Petrology, geochemistry, and evolution of the iron skarns along the northern contact of the Eğrigöz Plutonic Complex, Western Anatolia, Turkey

Turkish Journal of Earth Sciences
/>
Research Article

Turkish J Earth Sci
(2013) 22: 61-97
© TÜBİTAK
doi:10.3906/yer-1006-2

Petrology, geochemistry, and evolution of the iron skarns along the northern contact of
the Eğrigöz Plutonic Complex, Western Anatolia, Turkey
Tolga OYMAN*, İsmet ÖZGENÇ, Murat TOKCAER, Mehmet AKBULUT
Department of Geological Engineering, Faculty of Engineering, Dokuz Eylül University, Tınaztepe, Buca, TR−35100 İzmir, Turkey
Received: 01.06.2010

Accepted: 30.06.2011

Published Online: 04.01.2013

Printed: 25.01.2013

Abstract: The Çatak and Küreci skarn districts are located approximately 10 km NW of Emet (Kütahya) in Western Turkey. The
skarn and associated ore formations mainly occur at the contact between intrusive rocks of the Eğrigöz Plutonic Complex (EPC) and
calcareous pelitic schists with limestone lenses of the Sarıcasu Formation and meta-carbonate rocks of the Arıkaya Formation. The major,
trace, and rare earth element analysis of the igneous rocks indicate that they are high level, subalkaline, calc-alkaline, peraluminous to
metaluminous I-type intrusions, generated in a continental arc setting. Three distinct skarn-type mineralization, differing in their host
rocks and distance from the intrusive body, were chosen to establish the ore-forming conditions in different episodes of skarn formation.
The Küreci iron mineralization is hosted in a skarn zone with well-developed zoning from unaltered granodiorite and endoskarn,
andradite-diopside exoskarn, to diopside-wollastonite exoskarn towards a marble reaction front. In Sakari, the iron mineralization
and associated skarn have formed due to successive fracturing and infiltration processes. From early contact metamorphic rocks to
late prograde skarn at the Sakari prospect, the composition of clinopyroxene ranges from (Di50–70 Hd28–53 Jo1–2) to (Di19–73 Hd26–­77 Jo2–6)

and the composition of garnet ranges from (Ad95–99 Gr1–5) to (Ad40–61 Gr36–58), respectively. The presence of anisotropic grossular garnet
with high Fe2+/Fe3+ in crosscutting pyrrhotite-pyrite-bearing veinlets coupled with hedenbergitic pyroxene (Mg-poor clinopyroxene
with higher Fe2+/Fe3+) is consistent with reducing conditions during the later stage of prograde skarn alteration. The Çatak iron skarn is
characteristic, with its high sulphide content due to the presence of pyrrhotite, pyrite, and arsenopyrite, and low proportion of garnet
to pyroxene. The sulphur isotope (δ34S) compositions in the pyrrhotite-dominant skarn zones range between +0.84 to –2.23‰. We
interpret the bulk of the sulphur in the system as of igneous derivation and there has not been any significant sulphur contribution from
a crustal source. Fluid inclusion measurements conducted on skarn minerals of the proximal zone and distal zone+vein skarn revealed
high homogenization temperatures (371 to >600°C) and varying salinity values (10.5 to >70 wt% NaCl). The fluid inclusion data indicate
that there are at least three fluids associated with the genesis of the proximal skarn where the high garnet/pyroxene ratios are found.
Fluid inclusions that represent the early stages both in garnet and pyroxene plot in ‘Primary Magmatic Fluid’ and ‘Metamorphic Fluids’
fields. A magmatic fluid, presumably located at deeper parts of the system, mixed with a metamorphic fluid during its ascent. Over all
the Eğrigöz skarn a weak or moderate retrograde skarn alteration envelope formed, dominated by the incursion of meteoric waters
in the system, indicating limited fluid-rock interaction. Hydrofracturing resulted in pressure decrease and inclusions with Type III
(L+V+S) inclusions that plot in the ‘Secondary Magmatic Liquid’ and ‘Magmatic Meteoric Mixing’ fields.
Key Words: geochemistry, iron skarn, calc-silicate, Eğrigöz, Turkey

1. Introduction
In the Cenozoic copious magmatic activity took place in
Western Anatolia and the Aegean region. Magmatism was
most widespread and abundant during the oldest phase,
which began in the Late Eocene (about 37 Ma ago) and
ended in the Middle Miocene (about 14–15 Ma ago). It
is represented by volcanic and plutonic rocks of orogenic
affinity. The Eybek, Kozak, Alaçam, and Eğrigöz volcanoplutonic centres, predominantly consisting of intrusive
rocks, are the main examples of this early phase (e.g., Yılmaz
1990). The Eğrigöz Plutonic Complex (EPC) is situated
inland in Western Anatolia within the core and cover
*Correspondence:

sequences in the northeastern part of the Menderes Massif

(Figure 1). The EPC, with an outcrop area of approximately
550 km2, is one of the largest plutons in western Turkey,
and is associated with a number of mineral occurences
including iron skarns, Au-Ag-bearing mesothermal PbZn-Cu veins, skarns and gossans (Özgenç et al. 2006). The
district has been of economic interest since the second half
of the 20th century, and the magnetite resources around
the EPC are becoming increasingly important. Recent
geochronological studies focused on the crystallizing and
cooling ages of the Eğrigöz granite and yield ages around
20 Ma (Işık et al. 2004; Ring & Collins 2005; Hasözbek et

61


40

yu

Macedonia

Greece

Ductile shear
zones and
detachment
faults

Post Miocene
brittle faults

Menderes
Massif

İzmir-Ankara
zone

EĞRİGÖZ PLUTONIC
COMPLEX

TAŞBAŞI FORMATION

KIZILBÜK FORMATION

TERTIARY VOLCANICS
and TUFFS

EMET FORMATION

ALLUVIUM

GERNİ

TOKLAR

GILMANLAR

EVCİLER

İMRANLAR

ÖRENCİK

3

ÇALDİBİ

KIRKBUDAK FORMATION

BUDAĞAN FORMATION

İMRANLAR FORMATION

DAĞARDI MELANGE

ure

Fig

ÇOBANLAR

KIŞLAKÖY

DOLAYLAR FORMATION

SİMAV METAMORPHICS

SARICASU FORMATION

ARIKAYA FORMATION

0

N

MINERALIZATION

DETACHMENT
FAULT

THRUST FAULT

APPROXIMATE
FAULT

5 km

EĞRİGÖZ

MÜMYE

FAULT

MUSALAR

KÜRECİ

Figure 11

ÇAYIR

KIZILBÜK

DEĞERMİSAZ

Figure 1. Location of the study area in Turkey and geology of the study area (simplified after Akdeniz & Konak 1979; Işık & Tekeli 2001; Erkül 2010).

Granite

4 km

Hisarcık

Emet

Iraq

Iran

Eastern
stocks

Simav

Syria

Eğrigöz plutonic
complex

TURKEY

ANKARA

Miocene
volcanosedimentary
succession

Naşa

Kütahya

Armenia

Alluvium

Ko

e

İstanbul

nit

gra

nob
a

GEORGIA

QUATERNARY

CENOZOIC

40

MESOZOIC

30

PLIOCENE
TERTIARY
MIOCENE
PALEOCENE

CRETACEOUS
TRIASSIC JURASSIC

62
PALEOZOIC

Bulgaria

OYMAN et al. / Turkish J Earth Sci


OYMAN et al. / Turkish J Earth Sci

al. 2010). Although some studies have been carried out on
aspects of the economic potential of the district, studies
of ore and calc-silicate paragenesis and stable isotope

studies are lacking (Gümüş 1967; Özocak 1972; Taşan
& Cihnioğlu 1984). The aim of this study is to examine
the characteristics of this important, but otherwise little
known iron district and develop a model for its genesis.
For this purpose we focus on several iron deposits (Sakari,
Çatak, and Küreci) along the northern contact of the EPC,
which have representative ore and gangue paragenesis
for the mineralization of the northern contact. In these
deposits we describe the presence of changing redox
conditions in ore-forming magmatic-hydrothermal
systems based on the mineral chemistry of calc-silicates
(e.g., garnet, pyroxene, amphibole) and associated
sulphides of Cu, Fe, and As (e.g., chalcopyrite, pyrrhotite,
arsenopyrite). Compositional variations in calc-silicate
mineralogy reflect differences in magma chemistry, wall
rock composition, depth of formation, and oxidation
state (Burton et al. 1982; Gamble 1982; Meinert 1992,
1997). Fluid inclusion data, together with petrological and
isotopic information, may provide complete information
for knowing the P and T evolution of the skarn system.
Although the Sakari, Çatak, and Küreci iron skarns
are all spatially and genetically related to the EPC, the
differences in skarn texture, paragenesis, and geochemistry
are significant. Geochemical studies on ore and coexisting
calc-silicates in prograde stage or hydrous calc-silicate
overprint related to hydrothermal fluids give important
clues on the heat and fluid transfer from a cooling magma.
Rare earth element (REE) mobility is favoured by low
pH, high water/rock ratios, and abundant complex ions
(CO3–2, F–, Cl–, PO43–, SO42–) in the hydrothermal solutions

(Michard 1989; Lottermoser 1990, 1992). REE contents of
hydrothermally altered rocks in epithermal and porphyry
copper ore deposits indicate that fluid alteration is an
important agent in the mobilization of REE (Lottermoser
1990; Hopf 1993; Arribas et al. 1995; Bierlein et al. 1999;
Fulignati et al. 1999). Skarn systems adjacent to granite are
the most likely sources of the REE and speciation control on
the uptake and deposition of the REE from hydrothermal
fluids of different temperatures and composition (Smith et
al. 2000). Wang & Williams (2001) noted that REE were
transported in skarn-forming fluids and that their current
distribution is influenced by the occurrence of phases,
such as allanite and apatite in Cu-Au (Co-Ni) skarns in
the Cloncurry district in Queensland, Australia. REE
contents of ore samples from Pena Colorado iron skarns
in Colima, Mexico represent those of the andesitic tuffs
of the volcano-sedimentary rocks (Zürcher et al. 2001).
Trace, and rare earth element (REE) abundances provide
an opportunity to investigate the interaction between the
mineralizing fluids and the host rocks, with the chemistry

of ore-grade samples assisting us to identify and classify
the type of the deposit.
In this paper, we present new geochemical data from
the host granite and ore, including compositions of skarn
minerals, microthermometric studies on skarn minerals
and isotope determinations (S isotopes from pyrite and
pyrrhotite and O isotopes from magnetite, pyroxene, and
garnet). Based upon this data, we interpret the formation
conditions of the various iron skarns with different

paragenetic, spatial and temporal characteristics related to
the emplacement of the EPC.
2. Geologic and tectonic setting of Western Anatolia
The geological evolution of Western Anatolia was mainly
governed by Palaeo- and Neo-Tethyan events, which
preserve remnants of the Tethyan ocean. The Neotethys
Ocean was obliterated by the collision of the Eurasian and
African plates mainly during the Late Cretaceous–Tertiary
(Şengör & Yılmaz 1981; Şengör 1987). As a remnant of
Neotethys, the İzmir-Ankara Melange zone separates
the Sakarya Zone and the Anatolide-Tauride Block now
exposed in the metamorphic core complex of the Menderes
Massif. Intrusion of Palaeogene granitoids in the Sakarya
Zone, the İzmir-Ankara mélange zone, and the Menderes
Metamorphic core complex are linked by the subduction
of the East Mediterranean ocean floor, along the Hellenic
trench (Fytikas et al. 1984; Pe-Piper & Piper 1989; Gülen
1990; Delaloye & Bingöl 2000). The convergence has been
generated in a N–NE direction by subduction along the
Aegean and Cyprean arcs in the western and eastern
Mediterranean, respectively. The Menderes metamorphic
core complex has undergone five phases of metamorphism
(Bozkurt & Oberhänsli 2001). The age of the main
metamorphism affecting the whole massif is Palaeocene–
Eocene (Satır & Friedrichsen 1986; Hetzel & Reischmann
1996; Bozkurt & Satır 2000; Lips et al. 2001; Rimmele et al.
2003; Bozkurt 2004). Intracontinental N–S convergence
associated with the Palaeocene–Eocene collision along the
İzmir-Ankara suture zone continued until the Oligocene.
During the early Miocene, crustal thinning in the central

Menderes Massif was associated with the denudation
of the core complex in the footwalls of the Gediz and
Büyük Menderes detachment faults (Emre & Sözbilir
1997, 2007; Lips et al. 2001). These comprise mylonitised,
metamorphic, and granitic rocks lying below a low-angle
detachment fault, with associated chlorite brecciation and
two supradetachment basins containing a thick succession
of nonmarine strata (Hetzel et al. 1995; Emre & Sözbilir
1997; Koçyiğit et al. 1999; Sözbilir 2001, 2002; Seyitoğlu et
al. 2002; Işık et al. 2003, 2004; Bozkurt & Sözbilir 2004).
The footwall metamorphic rocks were progressively
mylonitised, exhumed, and intruded by syndeformational
granitoids (Turgutlu and Salihli granodiorites: the Salihli

63


64

4329217N

Kütahya-Simav

0490245E

4345601N

Ayazmant

Ayvalık

Kazdağ

500 000

5 750 000

estimation

0680634E

Kalkan

2 000 000

1 900 000

900

150 000

4400497N

4344044N

Kütahya-Simav

30–40 % Fe

78 000

No reserve

0683710E

Karaağıl

30–50 % Fe

18 000

0.6 % Cu;

46 % Fe and

pyrrhotite

Fe-Cu

Plutonic Complex

porphyries of Kozak

monzodioritic

granodioritic to

Pluton

diorite of Evciler

limestone lenses

intercalations, skarn after

carbonate lenses and

metamorphic rocks with

hornfels after regional

after marble

metamorphic rocks, skarn

hornfels after

rocks

Eğrigöz Pluton

granodiorite to quartz
Fe

metamorphic

calc-silicate hornfels after

dolomites

monzogranite of

granodiorite-

Eğrigöz Pluton

monzogranite of

skarn after limestone and

rocks

Eğrigöz Pluton
granodiorite-

metamorphic

calc-silicate hornfels after

monzogranite of

granodiorite-

lenses

Di-

Scp,

Di, Ad-Gr,

Hd, Scp, Pl

Ad, Gr,

Ad-Gr, Di-Hd

Po, Gn, Sp, various
Au-Ag-Te-Se

Or, Chl, Cal,
Qz,

Minerals

Bn, Mo, Go, Hem,

Mt, Cp, Py,

Po, Py, Cp

Cp)

Mt, Hem, (Py,

Mt, Py, Hem, Go

Oyman 2010

Öztürk et al. 2008

Öztürk et al. 2005;

Taşan & Cihnioğlu 1984

Taşan & Cihnioğlu 1984

1972

Gümüş 1967; Özocak

2003

Po, Cp, Py, Apy, Mt

Yücelay 1975; Tufan
(Apy, Cp, Ilm)

Tamer & Kurt 1982

Önal et al. 2009

Dora 1971;

Gümüş 1964;

Karaaslan & Başarı 1979

Reference

Gn, Sp, Py, Po,

Mt, Cp,

Py

Sp, Cp, Gn, Hem,

Sp, Gn)

Mt, Hem, (Py, Cp,

Ore minerals

Ep, Amp, Pl,

Chl, Qz

Ep, Cal, Amp,

Ep, Cal, Amp,

Scp

Di-Hd

Ep, Cal, Amp,

Di-Hd

Ant

Qz, Cal, Rt,

Or, Pl, Ep, Chl,

Tr, Cal, Ep

Cal, Tr

Ep, Cal

Late minerals

Ad-Gr,

Ad-Gr

after limestone blocks and

Di-

of Eybek Pluton

Hd

Ad-Gr,

Di-Hd, Scp

Wo, Ad-Gr,

Di-Hd,

Wo, Di-Hd,

Au (up to

Fe

Fe

Fe

Zn, Pb

Early
Minerals

metamorphic rocks, skarn

hornfels after

calc-silicate marble

amphibolite,

bearing schist

skarn after calcerous, marn

after limestone blocks

metamorphic rocks, skarn

hornfels after

Host rock

granite, granodiorite

Pluton

granodiorite of Şamlı

of Eybek Pluton

Cu

Fe-Cu

granite, granodiorite

of Eybek Pluton

granite, granodiorite

rocks

Associated igneous

Pb, Zn,

Fe

Metal

14 ppm) in

50 % Fe

49 % Fe

50 % Fe

4.10 % Pb

4.38 % Zn

58% Fe

1.25 % Cu

3 % Zn,

7 % Pb

50–60 % Fe

43 000

250 000

Grade

Size(t)

0485103E

4336320N

Kütahya-Simav

Evciler

0695750E

4404200N

0518300E

4408780N

0570560E

4398456N

0519986E

4390300N

0504800E

Location

Çatak

Karaaydın

Şamlı

Bağırkaç

Yaşyer

Deposit

Table 1. Major skarn mineralization in the Aegean region of Turkey.

OYMAN et al. / Turkish J Earth Sci


OYMAN et al. / Turkish J Earth Sci

granodiorite yielded 39Ar-40Ar amphibole isochron and
biotite plateau cooling ages of 19.5±1.4 and 12.2±0.4 Ma,
respectively) (e.g., Hetzel et al. 1995; Koçyigit et al. 1999;
Lips et al. 2001; Sözbilir 2001, 2002; Seyitoğlu et al. 2002;
Işık et al. 2003, 2004; Bozkurt & Sözbilir 2004). Different
models were proposed to explain the origin and timing of

the extension and granitoid intrusions in western Turkey.
The back-arc spreading (Le Pichon & Angelier 1979, 1981),
the tectonic escape (Şengör 1979; Şengör et al. 1985) and
the orogenic or gravitational collapse model (Dewey 1988;
Seyitoğlu & Scott 1996) are the proposed models. The
Tertiary magmatism in Western Anatolia consisted of
three geochemically distinct phases of magmatic activities
due to S–SW retreat of the active subduction zone
(Doglioni et al. 2002; Innocenti et al. 2005). The oldest
phase of the magmatic activity in western Anatolia, which
began in the Late Eocene (at about 37 Ma) and ended in
the Middle Miocene (at about 14–15 Ma), is represented
by volcanic and plutonic rocks of orogenic affinity. The
Eybek, Kozak, Alaçam, and Eğrigöz volcano-plutonic
centres, predominantly consisting of intrusive rocks, are
the main examples of this early phase (e.g., Yılmaz 1990).
Radiometric dating of the synextensional granites, which
intrude the Menderes Massif and Afyon Zone, are Early
Miocene in age. K-Ar age determination by Delaloye &
Bingöl (2000) shows that the subduction event must have
commenced before the Oligocene. Based on their K-Ar
age determinations on biotite and orthoclase from the
Eğrigöz granitoid, cooling ages were obtained of 20.0±0.7
to 20.4±0.6 Ma and 21.2±1.8 to 24.6±1.4 Ma, respectively.
The EPC was emplaced during the late stage of mylonitic
deformation of the extensional tectonic regime and was
deformed along the boundary of core rocks of the MCC
(Işık et al. 2004). On its northern and western boundary,
the Eğrigöz granitoid is separated from the cover series by
the Simav detachment fault (Işık et al. 1997; Işık & Tekeli

2001; Erkül 2010) (Figure 1). The cover sequences of the
MCC consist mainly of schist, recrystallized carbonates,
and ophiolitic mélange which experienced varying grades
of metamorphism. Intrusion and cooling of the Eğrigöz
granitoid occurred at 22.86±0.47 Ma (40Ar/39Ar ages, Işık
et al. 2004). More recently, U-Pb zircon analyses have
yielded crystallization ages of 19.4±4.4 Ma for the Eğrigöz
granite (Hasözbek et al. 2010). A cooling age of 18.77±0.19
Ma for the Eğrigöz granite was obtained by Rb–Sr (whole
rock, biotite) analyses.
3. Local geology
The oldest country rocks surrounding the Eğrigöz granitoid
are the Simav metamorphic sequence, which is exposed
southeast and west of the Eğrigöz Pluton (Figure 1). This
unit is comprised of biotite-muscovite schists, muscovitequartz schists, garnet schists, muscovite-quartz-biotite

schists, basic schists, quartzite and chlorite-bearing calcschists (Akdeniz & Konak 1979). The Balıkbaşı Formation
conformably overlies the Simav metamorphics in the south
of the EPC, and is comprised of laminated, bituminous
recrystallized limestone reflecting neritic facies conditions.
The Balıkbaşı Formation is uncomformably overlain by
Upper Palaeozoic–Lower Triassic Sarıcasu Formation,
mainly comprising schists that are metamorphosed
greenschist grade equivalents of detrital sediments with
basic tuffs and lavas. These schists were first identified
by Kaya (1972) as part of the İkibaşlı Formation and are
equivalent to the Sarıcasu Formation of Akdeniz & Konak
(1979). The Sarıcasu Formation is gradationally overlain
by the Arıkaya Formation that outcrops widely around
the village of Küreci. Based on lithological correlation

between the Arıkaya Formation and fossiliferous Permian
limestone and its boundary with overlying fossiliferous
Middle–Upper Triassic rocks, the age of the formation has
been thought to be Permian (Akdeniz & Konak 1979). The
limestone lenses in pelitic schists and the meta-carbonate
rocks of the Arıkaya Formation are important host rocks
for magnetite and pyrrhotite skarns.
3.1. Host rocks
Aureoles associated with EPC commonly show conspicuous
evidence of metasomatic processes related to local
injection of magma or hydrothermal fluids into country
rocks. The skarn zone between the Eğrigöz pluton and the
surrounding Palaeozoic Sarıcasu and Arıkaya formations
extends ~3.5 km along the contact and is ~10 to ~100 m
wide. The skarn bodies in Çatak are hosted by both the
Palaeozoic Sarıcasu and Arıkaya formations, but the bulk
of the ore bodies lie within the Sarıcasu Formation (Figure
1). This unit consists of muscovite-quartz-albite schist, and
muscovite-chlorite-calcite-quartz schist intercalated with
phyllite and crystalline limestone. Phyllite intercalated
with schists crops out near the village of Gürepınar. The
phyllite has well-developed cleavage, and is composed of
quartz, plagioclase, sericite, chlorite, titanite, and opaque
minerals. Recrystallized limestone lenses appear within
the schist sequence. One of the the largest outcrops of
these lenses is observed around the Elçekkaşı ridge. The
limestone is fine grained and grey. We speculate that
hydrothermal fluids migrated along pre-skarn fractures,
sedimentary contacts, and other permeable zones until
they reacted with the schist intercalated with carbonaterich intervals in the Sarıcasu Formation. Successive

fracturing and infiltration of more evolved mineralizing
fluids may have destroyed the original spatial zonation.
Along the northern border of the Eğrigöz Pluton
close to the contact, the rock is hornfelsed in a 10- to120-m-thick zone identified by its fine schistosity and
dark green to olive green colouration. Besides the more

65


OYMAN et al. / Turkish J Earth Sci

common porphyroblastic schists, minor mica-rich
varieties with lepidoblastic fabric and quartzo-feldspathic
varieties were observed. Albite porphyroblasts, quartz with
undulatory extinction, calcite, muscovite, chlorite, pyrite,
and magnetite are the main constituents of the schist. The
schistosity strikes mainly NE with variable dips of 30–60°
NW. Muscovite-quartz-albite schist was typically observed
along the eastern contact of the Eğrigöz pluton. The schist
generally possesses medium-high strength with medium
to coarse schistosity. The rock has a porphyroblastic texture
with quartz and albite porphyroblast lengths exceeding 3
cm. The porphyroblast-hosted matrix consists mainly of
muscovite and minor amounts of epidote, calcite, and
chlorite.
The Arıkaya Formation is a bedded package of
recrystallized limestone with a dense joint system,
intensely developed fold structures, and local brecciation.
Recrystallized limestones with grain sizes up to 2 mm
are common, and the limestone is locally dolomitic. The

ore-bearing calcic skarn is located at the northern contact
zones with granodiorite.
3.2. Plutonic rocks
The plutonic rocks in the region have a holocrystalline,
hypidiomorphic texture with quartz, plagioclase,
orthoclase, biotite, and hornblende as major rockforming minerals. Apatite, zircon, titanite, and muscovite
are common accessory phases. The opaque minerals
(e.g., magnetite, pyrite), chlorite, sericite, epidote, and
tourmaline are present, but less common. Fine-grained
plutonic varieties were observed close to the contact with
the skarn. Aplitic dykes of the EPC cross-cut plutonic
rocks associated with magnetite deposits at the eastern
and southern contacts, whereas in the centre and northern
part of the pluton, dykes are associated with polymetallic
veins.
Selected whole-rock analyses of 10 samples from
the EPC are listed in Table 2. In total 68 fresh, coarseto medium-grained phaneritic granitoid samples were
collected from the northern part of the EPC and 10 of
them were selected, based on their location near the
mineralization sites. In the QAP molecular normative
diagram (Streckeisen 1976), the rocks plot in the
monzogranite and granodiorite fields with a transitional
trend (Figure 2a) and in the Q-P diagram (Debon & Le Fort
1983), plot in the granodiorite and adamellite field (Figure
2b). The granitoids are plotted on the boundary between
the metaluminous and peraluminous granitoid fields in
an aluminum saturation diagram (Maniar & Piccoli 1989)
(Figure 2c). The granitoid rocks are calc-alkaline on the
Na2O+K2O–CaO versus SiO2 diagram (Frost et al. 2001)
(Figure 2d). The chondrite-normalized REE patterns

for EPC granitoids related to iron skarns are shown in
Figure 2e (normalizing values after Sun & McDonough

66

1989). The granitoids exhibit moderate LREE fractionated
patterns with a negative Eu anomaly (Eu/Eu*= 0.44–
0.61). They show nearly flat HREE patterns with Tb/Ybn
~1.2. Transitional trends from pre-plate collision to syncollision settings of granitoids are distinctive in Figure 2f.
The Rb vs Nb+Y diagram (Pearce et al. 1984) emphasizes
that the granitoids are volcanic arc granites (Figure 2h)
and plot in the volcanic arc + syn-collision field in the
Nb-Y diagram (Figure 2h).
4. Contact metamorphic assemblages
The limestone lenses in pelitic schists of the Sarıcasu
Formation and the meta-carbonate rocks of the Arıkaya
Formation are important host rocks for magnetite
and pyrrhotite skarns in the Çatak region. Early distal
isochemical metamorphism caused the formation of
calc-silicate hornfels and marble. Hornblende-biotite
hornfels contains quartz, hornblende, biotite, plagioclase,
muscovite, and andalusite. Pyroxene hornfels contains
diopside, garnet, plagioclase, K-feldspar, biotite, cordierite,
and sillimanite. Metamorphosed carbonate rocks include
marble. Metamorphism of pelitic schists and calcareous
rocks has produced an assemblage of andalusite,
cordierite, sillimanite, feldspar, biotite, tourmaline,
and quartz. The presence of cordierite, andalusite, and
sillimanite imply metamorphic temperatures of 500°C
and a maximum pressure of 2kb (Winkler 1967; Mason

1990). Stabilities of calc-silicate assemblages depend on
mole fractions of CO2 in the aqueous phase, as well as
on pressure, temperature, and the composition of solid
solution minerals (Sato 1980; Newberry 1982; Meinert
1982). If CO2 was near 0.1 kb during the metamorphism
and early skarn formation, a minimum temperature of
550°C is required for grossularite and wollastonite stability
at 2 kb (Greenwood 1967; Gordon & Greenwood 1971;
Meinert 1982). A minimum temperature of 550°C is also
required for the presence of diopside in pyroxene hornfels.
As prograde alteration reflects the protoliths, pelitic
schists are represented by hornblende hornfels, calcareous
rocks by pyroxene hornfels and meta-carbonate rocks by
pyroxene-garnet skarns.
The pyroxene hornfels facies is developed within a
restricted zone close to the contact with the pluton. The
presence of sillimanite indicates pyroxene-hornfels facies
conditions and aluminum rich protoliths (Winkler 1967).
The paragenesis in the pyroxene-hornfels facies comprises
cordierite, pyroxene, garnet, plagioclase, sillimanite,
orthoclase, biotite, titanite, apatite, chlorite, and quartz.
The hornblende hornfels facies is characterized by the
occurrence of andalusite, amphibole, biotite, muscovite,
quartz, and subordinate plagioclase, calcite, chlorite,
apatite, tourmaline, axinite, titanite, and zircon. Typical
pressures for the hornblende hornfels are less than 4


OYMAN et al. / Turkish J Earth Sci
Table 2. Composition of plutonic rocks.

Weight %

CT-1

CT1-8

CT1-13

CT1-14

CT2-22

CT2-23

CT3-7

OC1-26

KUR1-7

KUR-14

KB-1

SiO2
Al2O3
TiO2
Fe2O3
BaO
MnO

CaO
MgO
K2O
Na2O
P2O5
LOI
Total
ppm
Ba
Rb
Sr
Ga
Nb
Zr
Y
Th
Ni
Cr
V
Cu
Pb
Zn
La
Ce

70.93
14.33
0.32
2.82
0.09

0.03
2.09
0.73
4.13
3.51
0.03
0.94
100.05

71.09
13.92
0.32
2.43
0.09
0.04
2.12
0.73
4.23
3.24
0.12
1.1
99.52

71.16
13.87
0.34
2.89
0.09
0.04
2.03

0.89
4.62
3.13
0.01
0.8
99.95

67.13
15.08
0.45
3.55
0.11
0.06
2.9
1.22
3.95
3.51
0.14
0.85
99.04

72.15
13.76
0.29
2.69
0.09
0.05
2.1
0.8
4.16

3.24
0.11
0.61
100.15

65.9
15.52
0.56
4.22
0.13
0.08
3.13
1.49
4.08
3.38
0.1
0.88
99.55

66.56
15.37
0.51
4.09
0.12
0.06
3.16
1.35
3.74
3.69
0.198

0.85
99.78

67.38
15.2
0.47
3.78
0.11
0.06
3
1.27
3.75
3.61
0.11
0.73
99.57

67.96
14.78
0.43
3.58
0.12
0.05
2.84
1.12
3.9
3.41
0.05
0.88
99.22

70.01
14.42
0.35
2.9
0.11
0.05
2.42
0.9
4
3.5
0.03
1.11
99.9

69.08
14.47
0.41
3.39
0.13
0.08
2.4
0.99
4.26
3.37
0.13
0.77
99.6

785

164
230
18
14
145.5
26.5
16
15
690
15
25
25
25
39
72.5

1443
164.4
211.6
19.1
15
151.9
28.5
20.2
20
500
25
2.2
100
25

34.5
59.8

745
175
201
17
20
153.5
27.5
29
16
530
35
15
25
25
34.5
64

973
169
309
19
15
168.5
25
19
17
500

55
15
30
50
39
70.5

701
167
218
17
14
132
23.5
19
18
550
30
15
25
35
31
55.5

1100
152
330
20
17
212

31.5
19
19
440
65
15
25
55
44.5
80.5

1025
173
334
19
16
208
27
16
20
510
65
15
35
55
42
75.5

906
174

313
19
15
187.5
27
19
21
550
50
20
40
50
44
78.5

1005
153
300
18
14
175
25.5
58
22
550
50
15
25
30
37

65.5

937
153
282
17
14
157.5
24.5
21
23
550
35
15
25
35
39.5
70.5

1075
176.5
287
18
15
181
27
19
24
660
40

15
35
55
45
81

4
4.6
4.4
2.6
0.8
5.5
5
0.9
0.4
34
28.5
7.9
5.6
5
1.5
0.7
>0.5
0.4
7
2
2.7
>1

2.4

4.1
4.35
2.6
0.76
4.28
4.5
0.88
0.38
0.4
22.8
6.69
5.1
4
1.7
0.74
>0.5
0.4
6
1.1
2.66
>1

5
3.9
4.5
2.7
0.7
4.7
5
0.9

0.5
20
24.5
6.9
4.9
5
2.5
0.6
>0.5
0.4
8.5
3
3.2
>1

6
6.6
4.2
2.3
1
4.9
5
0.9
0.4
20
26.5
7.5
5
5
1.5

0.6
>0.5
0.4
6.5
3
2.6
>1

4.5
4.6
3.9
2.2
0.8
4.2
4
1
0.4
22
21.5
6.1
4.3
4
1.5
0.6
>0.5
0.4
5.5
5
2.7
>1

7.5
5.9
4.9
2.8
1.1
5.8
6
0.9
0.5
18
31
8.5
5.8
5
1.5
0.7
>0.5
0.4
4.5
5
3
>1

7
10.4
4.6
2.6
1.1
5.7

6
0.9
0.4
22
29
8
5.5
6
1.5
0.7
>0.5
0.4
6.5
3
2.8
>1

6.5
6.7
4.3
2.6
1
5.6
5
0.9
0.4
22
29
8.4
5.5

6
1.5
0.7
>0.5
0.4
6
3
2.8
>1

6
4.5
4.2
2.5
1
5
5
0.9
0.4
20
25.5
7
4.9
5
1.5
0.6
>0.5
0.4
12
3

2.7
>1

5
6.3
4.1
2.4
0.9
4.8
5
0.9
0.5
22
25.5
7.4
4.9
6
2
0.6
>0.5
0.4
6
4
2.9
>1

5.5
5.7
4.6
2.6

1.1
5.8
5
0.9
0.4
24
30
8.6
5.6
5
1.5
0.7
>0.5
0.4
6
4
2.8
>1

Co
Cs
Dy
Er
Eu
Gd
Hf
Ho
Lu
Mo
Nd

Pr
Sm
Sn
Ta
Tb
Tl
Tm
U
W
Yb
Ag

67


OYMAN et al. / Turkish J Earth Sci

kilobars, and temperatures range between 400 and 650°C.
The existence of andalusite indicates low- to intermediategrade metamorphism in the surrounding aureole. With
increasing distance from the contact aureole, greenschist
facies regionally metamorphosed rocks represent the
cover series.
5. Skarn mineralogy and paragenesis
More than 30 iron skarn occurences are known in the
Sarıcasu and Arıkaya formations, of which 12 had been
mined between 1950 and 1970. In Çatak district, the
area along the northern contact of the EPC contains
more than 15 iron deposits. The Çatak iron skarn district
consists of several bodies that include: Sakari, Çavdarlık,
Göğez, and Katranlı (Figure 3). Among these zones of

mineralization, the Sakari prospect differs because it
is a magnetite-dominated ore, compared to the other
prospects in the Çatak district. In the Katranlı, Göğez,
and Çavdarlık prospects, iron mineralization occurs
commonly as tabular bodies and lenses associated with
disseminated and stockwork-type deposition. Subordinate
small crosscutting veins and veinlets are also present. The
mineralization is closely associated with metasomatic
skarn consisting mainly of pyroxene, garnet, plagioclase,
amphibole, epidote, calcite, and quartz, preferentially
replacing pyroxene hornfels facies rocks between 10 to 100
m from the Eğrigöz granodiorite.
5.1. Mineralogy of skarn in the Katranlı, Göğez and
Çavdarlık districts
A narrow reaction zone (20 cm to 1.5 m thick) is developed
in the Göğez and Çavdarlık endoskarn toward the proximal
zone of the pluton (Figure 4). At the contact, the granite is
a darker greenish colour due to the metasomatic reaction
with the wall rock pyroxene hornfels. Chlorite, amphibole,
and epidote are the characteristic metasomatic minerals
in the endoskarn. Fracture-controlled metasomatism is
the most common replacement mechanism. Chlorite and
amphibole, as pseudomorphs of pyroxene, are the main
calc-silicate minerals associated with opaque minerals in
these fracture fillings. Disseminated anhedral to subhedral
opaque crystals are mantled by chlorite crystals in the
endoskarn. Epidote and plagioclase are associated with
chlorite and amphibole to a lesser extent. The plagioclase
is replaced by epidote and calcite. Biotite is also replaced
by chlorite and amphibole pseudomorphs. Some

disseminated anhedral ore minerals occur, which are
hematitized and limonitized.
The calcic exoskarn is composed chiefly of pyroxene
with subordinate garnet and amphibole. Microprobe
analyses were performed mainly on pyroxene, garnet,
amphibole, pyrrhotite, chalcopyrite, and arsenopyrite.
Details of the analytical methods of microprobe analysis
are given in the Appendix.

68

Garnet and pyroxene crystals are fractured and crosscut
by veinlets of late stage ore and retrograde minerals.
Euhedral to subhedral pyroxene is the earliest calc-silicate
mineral of the prograde stage with the grain size ranging
between 20 µm and 1 mm. Pyroxene grains within the
exoskarn range between hedenbergitic and diopsidic end
members with an average composition of Di43–53 Hd46–56
Jo1–2 (Figure 5a; Table 3). Optically and compositionally
zoned individual garnet grains typically are 10 µm – 3
mm in diameter. In the exoskarn zone, garnet is andradite
(Ad97–99) within a narrow compositional range (Figure
5b; Table 4). Pyroxene grains are extensively included
in garnet, magnetite, and pyrrhotite as relict crystals.
Pseudomorphic amphibole replacement after pyroxene
is the most common retrograde alteration, followed by
the replacement of the epidote, chlorite, and prehnite as
vein-fillings. The composition of amphiboles varies within
the calcic amphibole types (Na+K<0.5). Based on Leake
et al. (1997), amphibole compositions from Çatak plot as

actinolite-ferroactinolite and Mg hornblende. Amphiboles
of the Çatak pyrrhotite skarns have higher Mg, Si, and
lower Al values than those of the Sakari magnetite skarns.
Tourmaline with high magnesium and aluminum
contents (dravite) is common in the surrounding schist
and hornfels, suggesting a metasomatic origin related
to early contact metamorphism. Pyrrhotite, magnetite,
pyrite, and chalcopyrite in descending order are abundant
ore minerals. Accessory minerals include arsenopyrite,
gersdorffite, melnicovite pyrite, linnaeite, ilmenite, and
rutile. Magnetite follows coarse crystalline pyroxenegarnet skarn as an early ore phase. Exsolution lamellae
of ilmenite in magnetite appear as a function of oxygen
fugacity and temperature and are probably related to the
cooling stage of igneous activity.
Sulphides always post-date magnetite precipitation
(Figure 6a). Pyrrhotite developed between the euhedral
granular magnetite crystals and replaced magnetite along
their crystal edges. Lamellar intergrowth of pyrrhotite with
linnaeite is a common ore texture in the ore zone (Figure
6b). Pyrrhotite and/or pyrite-pyrrhotite veinlets crosscut
the early magnetite. In some samples, pyrrhotite forms as
a prograde texture as rhythmically banded features with
retrograde amphibole, which replaced prograde pyroxenes.
Chalcopyrite commonly replaces pyrrhotite and also fills
pyrrhotite-pyrite interstices (Figure 6c). As a late stage
event, melnikovite pyrite replaces both pyrrhotite and
chalcopyrite. Replacement of pyrrhotite by melnikovite
pyrite resulted in the development of well-developed ‘bird’s
eye’ texture (Figure 6d). Alteration of pyrrhotite along its
grain boundaries is a common hydrothermal process.

The porous texture of this secondary pyrite (melnikovite
pyrite) indicates appreciable volume decrease during
replacement (Figure 6c, d). Melnikovite pyrite also formed


OYMAN et al. / Turkish J Earth Sci

Q= Si/3-(K+Na+2Ca/3)

1000

300

200

4

100

8

b

a

12

3

7

11

2

6
10

1

5
9

0
-400

-300

-200

-100

0

100

200

300

P=K-(Na+Ca)

alkali feldspar granite
granite
granodiorite

MALI

trondjhemite
a

tonalite

a-c
quartz diorite
c-a

diorite

c

c

d
SiO2 (wt.%)

e

g

f

h

Figure 2. Plots comparing the major and minor element contents of plutonic rocks of the EPC (see Table 1 for data). (a) Normative
compositions of granitoids plotted on the classification diagram of Streckeisen (1976). Q= quartz; A= (Or); P= (Ab + An); (b) Plot
(after Debon & Le Fort 1983) displaying the mean composition of of the plutonic rocks; (c) A/NK vs A/CNK diagram, after Maniar &
Piccoli (1989); (d) Modified alkali versus silica plot showing the calc-alkaline affinity of the plutonic rocks, after Frost et al. (2001); (e)
Chondrite-normalized rare earth element patterns of plutonic rocks. Normalization values from Sun & McDonough (1989); (f) Majorelement geotectonic discrimination diagrams of the plutonic rocks; (g, h) Trace element geotectonic discrimination diagrams (Pearce
et al. 1984).

69


70

44

100

42

400

ÇAVDARLIK
PROSPECT

58

GÖĞEZ
PROSPECT

500 m

681000

300

42

200

35

73

43

682000

KATRANLI
PROSPECT

58

40

42

42

SAKARİ 75
PROSPECT

IMRANLAR
FORMATION
ARIKAYA
FORMATION
SARICASU
FORMATION

COLLUVIUM
AKDAĞ
VOLCANITE
EĞRİGÖZ
GRANITE

683000

HORNFELS and
SKARN ZONE

EXPLANATION

86

55

ALLUVIUM

33

45

64

42

FAULT

4366000

4367000

4368000

Figure 3. Çatak iron skarn mineralization, consisting of several skarn bodies including: Sakari, Çavdarlık, Göğez, and Katranlı (modified after Özocak 1972).

0

N

OYMAN et al. / Turkish J Earth Sci


OYMAN et al. / Turkish J Earth Sci

T1
-1

005

C

CT1-8
CT-1
CT1-5
SK-4

endoskarn

hornfels

skarn + ore

granodiorite

ore
pyrrhotite>>magnetite

0

6m

rubble

Figure 4. Cross section displaying contact between plutonic rocks and the pyrrhotite-dominated Göğez iron skarn in the Çatak district.

as a replacement of iron-rich calc-silicates (Figure 6e). Due
to the intense deformation euhedral to subhedral pyrites
and arsenopyrites are both fractured. Pyrite crystals are

rimmed with arsenopyrite which is also precipitated in
fractures of pyrite (Figure 6f).
Late stage hydrothermal events represented by fissurecontrolled veins clearly cut the sulphide precipitation and
surrounding skarn zone. Quartz veins with chalcopyrite
blebs have mantled the clasts of pre-existing ore minerals.
Native gold is associated with quartz in some samples.
Supergene effects include martitization of magnetite,
replacement of chalcopyrite and pyrrhotite by goethite
(Figure 7).
5.2. Sakari prospect
The geometry and the zonation of the skarn in Sakari
Tepe (Çatak District) is shown in Figure 8. At Sakari,
the mineralization and associated skarn were formed by
successive fracturing and infiltration processes, although
the contact with the intrusive rock is not exposed. The
skarn forms lenticular bodies and exhibits a gradual
contact with hornfelsic wall rocks. The deposit consists
of a zoned body, in which the central core is a pyroxenegarnet dominated, coarse-grained skarn associated with
massive ore. Pyroxene is the earliest calc-silicate mineral
that precipitated, due to interaction of a calcareous host
rock with ore-bearing metasomatic fluids (Figure 9a).
In the central core early pyroxene in magnetite-bearing
skarn is diopside (Di50–70 Hd28–53 Jo1–2) (Table 3, Figure
5a). In the central core early pyroxene is replaced by
magnetite and subordinate amounts of garnet (Ad95–99
Gr1–5) (Table 4). Pyrrhotite is the early sulphide mineral

that replaces prograde stage iron-bearing calc-silicates,
chiefly andradite (Figure 9b). The early stage pyroxenemagnetite-garnet association is both replaced and crosscut
by late stage, anisotropic garnet (Ad36–58 Gr40–61) (Figure

5b) and hedenbergitic pyroxene (Di19–73 Hd26–77 Jo2–6).
In the garnet-dominated coarse-grained magnetite
skarn, crosscutting pyroxene has a hedenbergite-rich
composition (Di19–73 Hd26–77 Jo2–6). In some places pyroxene
grains are replaced by amphibole, and both are replaced by
scapolite. Scapolite is also found as fracture and vug filling
fine-to coarse-grained crystals (>0.5 mm). Quartz, epidote,
chlorite, and rarely amphibole are the main retrograde
phases and they occur either in a fracture network,
which crosscuts or replaces the prograde skarn mineral
assemblage. Epidote is the most common retrograde
calc-silicate, and grossular is the dominant prograde calcsilicate (Figure 9c).
The main ore-bearing skarn zone grades into distal
banded hornfels, composed of pyroxene, amphibolebearing mafic layers and plagioclase, quartz, K-feldspar,
and cordierite-bearing felsic horizons. The mineralogy
probably reflects the lithological control of its protolith
during metasomatism. Pyroxene from the contact
metamorphic zone close to granodiorite is diopside
(Di49–76 Hd24–53), associated with epidote, albite, and
amphibole. Monomineralic diopside horizons probably
replace former thin carbonate intercalations, whereas
the polymineralic layers are developed over pre-existing
impure greywackes. Diopsidic hornfels is the early calcsilicate, which is replaced mainly by amphibole in bands.
The metasomatic fluids crystallizing pyroxene were not as

71


OYMAN et al. / Turkish J Earth Sci
Çatak skarn

Küreci skarn
Sakari hornfels
Sakari prograde skarn Stage I
Sakari prograde skarn Stage II

a
Hedenbergite

Diopside

Çatak skarn zone
Küreci skarn
Sakari prograde skarn
Sakari late prograde skarn

b
Grossular

Andradite

Figure 5. Ternary diagrams showing compositional variations of
clinopyroxene (a) and garnets (b) from the iron skarns
associated with the EPC.

enriched in Fe2+ and Mn as those of coarse-grained skarn
with massive ore. Fine-grained granoblastic plagioclase is
replaced by pseudomorphs of epidote in the plagioclase,
quartz, cordierite, sillimanite, and K-feldspar-bearing
bands of the hornfels.
The massive ore is dominated by euhedral magnetite

with crystal sizes ranging from 0.5 mm up to 2 mm.
Pyrrhotite, chalcopyrite, pyrite, and ilmenite are accessory
ore minerals. Sulphides constitute more than 10 percent of
the massive ore and post-date magnetite precipitation. We
speculate that the magnetite was deposited in at least two
stages. First, replacement of calc-silicates by magnetite is
the earliest and commonest event in the ore precipitation
stage. Following the deposition of prograde magnetite,
pyrrhotite appears, due to replacement of either magnetite
or iron-rich calc-silicates preferentially (Figure 9d). The
replacement textures in magnetite and/or pyrrhotite in
calc-silicates and pyrrhotite in magnetite are characteristic
features of deposition stages. The fractures resulting
from the intense deformation of magnetite and the
wall rock are filled by later sulphides, mainly hexagonal
pyrrhotite, chalcopyrite, and subordinate pyrite (Figure
9e, f). In this second stage, pyrrhotite was intergrown
with chalcopyrite and was subsequently replaced by
chalcopyrite. Where chalcopyrite is the only sulphide
phase, it appears to fill cavities, fractures and crystal
boundaries of granular magnetite. Exsolution lamellae
of ilmenite in metasomatic magnetite are a function of
oxygen fugacity and temperature related to the cooling
stage of igneous hornfels activity (Gasparrini & Naldrett
1972). Buddington & Lindsley (1964) first pointed out that

72

coexisting equilibrated pairs of titaniferous magnetite and
ilmenite may permit simultaneous determination of the

temperature and oxygen fugacity at the time of formation.
In Sakari rarely-observed fine-grained ilmenite crystals
were found as inclusions in magnetite and pyrrhotite.
Alteration of ilmenite to rutile is a common process.
Hematite and goethite after magnetite and goethite after
hematite and siderite are the main products of supergene
alteration (Figure 10). Hematite is relatively more abundant
and may form massive bodies as a replacement phase.
6. Skarns in the Küreci district
The geological map of the Küreci mineralization is
shown in Figure 11. The Küreci magnetite-specularite
mineralization consists of two main skarn bodies including
Maden Tepe and Karataş Tepe. In Maden Tepe, the extent
of the exoskarn zone may reach ~10 to ~50 m, whereas
the endoskarn zone is quite narrow up to a few metres.
The skarn zonation is: unaltered granodiorite, endoskarn,
andradite-diopside-magnetite skarn, magnetite, diopsidewollastonite skarn, and recrystallized limestone (Figure
12). The presence of wollastonite at the limestone front
suggests that the bulk of prograde skarn formation
occurred at temperatures between 550 and 600°C; within
the wollastonite stability field (Meinert 1982).
Granodiorite in the area can be affected by widespread
alteration due to the intense fracturing along the contact
with the Arıkaya Formation. The endoskarn is dominated
by plagioclase replacement after K-feldspar, chlorite after
amphibole, biotite, and plagioclase in the early skarn.
Fine-grained myrmekitic intergrowths occur in endoskarn
zones near contacts. Primary biotite and amphiboles in
the granodiorite are totally obliterated, whereas chlorite
and Fe-oxides were also precipitated along fractures and/

or crystal boundaries of plagioclase. The outer shell of the
granodiorite gains a darker greenish colour due to gradual
increase of chlorite towards the exoskarn. Magnetite
and pyrite appear as euhedral small grains disseminated
throughout the endoskarn (Figure 13a).
The host rock of the calcic skarn is mainly limestone
of the Permian Arıkaya Formation. The calcic skarn
assemblage is characterized by coarse, crystalline,
prograde garnet and pyroxene. The early stage of skarn
development is characterized by a well-developed garnet
skarn zone that extends ~25 m further out from the
contact of the intrusion. Garnets are andraditic (Ad95–99)
(Table 4). The massive garnet skarn locally contains
patches of magnetite and pyrite associated with retrograde
minerals (e.g., quartz, calcite, and chlorite) within cavities
and fractures (Figure 13b). The proportion of pyroxene
increases gradually towards the limestone (Figure 13c,
d). At the limestone contact, the skarn is represented by
a wollastonite-bearing zone, which can be several metres


12.64

24.71

0.00

0.02

0.00

0.00

100.0

MgO

CaO

NiO

Na2O

K 2O

Cr2O3

Total

100.05

0.02

0.00

0.04

0.00

23.83

11.64

0.67

10.77

0,803

100.32

0.05

0.00

0.31

0.00

16.26

9.73

0.58

20.19

3.14

0.08

49.980

OC14

0.991

0.000

0.007

0.000

0.000

Ca

Ni

Na

K

Cr

1.84

52.14

46.02

Johannsenite

Diopside

Hedenbergite

Mol percent end-members

0.535

Mg

0.487

0.472

0.019

Fe

Mn

0.009

0.006

Ti

48.03

49.92

2.06

0.000

0.000

0.009

0.000

0.997

0.506

0.021

0.000

0.000

Al

1.985

1.985

Si

44.68

53.36

1.96

0.000

0.000

0.005

0.000

0.995

0.544

0.020

0.456

0.005

0.000

1.987

structural formula on the basis of 6 oxygens

9.16

0.60

0.87

Al2O3

MnO

0.04

TiO2

FeOT

52.23

51.96

SiO2

0.04

OC14

OC14

75.83

21.18

3.00

0.001

0.000

0.013

0.000

0.982

0.211

0.030

0.757

0.015

0.000

1.996

100.19

0.02

0.00

0.17

0.00

22.88

3.54

0.88

22.59

0.31

0.00

49.81

OC17

73.32

20.73

5.96

0.000

0.000

0.009

0.000

0.945

0.217

0.062

0.766

0.006

0.000

1.998

99.1

0.0

0.0

0.2

0.0

21.6

3.00

1.20

22.9

0.70

0.00

49.50

OC17

76.66

19.17

4.17

0.000

0.000

0.020

0.000

0.937

0.200

0.043

0.799

0.075

0.001

1.948

99.44

0.00

0.00

0.25

0.00

21.54

3.30

1.26

23.52

1.56

0.05

47.60

OC17-8

37.69

60.81

1.50

0.000

0.000

0.005

0.000

0.984

0.627

0.015

0.389

0.006

0.000

1.986

99.99

0.01

0.00

0.07

0.00

24.07

11.02

0.48

12.17

0.14

0.01

52.03

OC24

48.88

49.36

1.76

0.000

0.000

0.005

0.000

0.977

0.504

0.018

0.499

0.008

0.000

1.994

100.24

0.02

0.00

0.06

0.00

23.61

8.75

0.55

15.44

0.18

0.02

51.62

OC24

23.78

75.55

0.67

0.001

0.000

0.010

0.000

0.991

0.765

0.007

0.241

0.006

0.001

1.991

99.81

0.04

0.00

0.14

0.00

24.71

13.71

0.22

7.69

0.13

0.02

53.17

OC24

5.49

94.39

0.12

0.001

0.000

0.006

0.000

0.993

0.948

0.001

0.055

0.007

0.000

1.994

99.68

0.03

0.00

0.09

0.00

25.42

17.45

0.04

1.81

0.17

0.00

54.67

KUR14

2.62

97.30

0.08

0.001

0.000

0.004

0.000

1.006

0.993

0.001

0.027

0.019

0.002

1.970

100.57

0,02

0,00

0,07

0,00

26.01

18.46

0.03

0.89

0.44

0.07

54.59

KUR14

1.68

98.32

0.00

0.000

0.000

0.010

0.000

0.997

1.005

0.000

0.017

0.028

0.002

1.965

100.55

0.01

0.00

0.15

0.00

25.84

18.71

0,00

0.57

0.65

0.07

54.55

KUR14

45.61

53.16

47.96

50.58

1.46

0.000

0.001
1.23

0.000

0.015

0.000

0.312

0.709

0.021

0.673

0.075

0.002

2.081

100.37

0.01

0.00

0.21

0.00

7.78

12.74

0.65

21.52

1.70

0.07

55.68

CT18

0.000

0.010

0.000

0.316

0.769

0.018

0.660

0.054

0.001

2.074

99.99

0.02

0.000

0.141

0.000

7.874

13.766

0.56

21,049

1.22

0.05

55.31

CT18

Table 3. Selected results of electron microprobe analyses of pyroxenes of the Sakari (OC), Çatak (CT), and Küreci (KUR) districts.

56.01

42.39

1.60

0.001

0.000

0.014

0.000

0.472

0.543

0.020

0.717

0.049

0.001

2.083

99.68

0.02

0.00

0.19

0.00

11.51

9.50

0.63

22.37

1.09

0.02

54.35

CT18

OYMAN et al. / Turkish J Earth Sci

73


OYMAN et al. / Turkish J Earth Sci
Table 4. Selected results of electron microprobe analyses of garnets of the Çatak, Sakari and Küreci districts.
CT-12

CT-13

CT14

OC-17

OC-19

OC-22

OC-34

KUR1-4

KUR1-6

KUR1-9

SiO2

36.19

36.28

36.46

38.03

37.94

37.83

35.85

35.10

35.68

35.65

TiO2

0.00

0.04

0.00

0.13

1.75

0.20

0.00

0.00

0.05

0.00

Al2O3

0.11

0.67

0.05

12.51

12.92

11.16

0.84

0.28

0.09

0.01

FeOT

28.68

28.25

28.57

13.93

11.68

15.35

26.94

27.03

27.79

27.62

MnO

0.36

0.42

0.36

1.27

1.05

1.05

0.11

0.26

0.13

0.13

MgO

0.19

0.08

0.17

0.08

0.05

0.07

0.12

0.44

0.21

0.25

CaO

33.5

33.71

33.18

33.28

34.49

33.56

33.45

33.16

33.38

33.38

Na2O

0.02

0.03

0.00

0.01

0.00

0.00

0.01

0.00

0.00

0.02

Cr2O3

0.03

0.00

0.00

0.00

0.00

0.01

0.00

0.01

0.01

0.03

Total

99.06

99.48

98.80

99.27

99.92

99.26

97.33

96.30

97.35

97.09

Structural formula on the basis of 24 oxygens
Si

5.981

5.960

6.043

6.001

5.940

5.997

6.006

5.944

5.992

6.001

Al

0.019

0.040

0.000

0.000

0.060

0.003

0.000

0.056

0.008

0.000

Ti

0.000

0.005

0.000

0.016

0.207

0.024

0.000

0.000

0.006

0.000

Fe+3

3.963

3.882

3.903

1.637

1.321

1.867

3.774

3.828

3.903

3.888

+2

0.000

0.000

0.057

0.202

0.209

0.169

0.000

0.000

0.000

0.000

Mn

0.050

0.058

0.052

0.170

0.140

0.141

0.016

0.038

0.018

0.018

Fe

Mg

0.046

0.020

0.043

0.020

0.014

0.017

0.032

0.113

0.053

0.063

Ca

5.931

5.935

5.892

5.626

5.785

5.700

6.004

6.017

6.006

6.019

Na

0.006

0.011

0.000

0.003

0.003

0.000

0.003

0.003

0.002

0.005

Cr

0.003

0.000

0.000

0.000

0.000

0.001

0.000

0.001

0.001

0.004

Mol percent end-members
Andradite

99.39

97.68

99.73

41.28

35.64

47.21

95.80

98.51

99.50

99.85

Grossular

0.00

1.00

0.00

53.31

61.58

48.86

3.39

0.00

0.00

0.00

Pyrope

0.52

0.33

0.27

0.33

0.24

0.29

0.53

1.45

0.46

0.05

Almandine

0.00

0.00

0.00

2.20

0.00

1.22

0.00

0.00

0.00

0.00

Spessartine

0.00

0.96

0.00

2.85

2.52

2.37

0.26

0.00

0.00

0.00

Uvarovite

0.08

0.00

0.00

0.00

0.00

0.03

0.00

0.03

0.03

0.09

thick. Massive magnetite ore and associated magnetite
veinlets are more common in pyroxene and wollastonitedominated skarns.
The Karataş Tepe mineralization differs from the Maden
Tepe district in its distinct emplacement, paragenesis, and
evolution. Specularite-quartz veins that cut the hornfels
have a N25°E strike with an average dip of 40°W. The
paragenesis of the vein is quite simple. Specularite is the
main ore mineral with subordinate amounts of pyrrhotite,
ilmenite, rutile, chalcopyrite, pyrite, goethite, and siderite.

The surrounding rock was intensively silicified along
its contact with the vein. Specularite occurs as elongate
euhedral crystals 0.5 to 0.9 mm long with lamellar
twinning. Fine needles of specularite interstitial with

74

quartz are widespread. Rutile is a common ore mineral
and pseudomorphous anatase after rutile is common.
Pyrrhotite is rare, with individual crystals situated
between elongate specularite. Their grain sizes typically
range between 5–100 µm. Textural relationships indicate
that specularite precipitation is sequentially overprinted
by fine-grained polygonal quartz crystals (Figure 13e).
Sericitization is superimposed and localized in the vein
and neighbouring fracture zones. Specularite is replaced
by goethite along its margin throughout the vein (Figure
13f). Rarely observed anhedral chalcopyrite and euhedral
pyrite is also replaced by goethite. The whole process
was terminated by the pseudomorphous replacement of
goethite by siderite.


OYMAN et al. / Turkish J Earth Sci

b

a

calcsilicate

ln

py

po
cpy

po
mt
0.06 mm

c

d

po

cpy

po

py
py
0.02 mm

0.06 mm

f

e
calcsilicate

py

py
mt

asp

0.06 mm

0.06 mm

Figure 6. Photomicrographs of skarn assemblages from the Çatak district. (a) Pyrrhotite (po) dominated sulphidefilled fractures cutting the pre-existing magnetite (mt); (b) Lamellar intergrowth of pyrrhotite (po) with linnaeite
(ln) and replacement of pyrrhotite by chalcopyrite (cpy) are common ore textures; (c) Infilling of interstices between
pyrrhotite (po) and pyrite (py) with bird eye texture by chalcopyrite (cpy); (d) The porous texture in melnikovite
pyrite (py) after pyrrhotite (po); (e) Melnikovite pyrite (py) formation as a replacement of iron-rich calc-silicates; (f)
Euhedral to subhedral pyrite (py) crystals are coated with later arsenopyrite (asp).

7. Fluid inclusion studies
Fluid inclusion measurements were conducted on primary
inclusion assemblages of clinopyroxene and garnet mainly
in proximal zones of the exoskarn. In the 10 samples from
the Çatak and Kureci areas, only primary fluid inclusions
yielded what we consider to be reliable homogenization
temperatures (Th) and final ice melting temperature (Tm-

ice) measurements. The results are summarized in Table
5 and Figure 14. Three different types of inclusions were

identified in garnets, based on room temperature phase
properties: (1) Type I inclusions are two phases (liquid +
vapor) and liquid-rich at room temperature; (2) Type II
inclusions contain two phases (liquid + vapor) and are
vapor-rich at room temperature and (3) Type III inclusions

75


OYMAN et al. / Turkish J Earth Sci

Mineral

Stage

Skarn stage
Prograde

Fluid inclusions in pyroxenes were two phase liquidrich inclusions with a gas-to-liquid ratio between 0.25
and 0.70. Fluid inclusions in pyroxene are small (5–15
μm), whereas the garnet-hosted inclusions have diameters
ranging from 10 to 30 μm (Figure 15). Type I and Type
II inclusions have been observed in the garnets associated
either with the magnetite or pyrrhotite ore. However Type
III inclusions were observed mainly in the magnetite ore
at Çatak.
In garnets from Küreci (KUR 1-4), formed along
the contact between granodiorite and limestone, the
homogenization temperatures of inclusions (Types I
and II) vary from 306.5 to over 600°C. The mean of the

frequency distribution yield temperatures of 424°C.
However the homogenization temperatures of inclusions
(Types I and II) in the Çatak garnets (CT-12) vary from
227 to over 600°C. The mean of the frequency distribution
yield temperatures of 473°C (Figure 14a).
The final dissolution of daughter crystals always
preceded the final homogenization. Halite dissolved at
temperatures ranging from 285 to 492°C in garnets from
Çatak (EMK-3). Ensuing homogenization to liquid after
halite dissolution occurs between 405 to over 600°C.
Salinities for Type I and Type II inclusions were,
based on final ice melting (T-m ice), computed. For Type
III inclusions (liquid + vapor + solid) from the Çatak
garnets (EMK-3), the salinity was determined from the
temperature of dissolution of the halite.
Final ice melting temperatures of inclusions in the
Çatak garnets (Types I and II in CT-12) range from –2.1
to –18.2°C (Table 5), which correspond to salinities of
3.5–21.1 wt% NaCl equivalent (Potter et al. 1978; Bodnar
1993). Final ice melting temperatures of Types I and II
inclusions of garnets from Küreci (KUR1-4) range from
–20.4 to –69°C (Table 5), which correspond to salinities of
10.4–22.6 wt% NaCl equivalent (Potter et al. 1978; Bodnar
1993).

Supergene
stage

Retrograde

Pyroxene
Garnet
Plagioclase
Apatite
Epidote
Amphibole
Axinite
Chlorite
Calcite
Quartz
Magnetite
Pyrrhotite
Linnaeite
Pyrite
Chalcopyrite
Arsenopyrite
Melnicovite pyrite
Gersdorffite
Rutile
Goethite
Hematite
Siderite

Figure 7. Schematic diagram showing paragenetic relationship
of skarn and ore assemblages of the pyrrhotite-bearing skarns.

are multiphase (liquid + vapor + solid) inclusions. They
contain solid crystalline phases known as daughter
minerals.

2
-2
C

C

O

O

O

6
O

C

-1

-1
C
O

O

C

-1
1

4

C

-1

-1

9

7

110

H
H

H

H H
H
H
H H

H
H

H

H

H

H

H

H

H
H
H

H H
H
H
H
H
H
H

H

skarn + ore

hornfels

skarn>>ore
0

ore
mag>>pyro

H H
H H

oxidized ore

4m

rubble

Figure 8. Geological section with sample locations illustrating the general setting of the Sakari magnetite skarn.

76


OYMAN et al. / Turkish J Earth Sci

a

b

calcsilicate

c

d

e

f

Figure 9. Photomicrographs of skarn assemblages from Sakari. (a) Replacement of early pyroxene (cpx) by magnetite; (b) textural
relationship between early calc-silicates (garnet chiefly) and ore minerals magnetite (mt) and pyrrhotite (po); (c) epidote (ep) is the most
common retrograde mineral in garnet (gt) dominated (gt>cpx) skarn; (d) precipitation of magnetite (mt) and pyrrhotite (po) following
the deposition of pyroxene in the main prograde stage; (e, f) the fractures of magnetite and the wall rock are filled by later sulphides,
mainly hexagonal pyrrhotite (po) and chalcopyrite (cpy).

Diopside-dominated pyroxenes in the proximal zone
are represented by a homogenization temperature of 424
to >600°C. Inclusions in pyroxene are characterized by a
salinity range of 6.2–10.6 wt% NaCl equivalent (Figure
14a, b).
The first ice-melting temperatures vary from –37 to
–67 °C, indicating the presence of CaCl2 in addition to
NaCl (Shepherd et al. 1985; Oakes et al. 1990).
Liquid-rich fluid inclusions (Type 1) locally occur
together with vapour-rich inclusions (type 2), suggesting

heterogeneous trapping of a boiling fluid during pyritearsenopyrite-pyrrhotite vein formation in the Çatak skarn
(Roedder 1984; Bodnar 1995).
8. Geochemistry
Geochemical analyses were performed on different rock
types, including the intrusion (n= 11), skarn (n= 10),
hornfels (n= 6) and ore facies (n= 8), to characterize
the mass transfer during mineralization and different
types of alteration. Trace and rare earth element (REE)

77

OYMAN et al. / Turkish J Earth Sci

Mineral

Stage

Skarn stage
Prograde

Retrograde

Supergene
stage

Pyroxene
Garnet
Wollastonite
Plagioclase
Scapolite
Apatite
Epidote
Amphibole
Prehnite
Chlorite
Calcite
Quartz
Ilmenite
Magnetite

Pyrrhotite
Pyrite
Chalcopyrite
Rutile
Goethite
Hematite
Siderite

Figure 10. Schematic diagram showing paragenetic relationship
of skarn and ore assemblages of the magnetite bearing skarns.

abundances provide an opportunity to investigate the
interaction between the mineralizing fluids and the host
rocks. Details of the analytical methods of whole-rock
geochemical analysis are presented in the Appendix. Data
for all samples are given in Table 6.
8.1. Hornfels
Chondrite-normalized REE patterns of hornfels samples
are shown in Figure 16a. Enrichment of LREE over HREE
is obvious, with slightly negative Eu, except in sample
OC1-14, which is a hornfels composed mainly of pyroxene,
amphibole, and plagioclase. The positive Eu anomaly
in OC1-14 is significant, with a relatively high Eu/Eu*
value (Eu/Eu*= 2.14). It seems likely that the interaction
of hydrothermal fluids with the plagioclase-rich rock
led to successive enrichment in Eu. OC-24 is a hornfels,
containing high proportions of pyroxene and subordinate
feldspar, quartz, amphibole, and chlorite. Tourmalinebearing metasomatized schist (CT4-3) surrounding
pyrrhotite-dominated skarns has high REE concentrations.
The degree of enrichment in this sample is higher in LREE

78

than in HREE with a negative Eu anomaly. Sample CT4-3
shows slightly negative Ce/Ce* (Ce/Ce*= 0.90) and high
La/Ybn (La/Ybn= 11.51) ratios, with greater enrichment of
LREE than HREE and a negative Eu anomaly (Eu/Eu*=
0.44). Trace-element data normalized to average crust are
plotted in Figure 16b. Hornfels samples are enriched in
U, Rb, Tb, Tm, Y and Yb and depleted in Ba, K, Sr, Zr, P
and Ti relative to average crust. The hornfels samples show
similar REE element abundances and patterns to intrusive
rocks, although with lesser negative Eu anomalies. This
could be indicator of the interaction between melt and
meta-pelitic rocks during the emplacement of intrusive
rocks.
8.2. Prograde skarn and ore samples
In the Maden area, endoskarn (KUR1-2) and exoskarn
samples KUR1-3, KUR1-4 and KUR1-6 are located
with increasing distance from the contact, and exhibit
concordant REE patterns reflecting depletion with
increasing distance from the pluton towards the
recrystallized limestone front (Figure 16c). The same
trend in trace elements is also evident in Figure 16d.
This depletion trend implies that trace and REE were
transported by magmatic hydrothermal fluids, although
their partition between the minerals was controlled by
distribution coefficients between fluids and minerals. The
REE pattern of KUR1-3 is pronounced with a positive Eu
anomaly, which reflects its higher Eu/Eu* ratio (Eu/Eu*=

1.23). The positive Eu anomaly, observed only in KUR1-3,
may be related to its zoned garnets that formed by lattice
diffusion and growth entrapment processes. Garnetdominated skarn samples (OC-19 and OCT-16) display
depletion of LREE and a relative enrichment in heavy
REE with a convex-up pattern (Figure 16Ee). Sample
OC-19 is a fine-grained skarn sample composed mainly
of garnet, pyroxene, magnetite, and epidote. OCT-16 is a
coarse-grained skarn sample composed mainly of garnet,
pyroxene, amphibole, chlorite, and magnetite. This type
of REE distribution in garnets is attributed to garnetites
(Whitney & Olmsted 1998). The positive Eu anomaly in the
hedenbergite-grossular-epidote-dominant skarn sample
(OC-17) is significant with the relatively high Eu/Eu*
value (Eu/Eu*= 3.45).
Garnet-bearing samples have relatively high HREE
patterns, and garnet appears to account for most of the
HREE. Due to its larger size (r= 1.26 A), Eu2+ can only be
hosted by a nearly pure andradite end-member (Whitney
& Olmsted 1998). Eu3+, transported in retrograde
hydrothermal fluids probably producing grossular
veinlets, may also account for the positive anomaly
(Whitney & Olmsted 1998). OC-17 has grandite garnet
(Gr53 An41) reflecting that the positive anomaly is related to
the enrichment of hydrothermal fluid in Eu3+ rather than
Eu2+ due to the substitution of Ca by Eu2+.


OYMAN et al. / Turkish J Earth Sci

6

87000

6

89000

6

N

91000

KARATAŞ
PROSPECT

43

65000

42

KÜRECİ
VİLLAGE

42

41

31

MADEN
PROSPECT

43
56

EXPLANATION
Akdağ
Volcanites
Eğrigöz
Plutonic
Complex

63000

Skarn Zone
Budağan
Formation

750

0

m

Kırkbudak
Formation
Sarıcasu
Formation

Fault

Figure 11. Simplified geology of the Küreci district showing skarn zonation (modified after Akdeniz & Konak
1979; Taşan & Cihnioğlu 1984).

Anhedral unzoned garnets in sample CT4-8 have
a significantly similar REE pattern with a positive Eu
anomaly, marked by a relatively higher Eu/Eu* ratio
(~2.17). The sample, composed mainly of andraditic
garnet (CT1-12) from the Çatak pyrrhotite skarn, has a
rather flat REE pattern that may be related to increasing
initial proportions of silicates to initial calcite (about 50%
initial calcite, Whitney & Olmsted 1998). Sample CT2-11
is a pyroxene-dominated, magnetite, amphibole, epidotebearing skarn that displays high LREE concentrations
(∑REE= 178.5 ppm) with a high La/Ybn ratio (La/Ybn=
30.5) and Ce/Ybn ratio (Ce/Ybn= 12.73).
8.3. Ore samples
Ore samples have depleted REE concentrations (Figure
16f). The magnetite-dominated ores of the Sakari Prospect
(OC-22, OC-25, OC1-6) are more depleted in LREE than
the others. Magnetite in prograde skarn sample OC16 and pyrrhotite in sample OC-22 have similar REE

patterns. The magnetite-rich sample OC-22 has a positive
Eu anomaly with high Eu/Eu* and Ce/Ce*, similar to
sulphide-rich skarns CT1-8 and CTa-4. Sample CTa-4 is
composed mainly of pyrrhotite, pyrite, chalcopyrite, and
arsenopyrite and has a significant Eu enrichment (Eu/
Eu*= 2.48) and a slight negative Ce anomaly (Ce/Ce*=
0.82). The Eu enrichment could be related to hydrothermal

fluid circulation which precipitated sulphur minerals
throughout the prograde skarn formation. Sample CT18 was collected from a sequence containing rhythmic
pyrrhotite bands in a retrograde amphibole-rich matrix,
and has moderate Eu and depleted Ce contents (Eu/Eu*=
0.87; Ce/Ce*= 0.43).
Pyrrhotite-rich sample CT1-5, collected close to the
granite contact, has a depleted REE pattern with relatively
low Eu and Ce contents (Eu/Eu*= 1.52; Ce/Ce*= 0.51).
Specularite-rich samples KUR-6 and KUR-10 with
quartz gangue have enriched LREE (La/Ybn 10.76 and

79


OYMAN et al. / Turkish J Earth Sci

KU
R1
-2

andradite-diopside-magnetite

KU
R
16

KU
R
14

KU
R
1-

3

105

magnetite

diopside+wollastonite

granodiorite &
endoskarn

skarn

Skarn + ore

ore
mag>>pyro

recrystallized
limestone

rubble

0

3m

Figure 12. Geological section showing skarn zonation in the Küreci prospect.

4.78 respectively) patterns relative to magnetite-rich ore
samples from Sakari Tepe. These rocks have negative Eu
anomalies (Eu/Eu*= 0.55 and 0.55) with lower Eu (Eu= 0.1
and 0.2) and Ce contents (Ce= 4.46 and 6.14).
8.4. Stable isotopes
To estimate the origin of the hydrothermal fluids,
sulphur and oxygen isotope analyses were carried out
at the University of Tübingen, Department of Geology
(Germany) and the University of Nevada, Reno (USA),
respectively. Details of the analytical methods of isotope
analysis are given in the Appendix.
Oxygen isotope compositions were determined
on coexisting magnetite-pyroxene, magnetite-calcite,
and pyroxene-quartz pairs to calculate fractionation
temperatures (see Table 7). Fractionation equations
of Bottinga & Javoy (1973, 1975) for the pyroxenemagnetite pair were used to calculate the equilibration
temperature. Pyroxene and magnetite from sample KAL-5
were extracted from intergrown pyroxene and magnetite
enclosed in euhedral garnet crystals. In this sample, the
magnetite value is higher than the pyroxene, indicating the
two minerals are not in equilibrium. The magnetite oxygen
isotope value of this sample is uncommonly high (9.34‰).
The ore consists mainly of magnetite that is oxidized
to hematite, which in some places is associated mainly
with calcite as the main gangue mineral. Using the
fractionation equations of Clayton & Keiffer (1991) for
magnetite-calcite, temperatures of 294°C and 511°C, were

estimated. Magnetite and associated calcite deviate only
slightly toward marine carbonate values.
Sulphur isotope analyses were performed on 8
monomineralic concentrates of sulphide minerals (4
pyrrhotite grains and 4 pyrite grains) taken from pyrrhotitedominated ore in the Göğez skarn. The sulphur isotope

80

compositions in the pyrrhotite dominant skarn zones are
tightly constrained within a δ34S range of –2.23‰ to +
0.84‰ (Figure 17, Table 7). Sulphur isotope temperatures
from pyrrhotite and pyrite range 405 to 525°C, and show
that pyrrhotite and pyrite are the earliest sulphur minerals
(Stage 1) of the sulphur stage.
9. Discussion
9.1. Intrusive compositions and geodynamics
Compositional variations in calc-silicate mineralogy
reflect differences in magma chemistry, wallrock
composition, depth of formation, and oxidation state
(Meinert 1997). Eğrigöz ore formation occured in different
stages of magmatic-hydrothermal activity associated with
an oxidized, I-type, high-level granitoid magma chamber.
Geochemical characteristics of the Egrigöz pluton indicate
an origin from partial melting of mafic lower-crustal source
rocks. In western Anatolia, the melt generation mechanism
for the intrusive rocks could be crustal extension and uplift
following collision (Özgenç & İlbeyli 2008).
9.2. Infiltrative skarn petrogenesis
In the early skarn stage, iron is present in calc-silicates and
associated magnetite ore; in the second stage, iron combined

with sulphur to form minerals such as pyrrhotite, pyrite,
chalcopyrite, and arsenopyrite. The oxidized character of
the magmatic fluid responsible for the garnet-dominant
prograde assemblage in the exoskarn is consistent with a
porphyritic environment in a continental arc setting. The
garnet/pyroxene ratio and their successive relationship
are related to the composition and oxidation state of the
wallrocks.
Elsewhere, in the Ocna de Fier-Dognecea orefield
(Banat-Romania), clinopyroxene is diopsidic when
associated with magnetite, although it has a more


OYMAN et al. / Turkish J Earth Sci

a

b

cpy

gt
cpx

mt
py
opaque
chl
pl

0.04 mm

c

0.3 mm

d

cal

amp

cpx
cpx
0.04 mm

e

0.3 mm

f
qtz

spc

qtz
spc
go
0.04 mm

0.04 mm

Figure 13. A to E crossed polars transmitted-light photomicrographs of skarn and ore textures in the Küreci district,
except F which is a reflected-light microphotograph. (a) Euhedral magnetite and pyrite crystals in the endoskarn
(KUR1-1); (b) pyroxene crystals in garnet (Ad98–100) from garnet-dominated skarn (KUR1-2); (c) diopsidic skarn
(Di94–98) where clinopyroxene is the dominant calc-silicate with less amounts of amphibole; (d) closer to the marble
front diopside after calcite; (e) specularite associated with quartz crystals; (f) Goethite (go) after specularite as a
result of supergene alteration.

hedenbergitic mole fraction when associated with garnet
(Ciobanu & Cook 2004). These authors also noted that
garnet from the Fe zone of calcic skarns is andraditic, but
where it is associated with pyroxene it has a greater grossular
component. Kwak (1994) noted that Mg-rich pyroxene and
andradite tend to occur in oxidized environments, whereas
hedenbergite and grossular tend to occur in reducing

conditions. Although the dominant garnet composition
is andraditic in composition and associated with Mg-rich
pyroxene, anisotropic and concentrically zoned grossularrich garnet compositions have been recorded in the Kara
magnetite-scheelite deposit (Zaw & Singoyi 2000).
The low to equal pyroxene/garnet abundance ratio
in the Sakari Tepe and Küreci mineralization indicates

81


OYMAN et al. / Turkish J Earth Sci
Table 5. Summary of fluid inclusion characteristics of the garnet and pyroxenes from the iron skarns associated with EPC. Th–
Homogenization temperature of vapor bubble; Tm-solid: Dissolution temperature of daughter minerals. Tm ice (°C): Melting of ice; Tm

first (°C): First melting temperature. The upper measurement limit of homogenization temperature is 600 °C.

Sample

Host mineral /
Inclusion type

Tm-solid (°C)

Th (°C)

Tm ice (°C)

Tm first (°C)

Salinity (equiv.
wt% NaCl)

KUR1-4

Küreci

Garnet / Type I>II

–

306.5 – >600

–20.4 / –69

–44.1 / 54.2

10.4 – 22.6

CT-12

Çatak

Garnet / Type I>II

–

226.7 – >600

–2.1 / –18.2

–42.3 / –54.8

3.5 – 21.1

EMK-22

Çatak

Pyroxene / Type I>II

–

423.7 – >600

–3.8 / –7.1

–37.4 / –52.4

6.2 – 10.6

EMK-3

Çatak

Garnet / Type III

285.3 / 492.4

405.2 – >600

–1.2 / –23.8

–44.2 / –67

37.1 – 58.7

dominantly oxidized conditions in wallrocks (Meinert
1998).
In oxidized environments, Mg-rich pyroxene and
garnet close to the andradite end-member are the most
stable calc-silicates. Unlike the other two skarns in the
northern contact of Eğrigöz, the Küreci occurrence is
hosted by marble. At Küreci, contact metasomatism of
limestones resulted in a skarn zone with an andraditediopside skarn towards the marble front. At Sakari, the

mineralization and associated skarn were formed by
successive fracturing and infiltration processes, but lack
a visible contact with intrusive rock in the field. The
contact metasomatic pyroxene is diopsidic. The system
tends to produce hedenbergitic pyroxene closer to the
proximal skarn in contact metamorphic rocks, where
the hydrothermal fluids would have migrated along preskarn fractures. In the skarn with iron mineralization
where hydrothermal fluids migrate, the final composition
of pyroxene is hedenbergite. At Sakari, it is possible to
offer a similar evolutionary trend in garnet composition
through time. The early garnet in the skarn association is
andradite, which is spatially and temporarily associated
with magnetite. The occurence of anisotropic grossular
in crosscutting veinlets coupled with hedenbergitic
pyroxene is consistent with reduced conditions during the
retrograde stage formation. From grossular-rich garnet
cores and accompanying hedenbergitic clinopyroxene, we
interpret that skarn formation was initially under relatively
reducing conditions (log fO2 ~21 to 23 bar at 1 kb),
perhaps 2±3 log units below the SO2/H2S buffer. Due to
the change of oxidation state towards reduced conditions,
sulphide saturation produces pyrrhotite, chalcopyrite, and
subordinate pyrite that post-dates magnetite and early
calc-silicate formation. Limited sulphide occurrences
at Sakari Tepe must be related to limited increase in fS2.
Mineralogical studies indicate that the magnetite-rich
prograde skarn is replaced and crosscut by a sulphide-

82

rich phase in the hydrothermal stage. Replacement of
pre-existing magnetite mainly by pyrrhotite and pyrite in
the skarn occurs by reactions which are independent of
oxygen fugacity (Burt 1972; Logan 2000). Sulphur isotope
temperatures from pyrrhotite and pyrite range from 405
to 525°C, and show that pyrrhotite and pyrite are the
earliest sulphide minerals (Stage 1) of the sulphur stage.
Arsenopyrite is an accessory phase in these samples. The
main arsenopyrite precipitation occurred in Stages 2 and
3. Most of the examined samples of arsenopyrites from
Stage 3 coexist with chalcopyrite and postdate pyrrhotite
(Figure 18). Chalcopyrite is not stable together with
pyrrhotite above 335°C. Cubanite replaces chalcopyrite
above this temperature. In some samples, former iron-rich
calc-silicate assemblages were replaced by pyrrhotite in
rhythmically banded textures. A variety of such textures,
involving magnetite and calc-silicate associations, were
also reported in the Ocna de Fier-Dognecea orefield
(Banat, Romania) (Ciobanu & Cook 2004). As in the
Ocna de Fier-Dognecea orefield, at least some pyrrhotite
may have precipitated during the prograde stage at Çatak.
Also, crosscutting veinlets of pyrrhotite and/or pyritepyrrhotite in early magnetite could indicate changing
redox conditions from oxidized to reduced. The presence
of arsenopyrite buffered by hexagonal pyrrhotite indicates
relatively low temperatures (<450±20°C) (Choi & Youm
2000).
As reactions in the carbonate protolith proceeded,
the fluid would have been changed from an oxidized to a
reduced condition with high pH. The ore-forming fluids in
the early magnetite stage (magnetite skarn) in the Küreci

sector were probably more oxidized and had lower pH
than at Çatak. In the Çatak sector, closer to the contact
zone, pyrrhotite, magnetite, pyrite, and chalcopyrite are
the dominant opaque minerals in a descending order of
abundance in which magnetite is markedly decreased.
This implies that the ƒO2 of the ore-forming fluid, from the


OYMAN et al. / Turkish J Earth Sci
20

a

18
16

ÇATAK
KÜRECİ

Garnet
Pyroxene
Garnet

N=59
N=18

Frequency

14
12

10
8
6
4
2
0
200 250 300 350 400 450 500 550 600
Homogenizationtemperature (°C)

34

ÇATAK

b

32

Garnet

30

N=52

28

ÇATAK

26

Type III

Garnet
N=7

24

KÜRECİ

22

Frequency

Type I>II

Pyroxene Type I>II

Type I>II

Garnet

20

N=18

18
16
14
12
10
8

6
4
2
0
0

5

10

15

20 25 30 35 40
Salinity (equiv.Wt.%NaCl)

45

50

55

60

Figure 14. Microthermometric data for fluid inclusions in
garnet and pyroxenes from the Çatak and Küreci skarns in the
Eğrigöz area. (a) Homogenization temperatures of primary fluid
inclusions in garnet and pyroxenes; (b) Average salinities wt%
NaCl eq. of primary fluid inclusions in garnet and pyroxenes.

magnetite skarn to the pyrrhotite skarn, was decreasing

whereas pH was increasing.
The sulphur isotope results reported here are consistent
with the interpretation that the bulk of the sulphur in the
system is of igneous derivation and implies that there has
been no significant contribution from crustal-sourced
heavy sulphur (e.g., Ohmoto & Goldhaber 1997). Similar
results for isotopically well-homogenized sulphur from
pyrrhotite, pyrite, chalcopyrite and arsenopyrite are
reported from other Cu, Au, W and Pb-Zn skarn deposits

(Figure 17) (Gray et al. 1995; Laouar et al. 2002; Chiarada
2003; Zhao et al. 2003).
9.3. Fluid inclusion constraints
Fluid inclusion measurements conducted on skarn
minerals in the proximal zone and the distal zone+vein
skarn in Çatak revealed high homogenization temperatures
(307 to >600°C) and varying salinity values (10.5 to 60
wt% NaCl). Fluid inclusions in the calc-silicates of the
prograde stage represent the composition of magmatic
fluids after reaction with the carbonate wall rocks. Most
of the inclusions in both garnet and pyroxene plot in the
‘Primary Magmatic Fluid’ and ‘Metamorphic Fluids’ fields
(Figure 19) (Bodnar 1999). A diagram of the temperature
of total homogenization versus salinity of all the fluid
inclusions studied reveals the existence of two fluids in the
Çatak region (Figure 19). An early fluid responsible for
magnetite deposition (corresponding to fluid inclusions of
Types I and II), is a high temperature fluid (temperature
of total homogenization: between 227–>600°C), with a
salinity ranging between 3.5–21.1 wt% eq. NaCl. At Çatak

a plot of salinity versus homogenization temperatures does
not define a continuous trend between the different stages
of skarn formation and may thus suggest intermittent flow
of external fluids during retrograde skarn formation at
different times, possibly in response to hydrofracturing.
The magmatic fluid, presumably originating from deeper
parts of the system, is mixed during its ascent with
metamorphic fluid rather than connate or groundwater
due to low fluid-rock ratio. Hydrofracturing in a transient
pressure increase subsequently results in a geostatic
pressure decrease and formation of Type III (L+V+S)
inclusions (temperature of total homogenization: between
405.2–>600°C), with a salinity ranging between 37.1–58.7
wt% eq. NaCl) that plot in the ‘Secondary Magmatic
Liquid’ and ‘Magmatic Meteoric Mixing’ fields (Bodnar
1999).
At Küreci the fluid (corresponding to Types I and II
fluid inclusions in garnets) responsible for magnetite
deposition along the contact between granodiorite and
marble is a high temperature fluid (temperature of total
homogenization: between 306.5–>600°C), with a salinity
ranging between 10.4–22.6 wt% eq. NaCl).
Calculated 18O values for anhydrous minerals from the
early prograde stage (garnet, magnetite and pyroxene) show
that they formed from fluids with a significant magmatic
component. Whitney et al. (1985) experimentally showed
that between 500 and 650°C, iron accounts for up to 50%
of the available chlorine in fluids in equilibrium with rocks
of granitic composition. Experimental work on magnetite
solubility indicates that iron is transported as FeCl2 in

sulphur-free, chloride-bearing supercritical fluids (Chou
& Eugster 1977; Boctor et al. 1980; Frantz et al. 1980). The
partitioning of metals between melt and magmatic fluid is

83


OYMAN et al. / Turkish J Earth Sci

b

a
V

L

L

V
10 m

c

10 m

d

L

V

V
S

S
L
10 m

10 m

Figure 15. Photomicrographs of primary fluid inclusions in garnet and pyroxene in skarn from Eğrigöz skarns. (a) Type I
inclusions are two phases (liquid + vapour) and liquid-rich at room temperature; (b) Type II inclusions contain two phases
(liquid + vapour) and are vapour-rich at room temperature; (c, d) Daughter-mineral (mainly halite)-bearing gas-rich, and
aqueous inclusions (Type-III) in garnet of Çatak skarn.

related to the Cl concentration of the magmatic fluid. In
the upper levels of the crust, the supercritical fluid consists
of vapour and hypersaline liquid phases (Sourirajan &
Kennedy 1962; Henley & McNabb 1978; Fournier 1987;
Shinohara & Fujimato 1994). Due to its low density,
the buoyant vapour phase separated from the dense
hypersaline liquid and began to differentially ascend.
Owing to the high crustal level of this emplacement, the
fluid began to boil during or very soon after its release,
promoting hydrofracturing and, in turn, further boiling.
Among CuCl3, CuCl, CuClOH, and CuCl2 the latter was
the dominant form of aqueous copper above 300°C. A
sudden increase in pH could also result from boiling,
due to the partitioning of acidic components (e.g., HCl,
CO2, SO2, H2S) into the vapour during phase separation

(Drummand & Ohmoto 1985). Copper transport by
magmatic vapour has been documented both in oxidized or
reduced porphyry Cu-Au deposits. However, Simon et al.
(2003) experimentally showed that the iron concentration

84

in the S-bearing magmatic volatile phase (1.1 molal) is
significantly higher than in the S-free magmatic volatile
phase (0.11 molal). The difference between ore-bearing
and barren porphyries may be that ore-bearing porphyries
crystallise from high oxygen fugacity melts, which do
not become depleted in Cu and Au when they become S
saturated, whereas barren porphyries crystallise from low
oxygen fugacity magmas that do.
9.4. Comparisons with other major iron deposits in the
Anatolides
Most of the important iron ore resources of Turkey occur
in Central-Eastern Anatolia, which has a geological
history of multiple orogenic and tectonic events.
Hercynian and pre-Hercynian phosphorus-rich (2% P)
magnetite deposits, such as Avnik (e.g., Helvacı 1984),
Bulam, Pınarbaşı, and Ünaldı, are hosted in Palaeozoic
metamorphic rocks of the Bitlis Massif. In central-eastern
Anatolia prolonged subduction (82.90±0.43 to 79.43±0.58
Ma) between Eurasian and Afro-Arabian plates was


OYMAN et al. / Turkish J Earth Sci
Table 6. Abundances of selected geochemical indicator elements in hornfelses, skarns and ores.

Çatak
Hornfels

Skarn+ore

Ore

%

CT2-1

CT4-3

CT2-10

CT2-11

CT1-12

CT4-8

CT1 - 5

CT1-8

CT - a4

Fe
Al2O3
CaO

MgO
S
TiO2
P2O5
ppm
Mo
Cu
Pb
Zn
Ni
As
Cd
Sb
Bi
Ag
Au * ppb
Hg
Tl
Co
Cs
Ga
Hf
Nb
Rb
Sn
Sr
Ta
Th
U
V

W
Zr
Y
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Tb/Ybn
La/Ybn
Eu/Eu*
Ce/Ce*

3.20
13.95
1.45
0.60
0.01
0.30
0.11

8.04
14.48
0.79
3.03
0.04
0.68
0.16

47.98
12.35
4.97
1.41
0.02
0.31
0.13

33.66
2.55
4.62
3.68
<0.4
0.213
0.046

42.46
0.73
33.25
0.69
0.03
0.02

0.03

27.68
0.79
4.56
5.06
16.15
0.035
0.03

64.93
0.13
0.16
0.12
30.13
0.03
0.02

28.38
0.74
13.62
3.86
14.52
0.02
0.04

39.25
0.83
6.81
2.75

17.41
0.05
0.03

1
2.8
19.5
26
3.1
2.7
0.1
0.1
0.3
0.1
0.8
0.01
0.3
3.3
10.8
18.8
4.7
15.8
208.6
3
182.1
1.8
20.5
7.4
22
4.4

143.1
34.3
36.8
60.8
6.89
24.2
5.1
0.88
5.06
0.81
4.99
0.98
3.1
0.47
3.27
0.47
1.13
8.07
0.53
0.92

3.2
14.5
101.8
212
34.5
12.3
2
0.5
0.4

0.3
0.7
0.02
0.5
5.3
11.3
27.5
5
10.8
270.5
14
115.7
0.8
11.6
6.5
94
7.7
170.2
23
36.9
66.1
8.46
32.4
6.6
0.8
4.69
0.75
4.38
0.73
2.07

0.34
2.3
0.32
1.48
11.51
0.44
0.90

0.4
19
5.6
14
2.1
4
<0.1
0.4
0.2
<0.1
<0.5
0.01
<0.1
6,8
1.3
22.9
4.5
11.2
27.3
20
161.9
1.2

12.3
5
28
0.5
178.6
17.3
14.1
19.7
2.19
9.2
2.5
0.59
2.64
0.48
2.55
0.55
1.75
0.24
1.93
0.26
1.13
0.24
0.70
0.85

0.17
111.31
4.67
41.5
5.7

2.9
<0.02
0.38
0.52
82
1.5
5
0.03
23.4
0.9
27.16
1.20
2.60
5.7
41.2
9
0.3
1.9
5.7
121
0.7
21.5
15.5
68
73.31
6.3
16.8
3
0.6
2.8

0.5
3
0.6
1.5
0.3
1.6
0.2
1.42
30.49
0.63
0.85

0.3
15.3
110
323
3.5
20.9
3.4
0.5
5.3
0.6
4
<0.01
<0.1
7.3
2.6
39
< 0.5
0.8

2.1
5405
11.6
<0.1
0.6
1.5
8
5.7
4.3
16.5
2.7
5.2
1.24
7.1
1.6
0.45
2.01
0.34
2.18
0.45
1.52
0.23
1.42
0.2
1.09
1.36
0.76
0.68

2.25

719.61
35.13
64.7
2.3
3.6
0.22
13.01
3.72
251
0.5
7
0.02>
4.9
1.1
5.09
0.12
0.84
1.1
5.7
13
0.1
0.6
3.8
5
22.3
3.7
4.4
4
6.23
1

3.3
0.7
0.5
0.7
0.1
0.6
0.1
0.3
0.1
0.3
0.1
1.52
9.56
2.17
0.75

0.38
1772.33
27.32
27.8
5.7
73.6
0.19
0.77
22
726
2.4
5>
0.02>
156.7

0.1
0.54
<0.02
0.16
0.4
1.2
2
<0.1
0.2
0.6
1
0.8
0.2
0.9
1
0.95
0.2
0.9
0.2
0.1
0.2
0.1
0.1
0.1
0.1
0.1
0.1
0.1
4.55
7.17

1.52
0.51

0.4
2.2
100.2
21
3.6
13.9
0.1
0.86
0.2
0.2
0.8
<5
0.2
2.4
4.1
19.1
4.5
15
164.4
4
211.6
1.7
20.2
6
25
1.1
151.9

28.5
6
4.39
1
3.6
0.7
0.2
0.7
0.1
0.7
0.1
0.5
0.1
0.7
0.1
1.26
9.30
0.87
0.43

1.08
390.18
107.93
48.9
4.0
2090.8
0.11
10.54
28.66
1825

52.8
<5
0.06
47.2
1
9.01
0.18
1.59
12.1
21.8
35
0.1
0.6
7.3
13
1.9
3.4
2.2
1
1.52
0.2
0.8
0.2
0.2
0.3
0.1
0.3
0.1
0.2
0.1

0.4
0.1
1.14
1.79
2.48
0.82

85