolia and the Easternmost Mediterranean Region.
Geol. Soc. Lond. Spec. Publ., 372 (http://dx.doi.org/10.1144/SP372.2).
Mascle, J., Le Cleach, A., Jongsma, D., 1986. The eastern Hellenic margin from Crete to
Rhodes: example of progressive collision. Marine Geology 73, 145–168.
22 J. Hall et al. / Marine Geology 349 (2014) 1–23
McClusky, S., Balassinian, S., Barka, A., Demir, C., Ergintav, S., Georgiev, I., Gurkan, O.,
Hamburger, M., Hurst, K., Kahle, H., Kastens, K., Kekelidze, G., King, R., Kotzev, V.,
Lenk, O., Mahmoud, S., Mishin, A., Nadariya, M., Ouzonis, A.M., Paradissis, D., Peter,
Y., Prilepin, M., Reilinger, R., Sanli, I., Seeger, H., Tealeb, A., Toksöz, M.N., Veis, G.,
2000. Global Positioning System constraints on plate kinematics and dynamics in
the eastern Mediterranean and Caucasus. Journal of Geophysical Research 105 (B3),
McClusky, S., Reilinger, R., Mahmoud, S., Ben Sari, D., Tealeb, A., 2003. GPS constraints on
Africa (Nubia) and Arabia plate motion. Geophysical Journal International 155,
Meijers, M.J.M., van Hinsbergen, D.J.J., Dekkers, M.J., Atlıner, D., Kaymarkcı, N., Langereis,
C.G., 2011. Pervasive Palaeogene remagnetization of the central Taurides fold-and-
thrust belt (southern Turkey) and implications for rotations in the Isparta Angle.
Geophysical Journal International 184, 1090–1112.
Monod, O., 1977. Recherces géologiques dans le Taurus occidentales au sud de Bey_ehir
(Turquie). (Thèse de Doctorat d'État ès Sciences) Université de Paris-Sud, Orsay,
France (442 pp.).
Monod, O., Kuzucuoğlu,C.,Okay,A.İ., 2006. A Miocene paleovalley network in the
western Taurus (Turkey). Turkish Journal of Earth Sciences 15, 1–23.
Moores, E.M., Twiss, R.J., 1995. Tectonics. W.H. Freeman, New York (415 pp.).
Morris, A., Robertson, A.H.F., 1993. Miocene remagnetization of carbonate platform and
Antalya Complex within the Isparta Angle, SW Turkey. Tectonophysics 220,
Ocakoğlu, F., 2002. Palaeoenvironmental analysis of a Miocene basin in the high Taurus
Mountains(southern Turkey) and its palaeogeographical and structuralsigniﬁcance.
Geological Magazine 139, 473–487.
Okay, A.I., Özgül, N., 1984. HP/LT metamorphism and the structure of the Alanya Massif.
In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution of the Eastern
Mediterranean. Geol. Soc. Lond. Spec. Publ., 17, pp. 429–440.
Papazachos, B.C., Papaioannou, Ch.A., 1999. Lithospheric boundaries and plate motions in
the Cyprus area. Tectonophysics 308, 193–204.
Piper, J.D.A., Tatar, O., Gürsoy, H., Koçbulut, F., Mesci, B.L., 2006. Paleomagnetic analysis of
neotectonic deformation in the Anatolian accretionary collage, Turkey. Geological
Society of America, Special Paper 409, 417–439.
Poisson, A., Wernli, R., Sağular, K., Temiz, H., 2003a. New data concerning the age of the
Aksu Thrustin the south of the Aksu valley,Isparta Angle (SW Turkey): consequences
for the Antalya Basin and the Eastern Mediterranean. Geological Journal 38, 311–327.
Poisson, A., Yağmurlu, F., Bozcu, M., Şentürk, M., 2003b. New insights on the tectonic set-
ting and evolution around the apex of the Isparta Angle (SW Turkey). Geological
Journal 38, 257–282.
Poisson, A., Orszag-Sperber, F., Koson, E., Bassetti, M.-A., Müller, C., Wernli, R., R ouchy,
J.-M., 2011. The Late Cenozoic evolution of the Aksu Basin (Isparta Angle: SW
Turkey). New Insights. Bulletin de la Societe Geologique de France 182, 133–148.
Price, S., Scott, B., 1994. Fault block rotations at the edge of a zone of continental
extension, southwest Turkey. Journal of Structural Geology 16, 381–392.
Robertson, A.H.F., 1998. Mesozoic–Tertiary tectonic evolution of the Easternmost
Mediterranean area: integration of marine and land-based evidence. In: Robertson,
A.H.F., Emeis, K.-C., Richter, C., Camerlenghi, A. (Eds.), Proceedings of the Ocean
Drilling Program, Scientiﬁc Results, 160, pp. 723–782.
Robertson, A.H.F., Eaton, S., Follows, E.J., Payne, A.S., 1995. Depositional processes and
basin analysis of Messinian evaporites in Cyprus. Terra Nova 7, 233–253.
Robertson, A.H.F., Poisson, A., Akinci, Ö., 2003. Developments in research concerning
Mesozoic–Tertiary Tethysand neotectonics in the IspartaAngle SW Turkey. Geological
Journal 38, 195–234.
Robertson, A.H.F., Woodcock, N.H., 1986. Therole of Kyrenia Range lineament, Cyprus, in
the geological evolution of the eastern Mediterranean area. Philosophical Transac-
tions of the Royal Society of London, Series A 317, 141–177.
Rouchy, J.M., Caruso, A., 2006. The Messinian salinity crisis inthe Mediterranean basin:a
reassessment of the data and an integrated scenario. Sedimentary Geology 188–189,
Ryan, W.B.F., Carbotte, S.M., Coplan, J.O., O'Hara, S., Melkonian, A., Arko, R., Weissel, R.A.,
Ferrini, V., Goodwillie, A., Nitsche, F., Bonczkowski, J., Zemsky, R., 2009. Global
multi-resolution topography synthesis. Geochemistry, Geophysics, Geosystems.
http://dx.doi.org/10.1029/2008GC002332 (10: Q03014).
Şafak, Ü., Kelling, G., Gökçen, N.S., Gürbüz, K., 2005. The mid-Cenozoic succession and
evolution of the Mut basin, southern Turkey, and its regional signiﬁcance. Sedimen-
tary Geology 173, 121–150.
Şaroğlu, F., Boray, A., Emre, O., 1987.Active faults of Turkey. Mineral Research Exploration
Institute (MTA), Turkey, unpublished report 8643, 394 pp.
Satur, N., Kelling, G., Cronin, B.T., Hurst, A., Gürbüz, K., 2005. Sedimentary architecture of a
canyon-style fairway feeding a deep-water clastic system, the Miocene Cingöz
Formation, southern Turkey: signiﬁcance for reservoir characterisation and model-
ling. Sedimentary Geology 173, 91–119.
Savaşçın,M.Y., Francalanci, L., Innocenti, T., Manetti, P., Birsoy,R., Dağ, N., 1995. Miocene–
Pliocene potassic–ultrapotassic volcanism of the Afyon–Isparta region (central-
western Anatolia, Turkey). Petrogenesis and geodynamic implications: International
Earth Sciences Colloquium on the Aegean Region (IESCA-1995), Proceedings, V-II,
Schildgen, T.F., Cosentino, D., Bookhagen, B., Dudas, D., Echtler,H., Niedermann, S., Radeff,
G., Strecker, M., Yıldırım, C., Hudáčková, N., VAMP, 2011. Changing ratesand patterns
of surface uplift at the southern margin of the Central Anatolian plateau (Turkey): new
data from marine stratigraphy and cosmogenic nuclide dating of river terraces. Geo-
physical Research Abstracts 13 (EGU2011-4113-1, 2011, EGU General Assembly 2011).
Şenel, M., 1997 a. Geological Map of Fethiye, L8 quadrangle, No:2, 1:100,000 General
Directorate of Mineral Research and Exploration, Ankara, Turkey, 22 pp.
Şenel, M., 1997 b. Geological Map of Fethiye, M8 quadrangle, No:4, 1:100,000 General
Directorate of Mineral Research and Exploration, Ankara, Turkey, 15 pp.
Şenel, M., Bölükbaşı, A.S., 1997. Geological Map of Fethiye, M9 quadrangle, No:5,
1:100,000. General Directorate of Mineral Research and Exploration, Ankara,
Turkey, 11 pp.
Şengör, A.M.C., Görür,N., Şaroğlu, F., 1985. Strike-slip faulting and related basin formation
in zones of tectonic escape: Turkey as a case study. Society of Economic Paleontolo-
gists and Mineralogists. Special Publication 37, 227–264.
Şengör, A.M.C., Yılmaz, Y., Ketin, I., 1980. Remnants of a pre-Late Jurassic ocean in
northern Turkey: fragments of Permian–Triassic paleo-Tethys? Bulletin of the Geo-
logical Society of America 91, 599–608.
Sunder, M., 1982. Kırka (Eskişehir) çevresinin jeolojisi ve Sarıkaya borat yataklarının
oluşumu. TÜBITAK 7. Bilim Kongresi, Yerbilimleri Seksiyonu Tebliğler Kitabı. 105–117.
Toprak, Y., Gül, M., Yaman, S., 2009. Miocene lacustrine succession of the Hoyran Lake
Basin, Isparta, southwest Turkey. Acta Geologica Polonica 59, 245–259.
Uffenorde, H., Lund, J.J., Georgi, K.H., 1990. Biostratigraphyof the Neogene in theIskenderun
Basin. Proceedings of the 8th Petroleum Congress of Turkey. Turkish Association of
Petroleum Geologists, pp. 363–370.
van Hinsbergen, D.J.J.,Schmid, S., 2012. Map view restoration ofAegean–West Anatolian
accretion and extension since the Eocene. Tectonics 31. http://dx.doi.org/10.1029/
van Hinsbergen, D.J.J., Krijgsman, W., Langereis, C.G., Cornée, J.J., Duermeijer, C.E., van
Vugt, N., 2007. Discrete Plio-Pleistocene phases of tilting and counterclockwise rota-
tion in the southeastern Aegean arc (Rhodos,Greece): Early Pliocene formationof the
south Aegean left-lateral strike-slip system. Journal of the Geological Society of
London 164, 1133–1144.
van Hinsbergen, D.J.J., Dekkers, M., Koç,A., 2010. Testing Miocene remagnetisation of Bey
Dağları: timing and amount of Neogene rotations in SW Turkey. Turkish Journal of
Earth Sciences 19, 123–156.
Waldron, J.W.F., 1984. Structural history of the Antalya Complex in the ‘Isparta Angle’,
southwest Turkey. In: Dixon, J.E., Robertson, A.H.F. (Eds.), The Geological Evolution
of the Eastern Mediterranean. Geological Society Special Publication, 17, pp. 273–286.
Weiler, Y., 1969. The Miocene Kythrea ﬂysch basin in Cyprus. Giornale di Geologia 35,
Williams, G.D., Ünlügenç, U.C., Kelling, G., Demirkol, C., 1995. Tectonic controls on strati-
graphic evolution of the Adana Basin, Turkey. Journal of the Geological Society of
London 152, 873–882.
Woodside, J.M., Mascle, J., Zitter, T.A.C., Limonov, A.F., Ergün, M., Volkonskaia, A.,
Yağmurlu, F., Savaşçın,Y.,Ergün,M.,1997.Relation of alkaline volcanism and active
tectonism within the evolution of the Isparta Angle, SW Turkey. Journal of Geology
Yalçın, N.M., Görür, N., 1984. Sedimentological evolution of the Adana Basin.In: Tekeli, O.,
Göncüoğlu, M.C. (Eds.), Proceedings of the International Symposium on the Geology
of the Taurus Belt, Ankara (165–172 pp.).
Yılmaz, Y., 1993. New evidence and model on the evolution of the southeast Anatolian
orogen. Geological Society of America Bulletin 105, 251–271.
Yılmaz, Y., Gürpınar, O., Yiğitbaş, E., 1988. Tectonic Evolution of the Miocene Basins at the
Amanos Mountains and the MaraşRegion, 1/1. Bulletin Turkish Association of
Petroleum Geologists 52–72.
Zitter, T.,Woodside, J., Mascle, J., 2003.The Anaximander Mountains:a clue to the tectonics
of southwest Anatolia. Geological Journal 38, 375–394.
23J. Hall et al. / Marine Geology 349 (2014) 1–23