Fluid inclusion and stable isotope constraints (C, O, H) on the Dağbaşı Fe–Cu–Zn skarn mineralization (Trabzon, NE Turkey)

Journal Pre-proofs Fluid Inclusion and Stable Isotope Constraints (C, O, H) on the Dağbaş ı Fe– Cu–Zn Skarn Mineralization (Trabzon, NE Turkey) Yılmaz Demir, Ali Dişli PII: DOI: Reference:

S0169-1368(18)30302-0 https://doi.org/10.1016/j.oregeorev.2019.103235 OREGEO 103235

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Ore Geology Reviews

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5 September 2018 30 October 2019 14 November 2019

Please cite this article as: Y. Demir, A. Dişli, Fluid Inclusion and Stable Isotope Constraints (C, O, H) on the Dağbaş ı Fe–Cu–Zn Skarn Mineralization (Trabzon, NE Turkey), Ore Geology Reviews (2019), doi: https://doi.org/ 10.1016/j.oregeorev.2019.103235

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1

Fluid Inclusion and Stable Isotope Constraints (C, O, H) on the Dağbaşı Fe–Cu–Zn

2

Skarn Mineralization (Trabzon, NE Turkey)

3 4 5 6 7 8 9

Yılmaz Demir1 and Ali Dişli1 1Recep

Tayyip Erdogan University, Department of Geological Engineering, 53100 Rize, Turkey ([email protected])

Abstract

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The Dağbaşı Fe–Cu–Zn skarn mineralization developed along the contact between the block

11

and lens shaped limestones of the Lower Cretaceous Berdiga Formation and the Upper Cretaceous

12

Dağbaşı Granitoid. The exoskarn-type mineralization is characterized by prograde stage garnet and

13

pyroxene, while the retrograde stage is characterized by epidote, tremolite, actinolite, and chlorite.

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Quartz and calcites were observed in both stages of the skarn development. The ore minerals mainly

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consist of magnetite and hematite, with a lesser amount of pyrrhotite, pyrite, chalcopyrite, sphalerite,

16

and minor galena. The homogenization temperatures (Th) and salinity values of the prograde stage

17

halite-bearing fluid inclusions are in the range of 412 to 514 °C and 48.8–61.8 wt% NaCl equ.,

18

respectively. The second stage liquid- and vapor-rich fluid inclusion assemblage reveals that boiling at

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temperatures of 353–458 °C took place after the formation of halite-bearing fluid inclusions. Final

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stage liquid-rich fluid inclusions were characterized by low Th (160 and 327 °C) and salinity values

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(0.5 and 6.2 wt% NaCl equ.). The decreasing salinity trend of the fluid inclusions versus Th indicated

22

that meteoric water was involved in the hydrothermal solutions. Eutectic temperatures (Te) of the

23

prograde stage fluid inclusions were found to be CaCl2 dominated, while retrograde stage inclusions

24

contained different salt combinations rather than a specific salt type. The minimum trapping pressures

25

of the early stage brine fluid inclusions were calculated to be between 710 and 884 bar, while later

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stage inclusions had much lower trapping pressures between ~195 and 445 bar. The δ18O isotopes of

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prograde stage quartz, garnet, and pyroxenes are close to the composition of the hydrothermal

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solutions of magmatic sources. Moreover, retrograde stage quartz, epidote, tremolite‒actinolite, and

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calcite minerals and their equilibrated solutions were found to be highly depleted by δ18O isotopes.

30

Therefore, the fluid inclusion and stable isotope constraints suggest that the hydrothermal solutions of

31

magmatic origin were responsible for the prograde skarn stage, while a mixture of magmatic and

32

meteoric solutions were responsible for the ore formation in a shallow skarn environment.

33 34 35 36 37 38

Keywords: Dagbası Fe–Cu–Zn skarn, Skarn deposits, Fluid inclusion, Stable isotopes, Trabzon, NE Turkey

1

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1. Introduction

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The eastern Black Sea region is one of the most important mining provinces in

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Turkey, located in the Alpine Metallogenic Belt, which extends approximately 500 km in

42

length and 100 km in width in an E–W direction. Many skarn-type mineralizations in this area

43

are accompanied by massive sulfide, porphyry, epithermal, and hydrothermal type ore

44

deposits. Some of the most well-known skarn deposits are Demirköy (Artvin), Kartiba,

45

Sivrikaya (Rize), Özdil, Ögene, Dağbaşı (Trabzon), Dereli, Kirazören (Giresun), Çambaşı

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(Ordu), Eğrikar, Düzköy, Arnastal and Camiboğazı (Gümüşhane) (Fig. 1).

47 48

The geological setting of this region is especially appropriate for the formation of

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skarn-type deposits, due to the large volume of Lower Cretaceous limestones of Berdiga

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Formation and the limestone layers in an Upper Cretaceous volcano-sedimentary unit that is

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in contact with younger granitic intrusions over an extensive area (Demir et al., 2017).

52

Therefore the skarn-type deposits in this region have attracted considerable academic research

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and mining activities over two decades. Most of the previous research has focused on the

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geological and mineralogical properties, skarn and granitoid geochemistry, and alteration

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mineralogy of these skarn ores (Aslan, 1991; Hasançebi, 1993; Sipahi, 1997, 2011; Saraç,

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2003; Çiftçi, 2011; Yılmaz, 2016; Sipahi et al., 2017; Demir et al., 2017; Demir, 2019).

57

According to these studies, some of the skarns in the region formed along the contact between

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the limestones of the Lower Cretaceous Berdiga Formation and the Upper Cretaceous

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granitoids (e.g., Dağbaşı, Özdil). However, some other skarns formed along the contact

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between Upper Cretaceous volcano-sedimentary units and Upper Cretaceous‒Eocene

61

intrusions

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mineralization has been reported from Özdil, Kartiba, Çambaşı, Kotana, Arnastal,

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Camiboğazı, and Sivrikaya, whereas both exoskarn- and endoskarn-type mineralization have

64

been reported from Kirazören, Eğrikar, and Ögene (Demir et al., 2017; Demir, 2019).

65

Magnetite and hematite are the main ore minerals at Özdil, Kartiba, Camiboğazı, and

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Sivrikaya, whereas sulfide minerals accompany magnetite and hematite at Arnastal, Kotana,

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Kirazören, Eğrikar, and Dağbaşı skarn deposits (Demir et al., 2017). These studies have

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shown that the skarn-type deposits in the region show differences in terms of the host rock

69

relationship, skarn type, and mineralogical properties. However, in these studies the reason

70

for these differences in skarn types and their mineralogical compositions are not sufficiently

71

explained.

(Kirazören,

Sivrikaya,

Arnastal-Camiboğazı,

72 2

Eğrikar).

Exoskarn-type

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Fluid inclusion and stable isotope studies provide useful information on the formation

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conditions of skarn deposits, the origin of hydrothermal solutions and fluid evolutions

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(Shepherd et al., 1985; Wilkinson, 2001; Bodnar et al., 2014). With the fluid inclusion studies,

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it is possible to determine the formation conditions (temperature, pressure, depth) of the

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solutions, which were formed in the early and late skarn stages, and the compositions

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(salinity, possible salt types) of these fluids. Using stable isotope studies, it is also possible to

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determine the source of the solutions by measuring the isotope compositions of each mineral

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phase. Isotope studies are also used to calculate the formation temperatures of equilibrated

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mineral phases and to determine the isotope composition of the solutions equilibrated with

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these minerals (Chiba et al., 1989; Zheng, 1993a, 1993b; Taylor, 1997).

83 84

The fluid inclusion and stable isotope characteristics of some of these skarns in the

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northeastern region have also been investigated (Kurt, 2014; Sipahi et al., 2017; Demir et al.,

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2017). Kurt (2014) indicated that the Kirazören skarn was formed at temperatures above 360–

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400 °C by magmatic originated hydrothermal solutions, and the salinity values were higher

88

than 26.3 wt% NaCl equ. A mixture of magmatic and meteoric solutions was suggested for

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the formation of the Eğrikar skarn deposit at a temperature range of 200–425 °C and salinity

90

values between 3.9–15.4 wt% NaCl equ. by Sipahi et al. (2017). A similar temperature

91

(between 166‒462 °C) and salinity range (between 0.35–14.3 wt% NaCl equ.) was also

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determined for the Sivrikaya skarn deposit by Demir et al. (2017), and a mixture of magmatic

93

and meteoric solutions was suggested for the skarn formation. The skarns in these studies

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were formed along the contact of Upper Cretaceous volcano-sedimentary units and Eocene

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granitoids. However, the Dağbaşı skarn deposit is different from the others by the host rock

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lithology, because it was formed along the Lower Cretaceous Berdiga limestones and Upper

97

Cretaceous Dağbaşı granitoid (Demir, 2019). Therefore, studies on the Kirazören, Eğrikar,

98

and Sivrikaya skarn deposits do not represent all of the skarn formations in the region, and do

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not demonstrate the formation conditions, fluid compositions and genesis of all the skarn

100

deposits.

101 102

Four different skarn deposits were observed in the Dağbaşı area, including the İpekçili,

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Köprüüstü, Kükürtlü and Dere Mahalle locations (Fig. 2). These deposits are not currently in

104

production, and there is currently no reserve record available. However, closed old adits and

105

an amount of mine waste around them indicate that there was once historical mining

106

production. In addition to ore-host rock relations and mineralogical properties, Demir (2019) 3

107

also investigated the composition of silicate, oxide, and sulfide mineral phase and

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geochemical properties of the Dağbaşı granitoid and its relation with the skarn types.

109

However, this earlier research did not sufficiently clarify the composition and the source of

110

solutions and the formation conditions of Dağbaşı skarns. Therefore, in this present study

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fluid inclusion and C–O–H isotopic characteristics of the Dağbaşı skarn mineralization are

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described in detail in order to investigate the formation conditions, the source of hydrothermal

113

solutions and its evolutions.

114 115

2. Regional geology

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The geological structure of northeastern Turkey, which lies in the Alpine‒Himalayan

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Orogenic Belt, was formed as a result of the long-lived subduction of the Tethyan Ocean

118

(Okay and Şahintürk, 1997). Early to Middle Carboniferous high T/low to medium-P

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metamorphic units which consist of quartz‒feldispathic gneiss, schist, amphibolite, phyllite,

120

chert, marble, and metaperidotites are the oldest units in the Hercynian basement (Topuz et

121

al., 2007; Dokuz, 2011; Dokuz et al., 2015). High-K, I-type granitoids of Carboniferous to

122

Early Permian age were emplaced into the Hercynian basement throughout the Sakarya Zone

123

(Topuz et al., 2010; Dokuz et al., 2017).

124 125

Paleozoic basement was overlain, transgressively, by Lower and Middle Jurassic

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volcanics and interbedded carbonate sediments of the Şenköy Formation, which represent an

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extensional arc environment (Okay and Şahintürk, 1997; Kandemir and Yılmaz, 2009). These

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volcano-sedimentary sequences grade into the Upper Jurassic‒Lower Cretaceous carbonates

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of the Berdiga Formation (Şengör et al., 2003; Karslı et al., 2010a). The Dağbaşı Skarn ore is

130

hosted by these platform carbonates.

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The Upper Cretaceous was composed of an acidic and basic succession of volcano-

133

sedimentary units, with a thickness of more than 2 km. This volcano-sedimentary sequence

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was separated into four different formations by Güven et al. (1998) – (from bottom to top)

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Çatak, Kızılkaya, Çağlayan, and Tirebolu. The Çatak Formation, which was conformably

136

overlain by the platform carbonates of the Berdiga Formation, mainly consisted of andesite,

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basalt and their pyroclastites, with interlayered sandstone, marl, shale, red biomicrite, and

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micritic limestone. The Kızılkaya Formation is separated from the Çatak Formation by dacitic

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and rhydacitic volcanites which contain clayey, sandstone, marl and red-biomicrite

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intercalations (Güven, 1993). Prismatic columnar dacites are common in this unit, and host 4

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numerous massive volcanogenic sulfide deposits across the region (Karslı et al., 2011). The

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Kızılkaya Formation is conformably overlain by the Çağlayan Formation, which consists of

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basic volcanic rocks and interlayered thin- to medium-bedded sandstones, red micritic

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limestones, marls, clayey limestones, and tuffs. The Tirebolu Formation, which represents the

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uppermost part of the Mesozoic sequence, conformably overlies the Çağlayan Formation, and

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comprises rhyodacite and pyroclastites, with pelagic micritic limestone, sandstone, and clayey

147

intercalations (Güven, 1993; Güven et al., 1998).

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Many different intrusions were emplaced into the northeastern region of Turkey from

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Early Jurassic to Late Eocene times. These intrusions developed in various geodynamic

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settings, and have different ages and compositions. The early Jurassic plutons (between

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176.95±0.49 and 178.41±0.44 Ma) are mainly composed of gabbro and gabbroic diorite, and

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characterized by relatively low SiO2 (47.09–57.15 wt%), moderate Na2O (1.19–3.92 wt%)

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and high Mg# (46–75) (Karslı et al., 2017). These gabbroic plutons show metaluminous

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geochemical character and belong to the slightly evolved I-type signature of continental

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magmatic arc setting. Early Jurassic (188.0±4.3 Ma) plutons were also reported by Dokuz et

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al. (2010) to exhibit a low-K, metaluminous to weakly peraluminous (ASI=0.94–1.11)

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character and granodiorite and, to a lesser extent, tonalite composition. On the contrary, late

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Jurassic plutons (153+3.4 Ma) consist of mostly quartz monzodiorite composition and show a

160

metaluminous (ASI=0.84–0.99) character with a medium to high K2O, and relatively high

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MgO content (Dokuz et al., 2010). These early and late Jurassic intrusions were interpreted to

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be the products of an arc continent collision event, in response to the closure of Paleotethys

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(Dokuz et al., 2010).

164 165

The composition of the Late Cretaceous plutons varies from low-K tholeiitic through

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calc-alkaline (rarely high-K) metaluminous and peraluminous leucogranites to alkaline

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syenites and monzonites (Okay and Şahintürk, 1997; Karslı et al., 2004; Karslı et al., 2007;

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Topuz et al., 2007; Kaygusuz and Aydınçakır, 2011). The age of these plutons was

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determined as between 88.1–86.0 million years by Kaygusuz et al. (2009) by the U–Pb

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SHRIMP method conducted on the zircon minerals. Late Eocene intrusions, on the other

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hand, are made up of granite‒granodiorite, with quartz‒monzonite‒tonalite porphyry

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associations, and compositions varying from low-K tholeiitic to high-K calc-alkaline (Moore

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et al., 1980; Boztuğ et al., 2004; Karslı et al., 2010b). The 40Ar-39Ar ages of hornblende and

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biotite separates and U–Pb zircon SHRIMP ages from these rocks range between 48 Ma and 5

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54 Ma (Karslı et al., 2010b; Topuz et al., 2011; Karslı et al., 2013). According to these

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studies, Early Cretaceous to Late Eocene intrusions were formed by arc-collisional through

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syncollisional crustal thickening to post-collisional extensional regimes.

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3. Host rock geology

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The basement around the Dağbaşı area is made up of Liassic volcanic rocks. This unit

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corresponds to the Şenköy Formation, which crops out along the northeastern part of Turkey

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(Kandemir and Yılmaz, 2009). This unit consists mainly of andesite, basalt, and their

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pyroclastites and includes interlayered sandstones, marls, and red biomicrite lenses, which are

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a few meters thick and several hundred meters in length. Based on the presence of

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foraminifera (Involitina liassica, Trocholina, Lenticulina, Spirillina, Vidalina martana,

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Lingulina, and members of the Lagenidea) found in the biomicrites, a Liassic age was

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determined by Güven et al. (1998).

189 190

The andesites and basalts are macroscopically characterized by gas cavities, highly

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fractured and altered. Brecciated, glassy, microlitic, microlitic porphyritic, amygdaloidal, and

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void textures were commonly observed in the microscopic scale. This unit consists of

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pyroxene, plagioclase, amphibole, biotite, augite and quartz. While the primary quartz content

194

does not exceed ~3%, this content increases due to silicification that has occured along the

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fractures and gas cavities. The silicification is accompanied by epidote, chlorite, and calcite

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along the contact with the granitoid.

197 198

The SiO2, MgO, and K2O values were reported between 55.74–57.57, 3.24–4.44, and

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0.07–2.19 wt%, respectively in the andesites, and the same values were between 47.4–51.59,

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2.62–8.52, and 0.13–1.58 wt%, respectively in basalts (Aydınçakır, 2006). These results

201

correspond to the trachyandesite, andesite, basaltic andesite and basalt composition on the

202

alkaline-silica diagram of Le Maitre et al. (1989). By using various trace element contents,

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Aydınçakır (2006) reported that the samples show calc-alkaline basalt composition, and

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volcanic arc related basalt features. The chondrite-normalized REE patterns of the samples

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also indicated that both andesites and basalts were derived from the same sources (Kaygusuz

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and Aydınçakır, 2009).

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This volcano-sedimentary unit contains block- and lens-shaped massive limestone

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layers that are up to 100 m thick and extend up to a few kilometers. These massive bedded

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limestones are Early Cretaceous platform carbonates, identified as Berdiga Formation by

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numerous researchers (Pelin, 1977; Kırmacı et al., 1996; Yılmaz and Kandemir, 2006;

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Kırmacı et al., 2018). According to Yılmaz and Kandemir (2006), these platform carbonates

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accumulated into the rift basin under the tectonically inactive conditions. Yılmaz et al. (2008)

214

also indicated that the Berdiga Formation formed in a shallow carbonate shelf environment.

215

These limestones show limited lateral extensions since they are formed in rift basins, and

216

display different lithofacies changing vertically. The lower parts of the formation consist of

217

thin to thick bedded dolomites, while the middle and upper parts consist of very thick bedded

218

to massive, locally laminated features displaying grainstone‒packstone and skeletal

219

wackestone lithofacies (Kırmacı et al., 2018). The limestones, with thicknesses varying

220

between 10 and 30 cm, are highly recrystallized close to the granitoid contact, and their initial

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texture is unrecognizable. Crystallized limestones contain garnet, epidote, magnetite,

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hematite, pyrite, pyrrhotite, and chalcopyrite in abundance throughout the skarn zones, and

223

silicification accompanies these minerals.

224 225

The Dağbaşı Granitoid was emplaced into both the volcano-sedimentary units of the

226

Senköy Formation and the massive banded limestone blocks of the Berdiga Formation (Fig.

227

2). The granitoid exhibits zonation from the central part towards the outer zone, in terms of

228

mineral size, abundance, and composition, with large crystals in the central part and smaller

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crystals towards the edges. The granitoid is characterized by higher orthoclase abundances at

230

the central part, and these orthoclase abundances decrease towards the edge, while quartz and

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plagioclase increase. Using the modal mineralogical abundances, the granitoid was

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subdivided into four different zones by Aydınçakır (2006). In this study, Aydınçakır (2006)

233

indicated that monzogranite composition was dominant in the central parts, while the

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granitoid became granodiorite, tonalite, and diorite composition towards the outer zone. The

235

granitoid is I-type, with a low to moderate K content, peraluminum and metaluminum

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transitional, having properties similar to volcanic arc type calc-alkaline granitoids

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(Aydınçakır, 2006; Demir, 2019). The granitoid contains some mafic microgranular enclaves

238

of surrounding volcanics, centimeters to 1 m in scale, along the contact with the volcanic

239

rocks. The age of pluton emplacement has been determined as between 82.9±1.3–88.1±1.7

240

Ma (Upper Cretaceous; Kaygusuz et al., 2009) using the U–Pb SHRIMP method conducted

241

on zircon minerals. 7

242 243

The youngest unit in the study area is the Uzuntepe Dacite, which consists of mainly

244

quartz megacrystals (up to cm in size) and a lesser amount of plagioclase, biotite, amphibole,

245

and fine-grained matrix. A number of Uzuntepe dacitic intrusions cut the basic volcano-

246

sedimentary units of Şenköy Formation. Their size ranges from a few hundred meters to 3 km.

247

The largest dacite outcrop is around the Osmanoğlu village and Demdemtaş hill, which

248

extends out to 3 km in size (Fig. 2). There was no contact between the Uzuntepe dacite and

249

the Berdiga limestones in the area. However the dacite dykes, which cut the Dağbaşı granitoid

250

in the vicinity of Sariot village, indicate that these dacites are younger than the Dağbaşı

251

granitoid (Aydınçakır and Kaygusuz, 2012).

252 253

4. Skarn mineralogy

254

Skarn type mineralizations have been observed along the border of Dağbaşı granitoid

255

and Berdiga limestones around the İpekçili, Köprüüstü, Kükürtlü and Dere Mahalle districts

256

(Fig. 2). Block and lens shaped Berdiga limestones are located in the host volcanites (Fig. 3).

257

Granitoid/limestone contact is not observable in the field, because the volcanic rocks outcrop

258

in a limited area between these units. All of these skarns developed as exoskarn type in the

259

closest border of the lens- and block-shaped limestones to the granitoid.

260 261

Both prograde and retrograde stage skarn mineral assemblages have been described in

262

the location of İpekçili and Köprüüstü ore. The prograde skarn mineral assemblage was

263

subdivided into early prograde and late prograde stages by Demir (2019). The early prograde

264

stage garnets and pyroxenes were described as having massive (Fig. 4a), rhythmically-banded

265

(Fig. 4b), nodular (Fig. 4c), and granular textures (Fig. 4d), while the late prograde stage

266

garnet and pyroxenes were described as having fracture-filling texture (Fig. 4e) in the

267

limestones. Demir (2019) indicated that these early prograde garnets were grossular

268

dominated, with the composition changing to andradite in the late prograde stage; whereas

269

early stage pyroxenes were hedenbergite in composition, becoming diopside in the late stage.

270

There was no relation between the early stage garnet and pyroxene with the ore minerals, but

271

the late stage garnets accompany magnetite and hematite in their growth zones (Fig. 4f). The

272

retrograde stage mineral assemblage in the İpekçili and Köprüüstü locations was represented

273

by hydrous silicate minerals, such as epidote, chlorite, and tremolite‒actinolite, accompanying

274

quartz and calcites minerals. Quartz, epidote, and calcites were formed along the fractures of

275

the volcanic host rocks (Fig. 4g) and limestones, while tremolite‒actinolite minerals were 8

276

commonly observed as massive textures (Fig. 4h). Retrograde stage tremolite and actinolite in

277

the Köprüüstü locations were also found to contain both macroscopic (Fig. 4i) and

278

microscopic (Fig. 4j) scale garnet and pyroxene inclusions.

279 280

Prograde stage minerals assemblages of garnet and pyroxenes were not determined in

281

the Kükürtlü and Dere Mahalle locations. In these locations the retrograde stage is

282

characterized by fracture-filling type (Fig. 4k) textures of epidote, quartz and calcite minerals

283

in the volcanic host rocks and limestones. In some cases, quartz inclusions, a few cm in size,

284

were also present in both the volcanic host rock and limestones (Fig. 4l).

285 286

In addition to magnetite and hematite, sulfide mineral assemblages were also

287

determined in the İpekçili and Köprüüstü locations, whereas the Kükürtlü and Dere Mahalle

288

locations were characterized by magnetite and hematite assemblage. Ore in the İpekçili and

289

Köprüüstü locations was characterized by irregularly shaped ore masses (Fig. 5a), and banded

290

ore textures along the weak zone of the limestone layers (Fig. 5b). Fracture-filling (Fig. 5c),

291

and breccia-filling (Fig. 5d) ore types were also common in the limestones and volcanic host

292

rocks. Taking into consideration the magnetite and hematite developments in the growth

293

zones of the late prograde stage garnets, Demir (2019) surmised that ore mineralization began

294

towards the end of the late prograde stage (Fig. 4f). However, the presence of these garnets as

295

inclusions in the magnetites and hematites indicates that the main ore development took place

296

after the development of garnets (Fig. 5e). Additionally, the presence of epidote, quartz, and

297

calcite accompanying the ore minerals (Fig. 5f) indicates that the main ore mineralization

298

formed during the retrograde skarn stage.

299 300

Magnetites and hematites also accompany sulfide ore minerals. In the polished

301

sections hematites were observed to be the filling between magnetite minerals (Fig. 5g).

302

Therefore, it is understood that magnetites were formed before hematites. These magnetites

303

and hematites are always surrounded by the sulfide minerals (Fig. 5h), which represent the

304

formation of oxide before sulfides. Chalcopyrite, and sphalerite formation along the pyrite

305

cracks also indicates the formation of pyrite before chalcopyrite minerals (Fig. 5i). Apart from

306

the presence of magnetite and hematite, there is some difference in terms of the sulfide

307

minerals between the İpekçili and Köprüüstü locations. Pyrrhotite is only present in the

308

İpekçili skarn zone, and accompany pyrite, chalcopyrite and sphalerites. This pyrrhotite, in

309

the İpekçili skarn, is characterized by a bird’s eye texture (Fig. 5j) and has been replaced by 9

310

late stage sphalerite. On the other hand, galena minerals, which were rarely observed in the

311

form of small inclusions in sphalerite (Fig. 5k), accompany pyrite, chalcopyrite, and

312

sphalerite in the Köprüüstü skarn ore.

313 314

The Kükürtlü and Dere Mahalle locations were characterized by magnetite and

315

hematite masses reaching up to a meter in size (Fig. 5l), and veins formed along the fractures

316

of the limestone and volcanic host rocks. In addition to the fracture-filling type veins, breccia-

317

filling type textures were also present. In this texture, breccias of magnetite and hematites

318

were filled up by later quartz and epidote minerals (Fig. 5m). In some cases, replacement

319

remnants of magnetite and hematite were also observed in the epidote- and quartz-bearing

320

samples. Disseminated magnetites and hematites were also observed in both macroscopic and

321

microscopic scale (Fig. 5n).

322 323

The presence of magnetite and hematites in the outer growth zone of late stage garnets

324

indicates that ore mineralization in the Dağbaşı area began at the end of the late prograde

325

stage. But the presence of garnet inclusions in the magnetite and hematite ores indicates that

326

the main ore mineralization was formed later than the prograde stage. The coexistence of

327

magnetite and hematites with the epidote-bearing samples and their contemporaneous

328

development (Fig. 5f) is consistent with the retrograde stage ore formation in the Dağbaşı

329

area. Ore mineralization in the Kükürtlü and Dere Mahalle district supports the retrograde

330

stage ore development because there was no evidence of prograde stage garnet and pyroxene

331

assemblage.

332 333

Magnetite and hematites were the main ore minerals in both the Kükürtlü and Dere

334

Mahalle locations. Limited amounts of pyrite and chalcopyrite minerals, a few microns in

335

size, were observed in these minerals (Fig. 5o). Therefore, the abundance of the sulfide

336

minerals in Kükürtlü and Dere Mahalle locations was much lower than in the İpekçili and

337

Köprüüstü ores. Similar findings were also reported by Demir (2019), and the generalized

338

mineral paragenesis and succession in the Dağbaşı Skarns are presented as shown in Figure 6.

339 340

5. Analytical techniques

341

Microthermometric measurements were performed on the approximately 200 µm-thick

342

double-polished sections using a Linkam THMG-600 heating‒freezing stage at the

343

Department of Geological Engineering of Recep Tayyip Erdoğan University, Turkey. This 10

344

equipment is suitable for temperature measurements between -180 and 600 °C. Liquid

345

nitrogen was used for cooling and freezing. The stage was calibrated using H2O– and H2O–

346

CO2 containing synthetic fluid inclusions (Sterner and Bodnar, 1984). The repeated

347

experiments indicate that the measurements were performed with an accuracy of ±0.2 °C for

348

the freezing, and ±0.5 °C for the heating processes. Salinity data, expressed as wt% NaCl

349

equ., was calculated from the melting temperature of the last ice crystal (Tm-ice) using the

350

equation reported by Bodnar and Vityk (1994). In this technique, the salinity is given as wt%

351

NaCl equ., because this calculation has an error margin of less than 5 wt%, even if there were

352

any other salt types. Isochores for the pressure estimations were also computed using the

353

equations reported by Bodnar and Vityk (1994). A total of 59 double-polished wafers were

354

prepared from garnet, epidote, quartz, and calcite samples. Prior to the microthermometric

355

measurements, fluid inclusions were petrographically investigated using the criteria of

356

Roedder (1984). The measurement process was performed according to Shepherd et al.

357

(1985).

358 359

Raman spectroscopic studies were carried out at the WITec GmbH (Ulm, Germany).

360

A 514 nm (green) laser mounted on the Alpha 300 confocal Raman spectrometer was used for

361

the measurements. The entire area of each of the fluid inclusions was scanned, using the 1 µm

362

intervals of the laser beam. Data was collected over a spectral range of 100 to 4000 cm-1.

363 364

A total of 22 samples from the skarn calcites were subjected to δ13C and δ18O isotope

365

analyses at the Cornell Stable Isotope Laboratory, USA, using a Thermo Scientific Delta V

366

Advantage mass spectrometer coupled to a Thermo Scientific Gas Bench II. A sample of

367

approximately 0.5–1 mg of calcite was reacted with phosphoric acid at 25 °C for 18 hours to

368

produce CO2. The produced CO2 gas was transferred into the isotope ratio mass spectrometer

369

and analyzed for its 13C/12C, and 18O/16O ratios. The laboratory stated that the precision of the

370

analyses was less than 0.20‰ in this method. The δ13C and δ18O isotope results were reported

371

relative to the Pee Dee Belemnite (PDB) standard. The results of the PDB standard were

372

converted to the Standard Mean Ocean Water (V–SMOW) standard using the empirical

373

relationship (Δ18OSMOW =1.03086*Δ18OPDB+30.86) of Friedman and O’Neil (1977).

374 375

The δ18O and δD analyses were carried out on the selected garnet, pyroxene, epidote,

376

quartz and tremolite samples. Before analysis, each sample was crushed to approximately 1

377

mm in size, and pure minerals were hand-picked using a binocular microscope. The carbonate 11

378

was removed with a treatment of 1 N HCl solution. The samples were cleaned with pure

379

water, and the purity was approved by X-ray diffraction.

380 381

The isotope analyses were performed at the Scottish Universities Environmental

382

Research Centre. In this laboratory, oxygen is extracted from a 1–2 mg sample by the CIF3

383

fluorination technique. The extracted oxygen was converted to CO2 with a reaction by hot

384

graphite rod. The isotope composition of the evolved CO2 was measured by a VG PRISM III

385

isotope ratio spectrometer. The results were reported relative to the V–SMOW standard. The

386

precision was ±0.2‰ (1σ) based on repeated analyses.

387 388

6. Fluid inclusion studies

389

Fluid inclusion studies were conducted on garnet, epidote, quartz, sphalerite, and

390

calcite minerals to determine their formation temperatures and the conditions under which

391

mineralization

392

microthermometric measurements, were commonly observed in quartz, calcite, and sphalerite

393

minerals, whereas fluid inclusions in garnet and epidote minerals were limited. Prior to the

394

freezing‒heating experiments, a genetic classification of the fluid inclusions in each crystal

395

was performed, according to the criteria suggested by Roedder (1984) and Shepherd et al.

396

(1985). Fluid inclusions randomly distributed in the host crystal or aligned parallel to the

397

grain boundaries are considered as primary, and inclusion trails occurring as planar groups

398

and askew to crystal boundaries are regarded as of secondary origin. Microthermometric

399

measurements were carried out on inclusions with a primary origin so as to avoid potential

400

post-mineralization effects in the secondary inclusions.

occurred

in

the

Dağbaşı

Skarns.

Fluid

inclusions,

suitable

for

401 402

Liquid-, solid-, and vapor-bearing (L+S+V) three-phase, and liquid- and vapor-bearing

403

(L+V) two-phase fluid inclusions were identified in the researched minerals. The fluid

404

inclusions were commonly polygonal and irregular in shape, while lesser amounts of rounded,

405

elliptical, tube- and pear-shaped morphologies were also present. In some cases, tabular-

406

shaped inclusions were also observed in calcite minerals. The size of the investigated fluid

407

inclusions generally varied between 10 and 60 µm. Inclusions larger than 30 µm in size were

408

rarely observed in the investigated minerals. None of the fluid inclusions contain separated

409

liquid CO2 phase at room temperature.

410

12

411

Although there are four different skarn locations around the Dağbaşı granitoid, no

412

decisive diversity was observed between these locations, in terms of the inclusion shape, size,

413

or microthermometric results. The results are presented with consideration to the inclusion

414

type and vapor/liquid ratio in each mineral phase, since the main differences vary according to

415

these properties. In terms of the number and volumetric proportions of the phases present at

416

room temperature, and the homogenization type of the fluid inclusions (to a liquid or vapor

417

phase), three major types of fluid inclusions were recognized from the Dağbaşı skarn ore. The

418

detailed properties of each type of inclusion are described below.

419 420

6.1.

Classification

421

Type I fluid inclusions are three-phase (L+S+V) inclusions. This type of inclusion was

422

identified in the early prograde stage garnet (Fig. 7a), quartz (Fig. 7b), and calcite (Fig. 7c)

423

minerals. Two-phase fluid inclusions (L+V), with vapor ratios higher than liquid phase

424

(V>L), were classified as Type II. Homogenization always occur to vapor phase in this type

425

of inclusions. These were identified only in late prograde stage garnet (Fig. 7d) and quartz

426

(Fig. 7e) minerals. Type III inclusions are two-phase (L+V), with vapor ratios always lower

427

than unity. These inclusions comprised three subgroups. Type IIIa have vapor bubbles that

428

occupy 30–40 vol% of the inclusions. Homogenization to the liquid phase occurs in these

429

types of inclusions. Type IIIa fluid inclusions were identified in the late prograde stage of

430

garnet and quartz minerals and were found to accompany Type II inclusions, even in the same

431

area of host crystals (Fig. 7f, g). Therefore, both Type II and Type IIIa inclusions were

432

grouped as a fluid inclusion assemblage. Type IIIa inclusions, in garnet and quartz minerals,

433

homogenized to liquid phase, whereas the Type II inclusions homogenized to the vapor phase

434

in the same minerals.

435 436

Type IIIb inclusions were observed in the retrograde stage epidote, calcite, sphalerite,

437

and quartz minerals (Fig. 7h‒j). Type IIIa and Type IIIb inclusions have similar shape, size

438

and liquid/vapor ratios. However, there was no coexistence of Type IIIb and Type II fluid

439

inclusions. Type IIIb inclusions were also characterized by their lower Th and salinity values

440

compared to the Type IIIa inclusions. In some cases, Type IIIb inclusions were observed to

441

participate in Type IIIc inclusions in the quartz minerals (Fig. 7k). Type IIIc inclusions are

442

characterized by a volumetrically smaller vapor/liquid ratio than in Type IIIa and Type IIIb

443

(Fig. 7k, l) that accounts for ~5–20 vol% of the inclusions, and this type of inclusion was

13

444

observed in all the investigated minerals of retrograde stage. Type IIIc fluid inclusions are

445

characterized by lower homogenization temperature (Th) and salinity values.

446 447

6.2.

Microthermometric results from the fluid inclusions

448

6.2.1. Eutectic temperatures

449

The eutectic temperatures (Te) of different types of inclusions in all the investigated

450

minerals were measured, and the results are presented in Table 1. The Te of the three-phase

451

Type I inclusions in the garnet, quartz, and calcite minerals was between -53.2 °C and -46.6

452

°C. For the gas-rich Type II and liquid-rich Type IIIa inclusions in late prograde stage garnet,

453

and quartz, Te temperatures were in the range of -57.4 °C to -34.6 °C, and -57.2 °C to -32.4

454

°C, respectively. Similar Te temperatures were measured in the fluid-rich Type IIIb inclusions

455

between -58.6 °C to -29.0 °C in the quartz, epidote and calcite minerals (Fig. 8).

456 457

When the Te of the Type I inclusions are compared to the eutectic temperatures of

458

various water–salt systems, this temperature range corresponds to the H2O–CaCl2 dominated

459

fluids. On the contrary, Te of the Type II, Type IIIa, and Type IIIb inclusions corresponded to

460

mixtures of different salt combinations rather than specific salt solutions. For instance, Te

461

temperatures close to -49.8 °C corresponded to the H2O–CaCl2 dominated solutions, while the

462

Te temperature range of -32.4 °C to -35 °C corresponded to the H2O–FeCl2–MgCl2 system.

463

Moreover, eutectic temperatures close to -56.6 °C corresponded to the melting temperature of

464

CO2 phase, rather than specific water–salt combinations. Although there were no separated

465

gas phases and clathrate occurrences in the fluid inclusions at room temperature, the values

466

approximating -56.6 °C indicated that a limited amount of dissolved CO2 may be present in

467

the composition of the gas- and fluid-rich inclusions.

468 469

The Te temperatures of the Type IIIc inclusions were measured between -35.2 °C and

470

-20.6 °C. Most of the composition of these inclusions, ranging between -28 °C and -20.6 °C,

471

was close to the eutectic temperatures of the H2O–NaCl system. Conversely, the composition

472

of a very few Type IIIc inclusions corresponded to the mixture of a H2O–FeCl2–MgCl2

473

system, with a Te range of -29 °C to -35.2 °C (Fig. 8).

474 475

6.2.2. Last ice melting, and halite melting temperatures and salinity values

476

The ice melting temperature (Tm-ice), which was used to calculate the salinity values

477

of the inclusions, was measured in the two-phase (L+V) Type II, Type IIIa, Type IIIb, and 14

478

Type IIIc inclusions. The halite melting temperature (Tm-halite) was measured in halite

479

bearing Type I inclusions to calculate salinity values. The Tm-ice and Tm-halite values,

480

measured from the different inclusion types, are presented in Table 1.

481 482

The halite melting temperatures of the brine (Type I) fluid inclusions were measured

483

in the range of 412 °C to 514 °C, and the salinity values of these inclusions was calculated

484

between 48.8 and 61.8 wt% NaCl equ. according to Bodnar and Vitky (1994) (Fig. 9). Tm-ice

485

temperatures of the Type II fluid inclusions were measured between -6.8 °C and -9.6 °C in

486

garnet, and -5.3 °C and -10.5 °C in quartz minerals. Accordingly, the salinity of these

487

inclusions was calculated between 10.3 and 13.7 wt% NaCl equ. in garnet, and 8.3 and 14.7

488

wt% NaCl equ. in quartz minerals. The salinity values in the Type II and Type IIIa fluid

489

inclusions in the garnet and quartz minerals were similar. But the salinity values of Type IIIb

490

inclusions in quartz, calcite, and epidote minerals were distinctively lower than the Type II

491

inclusions, with calculated values in the range of 4.8–10.1, 2.9–6.9, and 4.1–9.8 wt% NaCl

492

equ., respectively. The Type IIIc inclusions, observed in the retrograde stage of all

493

investigated minerals, had lower salinity values, in the range of 0.9–6.2, 0.5–4.8, and 1.1–5.1

494

wt% NaCl equ. in epidote, calcite, and quartz minerals, respectively. The salinity values of the

495

Type IIIb and Type IIIc inclusions in sphalerite minerals were very close to each other,

496

ranging between 0.7 and 5.1 wt% NaCl equ. (Fig. 9).

497 498

Taking into consideration the measured Tm-ice and Tm-halite temperatures, the

499

highest salinity values occurred in the Type I inclusions (61.8 wt% NaCl equ.), whereas the

500

lowest values were in the Type IIIc inclusions (0.3 wt% NaCl equ.). The variations in Figure

501

9 indicate a clearly decreasing trend between salinity and Th for all measured inclusions.

502 503

6.2.3. Homogenization temperature measurements

504

The homogenization temperatures (Th) of different types of fluid inclusions are given

505

in Table 1. Two different homogenization styles were observed in halite-bearing Type I

506

inclusions – melting temperatures of halite crystals (Tm-halite) and homogenization to liquid

507

phase by vapor disappearance (Th). Tm-halite was always higher than Th in these inclusions.

508

Homogenization of the vapor into the liquid phase occurred at between 352.5 °C and 368.6 °C

509

in garnet, 357 °C and 402 °C in quartz, and 348 °C and 446 °C in calcite minerals, but the

510

Tm-halite in these minerals ranged between 412 °C and 418 °C in garnet, 438 °C and 461 °C

511

in quartz, and 492 °C and 514 °C in calcite minerals (Table 1). 15

512 513

The homogenization temperatures of the Type II fluid inclusions, that homogenized to

514

vapor phase, were in the range of 355 °C–449 °C in garnet and 355.2 °C–452 °C in quartz

515

minerals. The Th of Type IIIa inclusions, which coexist with Type II inclusions in both

516

minerals, are in the range of 353 °C–458 °C in garnet, and 368 °C–451.2 °C in quartz. The

517

coexistence of both types of inclusions in the same fluid inclusion assemblage, with their

518

similar Th and salinity values, clearly indicates a boiling assemblage for both types (Bodnar et

519

al., 1985). The Th of Type IIIb inclusions in quartz, epidote, and calcite, which have no

520

coexistence with Type II inclusions, have a slightly lower Th range of 318 °C–412 °C.

521

The homogenization temperatures of Type IIIc inclusions in quartz, calcite, and

522

epidote minerals were in the range of 160 °C–327 °C, whilst Th of the Type IIIc inclusions in

523

all these minerals were much lower than the other fluid inclusion types. The homogenization

524

temperatures of the Type IIIb and Type IIIc inclusions in sphalerites were also close, between

525

231 °C and 327 °C. Although these two types of fluid inclusions occur in sphalerite, the Th

526

histogram of both types of fluid inclusions were represented by a unimodal distribution (Fig.

527

10) due to the close Th ranges of both types of inclusions.

528 529

6.3.

Pressure estimates

530

Estimates of the trapping pressure of the halite-bearing Type I fluid inclusions, which

531

homogenized by halite dissolution, can be calculated as a function of the liquid‒vapor Th and

532

Tm-halite, according to Sanchez et al. (2012). In this technique, after the calculation of the

533

salinity value of halite-bearing fluid inclusions, the halite liquidus line is drawn for the

534

calculated salinity value. The Th isochore for the measured Th temperature is also drawn by

535

the calculated salinity. The intersection of Th isochore with the halite liquidus gives the

536

minimum trapping pressure of each fluid inclusion. Based on this technique, an 884 bar

537

trapping pressure was calculated from the intersection of the halite liquid line, drawn from the

538

highest Tm-halite (514 °C) and Th isochore (446 °C) (Fig. 11). Accordingly, a 710 bar

539

trapping pressure was calculated from the intersection of the halite liquid line, from the lowest

540

Tm-halite (412 °C) and Th isochore (352 °C).

541 542

The presence of both vapor- and fluid-rich inclusions in the same FIA, the

543

homogenization of vapor-rich inclusions to vapor phase, and fluid-rich inclusions to liquid

544

phase, and their similar Th and salinity ranges (Table 1), are accepted as evidence of boiling

545

(Wilkinson, 2001; Bodnar, 2003). If boiling occurred, the homogenization of Type II and 16

546

Type IIIa inclusions would show on the two-phase critical curve (Fig. 11) according to Baker

547

and Lang (2003) and Sanchez et al. (2012). In this case, the measured pressure is equal to the

548

actual pressure, and no pressure correction was required (Roedder and Bodnar, 1980;

549

Wilkinson, 2001; Sanchez et al., 2012). Therefore, the minimum and maximum Th

550

temperatures of both the Type II and Type IIIa fluid inclusions, between 353 °C and 458 °C,

551

corresponded to the trapping pressure between 195 and 445 bar on the two-phase critical

552

curve (Fig. 11), respectively.

553 554

Type IIIb inclusions are characterized by moderate Th (318–412 °C, average 362 °C)

555

and salinity values (2.9–10.1 wt% NaCl equ., average 6.1 wt% NaCl equ.). The trapping

556

pressure for this type of inclusion can only be obtained where independent temperatures are

557

available. However, there were no independent temperatures for the calculation of the

558

trapping pressure of this kind of inclusion. Instead, the trapping pressure of Type IIIc fluid

559

inclusions, which were characterized by lower Th (160 °C–327 °C, average 240 °C) and

560

salinity values (0.5–6.2 wt% NaCl equ., average 3.2 wt% NaCl equ.), were calculated by the

561

independent temperatures. In this technique, the trapping pressure of the fluid inclusions was

562

defined by the intersection of independent temperatures with the fluid inclusion isochore,

563

which were drawn by using the average salinity and Th of fluid inclusions (Roedder and

564

Bodnar, 1980; Brown and Lamb, 1989). In this study, independent temperatures were

565

calculated between 135 °C and 281 °C from the retrograde stage quartz‒epidote, and

566

quartz‒tremolite mineral pairs. The highest trapping pressure of 310 bar was determined by

567

the intersection of the P–T isochron with the highest independent geothermometer of 281 °C.

568 569

6.4.

Raman studies

570

None of the fluid inclusions were found to contain separated gas phases, and clathrate

571

occurrences were also not observed during the microthermometric study. Although this

572

suggests that CO2 was not present in the composition of the fluid inclusions, the Te

573

temperatures of some fluid inclusions were very close to the melting temperature of CO2,

574

indicating that these inclusions may contain some dissolved CO2. Hedenquist and Henley

575

(1985) and Rosso and Bodnar (1995) indicated that gas species may not be detected by

576

microthermometric study if the bulk density is lower than 2 mol.%. Therefore, Raman

577

spectroscopy measurements were performed on representative fluid inclusions to confirm

578

whether there was evidence of separated CH4, CO2, or any other volatiles. No significant

579

spectra were obtained from the fluid inclusions in sphalerite, due to the absorption of the 514 17

580

nm laser wavelength, or in calcite, due to the fluorescence effect; however, meaningful

581

Raman measurements were achieved on the fluid inclusions of quartz, epidote, and garnet

582

minerals. The measurements were performed by the whole-area scanning technique because

583

there were no separated gas phases in the fluid inclusions.

584 585

According to the Raman measurements, the majority of the fluid inclusions had an

586

H2O composition that contained various dissolved salts, based on the 3100–3650 cm-1 band

587

range of the spectrum. It was not possible to determine the dissolved salt species in the fluid

588

inclusions, because the dissolved salt species in the solutions did not have characteristic peaks

589

in the Raman spectrum. They only affected the distribution range or intensity of the Raman

590

signals (Burke, 2001; Bakker, 2004; Baumgartner and Bakker, 2009; Frezzotti et al., 2012).

591

Despite that, Type IIIa inclusions in garnet (Fig. 12a), epidote (Fig. 12b), and quartz (Fig.

592

12c, d) were found to rarely contain CH4 according to the characteristic CH4 peak of spectra

593

at around 2917 cm-1. It has been confirmed from other studies that the given peak values

594

correspond to CH4 species in the fluid inclusions (Burke, 2001; Bakker, 2004; Lin et al.,

595

2007; Lin and Bodnar, 2010). These CH4 species were only detected in three of 18 measured

596

samples.

597 598

7. Carbon and oxygen isotope studies

599

In this study, δ13C and δ18O isotope analyses were performed on the prograde and

600

retrograde stage skarn calcites from the Köprüüstü, İpekçili, and Dere districts. These results

601

are listed in Table 2. The δ13C isotopic values of the prograde stage skarn calcites were

602

measured between -1.93 and 2.88‰, while retrograde stage skarn calcites were between -0.81

603

and 1.95‰ (Fig. 13a). The δ18O isotopic compositions of the prograde and retrograde stage

604

skarn calcites were measured between 9.02–16.18‰, and 7.33–9.72‰, respectively.

605 606

The δ13C isotope values of both prograde and retrograde stage skarn calcites, in the

607

range of -1.93 and 2.88‰, overlap both magmatic (0 to -10‰) and marine fields (-3 to 3‰,

608

Clarke and Fritz, 1997; Bowman, 1998; Hoefs, 2009). On the other hand, the δ18O isotope

609

compositions of these calcites range between the magmatic rocks and marine carbonates (Fig.

610

13a). The range of these values in Fig. 13a indicates that skarn calcites are considerably

611

depleted compared to the average δ18O and δ13C isotope compositions of the Berdiga

612

limestones (28.4‰SMOW and 0.6‰PDB, respectively, Kırmacı et al., 2018).

613 18

614

There is a depletion trend in the variations in the δ13C and δ18O isotopes of skarn

615

calcites, which follows a typical fluid‒rock interaction pattern (Bowman et al., 1985; Shin and

616

Lee, 2003). Therefore, the effect of fluid‒rock interaction on the decreasing δ13C and δ18O

617

isotopes was evaluated with the model of Taylor (1976). In this model the fluid is assumed to

618

be isotopically equilibrated with the rock in the closed system. The isotopic fractionation

619

factor was calculated according to the mineral‒water equation of Zheng (1993a). The initial

620

δ18O and δ13C isotope compositions of Berdiga limestones reported by Kırmacı et al. (2012)

621

were used for the calculation of the final δ18O and δ13C isotopes of skarn calcites. The M–A

622

mixing curves were calculated for varying X(CO2) and fluid rock ratios of 0.05, 0.2, 0.5 and

623

0.5, 2, 10, 40, 80, 100, respectively (Fig. 13b). The highest Th-halite temperature of halite-

624

bearing skarn calcites (514 °C) was used for the calculations.

625 626

The δ13C and δ18O isotopes measured from the prograde and retrograde stage skarn

627

calcites of Dağbaşı deposits rather follow the depletion curves of 0.05 and 0.1 X(CO2). The

628

isotopic depletion is shown in both oxygen and carbon isotopes. The fluid‒rock ratios of

629

prograde stage skarn calcites are highly variable and they range between 0.5 and 80. On the

630

contrary, fluid‒rock ratios of retrograde stage skarn calcites are in the range of 10–40.

631 632

8. Oxygen and hydrogen isotope studies

633

In order to clarify the metasomatic processes between the Dağbaşı granitoid and the

634

limestones, representative garnet, pyroxene, quartz, epidote, and tremolite minerals, from the

635

prograde and retrograde skarn stages of Köprüüstü and İpekçili, were analyzed for δ18O

636

isotopes. The results are given in Table 3. In addition, δ18O analysis of quartz and plagioclase

637

from the central part of the granitoid that was not affected by the skarn zones was performed

638

in order to compare with the skarn zones.

639 640

The δ18O isotope results of the quartz and plagioclase minerals from the central part of

641

the Dağbaşı Granitoid ranged from 11.8 to 12.4‰. These values are compatible with the

642

isotopic composition of magmatic waters given by published studies (Taylor, 1979; Hoefs,

643

2009). When the δ18O values of the skarn minerals are considered, the results of the prograde

644

skarn stage of quartz, pyroxene, and garnet range between 8.2 and 10.3‰, 1.9 and 5.3‰, and

645

3.0 and 4.8‰, respectively. The δ18O values of the retrograde stage were between 5.8–7.2‰

646

for quartz, 0.5–0.6‰ for tremolite, and -1.9 and -0.6‰ for epidote (Table 3). These values

647

indicate that the prograde stage skarn minerals were depleted when compared to the center of 19

648

the granitoid, and retrograde stage skarn minerals were also further depleted in comparison to

649

those at the center of the granitoid and the prograde stage skarn minerals. Therefore, the δ18O

650

isotope values show a gradually decreasing trend from the center of the granitoid to the

651

retrograde skarn stage (Fig. 14).

652 653

The δD isotope analyses were carried out on the retrograde stage epidote and tremolite

654

minerals from İpekçili and Köprüüstü in order to determine the origins of the hydrothermal

655

solutions that were effective in the formation of the skarns in the Dağbaşı area. The results are

656

given in Table 3. According to this, the δD isotope values for the epidote ranged between -52

657

and -84‰, and tremolite minerals between -49 and -63‰. The variation in the δD isotope

658

results are similar to those of magmatic water, but the variation in δ18O isotopes measured

659

from these minerals was highly depleted when compared to magmatic sources. This data,

660

which provides a distribution between the magmatic and the meteoric water line in Fig. 15,

661

can be explained by dilution with meteoric water, as found in various studies (Craig, 1961;

662

Taylor, 1974; Sheppard, 1984; Hoefs, 2009).

663 664

8.1.

Isotopic equilibration and thermometer

665

In this present study, the equilibrium temperatures of quartz‒plagioclase,

666

quartz‒pyroxene, quartz‒garnet, garnet‒pyroxene, quartz‒epidote and quartz‒tremolite

667

mineral pairs were calculated, and the results are given in Table 3. Fluid inclusion results and

668

calculated temperatures which were reported in previous studies (Kaygusuz and Aydınçakır,

669

2011) were used as independent thermometers to provide the isotopic composition of the

670

equilibrated solutions.

671 672

The equilibrium temperature of quartz and plagioclase, measured from the center of

673

the granitoid, was 998.3 °C, using the equilibrium equations of Zheng (1993a). By using the

674

mineral‒water equilibrium equations for quartz‒water and plagioclase‒water, the δ18O

675

composition of the solutions in equilibrium with quartz and plagioclase were calculated to be

676

10.11‰ and 11.44‰, respectively. The average temperature (827 °C) given by Kaygusuz and

677

Aydınçakır (2011) was used as an independent thermometer for the calculation of the isotopic

678

composition of the solutions. Both the measured isotopic compositions of the minerals

679

(between 11.8 and 12.4‰) and the calculated isotope compositions of the equilibrated

680

solutions (10.11 and 11.44‰) are characteristically distributed in the range of primary

681

magmatic water (Hoefs, 2009). 20

682 683

The δ18O isotope composition of the hydrothermal solutions, equilibrated with the

684

prograde stage garnet, pyroxene, and quartz, was calculated to be between 3.68 and 7.73‰ for

685

the average Th temperatures (456 °C) of the Type I fluid inclusions (equation of Zheng,

686

1993a). The δ18O isotope composition of the retrograde stage solutions, equilibrated with the

687

epidote, tremolite and quartz minerals, was also calculated between -0.68 and 2.53‰ for the

688

average Th temperatures (~321 °C) of the Type IIIa and Type IIIb fluid inclusions in all

689

investigated minerals, according to the equation provided by Zheng (1993b).

690 691

The equilibrium temperatures calculated between the quartz‒pyroxene, quartz‒garnet

692

and garnet‒pyroxene mineral pairs of the prograde stage from both Köprüüstü and İpekçili

693

skarns ranged from 374.9 to 459.2 °C (Table 3), with the exception of a single lower

694

equilibrium temperature of 163.8 °C. Additionally, the equilibrium temperatures between

695

quartz‒epidote, and quartz‒tremolite mineral pairs, representing the retrograde stage at the

696

same locations, were calculated at between 135.8 and 281.5 °C. These calculated equilibrium

697

temperatures are very close to the Th range of fluid inclusions measured from retrograde stage

698

quartz and epidotes.

699 700

9. Discussion

701

9.1.

702

Skarn type deposits typically occur along the contact between igneous intrusions and

703

carbonate host rocks. Therefore a variety of metasomatic processes involving fluids of

704

magmatic, metamorphic, meteoric, and marine origin were suggested for the skarn formations

705

(Einaudi and Burt, 1982; Kwak, 1986; Baker and Lang, 2003; Meinert et al., 2005). In these

706

studies, high salinity fluid inclusions (>50 wt% NaCl equ.) were generally linked to magmatic

707

sources. Moreover, Wilkinson (2001) indicated that these salinity values reach up to 70 wt%

708

NaCl equ. for the skarn type deposits. Therefore, the high salinity Type I fluid inclusions (in

709

the range of 48.8 to 61.8 wt% NaCl equ.), measured from early prograde stage garnet, quartz

710

and calcites, may refer to magmatic sources.

Origin of the hydrothermal fluids

711 712

However, it is necessary to evaluate the fluid inclusion results together with the

713

isotopic data in order to determine the source of hydrothermal solutions. The δ18O isotope

714

compositions of the prograde stage garnet, pyroxene, and accompanying quartz were

715

measured at between 1.9 and 10.3‰. The isotopic compositions of the hydrothermal solutions 21

716

equilibrated with these early stage minerals were calculated at between 3.68 and 7.73‰

717

according to the mineral‒water equation of Zheng (1993a), based on the independent

718

thermometer of 456 °C, which is the average Ths of the prograde stage (Type I) fluid

719

inclusions (Table 1). Both the δ18O isotope compositions of the skarn minerals and the

720

equilibrated hydrothermal solutions are predominantly consistent with magmatic sources

721

(Taylor and Shepherd, 1986). This is because δ18O isotopes of the garnet, pyroxene and

722

accompanying quartz in the range of 4–9‰ are accepted as an indication of their derivation

723

from magmatic sources (Bowman, 1998; Meinert et al., 2005).

724 725

A magmatic origin for the early stage skarn development is further supported by the

726

δ13C and δ18O isotope ratios of the skarn calcites. This is because the range of the δ18O

727

isotopes for the skarn calcites between magmatic rocks and marine carbonates suggests

728

magmatic originated hydrothermal solutions (O’Neil, 1977; Sheppard, 1984; Hoefs, 2009,

729

Fig. 13a). Besides, both the δ18O and δ13C isotopes show an isotopic depletion trend, which is

730

very consistent with the water‒rock interaction model (Taylor and O’Neil, 1977; Bowman et

731

al., 1985; Holness, 1997; Bowman, 1998; Shin and Lee, 2003). On this model, higher

732

fluid‒rock ratios of the prograde stage calcites (up to 80; Fig. 13b) were associated with

733

magmatic solutions in similar studies (Bowman, 1998; Shin and Lee, 2003; Vallance et al.,

734

2009; Oyman, 2010; Orhan et al., 2011). In these studies, on the other hand, fluid‒rock ratios

735

of retrograde stage skarn calcites at around 10–40 (Fig. 13b) were attributed to a mixture of

736

magmatic and meteoric solutions.

737 738

The lower salinity values of the retrograde stage Type IIIc fluid inclusions and their

739

decreasing salinity trend versus Th are clearly associated with the mixture of magmatic and

740

meteoric solutions (Fig. 9). The δ18O isotope compositions of epidote, tremolite, and

741

accompanying quartz which belonged to the retrograde stage were extensively depleted,

742

varying between -1.9 and 7.2‰, when compared to the magmatic-dominated prograde stage.

743

Accordingly, water in equilibrium with these late stage skarn minerals at temperatures of 321

744

°C has δ18O isotope values between -0.68 and 2.53‰ (Table 3), which is consistent with

745

predominantly meteoric solutions (Taylor, 1979).

746 747

9.2.

Formation conditions

748

The Ths of the prograde stage fluid inclusions in the garnet, quartz, and calcite

749

minerals were in the range of 412 and 514 °C (Table 1). The equilibrium temperatures of the 22

750

garnet–pyroxene, garnet–quartz, and pyroxene–quartz mineral pairs were also in the range of

751

375–451 °C, 399–459 °C, and 388 °C, respectively. Both the equilibrium temperatures of

752

stable isotope data and the Ths from the fluid inclusions range in close proximity to the

753

prograde skarn stage. The trapping pressure of this stage was calculated to be between 710

754

and 884 bar, based on the minimum and maximum salinity values of 48.8 wt% NaCl equ. and

755

61.8 wt% NaCl equ., respectively (Fig. 11).

756 757

The presence of brine-rich (Type I) and vapor-rich (Type II) fluid inclusions may

758

suggest immiscible entrapment; however, there is some evidence against immiscible fluids.

759

This evidence includes: 1) there are none coexisting within the same fluid inclusion

760

assemblage; 2) the Ths for Type I and Type II inclusion types differ greatly; and 3)

761

homogenization of the brine inclusions occurred via halite disappearance (Roedder and

762

Bodnar, 1980; Shepherd et al., 1985). In any case, coexisting vapor-rich inclusions with the

763

halite-bearing inclusions, and homogenization by halite dissolution, cannot represent an

764

immiscible pair because phase equilibrium does not permit the coexistence of both types of

765

inclusions (Bodnar, 2003). In addition, salinity values of the vapor-rich inclusions of up to

766

14.7 wt% NaCl equ. indicate that these inclusions cannot be immiscible pairs, but rather they

767

are a product of different stages. Bodnar et al. (1985) indicated that, in the case of the

768

coexistence of brine- and vapor-rich inclusions, the salinity of these vapor-rich inclusions

769

never exceeds 1 wt% NaCl equ.

770 771

On the other hand, the presence of both vapor-rich (Type II) and liquid-rich (Type

772

IIIa) inclusions in the same FIA likely indicates boiling. Similar Ths from vapor-rich and

773

liquid-rich inclusions of 355–452 °C, and 353–458 °C, respectively may support localized

774

boiling (Roedder and Bodnar, 1997). In the case of boiling, the homogenization of Type II

775

and Type IIIa inclusions would present on the two-phase critical curve. Taking into

776

consideration a minimum (353 °C) and maximum (458 °C) Th of this boiling fluid inclusion

777

assemblage, the trapping pressure is estimated to be between 195 and 445 bar, respectively

778

(Fig. 11).

779 780

The average salinity and Th temperatures of the Type IIIc fluid inclusions in the

781

retrograde stage quartz, epidote, and calcite minerals were calculated as 3.2 wt% NaCl equ.,

782

and 240 °C, respectively. The trapping pressure of these retrograde stage skarn minerals was

783

calculated as 310 bar based on the highest independent thermometer of 281 °C (Fig. 11). 23

784

However, the lower limit of the trapping pressure could not be calculated because the

785

minimum independent temperatures of 136 °C were lower than the average Th isochore of

786

this type of fluid inclusion in Fig. 11.

787 788

The calculated pressure values for the prograde skarn stage were around 884 bar, but

789

decrease as low as 195 bar through later stages. The calculated maximum pressure of the fluid

790

inclusions corresponded to the 3.3 km thick stratigraphic units overlying the mineral deposit

791

(assuming that the density of overlying units was 2.7 gr/cm3). On the other hand, a thickness

792

of 0.7 km was calculated based on the minimum pressure. Such a sudden change in the skarn

793

environment is unreasonable. However, Yardley and Lloyd (1995) indicated that a

794

metasomatic process in the skarn front caused the increasing hydraulic pressure and resulted

795

in the collapse of the skarn system when the hydraulic pressure exceeded the overlying

796

stratigraphic units. Such a collapse of the skarn system is a response to hydrofracturing and

797

brecciation which changes in permeability and porosity (Dipple and Gerdes, 1998).

798

Additionally, Clechenko and Valley (2003) stated that hydrofracturing and brecciation can

799

cause the mixture of meteoric waters in the skarn system. Such a mixture of meteoric water

800

may explain the boiling event in Dağbaşı skarn. Meinert et al. (2005) also pointed out that

801

fractured and brecciated host rocks in the skarn front refer to the shallow emplacement of

802

intrusion and skarn formation.

803

These statements indicate that the prograde stage of Dağbaşı skarn occurred in higher

804

hydraulic pressure conditions. Increasing hydrofracturing and brecciation resulted from the

805

mixture of meteoric water and boiling through later stages. The lower limit of the pressure

806

values calculated for the retrograde skarn stage indicates that the Dağbaşı skarns were formed

807

under shallow conditions. The similar emplacement depth of the Dağbaşı granitoid, which

808

was calculated at between 0.3 and 8 kbar by Kaygusuz and Aydınçakır (2011), is quite

809

compatible with the Dağbaşı skarns. Besides, the commonly observed ore textures of the

810

stockwork veining, brittle fracturing, and brecciation, as well as intensive hydrothermal

811

alteration in the Dağbaşı skarns, were considered as indicators of the shallow depths in

812

pioneering studies (Einaudi and Burt, 1982; Meinert et al., 2005), and highly compatible with

813

the shallow skarn formation.

814 815

9.3.

Source of methane-bearing fluid inclusions

816

On the basis of the petrographic and microthermometric study, there were no methane

817

and CO2 gas phases in the fluid inclusions. However, a minor amount of CH4 was detected in 24

818

the Type IIIa fluid inclusions by the Raman spectroscopy. Similar findings were also reported

819

by Hedenquist and Henley (1985) and Rosso and Bodnar (1995), that if the bulk density is

820

less than 2 mol% in the fluid inclusions these gas species may not be detected. The CH4 is not

821

present in all Type IIIa inclusions, because only three of the 18 measured Type IIIa

822

inclusions, which always coexist with Type II fluid inclusions of the boiling assemblage, were

823

found to contain CH4 phases.

824 825

Methane and trace concentrations of other hydrocarbons have been widely reported in

826

metamorphic, ultramafic, and igneous hosted ore deposits and hydrothermal systems (Welhan,

827

1988; Polito, 1999; Whiticar, 1999; Charlou et al., 2002; Fan et al., 2004; Phillips and Powel,

828

2010). According to Welhan (1988), the possible sources of methane in hydrothermal systems

829

are: (1) thermal degradation of hydrocarbons; (2) biological production; (3) outgassing of

830

juvenile carbons as CH4; and (4) inorganic synthesis in reactions at higher temperatures.

831 832

High salinity fluid inclusions of prograde stage skarn minerals do not contain CH4.

833

Therefore, the distinct lack of CH4 in the prograde stage was interpreted to suggest that the

834

CH4 in the fluid inclusions was not magmatic in origin. This is because, if this CH4 were

835

magmatic in origin, it would have been detected in the prograde stage Type I fluid inclusions.

836

The decarbonization process of fluid‒rock interaction, between low salinity fluids and

837

carbonaceous host rock, was also suggested in some studies as the basis for the origin of CH4-

838

bearing fluid inclusions in hydrothermal systems (Fu et al., 2014). However, this process is

839

also inappropriate for the explanation of CH4 in Dağbaşı skarns, because, if there was a

840

decarbonization process between hydrothermal solutions and carbonates, CO2 also would

841

have been detected in these fluid inclusions. However, the microthermometric study and

842

Raman measurements confirm that there were no CO2 phases. Additionally, if there was a

843

decarbonization process, the decarbonization trend of the data in Fig. 13a would have been

844

observed.

845 846

Biogenic origin was also suggested in some other studies (Whiticar, 1999; Ueno et al.,

847

2006; Potter and Longstaffe, 2007) by the incorporation of organic material into the high

848

temperature hydrothermal systems. Similarly, Shen et al. (2016) also reported that the organic

849

material of the host rocks produced the CH4 gas by thermal decomposition during the

850

emplacement of granitoid intrusions. Higher organic compounds in the limestones were also

851

indicated by Demir et al. (2017), and thermal degradation of organic materials was interpreted 25

852

as the most likely source of CH4 in the Sivrikaya skarn deposit. Therefore, it is suggested that

853

the CH4 at Dağbaşı skarn may have derived from the organic compounds of the limestone.

854

Therefore, the thermogenic origin of the CH4 at Dağbaşı skarn should be considered.

855 856

9.4. Comparisons with other skarn deposits

857

Skarn deposits in the northeastern region of Turkey are spatially associated with both

858

Early Cretaceous Berdiga limestones and Late Cretaceous carbonate layers in volcano-

859

sedimentary units (Saraç, 2003; Çiftçi, 2011; Demir et al., 2017; Sipahi, 2011; Sipahi et al.,

860

2017). The skarns associated with the Early Cretaceous Berdiga Formation are always

861

characterized by a considerable amount of sulfide phases (Aslan, 1991; Çiftçi 2011; Saraç,

862

2003). On the other hand, minor amounts of pyrite and chalcopyrite were accompanied by

863

magnetite and hematite along the skarn formation of the Upper Cretaceous carbonates (Saraç,

864

2003; Sipahi, 201; Demir et al., 2017). Reported ore reserves for these deposits are: three

865

hundred sixty thousand ton for Dereli (55 wt% Fe2O3), seven hundred fifty thousand ton for

866

Çambaşı (65.21 wt% Fe2O3), and five million ton for Kartiba (77.65 wt% Fe2O3), while there

867

are no reserve estimates for the other skarn locations (Saraç, 2003). However, old adits and

868

mine waste around these locations indicate that these skarns may also have significant ore

869

reserves to be mined. All these reserve estimates indicate that these skarns are relatively

870

small, but the polymetallic nature of these sulfide-bearing skarns makes them economically

871

significant because these metals needs to be recovered together with the main ore minerals of

872

magnetite and hematites.

873 874

Comparing the skarn type deposits of the eastern and western regions of Turkey by the

875

host rock lithology, they are related to Lower Cretaceous Berdiga limestones and Upper

876

Cretaceous volcano sedimentary units in the northeastern region (Demir et al., 2017).

877

However, in the western region, the Ayazmant and Evciler deposits were hosted by Early

878

Triassic metapelites‒metabasites and marbles (Öztürk et al., 2008; Oyman, 2010), and the

879

Kozbudaklar and Susurluk deposits were associated with the Triassic and Mesozoic

880

carbonates, respectively (Orhan et al., 2010; Orhan, 2017). Based on the skarn-related

881

intrusions, skarns in the northeastern region are related to the Late Cretaceous‒Eocene

882

granitoids, whereas skarn type deposits in the western region (i.e., Ayazmant, Kozbudaklar,

883

Susurluk, and Evciler) are related to the Eocene‒Miocene aged granitic intrusions (Öztürk et

884

al., 2008; Oyman, 2010; Orhan et al., 2010; Orhan, 2017). There were no porphyry systems in

885

close proximity to the skarn deposits of the northeastern region (Demir et al., 2017); however, 26

886

porphyry-bearing intrusions related to skarn deposits have been reported from the Ayazmant

887

and Kozbudaklar deposits of western Turkey (Oyman, 2010; Orhan, 2017). These

888

explanations indicate that the skarn-related granitoids in the eastern region are related to older

889

granitic intrusions compared to the western region. Therefore the absence of a porphyry

890

system in the eastern Black Sea region can be explained by the erosion of the older granitic

891

intrusions and accompanying porphyry deposits.

892 893

The δ18O and δD isotope results indicate that magmatic dominated solutions were

894

effective at the prograde stage of the Dağbaşı skarn deposits. On the other hand, the highly

895

depleted δ18O isotopes of the retrograde stage skarn minerals and their equilibrated solutions

896

indicate a dilution of magmatic solutions by the meteoric fluids. Highly depleted δ18O

897

isotopes of the skarn calcites and equilibrated solutions are also consistent with the mixture of

898

meteoric fluids. The δ18O isotopes, reported from Eğrikar skarn, are also consistent with the

899

magmatic source at the prograde stage, while the mixture of magmatic and meteoric water is

900

suggested for the retrograde skarn stage (Sipahi et al., 2017). According to the sulfur isotopes,

901

magmatic sources were demonstrated for the prograde stage of Kotana skarn deposit, and the

902

involvement of meteoric water was suggested at the retrograde stage (Çiftçi, 2011). Similar

903

findings were also reported from the Susurluk, Ayazmant and Evciler deposits of western

904

Turkey, where δ18O isotopes at these deposits were clearly linked to the magmatic source at

905

the early skarn stage, while the mixture of magmatic and meteoric solutions was suggested for

906

the later stages (Öztürk et al., 2008; Oyman, 2010; Orhan et al., 2017). When comparing the

907

skarn deposits of eastern and western Turkey, both of them are characterized by magmatic

908

dominated hydrothermal solutions at the early skarn stage, while later stages are characterized

909

by a mixture of magmatic and meteoric solutions.

910 911

The increasing meteoric effect on the Dağbaşı skarn deposit was also confirmed by the

912

fluid inclusion studies. The higher salinity values of the prograde stage fluid (up to 61.8 wt%

913

NaCl equ.) indicate that magmatic dominated fluids were effective on the skarn environment.

914

However, the decreasing salinity and Th trend of the retrograde stage fluid inclusions are

915

evidence of dilution with meteoric water. Salinity values reported from the Kotana (Çiftçi,

916

2011), Sivrikaya (Demir et al., 2017) and Eğrikar deposits (Sipahi et al., 2017) are somewhat

917

lower than the Dağbaşı skarn ore, up to 15, 15.4 and 14.3 wt% NaCl equ., respectively.

918

However, the decreasing salinity trend of the fluid inclusions in Sivrikaya and low salinity

919

values in the Eğrikar and Kotana deposits were accepted as evidence of dilution by meteoric 27

920

water. Moreover, higher salinity values of the prograde stage of the Susurluk and

921

Kozbudaklar skarn deposits of western Turkey, which reach up to 70 wt% NaCl equ. (Orhan

922

et al., 2010; Orhan, 2017), were also accepted as evidence of magmatic-related solutions. On

923

the other hand, decreasing salinity values of the retrograde stage were also accepted as

924

evidence of a meteoric effect on these deposits.

925 926

The homogenization temperatures of the fluid inclusions measured from the different

927

stages of skarn minerals were between 160–514 °C in the Dağbaşı skarn ore. These values

928

were slightly higher than the temperature ranges of the Eğrikar (between 200 and 425 °C;

929

Sipahi et al., 2017), Sivrikaya (166–462 °C; Demir et al., 2017), and Kotana (380–460 °C;

930

Çiftçi, 2011) skarn deposits. Much higher temperature values were reported from the Susurluk

931

(371–>600 °C; Orhan et al., 2010), Ayazmant (300–576 °C; Oyman, 2010), and Kozbudaklar

932

(308–>600 °C; Orhan et al., 2017) skarn deposits of western Turkey. When comparing these

933

temperature ranges of skarn deposits from eastern and western Turkey, it is apparent that the

934

western skarn deposits formed at higher temperature conditions. The much deeper formation

935

depths of the western skarn deposits may be one of the reasons for the higher temperature

936

conditions (Meinert et al., 2005) because the trapping pressure of the early stage fluid

937

inclusions in the Dağbaşı skarns was less than 884 bar, and there is no pressure estimate

938

reported from the other skarn locations of eastern Turkey. However, pressure estimates

939

reported from the Susurluk (~1000 bar) and Kozbudaklar (~2000 bar) skarns of western

940

Turkey are higher (Orhan et al., 2011; Orhan, 2017) than that of the Dağbaşı skarn ore.

941 942

10. Conclusions

943

Dağbaşı skarn deposits formed along the contact between the Lower Cretaceous

944

Berdiga limestone and the Upper Cretaceous Dağbaşı granitoid. The prograde stage was

945

represented by garnet and pyroxenes, while the retrograde stage was represented by epidote,

946

tremolite, actinolite, and chlorite. Quartz and calcites accompany both stages of the skarn

947

development. The ore minerals which were formed at the retrograde skarn stage mainly

948

consist of magnetite and hematite. Pyrrhotite, pyrite, chalcopyrite, sphalerite, and minor

949

galena are the accompanying sulfide phases in the ore.

950 951

The salinity values of the prograde stage fluid inclusions were measured to be between

952

48.8 and 61.8 wt%, while retrograde stage fluid inclusions were characterized by lower

953

salinity values. Higher salinity values of the prograde stage indicate magmatic dominated 28

954

solutions, whereas the lower salinity values of the retrograde stage and well-defined

955

decreasing salinity and Th trends of the fluid inclusions indicated a mixture of meteoric water

956

at the retrograde stage. The mixture of magmatic and meteoric solutions was further

957

supported by the boiling evidence of fluid inclusions.

958 959

The homogenization temperatures of the fluid inclusions in the prograde stage range

960

between 412 and 514 °C, and these measured temperatures were well within the appropriate

961

range for the calculated equilibrium temperatures of the prograde stage. On the other hand,

962

the equilibrium temperatures of the retrograde mineral assemblage (between 135 and 281 °C)

963

were slightly lower than the measured Th of the fluid inclusions (160–327 °C).

964 965

Thermal degradation of organic materials in carbonates was suggested as the source of

966

CH4 in fluid inclusions. This is because, if the CH4 were magmatic in origin, it would have

967

been determined in the first stage fluid inclusions. Moreover, if there was a decarbonization

968

process between the hydrothermal solutions and carbonates, CO2 would have been observed

969

in these inclusions.

970 971

The δ18O and δD isotopes of the prograde stage minerals and their equilibrated

972

solutions indicate that magmatic dominated solutions were effective at this stage. On the other

973

hand, δ18O and δD isotopes of the retrograde stage minerals and their equilibrated solutions

974

were found to be compatible with the meteoric origin.

975 976

Pressure estimates from the fluid inclusions indicate that early stage skarns were

977

formed in the range of 884 to 710 bar, whereas the formation pressure decreased to as low as

978

195 bar at the late skarn stage. Taking into consideration that ore formation takes place at the

979

retrograde stage, shallow skarn environments were suggested for the Dağbaşı skarn ore.

980 981

Acknowledgements

982

This study was financially supported by TÜBİTAK through project number 112Y331.

983

Special thanks are due to Andrea Jauss for providing the confocal Raman measurements at

984

WITec GmbH in Ulm, Germany. We would like to thank Serkan Şenkaya, Mustafa Aksu,

985

Kadir Bayraktar, and Mehdi İlhan for their assistance during the fieldwork and laboratory

986

studies. We are also grateful to anonymous reviewers whose valuable suggestions greatly

987

improved the earlier version of the manuscript. 29

988

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989 990

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991 992 993

Aydınçakır, E., 2006. Dağbaşı (Araklı-Trabzon) granitoyidi ve çevresinin petrografik, jeokimyasal ve petrolojik özelliklerinin incelenmesi. KTU, Fen Bilimleri Enstitüsü, Master Thesis, Trabzon, 120p.

994 995

Aydınçakır, E., Kaygusuz, A., 2012. Geç Kretase Yaşlı Dağbaşı (Araklı, Trabzon) volkanitlerinin petrografik ve jeokimyasal özellikleri. GÜFBED, 2, 2, 123-142.

996 997 998

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999 1000 1001

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1005 1006 1007

Bodnar, R.J., 2003. Introduction to aqueous fluid systems. In I. Samson, A. Anderson, D. Marshall, Eds. Fluid Inclusions: Analysis and interpretation. Mineral. Assoc. Canada, Short Course Series 32, 81–99.

1008 1009 1010

Bodnar, R.J., Reynolds, T.J., Kuehn, C.A., 1985. Fluid inclusion systematics in epithermal systems. In Geology and Geochemistry of Epithermal Systems (B.R. Berger and P.M. Bethke, eds) Society of Economic Geologist, Rev. Econ. Geol. pp. 2, 73–98.

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Bodnar, R.J., Vityk, M.O., 1994. Interpretation of microthermometric data for H2O-NaCl fluid inclusions. In Fluid Inclusions in Minerals, Methods and Appl. (ed. B. De Vivo and M.L. Frezotti), Virginia Tech 117–130.

1014 1015 1016

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Figure Captions: Figure 1. Major skarn occurrences and associated lithological units along the northeastern region of Turkey (modified from Güven, 1993). Figure 2. Geological map (modified from Kaygusuz and Aydınçakır, 2011) and skarn locations around the Dağbaşı (Trabzon) region. Figure 3. Geological map and skarn locations around the a) İpekçili, b) Köprüüstü, c) Kükürtlü, and d) Dere Mahalle districts (modified from Kaygusuz and Aydınçakır, 2011 and Demir, 2019). Figure 4. Macroscopic and microscopic scale ore texture in Dağbaşı skarn ore: a) early stage massive garnet in İpekçili skarn; b) rhythmically banded intergrowth of garnet and pyroxenes; c) nodular shaped garnet and pyroxenes in the carbonates; d) granular intergrowth of garnet and pyroxenes; e) late stage garnet developments along the cracks in the limestones (rarely including pyroxene); f) magnetite and hematite developments along the growth zone of late stage garnets; g) quartz, epidote and calcite developments in the volcanic rocks and carbonates during the retrograde skarn stage; h) massive tremolite‒actinolite intergrowth; i, j) macroscopic and microscopic scale garnet inclusions in the retrograde stage tremolite‒actinolite minerals; k) fracture-filling type quartz, epidote, and calcites in the volcanic host rocks; l) macroscopic scale quartz inclusions in the epidote-bearing samples (Grt: garnet; Pr: pyroxene; Mag: magnetite; Hem: hematite; Qz: quartz; Cal: calcite; Tr: tremolite; Act: actinolite; Ep: epidote). Figure 5. Macroscopic and microscopic scale ore texture in Dağbaşı skarn ore: a) magnetiteand hematite-bearing massive ore; b) banded ore along the carbonate layers; c) fracture-filling type magnetite veins in volcanites; d) breccia-filling type magnetite ore in volcanic host rocks; e) garnet inclusions in hematites represent later stage ore development; f) simultaneous ore formation with the retrograde stage epidotes; g) hematite formation around the magnetite minerals; g, h) magnetite and hematites were observed to be enclosed in chalcopyrites; i) chalcopyrite and sphalerite formation along the pyrite fractures; j) characteristic bird’s eye texture of pyrrhotites; k) galena inclusions in sphalerites; l) magnetite- and hematitedominated ore mass around the Kükürtlü and Dere Mahalle skarns; m) magnetite and hematite breccias were filled up by later stage quartz and epidotes; n) disseminated magnetite intergrowth in tremolites; o) minor amount of chalcopyrite inclusions in the hematites. Figure 6. Paragenetic diagram showing the mineralization stages in the Dağbaşı skarn (modified from Demir, 2019). Figure 7. Photomicrographs of primary fluid inclusions in different minerals: halite-bearing Type I inclusions in garnet (a), quartz (b) and calcite (c) minerals; vapor-rich Type II inclusions in garnet (d) and quartz (e) minerals; liquid-rich Type IIIa inclusions in garnet (f), and quartz (g) minerals; liquid-rich Type IIIb inclusions in epidote (h), calcite (i), sphalerite (j), and quartz (k) minerals; coexistence of Type IIIb and Type IIIc inclusions in quartz (k), Type IIIc inclusions in quartz (l). (S: solid, V: vapor, L: liquid phases.) 37

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Figure 8. Comparision of calculated salinity values (wt% NaCl equ.) of different types of fluid inclusions with the eutectic temperatures (Te, °C) of specific salt solutions. Figure 9. Distribution of Th temperatures of the different types of fluid inclusions versus salinity data from the Dağbaşı skarn ore (gray lines represent the surface fluid dilution trend of the fluid inclusion data). Figure 10. Histogram of homogenization temperature of fluid inclusions in garnet, epidote, quartz, calcite, and sphalerite minerals. Figure 11. Estimated pressure conditions of different fluid inclusion types. Type I inclusions, homogenized via halite dissolution, provide trapping pressure estimates between 884 and 710 bar based on isochores projected according to Sanchez et al. (2012) (lower limits of shaded area). Type II and Type IIIa inclusions are plotted on the liquid‒vapor curve according to the minimum and maximum Th ranges. Pressure estimates for these types of inclusions are between 445 and 195 bar. The isochores for Type IIIc inclusions were plotted using the equation of state of Brown and Lamb (1989) for average Th temperature (240 °C). The highest pressure estimates for Type IIIc inclusions were 310 bar for the highest independent geothermometer of quartz‒epidote mineral pairs (281 °C). The lower limit of pressure estimates cannot be plotted on the diagram because the independent geothermometer is much lower than the average Type IIIb isochore (gray arrows represent the pressure changes through the skarn development). Figure 12. Raman spectra of the CH4-bearing fluid inclusions in garnet (a), epidote (b) and quartz (c, d) minerals. Figure 13. a) Diagram of δ13CPDB vs. δ18OSMOW for the Dağbaşı skarn calcites, and comparison with the isotope composition of mostly known rock types (Sheppard, 1984; Hoefs, 2009). (Red empty circles represent the isotope composition of solutions equilibrated with prograde stage calcites for the Th temperatures of 514 °C. Blue empty symbols represent the isotope composition of solutions equilibrated with skarn calcites for the Ths of fluid inclusions of 311 °C). b) Plots of δ13C vs. δ18O isotopes of skarn calcites on the fluid‒rock interaction model of Taylor (1977). The curves calculated for different X(CO2) and fluid‒rock ratios of 0.01, 0.1, 05 and 0.5, 2, 10, 40, 80 and 100, respectively. Figure 14. δ18O composition of prograde and retrograde stage minerals from the Dağbaşı skarn ore (gray arrows represent decreasing isotopic trend from prograde to retrograde stage; plagioclase and two quartz minerals represent the center of the granitoid. Figure 15. δ18O and δD isotope compositions of retrograde stage epidote and tremolite minerals and the δ18O and δD isotopes of equilibrated solutions. The colored areas on the figure represent the isotope composition of different geological environments (modified from Taylor, 1974; Sheppard, 1984; Hoefs, 2009). (I: igneous, S: sedimentary, SMOW: standard mean ocean water.)

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1385 1386 1387 1388 1389 1390 1391 1392 1393 1394 1395 1396 1397 1398 1399 1400 1401 1402 1403

Table Captions: Table 1. Microthermometric fluid inclusion data from garnet, epidote, calcite, quartz and sphalerite minerals of Dağbaşı skarn ore. Table 2. Carbon and oxygen isotope composition of skarn calcites from the Dağbaşı Fe–Cu– Zn skarn mineralization. The isotope composition of the water, equilibrated with calcites, was also calculated for the highest Th temperature of fluid inclusions in prograde stage calcite (514 °C) and the average Th temperatures of the retrograde stage calcites (311 °C). The equation of Friedman and O’Neil (1977) was used for the calculations. Table 3. Oxygen and hydrogen isotope composition of the prograde and retrograde stage minerals from the Dağbaşı skarn ore. Plagioclase and two quartz minerals represent the center of granitoid (*The oxygen isotope composition of equilibrated solutions was calculated from different mineral‒water equations of Zheng (1993a, b).

Highlights: -

Both isotope and fluid inclusion data indicate that magmatic sources were responsible for the early stage skarn mineralization

1404 1405 1406

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Highly depleted isotope data and the decreasing salinity trend of the fluid

1407

inclusions correspond to the mixture of meteoric water during the later stage

1408

skarn development

1409 1410

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Decreases in the pressure conditions resulted in the boiling of hydrothermal solutions between prograde and retrograde stage skarn development

1411 1412 1413

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source of CH4 in the fluid inclusions.

1414 1415 1416 1417 1418 1419 1420 1421 1422 1423 1424 1425

Thermal degradation of organic materials in carbonates was suggested as the

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Declaration of interests ☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. ☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:

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1427 1428 1429 1430 1431

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