Multiple rifting pulses and sedimentation pattern in the C ameli Basin, southwestern Anatolia, Turkey
←
→
Page content transcription
If your browser does not render page correctly, please read the page content below
Sedimentary Geology 173 (2005) 409 – 431
www.elsevier.com/locate/sedgeo
Multiple rifting pulses and sedimentation pattern in the
Çameli Basin, southwestern Anatolia, Turkey
Mehmet Cihat Alçiçeka,*, Nizamettin KazancVb,c, Mehmet Özkula
a
Department of Geological Engineering, Pamukkale University, 20070 Denizli, Turkey
b
Department of Geological Engineering, Ankara University, 06100 Ankara, Turkey
c
Gebze Institute of Technology, 41400 Kocaeli, Turkey
Received 5 July 2002; received in revised form 11 November 2003; accepted 10 December 2003
Abstract
The neotectonic development of western Anatolia was characterized by the formation of numerous graben-type basins, which
have been well documented by general mapping, although the cause and timing of the Neogene regional tectonic extension remain
controversial. Previous interpretations of the origin and evolution of these Neogene basins were based mainly on regional-scale
tectonic inferences, rather than detailed basin-fill analysis. The present study of the terrestrial intramontane Çameli Basin in the
western Taurides combines detailed facies analysis with biostratigraphic dating (mammalian and molluscan fossils) and documents
three pulses of crustal extension that are reflected in changes in the palaeogeography and sedimentary architecture of the basin.
Development of the Çameli graben commenced in the Vallesian time (Early Tortonian), and is marked by alluvial-fan, fluvial
and lacustrine depositional systems, with freshwater molluscan fauna. A second pulse of tectonic extension occurred in the Late
Ruscinian time (Early–Middle Pliocene), producing a new normal fault that split the basin longitudinally into two
compartments. The lake environment expanded and deepened, coastal peat-forming mires developed and abundant mammal
fauna appeared by the Early Villanian time (Middle Pliocene), with the lacustrine deposits onlapping the basin-margin and
intrabasinal fault escarpments. The lacustrine environment subsequently shrank, as the progradation of axial river deltas and
basin-margin fan deltas caused water shallowing and shoreline regression. A third pulse of extension occurred at the end of
Villanian time (Late Pliocene), when the development of a new generation of normal faults further split the basin into still
narrower half-graben compartments. The third pulse of rifting is estimated to have accounted for little more than 10% of the
sub-basinal crustal extension, but caused the most striking changes in the basin palaeogeography and drainage pattern. The
inward development of the successive normal faults indicated a high-rate crustal extension. This is the first regional case study
of a terrestrial neotectonic graben employing detailed sedimentary facies analysis and mammal biostratigraphy and providing a
time-stratigraphic framework for the rifting pulses in western Anatolia.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Neotectonics; Rifting; Graben; Facies analysis; Neogene; Western Taurides
* Corresponding author. Tel.: +90 258 2134030x1513.
E-mail address: alcicek@pamukkale.edu.tr (M.C. Alçiçek).
0037-0738/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.sedgeo.2003.12.012
410 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
1. Introduction
The Early Miocene–Quaternary neotectonic exten-
sion in the western and southern to central parts of
Anatolia has led to the development of numerous fault-
bounded basins. Some of these extensional basins
were invaded by the sea in Burdigalian time (e.g., the
Aksu, Köprü, Manavgat, Adana, Mut and Ermenek
basins in southern Turkey; see Flecker et al., 1995,
2004; Glover and Robertson, 1998a; Tanar and
Gökçen, 1990; KarabVyVkoğlu et al., 2000; Satur et
al., 2000; Ilgar and Nemec, 2004) and hosted marine
sedimentation until the mid-Pliocene or even present
times (e.g., the outer Büyük Menderes Basin in
offshore western Turkey and the Cilicia Basin in
offshore southern Turkey; Aksu et al., 1987; Kelling et
al., 1987; Cronin et al., 2000), while others remained
fully terrestrial, hosting alluvial and lacustrine sed-
imentation (e.g., the inner Bqyqk Menderes, Burdur,
AlaYehir/Gediz, Simav, AkYehir-Afyon and Çameli
basins; see Price and Scott, 1991; Seyitoğlu and Scott,
1996; Seyitoğlu, 1997; Koçyiğit et al., 2000; Bozkurt,
2000; Alçiçek, 2001). Some of the grabens were
tectonically inverted by the mid-Miocene and/or Late
Pliocene pulses of Late- and post-orogenic compres-
sion (Flecker et al., 1995, 2004; Koçyiğit et al., 1999,
2000; KarabVyVkoğlu et al., 2000), some have been
cross-cut by younger faults (YVlmaz et al., 2000; Fig. 1. (A) Tectonic map of the NE Mediterranean region, showing
the location of the Çameli Basin (modified from Xengfr et al., 1985;
Xengfr, 1987), and most of these basins are no longer Armijo et al., 1996; Glover and Robertson, 1998b); (B) simplified
subsiding. The regional literature abounds in the geological map of the basin (based on Pamir, 1974; Xenel, 1997a).
mapping and stratigraphic documentation of these The area of the basin studied in detail is delineated by the dotted line.
Neogene basins (e.g., Becker-Platen, 1970; Erakman
et al., 1982; MeYhur and AkpVnar, 1984; Sfzbilir and extension, which all caused marked changes in the
Emre, 1990; Xenel, 1997a,b,c), but relatively few of basin’s internal palaeogeography and sedimentation
them have been studied in sufficient detail to pattern. This is the first regional case study of a
reconstruct the style of their development, subsidence terrestrial neotectonic graben involving mammal bio-
history and sedimentation pattern (e.g., Seyitoğlu and stratigraphic dating, which sheds new light on the
Scott, 1991; Price and Scott, 1991; Flecker et al., 1995, tectonic and sedimentation history of the extensional
2004; Glover and Robertson, 1998b; YVlmaz et al., Neogene basins in western Anatolia. Depending on the
2000; Bozkurt, 2000; KarabVyVkoğlu et al., 2000; regional tectonic model adopted (see Section 2), the
Koçyiğit et al., 2000; Alçiçek, 2001). chronostratigraphic framework derived from this
The present study documents the tectonic evolution study might be of regional or more local significance.
and basin-fill sedimentary architecture of the Çameli
Basin in southwestern Anatolia (Fig. 1), based on
detailed logging of outcrop sections, facies analysis 2. Neotectonic regional extension
and biostratigraphic dating. The basin hosted alluvial-
fan, fluvial and lacustrine depositional systems and The final stages of the Late Cretaceous–Miocene
evolved through three distinct phases of tectonic Tauride orogeny in southern Anatolia were accom-M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 411 Fig. 2. Geological map and transverse cross-section of the Çameli Basin; the geological map is modified from Xenel (1997a,b,c), and unpublished maps of the Mineral Research and Exploration Directorate of Turkey compiled by F. GöktaY and Y. Hakyemez, by permission of Y. Hakyemez, pers. commun., 1999. The numbers refer to the location of the measured sections shown in Fig. 7.
412 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
panied and directly followed by regional-scale
tectonic extension, which had a broadly N–S
direction in the western and southern part of
Anatolia and led to the development of numerous
fault-bounded intramontane basins (Xengör and
YVlmaz, 1981; Robertson and Dixon, 1984; Xengör
et al., 1985; Zanchi et al., 1993), including the
Çameli Basin discussed in the present paper (Figs. 1
and 2). The cause of this neotectonic crustal
extension has been widely disputed and remains
controversial (YVlmaz et al., 2000; Bozkurt, 2001;
Flecker et al., 2004). Some of the better-known
models for the neotectonic extension in western
Turkey are summarized below (see also Fig. 3):
According to the Dewey and Xengör (1979) model,
extension commenced in Late Serravallian time and
was caused by the westward tectonic escape of the
compound Anatolian craton (see also Xengör, 1979;
Xengör and YVlmaz, 1981; Xengör et al., 1985). In Le
Pichon and Angelier (1979) model, the Late Serraval-
lian onset of extension in Anatolia and the adjacent
Aegean region was related to the backarc tectonic
regime of the Cyprus and Hellenic subduction arcs to
the south. In the orogenic-collapse model of Dewey
(1988), Seyitoğlu and Scott (1991) and Gessner et al.
(2001), the extension began in the Late Oligocene–
Early Miocene time and was caused by the collisional
overthickening of crust along the İzmir–Ankara–
Erzincan Neotethyan suture, accompanied by a
pervasive crustal extension in the northern part of
the orogen due to the southeastward emplacement of
the Lycian allochthon. The combination of the
extension in the hinterland zone and coeval contrac-
tion in the Lycian foreland zone would then be a result
of the latest Oligocene–Miocene extensional collapse Fig. 3. The onset of neotectonic extension in western Anatolia, as
of the orogen, with the final southeastward movement postulated by Dewey and Xengfr (1979) and Dewey et al. (1986),
of the Lycian allochthon in the Late Miocene time by Le Pichon and Angelier (1979) and Meulenkamp et al. (1988),
(Collins and Robertson, 1998, 1999). According to and by Dewey (1988), Seyitoğlu and Scott (1991) and Gessner et al.
Xengör (1987), the north-trending grabens in western (2001). Note that the onset of extension in the Çameli Basin (present
study) corresponds best with the first two views.
Anatolia were formed under a N–S compressional
regime in Early Miocene time and were subsequently
cut by east-trending grabens in the Late Miocene. The NE-trending Çameli Basin began to subside in
However, more recent studies of the east-trending Early Tortonian time (Fig. 3), which might correspond
grabens (Seyitoğlu and Scott, 1992; Hetzel et al., to the onset of extension according to the first two
1995) and north-trending ones (Seyitoğlu and Scott, models or a relatively late, delayed orogen collapse
1994) indicate their simultaneous development in according to the third model. The present study shows
Early Miocene time (Seyitoğlu and Scott, 1996; that the development of the Çameli Basin involved
Seyitoğlu et al., 2002). three recognizable pulses of rifting, which have beenM.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 413
biostratigraphically dated and may provide a time- macro-mammal fossil was found near ElmalVyurt in the
stratigraphic framework for the evolution of the southern part of the basin (Xenel, 1997a; Saraç and
neotectonic grabens in western Anatolia. Xen, pers. commun., 1999) and determined to be of
Vallesian (Early Tortonian) age (Alçiçek, 2001). In the
present study, numerous other micro- and macro-
3. Basin stratigraphy mammal fossils have been found elsewhere in the
basin (Table 1 and Fig. 2; Alçiçek, 2001). These fossils
3.1. Basement units occur mainly in the lake-margin deposits of peat-
forming mires (the ÇamlVbel and Ericek localities),
The basement beneath the Çameli Basin consists of overbank deposits of the fluvial facies (the ElmalVyurt
the Lycian allochthon (Graciansky, 1972), which were locality) and distal fines of the alluvial fan facies (the
emplaced from the northwest in the end-Cretaceous to BVçakçV locality, see Table 1 and Fig. 7), and include
Neogene times and completed its southeastward rodent and fish remains (teeth, bones); some the bones
movement in Late Langhian times (Hayward, 1984; are fractured, but the fossils show no obvious evidence
Kissel et al., 1993; Collins and Robertson, 1998). The of reworking or redeposition. The lake margin deposits
ophiolite and marble thrust-sheets are locally uncon- are also typically rich in molluscan shells and plant
formably covered by Early Miocene deposits that was remains, which all indicate a freshwater palaeolake
first described by AltVnlV (1955), reaching up to 570 m environment. This new palaeontological evidence
in thickness, comprising alluvial redbeds overlain by indicates a Late Miocene (Tortonian) to Late Pliocene
shallow-marine sandstones, marls and fossiliferous (Gelasian) age of the Çameli Formation (Fig. 4).
limestones, which were deposited in a piggy-back The Çameli Formation has been divided into three
setting and transported jointly with the nappes. This lithostratigraphic subunits by Alçiçek (2001), referred
supra-allochthonous sedimentary cover is regarded to as the Derindere, KumafYarV and Değne Members,
here as a part of the bedrock succession (Fig. 2).
Table 1
3.2. The basin-fill succession Mammal fossils and their localities in the Çameli Basin (for areal
and stratigraphic locations, see Figs. 2 and 4)
The basin-fill deposits in the Çameli graben are Locality and Mammal fossils Faunal zone, time span
fully terrestrial, Late Miocene to Late Pliocene in age geographic and epoch (after
and ca. 500 m in thickness, unconformably overlying coordinates Steininger et al., 1996)
the allochthonous bedrock. The basin-fill succession BVçakçV Mimomys pliocaenius Zone MN 17
is primarily bounded by the graben’s main dip-slip N37800V53U Apodemus dominans 2.6–1.8 Ma
normal faults: the Dirmil Fault on the SE side and the E29817V57U Micromys praeminutus Late Villanian
Late Pliocene
Bozdağ Fault on the NW side. The deposits have been (Gelasian)
tilted by the successive pulses of rifting and graben ÇamlVbel Rodentia–Arviccolidae Zone MN 15–16
formation, and are unconformably overlain by non- N37810V27U Mimomys sp. 3.5–2.5 Ma
tilted Quaternary alluvial deposits, mainly less than 20 E29822V21U Late Ruscinian–Early
m in thickness (Fig. 2). Villanian
Middle–Late Pliocene
The bulk of the basin-fill succession is represented (Piacenzian–Gelasian)
by the Çameli Formation, which consists of alluvial- Ericek Mimomys occitanus Zone MN 15
fan, fluvial and lacustrine deposits. Its contact with the N37804V12U Apodemus dominans 3.8–3.2 Ma
bedrock involves both normal faults and an angular, E29811V55U Orientalomys similis Late Ruscinian
erosional unconformity (Fig. 2). This lithostratigraphic Pseudomeriones Early–Middle Pliocene
tchaltaensis (Zanclean–Piacenzian)
unit was previously mapped as the bNeogene coverQ of ElmalVyurt Perissodactyla–Equidae Zone MN 9–10
the nappes and assumed to be of Pliocene age (Becker- N36853V17U Hipparion cf. 10.8–8.5 Ma
Platen, 1970; Erakman et al., 1982; MeYhur and E29823V34U Primigenium sp. Vallesian
AkpVnar, 1984), although it neither dated nor sed- earliest Late Miocene
imentologically studied. More recently, a terrestrial (Early Tortonian)414 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
Fig. 4. Simplified stratigraphy of the Çameli Basin (not to scale).
which consist of alluvial-fan, fluvial and lacustrine deposition in the basin (Table 2). These facies form
deposits, respectively. The three members overlie one seven major spatial assemblages, or facies associa-
another in a layer-cake style in the central part of the tions, which are described below and attributed to
basin, but interfinger laterally with one another near different depositional environments of the evolving
the basin margins. basin.
4.1. Alluvial-fan facies association
4. Facies analysis
This sedimentary assemblage comprises facies
The deposits of the Çameli Formation have been Gm(a), Gm(s), Sh(s) and Sm (Table 2) and constitutes
divided into 18 sedimentary facies, ranging from the Derindere Member of the Çameli Formation (Fig.
subaerial to subaqueous and provide sedimentological 5A). The deposits form two coarsening-upward
information on the principal modes of sediment successions, 20–90 m thick (Figs. 5A and 6C, andM.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 415
Table 2 Table 2 (continued)
Lithofacies of the Çameli Formation (facies code modified from Facies Description Interpretation
Miall, 1977)
Sh(w) Medium- to coarse-grained Lake beach/shoreface
Facies Description Interpretation sandstones with planar parallel deposits (High wave
Gm(a) Matrix- to clast-supported, Cohesive debris-flow stratification and common energy)
mud-bearing, poorly sorted, deposits low-angle internal truncations;
reddish-brown, pebble to forming beds 5–20 cm thick,
coarse cobble conglomerates with a lateral extent of several
with scattered boulders; tens of metres, intercalated
forming beds 30–90 cm thick with lacustrine facies Sr(w)
and a few tens of metres in and Fl(w).
lateral extent, with sharp bases Sr(s) Fine- to medium-grained Overbank stream-flood
and no grading or weak, sandstones with unidirectional deposits
coarse-tail inverse grading; ripple cross-lamination;
commonly alternating with the forming lenticular or
mudstone facies Fm. wedge-shaped beds b20 cm
Gm(s) Poorly sorted, clast-supported, Stream channel-floor thick, lateral to channel-fill
sand-filled pebble to fine lags (pavements) deposits.
cobble conglomerates; forming Sr(w) Fine- to coarse-grained Lake beach/shoreface
beds 5–50 cm thick, non-graded, sandstones with bidirectional deposits (low wave
with erosional bases and ripple cross-lamination; energy)
sheet-like geometry. forming sheet-like beds
Gh Moderately sorted, faintly Gravelly sheet-flood b10 cm thick, intercalated
parallel-stratified, deposits with the lacustrine mudstone
clast-supported, pebble to fine facies Fl(w) and sandstone
cobble conglomerates; forming facies Sh(w).
beds b30 cm thick and a few Sl Medium- to coarse-grained Lateral accretion
tens of metres in lateral extent, sandstones with low-inclined deposits of stream
with erosional bases and no (b208) cross-stratification channels (point bars)
grading or weak normal transverse or normal to the
grading. palaeocurrent direction;
Gp Clast-supported, low-angle Longitudinal channel forming lenticular sets b2 m in
(b208) planar cross-stratified bars thickness, with erosional bases,
pebble conglomerates with underlain by the channel-lag
scattered cobbles; forming beds facies Gm(s).
5–20 cm thick and fining in Sp Medium- to coarse-grained Transverse bar deposits
downflow direction. sandstones, well sorted, with of stream channels
Sm Medium- to very Deposits of planar cross-stratification;
coarse-grained sandstones, hyperconcentrated forming beds 6–20 cm thick.
non-stratified, with scattered (pseudoplastic) Sf Foresets of parallel, tangential Gilbert-type delta
granules and/or pebbles at the subaerial flows sandstone cross-strata, 0.5–3 m foresets
base; forming beds 2–20 cm thick, including intrastratal
thick, with sharp, slightly slump features, overlying
erosional bases and common lacustrine mudstone facies
plant-root casts at the top. Fl(w) and covered with alluvial
Sh(s) Medium- to very Sandy sheet-flood topset deposits.
coarse-grained sandstones with deposits Fl(s) Massive to faintly laminated, Suspension fallout
planar parallel stratification; reddish-brown mudstones with deposits of waning
forming beds 5–15 cm thick, scattered sand grains and fine floods, accumulated in
with slightly erosional bases, granules; forming composite local overbank
a lateral extent of a few tens of units b15 m thick, intercalated flood-basins
metres and plant-root casts, with the assemblage of fluvial
intercalated with the alluvial facies Gm(s), Sl, Sh(s) and
facies Fl (s). Sr(s).
(continued on next page)416 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
Table 2 (continued) fluvial facies association in the basinward direction.
Facies Description Interpretation The deposits of facies Gm(a) predominate in a lower
Fl(w) Thinly laminated to massive, Offshore lacustrine part of the Derindere Member, abutting against the
bioturbated mudstones, clayey deposits (open lake) bedrock with a sharp and steep boundary, which
to silty, with marl and marly suggesting a fault contact (Fig. 2, along the Bozdağ
limestone interbeds and fresh
Fault, and Fig. 9B). In the middle and upper parts of
water molluscan shells;
forming beds 5–10 cm thick, the member, this basin-margin alluvium overlies the
stacked into successions b22 m bedrock with an erosional unconformity (Fig. 2, see
thick, intercalated with the east of SarVkavak and KarabayVr). This relation
wave-worked lacustrine facies indicates that the contact of the alluvial-fan deposits
Sh(w) and Sr(w).
with the margin varies from faulted to erosional and
Fm Massive, reddish-brown Waning-flood deposits
mudstones with dessication accumulated in dormant represents a depositional backlapping of the margin
cracks and/or calcrete nodules, alluvial-fan zones by the aggrading alluvial fans (Fig. 2, see near SuçatV).
forming beds 3–4 m thick,
intercalated with facies Gm(a). 4.2. Fluvial facies association
C Coal beds, 10–15 cm thick, Peat-forming mires in
intercalated with facies Fl(s) or stream overbank and
more commonly Fl(w), lake-margin areas These deposits, described below, constitute the
forming interbedded bulk of the KumafYarV Member of the Çameli
successions b5 m in thickness, Formation (Figs. 5D,E and 6A,B) and are attributed
containing rodent mammal to the graben’s axial system of perennial rivers. Based
remains and fish teeth and
on the range and spatial grouping of facies, three
bones.
P Laminated limestones with Offshore lacustrine subassociations have been distinguished and inter-
common bird-eye structures deposits preted to be the deposits of braided rivers, meander-
and gastropod fossils, forming (ephemeral lake) ing rivers and overbank peat-forming mires.
beds a few centimetre thick,
stacked into composite units
4.2.1. Braided-river deposits
13–30 m in thickness.
T Tufa deposits, associated with Deposits of fault-related This facies assemblage forms the lower part of
synsedimentary faults and coldwater springs the KumafYarV Member and consists of facies Fl(s),
found as two occurrences in the Sm, Sp, Gm(s), Gp and Gh (Table 2, Figs. 5D and
basin: a succession of 6A). The deposits form a succession 10–120 m
intercalated tufa and mudstone
thick, which overlies and laterally interfingers with
facies Fl(s), ca. 60 m thick, and
a pure tufa unit 7 m thick, both the alluvial-fan facies association, while passing
passing laterally into facies P distally into lake-margin deltaic deposits. The
and/or Fl(w). abundance of the channel-floor lag deposits of facies
Gm(s), transverse-bar deposits of facies Sp and
longitudinal-bar deposits of facies Gp indicates
the lower part of the ElmalVyurt and KarabayVr logs bedload-dominated braided rivers (Miall, 1985;
and the uppermost part of the ÇamlVbel and KavalcVlar Gibling et al., 1998).
log in Fig. 7). Facies Gm(a) predominates and has the
greatest lateral extent, on the order of several hundred 4.2.2. Meandering-river deposits
metres. Its bulk thickness increases toward the basin- This facies assemblage consists of mud-poor
margin faults. This facies association is relatively rich sandstones. Grain size varies from medium to very
in coarse conglomerates, abounds in the debris-flow coarse sand, but individual beds are moderately to
deposits of facies Gm(a), sheet-flood deposits of well sorted. The association consists of facies Sp,
facies Gh and Sh(s) and hyperconcentrated-flow Sl, Sr(s), Sh(s) and Fl(s) (Table 2, Figs. 5E and
deposits of facies Sm, and is thought to represent 6B), and constitutes a large portion of the upper
basin-margin alluvial fans (Rust, 1979; Evans, 1991; part of the KumafYarV Member. The evidence of
Ridgway and DeCelles, 1993). It intercalates with a lateral-accretion bars (facies Sl) and the well-M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 417 Fig. 5. Facies associations of the Çameli Formation: (A) alluvial-fan deposits of Derindere Member (lower part of the KarabayVr log in Fig. 7); (B) open-lake deposits (middle part of the Evciler log in Fig. 7); and (C) ephemeral-lake deposits of Değne Member (uppermost part of Fig. 6B and the SarVkavak log in Fig. 7); (D) braided-river deposits (upper part of Fig. 6A and middle part of the KavalcVlar log in Fig. 7); (E) meandering-river deposits of KumafYarV Member (lower part of Fig. 6A and the SarVkavak log in Fig. 7); (F) axial deltaic deposits (lower part of Fig. 6A and middle part of the KavalcVlar log in Fig. 7); and (G) fan-delta deposits of Değne Member (lower part of Fig. 6A and upper part of KavalcVlar log in Fig. 7) and (H) peat-mire deposits of KumafYarV Member (Pickaxe length is 80 cm).
418 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 Fig. 6. Selected logs from the Çameli basin-fill succession, showing main facies associations and their alternations. These logs provide more detail for the generalized logs displayed in Fig. 7: (A) the middle part of the KavalcVlar log, (B) the upper part of the SarVkavak log and (C) the upper part of KavalcVlar log. The mammal fossil localities are also indicated. See Figs. 2 and 8 for the locations of the logs. developed overbank deposits of facies Sr(s) and the lacustrine facies associations of the Değne Fl(s) indicate meandering rivers (Allen, 1983; Member. Collinson, 1996). The deposits form a fining- upward succession 150 m in thickness, intercalating 4.2.3. Peat-mire deposits with the braided-river facies association and passing This facies assemblage consists of the mudstone distally into lake-margin deltaic deposits. This facies facies Fl(s) and/or Fl(w) intercalated with the coal association is overlain by coal-bearing deposits and facies C (Figs. 5H and 6B, Table 2) and directly
M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 419 Fig. 7. Correlation of the measured outcrop sections of the Çameli basin-fill succession; for the locations of logs, see Figs. 2 and 8. The mammal fossil localities are also indicated. overlies the meandering-river facies association. The 4.3. Open-lake facies association deposits abound in plant remains, including leaves, twigs and root casts. Micro- and macro-remains of This facies association constitutes a large part of terrestrial mammal fauna (rodent bones and teeth), the Değne Member of the Çameli Formation (Fig. 5B) gastropod shells and the bones and teeth of freshwater and consists mainly of the muddy deposits of facies fishes are common in this facies association, which has Fl(w), with a varying amount of marl and clayey a thickness of up to 25 m and a lateral extent of several micritic limestone interbeds (Fig. 6B). Evaporites, hundred metres. The coal beds are autochthonous, such as gypsum or halite, are lacking. The deposits underlain by seatearths and the assemblage as a whole contain freshwater molluscan fossils, including Mela- indicates peat-forming mires developed in a slackwater nopsis (Lyrcaea) narzolina Bonelli, Pseudamnicola environment (Belt et al., 1984; McCabe, 1984), (Sandria) kochi Brusina, Pseudamnicola margarita developed as an overbank alluvial floodplain increas- Neumayer and Pseudamnicola margarita nuda Jeke- ingly inundated by an expanding lake. lius (S.K. YeYilyurt, pers. commun., 2001). These
420 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
deposits form a monotonous succession with a total foreset deposits of facies Sf, which commonly contain
thickness of up to 225 m and a lateral extent of a few fish teeth and bones and are overlain by fluvial topset
thousand meters. They overlie the peat-mire deposits, facies Gm(s), Gp, Sp, Sm, Sh(s) and Sr(s) (Fig. 5F).
interfinger laterally with deltaic deposits and are The foreset beds are subaqueous sediment gravity-
overlain by the ephemeral-lake facies association flow deposits, sandy to gravelly and a few centimetres
described below. The laminated mudstones and marls to a few decimetres thick, intercalated with wave-
are thought to have been deposited by sediment worked sandstone layers. The beds have a tangential
suspension fallout in quiet water. The carbonate geometry and an inclination from less than 58 to more
component is attributed to episodic chemical sed- than 308, averaging 208. The foreset itself is 5–12 m
imentation, rather than detrital provenance. The thick and, together with the fluvial topset, forms a
molluscan fauna suggests an open-lake, probably succession that reaches 45–55 m in thickness (Figs.
supplied with water by perennial rivers. The fauna 5F and 6). This facies assemblage passes laterally into
and lack of evaporites indicate low to moderate water open-lake deposits towards the basin centre and into
salinity, and the carbonate precipitation may thus fluvial deposits in the opposite direction. The topset
reflect periods of minimum sediment supply, perhaps consists of meandering-river deposits (Evciler area) or
corresponding to drier climatic phases or the lake- braided-river deposits (KavalcVlar area, Figs. 6A and
level highstands (cf. Ilgar and Nemec, 2004). 7). This facies association represents Gilbert-type
river deltas (see Colella and Prior, 1990) formed at
4.4. Ephemeral-lake facies association the margin of a relatively deep lake.
This facies association constitutes the upper part of 4.6. Fan-delta facies association
the Değne Member and consists of laminated mud-
stones and limestones of facies Fl(w) and P (Figs. 5C This facies assemblage, forming sandy successions
and 6B), intercalated with and passing laterally into 15–40 m thick, occurs in the uppermost part of the
the lake-shoreline sandstone facies Sr(w) or the deltaic Değne Member and interfingers with muddy deposits
deposits described further below. The deposits are (facies Fl) of open lake facies association. The
bioturbated, contain desiccation cracks, bird-eye association consists of the wave-worked sandy facies
calcite, plant remains and fewer freshwater gastropod Sh(w) and Sr(w), or subaqueous foreset facies Sf,
shells, and form a succession up to 20 m in thickness. overlain by alluvial-fan deposits, typically dominated
The succession is overlain by deltaic and fan-deltaic by facies Gm(a), Gh, Sh(s) and Sm (Figs. 5G and 6C).
deposits along the basin axis and margins, respec- This facies association passes laterally into the
tively. The sedimentary and palaeontological eviden- gravelly, reddish-colored alluvial fan deposits towards
ces point to an ephemeral lake environment (Anadón the basin margins. Based on this lateral relationship
et al., 1989). and the facies assemblage, these deposits are thought
to represent lacustrine fan deltas (see Nemec and
4.5. Deltaic facies association Steel, 1988), coeval with the deltaic deposits of the
previous facies association and formed by the
This facies assemblage occurs in the uppermost progradation of basin-margin alluvial fans into a lake
part of the Değne Member and consists of the sandy that occupied the basin centre. The subassociation
Fig. 8. Bedrock geology (A) and interpreted palaeogeographic evolution of the Çameli Basin (based on Xenel, 1997a): (B) the basin opens by a
rifting pulse in Vallesian (Early Tortonian) time (10.8–8.5 Ma), leading to deposition of the alluvial facies of Derindere Member, the fluvial
facies of KumafYarV Member and the ephemeral-lake facies of Değne Member; (C) the second pulse of rifting, in the Late Ruscinian (Zanclean-
Piacenzian), splits the graben axially into two compartments, leading to deposition of the lower tufa facies association of the basin succession
along the SarVkavak-KumafYarV Fault; (D) subsidence causes deepening and expansion of the lake and deposition of the open-lake facies
association of Değne Member; (E) the lake shrinks due to the progradation of alluvial fans and basin-axis fluvial systems, as typified by the fan-
delta and deltaic facies associations of Değne Member; and (F) the third pulse of rifting, at the end of the Villanian (Gelasian), splits the basin
into still narrower half-graben compartments, predominantly represented by the upper tufa facies association of the basin succession exposed
along the Uzunoluk-Çameli Fault.M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 421
422 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
composed of alluvial-fan deposits passing lakewards 5.1. The first rifting pulse and opening of the Çameli
into facies Sh(w) and Sr(w) is interpreted as a shoal- Basin (Late Miocene)
water fan delta, characterized by a wave-worked front,
whereas the subassociation involving foreset facies Sf The sedimentation in the Çameli Basin com-
overlain by alluvial-fan deposits is interpreted as a menced in the Vallesian (Early Tortonian) time
Gilbert-type fan delta, characterized by a steep, (10.8–8.5 Ma), as is indicated by the remains of
avalanching underwater slope. The two fan-delta mammal Perissodactyla-Equidae Hipparion cf. pri-
varieties, commonly coeval in the basin, reflect migenium sp. (MN zones 9–10) found in the lower-
different water depths of lake nearshore zone. most part of the basin-fill succession near ElmalVyurt,
ca. 25 km south of Çameli (Figs. 2 and 7). The graben
4.7. Tufa (spring-outflow) facies association began to subside by the formation of the Bozdağ Fault
to the northwest and the Dirmil Fault to the southeast
The carbonate tufa facies T (Table 2) occurs both (Figs. 8B and 9B,C), which are thought to be the
below and above the Değne Member, where it is earliest normal faults involved in the development of
intercalated with the laminated mudstones, clayey the Çameli Basin. To the southwest and northeast, the
limestones and marls of facies Fl(w) (Figs. 6A and 7). Çameli Formation unconformably overlies the bed-
The stratigraphically lower facies assemblage is rock and local Early Miocene deposits (Fig. 2). The
10–60 m in thickness and consists of facies T and graben, ca. 40 km wide and 60 km long, hosted an
Fl(w), whereas the upper assemblage consists mainly ephemeral-lake environment in the central part, where
of facies T and is ca. 6 m in thickness. Both these the facies associations of the Değne Member were
units have a limited lateral extent, are underlain and deposited. Axial fluvial systems supplied sediment
overlain by fluvial deposit (Fig. 7, see SarVkavak and predominantly from the NE and SW and alluvial fans
KavalcVlar logs), and their thickness increases prograded from the basin-margin fault escarpments,
towards the basin-margin SarVkavak-KumafYarV Fault depositing the facies associations of the Derindere
and Uzunoluk-Çameli Fault (Fig. 8C,F). This facies Member (Fig. 8B). The thickness of these members
association is attributed to the precipitation of indicates active, sydepositional subsidence. The
calcium carbonate from spring waters rising along thickening of the alluvial-fan deposits towards the
the basin-margin faults (cf. Guo and Riding, 1998). basin margins, and their fault contact with the latter,
Tufa and travertine (hotspring) deposits are generally indicate sedimentation related to growth faulting (Fig.
recognized to be related to faults (Heimann and Sass, 9A). The overall upward fining and basin-margin
1989), especially to extensional tectonic settings backlapping by the alluvial succession suggest a
(Altunel and Hancock, 1993). synrift to post-rift sedimentation. Constrained by
palaeontological dates, the thickness of deposits
suggests that the synrift (10.8–8.5 Ma) and post-rift
5. Basin evolution and changes in sedimentary of subsidence (until 3.8–3.2 Ma) created a basinal
environments accommodation depth of up to 210 m.
The tectonic structure of the basin (Fig. 2) shows 5.2. The second rifting pulse (Early–Middle Pliocene
a longitudinal dissection by successive generations of transition)
normal faults, which divided the basin into half-
graben compartments and caused major changes in The second rifting pulse created a NW-dipping
the basin palaeogeograpy and sedimentary environ- normal fault along the graben axis, referred to as the
ment. These intrabasinal changes are recognizable SarVkavak–KumafYarV Fault, which split the basin
from the spatial relationships among the main facies longitudinally into two compartments (Fig. 8C, see
associations described in the previous section. The also Fig. 2). The earlier deposits had been tilted
present section summarizes the successive rifting towards the southeast, although their contact with the
pulses and their effects on the sedimentation pattern overlying deposits shows no recognizable angular
in the basin. unconformity on a local outcrop scale, which suggestsM.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 423 Fig. 9. Structural features related to the development of the Çameli Basin: (A) Growth fault associated with the first pulse of rifting; (B) the Bozdağ Fault, bounding the basin to the NW; (C) the Dirmil Fault, bounding the basin to the SE; (D) the fault-related tilting of basin-fill deposits; and (E) the open-lake deposits overlying metamorphic basement, uplifted by the SarVkavak-KumafYarV Fault (see also Fig. 2). a gradual, progressive tilting of the basin floor, in a lake facies association (Fig. 8D). In the KavalcVlar area growth-fault style. This new axial fault caused the (Fig. 2), the open-lake deposits directly overlie the formation of the first unit of tufa deposits, exposed in bedrock that was elevated as the fault footwall (Fig. the SarVkavak–Ericek area and thickening towards the 9E). A peat-mire facies association deposited at the fault (Figs. 2 and 7, see the SarVkavak log; and Fig. lateral transition between the fluvial KumafYarV Mem- 8C). These are up to 60 m thick, passing basinward ber and the lacustrine Değne Member near ÇamlVbel, into an ephemeral-lake facies association. The tufa 13 km north of Çameli (Figs. 2 and 5H), contains facies association is overlain by fluvial and peat-mire micro-remains of mammal Rodentia–Arvicolidae facies associations (Fig. 7). The latter deposits abound Mimomys sp., which indicate a Late Ruscinian–Early in mammal micro- and macro-fossils, including the Villanian (Piacenzian–Gelasian) age of the deposits teeth and bones of Mimomys occitanus, Apodemus (MN zones 15–16, 3.5–2.5 Ma). The age difference dominans, Orientalomys similis and Pseudomeriones between the lacustrine deposits at ÇamlVbel and Ericek tchaltaensis, which indicate a Late Ruscinian (Zan- localities (Figs. 2 and 7) is attributed to block faulting clean–Piacenzian) age of the deposits (MN zone 15, and/or relative lake-level rise, either of which could 3.8–3.2 Ma). cause progressive inundation of the uneven topogra- This phase of subsidence involved also movement phy of the fault-created intrabasinal ridge. along the primary basin-margin faults, perhaps more During this second stage of sedimentation, a rapid, which created a deeper-water lake, caused monotonous open-lake succession of marls and marly inundation of the intrabasinal ridge and basin-margin limestones Fl(w) intercalated with wave-worked fault escarpments, and led to deposition of an open lacustrine facies Sh(w) and Sr(w) was deposited.
424 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
The net subsidence at this stage was probably rapid nius, Apodemus dominans and Micromys praeminu-
and driven chiefly by the two primary bounding faults tus in the lower/distal part of the alluvial-fan facies
of the graben, rather than by the younger growth fault assemblage that prograded onto the tufa deposits. This
(Fig. 8D, lower diagram). The open-lake facies episode of rifting is estimated to have accounted for
association in the basin’s central part is up to 220 m little more than 10% of the basin’s tectonic extension,
thick and shows no evidence of shallowing, passing but caused the most conspicuous palaeogeographic
into Gilbert-type river deltas to the southwest and changes in the basin (Fig. 8E,F).
northeast, and into Gilbert-type fan deltas towards the The Çameli Basin was completely filled with
primary basin-margin escarpments. The steep foreset fluvial and alluvial-fan deposits by the end of the
deposits with slump structures, overlain by terrestrial Villanian (Late Gelasian) stage, when tectonic sub-
topset facies, indicate a relatively deep lake and sidence apparently ceased. The subsequent phase of
limited progradation of the deltas. The lacustrine Quaternary sedimentation involved chiefly the inci-
facies association locally overlies the primary Dirmil sion of streams and development of small alluvial
and Bozdağ faults, notably near the villages of SuçatV fans, which prograded from the dormant fault escarp-
and ÇavdVr (Fig. 2), which indicates a post-rift ments and largely smoothed out the pre-existing
expansion of the lake or a shift in the locus of topographic relief of the basin.
faulting. The relatively thick open-lake facies succes-
sion eventually passes upwards into ephemeral-lake
deposits, displaying evidence of a basinward expan- 6. Discussion
sion of the advancing axial river deltas and basin-
margin fan deltas (Fig. 8E). The present study of the Çameli Basin, involving
detailed facies analysis of the basin-fill succession
5.3. The third rifting pulse (latest Pliocene) and biostratigraphic dating of its micro- and macro-
mammalian and molluscan fossils, provides a tenta-
After the shallowing of the lake (Fig. 8E), a new tive time-stratigraphic framework for the neotectonic
rifting pulse occurred, forming the AlcV–Kelekçi and rifting pulses that caused the development of Neo-
Uzunoluk–Çameli faults that split further the basin gene grabens in southwestern Anatolia. The tectonic
into still narrower half-graben compartments (Fig. cause of the formation of these basins is somewhat
8F). These faults are best recognizable where cutting controversial, as it has been variously attributed to
the open-lake sedimentary succession deposited the the westward tectonic escape of the Anatolian craton
previous, second rifting phase. The strike and north- (e.g., Dewey and Xengör, 1979), back-arc crustal
western dip of these new faults are similar to those of stretching (e.g., Le Pichon and Angelier, 1979) or
the pre-existing SarVkavak–KumafYarV Fault. This orogenic collapse (e.g., Dewey, 1988; Seyitoğlu and
faulting episode caused further southeastward tilting Scott, 1991; Gessner et al., 2001). On the other
of the segmented basin-fill succession (Fig. 9D) and hand, these various regional models are not neces-
was accompanied by the deposition of the second unit sarily mutually exclusive and may well be regarded
of tufa facies, associated with the Uzunoluk–Çameli as end-members of an actual spectrum of causal
Fault (Fig. 8F, see south of Çameli) and ca. 6 m thick factors. For example, Collins and Robertson (1998,
in the outcrops near KavalcVlar (Fig. 7, see the 1999) have documented that the emplacement of the
KavalcVlar log). The thickness of that second tufa Lycian allochthon took place episodically in the Late
unit increase towards the Uzunoluk–Çameli Fault, Cretaceous, Middle Eocene–Oligocene and Middle–
which suggests syndepositional faulting (Fig. 8F). The Late Miocene times and coeval with the formation of
tufa facies assemblage is overlain by fluvial and a foreland zone caused by extension in the hinter-
alluvial-fan facies associations (Fig. 6A and the land. Some of the basins were formed in arcuate
SarVkavak and KavalcVlar logs in Fig. 7). zone of crustal extension behind the Cyprus sub-
This third pulse of rifting occurred in the Late duction arc (Aksu, Köprü, Manavgat, Cilicia, Adana
Villanian (Gelasian) time (2.6–1.8 Ma), as is indicated and Mut-Ermenek basins), but the western troughs
by the remains of micro-mammals Mimomys pliocae- were affected also by the Lycian foreland compres-M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 425
sion, whereas those to the east were markedly Several of the Neogene basins in western Anatolia
influenced by the regional stress field generated by have recently been studied in considerable detail,
tectonic escape, which eventually led to their notably the Burdur (Price and Scott, 1989), AlaYehir/
tectonic inversion (Dreyer et al., 2005). Some other Gediz (Seyitoğlu et al., 2002; Xen and Seyitoğlu,
basins, such as the Burdur graben (KazancV, 1988, 2002), Büyük Menderes (Sözbilir and Emre, 1990;
1990), continued to subside in Quaternary time, thus Ünay et al., 1995; Seyitoğlu and Scott, 1992; Xen and
pointing to an important and varied role for the Seyitoğlu, 2002), Denizli (Sözbilir, 1997) and Tavas
tectonic escape process. In short, the southwestern Basins (Hakyemez, 1989). These studies are in an
part of Anatolia can be regarded as an interference important contribution to an overall understanding of
zone of at least three difference tectonic regimes, the regional pattern of tectonic extension and the
with differing stress fields and also somewhat formation of the Neogene basins (Fig. 10). However,
different time spans. The development of a particular relatively few studies have arrived in a coherent
basin thus probably involved the effects of more tectono-stratigraphic framework for the local rifting
than one regime, although one stress field might process. For example, the evolution of the Çameli
dominate in some cases or could prevail during Basin might be compared with that of the Tavas
certain periods. Basin, which involved Early Miocene sedimentation,
Fig. 10. Stratigraphic comparison of the Çameli Basin with other extensional basins in the western Anatolia.426 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
but the stratigraphic development of the latter basin shifted with time (e.g., Ziegler, 1988, 1994; Lundin
has not been sufficiently constrained by a structural and Doré, 1997; Skogseid et al., 2000; Reemst and
framework to provide the local history of rifting. Cloetingh, 2000). Several different models have been
The Çameli Basin was fully terrestrial and hosted suggested for the migration of rifting locus, all based
lacustrine, fluvial and alluvial-fan depositional sys- on the principle that the rifting occurs in places where
tems, including river deltas and fan deltas of both the lithosphere is weakest. One mechanism for a
shoal-water and Gilbert type. The sedimentation in the limiting extension at a particular location postulates
graben commenced with a rifting pulse in the that the cooling of the continental lithosphere during
Vallesian (Early Tortonian) time (10.8–8.5 Ma). The slow stretching may allow its local strength to
second pulse of rifting occurred in the Late Ruscinian increase, so that the locus of faulting will shift to a
(Zanclean–Piacenzian) time (3.8–3.2 Ma) and the zone of lower strength (England, 1983; Houseman
third pulse in the Late Villanian (Gelasian) time (2.6– and England, 1986; Sonder and England, 1989). In
1.8 Ma). These consecutive rifting pulses generated this model, no other factors, such as a possible change
new faults parallel to the graben axis, splitting the NE- in intra-plate or plate-boundary forces, are required to
trending basin into progressively narrower half-graben explain the shift.
compartments. Each rifting pulse reactivated the Kusznir and Park (1987, their Fig. 18) have
graben’s primary pair of bounding faults. suggested hypothetically that an inward shift of the
The third rifting pulse in the present case is locus of faulting in a rift system may be a direct result
estimated to have accounted for little more than of a high rate of lithospheric stretching, as opposed to
10% of the crustal extension, but caused the most a low rate, when the successive faults would tend to
dramatic change in the basin’s internal palaeogeog- form farther away from the rift-system axis. However,
raphy and sedimentation pattern. This evidence this notion is not supported by more recent modelling
implies that the magnitude of the intrabasinal changes studies (e.g., Sonder and England, 1989; Van Wijk
in palaeogeography and sedimentary environment in and Cloetingh, 2002), which show that an inward shift
an active graben may not necessarily reflect the actual may occur under a low-rate stretching, even if this
magnitude of lithospheric extension. does not culminate in a crustal break-up.
The post-rift phases of subsidence were probably Manatschal and Bernoulli (1999) have postulated
driven by subcrustal cooling and broader thermal that the cooling and strengthening of lithosphere
sagging, but tended also to involve the primary during stretching, if the latter is sufficiently slow,
bmasterQ faults of the graben, as is indicated by may force a shift in the rifting locus to previously non-
sediment thicknesses and facies relationship at the thinned or base-thinned zones. The process of syn-rift
margins. A similar phenomenon of post-rift fault cooling has also been proposed to explain the
activity has been recognized in many well-studied formation of some of the narrow-wide pairs of
extensional regions, such as the Mesozoic rift system conjugate plate margins of the South Atlantic (Davi-
of North-West Europe (Gabrielsen, 1986, 1990, 2001; son, 1997) and the migration of basin on the South
Roberts et al., 1990, 1993; Ziegler, 1990; Faerseth et Alpine rifted palaeomargin (Bertotti et al., 1997). The
al., 1995; Kyrkjebo et al., 2001). This activity modelling by Van Wijk and Cloetingh (2002) shows
typically involves some of the main and steepest that the deformation during a low-rate extension
faults, and has been attributed to the loading effect of localizes outside the first-formed graben, which in
the thick sediments accumulated during the rifting turn is uplifted and becomes a bcold spotQ zone in the
phase and/or a diachronous decline of the rifting pulse area. When syn-rift cooling, the lithosphere regains
along the graben, possibly in combination with a non- strength during stretching, instead of becoming
uniform thermal subsidence related to lithospheric weaker, and the lithospheric necking zone thus
structure. becomes stronger than the adjacent zones. In the
A striking aspect of the rifting process in the case-study models pertaining to the Mesozoic rifting
Çameli Basin was the progressively inward develop- in the Norwegian–Greenland Sea region (Van Wijk
ment of successive faults. There are numerous and Cloetingh, 2002), the transition velocity was
examples of basins in which the locus of extension about 8 mm/year and the locus of maximum thinningM.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431 427
migrated after some 50–60 Ma; these are time 5 Ma, the second of ca. 1 Ma and the third of 1.8 Ma,
intervals incomparably longer than in the present including the Quaternary period of sedimentation. It is
case, although the time frequency of shifting might uncertain as to which of the models might be most
possibly be higher if the stretching was faster and the appropriate in the present case, on this particular time
lithospheric properties were different. scale, because the lithosphere properties in the region
In the model postulated by Steckler and Ten Brink are little known and no modelling has thus far been
(1986), and derived from the Red Sea rift, it is the attempted. Firstly, it is uncertain whether the rifting
preexisting spatial variation in lithospheric strength pulses were merely short-term fluctuations of the
that controls the location and shifts of the rifting stretching process, or were separated by period of true
locus. The controlling factors include the thickness tectonic quiescence (cf. Cloetingh, 1988). Secondly, it
and composition of crust, the accumulated sediment is uncertain whether sufficient cooling could have
thickness and the geotherm distribution. occurred on the time scale considered (cf. Moretti and
In a model derived by Sawyer and Harry (1991) Friodevaux, 1986; Voorhoeve and Houseman, 1988).
from the Baltimore Canyon Trough on the U.S.A.
Atlantic margin, the axis of maximum thinning
migrates with time, and the switching of the rifting 7. Conclusions
locus is controlled by pre-existing weaknesses in both
the crust and the upper mantle, offset relative to each The Çameli Basin is a NE-trending Neogene
other. The heterogeneity creates an asymmetrical graben in southwestern Anatolia. Its deposits are
initial configuration of the lithosphere, with a zone completely terrestrial and represent river, alluvial-
of reduced and laterally changing strength. The fan, fan-deltaic and lacustrine settings, Late Mione to
modelling shows (Sawyer and Harry, 1991) that the Late Pliocene in age and ca. 500 m thick. The
lithosphere, when subject to extension, reacts with inception of this basin occurred later than the genesis
rifting that is first manifested mainly in the zone of of the marine basins of southern Anatolia (e.g., the
crustal weakness and, after some time, mainly in the Aksu, Köprü, Manavgat, Adana, Mut and Ermenek
zone of mantle weakness. Either of these two models basins) but was coeval with or slightly earlier than
might potentially be relevant in the present case. the grabens of western Turkey (Fig. 10). The activity
It has been often suggested also that the migration and/or filling of the Çameli basin ended in latest
of the locus of extension can be a direct consequence Pliocene while the Aegean grabens remained active
of multiple stretching phases, with intervening post- into the Quaternary. This shows that southwestern
rift periods during which the lithosphere is not under Anatolia played a critical role in the geological
tensile stress. In this model, which might apply in evolution of the eastern Mediterranean region by
the present case, the weakened lithosphere resulting acting as an interference zone between tectonic
from one stretching pulse requires sufficient time to regimes with differing stress fields and also some-
cool and regain strength before the next stretching what different time spans.
pulse occurs (Bertotti et al., 1997); this implies The most conspicuous feature of the basin is its
relatively long periods of tectonic quiescence be- longitudinal dissection by three successive genera-
tween consecutive rifting pulses. The thickness tions of normal faults, representing separate tectonic
proportion of the stronger mantle to the weaker pulses. These faults divided the basin into half-
crust in a thinned lithosphere has to be large, in grabens, leading to major changes in basin palae-
comparison to a non-thinned lithosphere (Bertotti ogeography and sedimentary environments (Figs. 2, 7
et al., 1997), and the model implies also sudden, and 8). The first tectonic pulse (Late Miocene)
time-dependent changes in the magnitude of the initiated the Çameli Basin while the second (Early–
intra-plate stress field as the actual cause of the Middle Pliocene) and the third pulse (latest Pliocene)
rifting and post-rifting phases. led to expansion of the basin by up to 10%. The rifting
In the present case, the three rifting phases are process in the Çameli Basin shows progressively
estimated to have had a duration on the order of 1–2 inwards development of successive faults but the
Ma, with the first post-rift relaxation phase of nearly causes of this development remain elusive.428 M.C. Alçiçek et al. / Sedimentary Geology 173 (2005) 409–431
Acknowledgements implications for the Late Cenozoic evolution of the Aegean.
Geophys. J. Int. 126, 11 – 53.
Becker-Platen, J.D., 1970. Lithostratigraphische Unterschungen
The study was supported by the Scientific and im Kanozoikum Südwest Anatoliens (Türkei)-(Kanozoikum
Technical Research Council of Turkey (TÜBİTAK und Braunkohlen der Turkei)-Beihefte zum. Geol. Jahrb. 97.
research grant YDABÇAG 100Y004). This paper 244 pp., Hannover.
arises of an unpublished PhD thesis by M. Cihat Belt, E.S., Flores, R.M., Warwick, P.D., Conway, K.M., Johnson,
K.R., Waskowitez, R.S., 1984. Relationship of fluviodeltaic
Alçiçek, completed at the Ankara University under the
facies to coal deposition in the lower fort union formation
supervision of Nizamettin KazancV. The manuscript (Paleocene), south-western North Dakota. In: Rahmani, R.A.,
was read and improved by Wojtek Nemec (University Flores, R.M. (Eds.), Sedimentology of Coal and Coal-Bearing
of Bergen), and benefited also from constructive Sequences, Spec. Pub. International Association of Sedimen-
reviews by Alastair H.F. Robertson (University of tologists, vol. 7, Blackwell, Oxford, pp. 177 – 195.
Edinburgh) and Alan S. Collins (Curtin University, Bertotti, G., Ter Voorde, M., Cloetingh, S.A.P.L., Picotti, V., 1997.
Thermochemical evolution of the south Apline rifted margin
Australia) and from helpful suggestions by Gilbert (north Italy): constraints on the strength of passive continental
Kelling (Keele University). Xevket Xen (CNRS) and margins. Earth Planet. Sci. Lett. 146, 181 – 193.
Gerçek Saraç (MTA) kindly helped in the biostrati- Bozkurt, E., 2000. Timing of extension on the Büyük
graphic determination and classification of mamma- Menderes Graben, western Turkey, and its tectonic
lian remains. Sevinç K. YeYilyurt (Onsekizmart implications. In: Bozkurt, E., Winchester, J.A., Piper, J.D.A.
(Eds.), Tectonics and Magmatism in Turkey and the Surround-
University) determined the molluscan fossils. Hüseyin ing Area, Spec. Pub. Geological Society vol. 173. Blackwell,
Erten and Hülya Alçiçek (Pamukkale University) London, pp. 385 – 403.
offered field assistance and helped to collect and Bozkurt, E., 2001. Neotectonic of Turkey—a synthesis. Geodin.
prepare the fossils for analysis. We are grateful also to Acta 14, 3 – 30.
Gürol Seyitoğlu and Ergun Gökten (Ankara Univer- Cloetingh, S.A.P.L., 1988. Intraplate stresses: a new element
in basin analysis. In: Kleinspehn, K.L., Paola, C. (Eds.),
sity), İbrahim Çemen (Oklahoma State University), New Perspectives in Basin Analysis. Springer-Verlag, New
Yavuz Hakyemez and NeYat Konak (MTA) and Fuat York, pp. 205 – 230.
Xaroğlu (TPAO) for their helpful comments and Colella, A., Prior, D.B. (Eds.), (1990). Coarse-Grained Deltas,
discussions. Spec. Pub. International Association of Sedimentologists vol.
10, p. 357.
Collins, A., Robertson, A.H.F., 1998. Processes of Late Cretaceous
to Late Miocene episodic thrust-sheet translations in the Lycian
References Taurides, SW Turkey. J. Geol. Soc. (Lond.) 155, 759 – 772.
Collins, A., Robertson, A.H.F., 1999. Evolution of the Lycian
Aksu, A.E., Piper, D.J.W., Konuk, T., 1987. Quaternary growth allochthon, western Turkey, as a north-facing Late Palaeozoic
patterns of Büyük Menderes and Küçük Menderes deltas, to Mesozoic rift and passive continental margin. Geol. J. 34,
western Turkey. Sediment. Geol. 52, 227 – 250. 107 – 138.
Alçiçek, M.C., 2001. Sedimentological investigation of Çameli Collinson, J.D., 1996. Alluvial sediments. In: Reading, H.G. (Ed.),
Basin (Late Miocene–Late Pliocene, Denizli, SW Anatolia). Sedimentary Environments: Processes, Facies and Stratigraphy.
PhD thesis, Ankara University, Ankara, Turkey. Blackwell Science, Oxford, pp. 37 – 82.
Allen, J.R.L., 1983. Studies in fluviatile sedimentation: bars, bar- Cronin, B.T., Gürbüz, K., Hurst, A., Satur, N., 2000. Vertical and
complexes and sandstone sheets (low-sinuosity braided streams) lateral organization of a carbonate deep-water slope marginal to
in the Brownstones (lower Ddevonian), Welsh Borders. Sedi- a submarine fan system, Miocene, southern Turkey. Sedimen-
ment. Geol. 33, 237 – 293. tology 47, 801 – 824.
AltVnlV, İ.E., 1955. The geology of southern Denizli. Rev. Fac. Sci. Davison, I., 1997. Wide and narrow margins of the Brazilian South
Univ. Istanb., Ser. B: Sci. Nat. 20 (1–2), 1 – 47. Atlantic. J. Geol. Soc. (Lond.) 154, 471 – 476.
Altunel, E., Hancock, P.L., 1993. Morphology and structural Dewey, J.F., 1988. Extensional collapse of orogens. Tectonics 7,
setting of Quaternary travertines at Pamukkale, Turkey. Geol. 1123 – 1139.
J. 28, 335 – 346. Dewey, J.F., Xengör, A.M.C., 1979. Aegean and surrounding
Anadón, P., Cabrera, L.I., Julia, R., Roca, E., Rosell, L., 1989. regions: complex multiplate and continuum tectonics in a
Lacustrine oil-shale basins in Tertiary grabens from NE Spain convergent zone. Geol. Soc. Amer. Bull. 90, 84 – 92.
(western European rift systems). Palaeogeogr. Palaeoclimatol. Dewey, J.F., Hempton, M.R., Kidd, W.S.F., Xaroğlu, F.,
Palaeoecol. 70, 7 – 28. Xengör, A.M.C., 1986. Shortening of continental litho-
Armijo, R., Meyer, B., King, G.C.P., Rigo, A., Papanastassiou, R., sphere: the neotectonics of eastern Anatolia—a young
1996. Quaternary evolution of the Corinth Rift and its collison zone. In: Coward, M.P., Ries, A.C. (Eds.), CollisionYou can also read
