Origin, evolution and geothermometry of the thermal ...

Journal of Volcanology and Geothermal Research 349 (2018) 1–30

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Origin, evolution and geothermometry of the thermal waters in the Gölemezli Geothermal Field, Denizli Basin (SW Anatolia, Turkey) Hülya Alçiçek a,⁎, Ali Bülbül a, Andrea Brogi b, Domenico Liotta b,c, Giovanni Ruggieri d, Enrico Capezzuoli e, Marco Meccheri f, İbrahim Yavuzer g, Mehmet Cihat Alçiçek a a

Department of Geological Engineering, Pamukkale University, TR-20070 Denizli, Turkey University of Bari, Department of Earth and Geoenvironmental Sciences, 70125 Bari, Italy IGG-CNR, Institute of Geosciences and Earth Resources, Via G. Moruzzi, 1, 56124 Pisa, Italy d IGG-CNR, Institute of Geosciences and Earth Resources, Via G. La Pira, 4, 50121 Firenze, Italy e University of Perugia, Department of Physics and Geology, 06123 Perugia, Italy f University of Siena, Department of Physics, Earth and Environmental Sciences, 53100 Siena, Italy g Denizli Governorship, Natural Sources and Cultural Heritage Directorate, TR-20070 Denizli, Turkey b c

a r t i c l e

i n f o

Article history: Received 6 February 2017 Received in revised form 26 July 2017 Accepted 28 July 2017 Available online 31 July 2017 Keywords: Hydrogeochemistry Isotopes Metamorphic CO2 Mixing process Water-rock interactions Tectonics

a b s t r a c t The Gölemezli Geothermal Field (GGF) is one of the best known geothermal fields in western Anatolia (Turkey). The exploited fluids are of meteoric origin, mixed with deep magmatic fluids, which interacted with the metamorphic rocks of the Menderes Massif. The geothermal fluids are channeled along Quaternary faults belonging to the main normal faults system delimiting the northern side of the Denizli Basin and their associated transfer zones. In this study, hydrochemical and isotopic analyses of the thermal and cold waters allow us to determine water-rock interactions, fluid paths and mixing processes. Two groups of thermal waters have been distinguished: (i) Group 1A, comprising Na-SO4 type and Ca-SO4 type and (ii) Group 1B, only consisting Ca-HCO3 type waters. Differently, two groups were recognized in the cold waters: (i) Group 2A, corresponding to Ca-HCO3 type and (ii) Group 2B, including Mg-HCO3 type. Their geochemical characteristics indicate interactions with the Paleozoic metamorphic rocks of the Menderes Massif and with the Neogene lacustrine sedimentary rocks. Dissolution of host rock and ion-exchange reactions modify thermal water composition in the reservoir of the GGF. High correlation in some ionic ratios and high concentrations of some minor elements suggest an enhanced water-rock interaction. None of the thermal waters has been reached a complete chemical re-equilibrium, possibly as a result of mixing with cold water during their pathways. Geothermal reservoir temperatures are calculated in the range of 130-210°C for the Gölemezli field. Very negative δ18O and δ2H isotopic ratios are respectively between −8.37 and −8.13‰ and −61.09 and −59.34‰ for the SO4-rich thermal waters, and ca. −8.40 and −8.32‰ and −57.80 and −57.41‰ for the HCO3-rich thermal waters. Low tritium (b1 TU) and low oxygen isotope values reflect a deep circuit and fluids of meteoric origin. Positive δ13CDIC ratios (+5.11 to +7.54‰) of all thermal waters imply a contribution of metamorphic origin. Heating is guaranteed by a deep circuit within an overheated continental crust, mainly affected by damaged rock volumes. Volatile ascent from deep magmatic sources through crustal structures can explain the occurrence of mantle volatiles at shallow depth in the Denizli Basin. The NW- and NE-trending fault systems, associated with their related fractures, played as hydraulic conduits underlining the strict link existing between fractures and fluid convection in the extensional settings. In this view, the GGF is a very good example of geothermal field associated to active tectonic setting and magmatism, as it is the case of the other geothermal fields occurring in the Denizli Basin. © 2017 Elsevier B.V. All rights reserved.

1. Introduction Western Anatolia (Fig. 1A) is characterized by low (20–70 °C), moderate (70–150 °C) and high (N150 °C) temperature of thermal fluids in the geothermal fields (Fig. 1B; Table 1). The Denizli Basin, located in ⁎ Corresponding author. E-mail address: [email protected] (H. Alçiçek).

https://doi.org/10.1016/j.jvolgeores.2017.07.021 0377-0273/© 2017 Elsevier B.V. All rights reserved.

the SW Anatolia, coincides with a wide geothermal area characterized by low- to high-temperature geothermal fields (Fig. 1B; Table 2). These are mostly aligned along the WNW-ESE trending Neogene-Quaternary normal fault systems related to the ongoing N-S extensional tectonic direction (e.g., Dumont et al., 1979; Şimşek, 1984; Altunel, 1994; Şimşek, 2003a, 2003b; Vengosh et al., 2002; Güleç et al., 2002; Güleç and Hilton, 2006, 2016). The Denizli Basin is well known for its thermal springs (i.e., Kızıldere, Tekkehamam, Yenice, Gölemezli, Karahayıt and

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H. Alçiçek et al. / Journal of Volcanology and Geothermal Research 349 (2018) 1–30

Fig. 1. (A) Location map of the Denizli Basin, SW Turkey (Bozkurt, 2003); (B) Schematic geological map of the squared area, indicated in (A). The main extensional basins of western Anatolia are shown together with location and classification (based on discharge temperatures) of the geothermal fields (based on Baba and Sözbilir, 2012; Akkuş et al., 2005).


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Table 1 Location map of geothermal fields in the western Anatolia (based on Akkuş et al., 2005). Çanakkale geothermal fields: (Ç1) Hıdırlar-Uyuz (40-84°C) (Ç2) Tuzla (51-174°C) (Ç3) Kestanbol-Akçakeçili (33-75°C) (Ç4) Çan (42-49°C) (Ç5) Ayvacık-Küçükçetmi (40-42°C) (Ç6) Biga-Kırkgeçit (52-58°C) (Ç7) Çan-Etili-Tepeköy-Bardakçılar (33-59°C) (Ç8) Lapseki-Kocabaşlar (36°C) (Ç9) Topaklar (Kum) (45-67°C) (Ç10) Bayramiç-Yukarı Palamut (35°C) (Ç11) Bayramiç-Külcüler (26-35°C) Balıkesir geothermal fields: (B1) Güre (33-58°C) (B2) Havran-Derman (25-65°C) (B3) Gönen (34-84°C) (B4) Kepekler (31-64°C) (B5) Kızık (49°C) (B6) Balya-Ilıca (Şamlı) (51-60°C) (B7) Pamukçu (26-59°C) (B8) Hisaralan (27-99°C) (B9) Hisarköy (25-98°C) (B10) Pelitköy (26-31°C) (B11) Ayvalık-Ilıca (31-34°C) (B12) İvrindi-Bozören (27°C) (B13) İvrindi-Ilıca-Gümeli (29-39°C) (B14) Kepsut-Eşeler (26-28°C) (B15) Savaştepe-Kirazköy Dağ (31-40°C) (B16) Susurluk-Gökçedere-Ömerköy (25-31°C) (B17) Susurluk-Yıldız (56-75°C) (B18) Emendre (33°C) Uşak geothermal fields: (U1) Emirfakı-Akbulak (31-38°C) (U2) Banaz-Hamamboğazı (37-72°C) (U3) Banaz-Kızılcaören (25°C) (U4) Ulubey-Aksaz (36-37°C) (U5) Ulubey-Hasköy (34°C) (U6) Örencik (30-34°C)

Manisa geothermal fields: (MA1) Turgutlu-Urganlı (51-80°C) (MA2) Salihli-Kurşunlu (35-168°C) (MA3) Alaşehir-Horzumsazdere (30-213°C) (MA4) Köprübaşı-Saraycık (30-74°C) (MA5) Kula-Emir-Şehitli (30-163°C) (MA6) Demirci-Eskihisar (29-42°C) (MA7) Alaşehir-Sarıkız (27°C) (MA8) Soma-Menteşe (57°C) (MA9) Sarıgöl (26°C) Kütahya geothermal fields: (K1) Simav-Eynal-Çitgöl-Naşa (43-162°C) (K2) Gediz-Abide (65-97°C) (K3) Muratdağı (37-39°C) (K4) Yoncalı (32-43°C) (K5) Emet (39-49°C) (K6) Yeniceköy (41°C) (K7) Dereli (40-42°C) (K8) Göbel (34°C) (K9) Ilıca (Harlek) (29-41°C) (K10) Hisarcık (40-51°C) (K11) Şaphane (26-46°C) İzmir geothermal fields: (İ1) Balçova (80-140°C) (İ2) Seferihisar (33-153°C) (İ3) Çeşme-Şifne (35-57°C) (İ4) Aliağa (İzmir) (55-96°C) (İ5) Aliağa-Reşadiye (29°C) (İ6) Bayındır-Ergenli (42-47°C) (İ7) Urla-Gülbahçe (33°C) (İ8-11) Bergama-Sucahlı-PaşaköyMahmudiye-Güzellik (26-57°C) (İ12) Dikili-Madra-BahçeliköyParastallı-Nebiler (25-57°C) (İ13) Dikili-Karadere-Çoban-KaynarcaBademli-Kocaoba (42-132°C)

Pamukkale fields; Fig. 2; Table 2) and for its geothermal potentiality (e.g., Şimşek, 1981, 1984, 1985; Gökgöz, 1994, 1998). In the last decades, the geothermal resources in this basin have been increasingly exploited and their features have been studied since 1967 (e.g., Şimşek, 1981, 1984, 1985; Gökgöz, 1994, 1998; Özgür, 2002; Vengosh et al., 2002; Şimşek, 2003a, 2003b; Şimşek et al., 2005; Alçiçek et al., 2016; Brogi et al., 2014, 2016; Tarcan et al., 2016 and references therein). The Gölemezli Geothermal Field (hereafter GGF) is one of the highest temperature geothermal fields occurring in the Denizli Basin and located between the Yenice and Karahayıt geothermal fields (Figs. 1B and 2). The GGF and its surrounding geothermal fields (e.g., Yenice, Karahayıt and Pamukkale; Fig. 1B) are well known for their colorful travertine formations (white color at Pamukkale and red color at Karahayıt) depositing nearby still active isolated thermal springs (Fig. 2) and/or along fissure-ridges (Bargar, 1978; Brogi and Capezzuoli, 2009; Altunel, 1994; Altunel and Hancock, 1993a, 1993b, 1996; Brogi et al., 2014, 2016; TUBITAK, 2016). Although almost 70% of the geothermal wells have been drilled in western Anatolia (Güleç and Hilton, 2016) where fluids feed greenhouses, thermal spas, geothermal power plants and carbon dioxide production sites, the understanding of the thermal water paths, their origin and recharge areas are still scarcely investigated. Several informations are still lacking in order to better plan sustainable exploitation and to prevent the economic risk of unproductive drillings. Many geological and hydrogeological contributions have been primarily focused on the hydrochemical properties, isotopic compositions and origin of thermal waters in the western Turkey

Afyon geothermal fields: (AF1) Ömer-Gecek-Kızık-Uyuz (46-99°C) (AF2) Gazlıgöl (38-74°C) (AF3) Heybeli-Çay (30-57°C) (AF4) Sandıklı (45-71°C) Aydın geothermal fields: (AY1-2) Ilıcabaşı-İmamköy (34-142°C) (AY3-5) Germencik-ÖmerbeyliBozköy-Çamur (36-232°C) (AY6-8) Sultanhisar-SalavatlıMalgaçemir-Güvendik (30-172°C) (AY9-10) Gümüş-Söke-Sazlıköy (27-39°C) (AY11) Buharkent-Ortakçı (51°C) (AY12) Nazilli-Gedik (32°C) (AY13) Kuşadası (26°C) (AY14) Davutlar (26-42°C) Denizli geothermal fields: (D1) Kızıldere (155-241°C) (D2) Babacık-Demirtaş (40-100°C) (D3) Tekkehamam-İnaltı-Uyuz (29-168°C) (D4) Bölmekaya (36°C) (D5) Yenice (Çizmeli)-Kamara (27-67°C) (D6) Gölemezli (50-88°C) (D7) Buldan-Efe (25°C) (D8) Karahayıt (35-51°C) (D9) Pamukkale (35°C) (D10) Çardak-Beylerli-Ilıcapınar (31-40°C) Muğla geothermal fields: (MU1) Yatağan-Bozhüyük (35-37°C) (MU2) Köyceğiz-Sultaniye-DelibeyRızaçavuş-Gelgirme (34-41°C) (MU3) Bodrum-Karaada (32°C) (MU4) Fethiye-Gebeler (37°C) (MU5) Ortaca-Çürükardı (29°C) (MU6) Datça-Gölbaşı-Ilıca (26-28°C)

geothermal systems (e.g., Mutlu and Güleç, 1998; Gemici and Tarcan, 2002; Vengosh et al., 2002; Tarcan, 2004; Şimşek, 2003a, 2003b; Şimşek et al., 2005; Baba and Sözbilir, 2012; Karakuş and Şimşek, 2013; Alçiçek et al., 2016; Güleç and Hilton, 2006, 2016). At the same time, hydrogeological and hydrochemical properties, and regional geological and tectonic settings of hydrothermal activities of some important geothermal fields of the Denizli Basin have been also studied in detail (e.g., Şimşek, 1984, 2003b; Şimşek et al., 2005; Çakır, 1999; Dilsiz, 2006; Uysal et al., 2007, 2009; Kele et al., 2011; De Filippis et al., 2013; Alçiçek et al., 2013; Brogi et al., 2014, 2016; Alçiçek et al., 2016). As it regards the study area, several studies were focused on geological, hydrogeological and hydrochemical topics (e.g., Şimşek, 1981, 1984, 1985; Gökgöz, 1994, 1998; Şimşek, 2003a, 2003b; Möller et al., 2004, 2008). Nevertheless, an integrated approach is still necessary. The aim of this paper is therefore to contribute to fill this gap by: (i) reviewing the tectonic setting, in order to contribute to the understanding of the relationships between the main structures and the thermal fluid-flow, (ii) defining the hydrogeological and hydrogeochemical properties of the thermal and cold waters, (iii) computing the maximum reservoir temperatures by means of several geothermometers, based on previous and new data, (iv) proposing a conceptual model of the GGF geothermal system, using different proxies and (v) evaluating the aquifer fluid composition and fluid-mineral equilibria. Furthermore, this integrated approach provides information for a better understanding of the drivers affecting the chemical composition of the thermal waters.

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Table 2 Hydrogeological properties of geothermal fields from the Denizli Basin (based on Akkuş et al., 2005; Yılmazer, 2009). Geothermal field

Temperature (°C)

Depth (m)

Flow (l/sec)

Geothermal potential (MWt)

Yenice (springs) Yenice (wells) Gölemezli (springs) Gölemezli (wells) Karahayıt (springs) Karahayıt (wells) Pamukkale (springs) Bölmekaya (spring) Buldan-Efe (spring) Kızıldere (wells) Tosunlar (wells) Babacık (springs) Demirtaş-Karataş Tekkehamam (springs) Tekkehamam (wells) İnaltı-Uyuz (springs) Ilıcapınar-Beylerli (spring) Ilıcapınar-Beylerli (well)

27–36 36–67 50–57 62–88 27–51 58–60 35 36 25 155–241 80–98 62 40–100 29–97.2 116–168 83.5–98 31 39.5

– 54–250 – 549–1500 – 452–568 – – – 368–2261 918–2653 – – – 615.5–2001 – – 53

0.3 4–140 6 15–140 0.5–75 15–80 30–130 0.2 0.6 81–186 39–144 2 – 30 11.6–15 30 – 28

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2. Geological outline Western Anatolia (Turkey) (Fig. 1A and 1B) is characterized by active extensional tectonics (McKenzie, 1978; Şengör and Yılmaz, 1981; Bozkurt, 2003; Kaymakçı, 2006; Ten Veen et al., 2009), thinned continental crust and lithosphere (e.g., Jolivet et al., 2013; Faccenna et al., 2014) and widespread geothermal anomalies (e.g., Şimşek, 1984; Güleç et al., 2002; Güleç and Hilton, 2006, 2016; Wiersberg et al., 2011) related to the emplacement of Neogene-Quaternary magmatic bodies at shallow depth (Vengosh et al., 2002; Güleç et al., 2002; Güleç and Hilton, 2006; Menant et al., 2016). The regional thermal flow in western Anatolia is very high, up to 2.6 HFU (~ 107 mW/m2; İlkışık, 1995), beyond the world average value of 1.5 HFU (60 mW/m2; Cermac and Hurtig, 1979). In western Anatolian domain, three magmatic associations are distinguished (e.g., Innocenti et al., 2005; Güleç and Hilton, 2006): (i) early-middle Miocene calc-alkaline and shoshonitic rocks, linked to the subducting oceanic litosphere (e.g., Francalanci et al., 2000); (ii) middle Miocene ultrapotassic shoshonitic and lamproidic rocks, derived from crustal streching (e.g., Francalanci et al., 2000; Doglioni et al., 2002; Innocenti et al., 2005) and (iii) late Miocene and late PlioceneQuaternary sodic alkaline rocks, accounted from a sub-lithospheric and lithospheric continental magmatic source, within the regional geodynamic frame of the African plate subduction (e.g., Doglioni et al., 2002; Innocenti et al., 2005). Such a typically calc-alkaline to shoshonitic volcanic centers are diffuse in the SE margin of the Denizli Basin (4.88–6.28 Ma; Ercan et al., 1983, 1985; Prelević et al., 2015; Semiz et al., 2012), as well as in the Söke Basin (6.99 Ma; Ercan et al., 1985), the Acıgöl Basin (5 ka; Sulpizio et al., 2013; Athanassas et al., 2017), the Bodrum domain (7.9-9.7 Ma; Ercan, 1982; Prelević et al., 2015), Bucak domain of the Burdur Basin (4.46–4.70 Ma) and Gölcük domain of the Isparta Basin (4.07–4.60 Ma; Prelević et al., 2015), the Selendi Basin (15.8 Ma; Ersoy and Helvacı, 2007) and the Uşak-Güre Basin (15.5 Ma; Karaoğlu and Helvacı, 2014) (Fig. 1B). The presence of more than 600 hot springs in western Anatolia denotes high permeability as an expression of the fracture network dissecting the metamorphic substratum (Çağlar, 1961; Akkuş et al., 2005). Low to high-temperature geothermal fields have been exploited since the 1960's and are mainly concentrated along the major bounding faults delimiting the structural depressions (Fig. 1B; Şimşek, 1981, 1984; Vengosh et al., 2002; Güleç et al., 2002) where thermal springs (up to 100 °C) are considered as the superficial expression of geothermal anomalies. These thermal waters reflect a significant geothermal potential in western Anatolia that has been mined through drillings by

58 16

140 – 2035 – 1037

–

the General Directorate of Mineral Research and Exploration of Turkey (MTA) (Çağlar, 1961; Akkuş et al., 2005). In the Denizli Basin, the geothermal exploration has been initiated in the Kızıldere Geothermal Field by the MTA-UNDP project during 1968 (Şimşek, 2003a, 2003b). This field has been used for carbon dioxide and electricity production. The Neogene Denizli Basin is 50 km wide and 70 km long. It is delimited by NW-trending faults, located at the junction with the E-trending Büyük Menderes and the NW-trending Gediz basins (Figs. 1B and 2; Şimşek, 1984; Sun, 1990; Koçyiğit, 2005; Kaymakçı, 2006; Alçiçek et al., 2007). The Neogene to Quaternary sedimentary succession of the basin is up to 1300 m thick and overlies unconformably the substratum made up of metamorphic rocks belonging to the Menderes Massif (Precambrian and Paleozoic-Mesozoic) and sedimentary and ophiolitic rocks of the Lycian Nappes (Mesozoic), the latter representing the westernmost part of the Tauride Orogen (Okay, 1989, 2001; Sun, 1990; Konak and Şenel, 2002; Alçiçek et al., 2007). The Menderes Massif includes a crystalline core and its metasedimentary cover exposed in a core-complex structure (Pamir and Erentöz, 1974; Şengör and Yılmaz, 1981; Okay, 1989; Sun, 1990; Bozkurt, 2001; Ten Veen et al., 2009; van Hinsbergen, 2010; van Hinsbergen and Schmid, 2012). In the Denizli Basin, the PrecambrianCambrian crystalline core rocks, up to 1000 m thick, are made up of augen gneiss, deformed metagranite, migmatite, gabbro with granulite and eclogite relics, and medium- to high-grade metamorphic schists (Şengör et al., 1984; Sun, 1990; Okay, 1989, 2001). The cover rocks contain three sequences (Okay, 1989, 2001; Sun, 1990; Gündoğan et al., 2008): (i) late Devonian-early Carboniferous of the Ortaköy Formation, composed of monotonous, brownish garnet micaschist, micaschist and quartzo-feldispatic schist (N 2000 m thick); (ii) Permo-Carboniferous Göktepe (İğdecik) Formation, made up of heterogeneous series of marble, metaquartzite, micaschist and phyllite, and thickness of 1400 m; (iii) Lycian Nappes, composed of Mesozoic dolomitic limestone, evaporite-dolomite, metacarbonate, and turbiditic sandstone, tectonically overlain by ophiolitic mélanges (Fig. 3). The Denizli Basin succession is lithostratigraphically described as the Denizli Group by Şimşek (1984) and biochronologically determined by Alçiçek et al. (2007) on the basis of terrestrial mammal fossil fauna (e.g., Saraç, 2003; Kaymakçı, 2006). The Denizli Group is subdivided into alluvial-fan to fluvial Kızılburun (early-middle Miocene), lacustrine Sazak (middle-late Miocene), lacustrine to fluvio-lacustrine Kolankaya (late Miocene-early Pleistocene) and alluvial-fan to fluvial Tosunlar (middle-late Pleistocene) formations (Fig. 3). The rocks of the Sazak and Kolankaya formations and Menderes Massif constitute the reservoir unit (shallow and deep reservoirs, respectively) in the


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Fig. 2. Geological map of the Denizli Basin, SW Turkey (based on Sun, 1990 and adapted from Konak and Şenel, 2002) with locations of the GGF and sampled waters.

GGF (Fig. 3), as well as in other geothermal fields of the Denizli Basin and its surroundings (Fig. 2). 2.1. Tectonic setting The Denizli Basin was initially subsided as a half graben during early Miocene along the NW-SE trending Babadağ Fault (Alçiçek et al., 2007). During early Quaternary, initial half graben was turned into a full graben due to the subsequent activity of the Pamukkale, Akköy and Tripolis (Yenice) faults, terminating the Denizli Basin to the north along which hot springs induced travertine/tufa precipitation (Altunel and Hancock, 1993a, 1993b; Altunel, 1994; Hancock et al., 1999; Çakır, 1999; Alçiçek et al., 2007; Brogi et al., 2014, 2016; Alçiçek et al., 2016; TUBITAK, 2016). It is suggested that travertine masses were formed where dip-slip normal fault segments has been displayed as the stepover zones (e.g., Şimşek, 1984; Çakır, 1999; Kaymakçı, 2006) and at the intersection with transfer zones delimiting areas with differentiated amount of extension (Brogi et al., 2014, 2016). In the Denizli Basin, high values of permeability are commonly associated to highly fractured rock volumes (e.g., Şimşek et al., 2005). Most of the geothermal fields in the Denizli Basin are associated to the NWtrending fault traces (Fig. 2; Table 2). Thermal springs are located in a wide area where the Tripolis normal fault segment is interrupted by a nearby orthogonal (NE-trending) transtentional fault segment crossing Büyük Menderes Valley, where the Kamara fissure ridge is located (Fig. 2). The northern part of the Denizli Basin is delimited by NW-trending normal faults, composed by a set of normal fault segments dipping toward SW. In the Gölemezli area, these faults display off-sets of several tens of meters and juxtapose the metamorphic rocks of the Menderes Massif complex to the Neogene-Quaternary basin-fill succession. In

some parts, the fault segments define step-over and relay ramps between overlapping fault segments such as it is the case (Fig. 2) of the Akköy and Tripolis segments (Çakır, 1999). Nevertheless, these faults are locally interrupted and/or dissected by NE-trending fault zones, determining an increasing of the fracture density at their intersection with the NW-trending faults (Kaymakçı, 2006; Brogi et al., 2014). In the Gölemezli domain, two fault systems demonstrate different kinematics (Brogi et al., 2016): the NE-trending faults are characterized by dominant oblique-slip movements and subsequent normal movements mainly developed in a later stage. Differently, the NW-trending faults are characterized by dominant normal movements although an oblique-slip component can be recognized as a subsequent faulting activity. On the basis of these cross-cutting relationships, both systems are considered to be active during Miocene and Quaternary (Alçiçek et al., 2013; Brogi et al., 2016). During Pleistocene and Holocene, hydrothermal fluid flows were concentrated along the fault segments belonging to both fault systems, as indicated by the location and structural features of the travertine bodies (fissure-ridges and mounds) and the thermal springs (Altunel and Hancock, 1993a, 1993b; Altunel, 1996; Çakır, 1999; Brogi et al., 2014, 2016; Alçiçek et al., 2016).

3. Hydrogeology The GGF is one of the moderate temperature geothermal fields in the Denizli Basin (Fig. 2). This field lies between longitudes 4204000 to 4210000 N and latitudes 0675000 to 0680000 E, with an area of ~ 15 km2. The climate of the area is semi-arid, with rainfall maximum in autumn and winter, and minimum during the summer. The mean annual values for precipitation and temperature are 584 mm and 15 °C, respectively.

6


In the GGF, one thermal spring and several thermal wells are situated on the Tripolis (Yenice) fault footwall (Fig. 2). The spring discharges 6 l/s of water at 42 °C. Five exploration wells (DG1-5; Fig. 4A–E) were drilled by the General Directorate of Mineral Research and Exploration of Turkey (MTA) during the 2001–2003 (Akkuş et al., 2005; Akıllı and Bülbül, 2006). In addition, two exploration wells were also drilled by the Denizli Governorship during 2008 (Göl-1 well; Fig. 4F) and 2010 (Göl-2 well; Fig. 4G). Moreover, additional wells were also drilled by private enterprises, reaching depths of 140–600 m. All GGF wells are commonly used for spa and greenhouse facilities. In the GGF, two rock units (the Neogene formations and Menderes Massif) can display local high permeability and these are considered as reservoirs (Fig. 3; Şimşek, 1984, 2003b; Şimşek et al., 2005): The Neogene formations (i.e. the Denizli Group) are locally good aquifer for the thermal waters representing the shallow (first) reservoir when intense fracturing developed. The Denizli Group is consisted of four formations, from bottom to top: The Kızılburun Formation is located below the Sazak Formation (Fig. 3). This formation with a wedge-shaped geometry (up to 450 m thick), overlies unconformably the bedrock of the Menderes Massif. The thickness of this formation increases towards the basinmargin faults (Pamukkale-Yenice Fault in the northern margin and Babadağ Fault in the southern margin, Fig. 2). The lateral extent decreases toward the basin center (several tens of kilometers). It is composed of two subunits from bottom to top: proximal-medial alluvial-fan (conglomerate, sandstone, siltstone and mudstones alternations) and distal alluvial-fan (sandstone, siltstone, mudstones, coal and clayey limestone alternations). Its mineral association is calcite + dolomite + illite + smectite. The Sazak Formation (up to 300 m thick) is subdivided into three subunits that are from the bottom to top (Fig. 3): lake-margin unit (limestone, marl, mudstone and clayey limestone alternations; calcite + dolomite + gypsum + celestite + barite + strontianite + quartz + illite + smectite), shallow lake unit (cherty limestones and dolostones; calcite + dolomite + strontianite + celestite + barite + quartz + illite + smectite + tridymite), and playa/saline lake unit (gypsarenite, selenite, gypsiferous mudstone and bituminous shale alternations; calcite + dolomite + gypsum + halite + quartz + cristobalite + strontianite + illite + smectite + heulandite + tridymite. The thickness of this formation increases toward the basin centre. In the study area, this unit extends laterally several tens of kilometers (Fig. 2) and is up to 200 m thick (DG-1 well; Fig. 4A). The Kolankaya Formation (up to 500 m thick) rests conformably on the Sazak Formation and overlies unconformably the metamorphic bedrock in the northern part of the basin (Figs. 2 and 3). It extends laterally over tens of kilometers and thickens towards the northern and southern basin margins. This formation consists of four subunits, that are, from bottom to top: shallow lake unit (mudstone-siltstone and marl alternations), sublittoral to profundal lake unit (alternating marl-claystone, sandstone, and clayey limestone alternations), littoral unit (conglomerate, sandstone and siltstone alternations) and alluvial fan deposits (conglomerate, sandstone, siltstone and mudstone alternations). Mineral association of the shallow lake and sublittoral/profundal lake unit is calcite + dolomite + strontianite + quartz + illite + gypsum + smectite + illite. The Tosunlar Formation (up to 150 m thick) unconformably overlies the older formations (Fig. 3). It can extend for a few kilometers toward the basin centre. This formation includes two units that are, from the bottom: proximal and medial alluvial-fan unit (conglomerate, sandstone, siltstone and mudstone alternations) and fluvial unit (sandstone, siltstone, mudstones and marl alternations). The deep (second) reservoir unit is hosted within the Paleozoic Ortakoy and İğdecik formations, which is mostly composed of quartzite, marble and schist alternations (Figs. 3 and 4A–G). Mineral assemblage is

biotite + muscovite + plagioclase + quartz + garnet + chloritoid, for the late Devonian-early Carboniferous Ortakoy Formation and chloritoid + muscovite + quartz + graphite + rutile + ilmenite + garnet, for the Permo-Carboniferous Göktepe (Igdecik) Formation (Okay, 1989; 2001). Similar to the first reservoir, this unit is widely recognized in the subsurface (tens of kilometers) and was drilled down to 850 m in the study area (DG-1 well; Fig. 4A). This formation has a relatively high secondary permeability when affected by faults and fractures, permitting circulation of the thermal waters. Thus, this unit locally plays the role of a good reservoir. 4. Sampling and methods Ten thermal and seven cold water samples were collected during 2014 and 2015 for chemical analyses. Further 18 thermal and 10 cold waters datasets were compiled from previous studies (Tables 3 and 4). Locality data and hydrochemical properties are listed in Tables 3 and 4. Samples were stored in two polyethylene bottles. One of the bottles was acidified with suprapure HNO3 for determination of cations and SiO2 analyses. The other one was kept unacidified for anion analyses. The charge balance error for the reported analyses (calculated with PHREEQC) is b 5% for all the samples. EC (electrical conductivity), pH and temperature were measured in the field, while alkalinity as HCO3 was defined by titration with HNO3 (0.1 M) when pH value reaches 4.2 (for HCO3) on the day of sampling. Major ion and trace element contents (10 thermal and 7 cold waters) were determined at the Bureau Veritas Laboratories (Ankara, Turkey) using Inductively Coupled Plasma Mass Spectrometry (ICP-MS). Oxygen and deuterium isotopic composition of water samples (10 thermal and 7 cold) was analyzed with a Finnigan Mat 252 Isotope Ratio Mass Spectrometer at SIRFER Laboratory (Utah, USA). Isotopic composition of dissolved inorganic carbon (DIC) from water samples (10 thermal and 7 cold) was measured by reaction with pure phosphoric acid to bring all DIC to gaseous CO2 during 48 h at 25 °C constant temperature, and then measured in an Europa Scientific 20-20 Isotope Ratio Mass Spectrometer at Iso-Analytical Laboratory (Cheshire, UK). The precision of the analyses is ±0.15‰ for δ18O and ±1‰ for δ2H (VSMOW) and ±0.05‰ for δ13C (VPDB). Tritium isotope analyses of 7 thermal and 2 cold waters were conducted at the Water Chemistry Laboratory of the Hacettepe University (Ankara, Turkey). 5. Hydrogeochemistry The hydrochemical properties of the thermal and cold waters from the GGF are described on the basis of their physicochemical and isotope data presented in Tables 3 and 4. 5.1. Physicochemical characteristics of the waters All thermal water types contain different pH values, discharge temperatures, electrical conductivities (EC) and major ion chemistry parameters. As seen in Fig. 5, the Gölemezli thermal waters are plotted in the magmatic steam and steam-heated waters field. 5.1.1. Thermal waters (Group 1) According to their major ionic abundances and position in the ClSO4-HCO3 (Fig. 5) and Piper (Fig. 6) ternary diagrams, the GGF thermal waters plot in different fields and can be divided into the following two chemical groups (Table 3): 5.1.1.1. Group 1A thermal waters. This group corresponds to SO4-rich alkaline-sulfate waters and consists of two chemical types: Na-SO4 type – These waters include SO4 (mean 1528 mg/l) and HCO3 (mean 1260 mg/l) as dominant anions, Cl (mean 77 mg/l) as minor anion; Na (mean 554 mg/l) and Ca (mean 378 mg/l) as dominant cations and Mg (mean 130 mg/l) and K (mean 60 mg/l) as minor cations


Fig. 3. Stratigraphy of the Denizli Basin (Şimşek, 1984; Okay, 1989; Sun, 1990; Saraç, 2003; Kaymakçı, 2006; Alçiçek et al., 2007; Gündoğan et al, 2008; Alçiçek et al., 2015).

7

8


Fig. 4. Schematic stratigraphic borehole logs within the GGF (Akkuş et al., 2005; Akıllı and Bülbül, 2006). The main lithological units are indicated. In the DG3 and DG2 wells, a temperature between 549 m (66 °C) and 596.8 m (73 °C) respectively was measured. Differently, temperatures in the DG4 and DG5 wells (70 °C and 62 °C, respectively) are reversed at 750 m in depth, suggesting movement of cooler water to depth. Contrastingly, temperature of DG1 well (88 °C) increases to depth, at 1500 m.

(Table 3). They show relatively high electrical conductivity (mean 3910 μS/cm), slightly acidic pH (mean 6.4) and discharge temperatures of mean 59 °C (49.8 to 88 °C) (Table 3).

Ca-SO4 type – Dominant anion concentrations of these waters (SO4: mean 1817 mg/l and HCO3: mean 1465 mg/l) are higher than Na-SO4 type, whereas minor anion contents (Cl: mean 77 mg/l) are similar to

Table 3 Chemical and isotope composition of sampled spring and drillhole waters from GGF (n.a.: not analyzed). Group Sample Name no

Ca-SO4 type G9 G10 G11 G12 G13 G14 G15 G16 G17 G18 G19 Mean

Gölemezli spring (Akkuş et al., 2005) Gölemezli spring (Tamgaç et al., 1995) Gölemezli Şanlıalp (Şengün, 2011) Gölemezli hamamı (Kıymaz, 2012) Kocabaylar 1 (KB) (this study) Kocabaylar 2 (KB) (this study) Kocabaylar 3 (KB) (this study) Boss (this study) N. Erdemir (NE) (this study) Göl 2.1 (this study) Gölemezli mud spring (this study)

Group 1B Thermal waters Ca-HCO3 type G20 Gölemezli MTA (Yaman, 2005) G21 Gölemezli DG 3 (Demirel and Kahraman, 2003) G22 Gölemezli DG 4 (Demirel and Kahraman, 2003) G23 Gölemezli DG 5 (Demirel and Kahraman, 2003) G24 Gölemezli Göl 1 (Şengün, 2011) G25 Gölemezli well 1 (Kıymaz, 2012) G26 Gölemezli well 2 (Kıymaz, 2012) G27 Gölemezli Göl 1 (this study) G28 Gölemezli DG3 (this study) Mean

T pH (°C)

SO4 K Ca Na Mg SiO2 EC δ18O δ2H δ13C Tritium Water HCO3 Cl (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (μS/cm) (‰ VSMOW) (‰ VSMOW) (‰ VPDB) (TU) composition

51.0 49.8 50.0 51.0 88.0

1970 1988 1992 1998 2001

Thermal spring Thermal spring Thermal spring Thermal drill Thermal drill

6.30 6.30 5.90 6.11 6.60

1208 1310 1330 1030 1506

78 76 79 75 73

1649 1590 1662 1665 1504

51 58 60 61 55

418 414 379 391 367

535 539 560 543 750

129 134 121 135 72

48 51 n.a. n.a. 48

n.a. n.a. 4000 4330 4227

n.a. n.a. n.a. n.a. n.a.




Na-Ca-SO4-HCO3 Na-Ca-SO4-HCO3 Na-Ca-SO4-HCO3 Na-Ca-Mg-SO4-HCO3 Na-Ca-SO4-HCO3

2002 2002 2015

Thermal spring 58.4 6.29 1549 Thermal drill 73.0 7.60 544 Thermal drill 52.8 6.08 1605 59.3 6.40 1260

86 55 90 77

1740 651 1767 1528

80 51 68 60

426 106 522 378

646 245 612 554

178 111 158 130

43 73 52 52

4340 2030 4530 3910

−8.32 n.a. −8.37 −8.35

−60.40 n.a. −61.09 −60.75

n.a. n.a. 5.68 5.68

1 n.a. 0.9 0.95

Na-Ca-Mg-SO4-HCO3 Na-Mg-Ca-SO4-HCO3 Na-Ca-SO4-HCO3

1976 1994 2011 2012 2014 2015 2015 2015 2015 2015 2015

Thermal spring Thermal spring Thermal drill Thermal drill Thermal drill Thermal drill Thermal drill Thermal drill Thermal drill Thermal drill Thermal spring

52.0 49.6 48.8 59.0 44.5 50.0 47.8 58.2 45.0 65.0 41.9 51.1

1030 1537 1263 1251 1434 1583 1576 1785 1467 1548 1636 1465

65 86 66 71 71 80 81 81 87 79 78 77

1820 1635 1836 1664 1806 1887 1842 1821 1977 1926 1773 1817

46 60 51 45 66 70 68 70 63 71 64 61

530 511 423 464 514 559 547 579 600 559 535 529

400 570 558 432 546 589 584 613 561 602 557 547

136 161 138 110 167 170 167 169 162 166 183 157

120 121 115 126 115 126 134 141 119 132 121 125

n.a. 3590 4080 4460 4210 3760 4290 4610 4570 4730 4210 4251

n.a. n.a. n.a. −8.32 −8.30 −8.13 −8.20 −8.32 −8.13 −8.26 −8.28 −8.24

n.a. n.a. n.a. −60.40 −59.97 −59.88 −59.68 −60.85 −59.34 −60.46 −60.54 −60.14

n.a. n.a. n.a. n.a. 6.63 7.00 7.54 5.39 6.83 5.11 6.04 6.36

n.a. n.a. n.a. 1 n.a. n.a. 0.08 0.05 n.a. 0.21 n.a. 0.34

Ca-Na-SO4-HCO3 Ca-Na-Mg-SO4-HCO3 Ca-Na-SO4-HCO3 Ca-Na-SO4-HCO3 Ca-Na-Mg-SO4-HCO3 Ca-Na-Mg-SO4-HCO3 Ca-Na-Mg-SO4-HCO3 Ca-Na-SO4-HCO3 Ca-Na-SO4-HCO3 Ca-Na-SO4-HCO3 Ca-Na-Mg-SO4-HCO3

2002 2002

Thermal drill Thermal drill

70.7 6.46 2340 66.0 7.10 1361

40 31

610 485

100 73

436 219

494 240

142 93

n.a. 49

3760 2220

−8.32 n.a.

−57.80 n.a.

n.a. n.a.

0.50 n.a.

Ca-Na-Mg-HCO3-SO4 Ca-Na-HCO3-SO4

2003

Thermal drill

70.0 6.70 1243

28

471

62

222

243

77

51

2695

n.a.

n.a.

n.a.

n.a.

Ca-Na-HCO3-SO4

2003

Thermal drill

62.0 6.60 1653

37

635

80

377

315

153

54

3120

n.a.

n.a.

n.a.

n.a.

Ca-Na-Mg-HCO3-SO4

2011 2012 2012 2015 2015

Thermal drill Thermal drill Thermal drill Thermal drill Thermal drill

57.0 67.0 69.0 58.5 53.0 63.7

35 31 27 42 29 33

475 378 432 552 432 497

51 52 43 76 56 66

529 449 456 591 392 408

370 248 208 346 224 299

107 72 84 124 55 101

53 64 52 56 47 53

3520 2420 2470 3760 2920 2987

n.a. n.a. n.a. −8.40 −8.33 −8.35

n.a. n.a. n.a. −57.80 −57.41 −57.67

n.a. n.a. n.a. 6.18 6.98 6.58

n.a. n.a. n.a. 0.23 0.40 0.38

Ca-Na-HCO3 Ca-Na-HCO3 Ca-Na-HCO3-SO4 Ca-Na-HCO3-SO4 Ca-Na-HCO3-SO4

6.30 6.60 6.60 6.28 6.66 6.42 5.60 6.18 6.01 6.34 5.90 6.26

6.95 6.89 6.69 6.69 6.70 6.75

2460 1848 1765 2564 1470 1856


Group 1A Thermal waters Na-SO4 type G1 Gölemezli (Gökalp, 1971) G2 Gölemezli spring (Guidi et al., 1990) G3 Gölemezli spring (Filiz et al., 1992) G4 Gölemezli (Özgür, 2002) G5 Gölemezli DG1 (Yıldırım and Güner, 2002) G6 Gölemezli thermal spa (Yaman, 2005) G7 Gölemezli DG 2 (Akkuş et al., 2005) G8 H. Demirci (HD) (this study) Mean

Sampling date

9

Chloride is termed a “relatively conservative element” and serves as a very good tracer element in fluids from geothermal systems by comparing its concentration to the one from other ions in solution (Michard, 1990; Motyka et al., 1993). The major ion contents and EC values of all GGF of thermal and cold waters were plotted against Cl concentrations (Fig. 7A–H). Accordingly, Na-SO 4 type and CaSO4 type display similar trends (Group 1A), whereas Ca-HCO3 type (Group 1B) shows different trends from Group 1A. Therefore, the Na, K, Ca, Mg, HCO3 and SO4 contents from Group 1A and 1B display two different linear correlations versus Cl, with distinct slopes.

Mg-Ca-HCO3 Mg-Ca-HCO3-SO4 Mg-Ca-HCO3-SO4 0.39 n.a. n.a. 0.39 −9.87 −4.09 −9.74 −7.90 −50.61 −50.01 −52.95 −51.19 −7.68 −7.24 −7.86 −7.59 530 920 1180 877 62 22 18 34 43 52 119 71 11 49 47 35 52 85 87 75 2 4 4 3 15 213 255 161 14 42 30 29

n.a. 17.4 n.a. n.a. n.a. n.a. 12.8 12.3 14.8 10.6 n.a. n.a. 2.73 n.a. 11.77 n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. −10.38 −8.82 −9.71 −8.16 −9.27 n.a. −56.00 n.a. n.a. n.a. −55.10 −51.70 −52.10 −55.50 −52.70 −54.70 −53.57 −54.19 −54.13 −53.97 n.a. −8.45 n.a. n.a. n.a. −8.80 −8.80 −8.70 −8.10 −8.20 −8.44 −8.21 −8.51 −7.90 −8.41 329 260 260 378 370 800 343 393 290 231 390 420 380 460 379 26 4 5 4 5 5 4 16 4 5 17 20 24 19 11 7 2 4 3 4 16 3 6 8 7 4 7 8 5 6 4 11 2 9 6 16 4 4 8 13 6 4 3 6 7 85 70 53 85 80 128 116 71 67 36 62 87 76 101 80 1 1 1 2 3 2 1 1 2 1 1 1 0 2 1 37 5 5 9 10 14 12 3 8 21 12 9 12 15 12 9 11 7 15 13 18 8 7 17 11 4 2 1 9 9 207 235 167 240 249 427 350 251 222 134 203 297 264 316 254 7.80 7.20 7.50 6.79 7.41 7.40 7.25 7.30 7.80 6.50 7.62 7.70 7.68 8.03 7.43 14.8 12.0 12.5 12.0 13.0 14.0 11.2 15.0 10.0 15.3 11.5 10.7 13.4 8.7 12.4 Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring Cold spring

Cold spring 13.9 7.48 358 Cold spring 16.6 7.77 704 Cold spring 14.2 7.40 612 14.9 7.55 558 2015 2015 2015 Group 2B Cold waters Mg-HCO3 type C15 Eymir (this study) C16 Yeniköy-Kocapınar (this study) C17 Irlıganlı-Çobanlı (this study) Mean

5.2. Circulation pathways and mixing processes

1994 1995 1995 1995 1995 2000 2000 2000 2000 2000 2015 2015 2015 2015

5.1.2.2. Group 2B cold waters. This group belongs to the Mg-HCO3 type. Dominant anion concentrations (HCO3: mean 558 mg/l and SO4: mean 161 mg/l) and minor anions Cl (mean 29 mg/l) are higher than the Ca-HCO3 type. Similarly, minor cations Mg (mean 71 mg/l) and K (mean 3.4 mg/l) are also higher than the Ca-HCO3 type (Table 4). Differently, dominant cation Ca contents (mean 75 mg/l) are lower than those detected in the Ca-HCO3 type. These waters have the highest electrical conductivity (mean 877 μS/cm), neutral pH (mean 7.6) and high discharge temperatures (13.9 to 16.6 °C; mean 15 °C) (Table 4), indicating a longer circulation and residence times if compared to the results from the Ca-HCO3 type.

Group 2A Cold waters Ca-HCO3 type C1 Uzunpınar (Tamgaç et al., 1995) C2 Kurtluca CK (Dilsiz, 2006) C3 Kurtluca CK (Dilsiz, 2006) C4 Uzunpınar CU (Dilsiz, 2006) C5 Uzunpınar CU (Dilsiz, 2006) C6 Güzelpınar (Yıldırım and Güner, 2002) C7 Sepetpınarı (Yıldırım and Güner, 2002) C8 Uzunpınar (Yıldırım and Güner, 2002) C9 Ören (Yıldırım and Güner, 2002) C10 Gölyeri (Yıldırım and Güner, 2002) C11 Uzunpınar (this study) C12 Akçapınar sondaj (this study) C13 Akçapınar kaynak (this study) C14 Güzelpınar (this study) Mean

5.1.2.1. Group 2A cold waters. This group is Ca-HCO3 type with predominant anions as HCO3 (mean 254 mg/l) and SO4 (mean 12 mg/l) and minor anion as Cl (mean 9 mg/l). The predominant cations are Ca (mean 80 mg/l) and Na (mean 7 mg/l) and minor cations are Mg (mean 6 mg/l) and K (mean 1.4 mg/l) (Table 4). This group have the lowest electrical conductivity (mean 379 μS/cm), neutral pH (mean 7.4) and low discharge temperatures (8.7 to 15.3 °C; mean 12 °C) (Table 4).

Table 4 Chemical and isotope composition of sampled cold waters from the GGF (n.a.: not analyzed).

5.1.2. Cold waters (Group 2) Cold waters are composed of two chemical group types, referred to as Group 2A (Ca-HCO3) and Group 2B (Mg-HCO3) (Table 4). These waters plot in the field of peripheral and shallow waters (Fig. 5). Anion and cation contents in the cold waters are lower than in the thermal waters. Furthermore, EC and temperature values of the cold waters are much lower than those recognized in the other types of thermal waters. Comparatively, pH values are always higher (Table 4). This accounts for a shorter circulation and residence time period, if compared to all other thermal waters.

Sampling Type date

5.1.1.2. Group 1B thermal waters. This group is only composed of CaHCO 3 water type. These waters have the highest HCO 3 contents (mean 1856 mg/l) and K (mean 66 mg/l), and the lowest SO 4 (mean 497 mg/l), Cl (mean 33 mg/l), Na (mean 299 mg/l) and Mg (mean 101 mg/l). The dominant cation Ca content (mean 408 mg/l) has moderate values (Table 3). This type water has the lowest electrical conductivity (mean 2987 μS/cm), the highest discharge temperatures (53 to 70.7 °C; mean 64 °C) and a neutral pH (mean 6.8) (Table 3). This reflects that Ca-HCO 3 type thermal waters have shorter circulation and residence times, if compared to Na-SO4 and Ca-SO4 types.

T pH (°C)

δ2H δ13C Tritium Water HCO3 Cl SO4 K Ca Na Mg SiO2 EC δ18O composition (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (mg/l) (μS/cm) (‰ VSMOW) (‰ VSMOW) (‰ VPDB) (TU)

those of Na-SO4 type. Similarly, dominant cations Ca (mean 529 mg/l) and minor cations Mg (mean 157 mg/l) are also higher than Na-SO4 type (Table 3). Differently, dominant cation Na contents (mean 547 mg/l) are lower than and minor cation K (mean 61 mg/l) are slightly higher than Na-SO4 type. They have relatively high electrical conductivity (mean 4251 μS/cm), a slightly acidic pH (mean 6.3) and lower discharge temperatures (41.9 to 65 °C; mean 51 °C) (Table 3), thus suggesting a longer circulation and residence time period, if compared with the indications from Na-SO4 waters.

Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3 Ca-HCO3


Group Sample no Name

10


11

Fig. 5. Cl-SO4-HCO3 triangular diagram of the GGF thermal and cold waters (based on Nicholson, 1993).

Consequently, the two trends of Fig. 7A–H, may be explained considering different circulation pathways for Group 1A and 1B thermal waters.

The good positive correlations for both SO4 type and HCO3 type thermal waters in the EC vs. Cl (Fig. 7A), Na vs. Cl (Fig. 7B) and K vs. Cl (Fig. 7C) suggest mixing of thermal and cold waters. The correlations

Fig. 6. Piper diagram for the studied water samples from the GGF.

12

H. Alçiçek et al. / Journal of Volcanology and Geothermal Research 349 (2018) 1–30 5000

800

4500

700

R² = 0.69

R² = 0.94

3000 2500 2000 1500

500 400 R² = 0.59

300 200

1000

100

500 0

0 0

20

40

60

80

100

Cl (mg/l)

A

0

20

40

60

80

100

Cl (mg/l)

B

120

700

100

600 500

80

R² = 0.92

Ca (mg/l)

K (mg/l)

R² = 0.91

600

3500

Na (mg/l)

EC (µS/cm)

4000

60 R² = 0.55 40

400 R² = 0.87 R² = 0.56

300 200

20

100 0

0 0

20

40

60

80

Cl (mg/l)

C

0

100

20

40

60

200

80

100

80

100

Cl (mg/l)

D 3000

180 2500

Mg (mg/l)

140

R² = 0.89

120 100

R² = 0.72

80

HCO3 (mg/l)

160

60 40

2000

R² = 0.68

R² = 0.92

1500 1000 500

20 0

0 0

20

40

60

80

100

Cl (mg/l)

E

0

20

40

60

Cl (mg/l)

F 10

2500

9 R² = 0.91

2000

R² = 0.91

8 B (mg/l)

SO4 (mg/l)

7 1500 1000

6 5 4

R² = 0.52

3

500

2

R² = 0.70

1 0

0 0

20

40

60

80

Cl (mg/l)

G

0

100

2500

40

60

80

100

Cl (mg/l)

18 16

R² = 0.93

2000

14

R² = 0.97

12

Sr (mg/l)

SO4 (mg/l)

20

H

1500 1000 R² = 0.75

10 8

R² = 0.75

6 4

500

H2S oxidation

2 0

0 0

200

400

600

800

1000

1200

1400

1600

Ca (mg/l)

I

0

500

1000

1500

2000

2500

SO4 (mg/l)

J

Group 1A:

Thermal waters: Na-SO4

Thermal waters: Ca-SO4

Group 1B:

Group 2A:

Cold waters: Ca-HCO3

Group 2B:

Cold waters: Mg-HCO3

Thermal waters: Ca-HCO3

Fig. 7. Relations among various ions versus chloride for Gölemezli thermal and cold waters from the GGF (A) EC vs. Cl (B) Na vs. Cl, (C) K vs. Cl, (D) Ca vs. Cl; (E) Mg vs. Cl; (F) HCO3 vs. Cl; (G) SO4 vs. Cl, (G) B vs. Cl, (I) SO4 vs. Ca, (J) SO4 vs. Sr.


of Na vs. Cl and K vs. Cl also indicate progressive reaction with feldspars in the deep reservoir, implying an independent indicator of residence times (Han et al., 2010). Good Ca vs. Cl (Fig. 7D), Mg vs. Cl (Fig. 7E), HCO3 vs. Cl (Fig. 7F) and SO4 vs. Cl (Fig. 7G) correlations in all thermal waters also reflect dilution (mixing) of thermal waters by shallow (colder) meteoric waters and/or a lesser degree of water-rock interaction at lower temperature (Vengosh et al., 2002; Han et al., 2010). High Na/Cl ratios (Table 5) in the Group 1A and 1B thermal waters may be controlled by water-rock interactions, and in specific dissolution of Na-feldspar in the deep reservoir metamorphic rocks. The high concentrations of Na and Ca are indicative of a significant interaction with Na- and Ca-rich minerals, possibly occurring in the reservoir and/or in the recharge area. High Na contents in the Na-SO4 type waters of Group 1A thermal waters indicate release by alteration of Na-rich feldspars, reasonably from the schists that are part of the deeper reservoir. Ca/Na ratios of the thermal waters are low, whereas Ca/Mg ratios are high (Table 5) which reflects ion exchange processes. Ca-SO4 type waters of Group 1B are produced by ion exchange reactions between Na and Ca on clayey horizons in the schists and Neogene sedimentary rocks. As presented in Table 5, the Na/Cl and K/Cl ratios of the Group 1A and 1B thermal waters are higher than those of Group 2A and 2B cold waters which explained as a consequence of its longer and deeper flow path (Han et al., 2010). The Ca/Cl ratios of thermal waters Group 1A are lower than the ones belonging to Group 1B (Table 5). Moreover, the Ca/Cl ratios of cold waters Group 2A (Ca-HCO3 type) are significantly higher than values from Group 1A, 1B and 2B (Table 5). Low Ca/Cl ratios in the thermal waters imply calcite precipitation along the flow path (Han et al., 2010). High Ca/Cl ratios in the cold waters indicate calcite or dolomite dissolution along with weathering of Ca silica minerals such as plagioclase in the metamorphic rocks (Han et al., 2010). K/Cl and Mg/Cl ratios of the thermal water from all groups are low, but still slightly higher than those of cold waters (Table 5). This indicates that K- and Mg-bearing minerals are scarce in the metamorphic bedrocks. Low K/Cl and Mg/Cl ratios are also probably due to the formation of K-bearing clays (illite) and Mg-bearing clays (smectite) (respectively) in the Neogene formations (Alçiçek, 2007). Thermal waters of Group 1B are characterized by Ca/Mg ratios higher than Group 1A, but the Ca/Mg ratios from Ca-HCO3 type of Group 2A are also higher than Group 1A and 1B (Table 5). This situation is attributed to related with Mg depletion (Müller, 1967), leading to the occurrence of clay minerals (illite and smectite) in the Neogene formations (Alçiçek, 2007). HCO3/Cl ratios in the Group 1B thermal waters are much higher than those from the Group 1A thermal and Group 2A and 2B cold waters (Table 5). This suggests a shorter flow path and a faster water cycle (Han et al., 2010; Majumdar et al., 2009). Moreover, high HCO3/Cl ratios (HCO3/Cl N 1) are found in all water groups, implying a mixing process between ascending thermal and descending cold waters (Giggenbach, 1988). The pH values below 7 (Table 3) of the Group 1 reflect carbonate dissolution. Dissolution of carbonate from the Neogene formations and from the marbles of the Paleozoic formations allows for an increasing of the Ca and HCO3 contents. High HCO3 concentrations in the Group 1B thermal waters indicate the reaction of CO2-rich waters with carbonate and marble. Group 1 thermal waters have the highest SO4/Cl ratios, whereas Group 2 cold waters have the lowest SO4/Cl ratios (Table 5). SO4/Cl ratios in the Group 1A thermal waters are higher than those of the Group 1B thermal and Group 2A and 2B cold waters. SO4 is the major anion in the Group 1A thermal waters. In the Group 1A thermal waters, Ca/SO4 molar ratios are above stoechiometric ratio (Fig. 7I), suggesting that the SO4 composition in the Group 1A was not only derived from evaporite deposits (leaching of sulfate minerals), but

13

also contributed by sulfur from sulfide deposits and/or (e.g., pyrite) hydrothermal gases (especially H2S and SO2) indicating a deep sulfate source (Levet et al., 2002; Cinti et al., 2011). These thermal waters have strong H2S smell at emergence, reflecting that a part of sulfate was produced by H2S oxidation during the thermal water rising to the wellhead. The Group 1B thermal waters exhibit a good correlation in Ca vs. SO4 with an equimolar Ca/SO4 ratio due to the dissolution of sulfate minerals (Fig. 7I). The dissolution of sulfates in the thermal waters is also supported by the positive correlation between SO 4 and Sr (Fig. 7J). Besides, the low SO4/Cl ratios in the cold waters reflect a loss of sulfur through precipitation of sulfides (pyrite), sulfates (anhydrite) or elemental sulfur or leaching of sulfate-poor rocks (Levet et al., 2002). To determine hydrochemical facies of the thermal and cold waters, base exchange indices (bei) are commonly used (Schoeller, 1934). The GGF thermal waters have very negative bei values, whereas cold waters display slightly negative to positive values (Table 6). Negative values are indicative for waters originating from metamorphic and sedimentary rocks with alkaline ions released by alteration of silicate minerals (Şahinci, 1991). This evidence supports that the GGF thermal waters circulated within the Paleozoic metamorphic rocks and the lacustrine deposits of the Neogene formations. 5.3. Trace and rare element concentrations Trace element and rare earth element (REE) compositions in geothermal waters are used to understand tracing the origin of fluids, the state of equilibrium in water-rock interaction, and changes of fluid composition (Möller, 2000; Göb et al., 2013). Hydrochemical features of geothermal fluids are very different from those derived from non-thermal fluids as a consequence of release/continuing mobility of some conservative elements such as Si, Fe, Li, B, Sr, Ba, F, Mn, Zn, Al and As (Stüben et al., 2003; Das et al., 2005), favoring temperature. Furthermore, high temperatures values correspond to increased minor element contents (Tarcan and Gemici, 2003). Mean Li, B, Sr, Si, Ba, F, Mn, Zn and Al (excluding Fe and As) contents of Group 1A thermal waters (Na-SO4 and Ca-SO4 types) are higher than those of Group 1B (Ca-HCO3 type). Comparatively, the trace element contents of Group 2 cold waters (Ca-HCO3 and Mg-HCO 3 type) are lower than those of Group 1 thermal waters (Table 7). This likely reflect a different degree of water-rock interaction. The abundance of minor elements in thermal waters, if compared to cold waters, indicate thermal waters with a higher reactivity, permitting the increase of leaching of the minor elements from the host rock (Ma et al., 2011). In addition, the contrast in trace element contents is likely determined by shorter residence time of colder waters and their continuous dilution with meteoric or river waters (Ma et al., 2011). The Sr contents of the Group 1A (mean 14.7 mg/l) thermal waters are lower than those of Group 1B (mean 12.3 mg/l). Similarly, the Sr contents of the Group 2A (0.17 mg/l) and 2B (1.60 mg/l) cold waters are the lowest (Table 7). The high Sr contents of the Group 1A and 1B thermal waters reflect exchange between rising fluids and Menderes metamorphics (103–1515 ppm, mean of 322 ppm; Koralay and Kılınçarslan, 2015) and Neogene formations with high Sr contents (104–15995 ppm, mean of 2066 ppm; Alçiçek, 2007). Similarly, high Sr values have been also reported in the late Miocene-Pliocene (Paton, 1992) Denizli volcanics, in the southeast of the Denizli Basin by Prelević et al. (2015) (2231–3220 ppm, mean of 2601 ppm) and Semiz et al. (2012) (2231–3505 ppm, mean of 2878 ppm) and the Quaternary travertines at Gölemezli, Karahayıt and Pamukkale geothermal fields (484–1756 ppm, mean of 904 ppm; Özkul et al., 2013). The Si contents of the Na-SO4 type thermal waters of Group 1A (mean 18 mg/l) are lower than those of Ca-HCO3 type of Group 1B (mean 25 mg/l). Besides, Ca-SO4 type thermal waters have the

14


Table 5 Average ionic ratios of the GGF thermal and cold water samples. Group

Water samples

Group 1 Group 1A

Thermal waters Na-SO4 type Ca-SO4 type Ca-HCO3 type Cold waters Ca-HCO3 type Mg-HCO3 type

Group 1B Group 2 Group 2A Group 2B

Na/Cl

Ca/Na

Ca/Mg

Ca/Cl

K/Cl

Mg/Cl

HCO3/Cl

SO4/Cl

6.75 6.68 8.13

0.78 0.77 0.68

2.78 3.31 4.5

4.86 6.68 12.48

0.79 0.76 1.87

1.8 2.0 2.8

16.0 18.4 52.2

19.72 22.75 13.96

1.34 1.17

0.66 0.43

15.9 1.20

15.09 2.88

0.22 0.12

1.3 2.8

49.72 20.91

3.60 4.88

highest Si contents (mean 58 mg/l) (Table 7). This probably indicates flow along different lithologies with different amount of Si-bearing minerals in the metamorphic rocks and Neogene formations (Han et al., 2010). Cl and B concentrations are higher in all thermal water groups (Group 1A and 1B) with respect to the content in the cold waters (Group 2A and 2B) (Tables 4 and 7). High B and Cl contents of the thermal waters indicate relatively deep circulation paths and interaction with leached host rocks. Boron in thermal waters probably was provided by interaction with the various geological formations and degassing of magma intrusions (Gemici and Tarcan, 2002). B and Cl are generally considered as conservative elements, which are used for following the groundwater flow paths and mixing processes of distinct waters (Motyka et al., 1993). There are high correlations of B and Cl concentrations in the thermal waters (r-values N 0.8 for Na-SO4 and Ca-SO4 types of the Group 1A and r-values N 0.7 for Ca-HCO3 type of the Group 1B; Fig. 7H). This confirms that the thermal waters have interaction with leached host rocks and, moreover, it accounts for the presence of mixing processes between thermal and cold waters in the GGF. The Group 1A and 1B thermal waters have similar high F contents (2.63 mg/l and 2.55 mg/l, respectively; Table 7), indicating that the uprising thermal waters interacted with F-bearing minerals belonging to the metamorphic rocks. High Li concentrations in the Group 1A and 1B thermal waters (1.58 mg/l and 1.63 mg/l, respectively; Table 7) are formed by exchange with the clays (such as smectite) from the Neogene formations (Alçiçek, 2007). Möller et al. (2004) reported that the GGF thermal waters contain high REE contents which are predominantly controlled by schists and marbles of the Menderes Massif. However, the Neogene formations are also rich in REE (Table 8, Alçiçek, 2007). As seen in Table 8 and Fig. 8, all data show similar REE trends and therefore, it may be considered that the GGF thermal waters are affected by both Neogene formations and the Menderes metamorphic rocks.

5.4. δ18O and δ2H isotopes The δ18O and δ2H isotopic compositions of thermal and cold waters are very useful tools for tracing the origin of the groundwater system and identifying the recharge elevation of meteoric waters (Craig, 1961; Giggenbach et al., 1983). They are also representative of the degree of water-rock interaction, and of boiling and mixing processes (Clark and Fritz, 1997). The δ18O-δ2H thermal and cold water compositions from the GGF is compared to the Local Meteoric Water Line (LMWL; δ2H = 8δ18O + 16; Şimşek, 2003a), the Global Meteoric Water Line (GMWL; δ2H = 8δ18O + 10; Craig, 1961) and to the Eastern Mediterranean Meteoric Water Line (EMMWL; δ2H = 8δ18O + 22; Gat and Carmi, 1970) (Fig. 9A). All GGF thermal and cold waters are commonly situated between EMMWL and GMWL lines (Fig. 9A), providing strong evidence for a predominant origin from meteoric water. All thermal waters fall to the right side of the GMWL line (Fig. 9A), indicating a slight evaporation effects in a semi-arid climate. This reflects the interaction of ions in solution and water molecules in a fluid system at high salinity, even at low temperature (Gonfiantini, 1986). The δ18O and δ2H values of Na-SO4 thermal waters (mean of ‐ 8.35‰ and − 60.75‰, respectively) are slightly lower than those of Ca-SO4 thermal water (mean of −8.24‰ and −60.14‰, respectively); values of Ca-HCO3 waters (mean of −8.35‰ and −57.67‰, respectively) are slightly higher (Table 3). Besides, the cold waters are located between EMMWL and GMWL lines (Fig. 9A), indicating a meteoric origin. The δ18O values of the cold waters also show slightly positive shift as a result of water-rock isotope exchange at relatively high temperature. The δ18O and δ2H values of CaHCO3 cold waters (mean of −8.41‰ and −53.97‰, respectively) are slightly lower than those of Mg-HCO3 cold waters (mean of − 7.59‰ and −51.19‰, respectively) (Table 4). The δ18O and δ2H isotopes have also important implications due to their relationship with the altitude (e.g., Craig, 1961). The cold waters

Table 6 Base exchange indices (Bei) of the GGF samples (Bei = (Cl − (Na + K) / Cl); Schoeller, 1934). Thermal waters

Cold waters

Na-SO4 type

Bei

Ca-SO4 type

Bei

Ca-HCO3 type

Bei

Ca-HCO3 type

Bei

Mg-HCO3 type

Bei

G1 G2 G3 G4 G5 G6 G7 G8

-6.52 -6.81 -6.83 -6.88 -9.33 -8.29 -3.35 -7.00

G9 G10 G11 G12 G13 G14 G15 G16 G17 G18 G19

-5.86 -6.32 -8.29 -4.89 -8.03 -7.82 -8.01 -7.52 -6.65 -7.33 -6.21

G20 G21 G22 G23 G24 G25 G26 G27 G28

-13.89 -9.09 -7.90 -11.15 -14.45 -7.93 -6.67 -11.34 -8.72

C1 C2 C3 C4 C5 C6 C7 C8 C9 C10 C11 C12 C13 C14

0.46 0.06 0.50 0.50 0.89 0.03 0.08 0.08 1.04 -0.51 -0.25 -0.27 -0.73 0.93

C15 C16 C17

0.93 0.27 -1.33


15

Table 7 Some minor element concentrations of the GGF thermal and cold water samples. Group

Water samples

Group 1 Group 1A

Thermal waters Na-SO4 type Ca-SO4 type Ca-HCO3 type Cold waters Ca-HCO3 type Mg-HCO3 type

Group 1B Group 2 Group 2A Group 2B

Fe (mg/l)

Li (mg/l)

B (mg/l)

Sr (mg/l)

Si (mg/l)

Ba (mg/l)

F (mg/l)

Mn (mg/l)

Zn (mg/l)

Al (mg/l)

As (mg/l)

0.08 0.14 0.98

1.58 1.63 0.66

7.51 7.00 4.23

14.7 12.3 5.5

18.4 58.2 24.9

0.14 0.08 0.07

2.63 2.55 1.06

0.120 0.054 0.073

0.020 0.011 0.010

1.590 0.016 0.009

0.010 0.010 0.016

0.10 0.09

0.05 0.21

0.06 0.21

0.17 1.60

5.20 15.9

0.05 0.07

0.21 0.20

0.003 0.005

0.002 0.012

0.002 0.004

0.006 0.003

display higher and variable δ18O and δ2H values (Fig. 9A), indicating a lower altitude of meteoric recharge. In the GGF, it is obtained that δ18O values decrease − 0.20‰ and δ2H values decrease − 0.8‰, per 100 m rise in altitude. This altitude effect is an indicator for recharge areas and used to calculate recharge areas for thermal waters by cold waters (Clark and Fritz, 1997). The equations are determined as “altitude (m) = − 548.3 ∗ δ18O − 3542” in the oxygen-altitude and “altitude (m) = − 131.8 ∗ δ2H − 6136” in the deuterium-altitude graphs. Accordingly, the recharge elevation of the GGF thermal waters are calculated as 780–840 m a.s.l., indicating that the recharge area is at higher altitudes of the Yenice Horst. The combination of relatively high δ18O values (mean of −8.22‰) and low EC (mean of 467 μS/cm) and low HCO3 (mean of 308 mg/l) values for all cold water types (Fig. 9B and C) reflects a low elevation recharge and a shallow circulation path. 5.5. Tritium (3H) isotopes Tritium is used to provide residence time and mixing of waters in geothermal systems (Panichi and Gonfiantini, 1978). Also, they can separate modern (b 50 years in age) from ancient (N 50 years in age) origin of groundwater (Clark et al., 1997; Clark and Fritz, 1997). Tritium values below 1 TU are commonly attributed to represent ages older than 50 years and values above 1 TU are characterized by modern groundwater (Ravikumar and Somashekar, 2011). The tritium contents ranging from 1 to 8 TU is considered as an admixture of recent water with old groundwater and groundwater having been subjected to radioactive decay (Ravikumar and Somashekar, 2011). In the GGF, tritium values from all groups of thermal waters are between 0.05 to 1 TU, indicating ancient or pre-modern waters (N50 years in age) (Table 3). The tritium contents of the sample C13 from Ca-HCO3 type cold water (2.73 TU) indicate radioactive decay (1–8 TU), thus indicating a mixture of old and new water recharge. Differently, C2 and C7–C10 samples from Ca-HCO3 type cold waters (10.6–17.4 TU) reflect recent water recharge (Ravikumar and Somashekar, 2011) (Table 4). Moreover, sample C15 from Mg-HCO3 type cold water (0.39 TU) shows ancient waters, with N 50 years in age. Tritium-Cl and tritium-EC relationships were used to separate shallow from deep circulating waters (Ravikumar and Somashekar, 2011). The combination of low tritium values with high EC and Cl values in the

Gölemezli thermal waters (Table 3, Fig. 9D and 9E) suggests deep circulation. Moreover, the Group 1A (SO4-rich) thermal waters indicate a deeper circulation with respect the one suggested by the Group 1B (HCO3-rich) thermal waters. Contrastingly, high tritium values and low EC and Cl contents of cold waters (Table 4, Fig. 9D and E) imply younger and shallower circulation, with a short residence time. 5.6. δ13C isotopes To define the evolution of dissolved inorganic carbon (DIC) and sources of carbon in waters, δ13C isotopes are commonly used (e.g., Mook and Tan, 1991; Truesdell and Hulston, 1980). The main source of dissolved inorganic carbon (DIC) in the natural waters is CO2 deriving from decaying organic matter in soils, dissolution of carbonates, metamorphic reactions, and from magma and/or mantle degassing (e.g., Hoefs, 2009). As seen in Table 3, the δ13CDIC ratios of the Group 1A (+ 5.11 to + 7.54‰ PDB, mean of + 6.28‰) and Group 1B thermal waters (+ 6.18 to + 6.98‰ PDB, mean of + 6.58‰) are similar and display positive values (Fig. 10 and Table 3). In contrast, the Group 2A cold waters have very negative ratios (− 10.38 to − 8.16‰, mean of − 9.27‰), whereas the δ13CDIC ratios of the Group 2B cold waters (−9.87 to −4.09‰, mean of − 7.90‰) are higher than the Group 2A waters (Fig. 10 and Table 4). As seen in Fig. 10, Group 1A and 1B thermal waters with positive δ13CDIC ratios indicate that the source of carbon may be mainly metamorphic CO2 (decarbonation of carbonate rocks), whereas the negative δ13CDIC ratios of the Group 2A and 2B cold waters are groundwater DIC in origin. The δ13CDIC ratios in travertine (+3.7 to +5.0‰; Özkul et al., 2013) of the Gölemezli field is consistent with the δ13CDIC ratios of the Group 1A and 1B thermal waters, reflecting a thermogene type origin (Gandin and Capezzuoli, 2008; Capezzuoli et al., 2014). Similarly, the banded and bedded travertines in the Gölemezli have positive δ13CDIC ratios ranging from +3.11 to +5.07‰ and +5.46 to +6.14‰, respectively (TUBITAK, 2016) (Fig. 9F). The thermogenic travertine with relatively high δ13CDIC values (−3 to +8‰) accumulates rapidly from high-temperature waters during cooling and displays a low amount of organic material and a massive structure, mostly reflecting tectonic activities (Pentecost, 2005). The endogenic travertine deposition is commonly controlled by active tectonic process interactions with fluid pressure and paleoclimatic conditions (Rihs et al., 2000; Faccenna et al., 2008).

Table 8 REE concentrations of the GGF thermal waters, Neogene formations and bedrock samples. La (ppm) GGF thermal waters (this study) Neogene formations (Alçiçek, 2007)

Menderes Massif metamorphics (Möller et al., 2004)

Ce (ppm)

Pr (ppm)

Nd (ppm)

Sm (ppm)

Eu (ppm)

Gd (ppm)

Tb (ppm)

Dy (ppm)

Y (ppm)

Ho (ppm)

Er (ppm)

Yb (ppm)

Lu (ppm)

0.00065 0.00013 0.00015 0.00006 0.00002 0.00001 0.00025 0.00001 0.00002 0.00001 0.00001 0.00003 0.00015 0.00001 Kızılburun Fm. Sazak Fm. Kolankaya Fm. Mean Schist + marble

0.50 3.77 14.60 6.29 20.10

0.60 7.27 29.20 12.36 37.68

0.08 0.71 3.28 1.36 4.66

0.40 3.23 11.90 5.18 16.39

0.10 0.70 2.42 1.07 3.04

0.05 0.19 0.50 0.25 0.76

0.08 0.55 1.97 0.87 2.84

0.01 0.10 0.32 0.14 0.43

0.09 0.51 1.81 0.80 2.59

0.30 2.57 10.50 4.46 15.54

0.05 0.16 0.34 0.18 0.52

0.05 0.32 1.02 0.46 1.49

0.05 0.32 0.94 0.44 1.32

0.01 0.05 0.15 0.07 0.19

16

H. Alçiçek et al. / Journal of Volcanology and Geothermal Research 349 (2018) 1–30 40 35 30 25 20

GGF thermal waters (This study) Menderes metamorphics (Möller et al., 2004) Kızılburun Fm. (Alçiçek, 2007)

15

Sazak Fm. (Alçiçek, 2007) 10 5

Kolankaya Fm. (Alçiçek, 2007)

0

Neogene formations (mean)

-5

Fig. 8. Comparison of REE concentrations of the GGF thermal waters, Neogene formations and bedrock samples.

The calculated δ13CCO2 values in the Gölemezli banded and bedded travertine are between − 3.30 and − 2.35‰ (mean of − 2.99‰) (TUBITAK, 2016). Likewise, the δ13CCO2 values have negative values with −3.5 and − 2.8‰ for the Yenice and with − 4.1 and − 3.5‰ for the Pamukkale and −4.2‰ for the Karahayıt geothermal fields nearby the GGF (Filiz, 1982, 1989; Ercan et al., 1994) (Fig. 2). On the other hand, the CO2 gases from the Kızıldere and Tekkehamam geothermal fields in the Denizli Basin (Fig. 2) have also slightly negative δ13CCO2 values (−0.62‰, Mutlu et al., 2008 and −1.2‰; Ercan et al., 1994, respectively). These values indicate mantle-derived magmatic origin (e.g., Filiz, 1982; Hilton et al., 2002; McCollom et al., 2010). Therefore, it can be assumed that the origin of the CO2 gases in the Gölemezli field is also mantle-derived. 6. Geothermometry applications: reservoir temperature estimation In geothermal systems, various geothermometers have been developed to estimate reservoir temperature (e.g. Nicholson, 1993). In this study, silica and cation geothermometers, mixing models, anyhdrite/ chalcedony saturation indexes and mineral saturation states (SI) methods are applied to the GGF. 6.1. Silica geothermometers Silica geothermometers (quartz, chalcedony, SiO2 vs. K2/Mg equilibrium diagram and silica enthalpy mixing model) are extensively governed by dissolved silica content and by the solubility of different silica species (e.g., Fournier, 1977). Such geothermometers are commonly used to identify the temperature of water–rock interaction processes at depth in ascending fluids, before their discharge (e.g., Fournier, 1977; Nicholson, 1993). The results of the quartz geothermometer give reservoir temperatures varying from 95 to 120 °C for Na-SO4 type, 145 to 157 °C for CaSO4 type and 99 to 114 °C for Ca-HCO3 type (Table 9). Additionally, temperatures estimated by the chalcedony geothermometer are encompassed between 64 and 92 °C for Na-SO4 type, 119 and 133 °C for Ca-SO4 type and 69 and 85 °C for Ca-HCO3 type. To determine silica species in the thermal waters and estimate reservoir temperature, the log (SiO2) vs. log (K2/Mg) equilibrium diagram is used (Fig. 11A). Ca-SO4 type thermal waters are scattered around the line of chalcedony (Fig. 11A). This indicates that chalcedony occurred in Ca-SO4 type thermal waters and might control dissolved silica. Besides, Na-SO4 and Ca-HCO3 type thermal waters are plotted between the lines of chalcedony and quartz (Fig. 11A), reflecting that both silica

species formed in these waters, affecting dissolved silica. Consequently, in the GGF, the silica solubility is governed by chalcedony and quartz, supported by the SIs for chalcedony and quartz are very close to zero, indicating their equilibriums with thermal waters (Table 10). According to this method, reservoir temperature of the GGF was obtained varying from 70 to 90 °C (Fig. 11A). 6.2. Mixing models Silica enthalpy mixing model is used to obtain the reservoir temperature of geothermal fields and assess the effects of the mixing processes (e.g., Nicholson, 1993). Fig. 11B presents the silica-enthalpy mixing model according to chalcedony and quartz solubilities. Two end member fluids have been given in this model: a cold water sample (C13 sample; temperature: 13.4 °C and SiO2: 24 ppm; Table 4) as one end member and the thermal waters as the other end member. In this model, thermal waters occur as the result of mixing of thermal water with cold water, assuming no steam or heat loss. The intersection point with the solubility curve for quartz (black curve line) gives the reservoir temperature, reflecting 145 °C for Na-SO4 and Ca-HCO3 type (Fig. 11B) and 290 °C for Ca-SO4 type (Fig. 11C) thermal waters. The pink curve line of chalcedony between the cold and thermal waters intersects the chalcedony solubility curve, suggesting the estimated reservoir temperature of 100 °C for Na-SO4 and Ca-HCO3 type (Fig. 11B) and 230 °C for Ca-SO4 type (Fig. 11C) thermal waters. Chloride enthalpy model is also one of the mixing models and commonly used to reveal the hydrologic conditions of a geothermal field (Fournier, 1979). In this model (Fig. 11D), the enthalpies of the thermal waters are plotted against their corresponding chloride contents. The reservoir temperature is found by drawing (i) a line from the hot spring with highest chloride content to the steam point (=boiling line) and (ii) a perpendicular line from the hot spring with equal enthalpy, but less chloride content, to the boiling line. This intersection refers to the minimum enthalpy/temperature value of the reservoir (Enthalpy = 898 kj/kg, T = 210°C). 6.3. Cation geothermometers Cation geothermometers (Na-K, Na-K-Ca-Mg, K-Mg, Na-K-Mg and Na-K/Mg-Ca) are based on chemical equilibration between phases involving ion-exchange reactions (e.g., Fournier, 1991). Na-K geothermometer yields very high results in a wide range between 208 and 297 °C for Na-SO4 type, 225 and 248 °C for Ca-SO4 type and 258 and 339 °C for Ca-HCO3 type for the GGF (Table 9).

H. Alçiçek et al. / Journal of Volcanology and Geothermal Research 349 (2018) 1–30 -12

-10

-8

-6

-4

-2

0

2

4

0

6

500

1000

1500

2000

17 2500

3000

3500

4000

4500

60

5000 0 -1

40 -2

H2Sexchange

0

-20

-9

-8

-7

-6

-4

-7

-9

Water-rock interaction

-80

(‰ VSMOW)

-5

-6

-8

CO2exchange

-10

-5

-60

Evaporation

A

-4

-40

Hydration of silicates

δ 18O

-3 δ 18O (‰)

δ 2H (‰ VSMOW)

20

-3

-2

-1

0

-10

EC (µS/cm)

B 0

2

4

6

8

10

12

14

16

18

3000

20 100

90 2500

80

70

1000

60

Deepcirculation

1500

Cl (mg/l)

HCO3 (mg/l)

2000

50 40 30

Shallow circulation 20

500

10 0

δ 18O (‰ SMOW)

C 0

500

1000

1500

2000

2500

3000

3500

4000

4500

5000 20

-40

0

Tritium (TU)

D -35

-30

-25

-20

-15

-10

-5

0 8

18

7

16

6 12 10 8

Shallow circulation

δ 13C (‰ VPDB)

Tritium (TU)

14 5 4 Kızılburun Fm.

3

Sazak Fm.

Deep circulation

6

Kolankaya Fm.

4

2

Gölemezli bedrock

1

Gölemezli banded travertine

2

Gölemezli bedded travertine

EC (µS/cm)

E Group 1A:

Thermal waters: Na-SO4

Thermal waters: Ca-SO4 Group 1B:

0

F Thermal waters: Ca-HCO3 Group 2A:

δ 18O (‰ VPDB)

Cold waters: Ca-HCO3

Group 2B:

0

Cold waters: Mg-HCO3

Fig. 9. (A) Plot of δ18O-δ2H for some waters in the GGF. Arrow shows salinity mixing line, GMWL: Global Meteoric Water Line (Craig, 1961), EMMWL: Eastern Mediterranean Meteoric Water Line (IAEA, 1981), LMWL: Local Meteoric Water Line (Şimşek, 2003a); (B) δ18O-EC diagram; (C) δ18O-HCO3 diagram; (D) Tritium-Cl diagram; (E) EC-Tritium diagram; (F) δ18O-δ13C diagram.

The Na-K-Ca-Mg geothermometer is used to eliminate the possible effects of Ca contents on the Na-K geothermometer (Fournier and Truesdell, 1973). It also provides very high values ranging from 191 to 239 °C for Na-SO4 type, 198 to 213 °C for Ca-SO4 type and 212 to 262 °C for Ca-HCO3 type; Table 9). The K-Mg geothermometer yields very low temperatures varying from 74 and 83 °C for Na-SO 4 type, 71 and 79 °C for Ca-SO 4 type and 75 and 90 °C for Ca-HCO 3 type. These temperatures are similar to the discharge temperatures (49.8 to 88 °C for Na-SO 4 type; 41.9 to 65 °C for Ca-SO 4 type and 53 to 70.7 °C for Ca-HCO 3 type) (Tables 3 and 9). As seen in Fig. 12, the Na-K-Mg diagram (Giggenbach, 1988) reflects that all GGF samples plot in the immature water field, due to the disequilibrium nature of the GGF waters. This case indicates the effects

of processes such as re-equilibrium and strong dilution/mixing with shallow cold water (Giggenbach, 1988). This method gives anomalously high temperature results, indicating a reservoir temperature of 260 °C for Na-SO4 and Ca-SO4 type and 300 °C for Ca-HCO3 type waters (Fig. 12). Furthermore, Na-K/Mg-Ca diagram (Fig. 13) of Giggenbach (1988) also applied to the GGF thermal waters. In this method, 10Mg/(10Mg + Ca) vs. 10K/(10K + Na) ratios of the GGF thermal waters are plotted. Accordingly, it is not observed any equilibrium between the reservoir rocks and thermal waters. In addition, this diagram indicates that the solutes in the GGF thermal waters are originated from the dissolution of reservoir rocks. Overall, reservoir temperature estimated using this method is 110–210 °C for the GGF.

18


Fig. 10. Carbon isotopes ratios of the GGF thermal waters (based on Trumbore and Druffel, 1995).

6.4. Anyhdrite/chalcedony saturation indexes

6.5. Mineral saturation states (SI)

To estimate the equilibrium temperature of thermal waters, anhydrite/chalcedony saturation indexes were also applied (Chiodini et al., 1995; Pastorelli et al., 1999; Levet et al., 2002). This method is used for both anhydrite and chalcedony (showing a retrograde and a prograde solubility, respectively) to estimate aquifer temperatures in evaporitic formations. Equilibrium temperatures can be recognized by the saturation indices (SI) of both minerals. These are equal and close to zero at the same temperatures. In this study, this method is applied to the GGF thermal waters that circulated within the evaporite-bearing rocks of the Sazak and Kolankaya formations. The anhydrite and chalcedony give wide temperature range and are both close to equilibrium in the temperature range (72–127 °C for Na-SO4 type, 67–130 °C for Ca-SO4 type and 78–116 °C for Ca-HCO3 type) of all types of waters (Fig. 14).

In the groundwater systems, mineral equilibrium calculations are used to evaluate the presence of reactive minerals and to estimate the mineral reactivity. The saturation index method is considered to predict the reactive minerals in the subsurface from the groundwater chemical data, without examining rock samples (e.g., Deutsch, 1997; Reed and Spycher, 1984). In order to predict the precipitation or dissolution phases in the flow paths of thermal waters, the mineral saturation indices can be used for some carbonate (calcite, aragonite and dolomite), sulfate (gypsum and anhydrite), and silica (quartz and chalcedony) minerals (Reed and Spycher, 1984). A saturation index of zero indicates a saturated solution in comparison to the solid phase, whereas a positive saturation index displays an oversaturated solution and a negative saturation index implies an undersaturated solution. In this study, the

Table 9 Reservoir temperatures (°C) estimated using Na-K, chalcedony, quartz, K-Mg and Na-K-Ca geothermometers. Na-SO4 type (Group 1A) SiO2 (quartz-no steam loss) SiO2 (quartz-maximum steam loss) SiO2 (chalcedony) Na-K Na-K-Ca, Mg corrected K-Mg

Fournier, 1977 Fournier, 1977 Fournier, 1977 Giggenbach, 1988 Fournier and Truesdell, 1973 Giggenbach, 1988

Ca-SO4 type (Group 1A) SiO2 (quartz-no steam loss) SiO2 (quartz-maximum steam loss) SiO2 (chalcedony) Na-K Na-K-Ca, Mg corrected K-Mg


Ca-HCO3 type (Group 1B) SiO2 (quartz-no steam loss) SiO2 (quartz-maximum steam loss) SiO2 (chalcedony) Na-K Na-K-Ca, Mg corrected K-Mg


G1

G2

G3

G4

G5

G6

G7

G8

Mean

100 101 70 229 199 74

103 103 73 238 206 77

– – – 238 208 79

– – – 242 210 78

100 101 70 208 191 83

95 96 64 250 219 81

120 119 92 297 239 76

104 104 74 241 210 79

104 104 74 243 210 78

G9

G10

G11

G12

G13

G14

G15

G16

G17

G18

G19

Mean

148 141 122 244 203 71

148 141 122 236 205 75

145 140 120 225 198 73

151 144 126 236 200 73

145 139 119 248 213 77

150 144 125 246 213 79

154 147 130 245 212 78

157 149 133 243 211 79

147 141 122 242 208 77

153 146 128 246 213 79

148 142 123 244 210 75

150 143 125 241 208 76

G20

G21

G22

G23

G24

G25

G26

G27

G28

Mean

– – – 296 246 90

101 102 71 339 262 87

103 104 73 320 250 85

105 105 75 320 252 83

105 105 75 258 212 76

114 113 85 300 233 82

104 104 74 298 227 75

107 107 77 304 242 84

99 100 69 319 243 87

105 105 75 306 241 83


A

B

C

D

19

Fig. 11. (A) The SiO2 log versus K2/Mg log binary plot (based on Giggenbach et al., 1994); (B, C) Silica enthalphy models for Na-SO4/Ca-HCO3 and Ca-SO4 type thermal waters, respectively, assuming no steam or heat steam loss before mixing; (D) Chloride-enthalpy diagram for the GGF thermal waters.

mineral saturation indices were calculated for many common minerals, at the measured discharge temperatures (Table 9). Mineral saturation indices for the GGF thermal waters are calculated using the “Aquachem-Phreeqc” software (Parkhurst and Appelo, 1999). The results show that all thermal water samples are mainly undersaturated with respect to anhydrite, gypsum, amorph silica and fluorite; and oversaturated with respect to calcite, aragonite, dolomite, quartz and chalcedony (Table 10). As a consequence, it should predict a sequence of deposition as calcite, aragonite, dolomite, hematite, quartz and chalcedony minerals in the Gölemezli system. The equilibrium state among waters and minerals is temperature-dependent, assessing for temperature calculations for the flow path (Reed and Spycher, 1984). Saturation indices of the Gölemezli thermal waters (Fig. 15A–F) are estimated at discharge temperatures and pH measured values. The hydrogen mass-balance causes changes of temperatures (Kharaka and Mariner, 1988) and saturation indices are recalculated to provide the equilibrium states of some hydrothermal minerals at distinct temperatures. If the equilibrium lines of a group of estimated minerals converge to a common point, this indicates a temperature corresponding to the most probable reservoir temperature (Tole et al., 1993). Considering the samples G2 and G8 from Na-SO 4 type, saturation indices with respect to chalcedony and gypsum minerals tend to get closer to zero (SI = 0), around the temperature of 147 °C (Fig. 15A) and 127 °C (Fig. 15B), respectively. In the samples G14 and G19 from Ca-SO4 type, chalcedony and gypsum minerals are saturated at temperature of 127 °C (Fig. 15C) and 135 °C (Fig. 15D), respectively. Regarding samples G27 and G28 from Ca-HCO 3 type, anhydrite and quartz minerals tend to be near to zero (SI = 0) around temperature of 110 °C (Fig. 15E and F, respectively). The assessment of the saturation indices of the minerals of Fig. 15 gives a reservoir temperature between 110 °C (anhydrite and quartz) and 147 °C (chalcedony and gypsum) for the GGF.

6.6. Summary of geothermometer applications In summary, the calculated reservoir temperatures by different geothermometers for the GGF thermal waters cover a wide range of values (Fig. 16). The log (SiO2) vs. log (K2/Mg) and K-Mg geothermometers yield very low results which are similar to the discharge temperatures. Thus, they cannot be considered realistic. The Na-K geothermometer can give erroneous results at temperatures below 180°C and is unsuitable for waters with high Ca contents, because temperature-dependent exchange equilibrium between feldspars and thermal waters is not attained at low temperatures and the Na/K ratio in the waters are controlled by leaching rather than chemical equilibrium (Giggenbach, 1988). The Na-K-Ca-Mg geothermometer gives similar values to the Na-K results. Therefore, both geothermometers cannot apply to the GGF. The Na-K-Mg geothermometer also gives unreasonable results due to mixing and dilution of thermal waters, as suggested by Nicholson (1993). But, it is observed that the chalcedony, quartz, silica and chloride enthalpy mixing models, Na-K/Mg-Ca, anhydrite-chalcedony and mineral saturation states methods provide reasonable results which best-fit temperatures ranging from 130 to 210 °C, and thus considered as the most reliable as reservoir temperatures. It is seen that the reservoir temperatures of the GGF cannot reach up to 250 °C, supported by the lack of strongly positive Eu anomaly (Fig. 8; Möller, 2000). 7. Discussion 7.1. Conceptual model for the GGF based on the chemical–physical data The hydrochemical properties and isotopic compositions of the GGF thermal and cold waters allow drawing of a schematic cross-section showing the hydrothermal circulation pathways (Figs. 17 and 18). As observed in Fig. 17, the GGF is mostly recharged by widespread infiltrations of meteoric waters, trapped in sectors of the Paleozoic formations (deep reservoir) and of the Neogene (shallow reservoir) formations

20

Table 10 Saturation index in discharge temperature of the GGF thermal and cold water samples. Sample no

Group 1

Thermal waters Na-SO4 type G5 G7 G8 Ca-SO4 type G13 G14 G15 G16 G17 G18 G19 G20 Ca-HCO3 type G25 G27 G28 Cold waters Ca-HCO3 type C11 C12 C13 C14 C15 C16 C17

Group 1A

Group 1B

Group 2 Group 2A

Group 2B

si_Calcite si_Aragonite si_Dolomite si_Anhydrite si_Gypsum si_Albite si_Anorthite si_Strontianite si_Talc si_Hematite si_Halite si_Chrysotile si_SiO2(a) si_Quartz si_Chalcedony si_Fluorite

0.32 0.40 0.42

0.42 0.29 0.30

0.65 0.74 0.75

−0.22 −0.20 −0.20

−0.18 −0.17 −0.19

−0.21 −0.20 −0.21

−2.75 −2.45 −2.55

−0.94 −0.80 −0.84

−1.44 15.55 −1.30 16.40 −1.34 17.33

−5.91 −5.90 −5.99

−5.51 −6.51 −6.52

−0.17 −0.15 −0.18

0.85 0.90 0.92

0.55 0.52 0.57

−1.02 −1.01 −1.03

0.85 0.74 −0.11 0.66 0.27 0.82 0.13 −0.63

0.72 0.61 −0.24 0.54 0.14 0.70 0.00 −0.78

1.64 1.40 −0.30 1.20 0.40 1.48 0.23 −1.33

−0.28 −0.19 −0.22 −0.11 −0.19 −0.02 −0.31 −1.18

−0.18 −0.14 −0.16 −0.15 −0.10 −0.12 −0.18 −0.83

−1.61 −1.80 −3.67 −1.94 −2.41 −3.81 −2.54 −2.44

−5.85 −5.96 −10.00 −6.02 −7.44 −7.10 −8.41 −6.57

−0.37 −0.52 −1.35 −0.67 −1.01 −0.53 −1.06 −1.80

1.38 0.57 −4.43 0.25 −2.49 −0.43 −2.81 −6.05

18.21 17.83 16.59 17.38 17.34 17.43 17.20 16.44

−6.13 −6.05 −6.05 −6.04 −6.03 −6.06 −6.08 −7.10

−3.81 −4.65 −9.73 −5.12 −7.71 −4.67 −8.27 −10.59

−0.15 −0.15 −0.10 −0.10 −0.13 −0.67 −0.01 −0.40

1.00 0.97 1.03 0.98 1.01 0.37 1.15 0.90

0.63 0.61 0.67 0.64 0.64 0.05 0.78 0.45

−0.95 −0.98 −0.97 −1.04 −0.90 −1.10 −0.92 −0.84

1.45 1.41 1.04

1.28 1.29 0.92

2.35 2.63 1.72

−0.50 −0.54 −0.71

−0.68 −0.58 −0.69

−2.61 −3.48 −1.72

−4.63 −6.66 −3.06

−0.25 −0.20 −0.50

0.50 18.50 1.52 17.88 −0.04 18.32

−5.25 −6.55 −6.87

−3.95 −2.91 −4.37

−0.58 −0.56 −0.60

0.50 0.51 0.50

0.15 0.17 0.16

−0.93 −0.95 −0.91

0.09 0.44 0.36 0.81 0.10 0.81 0.34

−0.06 0.29 0.21 0.65 −0.05 0.66 0.19

−0.86 −0.07 −0.06 0.44 0.31 1.64 1.02

−2.91 −2.96 −2.85 −2.71 −2.98 −1.75 −1.75

−2.46 −2.49 −2.42 −2.23 −2.56 −1.36 −1.33

−2.82 −2.51 −2.89 −1.85 −0.58 −1.52 −1.36

−4.89 −4.01 −4.58 −3.30 −3.02 −3.93 −3.45

−2.17 −1.73 −1.85 −1.38 −1.07 −0.22 −0.99

−3.29 −1.97 −1.22 −0.68 1.40 1.56 −0.24

−9.14 −9.68 −10.12 −8.79 −8.37 −7.28 −7.45

−7.12 −5.94 −5.33 −4.66 −3.55 −2.45 −4.09

−0.72 −0.65 −0.59 −0.64 −0.18 −0.66 −0.73

0.64 0.71 0.75 0.74 1.17 0.67 0.62

0.16 0.24 0.29 0.26 0.70 0.21 0.15

−4.81 −4.69 −4.78 −4.61 −5.03 −5.00 −5.02

18.00 18.02 17.58 18.60 17.39 17.79 17.74


Group


21

Fig. 12. Distribution of the thermal and cold waters from the GGF in a Na-K-Mg1/2 triangular diagram (based on Giggenbach, 1988).

(Şimşek, 1984, 1990). The recharge area is 780–840 m.a.s.l. of altitude, in the northern part of the study area (the Yenice Horst; Fig. 2). The NW-trending normal and NE-trending orthogonal fault systems and associated fractures, bounding the basin to the north (Figs. 17 and 18) are considered as the main conduit for both the cold water flowing down and the hot water circulating at depth. After having been heated at depth, the GGF thermal waters ascend along the major NW-trending

bounding faults and emerge as thermal springs flowing towards the center of the basin. They mix eventually with shallow colder groundwater. Their vertical connecting vertically due to the NW-trending normal faults developed inside the minor horst and graben structures bounded the Neogene formations. Although knowledge about depths at which thermal springs formed is difficult to estimate, the average geothermal gradient in many regions is generally 30 °C/km and a mean surface

Fig. 13. Na-K/Mg-Ca diagram for thermal waters from the GGF (based on Giggenbach, 1988).

22


Na-SO4 type water si_Anhydrite G1

1.5

si_Chalcedony G1 si_Anhydrite G2

1

si_Chalcedony G2

Log Q/K

si_Anhydrite G3

0.5


0

si_Chalcedony G4 si_Anhydrite G5 si_Chalcedony G5

-0.5

si_Anhydrite G6 si_Chalcedony G6

-1

127

72

-1.5 25

45

65

85

105

125

145

si_Anhydrite G7 si_Chalcedony G7 si_Anhydrite G8 si_Chalcedony G8

Temperature°(C)

A

Ca-SO4 type water si_Anhydrite G13

1.5


1


Log Q/K

0.5


0


-0.5


-1

si_Chalcedony G18

130

67

-1.5 25

45

65

85

105

125

145

Temperature°(C)

B

Ca-HCO3 type water


si_Anhydrite 21 si_Chalcedony G21

1.5

si_Anyhdrite G22 si_Anydrite G22

1

Log Q/K

si_Anhydrite G19

si_Anydrite G23 si_Chalcedony G23

0.5


0


-0.5

si_Anydrite G26

25

C

45

65

85

si_Chalcedony G26

116

78

-1

105

Temperature°(C)

125

si_Anhydrite G28

145


Fig. 14. Log SI of the Gölemezli thermal waters with respect to the combination of anhydrite and chalcedony saturation indexes (based on Pastorelli et al., 1999).

(atmospheric) temperature is 15 °C. Therefore, ca. 5 km reservoir depth is required to reach temperatures of up to 130–210 °C. This result may indicate the presence of possible third reservoir within the Precambrian gneisses in the GGF, as well as in the Kızıldere Geothermal Field in the Denizli Basin as reported by Şimşek et al. (2005). Further researches are still required to demonstrate the presence of such reservoir. Differences in the GGF thermal water compositions can be explained by several processes: (i) water-rock interaction in shallow and deep reservoir rocks, (ii) mixing between shallow meteoric and deep thermal waters, (iii) evaporation processes and (iv) input of magmatic fluids. These processes affecting the thermal waters of the GGF are based on the chemical and isotopic compositions and physicochemical characteristics (e.g., temperature, pH, electrical conductivity). The SO4-rich Group

1A thermal waters (Na-SO4 and Ca-SO4 types) are mainly governed by oxidation of H2S to H2SO4 in the steam-heated aquifer (Fig. 5) (Nicholson, 1993) and/or input of magmatic H2S and SO2 volatiles, as confirmed by the presence of magmatic CO2 (Ellis and Mahon, 1977; Giggenbach, 1988). They are also driven by leaching of sulfate-bearing reservoir rocks (saline lake units; Fig. 3) of the Sazak Formation. Besides, the HCO3-rich Group 1B thermal waters (Ca-HCO3 type) are probably affected by input of CO2-rich magmatic volatiles in the periphery of geothermal fields (Fig. 5), controlled by leaching of shallow reservoir carbonate rocks of the Neogene formations and deep reservoir marble and schist of the Menderes Massif (Fig. 3). Very high Sr contents in the Menderes metamorphics, Neogene formations, late Miocene-Pliocene volcanics (Paton, 1992) and Quaternary travertines (Khatib et al.,


G8- Na-SO4 water type

G2- Na-SO4 water type 2

2 si_Calcite 1

si_Dolomite

0.5

si_Anhydrite

si_Calcite

1.5

si_Aragonite

si_Aragonite

1

si_Dolomite SI

SI

1.5

0.5 si_Anhydrite 0

si_Gypsum

si_Gypsum

0

-0.5

si_Strontianite

-0.5 -1 25

75

125

Temperature (°C)

A

si_Chalcedony

si_SiO2(a)

-1.5

si_Quartz 147°C

si_Strontianite

-1

si_SiO2(a)

25

127°C

si_Quartz si_Chalcedony

2 si_Calcite

2

si_Dolomite

0.5

si_Anhydrite

0

si_Dolomite

25

75 Temperature (°C)

C

125 127°C

si_Strontianite si_SiO2(a)

-1.5


si_Gypsum

-1

si_SiO2(a)

-1.5

si_Anhydrite -0.5

si_Strontianite

-1

0.5 0

si_Gypsum

-0.5

si_Aragonite

1 SI

1

si_Calcite

1.5

si_Aragonite

1.5 SI

125

G19- Ca-SO4 type

G14-Ca-SO4 type

25

3.5 3 2.5 2 1.5 1 0.5 0 -0.5 -1 -1.5

si_Aragonite

si_Anhydrite

SI

si_Dolomite

si_Gypsum si_Strontianite si_SiO2(a)

75 110°C Temperature (°C)

125

125 135°C


G28-Ca- HCO3 type si_Calcite

25


D

G27-Ca-HCO3 type

SI


B

2.5

E

23


3 2.5 2 1.5 1 0.5 0 -0.5 -1 -1.5

si_Calcite si_Aragonite si_Dolomite

si_Anhydrite si_Gypsum si_Strontianite si_SiO2(a)

25

F

75 110°C Temperature (°C)

125


Fig. 15. Mineral equilibrium diagram for thermal waters from the GGF: (A, B) Samples G2 and G8 from Na-SO4 type; (C–D) Samples G14–G19 from Ca-SO4 type, and (E–F) Samples G27G28 from Ca-HCO3 type (respectively).

2014; Lebatard et al., 2014a, b) confirm that the GGF thermal waters originate both the interaction between mantle-derived gas and shallow/deep reservoir rocks. In view of the different trends of cation/Cl ratios in the waters, SO4rich Group 1A and HCO3-rich Group 1B thermal waters appear to be separated from the hydrothermal circulation in two distinct ascending pathways (Fig. 7A–J). Along these two pathways, both the shallow reservoir within the Neogene formations and the deep reservoir hosted in the Menderes metamorphics control the chemical and isotopic compositions of the Group 1A and Group 1B thermal waters. Both regional and local flow systems manage emergence of the Gölemezli thermal springs of which chemical and isotopic compositions reflect the mixing with cold shallow groundwater during their ascent (Fig. 9A). Wiersberg et al. (2011) documented that noble gas abundance also reflects mixing with cold meteoric water in the reservoirs. As seen in Fig. 9A, all types of thermal waters plotted to the right side of the GMWL line, indicating slight evaporation effects in a semi-arid climate. This case also reflects that oxygen and hydrogen isotopic values are not significantly affected by isotopic exchange reactions with the reservoir

rocks at depth. The low oxygen isotope and low tritium values (ca.− 8‰ and b 1 TU, respectively; Table 3) confirm the mixing of deep thermal fluids with shallow waters. This mixing mechanism can result from direct infiltration of local meteoric water in the thermal springs in the GGF. Besides, the negative δ18O and high tritium isotope compositions (mean of −8‰ and N1 TU, respectively; Table 4) of the cold waters suggest shallow and rapid circulation linked to recent precipitations. The GGF thermal waters with high electrical conductivity (EC), high trace element concentrations and the low δ3H (b 1 TU) values indicate that the waters circulate at major depth, as supported by the discharge and well temperatures of the springs (Table 3). Therefore, they are envisaged to have a longer time of residence in the reservoirs. Higher EC values reflect longer (and deeper) circulation and residence time compared to lower EC values. Accordingly to the northeast, the Ca-SO4 type waters have the highest electrical conductivity (4251 μS/cm), whereas those belonging to the Na-SO4 type waters display moderate electrical conductivity values (3910 μS/cm) indicating circulated for a shorter time period. Similarly to the northeast, the Ca-HCO3 type waters

24


Fig. 16. Comparison of all calculated geothermometers from the GGF thermal waters.

with the lowest electrical conductivity (2987 μS/cm) account for a minor residence time and faster circulation if compared to those with the SO4-rich thermal waters (Fig. 17). This case indicates that the residence time increases from NE to SW and the circulation of the thermal waters is from SW to NE in the study area. 7.2. Origin of CO2 in the GGF In the Denizli Basin, the thermal waters are commonly characterized by high CO 2 contents with minor amount of N2 , O 2 and H 2 (Ercan et al., 1994; Akan and Şimşek, 1997; Güleç, 1988; Wiersberg et al., 2011), similar to those of other western Anatolian basins (Mutlu et al., 2008; Güleç and Hilton, 2016). The CO2 is also major component for geothermal gases of both shallow and deep reservoirs in the GGF occupying 72 to 99% in volume proportion of whole gases (Ercan et al., 1994; Akan and Şimşek, 1997). The volume proportions of CO2 in the geothermal gas samples reach up to 69-86% in the Pamukkale (Akan and Şimşek, 1997), 66-91% in the Karahayıt (Akan and Şimşek, 1997), 96-98% in the Kızıldere (Mutlu et al., 2008; Wiersberg et al., 2011) and 83-97% in the Tekkehamam (Güleç, 1988; Wiersberg et al., 2011) geothermal fields. Likewise, in the Büyük Menderes Basin, western neighbor of the Denizli Basin, the geothermal vapor samples have also high CO 2 volume, ranging from 93 to 99% (Karakuş and Şimşek, 2013). Helium isotope ratios have been used to define the origin of volatiles, crust-mantle relationships and 3He/4He and 3He/20Ne ratios of the thermal waters (Güleç et al., 2002; Mutlu et al., 2012). Güleç and Hilton (2006, 2016) suggested that the helium isotope compositions of the western Anatolian geothermal fields indicate mixing mantle and crustal helium components. Helium isotope data of CO2 from the Pamukkale Geothermal Field, southeastern neighbor of the GGF (Fig. 2) indicate no negligible mantle-derived helium (46%, Ercan et al., 1994). Similar data have been obtained in other geothermal fields in the Denizli Basin (12-26% for Kızıldere: Mutlu et al., 2008; Wiersberg et al., 2011 and 30-36% for Tekkehamam: Güleç, 1988; Wiersberg et al., 2011; Fig. 2). High 3He/4He ratios reflect the mantle derived He (e.g., Graham,

2002), while low 3He/4He ratios are the typical of the crustal He (e.g., Ballentine and Burnard, 2002). The highest mantle-CO2 (CO2/3He ratio= 12x109) and air-corrected R/Ra ratios (3.69) is recorded in the Pamukkale Geothermal Field (Fig. 2; Ercan et al., 1994), indicating high mantle values (Karakuş and Şimşek, 2013). On the other hand, the lowest mantle-CO2 (CO2/He3 ratio = 180x109-920x109) and aircorrected R/Ra ratios (0.955-2.050) is obtained from the Kızıldere Geothermal Field (Fig. 2; Wiersberg et al., 2011), suggesting low crustal values (Karakuş and Şimşek, 2013). This trend may indicate thermal water circulation system from east to west through the Denizli Basin. From this point of view, it can be considered that the volatiles in the GGF are mostly originated from the mantle. The δ13C values of gases of the geothermal fields of Pamukkale, Karahayıt, Yenice, Kızıldere and Tekkehamam (Fig. 2) range from -4.2 to -0.62‰ (Filiz, 1982, 1989; Ercan et al., 1994; Mutlu et al., 2008), which are typical of magmatic origin. This phenomenon has already been supported by similar δ13CCO2 values (-3.30 to -2.35‰, mean of -2.99‰) in the Gölemezli banded and bedded travertines (TUBITAK, 2016). Specifically in the western Anatolia, high CO2 reflects mantle origin based on δ13CCO2 and He isotope data (e.g., Filiz, 1984; Ercan et al., 1994; Güleç et al., 2002). The presence of high CO2 gas concentration (approximately 700-800 mg/L) in the GGF thermal waters confirms a mantle contribution. Tracing the origin of CO2 is crucial to obtain the geochemical properties of CO2-rich thermal waters, and for those stable carbon isotopes have been commonly used as a tracer (e.g., Becker et al., 2008). CO2 is produced by many processes in orogenic belts and high heat flow systems, and therefore it can be originated from four main sources having different, partially overlapping, carbon isotope signatures in geothermal systems: (i) thermo-metamorphic decarbonation of carbonates, forming CO2 with δ13C values ranging -2 to +2‰; (ii) dissolution of carbonate rocks with δ13C values around 0‰; (iii) organic processes (primary biogenic carbon) with values lower than -22‰ and (iv) upper-mantle degassing having values between -8 and -3‰ (e.g., Truesdell and Hulston, 1980; Evans et al., 2008; Becker et al., 2008; Shipton et al., 2004). In contrast to δ13C values of CO2, the δ13CDIC values of the GGF thermal waters have positive values ranging


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Fig. 17. Simplified geological section through Gölemezli field and conceptual model of the fluid circulation in the GGF. The location of the investigated thermal wells and springs are also indicated.

from +5.11 to +7.54‰, reflecting metamorphic CO2 (Fig. 10). Likewise, in the Gölemezli field, positive δ13CDIC ratios are also obtained as +3.11 to +5.07‰ for the banded and +5.46 to +6.14‰ for the bedded travertines (TUBITAK, 2016), +0.78 to +5.04‰ for carbonates of the Neogene formations (Alçiçek et al., 2007) and +0.7 to +2.9‰ (Kele et al., 2011) and +0.20 to +2.98‰ (Brilli et al., 2015) for the metamorphic bedrocks (Fig. 9F). In view of entire carbon isotope data from the gas, carbonates and thermal waters, the CO2 content in the Denizli Basin is thus probably to be produced from the combination of several processes. In the study field, evolution of isotopically heavy carbon isotopes of the thermal waters and travertines may be explained by four stages, leading to modify their original carbon isotopic signature (Fig. 17): (i) Stage 1: CO2 production by metamorphic decarbonation reactions, hydrothermal degradation of organic matter (Evans et al., 2008; Becker et al., 2008) and carbonate dissolution controlled by silicate hydrolysis (Shipton et al., 2004; Koh et al., 2008) processes at the deep metamorphic reservoir ca. 5 km of depth; (ii) Stage 2: Phase separation (immiscibility/exsolution) of metamorphic fluids during decompression and temperature loss (Becker et al., 2008); (iii) Stage 3: Isotopically heavy CO2 production by carbonate dissolution within the Neogene formations constituting the shallow reservoir and (iv) Stage 4: Degassing of mantle-derived and metamorphic CO2 from the groundwater during rise to the surface: During the first stage, a large volume of isotopically heavy CO2 derives mainly by progradation of metamorphic reactions and hydrothermal degradation of organic matter within the deep reservoir (Fig. 17). Therefore,

the positive δ13CDIC ratios of the GGF thermal waters are mostly attributed to a 13C-rich CO2 contribution released during thermo-metamorphic decarbonation and/or dissolution of the metamorphic bedrock. However, these processes cannot produce very heavy δ13C in the thermal waters due to the moderate δ13CDIC isotopic values (+0.7 to +2.9‰ ; Kele et al., 2011 and +0.20 to +2.98‰; Brilli et al., 2015) of the bedrock (Figs. 9F and 17). Ferry (1994) and Koh et al. (2008) suggested that metamorphic decarbonation reactions may occur above 300°C. The final isotope composition of dissolved inorganic carbon (DIC) is governed by isotope enrichment factors between free CO2 and carbonate species in the solution, because an enrichment of isotopic composition of DIC may occurs as a result of water-gas interaction (Mook et al., 1974). Karakuş and Şimşek (2013) reported that estimated proportions of carbon from the geothermal fields in the vicinity of GGF (Pamukkale, Tekkehamam and Kızıldere; Fig. 2) are 0.5–12.4% MORB, 78.6–97.5% limestone and 2–9% sediment components, indicating that the carbon sourced mainly from the marbles and calcschists of the Paleozoic metamorphics at the deeper reservoir. The similar result has been obtained by Mutlu et al. (2008), suggesting that the carbon derived mainly from crustal marine limestone (N 72%) with minor amount of mantle (~4%) and sediment (~1%) constituents from the bedrocks (gneiss, schists and marble of the Menderes Massif) throughout the western Anatolian basins. Wiersberg et al. (2011) documented that the gas samples from Kızıldere and Tekkehamam geothermal fields contain high amounts of methane (up to 0.55 vol.%) generated mostly by abiogenic processes including hydrothermal degradation

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Fig. 18. 3D schematic illustrations of the conceptual tectonic model of the GGF in the Denizli Basin.

of organic matter, with minor amount of biogenic processes (microbial and thermal degradation of organic matter). Cruse and Seewald (2006) suggested that the hydrothermal degradation can lead 13C-enrichment in organic matters. The second stage (Fig. 17) is characterized by phase separation (immiscibility) of the supercritical H2O-CO2 fluid due to decompression and loss of temperature (~300 °C) (Becker et al., 2008). During the third stage, carbonate dissolution in the Neogene formations allows to lead isotopically heavy δ13C DIC values in the thermal waters due to the high δ13CDIC isotopic values (+ 3.11 to + 5.07‰) of carbonates of the Kızılburun, Sazak and Kolankaya formations during carbon isotope exchange process (Figs. 9F and 17). Finally (four stage), the supersaturated deep metamorphic fluids reach to the surface and metamorphic CO2 discharge into the atmosphere (Fig. 17). However, at this stage, degassing of mantle deriven CO2 up to the surface also occurs as supported by helium isotope data (Mutlu et al., 2008; Wiersberg et al., 2011). Degassing of both metamorphic and mantle derived CO2 can be explained by two possible mechanisms: (i) magmatic volatiles (e.g., CO2, He) formerly trapped within subsurface rocks and thus metamorphic fluids, as reported by Koh et al. (2008) and therefore the process results in degassing of isotopically heavy CO2. Such notion is supported by the mantle originated fluid inclusions trapped within the travertine in the GGF (TUBITAK, 2016) and (ii) an open system degassing at temperatures below ~125 °C allows an enrichment of 13C isotopes within the thermal waters and travertine (Mook et al., 1974; Evans et al., 2008) due to separation of isotopically light CO2 (Evans et al., 2008). The significant CO2 degassing leads to oversaturation and travertine deposition in the GGF, notably in the Pamukkale, Karahayıt and Yenice geothermal fields (Fig. 2). Isotopically enriched δ13CDIC values of thermal waters accompanied with travertines also reported in many geothermal fields such as

Yellowstone (Friedman, 1970), Utah (Shipton et al., 2004), Central Nepal (Evans et al., 2008; Becker et al., 2008) and Italy (Pentecost, 1995). In western Anatolia, the presence of mantle He in the thermal waters indicates that the fracture network is deeper enough to allow the rising of helium from the deeper structural levels (e.g., Güleç et al., 2002). Mantle-derived helium has been observed to escape along areas of continental crust undergoing extension and recent volcanic activity (e.g., Kennedy and van Soest, 2006, 2007). Mantle 3He can be carried across the ductile continental crust either by direct intrusion and degassing of mantle-derived magmas (Ballentine et al., 2005) or moved into the roots of deep shear zones by “geo-pressured” mantle fluids concentrated along the ductile lower crust and then transferred to the shallow crustal level by fault-controlled fluid flow or by diffusion through the brittle-ductile boundary (Kennedy and van Soest, 2007). In addition, the active extensional deformation can also accelerate permeability, and therefore, mantle helium can diffuse the ductile zone in regions even where there is no significant magmatism (Kennedy and van Soest, 2007), as well as in the study area. In that case, degassing of mantle-helium occur in the Gölemezli field likely related to NW-trending Tripolis fault (with strong historical seismic events) associated with the NE-trending faults associated to the transfer zone (Figs. 2 and 18). Both heat and helium transfer is associated to the current extension in western Anatolia (Güleç et al., 2002; Güleç and Hilton, 2006). These mantle melts emplace at crustal levels with no surface (volcanic) outcrops (Güleç and Hilton, 2006), as well in GGF. Although there is no surface volcanism in the Denizli Basin (except to the southeast margin), the high temperature geothermal fields could be closely related with the young alkaline volcanics (Güleç and Hilton, 2006). The mechanism of the ascent and emplacement of mantle volatile at shallow depths is attributed to adiabatic mantle melting, followed by degassing of magma and mantle volatile input along deep rooting in the extensional tectonic


regimes (Ballentine et al., 2005; Kennedy and van Soest, 2006, 2007), as well as in the Denizli Basin (Wiersberg et al., 2011; Güleç et al., 2014). Consequently, the youngest Quaternary volcanism is closely linked to the current extensional tectonic activity and therefore, it serves as the source of heat for the GGF. 8. Conclusions The Gölemezli Geothermal Field is one of the largest geothermal field in the western Anatolia. Detailed hydrogeochemical properties of the GGF waters were investigated and concluded as follows:

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from thermo-metamorphic reactions in the deep and shallow reservoirs, whereas the cold waters (between − 10.38 and − 4.09‰) indicate groundwater DIC in origin. (8) This study indicates volatile ascent from magmatic sources through fault planes and fracture networks of the basin is the main mechanism to explain the presence of heat and mantle volatiles at shallow depths of the Denizli Basin. Thus, the GGF appears to be affected by the tectonism and active volcanism events, as well as in other geothermal fields in the Denizli Basin.

Acknowledgements (1) The thermal waters in the GGF are affected by steam-heated processes and classified into two chemical groups: (i) Group 1A, comprising Na-SO4 type and Ca-SO4 type which are fed by a steam-heated aquifer directly influenced by input of magma originated sulphur and/or oxidation of H2S to H2SO4 and (ii) Group 1B, only consisting Ca-HCO3 type water which reflects discharging from peripheral aquifer, due to capture of CO2-rich vapors in subsurface groundwater. Two distinct groups are also identified in the cold waters: (i) Group 2A, corresponding to Ca-HCO3 type and (ii) Group 2B, including Mg-HCO3 type, reflecting interaction of CO2 charged fluids at lower temperatures and migration path mixing with local groundwater. (2) The ions of the GGF thermal waters are mainly derived from the water-rock interaction among the metamorphic rocks of the Paleozoic formations (where the main and deep reservoir is hosted), Neogene formations (where the shallow reservoir is located) and thermal waters. High correlation in some ionic ratios and high contents of some minor elements in the studied thermal waters support enhanced water–rock interaction. (3) The stable oxygen and deuterium isotopic compositions of the GGF thermal waters indicate that the thermal waters originated from the meteoric waters. According to Giggenbach's method, all the GGF waters are far from the full equilibrium, confirming mixing with colder meteoric waters. Mixing processes between thermal and cold waters takes place during the upward migration of thermal waters. (4) As presented by both chemical geothermometries and temperatures obtained in drilling wells and springs, the GGF is a moderate-temperature geothermal field. Temperature of the GGF geothermal reservoir is encompassed between 130 and 210 °C. Because of the presence of mixing processes, this temperature range should be considered as the minimum estimated reservoir temperatures. According to, as a maximum, the average depth of the deep reservoir is about 5 km, which may allow us to reach the depth of the Precambrian gneiss, indicating a possible third reservoir, as well as in the Kızıldere Geothermal Field in the Denizli Basin. (5) The GGF thermal waters are oversaturated at discharge temperatures for carbonate (calcite) and silica minerals (chalcedony and quartz) allowing the enrichment in carbonate and silica, determining travertine/tufa precipitation in the discharge area. (6) A conceptual hydrogeochemical framework for the GGF has been established based on its hydrochemical and isotopic compositions. The recharge elevation of thermal waters is 780–840 m.a.s.l., located in the area of southern piedmont of Yenice Horst, as part of the regional flow system. The δ18O (ca. −8‰), δ2H (ca. −60‰) and δ3H (b 1 TU) values of Gölemezli thermal waters indicate their deep-circulating meteoric origin. These waters likely originated from the infiltration of rainwater through fracture network and faults to depth. Subsequent heating occurred by conduction. Consequently, the waters rise to the surface through faults and fractures that acted as hydrothermal pathways. (7) The high δ13CDIC values (between + 5.11 and + 7.54‰) reflect that the carbon in the thermal waters mainly originated

This paper has been the result of an international bilateral cooperation between TUBITAK (Scientific and Techonological Research Council of Turkey) and CNR (National Research Council of Italy): grant number 113Y551. We are indebted to Mr. Kenan Inan and Mr. Said Bozoğlu from Denizli Governorship for their valuable help and interest. The enthusiasm and vitality of our friend and colleague Marco Meccheri has been at the base of our common research activity carried out during three periods of field work (2013-2015). Marco prematurely passed away and we have just integrated his work in this paper, honoring his memory. We are grateful to Mr. Barbaros Ozcelik for having permitted us to visit his quarry. We thank to A.İ. Okay (ITU) for his helpful informations. We are also grateful to two anonymous reviewers for their useful comments and suggestions. MCA is indebted to the Turkish Academy of Sciences (TUBA) for a GEBIP (Young Scientist Award) grant. References Akan, S.B., Şimşek, Ş., 1997. Study of the gas discharges in the Denizli Pamukkale geothermal area (SW Turkey). Proceedings International Conference on Water Problems in the Mediterranean Countries. Near East University, Nicossia, North Cyprus. Akıllı, H., Bülbül, E., 2006. Denizli ili jeotermal kaynakları değerlendirme raporu. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 10874, Ankara, Turkey (in Turkish). Akkuş, İ., Akıllı, H., Ceyhan, S., Dilemre, A., Tekin, Z., 2005. Geothermal Inventory of Turkey. General Directorate of Mineral Research and Exploration, No. 201, Ankara, Turkey (in Turkish). Alçiçek, H., Wesselingh, F., Alçiçek, M.C., 2015. Paleoenvironmental evolution of the late Pliocene-early Pleistocene fluvio-deltaic sequence of the Denizli Basin (SW Turkey). Palaeogeogr. Palaeoclimatol. Palaeoecol. 437, 98–116. Alçiçek, H., 2007. Sedimentological Investigations of Neogene Deposits of the Denizli Basin (Sarayköy-Buldan Area, SW Turkey). (Ph.D. Thesis). Ankara Univ., Turkey (304 p). Alçiçek, H., Varol, B., Özkul, M., 2007. Sedimentary facies, depositional environments and palaeogeographic evolution of the Neogene Denizli Basin of SW Anatolia, Turkey. Sediment. Geol. 202, 596–637. Alçiçek, M.C., Brogi, A., Capezzuoli, E., Liotta, D., Meccheri, M., 2013. Superimposed basin formation during Neogene-Quaternary extension in SW-Anatolia (Turkey): insights from the kinematics of the Dinar fault zone. Tectonophysics 608, 713–727. Alçiçek, H., Bülbül, A., Alçiçek, M.C., 2016. Hydrogeochemistry of the thermal waters from the Yenice Geothermal Field (Denizli Basin, southwestern Turkey). J. Volcanol. Geotherm. Res. 309, 118–138. Altunel, E., 1994. Active Tectonics and the Evolution of Quaternary Travertines at Pamukkale, Western Turkey. (PhD thesis). University of Bristol, p. 236. Altunel, E., 1996. Pamukkale travertenlerinin morfolojik özellikleri, yaşları ve neotektonik önemleri. MTA Dergisi 118, 47-64. Altunel, E., Hancock, P.L., 1993a. Active fissuring, faulting and travertine deposition at Pamukkale (W Turkey). In: Stewart, I.S., Vita-Finzi, C., Owen, L.A. (Eds.), Neotectonics and Active Faulting, Zeitschrift fur Geomorfologie Supply. 94, pp. 285–302. Altunel, E., Hancock, P.L., 1993b. Morphological features and tectonic setting of Quaternary travertines at Pamukkale, western Turkey. Geol. J. 28, 335–346. Altunel, E., Hancock, P.L., 1996. Structural attributes of travertine-filled extensional fissures in the Pamukkale Plateau, Western Turkey. Int. Geol. Rev. 38, 768–777. Athanassas, C.D., Modis, K., Alçiçek, M.C., Theodorakopoulou, K., 2017. Contouring the cataclysm: a geographical analysis of the effects of the Minoan eruption. Environ. Archaeol. https://doi.org/10.1080/14614103.2017.1288885. Baba, A., Sözbilir, H., 2012. Source of arsenic based on geological and hydrogeochemical properties of geothermal systems in Western Turkey. Chem. Geol. 33, 364–377. Ballentine, C.J., Burnard, P.G., 2002. Production, release and transport of noble gases in the continental crust. In: Porcelli, D.R., Ballentine, C.J., Weiler, R. (Eds.), Noble Gases in Geochemistry and Cosmochemistry: Reviews in Mineralogy and Geochemistry. Washington D.C., Geochemical Society and Mineralogical Society of America, pp. 481–538. Ballentine, C.J., Marty, B., Lollar, B.S., Cassidy, M., 2005. Neon isotopes constrain convection and volatile origin in the Earth's mantle. Nature 433, 33–38.

28


Bargar, K.E., 1978. Geology and thermal history of Mammoth hot springs Yellowstone National Park, Wyorning. U.S. Geol. Surv. Bull. 1444, 55p. Becker, J.A., Bickle, M.J., Galy, A., Holland, T.J.B., 2008. Himalayan metamorphic CO2 fluxes: quantitative constraints from hydrothermal springs. Earth Planet. Sci. Lett. 265, 616–629. Bozkurt, E., 2001. Neotectonic of Turkey a synthesis. Geodin. Acta 14, 3–30. Bozkurt, E., 2003. Origin of NE-trending basins in western Turkey. Geodin. Acta 16, 61–81. Brilli, M., Giustini, F., Conte, A.M., Mercadal, P.L., Quarta, G., Plumed, H.R., Scardozzi, G., Belardi, G., 2015. Petrography, geochemistry and cathodoluminescence of ancient white marble from quarries in the southern Phrygia and northern Caria regions of Turkey: considerations on provenance discrimination. J. Archaeol. Sci. Rep. 4, 124–142. Brogi, A., Capezzuoli, E., 2009. Travertine deposition and faulting: the fault-related travertine fissure-ridge at Terme S. Giovanni, Rapolano Terme (Italy). Int. J. Earth Sci. (Geol. Rundsch.) 98, 931–947. Brogi, A., Capezzuoli, E., Alçiçek, M.C., Gandin, A., 2014. Evolution of a fault-controlled travertine fissure-ridge in the western Anatolia extensional province: the Çukurbağ fissure-ridge (Pamukkale, Turkey). J. Geol. Soc. 171, 425–441. Brogi, A., Alçiçek, M.C., Yalçıner, C.Ç., Capezzuoli, E., Liotta, D., Meccheri, M., Rimondi, V., Ruggieri, G., Gandin, A., Boschi, C., Büyüksaraç, A., Alçiçek, H., Bülbül, A., Baykara, M.O., Shen, C.-C., 2016. Hydrothermal fluids circulation and travertine deposition in an active tectonic setting: insights from the Kamara geothermal area (western Anatolia, Turkey). Tectonophysics 680, 211–232. Çağlar, K.Ö., 1961. Türkiye maden suları ve kaplıcaları MTA Enstitüsü Yayını, No: 107, Fasikül 4, Ankara (in Turkish). Çakır, Z., 1999. Along-strike discontinuity of active normal faults and its influence on Quaternary travertine deposition: examples from western Turkey. Turk. J. Earth Sci. 8, 67–80. Capezzuoli, E., Gandin, A., Pedley, H.M., 2014. Decoding tufa and travertine (freshwater carbonates) in the sedimentary record: the state of the art. Sedimentology 61, 1–21. Cermac, V., Hurtig, E., 1979. Heat flow map of Europe. In: Cermac, V., Rybach, L. (Eds.), Terrestrial Heat Flow in Europe. Springer-Verlag, Berlin. Chiodini, G., Frondini, F., Ponziani, F., 1995. Deep structures and carbon dioxide degassing in central Italy. Geothermics 24, 81–94. Cinti, D., Procesi, M., Tassi, F., Montegrossi, G., Sciarra, A., Vaselli, O., Quattrocchi, F., 2011. Fluid geochemistry and geothermometry in the western sector of the Sabatini Volcanic District and the Tolfa Mountains (Central Italy). Chem. Geol. 284, 160–181. Clark, I.D., Fritz, P., 1997. Environmental Isotopes in Hydrogeology. Lewis Publishers, New York. Clark, W.B., Jenkins, W.J., Top, Z., 1997. Determination of tritium by mass spectrometric measurements. J. Appl. Radioact. Isot. 27, 515–522. Craig, H., 1961. Isotopic variation in meteoric waters. Science 133, 1702–1703. Cruse, A.M., Seewald, J.S., 2006. Geochemistry of low-molecular weight hydrocarbons in hydrothermal fluids from Middle Valley, northern Juan de Fuca Ridge. Geochim. Cosmochim. Acta 70, 2073–2092. Das, R., Das, S.N., Misra, V.N., 2005. Chemical composition of rainwater and dustfall at Bhubaneswar in the east coast of India. Atmos. Environ. 39, 5908–5916. De Filippis, L., Faccenna, C., Billi, A., Anzalone, E., Brilli, M., Soligo, M., Tucccimei, P., 2013. Plateau versus fissure ridge travertines from Quaternary geothermal springs of Italy and Turkey: interactions and feedbacks between fluid discharge, paleoclimate, and tectonics. Earth-Sci. Rev. 123, 35–52. Demirel, V., Kahraman, S., 2003. Denizli Belediyesi adına yapılan Denizli-Gölemezli (DG-3, DG-4, DG-5) kuyularına ait kuyu bitirme raporu. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 10660, Ankara, Turkey (in Turkish). Deutsch, W.J., 1997. Groundwater Geochemistry: Fundamentals and Application to Contamination. Lewis publisher, USA. Dilsiz, C., 2006. Conceptual hydrodynamic model of the Pamukkale hydrothermal field, southwestern Turkey, based on hydrochemical and isotopic data. Hydrogeol. J. 14, 562–572. Doglioni, C., Agostini, S., Crespi, M., Innocenti, F., Manetti, P., Riguzzi, F., Savasçın, Y., 2002. On the extension in western Anatolia and the Aegean Sea. J. Virtual Explor. 7, 167–181. Dumont, J.F., Uysal, Ş., Şimşek, Ş., Karamanderesi, İ.H., Letouzey, J., 1979. Formation of the grabens in southwestern Turkey. Bull. Miner. Res. Expl. Direct. Turkey (MTA) 92, 7–18. Ellis, A.J., Mahon, W.A.J., 1977. Chemistry and Geothermal Systems. 392. Academic Press. Ercan, T., 1982. Batı Anadolu'nun Genç Tektoniği ve Volkanizması. Batı Anadolu'nun Genç Tektoniği ve Volkanizması Paneli, TJK Yayını. pp. 5–14. Ercan, T., Günay, E., Baş, H., 1983. Denizli volkanitlerinin petrolojisi ve plaka tektoniği açısından bölgesel yorumu. Bull. Geol. Soc. Turk. 26, 153–160 (inTurkish). Ercan, T., Satır, M., Kreuzer, H., Türkecan, A., Günay, E., Çevikbaş, A., Ateş, M., Can, B., 1985. Batı Anadolu Senozoyik volkanitlerine ait yeni kimyasal, izotopik ve radyometrik verilerin yorumu. Türkiye Jeol. Kur. Bült. 28 pp. 121–136. Ercan, T., Matsuda, J., Nagao, K., Kita, I., 1994. Anadolu'daki sıcak sularda bulunan doğal suların izotopik bileşimleri ve karbondioksit gazının enerji açısından önemi. Türkiye 6. Enerji Kongresi, İzmir, Türkiye. pp. 197–207. Ersoy, Y., Helvacı, C., 2007. Stratigraphy and geochemical features of the Early Miocene bimodal (ultrapotassic and calcalkaline) volcanic activity within the NE-trending Selendi basin, western Anatolia, Turkey. Turk. J. Earth Sci. 16, 117–139. Evans, M.J., Derry, L.A., France-Lanord, C., 2008. Degassing of metamorphic carbon dioxide from the Nepal Himalaya. Geochem. Geophys. Geosyst. 9, Q04021. https://doi.org/ 10.1029/2007GC001796. Faccenna, C., Soligo, M., Billi, A., De Filippis, L., Funiciello, R., Rossetti, C., Tuccimei, P., 2008. Late Pleistocene depositional cycles of the Lapis Tiburtinus travertine (Tivoli, central Italy): possible influence of climate and fault activity. Glob. Planet. Chang. 63, 299–308.

Faccenna, C., Becker, T.W., Auer, L., Billi, A., Boschi, L., Brun, J.P., Capitanio, F.A., Funiciello, F., Horvath, F., Jolivet, L., Piromallo, C., Royden, L., Rossetti, F., Serpelloni, E., 2014. Mantle dynamics in the Mediterranean. Rev. Geophys. 52, 283–332. Ferry, J.M., 1994. Overview of the petrologic record of fluid flow during regional metamorphism in northern New England. Am. J. Sci. 294, 905–988. Filiz, Ş., 1982. Ege bölgesindeki önemli jeotermal alanların O18, H2, H3 ve C13 izotoplarıyla incelenmesi. Ege Ü., Yer Bil. Fak., Doçentlik Tezi, İzmir, Turkey (in Turkish). Filiz, Ş., 1984. Investigation of the important geothermal areas by using C, H, O, isotopes. Seminar on the utilization of Geothermal Energy for Electric Power Generation and Space Heating, 14–17 May 1984, Florance, Italy. Seminar ref. No: EP/ SEM. 9/R.3. Filiz, Ş., 1989. Isotopic Analyses of CO2 and Travertines in the Denizli Geothermal Province (West Anatolia). U.N. Seminar on New Developments in Geothermal Energy. 22–25 May 1989, Ankara. Filiz, Ş., Gökgöz, A., Tarcan, G., 1992. Hydrogeologic Comparisons of Geothermal Fields in the Gediz and Büyük Menderes Grabens. XI. Congress of World Hydrothermal Organization, İstanbul. Fournier, R.O., 1977. Chemical geothermometers and mixing models for geothermal systems. Geothermics 5, 41–50. Fournier, R.O., 1979. Geochemical and hydrologic considerations and the use of enthalpychloride diagrams in the prediction of underground conditions in hot-spring systems. J. Volcanol. Geotherm. Res. 5, 1–16. Fournier, R.O., 1991. Water geothermometers applied to geothermal energy. In: D'amore, F. (Ed.), Application of Geochemistry in Geothermal Reservoir Development. UNITAR/ UNDP Publications, Rome, pp. 37–69. Fournier, R.O., Truesdell, A.H., 1973. An empirical Na-K-Ca geothermometer for natural waters. Geochim. Cosmochim. Acta 37, 1255–1275. Francalanci, L., Innocenti, F., Manetti, P., Savasçın, M.Y., 2000. Neogene alkaline volcanism of the Afyon-Isparta area, Turkey: petrogenesis and geodynamic implications. Mineral. Petrol. 70, 285–312. Friedman, I., 1970. Some investigations of the deposition of travertine from Hot Springs-I. The isotopic chemistry of a travertine-depositing spring. Geochim. Cosmochim. Acta 34, 1303–1315. Gandin, A., Capezzuoli, E., 2008. Travertine versus Calcareous tufa: distinctive petrologic features and stable isotopes signatures. Ital. J. Quat. Sci. 21, 125–136. Gat, J.R., Carmi, I., 1970. Evolution in the isotopic composition of atmospheric waters in the Mediterranean Sea area. J. Geophys. Res. 75, 3039–3048. Gemici, Ü., Tarcan, G., 2002. Hydrogeochemistry of the Simav geothermal field, western Anatolia, Turkey. J. Volcanol. Geotherm. Res. 116, 215–233. Giggenbach, W.F., 1988. Geothermal solute equilibria - derivation of Na-K-Mg-Ca geoindicators. Geochim. Cosmochim. Acta 52, 2749–2765. Giggenbach, W.F., Gonfiantini, R., Jangi, B.L., Truesdell, A.H., 1983. Isotopic and chemical composition of Parbati valley geothermal discharges, NW-Himalaya, India. Geothermics 12, 199–222. Giggenbach, W.F., Sheppard, D.S., Robinson, B.W., Stewart, M.K., Lyon, G.L., 1994. Geochemical structure and position of the Waiotapu geothermal field, New Zealand. Geothermics 23, 599–644. Göb, S., Loges, A., Nolde, N., Bau, M., Jacob, D.E., Markl, G., 2013. Major and trace element compositions (including REE) of mineral, thermal, mine and surface waters in SW Germany and implications for water–rock interaction. Appl. Geochem. 33, 127–152. Gökalp, E., 1971. Denizli vilayeti, Yenice-Gölemezli-Karahayıt kaplıcaları çevresi jeolojik etüdleri ve jeotermik enerji imkanları. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 4571, Ankara, Turkey (in Turkish). Gökgöz, A., 1994. Pamukkale-Karahayıt-Gölemezli Hidrotermal Karstının Hidrojeolojisi. Süleyman Demirel Univ. (PhD Thesis, Isparta, Turkey, in Turkish). Gökgöz, A., 1998. Geochemistry of the Kızıldere-Tekkehamam-Buldan-Pamukkale Geothermal Fields. UNU, Geothermal Training Programme, Reports. pp. 115–156 (Reykjavik, Iceland). Gonfiantini, R., 1986. Environmental isotopes in lake studies. In: Fritz, P., Fontes, J.Ch. (Eds.), Handbook of Environmental Isotope Geochemistry 2. Elsevier, New York, pp. 113–168. Graham, D.W., 2002. Noble gas isotope geochemistry of mid-ocean ridge and ocean island basalts; characterization of mantle source reservoirs. In: Porcelli, D., Wieler, R., Ballentine, C. (Eds.), Noble Gases in Geochemistry and Cosmochemistry, Reviews in Mineralogy and Geochemistry. Mineral. Soc. Amer., Washington, D.C., pp. 247–318. Guidi, M., Marini, L., Principe, C., 1990. Hydrogeochemistry of Kızıldere geothermal system and nearby region. Geotherm. Resour. Counc. Trans. 14, 901–908. Güleç, N., 1988. He-3 distribution in western Turkey. Bull. Miner. Res. Explor. Inst. Turk. 108, 35–42. Güleç, N., Hilton, D.R., 2006. Helium and heat distribution in Western Anatolia, Turkey: relationship to active extension and volcanism. In: Dilek, Y., Pavlides, S. (Eds.), Post-Collisional Tectonics and Magmatism in the Eastern Mediterranean Region. Geological Society of America, GSA Special Paper 409, pp. 305–319. Güleç, N., Hilton, D.R., 2016. Turkish geothermal fields as natural analogues of CO2 storage sites: gas geochemistry and implications for CO2 trapping mechanisms. Geothermics 64, 96–110. Güleç, N., Hilton, D.R., Mutlu, H., 2002. Helium isotope variations in Turkey: relationship to tectonics, volcanism and recent seismic activities. Chem. Geol. 187, 129–142. Güleç, N., Mutlu, H., Hilton, D.R., 2014. Gas geochemistry of Turkish geothermal fluids: He-CO 2 systematics in relation to active tectonics and volcanism. In: Baba, A., Bundschuh, J., Chandrasekharam, D. (Eds.), Geothermal Systems and Energy Resources: Turkey and Greece. Taylor and Francis, The Netherlands, pp. 13–23.

H. Alçiçek et al. / Journal of Volcanology and Geothermal Research 349 (2018) 1–30 Gündoğan, İ., Helvacı, C., Sözbilir, H., 2008. Gypsiferous carbonates at Honaz Dağı (Denizli): first documentation of Triassic gypsum in western Turkey and its tectonic significance. J. Asian Earth Sci. 32, 49–65. Han, D.N., Liang, X., Jin, M.G., Currell, M.J., Song, X.F., Liu, C.M., 2010. Evolution of groundwater hydrochemical characteristic and mixing behavior in the Daying and Qicum sgeothermal systems, Xinzhou Basin. J. Volcanol. Geotherm. Res. 189, 99–104. Hancock, P.L., Chalmers, R.M.L., Altunel, E., Çakır, Z., 1999. Travitonics: using travertines in active fault studies. J. Struct. Geol. 21, 903–916. Hilton, D.R., Fischer, T.P., Marty, B., 2002. Noble gases in subduction zones and volatile recycling. Rev. Mineral. Geochem. 47, 319–370. van Hinsbergen, D.J.J., 2010. A key extensional metamorphic complex reviewed and restored: the Menderes Massif of western Turkey. Earth-Sci. Rev. 102, 60–76. van Hinsbergen, D.J.J., Schmid, S.M., 2012. Map-view restoration of Aegean-west Anatolian accretion and extension since the Eocene. Tectonics 31, TC5005. Hoefs, J., 2009. Stable Isotope Geochemistry. 6th edition. Springer, Berlin. IAEA, 1981. Stable isotope hydrology. Deuterium and oxygen-18 in water cycle. In: Gat, J.R., Gonfiantini, R. (Eds.), International Atomic Energy Agency Technical Report No. 210, Vienna, p. 339. İlkışık, O.M., 1995. Regional heat flow in western Anatolia using silica temperature estimates from thermal springs. Tectonophysics 244, 175–184. Innocenti, F., Agostini, S., Divincenzo, G., Doglioni, C., Manetti, P., Savaşçın, M.Y., Tonarini, S., 2005. Neogene and Quaternary volcanism in Western Anatolia: magma sources and geodynamic evolution. Mar. Geol. 221, 397–421. Jolivet, J., Faccenna, C., Huet, B., Labrousse, L., Le Pourhiet, L., Lacombe, O., Lecomte, E., Burov, E., Denèle, Y., Brun, J.P., Philippon, M., Paul, P., Salaün, G., Karabulut, H., Piromallo, C., Monié, P., Gueydan, F., Okay, A.I., Oberhänsli, R., Pourteau, A., Augier, R., Gadenne, L., Driussi, O., 2013. Aegean tectonics: strain localisation, slab tearing and trench retreat. Tectonophysics 597-598, 1–33. Karakuş, H., Şimşek, Ş., 2013. Tracing deep thermal water circulation systems in the E–W trending Büyük Menderes Graben, western Turkey. J. Volcanol. Geotherm. Res. 252, 38–52. Karaoğlu, Ö., Helvacı, C., 2014. Isotopic evidence for a transition from subduction to slabtear related volcanism in western Anatolia, Turkey. Lithos 192-195, 226–239. Kaymakçı, N., 2006. Kinematic development and paleostress analysis of the Denizli Basin (Western Turkish): implications of spatial varition of relative paleostress magnitudes and orientations. J. Asian Earth Sci. 27, 207–222. Kele, S., Özkul, M., Gökgöz, A., Fórizs, I., Baykara, M.O., Alçiçek, M.C., Németh, T., 2011. Stable isotope geochemical and facies study of Pamukkale travertines: new evidences of low-temperature non-equilibrium calcite-water fractionation. Sediment. Geol. 238, 191–212. Kennedy, B.M., van Soest, M.C., 2006. A helium isotope perspective on the Dixie Valley, Nevada hydrothermal system. Geothermics 35, 26–43. Kennedy, B.M., van Soest, M.C., 2007. Flow of mantle fluids through the ductile lower crust: helium isotope trends. Science 318, 1433–1436. Kharaka, Y.K., Mariner, R.H., 1988. Chemical geothermometers and their application to formation waters from sedimentary basins. In: Naeser, N.D., McCulloh, T. (Eds.), Thermal History of Sedimentary Basins. Springer Verlag, New York, pp. 99–117. Khatib, S., Rochette, P., Alçiçek, M.C., Lebatard, A.-E., Demory, F., Saos, T., 2014. Etude stratigraphique, sédimentologique et paléomagnétique des travertins de Denizli (Turquie) contenant des restes fossiles quaternaires (Stratigraphic, sedimentological and paleomagnetic study of the Kocabaş travertines, Denizli Basin, Anatolia, Turkey). Anthropologie 118, 16–33. Kıymaz, İ., 2012. Karahayıt (Denizli) yöresinin jeotermal potansiyeli. Süleyman Demirel Univ (MSc Thesis, Isparta, Turkey, in Turkish). Koçyiğit, A., 2005. The Denizli graben-horst system and the eastern limit of western Anatolian continental extension: basin-fill, structure, deformational mode, throw amount and episodic evolutionary history, SW Turkey. Geodin. Acta 18, 167–208. Koh, Y.-K., Choi, B.-Y., Yun, S.-K., Choi, H.-S., Mayer, B., Ryo, S.-W., 2008. Origin and evolution of two contrasting thermal groundwaters (CO2-rich and alkaline) in the Jungwon area, South Korea: hydrochemical and isotopic evidence. J. Volcanol. Geotherm. Res. 178, 777–786. Konak, N., Şenel, M., 2002. Geological Map of Turkey in 1/500.000 Scale: Denizli Sheet. Publication of Mineral Research and Exploration Directorate of Turkey (MTA), Ankara. Koralay, T., Kılınçarslan, S., 2015. Minero-petrographic and isotopic characterization of two antique marble quarries in the Denizli region (western Anatolia, Turkey). Period. Mineral. 84, 263–288. Lebatard, A.-E., Alçiçek, M.C., Rochette, P., Khatib, S., Vialet, A., Boulbes, N., Bourlès, D.L., Demory, F., Guipert, G., Mayda, S., Titov, V.V., Vidal, L., de Lumley, H., 2014a. Dating the Homo erectus bearing travertine from Kocabaş (Denizli, Turkey) at least 1.1 Ma. Earth and Planet. Sci. Lett. 390, 8–18. Lebatard, A.-E., Bourlès, D.L., Alçiçek, M.C., 2014b. Datation des travertins de Kocabaş par la méthode des nucléides cosmogéniques 26Al/10Be (Dating of the Kocabaş travertines with the 26Al/10Be cosmogenic nuclide method). Anthropologie 118, 34–43. Levet, S., Toutain, J.P., Munoz, M., Berger, G., Negrel, P., Jendrzejewski, N., Agrinier, P., Sortino, F., 2002. Geochemistry of the Bagne'res-de-Bigorre thermal waters from the North Pyrenean Zone (France). Geofluids 2, 25–40. Ma, R., Wang, Y., Sun, Z., Zheng, C., Ma, T., Prommer, H., 2011. Geochemical evolution of groundwater in carbonate aquifers in Taiyuan, northern China. Appl. Geochem. 26, 884–897. Majumdar, N., Mukherjee, A.L., Majumdar, R.K., 2009. Mixing hydrology and chemical equilibria in Bakreswar geothermal area, Eastern India. J. Volcanol. Geotherm. Res. 183, 201–212. McCollom, T.M., Lollar, B.S., Lacrampe-couloume, G., Seewald, J.S., 2010. The influence of carbon source on abiotic organic synthesis and carbon isotope fractionation under hydrothermal conditions. Geochim. Cosmochim. Acta 74, 2717–2740.

29

McKenzie, D.P., 1978. Active tectonics of the Alpine-Himalayan belt: the Aegean Sea and surrounding regions (tectonics of the Alpine region). Geophys. J. R. Astron. Soc. 55, 217–245. Menant, A., Sternai, P., Jolivet, L., Guillou-Frottier, L., Gerya, T., 2016. 3D numerical modeling of mantle flow, crustal dynamics and magma genesis associated with slab rollback and tearing: the eastern Mediterranean case. Earth Planet. Sci. Lett. 442, 93–107. Michard, G., 1990. Behaviour of the major elements and some trace elements (Li, Rb, Cs, Sr, Fe, Mn, W, F) in deep hot waters from granitic areas. Chem. Geol. 89, 117–134. Möller, P., 2000. Rare earth elements and yttrium as geochemical indicators of the source of mineral and thermal waters. In: Stober, I., Bucher, K. (Eds.), Hydrology of Crystalline Rocks. Kluwer Acad. Press, pp. 227–246. Möller, P., Dulski, P., Savasçın, Y., Conrad, M., 2004. Rare earth elements, yttrium and Pb isotope ratios in thermal spring and well waters of West Anatolia, Turkey: a hydrochemical study of their origin. Chem. Geol. 206, 97–118. Möller, P., Dulski, P., Özgür, N., 2008. Partitioning of rare earths and some major elements in the Kızıldere geothermal field, Turkey. Geothermics 37, 132–156. Mook, W.G., Tan, F.C., 1991. Stable carbon isotopes in rivers and estuaries. In: Degens, E.T., Kempe, S., Richey, J.E. (Eds.), Biogeochemistry of Major World Rivers, Scope Report 42. Wiley, New York, pp. 245–263. Mook, W.G., Bommerson, J.C., Staverman, W.H., 1974. Carbon isotope fractionation between dissolved bicarbonate and gaseous carbon dioxide. Earth Planet. Sci. Lett. 22, 169–176. Motyka, R., Nye, C., Turner, D., Liss, S., 1993. The geyser Bight geothermal area, Umnak Island, Alaska. Geothermics 22, 301–327. Müller, G., 1967. Diagenesis in argillaceous sediments. In: Larsen, G., Chilingar, G.V. (Eds.), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 127–177. Mutlu, H., Güleç, N., 1998. Hydrogeochemical outline of thermal waters and geothermometry applications in Anatolia, Turkey. J. Volcanol. Geotherm. Res. 85, 495–515. Mutlu, H., Güleç, N., Hilton, D.R., 2008. Helium-carbon relationships in geothermal fluids of western Anatolia, Turkey. Chem. Geol. 247, 305–321. Mutlu, H., Güleç, N., Hilton, D.R., Aydın, H., Halldorsson, S.A., 2012. Spatial variations in gas and stable isotope compositions of thermal fluids around Lake Van: implications for crust–mantle dynamics in eastern Turkey. Chem. Geol. 300, 165–176. Nicholson, K., 1993. Geothermal Fluids: Chemistry and Exploration Techniques. SpringerVerlag, Berlin, Heidelberg, New York. Okay, A.İ., 1989. Denizli'nin güneyinde Menderes masifi ve Likya naplarının jeolojisi. Bull. Miner. Res. Expl. Direct. Turkey (MTA) 109, 45–58. Okay, A.İ., 2001. Stratigraphic and metamorphic inversions in the central Menderes Massif: a new structural model. Int. J. Earth Sci. 89, 709–727. Özgür, N., 2002. Geochemical signature of the Kızıldere geothermal field, Western Anatolia, Turkey. Int. Geol. Rev. 44, 153–163. Özkul, M., Kele, S., Gökgöz, A., Shen, C.-C., Jones, B., Baykara, M.O., Fόrizs, I., Németh, T., Chang, Y., Alçiçek, M.C., 2013. Comparison of the Quaternary travertine sites in the Denizli extensional basin based on their depositional and geochemical data. Sediment. Geol. 294, 179–204. Pamir, H.N., Erentöz, C., 1974. Geological Maps of Turkey in 1:500.000 Scale: Denizli Sheet. Mineral Res. Expl. Direct. Turkey (MTA), Ankara, Turkey. Panichi, C., Gonfiantini, R., 1978. Environmental isotopes in geothermal studies. Geothermics 6, 143–161. Parkhurst, D.L., Appelo, C.A.J., 1999. User's Guide to PHREEQC (Version 2) - A Computer Program for Speciation, Batch-Reaction, One-Dimensional Transport, and Inverse Geochemical Calculations: U.S. Geological Survey Water-Resources Investigations Report 99-4259. p. 312. Pastorelli, S., Marini, L., Hunziker, J.C., 1999. Water chemistry and isotope composition of the Acquarossa thermal system, Ticino, Switzerland. Geothermics 28, 75–93. Paton, S.M., 1992. The relationship between extension and volcanism in western Turkey, the Aegean Sea, and central Greece. Unpubl. PhD Thesis, Cambridge University, p. 300. Pentecost, A., 1995. Geochemistry of carbon dioxide in six travertine-depositing waters of Italy. J. Hydrol. 167, 263–278. Pentecost, A., 2005. Travertine. Springer-Verlag, Berlin. Prelević, D., Akal, C., Romer, R.L., Mertz-Kraus, R., Helvacı, C., 2015. Magmatic response to slab tearing: constraints from the Afyon alkaline volcanic complex, Western Turkey. J. Petrol. 56, 527–562. Ravikumar, P., Somashekar, R.K., 2011. Environmental tritium (3H) and hydrochemical investigations to evaluate groundwater in Varahi and Markedenya river basins, Karnataka, India. J. Environ. Radioact. 102, 153–162. Reed, M.H., Spycher, N.F., 1984. Calculation of pH and mineral equilibria in hydrothermal waters with application to geothermometry and studies of boiling and dilution. Geochim. Cosmochim. Acta 48, 1479–1492. Rihs, S., Condomines, M., Poidevin, J.L., 2000. Long-term behaviour of continental hydrothermal systems: U-series study of hydrothermal carbonates from the French Massif Central (Allier Valley). Geochim. Cosmochim. Acta 64, 3189–3199. Şahinci, A., 1991. Doğal Suların Jeokimyası. Reform Matbaası, İzmir, Türkiye (in Turkish). Saraç, G., 2003. Türkiye Omurgalı Fosil Yatakları. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 10609, Ankara, Turkey (in Turkish). Schoeller, H., 1934. Les echanges de bases dans les eaux souterraines; trois exemples es Tunisie. Bull. Soc. Geol. Fr. 4, 389–420. Semiz, B., Çoban, H., Roden, M.F., Özpınar, Y., Flower, M.F.J., McGregor, H., 2012. Mineral composition in cognate inclusions in Late Miocene-Early Pliocene potassic lamprophyres with affinities to lamproites from the Denizli region, Western Anatolia, Turkey: implications for uppermost mantle processes in a back-arc setting. Lithos 134-135, 253–272. Şengör, A.M.C., Yılmaz, Y., 1981. Tethyan evolution of Turkey: a plate tectonic approach. Tectonophysics 75, 181–241.

30


Şengör, A.M.C., Satır, M., Akkök, R., 1984. Timing of tectonic events in the Menderes Massif, western Turkey: implications for tectonic evolution and evidence for Pan-African basement in Turkey. Tectonics 3, 693–707. Şengün, R., 2011. Yenicekent, Gölemezli ve çevresi hidrojeolojisi. (BSc Thesis). Dokuz Eylül University, İzmir, Turkey (in Turkish). Shipton, Z.K., Evans, J.P., Kirschner, D., Kolesar, P.T., Williams, A.P., Heath, J., 2004. Analysis of CO2 leakage through ‘low-permeability’ faults from natural reservoirs in the Colorado Plateau, East-central Utah. Geol. Soc. Lond., Spec. Publ. 233, 43–58. Şimşek, Ş., 1981. Aydın (Germencik) alanının jeotermal enerji olanakları: Yeni ve Yenilenebilir Enerji Simpozyumu Yayını 5, Ankara, Türkiye (in Turkish). Şimşek, Ş., 1984. Denizli, Kızıldere-Tekkehamam-Tosunlar-Buldan-Yenice alanının jeolojisi ve jeotermal enerji olanakları. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 7846, Ankara, Turkey (in Turkish). Şimşek, Ş., 1985. Geothermal model of Denizli, Sarayköy-Buldan area. Geothermics 14, 393–417. Şimşek, Ş., 1990. Karstic hot water aquifers in Turkey. In: Günay, G., Johnson, I., Back, W. (Eds.), Proceedings International Symposium on Hydrogeological Processes in Karst Terranes. IAHS Publication 207, pp. 173–184. Şimşek, Ş., 2003a. Hydrogeological and isotopic survey of geothermal fields in the Büyük Menderes graben, Turkey. Geothermics 32, 669–678. Şimşek, Ş., 2003b. Present Status and Future Development Possibilities of Aydın-Denizli Geothermal Province. International Geothermal Conference, Reykjavík, Session 5, Paper. 034 pp. 11–16. Şimşek, Ş., Yıldırım, N., Gülgör, A., 2005. Developmental and environmental effects of the Kızıldere geothermal power project. Geothermics 34, 239–256. Stüben, D., Berner, Z., Chandrasekharam, D., 2003. Arsenic enrichment in groundwater of West Bengal, India: geochemical evidence for mobilization of As under reducing conditions. Appl. Geochem. 18, 1417–1434. Sulpizio, R., Alçiçek, M.C., Zanchetta, G., Solari, L., 2013. Recognition of the Minoan tephra in the Acıgöl Basin, western Turkey: implications for inter archive correlations and fine ash dispersal. J. Quat. Sci. 28, 329–335. Sun, S., 1990. Denizli-Uşak Arasının Jeolojisi ve Linyit Olanakları. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 9985, Ankara, Turkey (in Turkish). Tamgaç, Ö.F., Yıldırım, N., Çetiner, H.L., 1995. Denizli-Karahayıt-Pamukkale ve çevresi sıcak su kaynaklarının korunması ve geliştirilmesine ait hidrojeoloji etüt raporu. Mineral Res. Expl. Direct. Turkey (MTA), Scientific Report No: 9942, Ankara, Turkey (in Turkish). Tarcan, G., 2004. Mineral saturation and scaling tendencies of waters discharged from wells (N150 °C) in geothermal areas of Turkey. J. Volcanol. Geotherm. Res. 142, 263–283.

Tarcan, G., Gemici, Ü., 2003. Water geochemistry of the Seferihisar geothermal area, İzmir, Turkey. J. Volcanol. Geotherm. Res. 126, 225–242. Tarcan, G., Özen, T., Gemici, Ü., Çolak, M., Karamanderesi, İ.H., 2016. Geochemical assessment of mineral scaling in Kızıldere geothermal field, Turkey. Environ. Earth Sci. 75, 1317. Ten Veen, J.H., Boulton, S.J., Alçiçek, M.C., 2009. From palaeotectonics to neotectonics in the Neotethys realm: The importance of kinematic decoupling and inherited structural grain in SW Anatolia (Turkey). Tectonophysics 473, 261–281. Tole, M.P., Armansson, H., Pang, Z., Arnorsson, S., 1993. Fluid/mineral equilibrium calculations for geothermal fluids and chemical geothermometry. Geothermics 22, 17–37. Truesdell, A.H., Hulston, J.R., 1980. Isotopic evidence on environments of geothermal systems. In: Fritz, P., Fontes, J.C. (Eds.), Handbook of Environmental Isotope Chemistry. 1. The Terrestrial Environment, Elsevier, pp. 179–226. Trumbore, S.E., Druffel, E.R.M., 1995. Carbon isotopes for characterizing sources and turnover of nonliving organic matter. In: Zepp, R.G., Sonntag, C. (Eds.), The Role of Nonliving Organic Matter in the Earth's Carbon Cycle. John Wiley & Sons, Chichester, pp. 7–22. TUBITAK, 2016. Integrated Approach to Characterize the Structural Setting and the Related Fossil Geothermal System Along the Northeastern Boundary of the Quaternary Denizli Basin. Scientific Project Report of International Bilateral Cooperation With CNR (Grant Number of 113Y551) (130 pp., Ankara, In Turkish with English abstract). Uysal, I.T., Feng, Y., Zhao, J.X., Altunel, E., Weatherley, D., Karabacak, V., Cengiz, O., Golding, S.D., Lawrence, M.G., Collerson, K.D., 2007. U-series dating and geochemical tracing of late Quaternary travertine in coseismic fissures. Earth Planet. Sci. Lett. 257, 450–462. Uysal, I.T., Feng, Y., Zhao, J.X., Işık, V., Nuriel, P., Golding, S.D., 2009. Hydrothermal CO2 degassing in seismically active zones during the Late Quaternary. Chem. Geol. 265, 442–454. Vengosh, A., Helvacı, C., Karamanderesi, İ.H., 2002. Geochemical constraints for the origin of thermal waters from western Turkey. Appl. Geochem. 17, 163–183. Wiersberg, T., Süer, S., Güleç, N., Erzinger, J., Parlaktuna, M., 2011. Noble gas isotopes and the chemical composition of geothermal gases from the eastern part of the Büyük Menderes Graben (Turkey). J. Volcanol. Geotherm. Res. 208, 112–121. Yaman, D., 2005. Menderes Masifi kıtasal rift zonlarında yer alan jeotermal sulardaki yüksek bor değerlerinin kökeni. Süleyman Demirel Univ. (PhD Thesis, Isparta, Turkey, in Turkish). Yıldırım, N., Güner, İ.N., 2002. Büyük Menderes grabeninin doğusunda yeralan jeotermal sahalarda bulunan suların izotopik ve hidrojeokimyasal özellikleri. Hidrojeolojide İzotop Tekniklerinin Kullanılması Sempozyumu, Ankara, Turkey. pp. 69–87 (in Turkish). Yılmazer, S., 2009. Batı Anadolu`nun olası jeotermal potansiyelinin belirlenmesi. Türkiye 11. Enerji Kongresi, İzmir, Turkey. pp. 1–12 (in Turkish).

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Origin and Evolution of the Cepheus Bubble

The Origin and Evolution of Photosynthetic Oxygen

The Origin and Evolution of Religious Prosociality

The origin and evolution of human glutaminases

Origin and Evolution of the Sodium -Pumping

Review The Origin and Early Evolution of

THE ORIGIN AND EVOLUTION OF CELESTIAL BODIES ...

Origin and evolution of the peroxisomal proteome

Origin and evolution of mineralizing fluids and

palaeopathology: studying the origin, evolution and ...

The evolution of HIV-1 and the origin of AIDS

The Multicellular Origin of Cancer and the Evolution of Oncogenesis

Origin and Evolution of Planetary Atmospheres - Core

Origin and Evolution of Near-Earth Objects

Evolutionary Origin of Politics and Political Evolution

Origin and Evolution of Processed Pseudogenes

Origin and Evolution of Continents - EAS

Origin, evolution and paleoepidemiology of brucellosis

Origin and Evolution of Dishevelled - Caltech Authors

Cosmology: The Origin, Evolution & Ultimate Fate of