Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on water-rock interaction

VOLGEO-106708; No of Pages 18 Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

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Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on water-rock interaction Harun Aydin a,⁎, Hüseyin Karakuş b, Halim Mutlu c a b c

Hacettepe University, Hydrogeological Engineering Program, Beytepe, 06800, Ankara, Turkey Kütahya Dumlupınar University, Department of Geological Engineering, 43100 Kütahya, Turkey Ankara University, Department of Geological Engineering, Gölbaşı, 06100, Ankara, Turkey

a r t i c l e

i n f o

Article history: Received 18 September 2019 Received in revised form 2 November 2019 Accepted 2 November 2019 Available online xxxx Keywords: Hydrogeochemistry Stable isotope Geothermal water Eastern Turkey

a b s t r a c t The hydrogeochemistry of geothermal waters from 31 geothermal fields in eastern Turkey is investigated with regard to major ion compositions, stable (δ18O-δ2H-δ34S) and tritium (3H) isotope systematics. Discharge temperature of studied waters varies from 24 to 65 °C. Four different geochemical processes were found to control the major ion concentrations of waters which include dissolution of carbonates, fluid-mineral interaction, oxidation of sulfur-bearing minerals, and chloride enrichment. The northern, central and southern provinces are represented by different local meteoric water lines (LMWL) with deuterium excess of 15.0, 13.9 and 16.5‰ VSMOW, respectively. The stable isotope values of the thermal waters are close to LMWLs and indicate a meteoric origin. The enrichment in oxygen isotope composition (0.6 to 7.7‰ V-SMOW) in some thermal waters resulted from water–rock interaction process and 18O-exchange process between CO2 and H2O. The δ34S and δ18O contents in dissolved sulfate cover a wide range from 6.2 to 32‰ V-CDT and from −2.5 to 14.8‰ V-SMOW, respectively, indicating that the sulfate isotope systematics of the majority of waters is governed by dissolution of terrestrial sulfate and marine evaporites. The reservoir temperatures estimated by chemical and isotopic geothermometers of K\\Mg (27–127 °C), silica (29–179 °C) and 18OSO4-H2O (51–196 °C), and by the silicaenthalpy mixing model (130 to 235 °C) yielded inconsistent results. The geological factors (e.g., relatively thick crust, low surface heat flux, absence of ideal cover units) in eastern Turkey have resulted in the development of low- or moderate-temperature geothermal systems. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Geothermal fields in Turkey are generally distributed along the west Anatolian graben systems (e.g. Baba and Sözbilir, 2012; Karakuş and Şimşek, 2013; Karakuş, 2015), major active fault zones such as North Anatolian and East Anatolian Fault Zones (e.g. Süer et al., 2008; Italiano et al., 2013) and in areas of Neogene volcanics (e.g. Mutlu and Güleç, 1998; Aydın et al., 2013; Mutlu et al., 2013). Nearly 80% of geothermal fields in Turkey are located in the western Anatolia where nearly 50 geothermal power plants are currently in operation with an installed capacity of about 1450 MWe. Although the highest temperature (295 °C) in these fields was encountered in an exploratory well in Cappadocia, the geothermal potential of the central and eastern Anatolian regions has not yet received sufficient attention for use rather than balneological and greenhouse purposes and limited amount of dry ice production. In the eastern Anatolia, distribution of geothermal fields is

⁎ Corresponding author. E-mail addresses: [email protected] (H. Aydin), [email protected] (H. Karakuş), [email protected] (H. Mutlu).

chiefly controlled by several active faults (e.g. North Anatolian Fault, Karayazı Fault, Northeast Anatolian Fault, Dumlu Fault, Çobandede Fault, Horasan Fault, Kağızman Fault, and Çaldıran Fault) and a widespread Neogene volcanism (Fig. 1). Most of thermal waters in the region are manifested along these structures with flow rates up to 100 L/s and discharge temperatures in the range of 24 to 65 °C. In this study, we present new data on chemical and isotopic compositions of a total of 70 thermal and cold waters collected from 31 geothermal fields in the eastern Turkey covering an area of about 106,000 km2 (~14% of Turkey). Given that only a limited number of hydrogeochemical surveys were carried out in this region (e.g. Akkuş et al., 2005; Firat Ersoy and Çalik Sönmez, 2014; Pasvanoğlu, 2014; Hatipoğlu Temizel and Gültekin, 2018), the results of this study will greatly contribute not only to establishment of a comprehensive geochemicalisotopic database but also assessment of fields with promising geothermal energy potential. In line with this objective, the chemical and isotopic compositions of thermal and cold waters were used to examine the role of lithological controls on water type, and to estimate reservoir temperatures using chemical and isotopic geothermometers. Stable isotope variations (δ18O, δ2H and δ34S) provided supplementary constraints on the fluid origin.

https://doi.org/10.1016/j.jvolgeores.2019.106708 0377-0273/© 2019 Elsevier B.V. All rights reserved.

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

Fig. 1. Maps showing (a) location and (b) geology of eastern Turkey (modified from Gattinger et al. (1962), Altınlı et al. (1963, 1964), and Erentöz and Ketin (1974)). Active tectonic lines modified from Bozkurt (2001) and Koçyiğit (2013). The post-collision volcanic centers are from Keskin (2007). Name of each geothermal field is given Tables 2 and 3. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

2. Geological outline The collision between the Eurasian and Arabian Plates – one of the best examples of continental collision – exerts a great control on the generation of widespread Cenozoic volcanism and thickening of crust beneath the eastern Turkey (Fig. 1). The earlier studies suggested that the collision was initiated in the Middle Miocene (Şengör and Kidd, 1979; Şengör and Yilmaz, 1981; Dewey et al., 1986) although recent works are in the favor of Late Oligocene to Early Miocene (Okay et al., 2010; Karaoğlan et al., 2016; Açlan and Altun, 2018; Oyan, 2018). The ongoing convergence between these plates along the Bitlis–Zagros Thrust Zone resulted in uplift of the Turkish-Iranian Plateau and associated crustal thickening (Şengör and Kidd, 1979). As is common to most part of Anatolia, the nature of the tectonic regime in eastern Turkey changed from compressional-contractional (post-collisional convergence) in the Middle Miocene-Early Pliocene to compressional-extensional (tectonic escape) starting from Late Early Pliocene onward (Koçyiğit et al., 2001).

The geological framework of the eastern Turkey is made up of several lithological units. The Paleozoic-Mesozoic granulite-facies metamorphic rocks (Göncüoğlu and Turhan, 1984) exposing at south of the Lake Van comprise the basement (Fig. 1). The overlying JurassicLower Cretaceous marine limestones and Upper Cretaceous ophiolitic mélange represent the accreted rocks of a subduction complex (Şengör et al., 2003). Paleocene-Eocene intrusions of quartz diorite and granodiorite composition cropped out only along the Black Sea coast (Fig. 1). These rocks belonging to the eastern Pontide orogenic belt were formed during the final collision following the subduction and closure of the Neotethys Ocean between the Pontides and Taurides (Eyuboglu et al., 2011). They are stratigraphically overlain by terrestrial deposits characterized by cyclic alternations of the lacustrine and fluvial environments ranging in age from Paleocene to Lower Miocene (Varol et al., 2016). Additionally, Late Eocene to Miocene terrestrial and marine evaporite deposits are recorded in several localities of the region (Abdioğlu et al., 2015; Varol et al., 2016; Gökmen, 2017). The volcanism in eastern Turkey is represented by several eruptive pulses with a great variability in composition and continued

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

uninterruptedly until historic times. The Upper Miocene-Pliocene volcanic units are generally confined to the Erzurum-Kars Plateau (Fig. 1) which hosts an almost complete record of collision-related volcanism (Şengör and Kidd, 1979; Keskin et al., 1998). According to Keskin et al. (1998), the volcanism in the area between Erzurum and Kars districts can be subdivided into three stages regarding composition and eruption time. The early stage (from 11 to 6 Ma) characterized by a bimodal volcanism is comprised by mafic-intermediate lavas and acid pyroclastic rocks. The middle stage (from 6 to 5 Ma) with a unimodal volcanism is composed of andesitic-dacitic lavas. The late stage (mostly 5–2.7 Ma) of bimodal volcanism consists of plateau basalts and basaltic andesite lavas and felsic domes (Keskin et al., 1998). The Pliocene-Quaternary volcanism in the region generated a number of stratovolcanoes (e.g. Mt. Ararat, Mt. Tendürek, Mt. Süphan, and Mt. Nemrut) with alkaline to subalkaline composition (Yılmaz et al., 1998). Most of these volcanoes trending NE–SW direction along the northern part of Lake Van are historically active (Fig. 1). The volcanic succession in eastern Turkey is unconformably overlain by the Pliocene clastic sediments and Quaternary alluvium deposits. The continuous convergence of the Eurasian and Arabian plates has given rise to development of NW and NE-extending active strike-slip faults, N-S to NNWtrending fissures and an undeformed Plio-Quaternary continental volcano-sedimentary sequence accumulated in several strike-slip basins (Koçyiğit et al., 2001). Indeed, 1939 Erzincan (Ms: 7.8), 1966 Varto (Mw: 6.7), 1976 Çaldıran (Mw: 7.3), 1983 Horasan (Ms: 6.9) and 2011 Van (Mw: 7.2) earthquakes are the latest manifestations of tectonic unrest in the region.

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model Los Gatos Liquid Water Stable Isotope Analyzer. Analytical errors of analysis are 0.2‰ V-SMOW for δ18O and 1‰ V-SMOW for δ2H. Tritium (3H) was measured by a Quantulus 1220 model ultra-low-level beta scintillation spectrometer. The precision of analysis is 0.3 tritium units (TU). δ34SSO4 and δ18OSO4 analyses were performed at the Isotope Tracer Technologies Lab., Waterloo (Canada). Prior to analysis, water samples were filtered, acidified (10% HCl) and barium chloride was added to precipitate BaSO4. The precipitate then was collected and washed using distilled water until a neutral pH value is reached. Dried samples were weighed into tin capsules for separate δ34SSO4 and δ18OSO4 analyses. For δ18OSO4 analysis, approximately 0.1 mg of sample was combusted at 1430 °C and purified by Gas Chromatography (GC) before the measurement by Continuous Flow Isotope Ratio Mass Spectrometry (CFIRMS). The analysis was carried out using a Finnigan Mat DeltaPlus XL IRMS coupled with a Thermo Scientific TC/EA. Approximately 0.3 mg sample was used for δ34SSO4 analysis, with 3 mg of niobium pentoxide added to each sample to ensure complete sample combustion. Samples were loaded into a Fisons Instruments Elemental Analyzer (EA) to be flash combusted at 1100 °C. Released gases were carried by ultrapure helium through the analyzer, then separated by GC. Clean SO2 gas was carried into the Mat 253 IRMS for analysis. The data obtained for δ18OSO4 and δ34SSO4 were corrected using international standards and normalized to V-SMOW and V-CDT scales, respectively. The analytical precision for δ34SSO4 and δ18OSO4 are ±0.5‰ V-CDT and ± 0.5‰ VSMOW, respectively. 4. Results

3. Sampling and analysis In-situ measurements and sampling of waters from 31 different localities were conducted during the August–September 2015. Sampling locations are shown in Fig. 1 and sample volumes and sampling methods for chemical and isotopic analyses are summarized in Table 1. Samples were kept cold with ice molds (4 °C) until analysis. Chemical analyses (major ions) and tritium were performed at the Water Chemistry and Environmental Tritium Laboratories of the Hacettepe University. Major ions (excluding HCO3) were analyzed by ion chromatography. Alkalinity was determined by using titrimetric method – samples were titrated with strong acid (0.1 N H2SO4) to the methyl orange endpoint. SiO2 was analyzed with spectrophotometric method (ASTM D859) by a precision of 5% relative standard deviation (RSD) at Water Chemistry Laboratory of the Department of Environmental Engineering at Van Yüzüncü Yıl University. The charge balance errors of the chemical analyses are within ±5% (Table 2). The saturation states of waters with respect to calcite were estimated at their discharge temperatures with the use of PHREEQC program (Parkhurst and Appelo, 2013). The δ18O and δ2H analyses of the water samples were carried at the Stable Isotope Laboratory of the Hacettepe University using LGR DT-100 Table 1 Summary of sampling and analysis. Parameter

Sample quantity and sampling method

pH, T, EC

In-situ measurement carried out with YSI 556MP with an accuracy of ±0.2 units, ±0.15 °C and ± 1 μS/cm, respectively. 125 mL HDPE (high density polyethylene) bottle, filtered (0.45 μm), untreated. Total alkalinity is determined with the titrimetric methods. 125 mL HDPE bottle, filtered (0.45 μm) and acidified (pH b2) with HCl. 125 mL HDPE bottle, diluted as 1/5–1/10 with ultrapure deionized water. 10 ml glass vials, filtered (0.45 μm), untreated. 500 mL HDPE bottle, filtered (0.45 μm), untreated. 1000 mL HDPE bottle, filtered (0.45 μm), untreated.

Anions

Cations SiO2 δ18O and δ2H 3 H δ34SSO4 and δ18OSO4

Table 2 shows the results of in-situ measurements and chemical analyses of thermal and cold waters collected from geothermal fields in the eastern Turkey. δ18O, δ2H, 3H, δ34S(SO4) and δ18O(SO4) isotope compositions of samples are given in Table 3. Data on Çaldıran (Aydın et al., 2013) and Diyadin (Mutlu et al., 2013) geothermal fields are also included in these tables. 4.1. In-situ measurements In this study, based on discharge temperatures and the specific electrical conductivity (EC25) values, the sampled waters were classified as thermal, mineral and cold waters. The categorization of thermal and non-thermal waters was made comparing their discharge temperatures with the average summer air temperature (Pentecost et al., 2003; Jiang et al., 2018). Using the average summer temperature of eastern Turkey (20.4 °C) waters of 32 springs and 10 wells are classified as thermal water. The rest of the samples (28 out of 70) are classified with respect to EC25 values. Waters with EC25 values higher than 1000 μS/cm are regarded as mineral waters (6 springs and 3 wells) and others as cold waters (12 springs, 4 wells and 3 rivers). The outlet temperatures of thermal, mineral, and cold waters are from 24.3 °C (SUS) to 65.2 °C (MLK), 11.6 °C (AKY) to 19.5 °C (BUR), and 5.5 °C (ADR) to 18.7 °C (İKD), respectively (Table 2). Accordingly, the studied waters can be regarded as low-enthalpy (b90 °C; Muffler and Cataldi, 1978) thermal waters. The pH values of thermal waters fall in a wide range from 5.6 to 7.6 although some waters show exceptionally high (9.2 to 9.5; samples AYD, BHS, and GRM) and low pH (2.4, sample EBU). pH of mineral waters (from 5.3 to 7.9) is in the similar range of thermal waters, except for one sample from the Tatvan area (TAT pH = 10.4; Table 2). The specific electrical conductivity (EC25) values of thermal waters vary in a wide range from 203 to 10,434 μS/cm. The lowest and highest EC25 values were measured in Ayder (AYD: 55.1 °C) and Özalp (CAY: 53.5 °C) waters (Table 2). Some thermal waters (e.g. AYD, BHS, OTG, and TAZ) have relatively low EC25 (203 to 690 μS/cm). The EC25 of mineral waters varies from 1244 to 4718 μS/cm. BUR (19.5 °C, 461 μS/cm) and SGZ (15.2 °C, 571 μS/cm) waters discharging as bubbling springs

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

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Field*

İkizdere (1) Ayder (2) Borçka (3) Alabalık (4) Şavşat (5) Ardahan (6) Olur (7) Susuz (8) Akyaka (9) Kağızman (10) Horasan (11)

Köprüköy (12) Pasinler (13)

Uzunahmet (14) Akdağ (15) Yeşilyayla (16) Ilıca (17)

Üzümlü (18)

Çat (19) Tekman (20)

Tutak (21) Diyadin (22)

Code

IKI IKD AYD ADR OTG OTD ALA ALB COR COY SGZ SPA BHS BHC SUS SEL AKY KOT EİE HNS HNG HNM DLC KPC HSK ASB NAC UZA UZS AKK AKC ARZ YSC AZY HRM KRZ PET ERC EKS GUN HOL GOK HAM MEM TUT BUR1 DIB1 DVT1 DYD1 EBU1 KOP1 MLK1 TAZ1 RAH1

Type

W-Thermal R-Cold W-Thermal S-Cold S-Thermal R-Cold S-Mineral S-Cold S-Thermal S-Cold S-Mineral W-Cold W-Thermal S-Cold S-Thermal R-Cold S-Mineral S-Thermal S-Cold W-Thermal W-Thermal S-Cold S-Thermal S-Cold W-Thermal S-Thermal W-Cold W-Thermal W-Cold S-Thermal S-Cold S-Thermal S-Cold W-Thermal W-Thermal W-Mineral W-Cold W-Thermal W-Mineral S-Cold S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Mineral S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Cold

Date

05/15 08/15 05/15 05/15 08/15 08/15 05/15 05/15 08/15 08/15 04/15 08/15 05/15 05/15 08/15 08/15 05/15 05/15 05/15 08/15 08/15 08/15 05/15 05/15 05/15 05/15 08/15 05/15 08/15 05/15 08/15 05/15 05/15 05/15 05/15 05/15 08/15 05/15 05/15 08/15 04/15 04/15 04/15 04/15 09/09 07/10 07/10 09/09 09/09 07/10 07/10 07/10 07/10 07/10

Long

40°36′36″ 40°36′36″ 41°05′53″ 41°07′03″ 41°54′18″ 41°54′18″ 42°04′25″ 42°04′29″ 42°24′05″ 42°21′47″ 42°39′38″ 42°40′14″ 42°10′09″ 42°10′10″ 43°04′51″ 43°05′13″ 43°37′53″ 43°00′07″ 42°42′24″ 42°14′27″ 42°14′29″ 42°14′27″ 41°51′13″ 41°51′19″ 41°40′58″ 41°36′11″ 41°45′26″ 41°26′15″ 41°26′10″ 41°21′36″ 41°21′36″ 41°16′54″ 41°16′35″ 41°06′19″ 41°06′24″ 41°06′56″ 41°02′30″ 39°35′29″ 39°36′57″ 39°38′46″ 40°52′06″ 41°15′23″ 41°08′24″ 41°09′24″ 42°46′29″ 43°37′48″ 43°37′01″ 43°39′13″ 43°39′13″ 43°42′39″ 43°38′58″ 43°34′21″ 43°37′55″ 43°44′45″

Lat

40°47′07″ 40°47′02″ 40°57′14″ 40°57′17″ 41°19′22″ 41°19′24″ 41°17′04″ 41°17′04″ 41°23′41″ 41°22′46″ 41°05′44″ 41°05′52″ 40°52′16″ 40°52′16″ 40°47′34″ 40°52′16″ 40°44′18″ 40°13′47″ 40°08′23″ 40°07′49″ 40°07′58″ 40°07′58″ 39°59′24″ 39°59′30″ 39°58′30″ 39°59′25″ 39°59′20″ 39°58′37″ 39°58′09″ 40°05′55″ 40°06′08″ 40°03′44″ 40°03′13″ 39°56′45″ 39°57′23″ 39°59′45″ 39°58′22″ 39°44′01″ 39°44′06″ 39°44′01″ 39°31′46″ 39°31′55″ 39°27′09″ 39°30′28″ 39°33′07″ 39°26′23″ 39°28′38″ 39°28′26″ 39°29′23″ 39°29′05″ 39°29′35″ 39°24′24″ 39°28′28″ 39°28′11″

Z

T

(m)

(°C)

855 848 1190 1367 1288 1285 859 969 1501 1711 1739 1799 1898 1878 1810 1919 1476 1434 1379 1624 1625 1645 1662 1662 1662 1697 1663 1830 1851 1820 1885 1846 1793 1766 1758 1784 1776 1162 1147 1393 1805 2090 1943 2045 1611 2067 1980 1999 1952 2051 1925 2059 1962 2095

63.2 18.7 55.1 5.5 32.5 17.5 11.7 13.1 36.6 6.0 15.2 17.4 36.1 9.2 24.3 14.7 11.6 27.0 11.9 36.5 57.3 14.5 24.9 13.1 37.7 28.6 12.5 57.0 13.6 28.5 12.9 33.4 11.6 36.0 38.4 19.2 14.5 33.4 18.6 13.7 25.3 30.3 46.9 41.0 24.6 19.5 48.3 64.2 53.7 25.8 50.6 65.2 39.8 11.5

pH

6.8 7.8 9.7 7.4 7.2 7.9 5.3 7.0 5.6 7.3 6.0 6.8 9.4 7.4 5.9 7.6 5.6 6.1 8.0 6.1 7.3 7.2 5.7 7.5 6.3 6.3 7.1 6.6 6.1 5. 9 7.3 6.2 7.2 6.2 6.5 7.9 7.4 6.5 5.9 7.5 6.5 6.4 6.6 6.5 6.4 5.4 7.0 6.6 7.0 2.4 7.2 6.6 6.9 7.8

EC25

Na

K

Ca

Mg

Alk.

Cl

SO4

SiO2

CB

(μS/cm)

(ppm)

(ppm)

(ppm)

(ppm)

(ppm)

(ppm)

(ppm)

(ppm)

(%)

4643 68 203 57 325 120 2399 290 9334 86 571 611 690 247 7438 112 2589 5007 511 8324 8553 519 1884 310 9580 4702 298 6117 764 2745 223 2495 269 7565 6626 4718 582 6672 2666 338 2824 1937 1321 1545 4676 461 2365 3557 2230 1978 3268 2125 575 681

1050 3.3 39 1.3 40 2.6 314 6.5 2050 4.7 48 46 130 20 1600 5.1 160 1100 42 2400 2550 16 240 16 1800 710 10 900 30 460 16 450 17 1700 1500 1300 14 570 80 2.1 190 86 84 46 770 22 160 120 160 34 170 160 150 17

140 0.8 0.9 0.3 0.7 0.4 19 5.2 51 0.5 8.5 8.2 1.4 0.6 20 1.5 13 46 3.6 73 82 3.2 34 3.8 59 53 4.2 120 6.5 17 1.6 21 1.2 42 40 17 1.1 19 8.4 0.7 9.9 21 14 10 46 34 70 52 59 71 69 76 70 1.9

410 11 13 6.2 31 22 240 43 370 9.0 33 36 18 25 260 15 260 260 41 220 290 51 120 31 520 240 39 450 72 82 20 93 21 150 160 61 75 420 67 2.1 510 230 160 240 150 49 93 130 180 70 69 130 78 48

82 1.3 0.4 0.5 0.1 2.1 45 7.1 64 2.9 33 33 1.6 9.0 170 5.1 150 78 21 88 130 29 51 13 160 120 7.5 240 37 96 9.6 100 10 81 100 15 27 600 330 43 52 110 60 70 110 9.3 81 84 69 32 76 52 78 30

3100 41 83 25 47 76 1100 150 3000 47 370 350 160 150 2500 81 1400 3050 290 5950 6700 310 950 190 2300 1550 160 2850 440 1400 140 1600 140 2150 2700 3500 360 4050 1900 210 1950 1400 950 1150 1800 130 880 920 990 0.0 590 710 750 310

460 0.7 2.8 0.4 4.9 0.4 150 4.5 1750 0.5 9.0 9.0 79 4.8 1750 1.1 180 320 6.1 660 590 4.1 140 2.5 2500 950 3.4 1150 7.3 250 5.6 160 1.1 1800 1150 25 2.7 770 36 2.9 170 65 64 19 560 19 160 120 120 9.0 180 270 170 1.0

290 4.8 37 1.1 110 3.4 230 13 240 2.1 6.1 8.5 63 3.3 2.5 1.5 110 150 21 160 150 9.6 4.7 4.0 5.2 0.2 8.1 0.3 12 0.1 1.8 2.3 5.3 0.8 0.4 0.0 7.0 37 60 12 20 53 32 44 68 120 140 100 160 670 130 120 150 8.2

150 3.5 79 2.7 50 8.2 120 na 120 6.0 68 na 62 na 120 na 110 100 6.9 140 120 18 80 6.1 180 100 7.5 170 7.5 100 na 140 na 110 140 68 5.3 180 110 na 51 80 95 75 85 86 68 153 60 140 56 120 79 5.2

4.8 3.3 3.7 2.3 −0.1 2.4 4.2 4.8 4.8 4.9 2.6 4.6 1.2 2.0 4.2 2.2 4.1 4.2 3.2 2.7 3.2 2.5 4.4 2.4 4.5 2.4 3.0 3.7 2.5 3.7 1.7 3.8 4.9 2.1 4.2 4.7 3.5 4.2 1.1 0.4 1.9 −3.8 −2.9 0.3 4.4 −0.5 −4.4 1.9 3.7 −3.1 4.6 −4.9 −4.2 3.7

Water Type NaHCO3 CaHCO3 NaHCO3 CaHCO3 NaSO4 CaHCO3 NaHCO3 CaHCO3 NaHCO3 CaHCO3 MgHCO3 MgHCO3 NaHCO3 CaHCO3 NaCl CaHCO3 CaHCO3 NaHCO3 CaHCO3 NaHCO3 NaHCO3 CaHCO3 NaHCO3 CaHCO3 NaCl NaCl CaHCO3 NaHCO3 CaHCO3 NaHCO3 CaHCO3 NaHCO3 CaHCO3 NaCl NaHCO3 NaHCO3 CaHCO3 MgHCO3 MgHCO3 MgHCO3 CaHCO3 CaHCO3 CaHCO3 CaHCO3 NaHCO3 CaSO4 NaHCO3 MgHCO3 CaHCO3 CaSO4 NaHCO3 NaHCO3 NaHCO3 MgHCO3

H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

Table 2 Physical and chemical compositions of geothermal and cold waters in eastern Turkey.

Çaldıran (25)

Özalp (26) Gürpınar (27) Başkale (28) Hizan (29) Tatvan (30)

Güroymak (31)

PAT HAD TAS AYR AYS2 BUG KOC2 ZYR2 CAY YUR CAM GRM KOK NHL TAT CKR

W-Mineral S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Thermal S-Cold S-Thermal S-Thermal S-Thermal S-Thermal S-Mineral S-Thermal S-Mineral S-Thermal

09/09 09/09 09/09 09/09 06/10 09/09 06/10 06/10 09/09 08/09 09/09 09/09 09/09 09/09 09/09 09/09

42°50′04″ 43°23′15″ 43°25′39″ 43°51′47″ 43°51′29″ 43°59′00″ 43°56′20″ 44°01′58″ 44°10′02″ 43°47′21″ 44°05′23″ 42°11′10″ 42°15′54″ 42°14′48″ 42°17′21″ 42°01′21″

39°14′28″ 39°13′32″ 39°17′01″ 39°07′17″ 39°06′46″ 39°00′52″ 39°04′45″ 39°02′26″ 38°31′01″ 38°13′51″ 37°56′49″ 38°17′49″ 38°23′13″ 38°38′45″ 38°31′48″ 38°39′13″

1645 1901 2102 2023 2015 2058 2037 2159 2135 2162 1894 1317 1743 2248 1678 1291

18.1 65.0 51.1 37.0 25.8 34.3 25.1 11.4 53. 5 25.4 34.0 46.8 16.3 51.6 14.2 34.5

6.2 6.5 7.1 7.0 6.1 7.6 6.7 7.2 7.6 6.7 7.2 9.2 6.3 6.9 10.4 6.5

1393 6037 4570 3807 1592 1329 3540 139 10,434 2489 5440 1078 3673 1661 1244 1905

180 1850 730 610 240 84 380 10 2600 40 900 210 140 320 120 300

20 190 70 170 77 24 120 0.7 170 4.3 160 4.2 26 33 13 52

92 330 220 200 52 99 140 20 180 180 110 20 390 70 100 130

12 64 75 30 50 71 170 2.8 69 120 120 1.5 190 15 62 94

400 1150 1200 1600 750 810 1950 85 5500 1150 2100 210 1200 1100 900 1550

190 2500 640 360 160 17 160 1.0 850 14 560 140 27 23 32 47

34 420 460 150 53 23 63 6.3 540 40 260 110 850 0.4 4.2 0.5

45 78 57 100 130 130 130 23 99 19 31 31 11 170 9.0 120

* Numbers in bracket indicate geothermal fields from Fig. 1; S: Spring; W: Well; R: River; 1 and 2: Data compiled from Mutlu et al. (2013) and Aydın et al. (2013), respectively; Alk: Total alkalinity; na: not analyzed.

4.6 4.9 4.5 3.8 3.4 3.1 4.2 4.6 4.4 3.2 3.9 2.9 3.8 4.2 0.4 3.3

NaHCO3 NaCl NaHCO3 NaHCO3 NaHCO3 MgHCO3 NaHCO3 CaHCO3 NaHCO3 MgHCO3 NaHCO3 NaCl CaSO4 NaHCO3 CaHCO3 NaHCO3 H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

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Patnos (23) Erciş (24)

5

6

H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

Table 3 Isotope ratios of geothermal and cold waters in eastern Turkey. Field*

Code

δ18OH2O (‰ V-SMOW)

(‰V-SMOW)

İkizdere (1)

IKI IKD AYD ADR OTG OTD ALA ALB COR COY SGZ SPA BHS BHC SUS SEL AKY KOT EİE HNS HNG HNM DLC KPC HSK ASB NAC UZA UZS AKK AKC ARZ YSC AZY HRM KRZ PET ERC EKS GUN HOL GOK HAM MEM TUT BUR1 DIB1 DVT1 DYD1 EBU1 KOP1 MLK1 TAZ1 RAH1 PAT HAD TAS AYR AYS2 BUG KOC2 ZYR2 CAY YUR CAM GRM KOK NHL TAT CKR

−12.5 −12.7 −14.2 −13.9 −14.0 −13.1 −13.7 −13.1 −13.4 −12.1 −13.3 −13.1 −14.5 −10.0 −12.7 −11.9 −13.6 −10.1 −12.7 −10.5 −10.7 −11.2 −13.8 −13.1 −13.2 −12.2 −11.0 −12.5 −12.8 −13.1 −12.5 −13.0 −12.9 −13.3 −13.6 −12.5 −11.0 −12.3 −12.2 −11.7 −12.0 −12.0 −12.2 −12.0 −12.1 −14.0 −12.6 −11.3 −11.5 −10.8 −12.5 −12.8 −12.8 −12.9 −11.4 −10.1 −11.3 −11.4 −12.2 −13.5 −11.6 −12.1 −3.4 −12.7 −10.7 −10.9 −10.7 −9.6 −9.8 −10.0

−96 −86 −99 −94 −95 −89 −96 −93 −104 −82 −93 −92 −101 −66 −93 −80 −100 −84 −92 −88 −88 −77 −98 −94 −97 −91 −75 −92 −92 −91 −84 −90 −88 −95 −97 −94 −77 −90 −84 −82 −82 −78 −79 −76 −87 −96 −94 −91 −92 −79 −95 −90 −94 −91 −83 −79 −82 −90 −85 −91 −85 −83 −72 −89 −82 −72 −69 −65 −66 −69

Ayder (2) Borçka (3) Alabalık (4) Şavşat (5) Ardahan (6) Olur (7) Susuz (8) Akyaka (9) Kağızman (10) Horasan (11)

Köprüköy (12) Pasinler (13)

Uzunahmet (14) Akdağ (15) Yeşilyayla (16) Ilıca (17)

Üzümlü (18)

Çat (19) Tekman (20)

Tutak (21) Diyadin (22)

Patnos (23) Erciş (24) Çaldıran (25)

Özalp (26) Gürpınar (27) Başkale (28) Hizan (29) Tatvan (30)

Güroymak (31)

δ2HH2O

δ34SSO4

δ18OSO4

(TU)

(‰V-CDT)

(‰V-SMOW)

0.68 ± 0.21 4.89 ± 0.28 0.06 ± 0.34 8.97 ± 0.53 0.16 ± 0.19 4.75 ± 0.29 5.48 ± 0.45 6.41 ± 0.48 0.45 ± 0.21 6.72 ± 0.33 0 ± 0.20 na 0.19 ± 0.34 7.85 ± 0.49 0.33 ± 0.20 na 0.30 ± 0.40 0.70 ± 0.22 na 0 ± 0.19 0 ± 0.21 4.01 ± 0.27 0 ± 0.21 na 0 ± 0.19 0.20 ± 0.35 na 0.01 ± 0.19 na 0.31 ± 0.39 na 0.16 ± 0.34 na 0.75 ± 0.36 0 ± 0.20 0 ± 0.19 5.41 ± 0.31 0.21 ± 0.19 1.92 ± 0.23 5.29 ± 0.29 2.13 ± 0.26 1.25 ± 0.21 0.05 ± 0.20 0 ± 0.19 0.25 ± 0.23 1.58 ± 0.25 b1.0 0 ± 0.25 0.12 ± 0.26 6.29 ± 0.29 b1.0 1.66 ± 0.23 b1.0 7.86 ± 0.30 0.48 ± 0.29 2.12 ± 0.30 1.35 ± 0.29 0.04 ± 0.27 3.25 ± 0.17 1.07 ± 0.29 1.70 ± 0.28 5.97 ± 0.31 0.20 ± 0.28 0.25 ± 0.28 0 ± 0.26 0 ± 0.26 0.71 ± 0.26 3.22 ± 0.32 1.12 ± 0.29 0.75 ± 0.27

15.9 na 10.5 na 8.3 na na na 10.6 na 8.9 na 29.6 na nm na 13.5 6.2 na 10.8 8.1 na 16.9 na 21.8 nm na nm na nm na nm na nm nm nm na 32.0 20.3 na 10.9 14.4 29.3 10.9 15.4 na na 20.0 13.7 na na na na na 15.6 18.2 18.1 19.6 na 12.2 na na 19.8 20.3 24.6 na 29.5 nm nm nm

4.3 na 4.8 na 6.9 na na na 9.7 na 5.5 na 14.8 na nm na 4.1 3.3 na 5.9 4.7 na 6.0 na 5.2 nm na nm na nm na nm na nm nm nm na 10.5 6.8 na 4.1 4.7 6.6 8.2 7.6 na na 8.1 2.0 na na na na na 7.7 −0.5 0.9 −2.5 na 1.5 na na 7.1 8.1 13.0 na 10.7 nm nm nm

3

H

* Numbers in bracket indicate geothermal fields from Fig. 1; 1 and 2: Data compiled from Mutlu et al. (2013) and Aydın et al. (2013), respectively; na: not analyzed; nm: not measured due to low SO4 concentration.

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

were classified as mineral water, although their discharge temperatures are below 20.0 °C and EC25 values are also b1000 μS/cm. The EC25 of cold waters is from 57 to 764 μS/cm. 4.2. Hydrochemical characteristics The geothermal waters of eastern Turkey show highly variable chemical characteristics. Even mineral water springs have distinct chemical composition. Most of studied thermal and mineral waters have relatively high concentrations of major cations such as Na (up to 2600 ppm) and Ca (up to 520 ppm). Mg concentrations, excluding sample ERC (610 ppm), fall in a wide range between 0.1 and 330 ppm. Potassium with relatively lower concentrations against other cations varies from 0.7 to 190 ppm. The lowest Mg (0.1–0.4 ppm) and K (0.7–0.9 ppm) concentrations are measured in waters collected from the Black Sea coast in northern part of the study area (samples AYD and OTG). HCO3 with a concentration range from 47 to 6700 ppm constitutes the main anion in thermal and mineral waters at most localities (except for the samples EBU, HSK, HAD and OTG) (Table 2). Cl and SO4 concentrations vary from 2.8 to 2500 ppm and 0 to 850 ppm. Low SO4 concentrations (0–5.2 ppm) are confined to Bitlis (CKR, NHL), Erzurum (AKK, ARZ, ASB, AZY, DLC, HRM, HSK, KRZ, UZA) and Kars (SUS) geothermal fields whereas high sulfate contents (420–540 ppm) are found in thermal waters of the Van region (CAY, HAD, TAS). Moreover, the highest SO4 concentrations were measured in EBU (670 ppm) and KOK (850 ppm) springs. The SiO2 falls in a wide range between 9 and 200 ppm and it is not correlated with discharge temperatures. 4.3. Isotope compositions 4.3.1. Oxygen and hydrogen isotopes Results of oxygen and hydrogen isotope compositions of water samples are shown in Table 3. In a general sense, δ18O and δ2H values of thermal waters are distributed over a wide range varying from −14.5 to −3.4‰ V-SMOW and from −104 to −65‰ V-SMOW, respectively. δ18O and δ2H of cold waters are nearly 0.04‰ and 2.6‰ V-SMOW lower than those of thermal waters. Considering the size of the studied region and the geographical differences among the sampling sites, three different local meteoric water lines (LMWL) were considered. These sub-regions are designated as northern, central and southern provinces. The northern province covers the Black Sea mountain range (İkizdere, Ayder, Borçka, Şavşat, Ardahan, Akyaka and Olur fields), the central province comprises geothermal fields in the Erzurum-Kars plateau (except for Çat, Olur and Tekman) and the southern province includes samples in the vicinity of Lake Van and the area around the volcanic centers (e.g. Nemrut Caldera, Mt. Tendürek) (Fig. 1). For the construction of local meteoric water lines for the northern and central provinces, δ18O and δ2H contents of cold-water springs and rainfall were used (see Supplementary data). δ18O and δ2H range from −14.5 to −12.5‰ V-SMOW and − 104 to −93‰ V-SMOW in the northern province, while they change from −13.8 to −10.1‰ V-SMOW and − 98 to −84‰ V-SMOW in the central province, and − 14 to −3.4‰ V-SMOW and − 96 to −65‰ V-SMOW in the southern province (Aydın et al., 2013; Mutlu et al., 2013), respectively. The results of oxygen‑hydrogen isotope compositions are consistent with previous studies (Aydın et al., 2009; Firat Ersoy and Çalik Sönmez, 2014; Pasvanoğlu, 2014; Hatipoğlu Temizel and Gültekin, 2018). Tritium contents of eastern Anatolian geothermal waters are between 0 and 6.29 TU and those of cold waters vary from 4.01 to 8.97 TU (Table 3). 4.3.2. Sulfur isotope systematics For some thermal waters from the Erzurum region (e.g. AKK, ARZ, ASB, AZY, HRM, KRZ and UZA), δ34SSO4 and δ18OSO4 analyses are not available due to low SO4 concentrations (b 5 ppm; Tables 2 and 3).

7

However, we compiled sulfate isotope data of several samples from Mutlu et al. (2012). δ34SSO4 values of waters vary in a wide range from 6.2 and 32‰ V-CDT, and δ18OSO4 falls in the range between −2.5 and 14.8‰ V-SMOW. Although sulfur isotope systematics are not generally correlated with geographical setting of samples, δ34SSO4 values of waters from northern part of the region are mostly higher than those of the southern part. Regarding the spatial distribution of δ18OSO4, the southern part usually has the lowest values (Table 3). In a previous study by Hatipoğlu Temizel and Gültekin (2018), the average δ34SSO4 and δ18OSO4 values of waters in the Pasinler (Erzurum) geothermal area are reported 19.3‰ V-CDT and 3.8‰ V-SMOW which are consistent with values obtained in this work (21.8‰ V-CDT and 5.2‰ V-SMOW). 5. Discussion 5.1. High pH waters in eastern Anatolia All sampled waters are of slightly acidic or near neutral character (Table 2). However, there are some exceptions to this. For example, a few thermal waters (AYD, BHS, and GRM) and one mineral water (TAT) show exceptionally high pH. Samples AYD and BHS with high pH values (~9.4) and low ionic contents (EC25: 200–610 μS/cm) are classified as Na-HCO3 type water whilst samples GRM (9.2, 1000 μS/cm) and TAT (10.4, 1250 μS/cm) are Na\\Cl and Ca-HCO3 type waters, respectively. Natural waters with high pH (~ 10 or above) are not usual (Hem, 1985). Processes that may cause pH to increase include (i) alteration of ferromagnesian minerals, such as serpentinization of olivine in ultramafic terranes (Hem, 1985), (ii) saline waters in soda lakes (Cioni et al., 1992), (iii) dissolution of silicate minerals in granitic or metamorphic terranes (e.g., Venturelli et al., 2003; Ramírez-Guzmán et al., 2004), and (iv) calcite precipitation and cation exchange processes along the flow path or within the aquifer (Krainov et al., 2001). The high-pH waters in the eastern Anatolia are not manifested through or not interacted with the ultramafic rocks. Therefore, the first process is unlikely to be effective on the high pH of these thermal waters. Secondly, the highest pH value in the region was measured in sample TAT located in the southwest of the Lake Van, which is one of the largest soda lakes on Earth. The pH values of Lake Van and groundwater in the carbonate aquifers in the region are reported 9.31 to 9.99 (Reimer et al., 2009) and 6.67 to 7.97 (Aydın et al., 2009), respectively. Since pH of sample TAT is out of these ranges, any mixing process for this sample is not likely. The mechanism controlling the high pH value in waters discharging from magmatic or metamorphic rocks is attributed to high water/rock ratio and low PCO2 content at the initial stage of rock dissolution (Ramírez-Guzmán et al., 2004). High-pH thermal and mineral waters are reported to manifest from granitic (T = 57.0 °C, pH = 9.40, Hem, 1985; T = 8.6 °C, pH = 10.06, Nordstrom et al., 1989), granitic-metamorphic (T = 38.3 to 43.1 °C, pH = 8.78 to 9.98, Ramírez-Guzmán et al., 2004), and volcanic terrains (T = 93.5 °C, pH = 9.48, Fournier, 1989). Similarly, high pH of thermal waters discharging from magmatic or metamorphic rocks in the region can be attributed to hydrolysis reactions of silicate minerals at low PCO2. The fact that AYD, BHS, GRM and TAT waters are supersaturated to calcite (see Section 5.5) implies that calcite precipitation may be also effective in the formation of high pH values. 5.2. Chemical composition of geothermal waters Chemical composition of thermal waters is modified by several physicochemical processes including fluid-mineral reactions, dissolution and precipitation, boiling, conductive cooling, and mixing. Among these, the solubility reactions are the primary process, which determine the abundance of a particular species that is retained in solution before the precipitation takes place (Nicholson, 1993). Consequently, the

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

temperature and the sort of lithology with which waters are interacted exert a great control on the type of waters. Geochemical processes affecting the chemistry of eastern Anatolian thermal waters are discussed in the Piper diagram (Fig. 2). In this graphic, we distinguished four different geochemical evolution paths for the samples: (i) Ca + Mg-HCO3 type waters – probably originated from the dissolution of carbonate rocks, (ii) Na-HCO3 type waters – possibly evolved by the dissolution of feldspars under relatively high temperature, or by calcite precipitation along the flow path or within the aquifer, (iii) SO4-rich waters – originated by dissolution of sulfates, by oxidation of sulfur-bearing minerals or -by the condensation of geothermal gases in oxygen-rich shallow groundwaters, and (iv) Na-Cl + HCO3 type waters that are initially of Na-HCO3 or Ca-HCO3 character and then progressively enriched by chloride during the deep circulation. Our classification based on major ion compositions are in good agreement with the results of previous studies in the region (e.g., Mutlu and Güleç, 1998; Taşkıran, 2006; Gültekin et al., 2007; Aydın et al., 2013; Mutlu et al., 2013; Firat Ersoy and Çalik Sönmez, 2014; Pasvanoğlu, 2014; Hatipoğlu Temizel and Gültekin, 2018). In addition, several element variation diagrams (e.g., Ca + Mg vs. HCO3, Na + K vs. HCO3, Ca vs. SO4 and Na vs. Cl) are used to examine possible hydrogeochemical processes and to distinguish different origins of waters (Mazor, 1991). The first group of thermal waters is represented by the Üzümlü (EKS, ERC), Çat (HOL), Tekman (GOK, HAM, MEM), Ardahan (SGZ), Akyaka (AKY), Çaldıran (BUG) and Gürpınar (YUR) localities which generally host carbonate and/or ultramafic rocks (Figs. 1 and 2). In this group of waters, Ca + Mg and HCO3 contents are positively correlated (R2 = 0.99; Fig. 3a). Dissolution of carbonate (e.g. calcite) and mafic minerals (e.g. olivine) that controls the chemistry of Ca + Mg-HCO3 type waters consumes carbon dioxide as shown in the following reactions:

Most of the Ca + Mg-HCO3 type waters are manifested as bubbling springs. As recognized in the vicinity of springs HAM, HOT and YUR, travertine deposits are formed as CO2 is lost during ascend of thermal waters to the surface. Another strong linear relationship (R2 = 0.939) was found between Na + K and HCO3 (Fig. 3b). Na-HCO3 type waters may occur by the dissolution of CO2 arising from mantle or crustal sources (Henley et al., 1984). The weak carbonic acid formed by the dissolution of CO2 alters the silicate minerals and generates Na-HCO3 type waters. The presence of deep-seated CO2 in the region has been reported by Nagao et al. (1989), Mutlu et al. (2012) and Aydın et al. (2015). In these studies, the major source of CO2 in fluids is suggested to be crustal origin (~85%) while the mantle and sedimentary CO2 have a limited contribution (~10% and ~5%, respectively). In addition, Nagao et al. (1989) reported that 80% of CO2 in samples from the Çaldıran geothermal area at northeast of Lake Van is originated from marine carbonates. Most Na-HCO3 type waters, except for sample AYD, are slightly acidic or have near neutral pH ranging from 5.7 to 7.9 (Table 2). Considering high geogenic CO2 degassing in the study area, another possible process that may lead to formation of Na-HCO3 waters might be precipitation of carbonate minerals (e.g. calcite, dolomite) along the flow path. SO4-rich waters (samples EBU, KOK, OTG and BUR) show a general enrichment in SO4 and do not preferentially favor any of the cations (Figs. 2 and 3c). They do not follow the Ca/SO4 = 1 line which is characteristic of dissolution of evaporite minerals (e.g. gypsum and anhydrite) (Fig. 3c). Moreover, the coefficient of determination for the linear relationship between Ca and SO4 concentrations of these samples is found 0.6 which indicates that other mechanisms rather than evaporite dissolution could be effective for the sulfate enrichment. Sample EBU (Diyadin area) with the lowest pH value (2.4) among the entire data set is probably originated from dissolution of H2S in an oxygenated shallow groundwater.

CaCO3 ðcalciteÞ þ CO2 þ H2 O↔Ca2þ þ 2HCO3 −

ð1Þ

H2 S þ 2O2 þ →2Hþ þ SO4 −2

Mg2 SiO4 ðforsteriteÞ þ 4H2 CO3 ↔2 Mg2þ þ H4 SiO4 þ 4HCO3 −

ð2Þ

Acid-sulfate type waters formed by this process may be best distinguished on a classical Cl-SO4-HCO3 ternary diagram of Giggenbach

ð3Þ

Fig. 2. Piper diagram for geothermal waters in eastern Anatolia. Cold water data from Aydın et al. (2017). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

9

HNG

HNG

80

100

Carbonate dissolution

HNS CAY

R² = 0.328

60

KRZ

KOT

R² = 0.741

R² = 0.990 Ca+Mg-HCO3

UZA

HRM

40

ERC

IKI

SUS

CAM

HSK

KOC KOK

AKY

SO4-rich Na-Cl

R² = 0.967 0 0

10

20

HNS CAY

80

R² = 0.939

30 40 Ca+Mg (meq/l)

50

18

60

40

EKS

KOC HOL

KOT

HRM

COR HSK

CAM SUS

T UT

20

(a)

KRZ UZA

T AS

R² = 0.649

HAD

(b)

0

70

0

(c)

KOK

ERC

R² = 0.835

60

Na-HCO3

EKS

NHL

20

HCO3 (meq/l)

HCO3 (meq/l)

100

20

40

60 80 Na+K (meq/l)

70

100

120

(d)

HAD HSK

R² = 0.609

9

CAY

50

Silicate mineral dissolution and/or cation exchange T AS HAD

R² = 0.280

6

CAM AYR

3

IKI

ASB

5

10

30

SUS

UZA

COR

HRM

ASB

CAY

ERC T AS CAM

R² = 0.519

HNS

T UT HNG

0

40

AZY

20 R² = 0.981

COR

ALA KOT

10

HNS

0

R² = 0.935 Reverse ion exchange

12

60

EBU

Cl (meq/l)

S04 enrichment

SO4 (meq/l)

15

SUS

15 Ca (meq/l)

HSK

UZA

20

0 25

IKI KOT KRZ

AYR AKK ARZ

0

20

40

HNG

Silicate mineral dissolution and/or cation exchange

60 Na (meq/l)

80

100

120

Fig. 3. Composition diagrams for geothermal waters in eastern Anatolia. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

(1991) shown in Fig. 4. In this diagram, the majority of samples, except for SO4-rich waters, plot in the field of peripheral waters that are represented by high HCO3 content. Samples EBU and OTG fall into the field of “steam-heated waters”, however the former plots very close to the SO4

apex (Fig. 4). The EBU water is manifested in the northern part of the historically active Tendürek volcano (Fig. 1). The interaction between acidic waters and host rocks may leach silica and metal cations and cause their concentration to increase (Nicholson, 1993). Regarding

Fig. 4. Giggenbach's (1991) Cl-SO4-HCO3 ternary diagram (the relative proportions were calculated from ppm units). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

sample EBU, high SiO2 concentration (142 ppm) even at low discharge temperature (25.8 °C) can be attributed to the leaching process. Differently from this, SO4-rich character of sample OTG, located in the vicinity of massive sulfide deposits along the Black Sea coast (Fig. 1), is undoubtedly due to dissolution of sulfide minerals which is best represented by the oxidation of pyrite: 2FeS2 ðpyriteÞ þ 7O2 þ 2H2 O→2Feþ2 þ 4SO4 −2 þ 4Hþ

ð4Þ

H+ and SO4 which are released by the oxidation reaction increase the acidity of water. However, near neutral pH of sample OTG (7.20) probably results from subsequent buffering reactions (e.g. dissolution of carbonate or silicate minerals). Samples BUR and KOK discharging in areas of young volcanic units (Fig. 1) are low-temperature bubbling springs with a slightly acidic character (pH = 5.4 and 6.3; Table 2). It is thought that oxidation of metallic sulfides and/or H2S emanation are the possible processes behind the observed SO4-rich character of these waters. The last geochemical evolution path in the region is evidenced by progressive enrichment in Cl and Na (Fig. 2). A linear relationship between these ions is highly apparent with a high coefficient of determination (R2) of 0.935 (Fig. 3d). EC25 and pH of this type of waters vary in a wide range; 740 to 11,900 μS/cm and 5.3 to 9.4, respectively. Geothermal waters with higher temperatures (~65 °C) (e.g., DVT, HAD, MLK) belong to this group (Table 2). Na\\Cl type waters are mostly formed as a result of dissolution of evaporite minerals such as halite or seawater intrusion/mixing. In a local scale, no evaporitic rock is exposed that could contribute Na and Cl ions to these waters, and all sampling points are far from the coastal area. As shown by rock dissolution experiments (Ellis and Mahon, 1964, 1967), some water-soluble constituents such as Cl and B largely exist as soluble salts on surfaces of grains, crystals, and micro-fissures in the rock. Therefore, Cl can be easily dissolved from such rocks without significantly altering the primary minerals (Arnórsson, 2000). The conservative behavior of chloride causes its concentration to increase during the deep circulation and/or water rock-interaction after dissolution from the rocks. In this context, the origin of Cl in the eastern Anatolian thermal waters can be attributed to such water-rock interaction process. 5.3. Isotopic compositions 5.3.1. 18O\\2H Relationship Using the stable isotope composition of rainfall and cold waters, three local meteoric water lines (LMWL) were proposed for the region. LMWL of the eastern Black Sea region (the northern province) is defined by the equation of δ2H = 8.0δ18O + 15.0 (R = 0.994) (Fig. 5a). The resulting line is consistent with the LMWL with deuterium excess of +16‰ V-SMOW that was reported for the same province (Ekmekçi and Gültekin, 2015). The LMWL of the central province is represented by δ2H = 7.9δ18O + 13.9 (R = 0.989; Fig. 5b) which is in good agreement with the equation of δ2H = 8δ18O + 14.9 proposed for the Şenyurt-Erzurum area (Sayın and Özcan Eyüpoğlu, 2005). Regarding the southern province, LMWL has deuterium excess of +16.5‰ VSMOW (Fig. 5c; Aydın et al., 2009). Most of studied thermal waters in these provinces plot close to LMWL or between LMWL and GMWL (δ2H = 8δ18O + 10), which is consistent with a meteoric origin with no significant geothermal shift (Fig. 5). Interestingly some samples in the northern (e.g. COR and IKI), central (e.g. HNG, HNS and KOT) and southern (e.g. AYR, CAM, CAY, DVT, DYD and HAD) provinces show slight to moderate positive shifts in δ18O which can be attributed to water-rock interaction and/or phase separation processes. Higher δ18O values imply enrichment in heavy isotope of oxygen due to exchange between thermal fluids and oxygen-bearing minerals. Enrichment in δ18O generally depends on reservoir mineralogy and temperature conditions. According to Truesdell and Hulston (1980), the largest geothermal shifts are common to

Fig. 5. δ18O–δ2H diagram for (a) the northern, (b) the central, and (c) the southern provinces in eastern Turkey. The cold and hot waters are designated by blue circle and red triangles, respectively. LMWL: Local Meteoric Water Line (for c Aydın et al., 2009); GMWL: Global Meteoric Water Line (Craig, 1961). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

carbonate-hosted high-temperature reservoirs (e.g. Lanzarote and the Salton Sea). This can be ascribed to the fact that δ18O content of carbonate minerals (29‰ V-SMOW) is higher than that of silicate minerals (8 to 12‰ V-SMOW; Clark and Fritz, 1997). The remarkable δ18O shift (nearly 7.7‰ V-SMOW) observed in Özalp (CAY) water in the southern province can be explained by isotopic exchange between thermal fluid and reservoir rock (Fig. 5c). The positive δ18O shifts recognized in other waters (e.g. HNG, HNS, KOT, DVT, DYD, AYR, CAM, COR, HAD and IKI) are in the range of 1.5 to 2.4‰ VSMOW. High EC25 values (N3500 μS/cm; Table 2) may reflect that water-rock interaction process played an important role for modification not only chemical but also isotopic composition of these waters. Alternatively, δ18O shifts may also be generated by exchange with CO2 as recently demonstrated by Karolyte et al. (2017) using the following

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-9

two-component mixing model (Johnson et al., 2011).


 δ18 O ¼ ð−0:003  elevation Þ−9:475 R2 ¼ 0:878

ð6Þ

Table 4 Temperature and isotope data of samples showing 18O enrichment. 18OiH2O is the initial 18O value obtained from LMWL's in Fig. 5. 18OfH2O is measured value. Δ18O is observed 18O shift (Δ18O = 18OfH2O − 18OiH2O). εCO2-H2O is the temperature dependent oxygen isotopic enrichment factor calculated from Bottinga (1968). X0CO2 is the fraction of thermogenic CO2.

ASB AYR CAM CAY CKR COR DIB DVT DYD HAD HNG HNS HSK IKI KOP KOT MLK NHL TAS TAZ UZA

T

18

(°C) 28.6 37.0 34.0 53.5 34.5 36.6 48.3 64.2 53.7 65.0 57.3 36.5 37.7 63.2 50.6 27.0 65.2 51.6 51.1 39.8 57.0

OiH2O

18

OfH2O

Δ18O

εCO2-H2O

(‰)

(‰)

(‰)

(‰)

−13.3 −13.3 −11.7 −11.1 −10.7 −14.9 −13.9 −13.4 −13.6 −11.9 −12.9 −12.9 −14.0 −13.9 −13.9 −12.4 −13.4 −10.1 −12.4 −13.9 −13.4

−12.2 −11.4 −9.8 −3.4 −10.0 −13.4 −12.6 −11.3 −11.5 −10.1 −10.7 −10.5 −13.2 −12.5 −12.5 −10.1 −12.8 −9.6 −11.3 −12.8 −12.5

1.1 1.8 2.0 7.7 0.7 1.5 1.2 2.1 2.1 1.8 2.2 2.4 0.8 1.4 1.4 2.3 0.6 0.6 1.1 1.1 0.9

39.4 37.8 38.4 34.9 38.3 37.9 35.8 33.2 34.9 33.1 34.3 37.9 37.7 33.3 35.4 39.8 33.0 35.2 35.3 37.3 34.3

X0CO2

0.28 0.34 0.59 1.00 0.29 0.21 0.15 0.21 0.24 0.20 0.25 0.49 0.13 0.14 0.17 0.88 0.05 0.11 0.16 0.17 0.10

-10

18 O

( ‰ V-SMOW)

18 O

-11

= (-0.003×Altitude) - 9.474 R² = 0.880

-12 IKI -13 OTG

ALA

-14

AYD -15 -55 0

250

500

750

1000 Altitude (m)

1250

1500

1750

(b)

-65 2H =

‰ V-SMOW)

where, δ18Oi and δ18Of denote initial and measured (final) oxygen isotope ratios. εCO2-H2O is temperature-dependent oxygen isotope enrichment factor, and X0CO2 is the CO2 fraction. This model was tested for samples showing an apparent 18O-shift (Fig. 5). The temperaturedependent oxygen isotope enrichment factor (εCO2-H2O) was calculated from measured temperatures using the equation proposed by Bottinga (1968). CO2 fraction (X0CO2) was estimated (Table 4) for three different initial δ18OiCO2 values; ~6‰ V-SMOW (MORB, White, 2013), 4–12‰ VSMOW (magmatic rocks, Bindeman, 2008), and 30‰ V-SMOW (thermogenic marine limestone, Clark and Fritz, 1997). Considering that CO2 is originated from MORB and magmatic rocks, the samples would be represented by an 18O depletion. However, this is inconsistent with what is shown from Fig. 5 which implies 18O-enrichment for the studies waters. On the contrary, the apparent 18O-enrichment observed in some samples (CAM: 2.0‰, CAY: 7.7‰, HNS: 2.4‰, KOT: 2.3‰) requires equilibrium with thermogenic CO2 with fractions between 0.49 and 1.0. CO2 fraction is found b0.4 for the rest of samples that show enrichment in 18 O varying from 0.6 to 2.2‰ (Table 4). The results show that, in addition to water-rock interaction, 18O-shift associated with 18O-exchange process between CO2 and H2O could also modify the stable isotope systematics of thermal waters of eastern Turkey. The calculated fraction values particularly for samples CAM, CAY, HNS and KOT are not likely realistic and require confirmation with the measured δ18OCO2 and δ18OH2O values. δ18O and δ2H values of waters tend to decrease by elevation, known as the altitude effect. In general, the decrease in δ18O content per 100 m of altitude varies from 0.15 to 0.50‰ V-SMOW whereas δ2H decreases from 1.0 to 4.0‰ V-SMOW (Clark and Fritz, 1997). Using altitude vs. δ18O or δ2H relationship, the average recharge elevation of a spring or groundwater can be estimated. The change in δ18O and δ2H of cold water and rainwater samples collected from the eastern Black Sea region (the northern province) with respect to elevation yields the following relations (Fig. 6a and b):

Code

(a)

h
 i h
i ð5Þ ¼ δ18 OiCO2 −εCO2 −H2 O  X OCO2 þ δ18 OiH2 O  1−X OCO2

2H (

18

δ OHf 2 O

11

(-0.024×Altitude) - 60.106 R² = 0.857

-75

-85

ALA

-95

OTG

IKI AYD -105 0

250

500

750

1000 Altitude (m)

1250

1500

1750

Fig. 6. (a) δ18O-altitude, and (b) δ2H-altitude relationships for geothermal waters for northern province. The cold and hot waters are shown by blue circle and red triangles. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)


 δ2 H ¼ ð−0:024  elevation Þ−60:138 R2 ¼ 0:858

ð7Þ

For the central and southern provinces neither δ18O nor δ2H is linearly correlated with the altitude, and therefore, recharge elevation of thermal waters in these regions could not be determined. From these equations (Eq. 6–7), the decrease in δ18O and δ2H of waters from the northern province is estimated 0.3‰ and 2.4‰ V-SMOW for 100 m increase in elevation. Assuming that oxygen isotope compositions of waters are not significantly affected by the interaction with rocks, the recharge elevations of Ayder (AYD), Alabalık (ALA) and Borçka (OTG) thermal waters are found between 1350 and 1600 m asl (Fig. 6a and b). Ignoring its δ18O enrichment (~1.5‰ V-SMOW), recharge altitude of Ikizdere (IKI) water corresponds to 980 m (Fig. 6b). The Eastern Black Sea mountain range extending in southern part of the Ikizdere and Ayder geothermal fields is home to several peaks with altitudes ranging from 3000 to 3900 m (e.g. Mt. Kaçkar). In the north of these areas, the elevation gradually decreases to the sea level. Considering the estimated average recharge elevations, thermal waters from these areas are likely recharged by rainwater falling on the northern front of the Eastern Black Sea mountain range. In a previous study, the depletions in δ18O and δ2H contents by about 0.09 to 0.85‰ VSMOW/100 m and 0.7 to 6.2‰ V-SMOW/100 m were compiled for various regions in Turkey with elevations up to 2500 m (Apaydın, 2018). These values are within the range of decreases in δ18O and δ2H estimated from the northeastern Anatolia. More specifically, for the Rize district (0.2‰ V-SMOW/100 m; Ekmekçi and Gültekin, 2015) and the area around Lake Van (0.18‰ V-SMOW/100 m; Aydın et al., 2009),

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altitude dependence of oxygen isotope corresponds well with the computed values for the investigated part of eastern Turkey.

a significant recent freshwater contribution that also recharges the crater lake.

5.3.2. Water circulation Tritium has a relatively short half-life (12.32 years; Lucas and Unterweger, 2000) compared to residence time of deeply circulating waters. Therefore, the tritium content of waters in deep geothermal systems is generally around 0 or below the detection limit. However, tritium provides an opportunity to determine the presence of recent recharge or any mixing between freshwaters and thermal waters. Most of samples have tritium contents close to 0 TU indicating minor or no recent freshwater contribution to the thermal waters of the eastern Anatolia. Some waters have N1 TU referring to freshwater mixing. Chloride behaves in a conservative manner and therefore Cl concentration in waters generally increases with increasing residence time. In this sense, tritium vs. Cl diagram is very useful for distinguishing young and deep groundwater flow systems (e.g. Karakuş and Şimşek, 2008; Apollaro et al., 2016). In order to examine possible mixing process between fresh and deep thermal waters, 3H vs. chloride graph was constructed (Fig. 7). Cold waters affected by a limited water-rock interaction have low chloride and high tritium contents and thus they can be classified as shallowly circulating young waters. As the residence time of the groundwater increases, concentrations of conservative constituents, such as chloride, increase, and tritium contents exponentially decrease. Chloride vs. tritium relationship along the deep circulation is pictured with a black dashed line in Fig. 7. Deep water component in the region is best represented by samples HSK, SUS, COR, HRM, UZA, HNS and AYR which have relatively high chloride concentration (N350 ppm) and tritium contents close to “0”. It is important to note that some thermal waters (HAD, AZY, TAS, IKI, and MLK) with tritium values between 0.68 and 2.12 TU show relatively high discharge temperatures (36.0 to 65.2 °C) and Cl concentrations (270 to 2500 ppm) (Tables 2 and 3). High ionic content (EC25: 2200 to 7600 μS/cm; Table 2) and apparent 18O enrichment of these waters (Fig. 5) indicate mixing with shallow freshwaters from rivers and creeks. For example, thermal springs of HAD and TAS emerge at the surface within the Zilan Stream valley (Van) and springs MLK and IKI; discharge along the Murat River (Agri) and İkizdere Creek (Rize), respectively. Likewise, although the hot water component is more dominant, spring waters AYS and HOL might also be influenced by mixing to some extent. The fact that EBU spring is manifested in close vicinity of cold water suggests that this sample has a minor thermal water component and, as mentioned earlier, it is an acid-sulfate type water, possibly evolved as a steam heated shallow local groundwater. Sample NHL, located in the crater lake of Nemrut volcano, has tritium content of 3.22 TU indicating

5.3.3. δ34S and δ18O in dissolved sulfate Sulfate in natural waters can originate from various reservoirs. Dissolution of terrestrial or marine evaporites, oxidation of sulfides, and hydrogen sulfide are the common sources. Both δ34SSO4 and δ18OSO4 are sensitive to biological and geochemical reactions, and therefore they highly fractionate in the biotic and abiotic environments (Clark and Fritz, 1997). In active volcanic geothermal systems, sulfate can be derived from hydrolysis of SO2 in magmatic vapor or hydrolysis of S from sulfur-bearing rocks. Sulfur from each of these sources has its own stable isotope signature. Sulfur of magmatic origin has δ34S value in a narrow range close to ~0‰ V-CDT (Thode, 1991). δ34S of dissolved sulfate in modern oceans is 21‰ V-CDT (Rees et al., 1978) which was not constant throughout the geological time and varied between 10.5‰ in the Permian and 31‰ in the Cambrian (Claypool et al., 1980). Marine and nonmarine evaporites, on the other hand, have a wider range of sulfur isotope composition varying from 5 to 35‰ (Thode, 1970). Processes affecting the sulfur and oxygen isotope composition of geothermal waters in the eastern Turkey are evaluated in the δ34SSO4 vs. δ18OSO4 plot (Fig. 8). The δ34SSO4 contents of samples ranging from 6.2 to 32.0‰ are higher than that of magmatic sulfur (0‰ V-CDT; Table 3) indicating high fractionation of δ34SSO4 among the sulfur compounds. The variation in sulfate isotopic compositions of studied geothermal waters can be explained by the following processes (i) mixing, and (ii) exchange of oxygen between sulfate and water which results in enrichment of δ18OSO4. The first process can be traced with the samples on a mixing line between the terrestrial sulfate and marine evaporite endmembers (red triangles in Fig. 8). The δ34S and δ18O contents of SO4 in these samples show a relatively modest correlation (R = 0.64). The presence of terrestrial (Varol et al., 2016; Gökmen, 2017) and marine evaporites (e.g. Abdioğlu et al., 2015) in various parts of eastern Turkey most probably explains the observed diverse stable isotope compositions of thermal waters. Secondly, thermal and mineral waters (AYR, BUG, DYD, HAD, TAS; blue squares in Fig. 8) in the northeast part of Lake Van have δ18OSO4 contents represented by the lowest values among the entire data set (−2.5 to 2.0‰ V-SMOW; Table 3). Such low δ18OSO4 values can be

Ayder

8

Diyadin

(SO4 ) (‰, VSMOW)

9

Marine Evaporites

Atmospheric Depositions

10

Fresh/young waters Shallow circulation

7

COR

MS

MEM

CAY DLC

SGZ

5

AYD HNG KOT

HOL

AKY

3H

(TU)

BUG

Old waters Deep circulation

5

0

GOK IKI

2

HOL GOK

MLK

1

CKR

IKI AYR

0 0

HAM

HSK

HAD AYR

AYS

3

EKS

DYD T AS

Terrestrial Sulphates

NHL

4

r = 0.640

PAT

OT G

18 O

EBU ALA

6

15

500

-5 -5

TAS CAM UZA CAY ASB HRM COR HNS 1000

AZY SUS

1500 2000 Cl (ppm)

HSK 2500

0

5

10 34 S

HAD

KOT

3000

Fig. 7. The 3H\ \Cl diagram. The cold (data from Aydın et al., 2017), and hot waters are shown by blue circle and red triangles. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

15 20 (‰, CDT)

25

30

35

(SO4 )

Fig. 8. The plot of δ34S-δ18O in dissolved sulfate. The red triangle and blue square represent mixing and sulfate-water oxygen exchange process which effects sulfate isotopic composition of samples, respectively. The end-member range of the terrestrial, atmospheric, and marine are −15‰ to 7‰ V-CDT, −5‰ to 10‰ V-CDT, and 9‰ to 34‰ V-CDT for δ34SSO4; and − 10‰ to 5‰ V-SMOW, 8‰ to 18‰ V-SMOW, and 10.3‰ to 19‰ V-SMOW for δ18OSO4, respectively; and the δ34SSO4 and δ18OSO4 values of modern seawater (MS) are 21‰ V-CDT and 9‰ V-SMOW, respectively (Clark and Fritz, 1997). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

explained by 18O isotopic exchange between water and dissolved sulfate under elevated temperatures (e.g. Lloyd, 1968; Mizutani and Rafter, 1969; Chiba and Sakai, 1985). According to Van Stempvoort and Krouse (1994), δ18OSO4 is significantly affected by sulfate-water oxygen exchange in high-temperature geothermal systems, and in equilibrium conditions the δ18OSO4 tends to increase relative to δ18OH2O. Using the equation of Mizutani and Rafter (1969), the relative oxygen isotope enrichment at equilibrium condition is calculated 16.6‰ VSMOW at 100 °C, 11.98‰ V-SMOW at 150 °C, and 8.76‰ V-SMOW at 200 °C. The δ18OH2O contents of AYR, BUG, DYD, HAD and TAS samples vary between −10.12‰ and − 13.49‰ V-SMOW (Table 3). Assuming equilibrium exchange reaction at 150 °C, the δ18OSO4 values of these samples were calculated around ±2‰ V-SMOW, which are consistent with the measured values. Therefore, it is concluded that oxygen isotopic composition of AYR, BUG, DYD, HAD and TAS samples is affected by sulfate-water oxygen exchange process.

5.4. Estimation of reservoir temperature In order to estimate the reservoir temperature of geothermal waters in the eastern Turkey, chemical (cations and silica) and isotopic (δ18OSO4-H2O) geothermometry methods were used. As a first approach, the applicability of geothermometry in the Na-K-Mg system can be evaluated using the ternary plot proposed by Giggenbach (1988). As shown in this diagram (Fig. 9), none of the thermal waters attains full equilibrium but only few samples (e.g. CAY, KRZ, HNS, and COR) are plotted close to the curve separating the “immature waters” from the “partially equilibrated waters”. Most of samples fall into the “immature waters” field. As pointed out by Giggenbach (1988), the suitability of thermal waters for the application of ionic solute geothermometers should be decided on the basis of “maturity index (MI)”. MI, the measure of the degree of water-rock equilibrium, is required ≥2 and this prerequisite is fulfilled by samples COR, BHS, SUS, HNS, HNG, AZY, KRZ, CAY and GRM. Using the Na\\K geothermometer of Giggenbach (1988), reservoir temperatures for samples from northern (e.g. COR, SUS), central (e.g. AZY, HNG), and southern provinces (e.g. CAY, GRM) are estimated in the range of 100–141 °C, 109–156 °C, and 130–200 °C, respectively (Table 5). Giggenbach (1988) stated that exchange reactions involving K\\Mg ions equilibrate much faster than those of Na\\K ions and, the estimated K\\Mg temperatures are likely to reflect conditions at shallow levels of a geothermal system and therefore K/Mg ratio appears to be a good indicator for the last temperature of water-rock equilibration (Fournier, 1989). This feature of the K\\Mg system demonstrated itself in the calculated temperatures that are much lower than Na\\K temperatures (Table 5). However, the highest reservoir temperature estimated by the K\\Mg geothermometer is 127 °C for sample AYR with MI value below 2. On the other hand, sample EKS has the lowest temperature with 27 °C, which is slightly higher than the discharge temperature.

Fig. 9. Na-K-Mg ternary plot (Giggenbach, 1988) for eastern Anatolia waters.

13

These findings suggest that results of this geothermometer should be evaluated with a particular attention. Silica geothermometers are based on experimentally determined solubility of various SiO2 phases. Because of the existence of different silica forms in nature (e.g. quartz, chalcedony and amorphous silica), different geothermometry equations have been proposed. In general, the most commonly applied silica geothermometers are quartz and chalcedony geothermometers. Quartz and chalcedony geothermometers (Fournier, 1977) yielded reservoir temperatures in the range of 36 to 179 °C, and 29 to 157 °C, respectively (Table 5). In this study, the majority of samples was collected from natural manifestations or shallow geothermal wells. Therefore, thermal waters are likely to be affected by the cold-water mixing, which can cause dilution of silica concentrations. The silica vs. enthalpy model (Fournier and Truesdell, 1974) is one of the most commonly techniques used for the estimation of reservoir temperature and mixing ratio of cold waters. The silica-enthalpy diagrams constructed on the basis of silica solubility for selected geothermal fields in eastern Turkey are shown in Fig. 10. From this model, reservoir temperatures predicted for the northern, central and southern provinces are in range of 150 to 200 °C, 170 to 235 °C, and 130 to 210 °C, respectively. Using these temperature ranges and discharge temperatures of respective thermal and cold waters, the mixing ratio is estimated between 56 and 93%. These values are not consistent with the findings from 3H\\Cl diagram (Fig. 7), such that some waters (e.g., AYD and IKI) have 3H contents close to “0” denoting negligible modern recharge contribution (Table 3). The δ18OSO4-H2O geothermometer depends on the 18O fractionation between water and dissolved sulfate (Lloyd, 1968; Mizutani and Rafter, 1969). Geothermometers based on these experiments provide estimates with a slight discrepancy with b10 °C. Seal et al. (2000) proposed the following linear regression for a temperature interval between 70 and 350 °C using the data of Lloyd (1968) and Mizutani and Rafter (1969).
 1000 lnα 18 OSO4 −H2 O ¼ 3:26 106 =T 2 −5:81

ð8Þ

Fractionation factor (α) in this equation defined as:

α 18 OSO4 −H2 O ¼

1000 þ δ18 OSO4 1000 þ δ18 OH2 O

ð9Þ

Reservoir temperatures computed by 18OSO4-H2O geothermometer are between 51 and 196 °C (Table 5). Most temperatures fall in the range of 70 to 150 °C (Fig. 11). Only for one sample (BHS), estimated temperature is found lower than the discharge temperature. For some samples the difference between the estimates are up to 90 °C and no significant correlation exists among the geothermometry results. This is probably due to different rate of equilibrium experimentally determined for each method. For example, exchange reactions required for Na\\K and δ18OSO4-H2O geothermometers progress very slowly at low temperatures and therefore, equilibrium conditions are attained in a longer time span. On the other hand, cooling of fluid during ascent to the surface results in underestimation of reservoir temperatures by K\\Mg and silica geothermometers. Reservoir temperatures estimated from K\\Mg, silica and 18OSO4-H2O geothermometers are in the range of 27 to 127 °C, 29 to 179 °C, and 51 to 196 °C, respectively. According to major classification systems based on average measured and/or calculated reservoir temperatures (Hochstein, 1990), geothermal systems in the eastern Turkey can be regarded as low (T b 125 °C) or moderate (T b 225 °C) temperature systems.

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Table 5 Estimated reservoir temperatures (°C) obtained by geothermometric methods. Code

T1

Na-K2

K-Mg2

SiO32

SiO42

18

O5

Code

T1

Na-K2

K-Mg2

SiO32

SiO42

18

O5

IKI AYD OTG ALA COR SGZ BHS SUS AKY KOT HNS HNG DLC HSK ASB UZA AKK ARZ AZY HRM KRZ ERC EKS HOL GOK HAM

63.2 55.1 32.5 11.7 36.6 15.2 36.1 24.3 11.6 27.0 36.5 57.3 24.9 37.7 28.6 57.0 28.5 33.4 36.0 38.4 19.2 33.4 18.6 25.3 30.3 46.9

* * * * 141 * 100 107 * * 152 156 * * * * * * 141 * 109 * * * * *

107 + 58 64 83 50 43 51 43 78 88 86 76 76 76 87 53 57 75 71 73 36 27 48 56 54

160 124 102 146 146 117 112 145 142 138 156 145 125 171 139 179 137 156 145 155 117 173 144 102 125 134

136 96 72 120 120 88 83 119 116 111 132 119 97 149 112 157 111 132 119 130 88 151 118 72 97 107

106 89 76 × 63 90 + × 98 139 110 119 83 93 × × × × × × × 64 89 112 107 90

MEM TUT BUR DIB DVT DYD EBU KOP MLK TAZ PAT HAD TAS AYR AYS BUG KOC CAY YUR CAM GRM KOK NHL TAT CKR

41.0 24.6 19.5 48.3 64.2 53.7 25.8 50.6 65.2 39.8 18.1 65.0 51.1 37.0 25.8 34.3 25.1 53.5 25.4 34.0 46.8 16.3 51.6 14.2 34.5

* * * * * * * * * * * * * * * * * 200 * * 130 * * * *

45 73 97 88 80 86 101 88 96 88 80 118 89 127 96 64 91 115 + 104 68 55 90 52 78

122 128 129 116 162 110 128 107 146 124 97 124 108 138 152 154 154 137 62 81 80 42 170 36 149

93 101 102 88 139 81 134 78 120 96 66 96 78 112 127 129 130 110 29 50 49 + 147 + 124

81 84 × × 86 137 × × × × 89 186 152 196 × 122 × 174 77 51 × 73 × × ×

1

: Discharge temperature; 2: Na\ \K and K\ \Mg geothermometer of Giggenbach (1988); 3 and 4: Quartz and chalcedony geothermometers of Fournier (1977), respectively; 5: δ18OSO4-H2O isotope geothermometer of Seal et al. (2000); *: Maturity index b2; +: ≤ discharge temperature; ×: No δ18OSO4 analysis. Name of the geothermal field can be seen from Tables 2 and 3.

5.5. Scaling potential Scaling caused by supersaturation of certain minerals, is a major problem encountered during the exploitation of thermal waters (Kristmannsdóttir, 1989). Cooling of thermal water due to pressure decrease or mixing with cold groundwater at shallow levels of a geothermal system might result in precipitation of various carbonate, sulfate and silica minerals. Calcite is the most common mineral with tendency to precipitate in thermal waters of high CO2 content.

Kağızman (10) Horasan (11) Pasinler Region (12, 13, 14) Erzurum Region (15, 16, 17)

İkizdere (1) Ayder (2) Borçka (3) Şavşat (5) Olur (7)

500

Therefore, in this section, calcite scaling potential of eastern Anatolian fluids is quantitatively assessed with the use of PHREEQC program (Parkhurst and Appelo, 2013). The majority of studied waters are supersaturated with respect to calcite and the degree of saturation slightly increases with increasing ionic strength which is a function of total molar concentration (Fig. 12). The presence of travertine deposits around the discharge sites of most waters implies that necessary precautions should be taken to prevent or manage the scale formation.

Tekman Region (19, 20) Diyadin Region (21, 22) Erciş Region (23, 24) Çaldıran (25) Özalp Region (26, 27, 28) Tatvan Region (30, 31)

~235 °C (93%)

~220 °C (88%) ~210 °C (78%)

~210 °C (81%)

300

~200 °C (87%)

SiO 2 (ppm)

~200 °C (85%)

SiO 2 (ppm)

SiO 2 (ppm)

400

~185 °C (73%)

~175 °C (91%)

~180 °C (79%) ~170 °C (67%)

~170 °C (72%)

200

~160 °C (72%)

~155 °C (81%) ~150 °C (66%)

~130 °C (56%)

100

(a) Northern province

0

0

200

400 600 Enthalpy (kJ/kg)

800

1000

(c) Southern province

(b) Central province 0

200

400 600 Enthalpy (kJ/kg)

800

1000

0

200

400 600 Enthalpy (kJ/kg)

800

1000

Fig. 10. Enthalpy vs. SiO2 diagrams based on quartz solubilities. The solid and dashed lines represent quartz solubility and maximum steam loss at 100 °C, respectively (Fournier and Truesdell, 1974). The numbers in the brackets (after temperature) indicate cold water mixing ratio. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

15

ERC

G 18 O(SO4 ) (‰, VSMOW)

10

KOK

COR

YUR MEM DVT T UT PAT OT G EKS HAM DLC HNS SGZ HSK GOK HNG IKI AYD HOL AKY KOT DYD BUG T AS

5

0

by a limited Plio-Quaternary volcanism (e.g., Kula volcanic field) and a series of east-west trending graben systems (e.g. Büyük Menderes, Alaşehir-Gediz, Simav) that are bounded by normal faults. Nevertheless, the majority of high-temperature geothermal reservoirs in Turkey is hosted in the Western Anatolian Graben Systems where volcanism is relatively less effective than the eastern Anatolia. The main factors that make a significant difference in the reservoir temperature of geothermal fields in eastern and western Turkey are (i) continuous thinning of the crust from east to west, and (ii) the presence of very thick impermeable cover rocks over the reservoirs within the western Anatolian graben systems. The first factor controls the surface heat flux. For the continental crust, surface heat flux is inversely correlated with crustal thickness (Condie, 1997; Bodri and Bodri, 1985). The thickness of the crust is 38–52 km in eastern Anatolia (Condie, 1997; Şengör et al., 2003; Pamukçu et al., 2007) and 36–47 km in the central Anatolia (Kuleli et al., 2004; Bekler et al., 2005). However, the crust is only 28–30 km thick in the western Anatolia and the thickness is gradually lowered to 25 km under the Aegean Sea (Zhu et al., 2006). Because of relatively thin crust in the western Anatolia, the surface heat flux is higher than the mean values recorded elsewhere in Turkey (74 mWm−2; Akın et al., 2014) and globally (81 mWm−2; Condie, 1997). For the western Anatolia, the average surface heat flux is 107 mWm−2 (Ilkişik, 1995) that is estimated from the silica temperatures of geothermal springs. On the other hand, based on Curie point depth from aeromagnetic maps, Pamukçu et al. (2014) predicted the surface heat flux in eastern Anatolia in the range of 20 to 100 mWm−2. Additionally, Akın et al. (2014) indicated that the average heat flux around Erzurum, Kars, eastern Black Sea and Van regions is 61 mWm−2 (min-max: 41–116 mWm−2), 54 mWm−2 (50–80 mWm−2), 52 mWm−2 (45–63 mWm−2), and 62 mWm−2 (42–87 mWm−2), respectively. Hence, the occurrence of moderate/low-temperature geothermal systems in the region is believed to be associated with relatively high crustal thickness and low surface heat flux. The second factor controls the heat loss from the reservoir and/or the rate of cold water mixing along the flow path. The cover rocks of geothermal fields in the western Anatolia are comprised by impermeable sedimentary units with thickness of 1.5–2 km (e.g. Paton, 1992; Sarı and Şalk, 2006) preserving the temperature of geothermal

CAM

BHS

CAY

HAD

AYR

-5

-10

-16

-14

-12

-10 -8 -6 G 18 O(H2 O) (‰, VSMOW)

-4

-2

0

Fig. 11. δ18OH2O vs. δ18OSO4 plot for geothermal waters in eastern Turkey. The red triangle and blue square represent mixing and sulfate-water oxygen exchange process which effects sulfate isotopic composition of samples, respectively. The dot lines represent temperature trajectories for equilibrium fractionation of δ18O between SO4 and H2O (Seal et al., 2000). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

5.6. Geological factors controlling the development of geothermal systems in Eastern Anatolia The reservoir temperatures of studied waters are generally lower than those of high-temperature geothermal systems in the Western Anatolian graben systems (Mutlu and Güleç, 1998) (e.g. Alaşehir 159–287 °C; Germencik 203–276 °C; Kizildere 148–242 °C; Salavatli 148–211 °C). This is attributed to the differences in structural and geological conditions (crustal thickness, presence of cover rock, tectonic regime) that control the formation of geothermal systems in the eastern and western Anatolian regions. High-temperature geothermal fields are known to occur along active volcanoes and in regions of tectonic unrest (e.g. Moeck, 2014; Chandrasekharam et al., 2016). The eastern Turkey is represented by active tectonism (e.g., Bitlis-Zagros Thrust Zone, North Anatolian Fault, East Anatolian Fault) and recent volcanism (e.g., Nemrut, Süphan, Tendürek, Ararat) whereas the western Anatolia is signified

2.5

TAT

2

CAY KRZ

1.5

Log SICalcite

0 -0.5 OTG

-1.5

IKI

UZA TAS AYR DYD ERC CAM DVT DIB HOL HAD KOP M EM HSK NHL YUR KOC HRM HAM HNS GOK M LK CKR KOK KOT TAZ ASB AZY COR ARZ TUT SUS AKK AKY DLC PAT EKS AYS ALA

AYD

0.5

-1

SGZ

Ca+Mg-HCO3

-2

Na-HCO 3

BUR

SO4 -rich Na-Cl

-2.5 -3 0.001

HNG

GRM BUG

BHS

1

15

0.01

Ionic Strength

0.1

1

Fig. 12. The plot of saturation index (calcite) vs. ionic strength for the geothermal waters in eastern Turkey. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708

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H. Aydin et al. / Journal of Volcanology and Geothermal Research xxx (xxxx) xxx

reservoirs. Differently, intensely fractured Pliocene-Quaternary volcanics of 0.8–1 km thickness (e.g. Innocenti et al., 1980, 1982) are the common cover units in the eastern Turkey (Fig. 1). It is thought that high secondary porosity and the absence of impermeable lithologies make geothermal reservoirs in the eastern Anatolia more vulnerable to cooling via mixing with cold water (Fig. 10). It is likely that cold waters penetrating the fractured volcanic rocks are mixed with deeply circulated thermal waters which eventually resulted in cooling and dilution of springs. Consequently, both the relatively thick crust and the low surface heat flux together with the absence of ideal cover units in eastern Turkey has resulted in the development of moderate/ low-temperature geothermal systems. 6. Conclusion The eastern Turkey has unique geological setting that hosts both Neogene-Quaternary volcanism and active tectonism providing a potentially rewarding locality to investigate the relationship between tectonomagmatic activity and geothermal systems. In this study processes controlling the hydrogeochemical and isotopic characteristics of thermal and cold waters collected from 31 geothermal fields in eastern Turkey were investigated and the main results are listed below. The geothermal systems in eastern Turkey comprise several thermal springs with discharge temperatures ranging from 24 to 65 °C. Thermal waters in the region are classified into four distinct groups based on their major ion compositions. Ca + Mg-HCO3 type waters are affected by the dissolution of carbonate and mafic minerals. Na-HCO3 and NaCl + HCO3 are the most common water types in the region. Composition of Na-HCO3 type water is essentially controlled by feldspar dissolution or calcite precipitation, while Na-Cl + HCO3 type water is generated consistently by the enrichment in chloride along the deep flow path. The SO4-rich type waters are observed only at four localities (BUR, EBU, KOK, OTG). Sample EBU from Diyadin area has the lowest pH in the region, and the SO4-rich character of this water is probably originated by the dissolution of H2S in an oxygenated shallow groundwater. Other SO4-rich waters have slightly acidic or neutral character and their chemical composition is controlled by oxidation of sulfur-bearing minerals. The cold waters are represented by Ca-HCO3 type. The origin of thermal waters is investigated by stable isotope (δ18O, δ2H, δ34SSO4 and δ18OSO4) and tritium (3H) compositions. Considering the size of the study area and the topographical differences, three different local meteoric water lines (LMWL) are proposed for the northern, central and southern provinces with deuterium excess of 15.0, 13.9 and 16.5‰ V-SMOW, respectively. The oxygen and hydrogen isotope compositions of thermal waters indicate a meteoric origin and tritium data (~0 TU) imply deep circulation. A remarkable δ18O shift (7.7‰) in sample CAY (Özalp) in the southern province and moderate δ18O shifts (1.5 to 2.38‰) recognized in several thermal waters (e.g. HNG, HNS, KOT, DVT, DYD, AYR, CAM, COR, HAD and IKI) are attributed to water– rock interaction process and 18O-exchange process between CO2 and H2O along deep flow paths. Except for waters in the Erzurum region and northeast part of Lake Van, the sulfur isotope composition of thermal waters is controlled by mixing between the terrestrial and marine evaporites. The enrichment of δ18O in sulfate phase is remarkable for thermal waters from northeast part of Lake Van (AYR, BUG, DYD, HAD, TAS). Reservoir temperatures calculated by K\\Mg (27–127 °C), silica (29–179 °C) and 18OSO4-H2O (51–196 °C) geothermometers vary in a wide range. The Na\\K geothermometer estimates a higher range of temperatures for samples from the northern (e.g. BHS, COR, SUS), central (e.g. AZY, HNS, HNG) and southern provinces (e.g. CAY, GRM): 100–141 °C, 109–156 °C, and 130–200 °C, respectively. Temperatures computed by the silica-enthalpy model are in range of 150 to 200 °C, 170 to 235 °C and 130 to 210 °C for the respective provinces. The chemical geothermometers and silica-enthalpy model yielded inconsistent reservoir temperatures suggesting that results of geothermometers

should be used cautiously. However, a detailed geophysical survey (e.g. electrical sounding, self-potential) is required to delineate depth, extent and fluid temperature of the reservoir. Consequently, results of this study revealed no concrete sign of high-temperature (N 225 °C) reservoirs in the eastern Turkey. It is thought that the relatively thick crust and low surface heat flux together with the absence of ideal cover units in the region has resulted in the development of moderate or lowtemperature geothermal systems. Declaration of competing interest 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. Acknowledgements This study was supported by the Scientific and Technological Research Council of Turkey (TUBITAK) (under grant #114Y067) and by the Eskişehir Osmangazi University (under grant #2009/15017). The authors are grateful to Franco Tassi and an anonymous reviewer for their critical comments and suggestions that helped to improve the manuscript. Appendix A. Supplementary data Supplementary file contains data for the construction of local meteoric water lines for the northern and central provinces in eastern Turkey. Supplementary data to this article can be found online at https://doi.org/10.1016/j.jvolgeores.2019.106708. References Abdioğlu, E., Arslan, M., Aydinçakir, D., Gündoğan, I., Helvaci, C., 2015. Stratigraphy, mineralogy and depositional environment of the evaporite unit in the Aşkale (Erzurum) sub-basin, Eastern Anatolia (Turkey). J. African Earth Sci. 111, 100–112. https://doi. org/10.1016/j.jafrearsci.2015.07.013. Açlan, M., Altun, Y., 2018. Syn-collisional I-type Esenköy Pluton (Eastern AnatoliaTurkey): an indication for collision between Arabian and Eurasian plates. J. African Earth Sci. 142, 1–11. https://doi.org/10.1016/j.jafrearsci.2018.02.019. Akın, U., Ulugergerli, E.U., Kutlu, S., 2014. The assessment of geothermal potential of Turkey by means of heat flow estimation. Bull. Miner. Res. Explor. 149, 201–210. https://doi.org/10.19111/bmre.58938. Akkuş, İ., Akıllı, H., Ceyhan, S., Dilemre, A., Tekin, Z., 2005. Potential of Geothermal Resources of Turkey. MTA Book Series 201 p. 849 Ankara. (in Turkish). Altınlı, İ.E., Pamir, H.N., Erentöz, C., 1963. 1/500.000 scale Geological map of Turkey-Erzurum, MTA Institution, p. 131, Ankara (in Turkish). Altınlı, İ.E., Pamir, H.N., Erentöz, C., 1964. 1/500.000 scale Geological map of Turkey-Van, MTA Institution, p. 90, Ankara (in Turkish). Apaydın, A., 2018. Changes on stable isotope contents of precipitation with elevation: An evaluation on some case studies in Turkey. HYDRO’2018-National Symposium on Hydrogeology and Water Resources, Ankara, Turkey, pp. 66–74 (in Turkish). Apollaro, C., Vespasiano, G., Muto, F., De Rosa, R., Barca, D., Marini, L., 2016. Use of mean residence time of water, flowrate, and equilibrium temperature indicated by water geothermometers to rank geothermal resources. Application to the thermal water circuits of Northern Calabria. J. Volcanol. Geotherm. Res. 328, 147–158. https://doi. org/10.1016/j.jvolgeores.2016.10.014. Arnórsson, S., 2000. The quartz and Na/K geothermometer: I. New thermodynamic calibration. World Geothermal Congress 2000, Kyushu–Tohoku, Japan, pp. 929–934. Aydın, H., Ekmekçi, M., Tezcan, L., Dişli, E., Aksoy, L., Yalçın, M.P., Özcan, G., 2009. Assessment of water resources potential of Gürpınar (Van) karst springs with regard to sustainable management. TUBITAK Project Report (no: 106Y040), Van, Turkey (in Turkish). Aydın, H., Mutlu, H., Kazancı, A., 2013. Hydrogeochemical properties of Çaldıran (Van) geothermal field. XI. Ulusal Tesisat Mühendisliği Kongresi ve Sergisi (TESKON 2013) Jeotermal Semineri, Izmir, Turkey, pp. 71–90 (in Turkish). Aydın, H., Hilton, D.R., Güleç, N., Mutlu, H., 2015. Post-earthquake anomalies in He-CO2 isotope and relative abundance systematics of thermal waters: the case of the 2011 Van earthquake, eastern Anatolia. Turkey. Chem. Geol. 411, 1–11. https://doi.org/ 10.1016/j.chemgeo.2015.06.019. Aydın, H., Karakuş, H., Mutlu, H., 2017. Hydrogeochemical and isotopic investigation of geothermal systems of north eastern Anatolia. TUBITAK Project Report (no: 114Y067), Van, Turkey (in Turkish). Baba, A., Sözbilir, H., 2012. Source of arsenic based on geological and hydrogeochemical properties of geothermal systems in Western Turkey. Chem. Geol. 334, 364–377. https://doi.org/10.1016/j.chemgeo.2012.06.006.

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Please cite this article as: H. Aydin, H. Karakuş and H. Mutlu, Hydrogeochemistry of geothermal waters in eastern Turkey: Geochemical and isotopic constraints on wa..., Journal of Volcanology and Geothermal Research, https://doi.org/10.1016/j.jvolgeores.2019.106708