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Lithos 84 (2005

Alpine high-P/low-T metamorphism of the Afyon Zone and

implications for the metamorphic evolution of

Western Anatolia, Turkey

Osman Candana, Mete Cetinkaplana, Roland Oberhanslib,*,

Gaetan Rimmelec, Cuneyt Akala

aDokuz Eylul University, Engineering Faculty, Department of Geological Engineering, 35100, Bornova Izmir, TurkeybInstitut fur Geowissenschaften, Universitat Potsdam, Postfach 601553, Potsdam 14415, Germany

cEcole Normale Superieure Laboratoire de Geologie UMR 8538 du C.N.R.S. 24, rue Lhomond 75 005 Paris, France

Received 25 February 2004; accepted 7 February 2005

Available online 24 June 2005

Abstract

Carpholite+chloritoid+pyrophyllite association occurs widely in the Triassic metaclastic rocks of the Afyon Zone in west-

central Turkey. Fe-Mg-carpholite is associated with rare aragonite pseudomorphs and glaucophane in marbles and metabasites,

respectively. The Afyon Zone consists stratigraphically of a Pan-African basement and an overlying Mesozoic cover sequence.

The Pan-African basement, which shows Barrovian-type amphibolite-facies metamorphism, comprises garnet–mica schists,

intruded by sodic amphibole-bearing metagabbros and leucocratic metagranites. It is unconformably overlain by a continuous

metasedimentary sequence extending from Triassic to early Palaeocene. This cover sequence begins with metaconglomerates,

which pass upwards into phyllites. Fe-Mg-carpholite occurs within this metaclastic sequence as rosette-like crystals in

metapelites and fibres in quartz segregations. The metaclastic rocks are succeeded by metamorphosed platform carbonates,

grading into Latest Mesozoic metamorphosed pelagic limestones, which in turn progress up to a Late Mesozoic–Early Tertiary

olistostrome. This sequence is tectonically overlain by the blueschists of the TavsanlV Zone. Fe-Mg-carpholite-bearing

assemblages imply temperatures of about 350 8C and minimum pressures of 6–9 kbar, corresponding to burial depths of

about 30 km for the Mesozoic passive continental margin sediments and the underlying Pan-African supracrustal metasediments

and metaintrusives.

The metamorphic rocks of the Afyon Zone are unconformably overlain by Upper Palaeocene–Lower Eocene sedimentary

rocks, indicating a Paleocene age for the regional HP/LT metamorphism. This implies continuous younging of HP/LT

metamorphism in the Anatolides related to northward subduction of the Anatolide–Tauride platform beneath the Sakarya

Zone. From north to south this involved the TavsanlV Zone (Campanian, 80F5 Ma), the Afyon Zone (Palaeocene?), the

0024-4937/$ - s

doi:10.1016/j.lith

* Correspondin

E-mail addre

) 102–124

ee front matter D 2005 Elsevier B.V. All rights reserved.

os.2005.02.005

g author.

ss: roob@geo.uni-potsdam.de (R. Oberhansli).

O. Candan et al. / Lithos 84 (2005) 102–124 103

Menderes Massif (Middle Eocene) and the Lycian Nappes (Late Cretaceous–Eocene?), all of which were probably derived from

the frontal part of the Anatolide–Tauride platform.

D 2005 Elsevier B.V. All rights reserved.

Keywords: HP/LT metamorphism; Neo-Tethys; Geodynamic evolution; Carpholite

1. Introduction

From the Riff in Morocco to Oman, relics of high-

pressure minerals give evidence of subduction related

processes. Fe-Mg-carpholite and Mg-rich chloritoid

were recognized as typical blueschist-facies minerals

in metapelitic rocks by Chopin and Schreyer (1983).

Since then carpholite has been reported from many

high-pressure metamorphic terrains in the Alpine–Hi-

malayan orogen, including the Betic Cordilleras

(Spain, Goffe et al., 1989; Azanon et al., 1998),

Monte Argentario (Italy, Theye et al., 1997), the

Western Alps (Goffe and Velde, 1984; Goffe and

Chopin, 1986), the Central Eastern Alps (Oberhansli

et al., 1995; Goffe and Oberhansli, 1992), Crete and

the Peleponnes, Greece (Seidel et al., 1982; Theye et

al., 1992), Samos island, Cycladic Complex (Okrusch

et al., 1985) and in Oman (Goffe et al., 1988, Michard

et al., 1994). Here we report the widespread occur-

rence of carpholite in the Afyon Zone in west-central

Turkey and discuss its petrologic and tectonic signif-

icance. This carpholite discovery complements those

recently reported from the Lycian Nappes and the

Menderes Massif (Oberhansli et al., 2001; Rimmele

et al., 2003a,b).

Western Turkey is divided into the Pontides in the

north, which is separated by the Izmir Ankara suture

from the Anatolide–Tauride Block in the south (Fig.

1), (Ketin, 1966; Sengor and Ylmaz, 1981; Okay et al.,

1996). The Anatolide–Tauride Block is a Gondwana-

related continental block, which during the Mesozoic

was bounded to the north and south by the branches of

Neo-Tethys. The northern margin of the Anatolide–

Tauride Block was deformed and regionally metamor-

phosed during the Alpine orogeny, and is subdivided

into several tectonic zones based on the age and type of

metamorphism. The main metamorphic units are from

north to south: the TavsanlV Zone, the Afyon Zone, the

Menderes Massif and the Lycian Nappes.

The NW–SE-trending Afyon Zone extends over a

distance of 600 km south of the TavsanlV Zone from

BalVkesir via Konya to Bolkardag. The Afyon Zone

was regarded to have undergone lower greenschist-

facies metamorphism (Okay, 1984; Okay et al., 1996;

Goncuoglu et al., 1992). However, here we report

sodic amphibole and widespread Fe-Mg-carpholite–

chloritoid–pyrophyllite assemblages in the basement

and cover sequence of the Afyon Zone, respectively.

The recognition of Fe-Mg-carpholite and associated

assemblages in the Afyon Zone, as well as the Lycian

Nappes and the Menderes Massif, requires us to re-

consider the tectono–metamorphic evolution of the

Anatolide–Tauride Block, which is related to the sub-

duction and subsequent collision processes of the

Neo-Tethys Ocean in Albian–Eocene times.

2. Geological setting

The Afyon Zone, described by Okay (1984), is one

of the main tectonic zones of the Anatolides. Al-

though name and boundaries are disputed (Okay,

1984: Afyon– Bolkardag Zone; Okay et al., 1996

and Tolluoglu et al., 1997: Afyon Zone; Ozcan et

al., 1990 and Goncuoglu et al., 1996: Kutahya– Bolk-

ardag Belt including both Afyon and TavsanlV Zones),

it is generally accepted that the Afyon Zone comprises

platform sediments that were deposited on the north-

ern passive continental margin of the Anatolide–Taur-

ide Block. Previous studies suggested that the

stratigraphy of the Afyon Zone is similar to that of

the TavsanlV Zone and to that of the Mesozoic–Early

Tertiary series of the Lycian Nappes, as well as to that

of the cover series of the Menderes Massif (Okay et

al., 1996).

The Paleozoic and Mesozoic units of the Afyon

Zone, especially around Konya and Afyon, have been

dated paleontologically despite the low-grade meta-

morphism (Ozcan et al., 1988; Goncuoglu et al.,

1992). Previous studies have compiled the general

lithostratigraphy as follows: in the central part of the

Afyon Zone Carboniferous to Permian basement

Fig. 1. Tectonic zones of the Anatolides and localities of carpholite occurrence in northwest Turkey. The study area shown in Fig. 2, is indicated.

O.Candanet

al./Lith

os84(2005)102–124

104

O. Candan et al. / Lithos 84 (2005) 102–124 105

schists (Ozcan et al., 1988; Goncuoglu et al., 1992)

are unconformably overlain by Mesozoic cover

sequences composed of basal conglomerates followed

by neritic platform sediments with an age interval

between Anisian to Early Maastrichtian (Ozcan et

al., 1988; Goncuoglu et al., 1992; Tolluoglu et al.,

1997). The basal conglomerates have been interpreted

to reflect rifting and rapid uplift on the northern

continental margin basement of Gondwana, related

to the opening of the Izmir Ankara branch of the

Neo-Tethys Ocean as a result of the southward sub-

duction of the Paleotethyan oceanic crust (Goncuoglu

et al., 1992). The Cretaceous platform carbonates pass

gradually into Upper Cretaceous marbles of pelagic

origin with chert layers, reflecting platform subsi-

dence (Goncuoglu et al., 1992). An Upper Maastrich-

tian to Lower Palaeocene olistostromal unit is

overlying this pelagic sequence. This olistostromal

unit is unconformably covered by non-metamorphic

Upper Palaeocene to Lower Eocene shallow water

sediments (Ozcan et al., 1988).

3. Lithostratigraphy of the central part of Afyon

Zone

The study area is located 30 km northeast of

Afyon, in the central part of the 600-km long, NW–

SE-trending Afyon Zone (Fig. 2). A generalized strati-

graphic column of this area is given in Fig. 3. The

succession can be divided into two groups separated

by an unconformity: 1) a lower assemblage of am-

phibolite-facies schists with mafic and felsic metaig-

neous rocks (Pan-African Basement) and 2) an upper

sequence of low-grade metasediments whose proto-

liths were deposited in continental to shallow-water

environments. In the study area, strongly deformed,

poly-metamorphic garnet–mica schist, chlorite–albite

schist and quartzite intercalations of over 4 km struc-

tural thickness form the upper greenschist to amphib-

olite-facies crystalline basement of this part of the

Afyon Zone. In the central part of the Afyon Zone

around Kutahya, KVrka, Bayat and Bolvadin, this

assemblage is exposed as scattered outcrops of vary-

ing size below Mesozoic and Tertiary sedimentary and

volcanic rocks. This schist unit extends over 150 km

and implies that Pan-African Basement is present

below the central part of the Afyon Zone.

The Pan-African schist unit is dominated by the

garnet–mica schists and is characterized by largely

homogenous compositions. Some minor gradual ver-

tical as well as lateral (decimeter to meter) changes in

composition precludes the identification of subunits

and detailed mapping. The schist unit is conspicuous-

ly free of carbonate layers. However, dark gray to

black quartzite layers, up to several meters thick are

common. The fine lamination and lack of micas in

these quartzite layers suggest a high maturity for their

sedimentary precursors, which represent beach depos-

its. Pre-Alpine, upper greenschist-facies metamor-

phism is suggested by the assemblage quartz+

albite+chlorite+white micaFgarnetF biotite. Al-

though no fossils and geochronological data have

yet been reported, a Pan-African age can be assumed

for this psammitic to pelitic sequence based on re-

gional correlations, especially with the Menderes

Massif (Dora et al., 2001; Koralay et al., 2002).

The schist unit also comprises minor volumes of

metabasites. These mafic rocks constitute about 5% of

the sequence and occur as stocks and boudinaged

dykes. Small bodies (b100 m) are formed by strongly

foliated dark green to black amphibolites. Whereas

stock-like bodies of up to 600 m diameter have foli-

ated amphibolitic margins and undeformed gabbroic

cores. The undeformed gabbroic cores, although com-

pletely recrystallized during the metamorphic over-

print, still exhibit medium to coarse-grained igneous

holocrystalline textures. Amphibolites are dominated

by Barrovian-type metamorphism with a common

assemblage of albite+hornblende+zoisite/epido-

teFgarnet. The occurrence of sodic amphiboles indi-

cates HP/LT blueschist-facies metamorphism that

affected the whole Pan-African basement, most prob-

ably during the Alpine event. Due to a later greens-

chist-facies overprint, these amphiboles only occur as

partially replaced relics (Fig. 4). In addition to the

basic inclusions, metamorphosed metagranites, mostly

occurring as sills parallel to the foliation, were recog-

nized in the Pan-African schists. Distinct deformation

features, e.g. metaleucogranites cutting the foliation of

the enclosing schist, favor a post-Pan-African relative

age for the emplacement of the leucogranite protolith.

Considering the close resemblances to the Early Tri-

assic leucocratic orthogneisses of the Menderes Mas-

sif (Koralay et al., 2001), a possible Triassic age can

be postulated for their protoliths.

Fig. 2. Geological map of the central part of the Afyon Zone and distribution of the HP–LT mineral assemblages. Inset illustrates the tectonic

zones of Pondites and Anatolides within NW Turkey.

O. Candan et al. / Lithos 84 (2005) 102–124106

Fig. 3. Generalized stratigraphic section of the central part of the Afyon Zone shown in Fig. 2.

O. Candan et al. / Lithos 84 (2005) 102–124 107

The Pan-African basement is unconformably over-

lain by the Mesozoic cover sequence. The Early Tri-

assic to Early Tertiary units include thick sequences of

metaconglomerates, phyllites and metacarbonates.

Two types of conglomerates have been distinguished

in terms of clast and matrix: carbonate metaconglo-

merates and quartz metaconglomerates. The carbonate

metaconglomerates cover an area of 3 by 30 km, in

NW–SE direction between KVrka and Bayat. Carbon-

ate-cemented metaconglomerates with marble pebbles

containing fibrous carbonate crystals pass gradually

into quartz metaconglomerates. Here, quartz metacon-

glomerate horizons, up to 50 m thick, are interbedded

with grayish to black phyllites, whereas the metacon-

glomerates in PasadagV constitute the lowest level of

the Mesozoic sequence and are typified by thick (up

to 300 m), massive quartz metaconglomerates. Pure

pyrophyllite patches and horizons or pyrophyllite-rich

yellowish-gray phyllites commonly occur as interca-

lations within the quartz metaconglomerates. Rock-

forming Fe-Mg-carpholite occurrences are mostly re-

stricted to such Al-rich zones which, most probably,

O. Candan et al. / Lithos 84 (2005) 102–124108

are derived from kaolinitized rhyolitic tuffs. Two

types of carpholite occurrences are recognized: 1)

Rosette-like carpholite aggregates of up to 3 cm in

diameter, occurring as a widespread rock-forming

mineral, and associated with chloritoid (b3 mm)

and pyrophyllite, and 2) Quartz segregations with

carpholite fibres reaching lengths of up to 10 cm.

Metaconglomerates gradually pass into phyllite

quartz–phyllite quartzite intercalations that can reach

up to 3 km in thickness. This psammitic to pelitic

sequence contains Lower Scythian–Lower Anisian

marble horizons up to 100 m thick. In the phyllites,

chloritoid appears as a widespread rock-forming min-

eral, forming individual tabular or rosette-shaped

crystals up to 1 cm in diameter. The lithofacies of

the sedimentary protoliths of this sequence (mud-

stones, sandstones, conglomerates and limestones)

indicates continental to shallow-marine depositional

environments.

The phyllitic sequence is conformably overlain by

100 m thick, reddish-gray quartz-phyllites, phyllites,

calcareous phyllites and marbles, and finally by 2000

m thick dolomitic platform metacarbonates. In the

transition zone, metabasites are present as lenses and

discontinuous horizons. At PasadagV, sodic amphi-

bole–epidote assemblages were identified in these

rocks. The lowest levels of the platform-type meta-

carbonates, ~100 m thick, as well as the marble layers

close to the contact are made up of rosette-like co-

lumnar and/or fibrous calcite crystals that are pseudo-

morphs after aragonite. The main lithology of the

platform carbonates consists of gray, massive to

thick-bedded dolomites.

In the northern part of the study area, the Creta-

ceous platform carbonates pass gradually into Upper

Cretaceous, thin-bedded pelagic marbles with chert

layers, reflecting platform subsidence. These pelagic

carbonates are characterized by rosette-like columnar

carbonate crystals of up to 20 cm in length, which are

very common in the Lycian Nappes (Rimmele et al.,

2003a). Although for the Afyon Zone, the original

contact between the pelagic marbles and the overlying

Upper Maastrichtian–Lower Paleocene olistostromes

is generally interpreted as transitional (Ozcan et al.,

1989), in the northern part of our study area this

contact is represented by a ductile shear zone, up to

10 m thick, which developed during the internal im-

brication of the Afyon Zone. This olistostromal unit

with a very low-grade matrix consisting of flysch-type

mudstone and sandstone intercalations, contains huge

blocks of pelagic and neritic limestones, blueschist,

ophiolite and serpentinite. The blocks were most

probably derived from the TavsanlV Zone and from

obducted ophiolites. This blocky unit with its conti-

nental flysch-type matrix is tectonically overlain by a

tectonic melange of the TavsanlV Zone that was meta-

morphosed under blueschist-facies conditions. This

tectonic melange, described as a volcano–sedimentary

complex by Okay (1984), is exposed over large areas

along the northern border of the Afyon Zone. Non-

metamorphic, Upper Paleocene–Lower Eocene shal-

low-water sediments consisting of siltstones, marly

limestones and limestones, unconformably overlay

all the units.

4. Petrography of the Afyon zone

4.1. Pan-African basement

Three metamorphic facies can be identified in

the Pan-African basement units: i) Barrovian-type

MP metamorphism in metapelites, ii) HP / LT blue-

chist-facies metamorphism in metagabbros and

amphibolites, and iii) MP greenschist-facies overprint

in both metapelites and metabasites. Based on the

mineral assemblages three types of schist are distin-

guished; garnet–mica schist, albite–schist and phen-

gite–quartz schist (Table 1). The garnet–mica schists

consist of albite, quartz, phengite, garnet and biotite

(Fig. 4a). These metapelites underwent a strong ret-

rograde Alpine overprint and garnet and biotite were

mostly replaced by chlorite. The albite–schists are

characterized by albite porphyroblasts of up to 0.5

cm in diameter, and their mineral assemblage com-

prises albite, phengite and quartz. The phengite–

quartz schists form the least common rock type of

the basement. The abundance of phengite and quartz

vary continuously from pure quartzites to quartz-rich

phengite schists. All of these metapelites contain ev-

idence for multi-stage deformation during Pan-Afri-

can (S1) and Alpine (S2 and S3) events. The main

foliation (S1) developed under MP conditions, and is

defined by a parallel alignment of muscovite and

retrogressed biotite, and by syn-tectonic growth of

albite porphyroblasts with helicitic graphite inclu-

Table 1

Estimated modes of representative metapelites and metabasites from the Pan-African basement and the cover sequence

Pan-African basement Cover sequence

Schist Metabasites Phyllite Metabasite

Metagabbro Amphibolite

01-210 01-187 01-150 01-304 00-123/5 01-408 01-182 01-187 02-111 02-110

Quartz 18 24 3 8 21 11 26 36 4 2

Plagioclase 36 19 26 17 – – 14 9 28 16

Biotite – 2 14 – – – – – – –

Muscovite 24 32 8 – 14 5 34 26 – –

Pyrophyllite – – – – 35 62 5 9 – –

Chlorite 16 12 7 12 – – 4 6 26 36

Garnet 6 11 – – – – – – – –

Chloritoid – – – – 12 8 17 14 – –

Carpholite – – – – 18 14 – – – –

Sodic amphibole – – – 6 – – – – 4 6

Barroisite – – – 8 – – – – 10 4

Hornblende – – 21 – – – – – –

Actinolite – – – 28 – – – – 8 24

Epidote – – 16 21 – – – – 14 10

Zoisite – – 5 – – – – – 6 2

Additional accessories may be zircon, tourmaline and rutile.

O. Candan et al. / Lithos 84 (2005) 102–124 109

sions. S2 is characterized by a penetrative foliation

composed of medium- to coarse-grained white micas

that are oriented at approximately 908 to the (S1)

foliation. During this deformation period (S1), most

probably related to Alpine metamorphism, garnets and

biotites were strongly altered to chlorites. In addition,

in mica-rich layers a set of discontinuous shear bands

(S3), marked by fine-grained white mica and Fe-oxi-

des, developed. These shear bands can be attributed to

the late stages of exhumation of the Afyon Zone. They

are similar to those recognized throughout the cover

sequence and indicate top-to-the south movements.

Metabasic rocks have massive to weakly de-

formed, coarse-grained cores of flaser metagabbro,

and foliated margins mainly consisting of amphibo-

lite. Although the igneous assemblages in the cores of

the stocks are completely replaced by metamorphic

minerals, relic holocrystalline textures can still be

recognized. The mineral assemblage of the flaser

gabbros is Ca-amphibole, zoisite, epidote, garnet,

white mica and chlorite. During HP metamorphism,

igneous plagioclase was pseudomorphosed by aggre-

gates of zoisite. Amphibolitic rims of stocks as well as

small amphibolite lenses contain relic HP assem-

blages. These blueschist assemblages consist of ferro-

glaucophane, epidote, zoisite, rutile and quartz.

Ferroglaucophane is mostly retrograded to barroisite

and chlorite (Fig. 4b). Albite forms euhedral crystals

overgrowing the foliation during the greenschist-fa-

cies overprint.

4.2. Cover sequence

Both metasediments and metabasites of the cover

sequence contain one or more of the high-pressure

minerals carpholite, pyrophyllite, chloritoid, aragonite

pseudomorphs and sodic amphibole (Table 1). Among

these rocks, chloritoid-bearing phyllites dominate.

Metaconglomerates and phyllites containing carpholite

occur in subordinate amounts. Some marbles contain

fibrous carbonate crystals inferred to be pseudomorphs

after aragonite. Sodic amphibole is only present in lens-

shaped metabasites within the phyllites.

Chloritoid-bearing metapelites are characterized by

a marked schistosity and a typical phyllitic luster, this

colour varying from silverish grey to reddish pink.

The mineral assemblage of the phyllites is pyrophyl-

lite, phengite, chloritoid, quartz and chlorite. Chlo-

ritoid occurs as individual prismatic and rosette-shaped

syn-to post-tectonic crystals with inclusions of quartz.

In some samples, chloritoid in association with phen-

gite is replaced by chlorite due to isothermal decom-

Fig. 4. Photomicrographs of the mica schist and amphibolitic meta-

gabbro from the Pan-African basement (plane polarized light). (a)

Garnet–mica schist with muscovite (Ms), garnet (Gr) and biotite

(Bi) partially replaced by chlorite (Chl). (b) Ferroglaucophane-

bearing amphibolite from the rim of a metagabbro stock. Ferroglau-

cophane (Fe-gl) is replaced by barroisite (Ba). (Ep: Epidote).

O. Candan et al. / Lithos 84 (2005) 102–124110

pression during the retrograde overprint. This can be

inferred from the reaction chloritoid+celadonite com-

ponent in phengite+water reacts to muscovite com-

ponent in phengite+chlorite+quartz (Oberhansli et

al., 1995). Fe-Mg-carpholite occurs as rock-forming

minerals within the main foliation in pyrophyllite-rich

phyllites, forming thin layers up to 1 m and irregular

patches in the quartz metaconglomerates. These alu-

minous rock types contain the unaltered high-pressure

assemblage Fe-Mg-carpholite, chloritoid and pyro-

phyllite. In addition, phengite may be present in

minor amounts. Rosette-shaped carpholite (1–10 mm

in size) is abundant in phyllites and quartz metacon-

glomerates (Fig. 5a). In most cases, carpholite is

brown, stained by hematite inclusions. Carpholite

seems to be in equilibrium with chloritoid, because

no disequilibrium textures at contact surfaces have

been observed. In some cases, carpholite occurs in

close association with chlorite, indicating partial re-

placement of carpholite. This textural relation is inter-

preted as HP cooling during the retrograde evolution

(reaction 4 in Fig. 7). Generally, the carpholite-bear-

ing phyllites are affected by at least three phases of

ductile deformation. A planar internal foliation (S1)

related to a first deformation (D1) is preserved in

carpholite in the form of trails of solid inclusions,

which are oblique to the foliation of the matrix. A

second deformation (D2) produced the main foliation

(S2) of the phyllites. S2 wraps around the carpholite

porphyroblasts with the development of pressure sha-

dows (Fig. 5b). D3 deformation, related to the exhu-

mation of the HP rocks, is marked by shear-band

cleavages and lacks HP minerals. Two generations

of chloritoid crystals or continuous syn-tectonic

growth of chloritoid porphyroblasts are commonly

observed in chloritoid phyllites.

Metaconglomerates and phyllites contain numer-

ous carpholite-bearing quartz segregations that ap-

pear as synfolial lens-shaped veins, as described from

all carpholite localities along the Peri-Tethyan domain

from Oman to the Riff. In such veins, carpholite

appears as mesoscopic long green fibres (3–12 cm

long and 0.1–1 mm wide) in quartz segregations and

as hair-like microfibres or very thin needles in quartz

in some segregations (Fig 5c). The main assemblage

in quartz segregations is quartz, carpholite, chloritoid

and calcite. Small rosette-shaped chloritoid grows in

both quartz segregations and, especially, within the

surrounding of green-coloured chloritoid and quartz-

rich domains. Pyrophyllite is not present in quartz

segregations but occurs in the surrounding rocks.

Metacarbonates occurring as horizons in phyllites

and at the base of the platform sediments contain up to

3 cm long columnar or fibrous carbonate crystals. The

macroscopic habit and optical properties fit those of

aragonite, but X-ray powder diffraction indicates cal-

cite. However, based on the high Sr and low Mg

contents as measured by microprobe in these carbonate

minerals, we suggest that most of these fibers consist

of a submicroscopic mixture of calcite and strontianite

and indeed represent pseudomorphs after aragonite.

The metabasites occur as discontinuous lenses

within the transition zone composed of phyllites to

Fig. 5. Petrographic features of the metapelites and metabasites of the cover sequence (plane polarized light). (a) Rosette-shaped Fe-Mg-

carpholite (Car) crystals in phyllites, (b) main foliation (S2) of the phyllites wrapping around carpholite porphyroblasts. Inclusion patterns in

carpholite represent (S1) foliation. (c) Quartz segregation (Qtz) including green fibres of carpholite (Car), (d) Blueschist metabasite with sodic

amphibole (Na-amph) which is partially replaced by actinolite (Ac) and chlorite (Chl). (Prl: pyrophyllite, Cld: chloritoid).

O. Candan et al. / Lithos 84 (2005) 102–124 111

platform carbonates. They are dark green, strongly

foliated and dominated by a greenschist-facies assem-

blage consisting of chlorite, albite, epidote, actinolite

and sphene. In the least retrograded metabasites, sodic

amphibole, zoisite, rutile and quartz, relics of a former

HP assemblage, were recognized. Well-preserved blue

amphibole needles (20–200 Am) occur as inclusions in

albite porphyroblasts and quartz. The chlorite-rich

matrix shows all stages of retrograde alteration of

sodic amphibole to actinolite and finally chlorite

(Fig. 5d). Zoisite is rimmed by Fe-rich epidote. Co-

rona structures of sphene around rutile, characteristic

of retrograded HP metabasites, are common.

5. Mineral chemistry and P–T estimates

Mineral analyses were obtained using two different

electron microprobes: a Cameca SX100 at the Geo

Forschungs Zentrum (Potsdam, Germany), and a

JEOL 8800 in the Humboldt Museum (Berlin, Ger-

many), both operated under standard conditions (15

kV, 10–20 nA, PAP correction procedure), using nat-

ural and synthetic standard minerals (in Potsdam:

Fe2O3 [Fe], rhodonite [Mn], rutile [Ti], MgO [Mg],

wollastonite [Si, Ca], fluorite [F], orthoclase [Al, K],

albite [Na]; in Berlin: anorthoclase [Si, Al, Na, K],

ilmenite [Ti, Mn], magnetite [Fe], Cr-augite [Mg, Ca],

apatite [F], tugtupite [Cl]). The analytical spot diam-

eter was set between 3 and 5 Am, keeping the same

current conditions. Representative mineral analyses

from both the Pan-African basement and the cover

sequence are given in Table 2.

5.1. Fe-Mg-carpholite

The Fe-Mg-carpholite structural formula (Fe, Mn,

Mg)Al2Si2O6(OH, F)4 was calculated on the basis of 5

cations for the calculation of Si and 3 cations for Al,

Fe, Mn and Mg, in order to account for the contribu-

Table 2

Representative mineral chemistry of garnet, sodic amphibole, chloritoid, carpholite and aragonite pseudomorphs from the Pan-African basement and the cover sequence

Pan-African basement Cover sequence

Garnet Sodic amphibole Barroisite Carpholite Chloritoid Aragonite pseudomorphs

Core Rim

01-210/1 01-210/2 01-304/3 01-304/5 01-304/4 123/5-1 123/5-3 123/5-14 123/5-20 123/6-5 123/5-3 123/5-8 123/5-9 124/1 124/3 124/4

SiO2 35.51 35.47 54.34 52.13 45.88 SiO2 38.84 39.21 38.94 38.98 24.63 26.36 24.50 24.86 MgCO3 0.0000 0.0000 0.0000

TiO2 0.00 0.08 0.04 0.12 0.27 TiO2 – – – – 0.00 0.00 0.02 0.03 Na2CO3 0.0000 0.0000 0.0399

Al2O3 21.58 21.53 10.09 10.66 9.94 Al2O3 33.81 33.73 34.15 33.87 42.78 42.47 43.27 43.78 CaCO3 99.6629 99.7505 99.5522

FeO 33.67 33.80 19.95 20.59 24.35 FeO 8.93 8.97 8.77 9.20 22.93 22.37 24.15 22.26 MnCO3 0.0343 0.0000 0.0000

MnO 2.02 2.41 0.23 0.27 0.41 MnO 0.05 0.06 0.04 0.06 0.28 0.24 0.09 0.20 SrCO3 0.3028 0.2047 0.4079

MgO 1.57 1.53 5.46 5.19 5.36 MgO 7.87 7.72 7.91 7.51 3.04 2.67 1.75 3.36 BaCO3 0.0000 0.0000 0.0000

CaO 4.82 4.50 1.56 2.57 6.51 CaO – – – – – – – – FeCO3 0.0000 0.0448 0.0000

Na2O 6.82 6.69 4.62 Na2O – – – – 0.01 0.00 0.00 0.01 Total 100 100 100

K2O 0.10 0.19 0.45 K2O – – – – 0.01 0.02

Cr2O3 0.00 0.01 0.04 0.01 0.02 F 0.04 0.05 0.11 0.23 – –

Total 99.17 99.33 98.63 98.41 97.81 Total 89.52 89.73 89.93 89.84 93.68 94.13

Si 7.269 6.293 7.737 7.532 6.894 Si 1.983 2.000 1.980 1.992 1.987 2.168

Ti 0.000 0.011 0.004 0.013 0.030 Ti – – – – 0.000 0.000

Al 5.208 4.503 1.693 1.816 1.761 Al 2.023 2.028 2.032 2.035 4.068 4.117

Fe 5.764 5.015 2.375 2.487 3.060 Fe3+ – – – – – –

Mn 0.349 0.362 0.028 0.033 0.052 Fe2+ 0.379 0.383 0.371 0.392 1.547 1.539

Mg 0.480 0.404 1.158 1.118 1.199 Mn 0.002 0.003 0.002 0.002 0.019 0.017

Ca 1.056 0.856 0.238 0.397 1.048 Mg 0.595 0.587 0.596 0.571 0.183 0.160

Na 0.000 0.000 1.883 1.873 1.346 Ca – – – – – –

K 0.000 0.000 0.018 0.035 0.086 Na – – – – 0.001 0.000

Cr 0.000 0.001 0.005 0.001 0.003 K – – – – 0.001 0.002

F 0.006 0.008 0.018 0.037 – –

Alm 68.6 73.0

Prp 8.0 6.7 XMg 0.61 0.60 0.62 0.59 0.10 0.09

Sps 5.8 6.0 XFe 0.39 0.39 0.38 0.41

Grs 13.4 12.5 XMn 0.00 0.00 0.00 0.00

And 4.2 1.8

Uva 0.0 0.0

O.Candanet

al./Lith

os84(2005)102–124

112

O. Candan et al. / Lithos 84 (2005) 102–124 113

tion of surrounding quartz when analysing fibres

smaller than the microprobe beam diameter (Goffe

and Oberhansli, 1992). Analyses showing an oxide

sum lower than 85 wt.% or greater than 90 wt.% were

rejected. The Fe3+ (Fe3+=2�Al) and Fe2+ contents

were calculated after Goffe and Oberhansli (1992).

Three phyllite samples (123-6, 123-5 and ro997)

were analysed, in which carpholite occurs as a rock-

forming mineral in the assemblage Fe-Mg-carpho-

lite–chloritoid–pyrophyllite–phengite. The composi-

tion of Fe-Mg-carpholite varies: XMg ranges from

0.37 to 0.64 (Fig. 6a). The Mn contents are very

low in all samples, and XMn ranges from 0.000 to

Fig. 6. Mg–Fe–Mn ternary diagrams showing the compositions of Fe–M

0.006. As compared to carpholite analyses reported

from the Central Alps (3–4.6 wt.%; Goffe and Ober-

hansli, 1992) and Betic Cordillera (max. 2.3 wt.%;

Azanon and Goffe, 1997), the fluorine content of

carpholite is generally very low (b0.05) but can

reach 0.23 wt.%. No chemical zoning was detected

in carpholite.

5.2. Chloritoid

The chloritoid formula (Fe, Mn, Mg)2Al4-Si2O10(OH)4 was calculated on the basis of 12 oxy-

gens. Fe3+ /Fe2+ (Fe3+=4�Al) was calculated after

g–carpholite (a) and chloritoid (b) of the Afyon metasediments.

O. Candan et al. / Lithos 84 (2005) 102–124114

Chopin et al. (1992). Analyses showing an oxide sum

lower than 90 wt.% or greater than 94 wt.% were

rejected. This mineral is richer in iron than carpholite

and shows a very limited compositional spread. The

XMg values of chloritoid from four samples range from

0.09 to 0.22 and, like Fe-Mg-carpholite, chloritoid

shows a very low manganese content ranging from

0.09–0.29 wt.% (Fig. 6b). Similar compositions of Fe-

Mg-carpholite and chloritoid have been recently

reported from the basal metasediments of the Lycian

Nappes (Oberhansli et al., 2001; Rimmele et al.,

2003a).

5.3. Aragonite pseudomorphs

Aragonite pseudomorphs can be recognized in the

field by their macroscopic habit. However, X-ray

powder diffraction studies yield characteristic pat-

terns of calcite, indicating that most of the fibrous

carbonates recrystallized to calcite during a retro-

grade overprint. In order to determine whether

these calcites are pseudomorphs after aragonite,

trace element contents were measured on fibrous

crystals. The Sr content of some crystals is very

high and ranges from 1630 to 4080 ppm, which is

consistent with aragonites described from the West-

ern Alps (Gillet and Goffe, 1988) and Betic Cordil-

lera (Azanon and Goffe, 1997). The Ba content is

low and ranges from 20 to 260 ppm. Mg has not

been detected in these crystals.

5.4. Sodic amphibole

Amphiboles from one sample from an amphibolitic

margin of a metagabbro occurring in the Pan-African

basement were analysed. The amphibole analyses

have been calculated on the basis of 23 oxygens

and 13 cations. The compositions of amphibole

correspond to sodic and sodic–calcic amphibole

after Leake et al. (1997). The maximum Ca and

Na occupancies in the M4 structural site of sodic

amphiboles are 1.624 and 0.664, respectively. The

compositional range of the sodic amphiboles is very

narrow and is restricted to the field of ferro-glauco-

phane. A representative sodic amphibole formula is

(Na0.14 K0.01)(Na1.90 Ca0.10)(Mn0.02Fe2+2.01 Mg1.11

Fe3+0.23 Al1.63)(Al0.12 Si7.88) O22 (OH)2. Amphiboles

replacing ferro-glaucophane belong to the sodic–cal-

cic amphiboles (Leake et al., 1997) and are barroi-

sitic and katophoritic.

5.5. Garnet

Garnet analyses from the garnet mica-schists (01-

210) of the Pan-African basement show distinct chem-

ical zoning (representative analyses of core: Alm 66.3

Prp 4.8 Sps 16.8 Grs 11.5 And 0.6 and rim: Alm 73.0

Prp 6.7 Sps 6.0 Grs 12.5 And 1.8) and consist pre-

dominantly of almandine, spessartine and grossular

components. The cores of all garnets are enriched in

Mn relative to Fe, and the rims are enriched in Fe and

depleted in Mn. The almandine content increases

regularly toward the rim from 66.3 to 73.0 mol%,

while the spessartine content decreases from 16.8 to

3.7 mol%. The amount of the pyrope component is

generally lower than 5 mol% in the core and increases

slightly toward the rim. The grossular component is

fairly constant within the individual crystals and

increases from 10.6 to 12.5 mol% toward the rim.

5.6. Pressure–temperature estimates

P–T conditions were estimated by calculation of

phase equilibria using PTAX (Berman and Perkins,

1987) with the internally consistent database of Ber-

man (1988). The thermodynamic data for Fe-Mg-car-

pholite and the activity model for the Mg end-member

are from Vidal et al. (1992); a(car)= (XMg)(XAl)2(XOH)

4

with XMg=Mg/ (Mg+Fe2++Mn), XAl= (2�Fe3+) /2

and XOH=((4�F) /4). The thermodynamic data used

for Mg-chloritoid are provisional. They were estimat-

ed by Patrick and Berman (unpublished data, 1989)

and have previously been used in many studies (Ober-

hansli et al., 1995; Jolivet et al., 1996; Azanon and

Goffe, 1997; Goffe and Bousquet, 1997; Bousquet et

al., 2002). The thermodynamic dataset is listed in

Vidal et al. (1999). Mg-chloritoid activities were cal-

culated after Theye et al. (1992); a(cld)=XMg, with

XMg=Mg/ (Mg+Fe2++Mn). Because the three sam-

ples analysed were chlorite-free, metamorphic condi-

tions were calculated using activities of a theoretical

chlorite composition deduced from both Fe-Mg-car-

pholite—chlorite and chloritoid—chlorite partitioning

(KD[car / chl] and KD[cld / chl]) in carpholite- and

chloritoid-bearing rocks of Crete and the Peloponnese

(Theye et al., 1992). These authors report KD[car / chl]

O. Candan et al. / Lithos 84 (2005) 102–124 115

values between 1.1 and 1.4 and KD[cld / chl] values

ranging from 6.2 to 8.5. We also used the log(Mg/

Fe)car / log(Mg /Fe)chl and log(Mg/Fe)cld / log(Mg/

Fe)chl graphs from Azanon and Goffe (1997) in

order to estimate theoretical XMg values of chlorite.

The equilibrium curves involving carpholite, chlori-

toid and chlorite for sample 123-5 is shown in Fig. 7.

PT conditions were estimated at 6–9 kbar at tempera-

tures of about 350 8C. This domain is located in the

pyrophyllite stability field and shows PT conditions

roughly similar to those from the Lycian Nappes

(Oberhansli et al., 2001).

200 250 300 3

Pre

ssur

e (G

Pa)

Temper

2.0

1.5

1.0

0.5

XMg = 0.0

1: Chl + 4Kln =

2: 2Qz + Kln =

3: 4Prl + Chl =

4: 5Car + 9Qz =

5: Car =

6: Prl =

7: 2Ms + Chl + 2Qz =

8: calcite =

52

14

carpholi

Fig. 7. PT conditions for the Fe–Mg–carpholite–chloritoid assemblage (

Oberhansli et al., 2003). Equilibrium curves are calculated with theoretica

chlorite partitioning values and from the Fe–Mg–carpholite/chlorite partitio

similar HP assemblages occurring in Crete and Peloponnese (for explanati

the observed mineral assemblages. Additional data: aragonite–calcite transi

Car: carpholite, Qtz: quartz, Ky: kyanite, Prl: pyrophyllite, Cld: chloritoid

6. Discussion

6.1. Comparison of the Afyon Zone with other tectonic

units of the Anatolides

The main tectonic units of the Anatolides, i.e. the

TavsanlV Zone, Afyon Zone, Menderes Massif and

Lycian Nappes, are regarded to have been derived

from the Anatolide–Tauride platform during the clo-

sure of Neo-Tethys by a continuous process of intra-

oceanic subduction, obduction of oceanic lithosphere,

accretion and, finally, continental collision (Okay et

50 400 450 500ature (°C)

chloritoid

Qz + 5Car + 2W

Prl + W

5Ctd + 14Qz + 3W

4Prl + Chl + 2W

Qz + Ctd + W

Ky + 3Qz + W

3Ctd + 2ACel + W

aragonite

5

3.30

3.20

3.10

3.05

7

6

XMg = 0.4 XMg

= 0.6 XMg

= 0.7

3

te

8

sample 123-5) from the metasediments of the Afyon Zone (after

l chlorite compositions, which were estimated from the chloritoid/

ning values (black dashed lines) reported in Theye et al. (1992) for

on see text). Reaction 7 is contoured for Si contents of phengite for

tion curve (Johannes and Puhan, 1971). (Kln: kaolinite, Chl: chlorite,

, W: water).

O. Candan et al. / Lithos 84 (2005) 102–124116

al., 1996, 2001). Fig. 8 shows a present-day N–S

section through the Mesozoic–Early Tertiary succes-

sions for these tectonic zones. They have striking

similarities, not only in terms of stratigraphy but

also distribution of HP assemblages, as characterised

by Fe-Mg-carpholite. Carpholite occurrences are

clearly restricted to the Al-rich Triassic continental

to shallow-water clastic metasediments in the tectonic

zones of the Anatolides, the Lycian Nappes (Oberhan-

sli et al., 2001; Rimmele et al., 2003a), the Afyon

Zone (Candan et al., 2002) and the Menderes Massif

(Rimmele et al., 2003b). However, no carpholite has

been reported from the clastic metasediments under-

lying the Early Mesozoic platform-type carbonates in

the TavsanlV Zone.

The most striking aspects of the evolution of the

platform are its subsidence and the ophiolite obduc-

tion. These events are related to the initiation of the

closure of the Neo-Tethyan ocean along a north-

ward-dipping, intra-oceanic subduction zone in

Late Cretaceous time (Sengor and YVlmaz, 1981;

Okay et al., 1996, 2001; Goncuoglu et al., 2000).

This event is marked by a transition of neritic to

pelagic carbonate sedimentation onto the Anatolide–

Tauride platform. Platform subsidence was followed

by deposition of the flysch-type clastic sediments

with huge carbonate and ophiolite blocks, docu-

menting the propagation of the obducted ophiolites

and internal imbrication of the platform sequence.

The original positions of the major tectonic zones of

the Anatolide–Tauride platform prior to internal im-

brication from north to south were: the TavsanlV

Zone, Lycian Nappes, Afyon Zone and the Menderes

Massif (Fig. 9).

In the TavsanlV Zone, the stratigraphic succession

starts with Upper Paleozoic–Lower Triassic jadeite-,

glaucophane- and chloritoid-bearing phyllitic series.

Lower Triassic to Albian sediments are represented by

massive, platform metacarbonates. The deposition of

the pelagic metacarbonates and conformably overly-

ing metashale–metachert–metabasite intercalations

was initiated during the Cenomanian (Okay, 1986).

This lithofacies succession suggests a northernmost

position on the Anatolide– Tauride platform. Olistos-

trome deposits are absent in the TavsanlV Zone (Okay

et al., 1996).

The allochthonous Lycian units occur as unrooted

nappe stack over a large area between the Menderes

Massif and the Beydaglar autochthon (Collins and

Robertson, 1997, 1998). Bernoulli et al. (1974) and

Cakmakoglu (1985) studied in detail the stratigraphy

of the western Lycian Nappes between Gulluk and

Mugla, where they form part of the Lycian thrust

sheets (Collins and Robertson, 1997). The lowest

levels of the succession are dominated by chloritoid-

bearing, reddish phyllites with spectacular Fe-Mg-

carpholite fibres in quartz segregations (Oberhansli

et al., 1998a, 2001; Rimmele et al., 2003a). The

Middle Triassic to Middle Liassic sequence is charac-

terized by shallow-water limestones and dolomites.

These carbonates are overlain by a thick sequence

of pelagic and turbiditic limestones with cherts, rang-

ing from Upper Liassic to Cenomanian in age. These

lithologies mark a distinct change in depositional

environment and the installation of a pelagic regime

over a wide area. At first sight, the timing of the

transition from a neritic to pelagic regime is not

identical to that in the TavsanlV Zone, the Afyon

Zone and the Menderes Massif (Fig. 9). Late Turo-

nian- Early Senonian flysch deposits grade upwards

into wildflysch (Karabogurtlen unit) with limestone

(not younger than Maastrichtian; Bernoulli et al.,

1974) and blueschist blocks, the latter most probably

derived from the TavsanlV Zone. Taking this into

consideration, a possible original setting for the Ly-

cian Nappes between the TavsanlV and the Afyon

Zones can be envisaged (Fig. 9).

Close stratigraphic similarities between the Afyon

Zone and Menderes Massif have been reported (Okay

et al., 1996; Goncuoglu et al., 1996). However, com-

pared to the Menderes Massif, the transition from a

neritic to pelagic regime in the Afyon Zone occurred

slightly earlier. This suggests that the original paleo-

geographic position of the Afyon Zone was north of

the Menderes Massif. It must be emphasized that the

time of transition from a neritic to pelagic environ-

ment, as represented by cherty limestones and pelagic

micrites, is Lower Maastrichtian. These sequences are

overlain by an olistostromal unit of Upper Maastrich-

tian–Lower Paleocene age, containing huge blocks of

neritic and pelagic limestones, serpentinites and blues-

chists, embedded in matrix of unmetamorphosed sand-

stone, shale and turbidite (Goncuoglu et al., 1992).

In the Menderes Massif, chloritoid phyllites with

quartz metaconglomerate horizons containing the Fe-

Mg-carpholite–chloritoid–kyanite assemblage (Rim-

Fig. 8. Generalized stratigraphic sections of the tectonic zones of the Anatolides and the locations of HP/LT minerals. The thicknesses of the units are very approximate. Data from (*)

Okay (1984, 1986), Okay and Kelley (1994); (**) This study, Ozcan et al. (1989), Goncuoglu et al. (1992); (***), Durr (1975), Konak et al. (1987), Candan et al. (1997), Ozer (1998),

Ozer et al. (2001); (****) Bernoulli et al. (1974), Camakoglu (1985).

O.Candanet

al./Lith

os84(2005)102–124

117

Fig. 9. Paleogeographic reconstruction showing a possible original position of the major tectonic zones of the Anatolides as deduced from the

continuous southward-younging of the pelagic carbonates and olistostrome deposits and the ages of HP metamorphisms. Data from (*) Okay

(1984,1986), Okay and Kelley (1994); (**) Bernoulli et al. (1974), Camakoglu (1985); (***) Ozcan et al. (1989), Goncuoglu et al. (1992);

(****) Durr (1975), Konak et al. (1987), Ozer (1998), Ozer et al. (2001); time scale by Gradstein and Ogg (1996).

O. Candan et al. / Lithos 84 (2005) 102–124118

mele et al., 2003b) constitute the lowest units of the

Mesozoic sequence. These clastic metasediments are

conformably overlain by Late Triassic–Cretaceous

metabauxite-bearing platform metacarbonates. The

Cretaceous to Early Tertiary stratigraphy of the Men-

deres Massif is paleontologically well dated by Durr

(1975); Konak et al. (1987) and Ozer (1998). The

Upper Cretaceous sequence of the Menderes Massif

is dominated by platform carbonates. It is typified by

well-preserved rudist fossils in two horizons of mid-

dle-late Cenomanian and Santonian–Campanian age

(Ozer, 1998). The foundering from platform to slope

and afterwards to basin conditions in the Menderes

Massif starts in the Late Campanian–Late Maastrich-

tian and is represented by reddish pelagic marbles

with marly–pelitic interlayers. The marbles with pe-

lagic precursors are overlain by an olistostrome with

marble and serpentinite blocks. This blocky unit,

possibly deposited in front of the obducted ophiolite

slab, is dated as Middle Paleocene near Milas and

Lower–Middle Eocene near Cal (Ozer et al., 2001).

6.2. Tectonic implications

Detailed studies of the Late Cretaceous–Early Ter-

tiary evolution of the Neo-Tethyan ocean and its

implications for Western Turkey have been carried

out by Sengor and Ylmaz (1981), Okay et al. (1996,

2001), Collins and Robertson (1997) and Goncuoglu

et al. (2000). However, most of these papers focused

either on the high-pressure metamorphism in the Tav-

sanl Zone (Okay et al., 1996; Sherlock et al., 1999;

Okay et al., 1998) or the southward translation of the

Lycian Nappes (Collins and Robertson, 1997, 1998,

O. Candan et al. / Lithos 84 (2005) 102–124 119

1999). Recent discoveries of Cretaceous to Early

Tertiary (?) high-pressure imprints characterized by

Fe-Mg-carpholite in the Lycian Nappes (Oberhansli

et al., 1998a, 2001; Rimmele et al., 2003a), the Afyon

Zone (Candan et al., 2002) and the Menderes Massif

(Rimmele et al., 2003b), necessitate a review of these

tectonic models. Although detailed radiometric dating

of the high-pressure events and/or the prograde and

retrograde P–T histories of the different tectonic units

have not yet been completed, the tectono–metamor-

phic evolution of the Afyon Zone, as well as the other

zones of the Anatolides as deduced from this HP

evidence can be suggested as follows.

An Albian age has been assumed for the initiation

of closure of the Neo-Tethyan ocean along a north-

ward-dipping intra-oceanic subduction zone (Sengor

and YVlmaz, 1981; Okay et al., 1998). The subduction

of most of the oceanic lithosphere must have been

completed in Albian–Turonian times. The Cenoma-

nian pelagic marbles and overlying metabasic–meta-

chert–metashale intercalations in the TavsanlV Zone

mark the subsidence of the Anatolide–Tauride plat-

form as it approached the subduction zone. Following

complete consumption of the oceanic lithosphere, the

northern passive continental margin of the Anatolide–

Tauride platform was subducted beneath an oceanic

mantle wedge and the underlying accretionary prism

(Okay et al., 1998). P–T estimates of the peak meta-

morphic conditions, 24F3 kbar and 430F30 8C(Okay, 2002), and Ar–Ar mica ages of 80F0.5 Ma

from the TavsanlV Zone indicate subduction of the

platform to depths of about 60 km and HP/LT meta-

morphism during the Campanian (Okay and Kelley,

1994; Sherlock et al., 1999).

While neritic carbonate deposition continued in the

Afyon Zone and the Menderes Massif during the Late

Turonian–Early Cenonian, flysch deposition in front

of the obducted ophiolite slab suggests a more north-

erly original position of the Lycian domain. The wild-

flysch of the Karabogurtlen unit containing blueschist

and peridotite blocks, formed during Maastrichtian–

(?) Early Tertiary. The presence of blueschist blocks

point to at least partial exhumation of the TavsanlV

Zone. A mechanism of coeval subduction and exhu-

mation by underplating (Sherlock et al., 1999) during

the Campanian (Okay et al., 1998) is suggested. In

this model, the preservation of the blueschists during

the exhumation is attributed to the underplated cold

continental material. Considering the original position

of the Lycian domain, these continental deposits must

have formed the Permo(?)–Triassic to Early Creta-

ceous passive continental margin sediments of the

Lycian Nappes. In the basal metapelites of the Kara-

ova unit of the Western Lycian Nappes, minimum

pressures of c. 8 kbar, corresponding to a lithostatic

load of 30 km of oceanic mantle and underlying

continental sediments, and maximum temperatures

of c. 400 8C, indicate blueschist-facies metamorphism

(Oberhansli et al., 2001; Rimmele et al., 2003a).

Similar to the situation in Oman (Michard et al.,

1994), a tectonic setting in which passive continental

margin sediments were imbricated while obduction of

oceanic lithosphere was going on, is suggested for the

Lycian Nappes (Oberhansli et al., 2001). Considering

the timing of exhumation of the TavsanlV Zone, the

arrival of the ophiolites and the deposition of the

olistostromes in the Lycian domain and the Afyon

Zone, a Latest Maastrichtian to Early Tertiary (Early

Paleocene?) age bracket can be postulated for the

high-pressure metamorphism. The high-pressure

metamorphism and internal imbrication suggest that

the main deformation of the Anatolide–Tauride plat-

form was probably initiated in Maastrichtian times.

The onset of the pelagic depositional environment

in the Afyon Zone started in the Early Maastrichtian

and was followed by the deposition of a Lower Pa-

leocene olistostrome in the front of the advancing

ophiolitic Nappes (Goncuoglu et al., 1992). The oc-

currence of carpholite-bearing assemblages in Early

Triassic metaclastics, in combination with P–T esti-

mates, show that the internal imbrication of the plat-

form continued during the Early Paleocene, and that

some of the tectonic slices of the Afyon Zone were

buried to depths of c. 35 km and underwent blues-

chist-facies metamorphism. Furthermore, in contrast

to the other tectonic zones, the presence of glauco-

phane in the Pan-African basement implies that the

crystalline basement of the Afyon Zone was buried

deeply and was metamorphosed during the same high-

pressure event. The absence of any HP evidence in the

matrix of the olistostrome and the presence of ductile

shear zones between olistostrome and underlying

metacarbonates implies that the uppermost units of

the Afyon Zone were not buried to deep enough to

cause the same HP/LT metamorphism. During exhu-

mation, aragonite-bearing carbonates were juxtaposed

O. Candan et al. / Lithos 84 (2005) 102–124120

with the olistostromes along ductile shear zones. Pa-

leontological evidence obtained from the olistostrome

(Lower Paleocene) and unmetamorphosed marine

sediments (Upper Paleocene) that unconformably

overlie the Afyon metamorphics as well as the ophio-

lites, constrain the timing of ophiolite emplacement,

burial of the continental material, high-pressure meta-

morphism and exhumation along a cool retrograde P–

T path allowing the preservation of carpholite, to the

interval Lower Paleocene–Upper Paleocene. Similar

to the Lycian Nappes (Oberhansli et al., 2001) a

continuous process with obduction, southward propa-

gation of the oceanic lithosphere, deformation and

internal imbrication of the Anatolide–Tauride plat-

form must be envisaged for the Afyon Zone.

Different ages ranging from Paleocene to Early

Eocene, are suggested for the continental collision

between the Anatolide–Tauride platform and the Pon-

tides (Sengor and YVlmaz, 1981; Goncuoglu et al.,

2000; Okay et al., 2001). However, in the Menderes

Massif this time interval is still marked by olistos-

trome deposition occurring in the uppermost unit of

the sequence (Gutnic et al., 1979; Ozer et al., 2001).

Along the northwest border of the Menderes Massif,

between Dilek Peninsula and Akhisar, tectonic rem-

nants of the Cycladic Complex are sandwiched be-

tween the Lycian Nappes and the Menderes Massif.

There, high-pressure relics (eclogites and blueschists)

in Mesozoic–Early Tertiary series were fist recognized

by Candan et al. (1997) and correlated with the

Cycladic Complex in terms of close lithostratigraphi-

cal similarities (Candan et al., 1997; Cetinkaplan et

al., 2000; Okay at al., 2001) and the age of the high-

pressure metamorphism, dated at 40F0.5 Ma by Ar /

Ar method (Oberhansli et al., 1998b). This Cycladic

sequence consists of two tectonic units: i) Triassic to

Upper Cretaceous passive continental margin sedi-

ments with blueschist relics and ii) a metaolistos-

trome, probably Early Tertiary in age, with

completely fresh eclogites and other HP blocks (Can-

dan and Cetinkaplan, 2000). Below this Cycladic

Complex, new evidence of high-pressure relics from

the Mesozoic cover series along the southern flank of

the Menderes Massif has been discovered recently

(Rimmele et al., 2003b). There, carpholite-bearing

assemblages (magnesiocarpholite–kyanite; 10–12

kbar /440 8C) occur in Triassic metaconglomerates.

This evidence shows that during continental collision

the Mesozoic–Early Tertiary cover of the Menderes

Massif suffered internal imbrication and burial to

minimum depths of about 30 km during the Middle

to Late Eocene. In contrast to the Afyon Zone, there is

no evidence that the Pan-African basement was deep-

ly buried during the Alpine event. The widespread

eclogite occurrences in the Pan-African basement of

the Menderes Massif are attributed to the Pan-African

orogeny (Oberhansli et al., 1997; Candan et al., 2001;

Oberhansli et al., 2002). For a long time it has been

emphasized that the age of the Barrovian overprint of

the Pan-African basement and the cover series is

consistent with the time of emplacement of the Lycian

Nappes (Sengor and YVlmaz, 1981; Sengor et al.,

1984; Dora et al., 1995). Rb /Sr and Ar /Ar cooling

ages of micas from the Pan-African basement and

Paleozoic cover units fall in the range of 50–27 Ma

with an average of 37F1 Ma (Satr and Friedrichsen,

1986) and 43–37 Ma (Hetzel and Reischmann, 1996)

for this medium temperature event. During the Late

Eocene– Late Oligocene, no sediments were deposit-

ed in western Anatolia, indicating that the region was

a subaerial landmass, possibly as a consequence of the

continuing convergence (YVlmaz et al., 2000). The

exhumation mechanism and its timing for the Men-

deres Massif are widely debated. Extensional exhu-

mation starting in Early Miocene time along low-

angle normal faults and a typical core-complex for-

mation was suggested for the Menderes Massif by

Bozkurt and Park (1994) and Bozkurt (2004). This

neglects the facts that i) a Pan-African intrusion age of

the protoliths of the augen gneisses (ca. 550 Ma) has

been proven by several methods (Hetzel and Reisch-

mann, 1996; Loos and Reischmann, 1999; Gessner et

al., 2001, 2004; Koralay et al., 2004) and ii) evidence

of a transition from ductile to brittle deformation

between the augen gneisses and schists along the

southern flank of the Menderes Massif is absent as

emphasized by Hetzel and Reischmann (1996). More-

over, the fluvial and lacustrine sediments, which form

the oldest sediments unconformably covering the

Menderes Massif, reveal that the uplift and exhuma-

tion of the Menderes Massif occurred before the Early

Miocene. Thus, an exhumation mechanism along

the low-angle thrust faults, back thrusts and asso-

ciated normal faults during Eocene–Oligocene time

was suggested by YVlmaz et al. (2000). The con-

vergent deformation continued until the Early Mio-

O. Candan et al. / Lithos 84 (2005) 102–124 121

cene in the external parts of the Taurides. It is

suggested that during Burdigalian–Seravallian time,

the Lycian Nappes and the Menderes Massif were

emplaced onto the Bey Daglar autochthon along an

intracrustal thrust zone (Sengor et al., 1984).

7. Conclusions

The widespread occurrence of the assemblage bFe-Mg-carpholite–chloritoid–pyrophylliteQ in Lower Tri-

assic quartz metaconglomerates and phyllites of the

Afyon Zone shows that these units were buried to a

depth of about 35 km and suffered HP/LT metamor-

phism under blueschist-facies conditions. Estimated

P/T conditions for these rocks give 350 8C and 6–9

kbar. The original unconformable contact and the

sodic amphibole relics in the amphibolitic metagab-

bros occurring in the underlying Pan-African base-

ment of the Afyon Zone suggest that the crystalline

basement also underwent this burial. During the con-

tinuous processes of the consumption of the Neo-

Tethyan ocean, a tectonic scenario similar to the Ly-

cian Nappes suggested by Oberhansli et al. (2001),

internal imbrication of the Anatolide–Tauride plat-

form and burial of the tectonic slices beneath the

obducted oceanic lithosphere, can be envisaged for

the Afyon Zone. Petrological and regional evidence

favor a syn-orogenic exhumation with a cool retro-

grade P–T path. The excellent preservation of HP

mineral assemblages, including carpholite, aragonite

and sodic amphibole, implies that the thermal gradient

remained cool during a significant part of the exhu-

mation process. Such conditions with a cold geother-

mal gradient can only be preserved, if cold material is

continuously subducted while exhumation in a tecton-

ic accretion channel (TAC) is proceeding. This corre-

sponds to a continental equivalent of a slab roll back

during subduction. The timing of the processes of the

ophiolite emplacement, internal imbrication of the

platform, burial of the continental material, high-pres-

sure metamorphism and exhumation can be con-

strained by fossil evidence from the olistostrome

(Lower Paleocene) and the oldest unmetamorphosed

marine sediments (Upper Paleocene) unconformably

covering the Afyon metamorphics as well as the

ophiolite nappes, to a Lower Paleocene–Upper Paleo-

cene time interval.

The ages of the pelagic carbonates and subse-

quently following olistostrome deposits change pro-

gressively from Cenomanian in the TavsanlV Zone in

the north to Early Middle Eocene in the Menderes

Massif in the south. This reflects progressive south-

ward implication of the Anatolide–Tauride platform

in subduction processes. In addition, continuing

compression during subduction and later continent–

continent collision caused continuous southward-

younging of high-pressure metamorphisms in the

major tectonic zones of the Anatolides, from Cam-

panian in the TavsanlV Zone to Middle Late Eocene

in the Menderes Massif.

Acknowledgements

We would like to thank O.O. Dora for his contin-

uous support of our work. This cooperation was and is

substantially supported by TUBITAK (grant

101Y022), the Volkswagen Stiftung, the German Sci-

ence Foundation, DFG (grant OB 80/22), the German

academic exchange organization, DAAD, and the

bDeutsch-Franzosische HochschuleQ. The support of

all these granting agencies is warmly acknowledged.

We thank R. Bousquet, A. Okay and P. O’Brien for

critically reading the manuscript.

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