Neotethys closure history of Anatolia

22 Pages·2013·1.42 MB·English

J.metamorphicGeol.,2013,31,585–606 doi:10.1111/jmg.12034
40 –39
Neotethys closure history of Anatolia: insights from Ar Ar
P–T
geochronology and estimation in high-pressure
metasedimentary rocks
A. POURTEAU,1 M. SUDO,1 O. CANDAN,2 P. LANARI,3 O. VIDAL3 AND R. OBERHA€NSLI1
1Institutfu€rErd- undUmweltwissenschaften, Universit€at Potsdam,Karl–Liebknecht–Straße24–25, Potsdam–Golm,14476,
Germany([email protected])
2DokuzEylu€lU€niversitesi, Mu€hendislikFaku€ltesi, JeolojiMu€hendislig(cid:2)iBo€lu€mu€,TınaztepeKampusu,Buca–Izmir, 35160,
Turkey
3Isterre,Universite(cid:3) JosephFourier Grenoble,CNRS,1381rue delaPiscine, GrenobleCedex 09, 38041,France
ABSTRACT The multiple high-pressure (HP), low-temperature (LT) metamorphic units of Western and Central
Anatolia offer a great opportunity to investigate the subduction- and continental accretion-related
evolution of the eastern limb of the long-lived Aegean subduction system. Recent reports of the HP–
LT index mineral Fe-Mg-carpholite in three metasedimentary units of the Gondwana-derived Anato-
lide–Tauride continental block (namely the Afyon Zone, the O€ren Unit and the southern Menderes
Massif) suggest a more complicated scenario than the single-continental accretion model generally
put forward in previous studies. This study presents the first isotopic dates (white mica 40Ar–39Ar
geochronology), and where possible are combined with P–T estimates (chlorite thermometry, phengite
barometry, multi-equilibrium thermobarometry), on carpholite-bearing rocks from these three HP–
LT metasedimentary units. It is shown that, in the Afyon Zone, carpholite-bearing assemblages were
retrogressed through greenschist-facies conditions at c. 67–62 Ma. Early retrograde stages in the O€ren
Unit are dated to 63–59 Ma. In the Kurudere–Nebiler Unit (HP Mesozoic cover of the southern
Menderes Massif), HP retrograde stages are dated to c. 45 Ma, and post-collisional cooling to c.
€
26 Ma. These new results support that the Oren Unit represents the westernmost continuation of the
Afyon Zone, whereas the Kurudere–Nebiler Unit correlates with the Cycladic Blueschist Unit of the
Aegean Domain. In Western Anatolia, three successive HP–LT metamorphic belts thus formed: the
northernmost Tav(cid:1)sanlı Zone (c. 88–82 Ma), the O€ren–Afyon Zone (between 70 and 65 Ma), and the
Kurudere–Nebiler Unit (c. 52–45 Ma). The southward younging trend of the HP–LT metamorphism
from the upper and internal to the deeper and more external structural units, as in the Aegean
Domain, points to the persistence of subduction in Western Anatolia between 93–90 and c. 35 Ma.
After the accretion of the Menderes–Tauride terrane, in Eocene times, subduction stopped, leading to
continental collision and associated Barrovian-type metamorphism. Because, by contrast, the Aegean
subduction did remain active due to slab roll-back and trench migration, the eastern limb (below
Southwestern Anatolia) of the Hellenic slab was dramatically curved and consequently teared. It
therefore is suggested that the possibility for subduction to continue after the accretion of buoyant
(e.g. continental) terranes probably depends much on palaeogeography.
Keywords: 40Ar–39Ar geochronology; Anatolia; chlorite–phengite thermobarometry; high-pressure
metasedimentary rocks.
Hinsbergen et al., 2005). This phenomenon, recog-
INTRODUCTION
nized in several other subduction systems worldwide,
The Mediterranean realm represents one of the great- like the Scotia–Sandwich, the Lesser Antilles and the
est natural laboratories for investigating subduction- Banda subduction zones, results in concave trenches
and continental accretion-related geodynamics. The and ‘spoon-shaped’ down-going slabs (Hsui &
integration of multi-disciplinary data has demon- Youngquist, 1985; Spakman & Hall, 2010), which in
strated that, due to subduction retreat, subduction of general are delimited laterally by sub-vertical slab
a single lithospheric slab could persist despite the tears, expressed, at the surface, by strike-slip faulting
accretion of buoyant micro-continental terranes to and specific volcanic activity (Govers & Wortel,
the trench (e.g. Malinverno & Ryan, 1986; Royden & 2005). Slab roll-back, trench migration and -curving,
Burchfiel, 1989; Jolivet & Faccenna, 2000; van and slab tearing, also reproduced by analogue and
©2013JohnWiley&SonsLtd 585
586 A. POURTEAU ET AL.
numerical modelling experiments (e.g. Husson et al., opened perspectives towards a better understanding
2009; Schellart, 2010), are processes commonly envis- of the subduction- and collision accretion-related tec-
aged for the evolution of retreating subduction sys- tonics in this region. To place their burial and exhu-
tems. However, what controls the location and mation into an accurate timeframe, the timing of
timing of lateral disruption is not fully resolved, as in subduction-related metamorphism in these units still
the Aegean Domain in particular. needs to be determined. In consequence, this study
In the Aegean subduction system, the retreat of the combines 40Ar–39Ar geochronology and, when possi-
Hellenic NW- to NE-plunging slab is notably illus- ble, thermobarometric calculations in white mica-,
trated by the younging, from the inner to the outer carpholite-bearing rocks and uses these results to dis-
parts of the upper plate, of extensional tectonics, arc cuss the significance of the isotopic dates. Based on
magmatic activity and high-pressure (HP), low-tem- these new data, a new evolutionary model for Wes-
perature (LT) metamorphic events (e.g. Ring et al., tern and Central Anatolia from c. 85 Ma to the pres-
2010; Jolivet et al., 2013). According to a generally ent situation is proposed.
accepted model, in response to the burial of continen-
tal terranes, subduction temporarily stalled, then the
GEOLOGICAL SETTINGS
slab rolled back, and the accreted crustal terranes
were decoupled from the underlying, down-going
Tectonicdomains of Western Anatolia
lithosphere leading to trench migration towards the
foreland (Jolivet & Faccenna, 2000; van Hinsbergen Western to Central Anatolia is composed of three
et al., 2005; Jolivet & Brun, 2008). As displayed on tectonic regions that experienced contrasting tectonic
tomographic images, the Hellenic slab is interrupted histories between the Late Cretaceous and Palaeogene
to the east by a slab tear located below Southwestern (Ketin, 1966). Northern Anatolia exposes a polystage
Anatolia (Bijwaard et al., 1998; Biryol et al., 2011; fold-thrust belt known as the Pontides, which are
Salau€n et al., 2012). According to its possible volca- characterized by pre-Jurassic metamorphic rocks with
nic expression in Western Anatolia, this tear might unconformable non-metamorphosed Jurassic to Cre-
have formed at c. 20 Ma (Dilek & Altunkaynak, taceous sedimentary and volcanic rocks (Altiner
2009). However, the understanding of the evolution et al., 1991; Yılmaz et al., 1995; Okay & Sahinturk,
of subduction at the eastern periphery of the Aegean 1997). Central Anatolia is composed of the Central
Domain lacks accurate temporal constraints on conti- Anatolian Crystalline Complex made of Late Creta-
nental-accretion events. Therefore, to gain insights ceous intrusions and Barrovian metamorphic rocks
into the disruption of the Hellenic slab, the evolution (e.g. Seymen, 1981; Go€ncu€o(cid:2)glu et al., 1997; Aydın
of the multiple HP–LT metamorphic units exposed et al., 1998; Whitney et al., 2001) with inferred early
in Western and Central Anatolia was investigated. Palaeozoic to Mesozoic protolith (Kocak & Leake,
In Western and Central Anatolia, subduction- and 1994; Go€ncu€o(cid:2)glu et al., 1997). In the Late Cretaceous
collision-related metamorphic events associated with and early Cenozoic, arc magmatism above the north-
the accretion of the Gondwana-derived Anatolide– dipping Neotethys subduction zone took place in the
Tauride Block to the southern composite margin of Pontides (S(cid:1)engo€r & Yılmaz, 1981; Okay & Satır,
Eurasia occurred in an interval greater than 40 Ma— 2006) and the Central Anatolian Crystalline Complex
88–82 Ma for the HP–LT metamorphism v. likely (Kadio(cid:2)glu et al., 2003; Ilbeyli et al., 2004).
45–30 Ma for the Barrovian metamorphism (see Western and Southern Anatolia consists of meta-
details below). For comparison, in other collisional morphosed and non-metamorphosed Precambrian to
belts, subduction-related metamorphism is followed Eocene rocks, known as the Anatolide–Tauride Block
by a collision-related overprint within 15–20 Ma or (Okay & Tu€ysu€z, 1999). This domain is characterized
less (e.g. Western Himalaya, de Sigoyer et al., 2000; by a Precambrian crystalline basement of Gondwana
Armenia, Rolland et al., 2009; Central Alps, Wie- affinity (Pan-African orogeny, late Ordovician glacia-
derkehr et al., 2009; Southeastern Turkey, Oberh€ansli tion; Monod et al., 2003), and Late Cretaceous to
et al., 2012). Although this peculiar timing was docu- early Cenozoic HP–LT metamorphism that affected
mented for more than a decade (Sherlock et al., its northern passive margin (e.g. S(cid:1)engo€r & Yılmaz,
1999), no tectonic reconstruction for Anatolia 1981; Pourteau et al., 2010).
accounts for this peculiar feature yet.
In the past few years, Fe-Mg-carpholite-bearing
assemblages, as evidence for HP–LT metamorphism, Anatolide–Taurideunits
were documented in several tectonic units of Western The term ‘Anatolides’ is used to designate the units
and Central Anatolia (Oberh€ansli et al., 2001; Rim- affected by Alpine (i.e. Late Cretaceous to early
mel(cid:3)e et al., 2003a,b, 2006; Candan et al., 2005; Pour- Cenozoic) regional metamorphism, and the term
teau et al., 2010) that were until then supposedly ‘Taurides’ designates the non-metamorphosed thrust
characterized by greenschist-facies metamorphism and folded external platform units (Fig. 1).
(e.g. Okay et al., 1996). These discoveries, although The Taurides, which consist of Precambrian to
shattering previous geodynamic reconstructions, have Eocene sedimentary rocks and Neotethyan ophiolites
©2013JohnWiley&SonsLtd
NEOTETHYS CLOSURE HISTORY OF ANATOLIA 587
€
Figure 1. TectonicunitsofWesternandCentralAnatolia(notetheOrenUnitklippenatoptheCycladicBlueschistUnitand
northernMenderesNappes)anddistributionofthecarpholite-bearingsamplesusedfor40Ar–39Argeochronologyinthepresent
study.OthercarpholitelocalitiesdiscussedintextareBah(cid:1)ceyakaintheKurudere–NebilerUnit(Rimmel(cid:3)eet al.,2003b),and
KonyainthecentralAfyonZone(Pourteauet al.,2010).Unitabbreviationsare:AlNa,Alada(cid:2)gNappes;BoZn,BornovaZone;
CyBs,CycladicBlueschistUnit;HBHNa,Hoyran–Bey(cid:1)sehir–HadimNappes;LyNa,LycianNappes;O€rUn,O€renUnit.Toponymic
abbreviationsare:Af,Afyonkarahisar;An,Ankara;De,Denizli;Iz,Izmir;Ka,Kayseri;Ko,Konya;Ku€,Ku€tahya;Me,Mersin;
Or,Orhaneli;Si,Sivrihisar;Ya,Yahyalı.
with 93–90 Ma–old metamorphic soles (C(cid:1)elik et al., volcanic formations (e.g. S(cid:1)engo€r & Yılmaz, 1981;
2006), were affected by multistage deformation Okay et al., 1996). The general Mesozoic lithostratig-
between latest Cretaceous and late Miocene times, raphy consists of a continuous succession of Triassic
leading to the formation of regional-scale nappe sys- clastic sedimentary rocks (with volcanic intercala-
tems (e.g. Lycian Nappes; Gutnic et al., 1979; Collins tions) followed by Upper Triassic to Upper Creta-
& Robertson, 1998). The NW edge of the Menderes ceous neritic carbonates followed by pelagic
Nappes is thrust below the Bornova Zone, represent- limestones, cherts, flysches, and olistostromes (e.g.
ing a non-metamorphosed, tectonized Maastrichtian– Bozkurt & Oberh€ansli, 2001). This sequence is inter-
Danian ‘megaolistostrome’, which might have formed preted as the continuous record of the development
along a sinistral transform zone at the edge of the of a continental passive margin from rifting in the
Anatolide–Tauride Block (Okay et al., 1996). Early Triassic (e.g. Akal et al., 2012) to oceanization
Here the main characteristics and nomenclature of in the middle Triassic, and later carbonate platform
the western Anatolides, where it has been best drowning in the Late Cretaceous, as it entered the
described, are summarized and this tectono-stratigra- trench of the north-dipping Neotethyan subduction
phy is expanded eastwards. For further details, the zone (Okay et al., 2001). Between the Late Creta-
reader is referred to Jolivet et al. (2013). The western ceous and the early Cenozoic, parts of the continental
Anatolides are composed of a metamorphosed Meso- margin were metamorphosed in a subduction zone,
zoic sedimentary sequence, and its substratum repre- as evidenced by the occurrence of widespread HP–
sented by Precambrian high-grade metamorphic and LT minerals in Lower Triassic sedimentary and vol-
intrusive rocks and Palaeozoic sedimentary and canic rocks (Okay et al., 1996; Pourteau et al., 2010).
©2013JohnWiley&SonsLtd
588 A. POURTEAU ET AL.
Owing to contrasting lithostratigraphies and meta- Table 1. SummaryofthepublishedisotopicdatesforHP–LT
morphic grades, the Anatolides are subdivided into metamorphismintheAnatolides.
several tectonic units (Fig. 1).
Dating Isotopic
The Tav(cid:1)sanlı Zone is a Late Cretaceous HP–LT Unit/locality Rocktype method date Error Reference
metamorphic belt (Okay, 1984), derived from the
most distal parts of the north-facing Anatolide-Tau- Tav(cid:1)sanlıZone*
Orhaneli Lws-bearing PhRb–Sr 79.7 1.6 Sherlocketal.
ride platform (Candan et al., 2005). In its western metabasite (1999)
part, the Tav(cid:1)sanlı Zone is characterized by well-pre- Orhaneli Metachert PhRb–Sr 78.5 1.6 Sherlocketal.
(1999)
served lawsonite-glaucophane-bearing metasedimenta-
Sivrihisar Lws-bearing PhRb–Sr 82.8 1.7 Sherlocketal.
ry and metavolcanic rocks (including local lawsonite metabasite (1999)
eclogite), for which peak P–T conditions were esti- Sivrihisar Metachert PhRb–Sr 80.1 1.6 Sherlocketal.
mated to be 2.0–2.6 GPa and 430–550 °C (Okay & Sivrihisar Cld-bearing PhAr–Ar 87.9 0.6 Sea(1to9n99e)tal.
Kelley, 1994; Okay, 2002; Davis & Whitney, 2006, schist (2009)
2008; C(cid:1)etinkaplan et al., 2008). Sporadic occurrence AfyonZone
of blueschist blocks in m(cid:3)elange localities near Konya, KKuu€€ttaahhyyaa CCaarr––qqzzvveeiinn PPhhAArr––AArr 6625..89 12..58 tthhiissssttuuddyy
in the Bolkar Mountains and near Kayseri (van der Ku€tahya Car–qzvein PhAr–Ar 61.5 8.0 thisstudy
Afyon Car-bearingphyllite PhAr–Ar 74.6 1.1 thisstudy
Kaaden, 1966; Droop et al., 2005) outlines the east-
Afyon Car-bearingphyllite PhAr–Ar 83.4 0.7 thisstudy
ward continuation of the Tav(cid:1)sanlı Zone into Central Yahyalı Car–qzvein PhAr–Ar 65.7 0.2 thisstudy
Anatolia (Okay, 1984; Pourteau et al., 2010). Lawso- Yahyalı Car–qzvein PhAr–Ar 64.8 0.4 thisstudy
nite blueschists from near Konya experienced 0.9– O€rYenahUynailtı Car–qzvein PhAr–Ar 66.7 0.4 thisstudy
1.1 GPa and 375–450 °C (Droop et al., 2005). The Milas Red-greenphyllite PhAr–Ar 70-90 Ring&Layer
good preservation of lawsonite throughout the unit O€ren Car–qzvein PhAr–Ar 62.6 0.4 thi(s2s0t0u3d)y
indicates cooling during decompression. Isotopic O€ren Car–qzvein PhAr–Ar 59.4 0.7 thisstudy
dates obtained on blueschists from the Tav(cid:1)sanlı Zone O€ren Car–qzvein PhAr–Ar 60.3 0.3 thisstudy
Kurudere–NebilerUnit
bracket peak- to early retrograde stages between 88
Kurudere Car–ky-bearingqz PhAr–Ar 45.9 2.0 thisstudy
and 78 Ma (40Ar–39Ar and 87Rb–86Sr on phengite; vein
Table 1 and references therein), and subsequent exhu- Nebiler Ky-bearing PhAr–Ar 26.5 0.8 thisstudy
microconglomerate
mation during the latest Cretaceous and the Palaeo-
CycladicBlueschistUnit(Anatolianpart)
cene (c. 60 Ma; 40Ar–39Ar on fine-grained phengite; Dilek Ep-blueschist PhAr–Ar 40.1 0.4 Oberh€ansli
Sherlock et al., 1999). East of Sivrihisar (Fig. 1), etal.(1998)
Ringetal.
blueschist facies minerals were replaced by greens- (2007)
chist- to amphibolite facies Barrovian-type assem-
blages (e.g. chloritoid (cid:1) staurolite (cid:1) sillimanite; *OtherphengiteAr–ArdatesbyOkayetal.(1998),Sherlocketal.(1999)andSherlock
&Kelley(2002)arenotincludedduetocommonexcessargon.
Whitney, 2002; Whitney et al., 2011) supposedly
formed c. 60 Ma (40Ar–39Ar on muscovite; Seaton
et al., 2009). Near Orhaneli and Sivrihisar, the stratigraphic relations leave a possible interval
Tav(cid:1)sanlı Zone is overlain by an ophiolitic complex, between the Campanian (c. 83–70 Ma) olistostrom at
composed of peridotites and an accretionary complex the uppermost stratigraphic levels of the Afyon Zone
€ €
(e.g. Okay, 1984). The tectonic contact between the (Ozer & Tansel Ongen, 2012), and the late Palaeo-
HP–LT rocks and the ophiolites is sealed by early cene (c. 58–55 Ma) non-metamorphic sedimentary
Eocene calc-alkaline intrusions derived from mantle cover (Dirik et al., 1999; Candan et al., 2005).
wedge melts (Harris et al., 1994; Delaloye & Bingol, The Menderes Massif, which occurs as a regional
2000). tectonic window below the other metamorphic units,
The Afyon Zone, which occurs south of and struc- exhibits a complicated internal structure (e.g. S(cid:1)engo€r
turally below the Tav(cid:1)sanlı Zone (Fig. 1), is a low- et al., 1984; Gessner et al., 2001c; Okay, 2001; Ring
grade HP metamorphic belt (Pourteau et al., 2010), et al., 2001). Yet, a general lithostratigraphy was
representing a Mesozoic continental passive margin, restored. It consists of a polymetamorphic (Pan-Afri-
including a polymetamorphic Precambrian substra- can and Alpine metamorphisms; Candan et al., 2011)
tum (Candan et al., 2005). Fe-Mg-carpholite, glauco- gneiss core overlain by monometamorphic Palaeozoic
phane, and calcite pseudomorphs after aragonite schist and Mesozoic marble covers, and a Palaeogene
documented throughout the Afyon Zone are evidence olistostrome (Schuiling, 1962; Du€rr, 1975; O€zer et al.,
for blueschist facies metamorphism (Candan et al., 2001). The Menderes Massif is generally character-
2005; Pourteau et al., 2010). Based on carpholite– ized by Barrovian-type metamorphism (Bozkurt &
chloritoid assemblages in metasedimentary rocks, Oberh€ansli, 2001; Gessner et al., 2001a; van Hinsber-
metamorphic conditions experienced by the Afyon gen, 2010) with P–T conditions (estimated in mono-
Zone were estimated to be 0.8–1.1 GPa and 250– metamorphic Palaeozoic schists from the southern
350 °C (Candan et al., 2005; Pourteau, 2011). No Menderes Massif) of 0.4–0.5 GPa and 330–530 °C
isotopic date information is available yet for the (monometamorphic Palaeozoic schists from the
metamorphic evolution of the Afyon Zone, but southern Menderes Massif; Whitney & Bozkurt,
©2013JohnWiley&SonsLtd
NEOTETHYS CLOSURE HISTORY OF ANATOLIA 589
2002; R(cid:3)egnier et al., 2003). In the same region, Fe- omorphs after aragonite in marbles are widespread
Mg-carpholite relicts were described at the base of evidence for blueschist facies metamorphism
the Mesozoic carbonate sequence, as evidence for (Oberh€ansli et al., 2001; Rimmel(cid:3)e et al., 2003a, 2006).
HP–LT metamorphism in this part of the massif Peak P–T conditions were estimated to be ~1.2 GPa
(Rimmel(cid:3)e et al., 2003b; Whitney et al., 2008). The and 400 °C (Rimmel(cid:3)e et al., 2005). Rocks near the
€
preservation of carpholite (i.e. the occurrence of re- base of the Oren Unit (including the klippen) experi-
licts vs. pseudomorphs) correlates with that of dia- enced isothermal decompression at ~400–450 °C,
spore in metabauxite horizon in platform-type whereas rocks at higher tectonostratigraphic positions
marbles (Rimmel(cid:3)e et al., 2003b), as also mentioned experienced cooling during decompression (Rimmel(cid:3)e
in the Cycladic Blueschist Unit (Candan et al., 1997). et al., 2005). This shows that the thermal evolution of
P-T estimates are 1.2–1.4 GPa/470–550 °C in Kuru- HP–LT rocks during their exhumation can signifi-
dere (location of sample Kuru0110; Fig. 1), 0.6– cantly vary within a single tectonic unit. 40Ar–39Ar
0.8 GPa/400–450 °C in Bah(cid:1)cekaya, and 0.9–1.1 GPa/ dates between 90 and 70 Ma were obtained on pheng-
380–480 °C in Nebiler (location of sample Nebil0101; ite from metasedimentary rocks as the only isotopic
Fig. 1) (Rimmel(cid:3)e et al., 2003b, 2005; Whitney et al., estimate for the metamorphism of the O€ren Unit
2008). This HP–LT metamorphic event probably (Ring & Layer, 2003). Stratigraphic relations bracket
affected the entire Mesozoic marble sequence, the HP–LT metamorphic event in a latest Cretaceous
described as stratigraphically continuous (Du€rr, 1975; to Eocene interval (Rimmel(cid:3)e et al., 2003a). Owing to
€
Ozer et al., 2001), but not the Palaeozoic schist its structural position, metamorphic evolution and
€
cover, which, in parts is only affected by low-grade lithostratigraphy, the Oren Unit might represent the
conditions (chlorite zone; Schuiling, 1962), lacks HP westerncontinuationoftheAfyonZone.
minerals. Therefore, in this region, only the marble In Western Anatolia, the Cycladic Blueschist Unit,
cover (including basal conglomerates containing which is otherwise widely exposed in Aegea, consists
carpholite relicts) is considered to have experienced of two HP metamorphic units (a lower platform
HP–LT metamorphism. Hence, the Menderes Massif sequence and an upper metamorphic m(cid:3)elange) that
sensu lato is sub-divided into the low- to middle-P lie tectonically over the western Menderes Nappes
Menderes Nappes (Gessner et al., 2001a), and the (Candan et al., 1997; Oberh€ansli et al., 1998; Ring
newly termed HP–LT Kurudere–Nebiler Unit et al., 1999; Okay, 2001). Peak P–T estimates are
(Fig. 1). The timing of this HP–LT metamorphism 1.1–1.5 GPa and 440–550 °C (C(cid:1)etinkaplan, 2002;
was not yet estimated by means of isotopic dating Ring et al., 2007). Available phengite 40Ar–39Ar dates
methods, but it is predated by the middle Palaeocene indicate exhumation between c. 40 and 32 Ma, along
depositional age of the olistostrome topping the Cre- an isothermal decompression path through the
taceous marble (O€zer et al., 2001), and probably greenschist facies (Oberh€ansli et al., 1998; Ring et al.,
post-dated by the Barrovian metamorphism the Men- 2007). This is consistent with HP metamorphic tim-
deres Massif underwent. Available isotopic dates for ing in the Cycladic Blueschist Unit in Aegea (e.g.
the metamorphic evolution of the Menderes Massif Jolivet & Brun, 2008; Ring et al., 2010), where the
sensu lato (40Ar–39Ar and 87Rb–86Sr on white mica metamorphic peak was dated to c. 52 Ma
and biotite) are scattered between c. 62 and 27 Ma (176Lu-176Hf on eclogitic garnet; Lagos et al., 2007).
(Satır & Friedrichsen, 1986; Bozkurt & Satır, 2000; Several tectonic units of the Anatolide–Tauride
Lips et al., 2001). Therefore, the HP–LT metamor- Block are topped by widespread ophiolites (Fig. 1)
phic history of the Kurudere–Nebiler Unit likely that were generated above a subduction zone (e.g.
€
occurred sometime between the late Palaeocene and Onen & Hall, 1993; Parlak et al., 1996). Greenschist-
the Eocene. It is noteworthy that Ring et al. (1999) toamphibolitefaciessub-ophioliticmetamorphicsoles
regarded most of the Mesozoic cover of the entire formed during subduction initiation were dated to
Menderes Massif sensu lato as the continuation of 93–90 Ma(e.g.Robertson,2002;C(cid:1)eliket al.,2006).
the Cycladic Blueschist Unit, even parts free of HP
relicts.
€ P–T ESTIMATION AND 40Ar–39Ar
The Oren Unit (Pourteau et al., 2010), formerly
GEOCHRONOLOGY
referred to as the metamorphosed part of the Lycian
Nappes (Oberh€ansli et al., 2001; Rimmel(cid:3)e et al., Fe-Mg-carpholite in metasedimentary rocks is the
2003a), lies structurally over the Kurudere–Nebiler main, or locally the only evidence of ‘Alpine’ HP–LT
€
Unit, the Menderes Nappes, and the Cycladic Blue- metamorphism in the Afyon Zone, the Oren Unit
schist Unit (Fig. 1). Kinematic indicators (Rimmel(cid:3)e and the Kurudere–Nebiler Unit. Phengite, typically
et al., 2003a, 2006) indicate that, during the late stage accompanied by chlorite, commonly occurs in carph-
€
ofitsexhumation,theOrenUnitwastransportedover olite-bearing rocks, offering the chance to determine
these units towards the ESE (after restoration from the timing of metamorphism by using 40Ar–39Ar geo-
the Neogene extensional deformation; Pourteau et al., chronology. Several studies previously showed that
2010; van Hinsbergen, 2010). Fe-Mg-carpholite in the this approach yields reliable metamorphic dates (e.g.
lower metasedimentary rocks and calcite pseud- Jolivet et al., 1996; Agard et al., 2002; Wiederkehr
©2013JohnWiley&SonsLtd
590 A. POURTEAU ET AL.
et al., 2009; Oberh€ansli et al., 2012). In these rocks, intergrown, but pseudomorphs are locally only com-
white mica may have formed before, during and after posed of phengite. In situ laser ablation therefore was
carpholite growth, so several generations can coexist the most suitable method to extract Ar from selected
in a single sample and thus provide various isotopic chlorite-free domains (Fig. 2a). The sample Afy0212
dates (e.g. Agard et al., 2002; Wiederkehr et al., was described by Pourteau (2011) for remarkable
2009). To evaluate the possibility of such situations, carpholite–chloritoid textures. Pyrophyllite, which
P–T conditions prevailing during the growth of the occurs as large crystals wrapping carpholite and chlo-
dated white mica can be estimated by chlorite–pheng- ritoid, contains discrete layers of celadonite-poor
ite thermobarometry (multi-equilibrium approach; (X = 0.04), muscovite-rich phengite. In sample
Cel
Vidal & Parra, 2000) (see below). Therefore, when Ku€t0815, carpholite is rather well preserved, but dis-
possible, 40Ar–39Ar dating was combined with P–T plays evidence for local breakdown into chlorite and
estimation in individual samples to decipher the sig- pyrophyllite that here again contains thin celadonite-
nificance of isotopic dating results. poor (X = 0.02) phengite interlayers. To extract
Cel
enough Ar for isotope analysis from these pyrophyl-
lite–phengite composite grains, stepwise heating was
Sampleselection
applied on separated grains from these two samples.
In the region of Afyon city, carpholite-bearing
Selectioncriteria
metapelitesandmetaconglomeratescontainabundant,
For 40Ar–39Ar geochronology, selected samples con- fine-grained phengite, intergrown with pyrophyllite,
tain white mica that (i) is rich in potassium and poor quartz and iron oxides, but notably no chlorite (Can-
in sodium (phengitic v. paragonitic mica); (b) is obvi- dan et al., 2005; Pourteau, 2011). Sample Bay0851
ously of metamorphic origin; (c) occurs as medium- (Fig. 1; Table S2) is a carpholite-bearing silvery phyl-
sized to coarse crystals (> 125 lm); (d) is not inter- lite (Fig. 2b). The notable absence of chlorite in this
grown with chlorite or pyrophyllite; and (e) was not sample, on the one hand, prevents its use for accurate
affected by late fluid circulation (i.e. no fractures, no P–T estimation, but, on the other hand, indicates that
late calcite crystallization). Bearing in mind that the peak mineral assemblage (represented by carpho-
carpholite-bearing quartz veins originate from the lite and chloritoid) was well preserved. Most matrix
crystallization of a fluid phase under HP–LT condi- phengite is poor in celadonite (X = 0.01–0.08). By
Cel
tions, white mica in such a rock type is exclusively of contrast, phengite in contact with carpholite fan-
metamorphic origin. Therefore, to avoid the risk of shaped aggregates yields X values up to 0.32, and is
Cel
the radiogenic Ar inheritance of detrital mica recrys- thus among the highest-P phengite found in the Af-
tallized during metamorphism, quartz vein samples yon Zone. This suggests that Ar isotope analysis in
were preferentially selected. this sample might provide near-peak metamorphic
dates. With the aim of analysing higher-P and lower-
P phengite distinctively, sample Bay0851 was pre-
Selectedsamples
paredforinsitulaserablation.
Because both fine grain size and common intergrowth In the Konya region (Fig. 1), some carpholite-
of white mica with chlorite critically restrict sample quartz veins also contain white mica, but this con-
selection, a total of nine samples were selected, six tains a significant paragonite component, preventing
from the Afyon Zone, one from the O€ren Unit, and any accurate P–T estimation and, a priori, hampering
two from the Kurudere–Nebiler Unit. Carpholite- 40Ar–39Ar geochronological investigations. Ar isotope
bearing rocks in the Afyon Zone are mostly quartz analysis was tentatively performed on a sample from
segregations (or ‘veins’), and also quartz–mica schists the region of Konya (Kon0316; Fig. 1; Table S2),
and pyrophyllitite (Candan et al., 2005; Pourteau, but laser ablation experiments did not release any sig-
2011). Veins typically consist of the assemblage nificant gas amount for Ar isotope analysis.
quartz + carpholite (cid:1) pyrophyllite (cid:1) chlorite (cid:1) chlo- In the easternmost regions of Yahyalı and Kayseri
ritoid. Phengite and paragonite are more rarely (Fig. 1; Table S2), carpholite, which is again
observed. Samples Afy0206, Afy0212 and Ku€t0815 observed only as isolated fibres in quartz, was sub-
were collected from the area near Ku€tahya (Fig. 1; stantially pseudomorphed by phengite and chlorite
Table S2). In sample Afy0206, phengite and chlorite through the reaction R1. Chloritoid also occurs
occur as needle-shaped associations that are up to locally among phengite–chlorite associations, proba-
2 cm in length, and carpholite is observed only as bly as a result from the progressive prograde reaction
tiny fibres in quartz (Fig. 2a). This suggests that large carpholite = chloritoid + quartz + water (R2; Chopin
needlesofcarpholitewereconsistentlypseudomorphed & Schreyer, 1983). In the sample Yah04 that was
to white mica and chlorite through the continuous, selected, phengite and chlorite are mostly intergrown,
decompression reaction carpholite + phengite = chlo- but also occur as isolated crystals, allowing easy min-
rite + muscovite + quartz + water (R1; e.g. Bousquet eral separation.
€
et al., 2002), whereas tiny needles, included in quartz, In the Oren Unit, carpholite-bearing quartz veins
were preserved. Phengite and chlorite are typically commonly contain peak to retrograde phengite and
©2013JohnWiley&SonsLtd
NEOTETHYS CLOSURE HISTORY OF ANATOLIA 591
(a) (b)
(c) (d)
Figure 2. Opticalmicrophotographs(crosspolars)ofthesamplesinvestigatedbyinsituUVlaserablation.a)Retrogradephengite
andchloriteoccursasneedle-likepseudomorphsaftercarpholite,preservedonlyasfibresincludedinquartz,inAfy0206(western
AfyonZone).b)Pre-foliation,near-peakphengiteisobservedamongquartzandcarpholite,whereassyn-foliation,retrograde
phengiteconstitutesmostoftheschistosematrixinphylliteBay0851(centralAfyonZone).c)Texturalequilibratedretrograde
phengite,kyaniteandchloriteformedduetothebreakdownofcarpholite(fibresinquartz)inKuru0110(westernKurudere–
NebilerUnit).d)Lateretrogradewhitemicagrewduringthedevelopmentofpost-kyanitefoliationinNebil0101(eastern
Kurudere–NebilerUnit).MineralabbreviationsareafterWhitney&Evans(2010).
chlorite (Oberh€ansli et al., 2001; Rimmel(cid:3)e et al., Azan~on & Goff(cid:3)e, 1997; Rimmel(cid:3)e et al., 2003b). The
€
2003a, 2005). The selected sample Oren001 (Fig. 1; eastern sample Nebil0101 is a metamorphosed quartz
Table S2) stemmed from a quartz vein encompassing micro-conglomerate containing kyanite + chlori-
well-preserved carpholite needles partly replaced by toid + phengite as the main assemblage. Chlorite was
phengite and chlorite (Rimmel(cid:3)e et al., 2003a). In this reported to be abundant in this sample (Rimmel(cid:3)e
sample, carpholite–chlorite–phengite–quartz–water et al., 2003b), but was not seen in the selected sample
multi-equilibrium calculations yielded P–T conditions fragment. Furthermore, no carpholite relict was
of 0.8–1.1 GPa and 320–380 °C (Rimmel(cid:3)e et al., 2005). observed in this sample, whereas kyanite–chlorite
From the Kurudere–Nebiler Unit, two samples pseudomorphs after carpholite were described from
previously studied by Rimmel(cid:3)e et al. (2003b, 2005) the same locality. In the selected rock fragment,
were selected: Kuru0110, located at the diaspore– scarce phengite layers among the recrystallized quartz
corundum isograd, and Nebil0101, located within the matrix define a weak foliation that wraps crystals of
corundum zone. Sample Kuru0110 is a quartz vein kyanite (Fig. 2d) and chloritoid, and therefore proba-
containing the main assemblage kyanite + chlo- bly post-date their formation.
rite + phengite + pyrophyllite + carpholite. Kyanite,
chlorite and phengite, observed in equilibrium tex-
P–Testimation– analytical methods
tures, whereas carpholite occurs only as fibres
included in quartz (Fig. 2c). Phengite therefore grew Metamorphic P–T conditions for samples Afy0206
coevally with the products of the reaction and Yah04 (western and eastern Afyon Zone, respec-
carpholite = chlorite + kyanite + quartz + water (e.g. tively) were estimated using chlorite–quartz–water
©2013JohnWiley&SonsLtd
592 A. POURTEAU ET AL.
thermometry (Vidal et al., 2005, 2006), phengite–
quartz–water thermobarometry (Dubacq et al., 2010), Chlorite–quartz–waterthermometry
and the chlorite–phengite–quartz–water multi-equilib- The evolution of chlorite composition with the tem-
rium approach first proposed by Vidal & Parra perature (e.g. Vidal et al., 2001) in equilibrium with
(2000) and subsequently applied in many different quartz and water can be modelled using the two inde-
geological contexts (e.g. Trotet et al., 2001, 2006; pendent equilibria (Vidal et al., 2005, 2006):
Parra et al., 2002; Arkai et al., 2003; Vidal et al.,
2006; Ganne et al., 2012; Grosch et al., 2012; Lanari 5Mg-amesiteþ4daphnite$4clinochlore
et al., 2012a,b). When possible (sample Yah04), P–T (1)
conditions were also estimated from chlorite–chlori- þ5Fe-amesite
toid–quartz–water assemblages. 4daphniteþ6sudoite$3Mg-amesiteþ5Fe-amesite
(2)
þ14quartzþ8H O
2
Mineral compositionanalysis
where Mg-amesite, Fe-amesite, daphnite, sudoite
Elementmapping(forSi,Al,Fe,Mg,KandNa)com-
and clinochlore are chlorite end-members. The activi-
bined with spot analysis was performed using an elec-
ties of these end-members were calculated using the
tron microprobe JEOL JXA8200 (Potsdam
thermodynamic data and the activity model used in
University), following the procedure described by De
Vidal et al. (2006). For any given pressure, the tem-
Andradeet al.(2006).Areasselectedformappingcon-
perature of crystallization and the ferric iron content
tain the coarsest grains possible and display sharp
(X ) of chlorite were simultaneously estimated
Fe3+
grain boundaries (i.e. limited intergrowths). Operating
using criterion based on the convergence of the equi-
conditions for maps were 15 keV accelerating voltage,
libria (1) and (2). Following the strategy previously
5 lm beam size, 100 nA beam current, for counting
detailed in Lanari et al. (2012a), T and X of
Fe3+
times of 300 ms on peak. Spot analyses were acquired
chlorite were estimated at a fixed pressure of 4 kbar
with 15 keV accelerating voltage, 5 lm beam size,
and a water activity equal to 1. Convergence was
10 nA beam current, for counting times of 20 s on
assumed when the temperature difference between
peak and 10 s on background. Standards of pyrope the two equilibria was < 30 °C. Each chlorite group
forSi, rutileforTi,pyropeforAl,hyperstheneforFe,
was considered individually to calculate a tempera-
fayalite for Mn (to analyse accurately very low ture using the chlorite–quartz–water equilibrium (at a
amount of Mn), diopside for Mg, diopside for Ca,
given pressure) following the method of Vidal et al.,
anorthiteforNa,andmicroclineforKwereused.
(2005); Vidal et al. (2006). Results are displayed as
A distinct chemical phase (chlorite, phengite,
maps of chlorite temperatures (Fig. 3b-b’).
quartz, Ti oxide; Fig. 3–b–b’) was attributed to each
pixel of the element maps using XMapTools software
(Lanari et al., 2012a,b, 2013; http://www.xmaptools. Phengite–quartz–waterbarometry
com). Then, profiles of individual spot quantitative
analyses traced across phengite–chlorite aggregates The equilibrium conditions of the assemblage pheng-
ite + quartz + water can be modelled using the fol-
were used as standards to obtain a quantitative anal-
lowing three equilibria (among which two are
ysis (in oxide weight percent) from each pixel of the
independent):
qualitative element maps. Therefore, in each investi-
gated area, a large number of spot analyses were 3celadoniteþ2pyrophyllite$2muscoviteþbiotite
acquired for each phase to ensure the precision of
þ11quartzþ2H O
this quantification. Structural formula and atom site 2
repartition for each pixel of chlorite and phengite (3)
were calculated using the functions Chl-StructForm
3celadoniteþ2pyrophyllite(cid:3)1H O
and Phg-StructForm in XMapTools, following the 2 (4)
solid solution models of Vidal & Parra (2000), Vidal $2muscoviteþbiotiteþ11quartzþ3H2O
et al. (2006) and Dubacq et al. (2010). Chlorite and
pyrophyllite(cid:3)1H O$pyrophylliteþH O (5)
phengite pixels were then distributed into various 2 2
composition groups using the K-means clustering where celadonite, pyrophyllite, pyrophyllite(cid:3)1H O,
2
method (Saporta, 1990) for each of which an average muscovite and biotite are the end-members of pheng-
composition was calculated. ite. The activities of these end-members were calcu-
In addition, the presence of very fine-grained chlo- lated using the thermodynamic data and activity
ritoid in Yah04, albeit not allowing element mapping, model used in Dubacq et al. (2010). Equilibrium con-
allowed calculation of P–T conditions prevailing for vergence at various P–T conditions was achieved by
local chlorite–chloritoid–quartz–water equilibrium changing the amount of interlayer water (i.e. the pro-
(see detail below). The composition of chloritoid and portion of pyrophyllite(cid:3)1H O). For the average
2
chlorite in equilibrium textures was determined composition of each phengite group (except for
through individual spot analysis. muscovite–paragonite textures), the method of Du-
©2013JohnWiley&SonsLtd
NEOTETHYS CLOSURE HISTORY OF ANATOLIA 593
(b)
(a)
(c)
(d)
(b’)
(a’)
(c’)
(d’)
Figure 3. Resultsofmulti-equilibriumthermobarometryforthesamplesAfy0206andYah04.a)anda’)Ternarydiagramsfor
whitemicaandchloritecompositions.Eachpointrepresentsthecompositionofonepixel.Coloursindicatethestatistical
compositiongroupsmentionedinthetext.MineralabbreviationsareafterWhitney&Evans(2010).b)andb’)Mapviewsof
phasedistribution,andmicapressure–andchloritetemperaturegroupsintheinvestigatedareas.“XF”standsforFe3+/Fe in
total
chlorite.c)andc’)chlorite–phengite–quartz-water(Afy0206)andchlorite–chloritoid–quartz–water(Yah04)multi-equilibrium
calculationsusingaveragegroupcompositions.d)andd’)Resultsummary.Greyfieldsrepresentthestabilityfieldsofcarpholite
withMg/(Mg+Fe) = 0.30(d)and0.35(d’),calculatedwithTheriakDomino(DeCapitani&Petrakakis,2010)followingPourteau
(2011).Seetextfordetails.
©2013JohnWiley&SonsLtd
594 A. POURTEAU ET AL.
bacq et al. (2010) was used to plot a P–T line corre- tions (C3 and c3) tend to occur in external parts of
sponding to the phengite–quartz–water equilibrium, chlorite aggregates (Fig. 3–b–b’). Because, at fixed
with variable hydration states (see the method bulk rock composition, the abundance of the sudoite
description above). Phengite equilibrium pressure and end-member (vacancy on the octahedral site [M1])
hydration state were then estimated at the tempera- decreases with increasing temperature (Cathelineau &
tures of the chlorite–quartz–water equilibria (Lanari Nieva, 1985; Hillier & Velde, 1991; Vidal et al.,
et al., 2012a). Results are displayed as maps of 2001), this zoning trend indicates that chlorite grew
phengite pressures in Fig. 3–b–b’. during cooling. This supports the statement that
chlorite is the product of carpholite retrogression. On
the other hand, celadonite-content in phengite, thus
Chlorite–phengite–quartz–water thermobarometry
its Si content, increases with pressure (Bousquet
Once chlorite–quartz–water thermometry and pheng- et al., 1998; Agard et al., 2001). Therefore, owing to
ite–quartz–water barometry were applied, thermody- higher celadonite-contents, phengite M1 formed at
namic equilibrium between each chlorite group and higher pressure than M2 (X = 0.16 v. 0.04), and
Cel
each phengite group was tested through a full chlo- m1 grew at slightly higher pressure than m2
rite–phengite multi-equilibrium approach involving (X = 0.04 v. 0.02) (Fig. 3–a–a’; Table 2). Since
Cel
X in both minerals. P–T equilibrium conditions phengite formed through the decompression reaction
Fe3+
of chlorite–phengite–quartz–water assemblages were R1, higher-P M1 and m1 phengites probably formed
calculated from the convergence of 64 equilibria earlier than M2 and m2, respectively. Carpholite in
obtained from the chlorite and phengite end-members Afy0206 and Yah04 gives X values of 0.5 and 0.3–
Mg
mentioned above (details in Vidal et al., 2006). Only 0.4, respectively. Chloritoid analysis in Yah04 reveals
equilibria showing a good convergence were selected. that, from grain to grain, X varies between 0.12
Mg
and 0.22 (e.g. Table 3).
Chlorite–chloritoid–quartz–water thermobarometry
P–T equilibrium conditions for chlorite–chloritoid– Thermobarometricresults
quartz–water assemblage were calculated from the In Afy0206, chlorite–quartz–water thermometry using
convergence of 39 (three independent) equilibria theC1,C2,andC3chloriteaveragecompositionspro-
obtained from the chlorite end-members mentioned vides temperatures of ~351, 258 and 197 (cid:1) 50 °C and
above with Mg-chloritoid and Fe-chloritoid (thermo- X of 6, 19 and 28% respectively (Table 2). The
Fe3+
dynamic data from JUN92.bs database, updated two phengite P–T equilibrium lines obtained from
from Berman (1988)). Only equilibria showing a good phengite–quartz–water barometry intersect only with
convergence were selected. This thermobarometer was the C1 chlorite line (T = 350 °C). Intersection points
applied to Yah04 using individual spot analyses of indicate for M1 and M2 mica pressures of 0.6 and
chlorite and chloritoid observed in equilibrium textures. < 0.3 GPa respectively. In addition, multi-equilibrium
calculations using group average compositions reveal
thermodynamic equilibrium between C1 chlorite and
P–Testimation – results M1 mica, and yield estimates of ~0.6 GPa and
~350 °C (Fig. 3c). These P–T conditions areconsistent
Mineral analysesin Afy206and Yah04 with those determined individually using chlorite–
Whereas under the optical microscope chlorite and quartz–waterandphengite–quartz–waterequilibria.
white mica seem to have formed during a single event In Yah04, chlorite–quartz–water thermometry
(carpholite breakdown reaction R1), several composi- yields temperatures of 384, 320 and 248 (cid:1) 50 °C (at
tion groups, and thence P- and T groups, can be dis- 1.0 GPa) and X of 6, 19, and 28% for the c1, c2
Fe3+
tinguished for each phase. In each sample, the and c3 chlorite groups respectively (Table 2). Here
statistical analysis gave chlorite pixel distributions of again, the two phengite P–T equilibrium lines
three compositional groups: C1 to C3 (upper case) in obtained for m1 and m2 intersect the highest-T chlo-
Afy0206, and c1 to c3 (lower case) in Yah04. On the rite (c1, at 380 °C; Fig. 3d’) at 0.7 and < 0.4 GPa,
other hand, three compositional groups were also dis- respectively. Unfortunately, no equilibrium was
tinguished for mica pixels: M1, M2, M + P in found between the c1 chlorite and the m1 mica, but
Afy0206, and m1, m2, m + p in Yah04, where M + P an equilibrium was identified between the c1 chlorite
and m + p stand for muscovite–paragonite mixtures. and average chloritoid composition (X = 0.14) at
Mg
Average compositions for each of these groups, ~0.6 GPa and ~350 °C (Fig. 3c’). In addition, P–T
except the muscovite–paragonite mixtures, are shown conditions calculated for individual chlorite–chlori-
in Table 2. In both samples, chlorite compositional toid pairs of spot analyses range between 0.82 and
groups display an increase in the sudoite-content 1.13 GPa, and between 383 and 425 °C (Table 3).
from C1 to C3 and from c1 to c3, but no significant Therefore, chloritoid growth, out of the carpholite
change in the proportion of the other end-members stability field (Fig. 3d’), probably represents the HP
(Fig. 3–a–a’ and Table 2). Sudoite-richest composi- thermal climax experienced by this sample.
©2013JohnWiley&SonsLtd