Contemporaneous alkaline and calc-alkaline series in ...

Journal of Volcanology and Geothermal Research 356 (2018) 56–74

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Original research article

Contemporaneous alkaline and calc-alkaline series in Central Anatolia (Turkey): Spatio-temporal evolution of a post-collisional Quaternary basaltic volcanism Gullu Deniz Dogan-Kulahci a,⁎, Abidin Temel a, Alain Gourgaud b, Elif Varol a, Hervé Guillou c, Catherine Deniel b a b c

Hacettepe University, Geological Engineering Department, 06800, Beytepe, Ankara, Turkey Blaise Pascal University, UMR-CNRS 6524, Campus Universitaire des Cézeaux, 6 Avenue Blaise Pascal, TSA 60026 - CS 60026- 63178, Aubiere Cedex, France LSCE/IPSL, Laboratoire CEA-CNRS-UVSQ, Domaine du CNRS, Bât. 12, Avenue de la Terrasse 91198 Gif sur Yvette, Paris, France

a r t i c l e

i n f o

Article history: Received 27 April 2017 Received in revised form 16 January 2018 Accepted 16 February 2018 Available online 23 February 2018 Keywords: Central Anatolia Quaternary basalts Mineralogy Geochemistry Petrology K-Ar ages

a b s t r a c t This study focuses on spatio-temporal evolution of basaltic volcanism in the Central Anatolian post-collisional Quaternary magmatic province which developed along a NE-SW orientation in Turkey. This magmatic province consists of the stratovolcanoes Erciyes (ES) and Hasandag (HS), and the basaltic volcanic fields of Obruk-Zengen (OZ) and Karapınar (KA). The investigated samples range between basic to intermediate in composition (48–56 wt% SiO2), and exhibit calc-alkaline affinity at ES whereas HS, OZ and KA are alkaline in composition. Based on new K\\Ar ages and major element data, the oldest basaltic rock of ES is 1700 ± 40 ka old and exhibits alkaline character, whereas the youngest basaltic trachyandesite is 12 ± 5 ka old and calc-alkaline in composition. Most ES basaltic rocks are younger than 350 ka. All samples dated from HS are alkaline basalts, ranging from 543 ± 12 ka to 2 ± 7 ka old. With the exception of one basalt, all HS basalts are ~100 ka or younger in age. K\\Ar ages range from 797 ± 20 ka to 66 ± 7 ka from OZ. All the basalt samples are alkaline in character and are older than the HS alkaline basalts, with the exception of the youngest samples. The oldest and youngest basaltic samples from KA are 280 ± 7 ka and 163 ± 10 ka, respectively, and are calc-alkaline in character. Based on thermobarometric estimates samples from OZ exhibit the highest cpx-liqidus temperature and pressure. For all centers the calculated crystallization depths are between 11 and 28 km and increase from NE to SW. Multistage crystallization in magma chamber(s) located at different depths can explain this range in pressure. Harker variation diagrams coupled with least-squares mass balance calculations support fractional crystallization for ES and, to lesser extend for HS, OZ and KA. All basaltic volcanic rocks of this study are enriched in large-ion lithophile elements (LILE) and light rare earth elements (LREE). The lack of negative anomalies for high field strength elements (HFSE; Y, Yb) and the La/Nb N1 favor a shallow lithospheric source for ES, HS, OZ and KA basaltic volcanic rocks, whereas some samples bear the trace element signature of an asthenospheric mantle source. The lithospheric mantle beneath Central Anatolia may have not been affected from asthenospheric mantle directly. Negative Nb-Ta-Ti anomalies and a positive Pb spike of ES, HS, OZ and KA may be ascribed to crustal contamination or as the imprints of the previous subduction processes. According to this study, and previous studies, the effect of subduction and/or crustal contamination in Central Anatolia decreased from the Miocene to the Quaternary, and the origin of the Quaternary basaltic rocks mainly derived from subduction-related magmas enriched with sediment input rather than to slab-derived fluids. Our calculated eruption ages for the four basaltic complexes show that spatial differences predominate, whereas temporal trends are difficult to discern due to limited age resolution. According to the available geochronological, petrological and geochemical data, alkaline and calc-alkaline volcanism occurred simultaneously from distinct parental magmas. © 2018 Elsevier B.V. All rights reserved.

1. Introduction The Central Anatolia plate is located between the Afro-Arabian and the Eurasian plates. During the Early Miocene, the tectonic regime ⁎ Corresponding author. E-mail address: [email protected] (G.D. Dogan-Kulahci).

https://doi.org/10.1016/j.jvolgeores.2018.02.012 0377-0273/© 2018 Elsevier B.V. All rights reserved.

changed as a result of the convergence between the Eurasian and Afro-Arabian plates, which lead to widespread volcanism following the collision (Şengör, 1980; Şengör and Yılmaz, 1981). In addition deformation of the continental crust of Anatolia started along the Bitlis Suture Belt in Eastern Anatolia (Fig. 1). The shortening and thickening of the crust with time was unable to compensate for the deformation during Miocene and the Anatolian block started to drift westward along two

G.D. Dogan-Kulahci et al. / Journal of Volcanology and Geothermal Research 356 (2018) 56–74

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Fig. 1. a) Location map of the studied areas (http://www2.jpl.nasa.gov/srtm/) derived from 90-m SRTM digital elevation model (DEM). The DEM is prepared by hillshade module of ArcGis (v.10.3) and it is artificially illuminated from northeast with an angle of 45°. b) Geological map of the Erciyes stratovolcano (modified from Şen et al., 2003 and Kürkçüoğlu, 2000), with sampling locations. Legend: 21–14: Koçdağ stage; 13–1: Erciyes stage. c) Geological map of the Hasandag stratovolcano (modified from Aydar and Gourgaud, 1998), with sampling locations. Legend: 15–11: Paleovolcano; 10–8: Mesovolcano; 7–1: Neovolcano. d) Geological map of the Obruk-Zengen basaltic volcanic field (from MTA 1/500.000 geological map), with sampling locations Legend: 6: Pliocene, 5: Upper Miocene–Pliocene, 4–1: Quaternary. e) Geological map of the Karapınar basaltic volcanic field (from MTA 1/500.000 geological map), with sampling locations. Legend: 5: Middle Jurassic–Cretaceous, 4–1: Quaternary.

strike-slip faults, the North Anatolian Fault (NAF) and East Anatolian Fault (EAF) (McKenzie, 1972; Şengör et al., 1985; Dewey et al., 1986). Intraplate deformation developed within the Anatolian Block during in addition to pull-apart basins, normal, strike-slip and thrust faults during this time (Şengör, 1980; Dhont et al., 1998; Seyitoğlu et al., 2009). During the Neogene-Quaternary period, post-collisional magmatism occurred in Turkey, Italy, Spain and North Africa (Halloul and Gourgaud, 2012). The geochemical characteristics of post-collisional magmatism are comprehensive and mainly subduction-related, although subduction processes terminated with the onset of continental collision (Wang et al., 2004; Seyitoğlu and Scott, 1992; Seyitoğlu et al., 1997). According to previous studies focused on Neogene Quaternary volcanism (Deniel et al., 1998; Temel et al., 1998; Alıcı Şen et al., 2004; Kürkçüoğlu, 2010; Genç and Yürür, 2010; Özsayın et al., 2013), the major tectonic elements (Tuzgölü fault zone and Ecemiş fault zone) and some local faults are still active in Central Anatolia, and volcanism is considered to be related to previous subduction. In addition, the subduction-related signatures are attributed to metasomatism by slabderived fluids of the mantle lithosphere prior to collision (Pearce et al., 1990; Turner et al., 1992, 1993, 1996; Platt and England, 1993; Peccerillo, 1999). Furthermore, geochemical data for the Tepeköy Volcanic Complex in Central Anatolia have been interpreted to indicate a subcontinental lithospheric mantle source enriched in incompatible

elements due to previous subduction processes (Kuşçu-Gençalioğlu and Geneli, 2010). Recent studies of the eastern and western Mediterranean on the European and Eurasian plates discuss various scenarios regarding crustal thickness, amounts and timing of uplift, subduction, slab detachment, slab delamination, slab rollback and slab break-off processes, geophysical as well as mantle-crust relation and asthenosphere-lithosphere interaction (Govers and Wortel, 2005; Biryol et al., 2011; Cosentino et al., 2012; Schildgen et al., 2012; Faccenna et al., 2006, 2013; Aydar et al., 2013; Fichtner et al., 2013; Vinnik et al., 2014; Delph et al., 2015; Kind et al., 2015; Govers and Fichtner, 2016; Reid et al., 2017; Bilim et al., 2017; Göğüş et al., 2017). One of the most significant datasets suggests that the eastern sector of the Cyprus slab is too deep (˃200 km) to have solely generated the young volcanism in Central Anatolia from the Late Miocene until the Quaternary period (Biryol et al., 2011). However, teleseismic P-wave tomograms of Biryol et al. (2011) show slow velocity perturbations about (1–2%) to 200 km below the Central Anatolian volcanic zone therefore, it is possible that these slow anomalies are due to the presence of ascending asthenosphere which is responsible Central Anatolian volcanism. Conversely, Bartol and Govers (2014) showed that the effect of the Cyprus slab, presently at 660 km depth, could not have generated Central Anatolian volcanism alone. Seismic tomography imaged a single slab moving horizontally from the west towards the east, and the authors suggest a model

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where instantaneous delamination occurred both beneath the Eastern Anatolian Plateau (EAP) and the Central Anatolian Plateau (CAP). Significant contributions have been made in the literature regarding the geochronology (Besang et al., 1977; Ercan et al., 1990, 1994; Notsu et al., 1995; Deniel et al., 1998; Kuzucuoğlu et al., 1998; Şen et al., 2003; Schmitt et al., 2014; Aydın et al., 2014; Sarıkaya et al., 2017; Reid et al., 2017), stratigraphy (Innocenti et al., 1975; Pasquarè et al., 1988; Le Pennec et al., 1994), petrology (Aydar, 1992; Aydar and Gourgaud, 1998) and geochemistry (Temel et al., 1998; Kürkçüoğlu et al., 1998; Alıcı Şen et al., 2004; Siebel et al., 2011) of Central Anatolian volcanism, however there is no adequate data on the spreading of basaltic volcanism over large areas in Central Anatolia. Therefore, the purpose of this paper is to fill this gap, and to reconstruct the spatiotemporal evolution of the post collisional Quaternary basaltic volcanism in Central Anatolia. Only undifferentiated basaltic rock samples, which give better information about the source, are considered. We present here new geochronological, mineralogical and petrological data on the basaltic units of Erciyes (ES) and Hasandag (HS) stratovolcanoes and of the Obruk-Zengen (OZ) and Karapınar (KA) basaltic volcanic fields (Fig. 1). 2. Geological background After the Early Miocene continental collision between Afro-Arabian and Eurasian plates, structural changes occurred in the Anatolian Block (Şengör, 1980; Şengör and Yılmaz, 1981; Dewey et al., 1986; Dhont et al., 1998; Seyitoğlu et al., 2009). For example the Sultansazlığı pullapart basins associated with ES formed during that time. Moreover, the Ecemiş and Tuzgölü strike-slip fault systems and the Derinkuyu normal faults are the most important structural elements in Central Anatolia (Fig. 1a). Volcanic rocks cover a wide surface and overlie the Paleozoic– Mesozoic metamorphic series. Volcanic centers comprise stratovolcanoes, maars, domes, lava flows, scoria cones and ignimbrite-forming calderas of Miocene-Quaternary age. In addition, felsic ignimbrites of the Mio-Pliocene have left spectacular and voluminous formations in Cappadocia (Temel, 1992; Le Pennec et al., 1994; Temel et al., 1998). Detailed stratigraphic studies of these volcanic sequences have been made by previous workers (e.g., Pasquarè et al., 1988; Ayrancı, 1969, 1991; Ercan et al., 1990; Aydar, 1992; Le Pennec et al., 1994; Temel et al., 1998; Şen, 1997).

southern (Dikkartın) flanks of the Erciyes stratovolcano were also dated by Sarıkaya et al. (2017) (Fig. 1b). Average exposure ages of 7.2 ± 0.9 ka and 7.7 ± 0.4 ka were obtained from Karagüllü and Perikartın, respectively. The Çarık lava flow has been dated at around 98.4 ± 3.6 ka and 36.1 ± 1.1 ka, and Dikkartın at 8.8 ± 0.6 ka. It has been suggested by Kürkçüoğlu (2000) that the dacitic and rhyodacitic units could have been followed by alkaline basaltic products in the Northern part of ES. The volcanic history of the HS is divided in three successive phases: Paleovolcano, Mesovolcano and Neovolcano (Aydar, 1992). The Paleovolcano phase comprises the Keçikalesi tholeiitic volcanic unit, dated between 13.7 ± 0.3 and 12.4 ± 4 Ma (Besang et al., 1977), and basaltic andesites, 7.21 ± 0.09 Ma in age (K/Ar; Aydar and Gourgaud, 1998). The most voluminous phase of this stage comprises rhyolitic ignimbrites which have been dated at 6.31 ± 0.20 Ma (Deniel et al., 1998). The Mesovolcano phase comprises calc-alkaline andesitic lava flows, domes and block-and-ash flow deposits. During the Neovolcano phase, effusive and explosive activity produced andesitic to rhyolitic domes, block-and-ash flows, phreatomagmatic breccias and rhyolitic ignimbrites (Aydar, 1992). The ages of the dome and blocky flows near Karakapı village (Fig. 1c) range between 0.08 and 0.58 Ma (Ercan et al., 1990). Following this stage, basaltic cones were emplaced to the northeast of the Gözlükuyu village between 120 and 65 ka (Ercan et al., 1990), and lava flows, 34 ka, (Aydar and Gourgaud, 1998) drape the flanks of Yıprak Hill (Fig. 1c). The youngest effusive activity is dated between 29 and 33 ka (Kuzucuoğlu et al., 1998). The youngest volcanic products of the HS comprise alkaline basaltic domes, block and ash flows (Aydar and Gourgaud, 1998), and explosively emplaced andesitic deposits dated at ca. 29 and 9 ka (Schmitt et al., 2014). The postcaldera andesitic dome is thought to be younger than 6 ka (Aydar and Gourgaud, 1998). The youngest products of ES and HS are mainly acidic, however radiometric dating of the latest basaltic volcanism is limited. Most of basaltic rocks display alkaline and calc-alkaline affinity in HS and ES, respectively (Fig. 2). 2.2. Obruk-Zengen (OZ) and Karapınar (KA) basaltic volcanic fields Monogenetic volcanoes are exposed over a ~2400 km2 area in Central Anatolia (Reid et al., 2017). The monogenetic volcanism of ObrukZengen (OZ) extends towards the south and southwest of Hasandag

2.1. Erciyes (ES) and Hasandag (HS) stratovolcanoes The volcanic history of the ES can be subdivided into two evolutionary stages, where the alkaline Koçdağ stage preceded the calc-alkaline Erciyes stage (Kürkçüoğlu et al., 1998; Şen et al., 2003). During the Koçdağ stage, volcanic activity was effusive with tholeiitic basalts, and followed by intermediate lavas (Fig. 1b). Later, monogenetic scoria cones of basaltic andesite composition were emplaced. The final products of the Koçdağ stage are pyroclastics dated at 2.8 ± 0.1 Ma (Innocenti et al., 1975), and dacitic domes and intra caldera andesitic lavas dated at 2.544 ± 0.306 Ma (Notsu et al., 1995). The Koçdağ stage ended with a caldera collapse. The Erciyes stage was preceded by andesitic eruptions, after which the volcanic activity became more explosive, resulting in dacitic and rhyodacitic eruptions (Şen, 1997). The dacites have been dated at 0.3 ± 0.1 Ma (Innocenti et al., 1975), alkali basalts and basaltic andesites at 0.15 ± 0.07 Ma (K/Ar age, Ercan et al., 1994) and at 0.171 ± 0.12 Ma (K/Ar age; Notsu et al., 1995). The volcanic activity continued until more recent times and resulted in base surges, plinian fallouts and rhyodacitic pumice flows which spread over 11.000 km2 around a felsic dome (Pasquarè et al., 1988). The age of these eruptions are 0.13 ± 0.02, 0.14 ± 0.02 and 0.115 ± 0.02 Ma (K/Ar; Ercan et al., 1994). One of the youngest dated volcanic products of the Erciyes is the Çarık lava flow, 40Ar/39Ar dated at 80 ± 10 ka, (Notsu et al., 1995). Young monogenetic parasitic lava domes, mostly dacitic and rhyolitic from the northern (Karagüllü, Perikartin, Çarık) and

Fig. 2. Total alkaline (Na2O + K2O wt%) vs. silica (SiO2 wt%) diagram of the volcanic rocks from ES, HS, OZ and KA (Le Bas et al., 1986) and alkali-subalkali discrimination (dashed line, according to Miyashiro, 1978).


stratovolcano (HS), and comprises basaltic scoria cones with associated lava flows (Fig. 1d). The monogenetic volcanism of Karapınar (KA), located to the southwest part of OZ, consists of basaltic scoria cones with related lava flows and maars, such as Meke (Fig. 1e). The first volcanic products of KA, located near Karapınar town, were trachyandesitic and andesitic lavas, followed mainly by basaltic lava flows and phreatomagmatic deposits (Ercan et al., 1990). The majority of rock samples collected from OZ and KA comprise alkaline basalts and trachybasalts, calc-alkaline basalt and basaltic-andesite (Fig. 2). A recent study (Reid et al., 2017) dated the Egrikuyu (named here OZ) samples between 0.6 and 0.2 Ma (40Ar/39Ar). KA samples dated in the same study are younger than 0.6 Ma.

3. Analytical techniques To obtain a representative set of samples spanning all the stratigraphic units our sampling was guided by previous studies establishing the geological framework of the ES (Innocenti et al., 1975; Ercan et al., 1994; Notsu et al., 1995; Şen et al., 2003; Kürkçüoğlu et al., 1998) and HS (Aydar and Gourgaud, 1998; Deniel et al., 1998; Ercan et al., 1990, 1994) stratovolcanoes, as well as the OZ (Ercan et al., 1994; Notsu et al., 1995) and KA (Ercan et al., 1990; Notsu et al., 1995). Unaltered samples are preferable because diagenesis may cause the gain and/or loss of K and Ar and as a consequence may lead to under or overestimated K\\Ar ages (e.g., Guillou et al., 2017). Therefore, the sample selection was also guided by loss-on-ignition values (LOI), macroscopic (hand specimen) and microscopic (thin section) observations. Faure and Mensing (1993) showed that for water contents of b1%, Rb, which is a more mobile element than K, remained stable in the volcanic rocks. The LOI of our samples range between −0.58% and 0.58% (Table 6). Therefore we consider all the dated samples unaltered as their LOI is b1%. Rock samples, of approximately 500 g were crushed and sieved to 125–250 μm. The samples were then ultrasonically washed in an acetic acid bath at 60 °C for 40 min. Successive magnetic and densitometric (diiodomethane) separations were used to remove phenocrysts, potential carriers of excess 40Ar*, from the groundmass using the procedure detailed in Guillou et al. (1998). Samples were then carefully rinced with acetone, and dried. The homogeneity and freshness of the groundmass aliquots were then checked via microscopic inspection. K analyses were made at the Centre de Recherches Pétrographiques et Géochimiques (CRPG), (Nancy, France). The isotopic composition and abundance of Ar were determined using an unspiked technique described by Charbit et al. (1998). In this technique, argon extracted from the sample is measured in sequence with purified aliquots of atmospheric argon at comparable working gas pressure in the massspectrometer, to suppress mass discrimination effects between the atmospheric reference and the unknown. Electron microprobe analyses of the cores and rims of phenocrysts, microcrysts and microlites were undertaken at the Laboratoire Magmas et Volcans, Université Blaise Pascal, (Clermont-Ferrand, France) using a Cameca SX-100. Beam parameters for phenocrysts, microcrysts and microlites were 15 kv, 10–12 nA and 10 s and for glass analyses 15 kv, 4 nA and 5 μm, 10 s. Whole-rock major element analyses were performed at Service d'Analyse des Roches et des Minéraux (SARM), CRPG–CNRS (Nancy, France) using a Jobin-Yvon JY 70 inductively coupled plasma atomic emission spectrometer (ICP-AES). Trace and rare earth elements were analyzed by inductively coupled plasma mass spectrometry (ICP-MS) at CRPG. Before solutions were analyzed by the automatic procedure installed in the system, 200 mg powdered samples, were molten by LiBO2 and dissolved by HNO3. Detection limits and uncertainties for major elements at given intervals are presented on the SARM-webpage (http://helium.crpg.cnrs-nancy.fr/ SARM/pages/roches.html).

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4. Mineralogy and petrography 4.1. Basaltic rocks of Erciyes and Hasandag stratovolcanoes The common mineral assemblage of the calc-alkaline ES basaltic rocks consists of plagioclase + olivine + clinopyroxene ± orthopyroxene and oxide. The majority of samples display porphyritic textures, while some phenocrysts contain glassy inclusions, reaction rims and exhibit strong zoning. Disequilibrium features are observed in most plagioclase phenocrysts, either as resorbed and fragmented crystals, or as zoned crystals with clear homogeneous cores and sieved rims (or vice versa), whereas other crystals only display sieve textures. Olivine crystals exhibit iddingsitization. Clinopyroxene crystals occur as microlites in the groundmass and as microcrysts of few hundred microns in size. Orthopyroxene is found only in ES calk-alkaline basaltic samples. The groundmass is blackish and consists mainly of microlites of the above mentioned minerals. The alkali basalts collected on the flanks of HS contain plagioclase + olivine + clinopyroxene and oxide minerals, exhibiting hypocrystallineporphyritic textures. Some plagioclase crystals exhibit sieve textures. Olivine phenocrysts and microcrysts are euhedral in shape and some display skeletal and resorbed morphology. Clinopyroxene crystals occur as phenocrysts, microcrysts and/or microlites in the matrix. Few of the basaltic samples exhibit “sector zoning” and serpentinization. The groundmass consists of microlites of plagioclase, olivine and clinopyroxene. 4.2. Basaltic rocks of Obruk-Zengen (OZ) and Karapınar (KA) basaltic volcanic fields Basalts of OZ include microcrysts of plagioclase + olivine + clinopyroxene and oxide minerals. Generally the basalts display a hypocrystalline porphyritic texture. Some of the minerals exhibit sieve textures. Euhedral plagioclases display sub-trachytic texture and olivine crystals display skeletal and resorbed features. Some olivines contain inclusions of Cr-spinel and have occasionally experienced iddingsitization. Clinopyroxene crystals are common as microlites (few millimeters in size) in the matrix, whereas some crystals are resorbed and fragmented. Groundmass microlites are represented by plagioclase, olivine, clinopyroxene and oxide minerals. The basaltic rocks of the KA volcanoes exhibit a mineral assemblage of plagioclase + olivine + clinopyroxene and oxide, having dominantly a hypocrystalline porphyritic texture. Plagioclase crystals are mostly subhedral and sometimes resorbed. Some plagioclase crystals display sieve texture. The olivine phenocrysts are embayed and resorbed, and some crystals contain inclusions of Cr-spinel. Clinopyroxene crystals are mainly as subhedral, and sometimes resorbed phenocrysts a few millimeters in size. The groundmass exhibits microlites of the afore mentioned minerals. 5. K\\Ar geochronology 5.1. The new K\\Ar ages of the Central Anatolian basaltic rocks K\\Ar ages of the ES, HS, OZ and KA are presented in Table 1a-d. Based on the new K\\Ar data, the oldest sample of ES is an alkaline trachybasalt (1700 ± 40 ka; ER 13), and the youngest sample is an alkaline basaltic trachyandesite (12 ± 5 ka; ER 29) (Fig. 1b). These dates are consistent with two 40Ar/39Ar ages of samples collected along the southwest slope of ES (580 ± 130 and 210 ± 180 ka) by Higgins et al. (2015). Our K\\Ar ages show that the youngest samples are located northeast and northwest of the ES summit. Dates obtained from cosmogenic 36Cl of samples collected by Sarıkaya et al. (2017) from the north of ES (Carık Tepe), located in the same area as our sample (ER11) (Fig. 1b), are consistent with our K\\Ar dates obtained from this study. The oldest sample from HS (HA23b) is an alkaline basalt located northwest of the main cone (Mahmutlu Valley), (Fig. 1a), and dated at

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Table 1 Unspiked K\ \Ar ages of the Cappadocia samples. a. Erciyes stratovolcano b. Hasandag stratovolcano c. Obruk-Zengen d. Karapınar. The isotopic composition and abundance of Ar were determined using the unspiked technique described in Charbit et al. (1998). Unspiked K\ \Ar analysis of each sample involved three independent determinations of potassium and two of argon. Based on replicate analysis of material references the potassium concentrations are determined with an uncertainty of 1% (1σ). The potassium concentrations are combined to yield a mean value. Ages of each sample are calculated using this value and the weighted mean of the two independent measurements of 40Ar*(radiogenic argon). Uncertainties for the Ar data are 1σ analytical only, and consist of propagated and quadratically averaged experimental uncertainties arising from the 40Ar (total), and 40Ar* determinations. Uncertainties on the ages are given at 2σ. Sampling locations are given by UTM/ED50. Sample ID experiment n°

ER11 8200 8221 ER29 8134 8165 ER17 8161 8177 ER1 8164 8180 ER20 8186 8201 ER10 8261 8328 ER3 6861 6877 ER32 8399 8415 ER13 8241 8257 Sample ID experiment n°

UTM coordinates (ED50)

Rock type

Weight molten (g)

K* (wt%)

40Ar* (%)

40Ar* (10−13 mol/g)

40Ar* weighted mean (± 1σ)

Age ± 2σ ka

0.288 ± 0.057

12 ± 5

E

N

712236

4276926

Ba.And

1.00758

1.046 ± 0.011

−0.393

b0

699444

4275215

Ba.Tr.And

1.00562 1.04062

“…………..” 1.345 ± 0.013

−0.097 0.085

b0 2.753

698720

4272507

Ba.And

1.50364 1.00525

“…………..” 1.287 ± 0.013

0.100 1.397

2.988 1.353

713819

4256265

Ba.And

1.50996 1.0115

“…………..” 1.038 ± 0.010

0.824 0.209

1.306 1.65

1.357 ± 0.058 1.669 ± 0.097

93 ± 11

702836

4265824

Ba.And

1.00877 1.00569

“…………..” 1.146 ± 0.011

0.199 0.716

1.722 2.829

2.922 ± 0.062

147 ± 7

716512

4280805

Ba.And

1.00353 0.49442

“…………..” 0.855 ± 0.009

0.671 1.033

2.978 4.619

4.405 ± 0.105

297 ± 15

718732

4283383

Ba.And

0.50032 1.36217

“…………..” 0.880 ± 0.009

1.177 4.773

4.373 5.359

5.373 ± 0.083

352 ± 9

725707

4271840

Ba.And

1.90831 1.01111

“…………..” 0.838 ± 0.008

6.506 6.538

5.380 10.062

10.282 ± 0.044

707 ± 15

692946

4268911

Tr. Basalt

1.5034 0.49576

“…………..” 0.830 ± 0.083

7.080 17.302

10.490 24.538

24.419 ± 0.105

1695 ± 37

1.01285

“…………..”

25.758

24.330

Rock Type

Weight molten (g)

K* (wt%)

40Ar* (%)

40Ar* (10−13 mol/g)


Age ± 2σ ka


59 ± 5

E

N

HA31 8138 8181 HA13b 8288 8289 HA20 8213 8237 HA19 8195 8227 HA18 8247 8260 HA23b 8396 8412

594308

4220374

Basalt

1.05637

0.847 ± 0.009

0.018

0.038

0.031 ± 0.052

2±7

599154

4228358

Basalt

1.50318 0.99537

“…………..” 1.079 ± 0.011

0.022 0.417

0.028 0.191

0.283 ± 0.042

15 ± 5

598739

4234457

Basalt

1.50806 1.00576

“…………..” 1.038 ± 0.005

0.797 1.150

0.343 0.906

1.030 ± 0.047

57 ± 5

594623

4235116

Basalt

1.00321 1.00486

“…………..” 1.129 ± 0.011

0.879 1.927

1.051 1.592

1.418 ± 0.041

72 ± 4

592415

4232594

Tr. Bas

1.28385 1.00058

“…………..” 1.212 ± 0.012

1.887 1.657

1.347 2.092

2.129 ± 0.052

101 ± 5

622176

4227901

Basalt

1.50051 0.99697

“…………..” 0.996 ± 0.010

2.412 9.860

2.164 0.930

9.379 ± 0.042

543 ± 12

2.00253

“…………..”

8.005

0.944

Sample ID experiment n°


HA17 6862 6878 HA27 8394 8410 HA15 8212 8236 HA14 8217 8233 HA26

Rock type

Weight molten (g)

K* (wt%)

40Ar* (%)

40Ar* (10−13 mol/g)


Age ± 2σ ka

E

N

593925

4187914

Basalt

1.37556

0.805 ± 0.008

1.130

0.993

0.919

66 ± 7

610725

4205756

Ba.Tr.And

2.12687 1.0037

“…………..” 1.619 ± 0.016

1.107 1.850

0.868 9.365

9.303 ± 0.046

331 ± 7

596074

4200670

Basalt

1.02372 1.02037

“…………..” 1.079 ± 0.011

1.616 5.091

9.216 7.561

7.583 ± 0.049

405 ± 10

598610

4196847

Tr.Bas

1.49874 0.50150

“…………..” 1.229 ± 0.012

5.675 4.090

7.588 8.457

8.873 ± 0.061

416 ± 10

615277

4192594

Basalt

1.03149 0.480206

“…………..” 1.039 ± 0.01

5.236 5.688

8.930 11.797

12.117 ± 0.070

672 ± 16


61

Table 1 (continued) Sample ID experiment n°

UTM coordinates (ED50) E

N

8666 8682 HA16 8137 8162

601553

4185909

Sample ID experiment n°


KA19 8397 8413 KA21 8245 8262 KA18 (CK89–19)

Rock type

Weight molten (g)

K* (wt%)

40Ar* (%)

40Ar* (10−13 mol/g)

0.93817 1.00134

“…………..” 0.656 ± 0.007

4.847 12.273

12.195 9.043

1.50996

“…………..”

4.975

9.098

Rock type

Weight molten (g)

K* (wt%)

40Ar* (%)

40Ar* (10−13 mol/g)

Basalt


Age ± 2σ ka

9.077 ± 0.071

797 ± 20


Age ± 2σ ka

E

N

557051

4171549

Tr.Bas.

1.00007

0.938 ± 0.009

−0.094

b0

4174038

Ba.And

2.01299 0.50028

“…………..” 1.121 ± 0.012

−0.112 0.756

b0 3.152

–

551182

0.50674

“…………..”

0.855

3.260

3.162 ± 0.091

552073

4160791

Ba.And

163 ± 10

280 ± 7

543 ± 12 ka; This age is consistent with the 40Ar/39Ar age (654 ± 24 ka, sample K-51) of samples obtained from the same area (Aydın et al., 2014). The youngest sample is also an alkaline basalt (2 ± 7 ka; HA31), located west of the main cone (Fig. 1c). All basaltic samples of HS, excluding HA23b, are ~100 ka or younger. These young ages agree with the data of Schmitt et al. (2014), who dated the Holocene activity of HS. The oldest OZ sample (HA16) is from Gaffar Hill, dated at 797 ± 20 ka. The youngest sample (HA17) is located to the west of Gaffar Hill and dated at 66 ± 7 ka (Fig. 1d). The K\\Ar age of sample HA16 is thought to be a youngest age, because the same unit has been dated by 40Ar/39Ar at 1040 ± 30 ka (Reid et al., 2017). In the same sector, a trachy-basalt with an 40Ar/39Ar age of 490 ± 47 ka (Aydın et al., 2014, sample K-81) is consistent with the age range defined in this study. The K\\Ar ages of samples HA14 and HA15, 416 ± 10 and 406 ± 10 ka respectively, are in good agreement with the 40Ar/39Ar ages of 450 ± 20 ka (sample R13EG40) and 490 ± 30 ka (sample R11EG08) reported by Reid et al. (2017) for the same unit. K\\Ar dating indicates that basalts of the OZ are older than those of HS, with the exception of sample HA17. Due to alteration, the number of KA samples suitable for dating was limited. Three 40Ar/39Ar ages constrain the volcanic activity in this area between 500 ± 20 and 290 ± 10 ka (Reid et al., 2017).

The oldest sample is a calc-alkaline basalt collected from south of Karapınar (KA 18). Sample KA18 is considered as equivalent to sample CK-89-19 of Olanca (1994) as they were collected from the same location (Table 1). The K\\Ar age obtained for sample KA 21, a calcalkaline basaltic andesite sampled east of Karapınar, of 163 ± 10 ka is younger than 40Ar/39Ar age of Reid et al. (2017). 6. Mineral chemistry Average mineral compositions are given in Table 2, and selected electron microprobe analyses of single clinopyroxenes, orthopyroxenes, plagioclases, olivines and oxides are shown in Table 3 and Table 4. 6.1. Clinopyroxene In ES basaltic rocks, the clinopyroxene phenocrysts composition ranges from diopside to augite. The composition of clinopyroxene is Wo(33–46), En(42–52), Fs(9–22) for phenocrysts and microcrysts, and Wo(23–47), En(31–52), Fs(7–22) for microlites (Table 2). The Mg-numbers of are wide ranging from 59 to 87, and oscillatory-zoned clinopyroxenes are common. Normal-zoned clinopyroxene phenocrysts with Mg-rich cores coexist with reversed-zoned clinopyroxene phenocrysts with

Table 2 Average mineral compositions of basaltic rocks from ES, HS, OZ and KA. (Plg: plagioclase, Cpx: clinopyroxene, Opx: orthopyroxene, An: anorthite, Fo: forsterite, Wo: wollastonite, En: enstatite, Fs: ferrosillite, Usp: ulvo-spinel, Ilm: ilmenite).

Plg (Phenocrysts & microcrysts) Plg (Microlites) Olivine (Phenocrysts & microcrysts) Olivine (Microlites) Cpx (Phenocrysts & microcrysts)

Cpx (Microlites)

Opx(Phenocrysts & microcrysts)

Opx (Microlites)

Ti-Magnetite Ilmenite Spinel

ES

HS

OZ

KA

An (35–85) An (39–70) Fo% (57–86) Fo% (66–83) Wo(33–46) En(42–52) Fs(9–22) Wo(23–47) En(31–52) Fs(7–22) Wo(2−13) En(65–83) Fs(14–30) Wo(3–9) En(66–79) Fs(18–29) Usp(35–73) Ilm(98) *

An (37–85) An (52–72) Fo% (68–86) Fo% (53–75) Wo(42–47) En(41–47) Fs(8–16) Wo(41–45) En(39–46) Fs(10–17) –



–

–

–

Usp(57–62) Ilm(94) *

Usp(38–58) * *

Usp(40–63) * *

62


Table 3 Selected microprobe analyses of clinopyroxene and orthopyroxene from ES, HS, OZ and KA (c:center, r:rim, mic:microlite). Samples

ER1

ER13

HA19

HA13b

HA14

HA16

KA19

ER1

Clinopyroxene

SiO2 (wt%) TiO2 Al2O3 FeO MnO MgO CaO Na2O Total Si Alıv Alvı Fe+3 Fe+2 Mn Mg Ca Na Total Ae Aug Wo En Fs Mg#

ER3

ER17

Orthopyroxene

Microcryst

Microlite

Microlite

Phenocryst

Microlite

Phenocryst

Phenocryst

Microlite

Phenocryst

Microcryst

69—c

70—r

57—mic

102—mic

24—c

25—r

77—mic

30—c

31—r

123—c

124—r

56—mic

3—c

4—r

121—r

51.50 0.90 3.62 7.31 0.22 17.26 18.90 0.25 100.06 1.89 0.12 0.04 0.04 0.18 0.01 0.94 0.74 0.02 4.00 4.35 91.41 38.71 49.22 12.07 80.79

50.08 0.78 4.50 6.23 0.21 15.81 20.37 0.35 99.01 1.86 0.15 0.05 0.06 0.14 0.01 0.87 0.81 0.03 4.00 5.99 88.66 42.96 46.41 10.63 81.89

47.25 3.46 4.68 11.83 0.17 10.61 20.91 0.73 99.72 1.80 0.20 0.01 0.05 0.33 0.01 0.60 0.85 0.05 4.00 5.70 93.40 46.43 32.77 20.79 61.49

52.26 0.85 2.18 6.76 0.19 15.99 20.87 0.31 99.64 1.93 0.07 0.03 0.02 0.19 0.01 0.88 0.83 0.02 4.00 1.86 95.45 43.00 45.81 11.19 80.81

50.14 1.18 3.97 6.81 0.16 14.99 21.67 0.36 99.64 1.86 0.14 0.03 0.06 0.15 0.01 0.83 0.86 0.03 4.00 6.62 90.33 45.16 43.48 11.36 79.67

51.12 1.01 3.45 7.03 0.17 15.34 21.97 0.40 100.74 1.87 0.13 0.02 0.08 0.14 0.01 0.84 0.86 0.03 4.00 8.03 89.99 44.91 43.61 11.48 79.54

50.38 1.30 2.55 8.28 0.20 14.38 21.61 0.40 99.11 1.89 0.11 b.d. 0.07 0.19 0.01 0.80 0.87 0.03 4.00 7.52 92.48 44.81 41.50 13.70 75.61

52.53 0.55 3.10 5.00 0.14 16.61 20.86 0.42 99.39 1.93 0.07 0.06 0.01 0.15 b.d. 0.91 0.82 0.03 4.00 0.61 93.03 43.49 48.15 8.37 85.51

50.80 0.73 4.65 4.71 0.11 15.75 22.03 0.39 100.20 1.86 0.15 0.06 0.05 0.10 b.d. 0.86 0.86 0.03 4.00 5.07 89.12 46.20 45.93 7.88 85.61

48.31 1.45 6.18 8.92 0.19 14.12 20.20 0.50 99.93 1.79 0.21 0.06 0.11 0.16 0.01 0.78 0.80 0.04 4.00 11.84 82.08 43.02 41.84 15.15 73.84

49.36 1.29 4.48 7.90 0.06 14.01 21.96 0.47 99.64 1.83 0.17 0.03 0.10 0.15 b.d. 0.78 0.87 0.03 4.00 10.22 86.65 46.07 40.91 13.02 76.00

54.00 0.48 0.92 16.20 0.46 25.28 2.26 0.06 99.65 1.97 0.03 0.01 b.d. 0.49 0.01 1.38 0.09 b.d. 4.00 b.d. 98.97 4.46 69.76 25.77 73.57

54.35 0.16 2.67 12.13 0.25 28.72 1.46 0.06 99.90 1.93 0.07 0.04 0.02 0.35 0.01 1.52 0.06 b.d. 4.00 1.52 94.01 2.88 78.16 18.96 80.82

55.43 0.12 2.06 10.13 0.21 29.82 1.38 0.04 99.46 1.97 0.03 0.05 b.d. 0.30 0.01 1.58 0.05 b.d. 4.00 b.d. 94.82 2.74 81.45 15.81 84.01

54.47 0.38 0.50 14.94 0.47 25.17 3.65 0.02 99.73 1.98 0.02 0.01 b.d. 0.46 0.01 1.37 0.14 b.d. 4.00 b.d. 99.38 7.18 69.11 23.71 75.03

b.d.' (below detection).

Mg-rich rims. In the HS samples, clinopyroxene phenocrysts, microcrysts and microlites are represented by diopside and augite. The composition of clinopyroxene is Wo(42–47), En(41–47), Fs(8–16) for phenocrysts and microcrysts, and Wo(41–45), En(39–46), Fs(10–17) for microlites (Table 2). The ranges of Mg-numbers are narrower from 72 to 85, than for ES basaltic rocks. Overall, clinopyroxenes display normal zoning, but some reverse-zoned clinopyroxene phenocrysts

and microcrysts can also be found. Similarily, clinopyroxenes of OZ are all diopside and augite in composition, and the composition of clinopyroxene is Wo(40–49), En(41–48), Fs(8–17) for phenocrysts and microcrysts (Table 2). The composition of microlites is Wo(40–48), En(35–46), Fs(10–19) and their Mg-numbers range widely between 65 and 84. The clinopyroxenes of KA are diopside and augite, and the composition of clinopyroxene is Wo(41–46), En(40–48), Fs(9–17) for phenocrysts

Table 4 Selected microprobe analyses of plagioclase, olivine and oxides from ES, HS, OZ and KA (c: center, r: rim, mic: microlite). Samples

ER1

ER10

HA20

HA31

HA14

HA16

HA17

KA19

KA21

Plagioclase

SiO2 (wt%) Al2O3 FeO MgO CaO Na2O K2O Total Si Al Fe Mg Ca Na K An Ab Or

Phenocryst

Phenocryst

Microcryst

Microlite

Microlite

Microcryst

Microlite

Phenocryst

Phenocryst

59—c

60—r

26—c

27—r

129—c

130—r

89—mic

119—mic

2—c

3—r

132—mic

116—c

118—r

54—c

55—r

52.78 29.73 0.50 0.05 12.89 4.06 0.09 100.09 9.57 6.36 0.08 0.01 2.50 1.43 0.02 63.38 36.10 0.52

51.80 29.88 0.48 0.08 13.28 3.98 0.13 99.62 9.47 6.43 0.07 0.02 2.60 1.41 0.03 64.34 34.93 0.73

52.41 29.66 0.46 b.d. 12.80 4.19 0.24 99.79 9.55 6.37 0.07 b.d. 2.50 1.48 0.06 61.94 36.68 1.38

51.30 30.63 0.56 b.d. 13.81 3.75 0.18 100.28 9.34 6.57 0.09 b.d. 2.69 1.32 0.04 66.35 32.61 1.04

54.80 27.88 0.90 0.11 10.72 5.03 0.51 100.03 9.93 5.96 0.14 0.03 2.08 1.77 0.12 52.49 44.54 2.97

52.65 28.90 0.94 0.12 12.30 4.26 0.29 99.47 9.64 6.23 0.14 0.03 2.41 1.51 0.07 60.43 37.86 1.70

51.84 29.61 1.11 0.09 12.91 4.06 0.21 99.95 9.47 6.38 0.17 0.02 2.53 1.44 0.05 62.99 35.82 1.19

52.31 28.96 0.88 0.09 12.69 4.30 0.35 99.75 9.57 6.25 0.13 0.02 2.49 1.52 0.08 60.77 37.24 2.00

50.35 30.73 0.78 0.11 14.17 3.24 0.12 99.55 9.24 6.65 0.12 0.03 2.79 1.15 0.03 70.28 29.03 0.69

50.59 30.38 0.79 0.14 14.18 3.33 0.13 99.68 9.28 6.57 0.12 0.04 2.79 1.18 0.03 69.63 29.58 0.79

50.41 30.80 0.84 0.17 14.74 3.13 0.17 100.42 9.20 6.62 0.13 0.04 2.88 1.11 0.04 71.53 27.50 0.98

48.54 32.35 0.63 0.05 15.83 2.65 0.09 100.13 8.90 6.99 0.10 0.01 3.11 0.94 0.02 76.37 23.12 0.50

50.82 30.41 0.63 0.09 13.93 3.50 0.20 99.65 9.31 6.57 0.10 0.02 2.74 1.25 0.05 67.91 30.92 1.17

50.71 30.80 0.51 0.10 14.20 3.50 0.11 100.00 9.26 6.63 0.08 0.03 2.78 1.24 0.03 68.72 30.63 0.65

50.62 31.25 0.66 0.15 14.57 3.06 0.12 100.45 9.21 6.70 0.10 0.04 2.84 1.08 0.03 71.95 27.32 0.73

“b.d.” (below detection).


63

and microcrysts, and Wo(41–46), En(40–48), Fs(9–16) for microlites (Table 2). Their Mg-numbers are represented by a narrow range from 71 to 84. Phenocrysts and microcrysts are normally zoned with Mgrich cores and Fe rich rims. 6.2. Orthopyroxene Orthopyroxene occurs only in ES, being of enstatite and pigeonite composition with Wo (3–9) , En (65–82) , Fs (15–30) . Sample ER32 only contains enstatite composition orthopyroxene. Normallyzoned orthopyroxene phenocrysts with Mg-rich cores coexist with reverse-zoned orthopyroxene phenocrysts with Mg-rich rims. The Mg-numbers of orthopyroxene range between 69 and 84. 6.3. Feldspars Plagioclase phenocrysts An(35–85) and microlites An(39–70) exhibit a wide compositional range from andesine to bytownite at ES. Plagioclase compositions vary between andesine to bytownite from HS, having anorthite contents of plagioclase with microcrysts and microlites of An(37–85) and An(52–72) respectively (Table 2). In the OZ, compositions of plagioclase phenocrysts range between andesine to bytownite An (35–78) for phenocrysts and microcrysts, and An(52–73) for microlites (Table 2). Plagioclases of KA samples display a wide compositional range between andesine and bytownite, with An (34–76) for phenocrysts and microcrysts, and An (45–70) for microlites (Table 2). In all four volcanic areas plagioclases exhibit both normal and reverse zoning (Table 3). 6.4. Olivine Embayed and euhedral olivines coexist in all basalt samples, in addition to normal-zonning with Fo-rich cores. Olivines in ES rocks exhibit a wide compositional range of. Fo(57–86) for phenocrysts and microcrysts, and Fo(66–83) for microlites (Table 2). Most olivines are normally-zoned with respect to Mg and Fe, while some crystals exhibit reverse zoning. In HS rocks, the compositional range of olivine is Fo(68–86) for phenocrysts and microcrysts, and Fo(53–75) for microlites. In all OZ basalts, olivine is the most abundant mineral phase with the composition varying between Fo(59–90) for phenocrysts and microcrysts, and Fo(57–87) for microlites. In comparison the composition of olivine ranges between Fo(77–89) for KA samples. 6.5. Oxide minerals Oxide minerals, such as magnetite, ilmenite and Cr-spinels, are generally observed as microlites in the groundmass of ES basaltic rocks (Table 2). In samples from HS and OZ, oxides occur as magnetite and ilmenite microlites. Oxide minerals occur as magnetite, Cr-spinel and Alchromite microlites in OZ and KA samples. 7. Geochemistry 7.1. Major element characteristics and classification of the basaltic rocks Major element results of ES, HS, OZ and KA basaltic volcanic centers are shown in Table 6. Most of the ES samples are calc-alkaline whereas those of HS, OZ and KA are alkaline (Fig. 2). The majority of the samples are basalt and basaltic-andesite, and few classified as trachybasalt and basaltic-trachyandesite according to the international classification of Le Bas et al. (1986) (Fig. 2). The LOI values vary between 0.43% and 0.58% (Table 6). All selected samples are basaltic rocks with compositional range of 48–56% SiO2. MgO contents range between 4 and 7 wt%, with Mg# between 53 and 69. Trachybasalt and basaltictrachyandesite samples are hawaiites (Na2O – 2 N K2O) and mugearites (Na2O - 2 N K2O), respectively. The subalkaline samples of ES are calc-

Fig. 3. a) K2O (wt%) vs. SiO2 (wt%) diagram (Peccerillo and Taylor, 1976) b) K2O (wt%) vs. Na2O (wt%) diagram (Middlemost, 1975). Symbols are as in Fig.7.

alkaline (Fig. 3a), while the alkaline samples of HS, OZ and KA exhibit a sodic character (Fig. 3b). According to the Al2O3 N Na2O + K2O + CaO values, all basaltic volcanic rocks display strongly peraluminous features. Harker diagrams for major element oxides indicate that SiO2, Al2O3 and Na2O values increase with decreasing MgO content; however, CaO follows a positive trend (Fig. 4). 7.2. Trace element characteristics Trace element analyses of the basaltic rocks are listed in Table 7, and Harker diagrams for selected trace elements versus Th are plotted in Fig. 4. The samples of ES and HS display a positive trend for the incompatible elements U and Ba, whereas Y and Eu decreasing with increasing Th. Conversely, OZ and KA samples lack systematic trends. The corresponding multi-element plots and the chondrite-normalized rareearth element (REE) patterns are shown in Fig. 5. Multi-element patterns of ES samples show enrichment in LILE (Rb, Ba, Th, U and Pb) and depletion in Nb, Ta and Ti. Only the oldest sample ER 13, (1700 ka old trachybasalt) exhibits the opposite trend with Rb, Ba, U and Th depletions and Y, Yb and Lu enrichments. HS samples are enriched in LILEs and show negative anomalies for Nb, Ta and Ti, while having a positive Pb anomaly. Samples from OZ are enriched in LILEs, showing higher negative anomalies for Nb, Ta and Ti than the HS samples. The MgO-enriched OZ basalt samples HA 16 (797 ka) and HA 17 (66 ka) are characterized by high negative anomalies for Nb, Ta and Ti. KA samples are enriched in LILE and show negative anomalies for Nb, Ta and Ti. With the exception of the oldest sample (ER13), all ES samples are significantly enriched in LREE with their (La/Yb)N varying from 4.7 to 10.3 (Fig. 5). Sample ER13 is more enriched in heavy rare earth elements (HREE) than other ES rocks. HS samples exhibit an enrichment in LREE and have (La/Yb)N ratios between 6.6 and 9.5.

64


Fig. 4. Selected major (wt%) and trace element (ppm) Harker variation diagrams of the ES, HS, OZ and KA.

The OZ samples are characterized by an enrichment in LREE with (La/Yb)N of between 6.9 and 9.7. The REE patterns of the KA samples are enriched in LREE and exhibit higher (La/Yb)N than those from ES, HS and OZ, ranging from between 7.8 and 10.6. A significant Eu anomaly has not been observed in any samples from ES, HS, OZ and KA. Different geochemical signatures for the four volcanic areas are displayed on the Zr/Nb vs. La/Nb diagram (Fig. 6). ES and KA rock compositions do no overlap. ES data exhibit more variations in Zr/Nb, while KA samples show more variations in their La/Nb ratios. Trachy-basalt ER13 possesses the highest Zr/Nb among all basaltic samples. HS and OZ basalts show scattered distributions with variations in both ratios. As a reference, low La/Nb (0.64–1.19) and Zr/Nb (3.2–11.4) are observed in HIMU-type (high U/Pb), EMI and EMII-type OIB basalts, where La/Nb and Zr/Nb ratios in MORB vary from 1.07 and 30, and in

the continental crust from 2.2 and 16.2 (Rudnick and Gao, 2003). The La/Nb of ES, HS and OZ rocks plot between enriched mantle (EMI) and crustal components, whereas KA samples have higher ratios than crustal components.

8. Discussion 8.1. Spatio-temporal evolution The majority of alkaline basalts and basaltic andesites from HS, OZ and KA, and the calc-alkaline basaltic andesites from ES have been emplaced contemporaneously during the period of post-collisional Quaternary Central Anatolian basaltic volcanism (Table 1). The La/Nb vs. age diagram in Fig. 7a indicates that:


65

Fig. 5. Primitive mantle-normalized (PM; Sun and McDonough, 1989) multi-element diagrams of the a) ES, (c) HS, (e) OZ, g) KA and Chrondrite-normalized spidergrams of the b) ES, d) HS, f) OZ, h) KA.

(1) Except for the oldest basaltic volcanic rock (ER13) from western flank of ES, contemporaneous alkaline and calc-alkaline basaltic rocks occur. (2) The KA basaltic volcanic field has the highest La/Nb, and thus separates itself from the other three studied volcanic areas (Fig. 7b).

(3) The basaltic rocks of OZ have similar La/Nb and are all older than N300 ka, with the exception of sample HA16 (from south of Gaffar Hill). (4) The majority of the basaltic rocks of ES and HS are younger than ˂350 ka and ˂100 ka, respectively.

66


Fig. 6. Zr/Nb vs. La/Nb diagram. Symbols are as in Fig. 4.

The K\\Ar ages of basaltic rocks from all four volcanic areas show that spatial diversity is more pronounced than temporal evolutionary trends. The trend between geochronology and geochemistry can only be seen in trace element ratio (La/Nb) from samples HA18, HA19, HA20 located northwest of the HS complex (Fig. 1c). This study has revealed much younger ages of eruptive products from the ES, HS and KA complexes. The exposure ages from north of ES (Sarıkaya et al., 2017) are compatible with our new findings. These new ages reveal that historical eruptions have occurred in Cappadocia. Such data are very important for assessment of volcanic hazards as mentioned before in Schmitt et al. (2014). 8.2. Interpretation of the mineral data The common mineral assemblage of ES, HS, OZ and KA basaltic rocks consists of plagioclase + olivine + clinopyroxene ± orthopyroxene and oxides. Opx phenocrysts occur in calc-alkaline mafic and intermediate lavas and not in alkaline lavas. Overall ES samples have both cpx and

opx, while samples having either cpx or opx crystals are mainly located in the north and northwest of the ES. The 707 ka old sample ER32 has only opx in its mineral assemblage,however it is located east of the main edifice (Fig. 1b). This information does not provide any indication that there are systematic relationships between locations, ages and the mineral assemblages of ES rock samples. 8.2.1. Mineral-melt equilibrium Magmatic systems are sensitive to changes in the composition of phenocrysts and in the mineral-liquid equilibrium. If phenocrysts are in equilibrium with the liquid, they can be used to estimate the temperature (and sometimes pressure) of the magma (Damasceno et al., 2002). During magma storage and ascent, melt composition changes occur due to fractional crystallization or magma mixing, thus affecting the physical conditions of crystallization due to cooling, decompression or magma mixing (Aldanmaz, 2006). Mixing processes can be determined by compositional and textural zoning of phenocrysts (Hibbard, 1981; Davidson et al., 1998; Ginibre et al., 2002; Troll and Schmincke,

Fig. 7. a) La/Nb vs. Age (ka) all basaltic volcanic rocks from ES, HS, OZ and KA b) La/Nb vs. Age (ka) except the oldest basaltic volcanic rock (ER13) from ES. Symbols are as in Fig. 4.

G.D. Dogan-Kulahci et al. / Journal of Volcanology and Geothermal Research 356 (2018) 56–74 Table 5 Estimated temperature (°C), pressure (kbar) and depth (km) values for ES, HS, OZ and KA rock samples, calculated after Putirka (2008). Studied Area

ES HS OZ KA

Temperature (°C)

Pressure (kbar)

Eq. (32d-33-34) cpx-liq ± 45

Eq. 23 plg-liq ± 43

cpx-liq ± 3

1138–1167 1208 1120–1233 1172

1146–1183 1250 1164–1207 1171–1228

4–5 5 8 5

Depth (km)

11–17 18 28 21

2002). Several petrographic features as well as compositional zoning in phenocrysts provide clues about the chemical and physical conditions of the melt. The occurrence of both equilibrium and disequilibrium minerals in samples from all four volcanic centers, emphasizes that the formation conditions are variable and related to several processes. Whereas coexisting normal and reverse-zoning in plagioclase and pyroxene indicate to disequilibrium conditions. Sieve-textured plagioclases are the petrographic signs of the variable conditions. According to Nelson and Montana (1992) sieve-textured plagioclases can be related to rapid decompression rather than magma mixing. Rapid magma ascent may preserve melt inclusions as pockets of glass within phenocrysts (Gill, 2010).

8.2.2. Clinopyroxene-liquid equilibrium According to experimental studies (Grove and Bryan, 1983; Toplis and Carroll, 1995), the distribution coefficient between clinopyroxene and basaltic melts is 0.23 ± 0.05 (Hoover and Irvine, 1977; Damasceno et al., 2002). In Fig. 10a, the minority of the samples of ES, OZ and KA are in the equilibrium field.

8.2.3. Plagioclase-liquid equilibrium The plagioclase-equilibrium conditions are assumed by Sisson and Grove (1993); Feeley and Dungan (1996) and Ferla and Meli (2006) to be almost dry (KD = 1–3) and at low pressures (b2 kbar).The composition of the majority of the plagioclase phenocryst cores in ES, OZ and KA samples lie in the liquid equilibrium field (Fig.10b).

67

8.2.4. Estimated temperatures and pressures Equilibrium cores of clinopyroxene and plagioclase phenocrysts were used to estimate temperatures, pressures and depths (Table 5). Instead of glass compositions, the whole rock compositions were used for liquid composition. The geothermometer calculations were undertaken according to Putirka (2008). The highest cpx-liq temperature of 1120 to 1233 °C and pressure of 8 kbar were estimated for the OZ complex, whereas the lowest plg-liq temperature of 1146 to 1183 °C and pressure of 4–5 kbar were given by ES samples. Estimated crystallization depths for ES are the lowest (11–17 km) and for OZ the highest (19–28 km). The mean value of estimated crystallization depths are increasing from NE to SW, i.e. from the ES to KA complex. Thus, the ES magma chamber, located in the pull-apart basin, lie relatively close to the surface. Therefore, the differences between the four volcanic areas (ES, HS, OZ and KA) may be explained by multistage crystallization of minerals in several magma chamber(s) located at different depths.

8.3. Magmatic changes in the mantle sources Recent studies on the Anatolian, eastern and western Mediterranean area revealed various scenarios regarding crustal thickness, uplift and uplift time, subduction, slab detachment, slab delamination, slab rollback, and slab break-off processes, in addition to mantle-crust relation and asthenosphere-lithosphere interaction (Govers and Wortel, 2005; Biryol et al., 2011; Cosentino et al., 2012; Schildgen et al., 2012; Faccenna et al., 2006, 2013; Aydar et al., 2013; Vinnik et al., 2014; Delph et al., 2015; Kind et al., 2015; Govers and Fichtner, 2016; Reid et al., 2017; Göğüş et al., 2017).The Quaternary basaltic rocks of this study show a complex geochemical pattern. A low La/Nb ratio is indicative of within-plate enrichment or an asthenospheric source, whereas a high La/Nb ratio (N1) indicates subduction enrichment or a lithospheric source (De Paolo and Daley, 2000; Huang et al., 2000). High Nb/La (N1; Abdel-Rahman and Lease, 2012) and Nb/U (~47; Hofmann, 1997) values are typical characteristics of rocks that evolved from an OIBlike asthenospheric mantle source and that were not contaminated by the continental crust. In the ES, HS, OZ and KA rocks Nb/La ratios vary between 0.47 and 0.66, 0.45–0.58, 0.43–0.61, 0.28–0.38, whereas Nb/U ratios vary between 6.90 and 23.34, 12.43–21.01, 5.83–18.15, 4.08–8.86 for ES, HS, OZ and KA rocks, respectively. Primitive-mantle normalized multi-element concentrations (Fig. 5), negative Nb-Ta-Ti anomalies and positive Pb spikes suggest that the four studied volcanic

Fig. 8. Comparison of ES, HS, OZ and KA basaltic rocks with Global subducting sediment GLOSS (average) (Plank and Langmuir, 1998). Symbols are as in Fig. 4.

68


areas may have been affected by both crustal contamination and/or subduction related enrichment to different degrees. These anomalies increase gradually from the HS complex in the northest to the KA complex in the southwest. Negative Nb\\Ta and positive Pb anomalies can be interpreted as the result of two different processes: i) crustal contamination occurring at shallow depth, where subducted sediments are added to the mantle source and ii) subduction related fluid-induced processes (Lustrino and Wilson, 2007). There are significant differences between a source enriched by fluid components, which are separated from the subduction slab, and a source enriched by subducted sedimentary components (e.g., Hawkesworth et al., 1997; Elburg et al., 2002; Chiaradia et al., 2011). According to Chiaradia et al. (2011), enrichment in aqueous fluid-immobile incompatible elements (Th, La, Zr, Nb) suggests a significant difference in the slab component or in the mantle source. High Th/La ratios (0.28–0.36) and Th values (4–24 ppm) indicate the existence of a subducted sediment component in the mantle source (Plank, 2005). In ES samples, Th/La ratios vary between 0.08 and 0.24, and Th values range from 1.8–5.8 pmm. The Th/La values

Fig. 9. a) Th/Y-NbYb variation diagram of ES, HS, OZ and KA. b) Th/Yb-Ta/Yb variation diagram of ES, HS, OZ and KA. Data sources (Ormerod et al., 1991; Çubukçu, 2008; Özdemir, 2011; Keskin et al., 2012; Ali and Ntaflos, 2011; Tang et al., 2006; Downes et al., 1995; Alıcı Şen, 2002).

from HS, OZ and KA samples are generally lower (0.13–0.35) and Th values are higher (3.5–9.2 ppm) than the above given reference ranges by Plank (2005). Subduction-related magmas modified by subducted sediments are separated from magmas affected by fluid metasomatism at the mantle source, due to the high Th/Ce N0.15 (Hawkesworth et al., 1997; Guo et al., 2006) and low Ba/La (Sheppard and Taylor, 1992; Hawkesworth et al., 1997) ratios. The values of Th/Ce and Ba/La for ES are 0.04–0.12 and 10.04–14.30 respectively. Whereas, the Th/Ce values for HS, OZ and KA are 0.07–0.09, 0.07–0.18 and 0.11–0.16 respectively; and 12.35–13.28, 11.18–16.81 and 11.48–12.45.These ratios show that the ES magma was less affected by subducted sediments than HS, OZ and KA magmas. Ba/La vs. Th/Yb diagrams for all studied volcanic complexes (Fig. 8) confirm that KA samples follow a sediment input trend, while the Th/Yb and Ba/La values are generally lower than the Global subducting sediment (GLOSS(average)). Nb/Y and Th/Y ratios are geochemical tracers for mantle source heterogeneity and crustal contamination effects (Pearce, 1983), while Th/Nb ratio is a good tracer for arc lavas

Fig. 10. a) Variation of clinopyroxene compositions versus whole rock geochemistry in a cpx - melt Fe/Mg equilibrium diagram. Whole rock Mg-numbers versus Mg-numbers of cpx Mg-number = 100*Mg+2/(Mg+2 + Fe+2) is calculated assuming Fe+3/Fe+2 = 0.1. The equilibrium field for Fe/Mg exchange between cpx and the melt (0.23 ± 0.05; Toplis and Carroll, 1995) is shown. Arrows indicate the relative effects of xenocryst addition and groundmass crystallization. b) Plagioclase - melt equilibrium diagram; KD values are from Sisson and Grove (1993).


(Brenan et al., 1995; Elliot et al., 1997). If the mantle was enriched by subduction or the mantle derived magma was affected by crustal contamination, Th/Y should exceed Nb/Y. However, if the mantle was enriched by an earlier and weak partial melting process, it could exhibit withinplate enrichment trends having high Th/Y and Nb/Y ratios. All basaltic rocks from ES, HS, OZ and KA are above the Th/Nb = 0.1 (Fig. 9a). Accordingly, high Th/Y and Nb/Y ratios in HS and OZ samples are indicative of within-plate enrichment, whereas increasing Th/Y and decreasing Nb/Y in KA suggests mainly subduction enrichment. Recent studies investigated the geophysical context of Arabian, Anatolian and Aegean microplate interaction (e.g., Biryol et al., 2011; Cosentino et al., 2012; Faccenna et al., 2013; Delph et al., 2015; Kind et al., 2015). The Afar plume is the most important model for the Arabian and the Anatolian systems, and hence its effect on Anatolia needs to be determined (Faccenna et al., 2013). According to Cosentino et al. (2012) and Faccenna et al. (2013), a slab window formed at ~8–10 Ma widened westward, and may be the reason for hot material migration from Arabia towards the southeast of Turkey. The Karacadağ volcanic complex, located in the southeast region of Turkey, displays both asthenospheric and enriched sub-continental lithospheric mantle source signatures which can be explained by the Afar plume (Keskin et al., 2012). However, in the Th/Yb vs. Ta/Yb diagrams of Fig. 9b, Karacadağ and ES, HS, OZ and KA basaltic rock data do not follow the same trends. Reid et al. (2017) presented data related to the lithosphere-asthenosphere interaction in Central Anatolia using shear velocity profiles. According to their profiles,

69

the positive correlation between depth and shear velocity below ~45 km could be explained by upwelling asthenosphere undergoing decompression melting, and depths of 60–70 km could be the regional depth for stagnation of a rising mantle and melt accumulation zone in Central Anatolia (Reid et al., 2017). If this is the base of the mantle lithosphere relict after slab rollback-induced delamination, this residual mantle lithosphere might have undergone small-scale convection and associated decompression melting. This convection could permit intermittent mantle upwelling to shallow depths and draining of decompression melts. Consequently, basalts in Central Anatolia, especially the HS could have been hybridized by both subducted modified lithosphere and shallow convecting asthenosphere components. This study and tomographic images of recent studies neither confirm the occurrence of a present-day mantle plume nor the contemporaneous subduction beneath Central Anatolia (Biryol et al., 2011). Bartol and Govers (2014) stated that on geological time scales, delamination progress is too fast to leave a migration footprint. The geochemical data presented here suggests that the four volcanic areas studied have mainly evolved from a lithospheric mantle source, and traces of the asthenospheric mantle may have affected the lithospheric mantle. As a general observation, mafic and felsic volcanic rocks in Central Anatolia (Niğde Volcanic Complex, Acigöl) isotopically and geochemically resemble each other and fractional crystallization and crustal contamination play a role in their genesis (Siebel et al., 2011; Aydın et al., 2014). Besides, it was confirmed that parental magmas derived from an isotopically

Fig. 11. a) Yb vs. La trace element diagram (Garnet peridotite data was from Frey (1980) and spinel peridotite data from McDonough, 1990). In order to minimize the effects of fractionation, the compositions of the ES samples were corrected to primary magma compositions by using the least-squares linear regression method proposed by Turner and Hawkesworth (1995). A fractionation correction has been carried out at a constant MgO of 7% for ES samples b) Al2O3/CaO vs. SiO2 variation diagram of ES, HS, OZ and KA.

70


8.4.2. Fractional crystallization The Harker diagram trends reveal olivine, plagioclase and pyroxene fractionation (Fig. 4). To further quantify fractionation fractional crystallization (FC) modeling was undertaken using least-squares mass balance calculations (Byran et al., 1969). For modeling purposes, samples with high Mg#, Cr, Ni, Fe2O3 and MgO contents as well as low SiO2 were selected as end members, as the element contents of these samples were assumed to represent the original liquid. The fractionation degrees of plagioclase, clinopyroxene and olivine were calculated in percentages (Table 8). Fig. 11b shows Al2O3/CaO correlation indicating that clinopyroxene fractionation is more effective than plagioclase fractionation. Furthermore, plagioclase fractionation has been more effective in HS than at the other areas where clinopyroxene fractionation is dominant. According to the FC calculations, OZ samples exhibit little evidence for plagioclase fractionation, whereas clinopyroxene and olivine fractionation can be observed (Fig. 11b). The results indicate that FC modeling is compatible with the data because the residual sum of squares (∑R2) is lower than one. Samples HA31 (HS), HA27 (OZ) and KA19 (KA) have the lowest Mg# values (60–61), whereas OZ sample HA16 and HA17 have the highest Mg# values (73 and 75, respectively; Table 6). According to Frey et al. (1978), Mg# values between above 68 and 70 are related to primary magmas. Considering that many of the samples have similar Mg# values, it can be implied that the effect of fractionation was limited and/or negligible in those cases. Fractional crystallization was more effective at ES and HS stratovolcanoes than at the OZ and KA basaltic volcanic fields.

depleted and chemically enriched lithospheric mantle, which was metasomatized by previous slab-derived fluids, with a lesser amount of asthenospheric mantle material (Aydın et al., 2014).

8.4. Effects of the magmatic processes 8.4.1. Partial melting Enrichment in LILE and LREE can be explained by subduction, crustal contamination or low degree partial melting (Pearce et al., 1990). Generally, ES, HS, OZ and KA basaltic rocks have high LILE and LREE contents, but low HFSE and HREE content (Fig. 5). This may be related to magmatic differentiation processes, fractional crystallization, magma mixing, crustal contamination, or may show the tendency of LILEs and LREEs to partition into the liquid phase during partial melting. A higher or lower LREE/HREE ratio provides information about the degree of melting at the source. In addition, this ratio also strongly depends on the minerals in the restite. For low degrees of partial melting, low Zr/Nb, Zr/Ba and high Th/Y, Ta/Th and La/Nb ratios are expected, whereas high Zr/Nb and Th/Y,low Zr/Ba and Ta/Th, and moderately elevated La/Nb ratios are observed in KA. Multi-element diagrams (Fig. 5) lack negative anomalies for the HFSE in all the rock samples. All Dy/Yb ratios are ˂ 2, which indicates that garnet was absent in the restite (Jung et al., 2006). Yb vs. La (Fig. 11a) bulk melting modeling was conducted using spinel and garnet peridotite as starting components (Shaw, 1970). Hence, calculated degrees of partial melting vary between 6 and 15% for ES 8–9% for HS, 8–12% for OZ and 8% for KA rock samples. These values are similar to spinel peridotites, lying between the spinel peridotite and garnet peridotite curves. Only ES sample ER13 plots close to the garnet peridotite curve with a partial melting degree of 8–10%.

8.4.3. Crustal contamination La/Nb N1.5 is often used as an indicator for crustal contamination (Hart et al., 1989). La/Nb ratios range between 1.52 and 2.73 (ES),

Table 6 Major element analysis of ES, HS, OZ and KA volcanic rocks. Major element (wt%)

Erciyes stratovolcano ER1

ER2

ER3

ER10

ER11

ER13

ER17

ER20

ER26

ER27

ER29

ER30

ER32

SiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O TiO2 MnO P2O5 LOI Mg# Total

54.34 17.17 6.88 6.10 8.32 3.48 1.24 0.96 0.11 0.36 0.23 69.06 99.19

55.33 17.32 6.97 5.62 8.13 3.66 1.37 1.04 0.11 0.33 0.28 66.99 100.15

53.13 16.52 8.24 6.54 8.54 3.62 1.02 1.30 0.13 0.37 −0.16 66.61 99.25

53.09 16.95 8.39 6.65 8.54 3.81 0.96 1.36 0.13 0.40 −0.10 66.60 100.18

56.70 17.65 7.22 4.04 7.06 4.08 1.24 1.05 0.12 0.45 0.30 59.28 99.91

48.89 17.06 11.99 5.44 8.46 4.17 0.94 2.40 0.18 0.51 −0.32 53.29 99.73

55.46 17.57 7.27 4.95 7.54 3.97 1.54 1.11 0.12 0.38 0.30 63.17 100.20

54.53 17.56 7.24 5.88 8.22 3.70 1.36 1.10 0.12 0.36 0.29 67.13 100.36

56.43 17.98 6.57 4.59 8.20 3.80 1.31 1.04 0.11 0.34 0.08 63.72 100.42

55.01 17.42 6.81 5.62 8.28 3.64 1.14 1.01 0.11 0.33 0.42 67.27 99.79

52.45 17.68 8.96 4.78 7.54 4.31 1.65 1.54 0.15 0.66 0.15 58.18 99.87

55.07 17.40 7.94 4.84 7.42 4.06 1.62 1.39 0.13 0.41 0.50 61.34 100.79

56.22 17.78 6.55 5.12 7.93 3.65 0.95 0.94 0.10 0.28 0.58 66.31 100.09

Major element (wt%)

SiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O TiO2 MnO P2O5 LOI Mg# Total “b.d.” below detection.

Hasandag stratovolcano

Obruk-Zengen

Karapınar

HA13 b

HA18

HA19

HA20

HA23b

HA31

HA14

HA15

HA16

HA17

HA26

HA27

KA18

KA19

KA21

49.33 16.08 8.88 8.52 9.78 3.44 1.15 1.19 0.15 0.44 −0.14 69.16 98.72

51.06 16.86 8.72 7.31 8.76 3.86 1.46 1.43 0.14 0.43 −0.43 67.82 99.88

50.43 16.56 8.87 7.99 9.42 3.64 1.26 1.33 0.15 0.44 −0.58 67.79 99.72

50.36 16.25 9.14 8.75 9.85 3.64 1.12 1.21 0.15 0.43. −0.27 69.11 100.64

50.90 16.22 8.86 7.35 9.32 3.68 1.10 1.23 0.15 0.37 b.d. 65.98 98.97

49.24 17.80 9.57 6.01 10.68 3.70 0.97 1.34 0.16 0.38 −0.06 59.50 99.51

48.99 16.24 9.01 8.52 9.28 3.65 1.45 1.40 0.15 0.43 −0.35 70.41 98.69

49.38 16.40 9.42 8.71 8.98 3.71 1.23 1.50 0.15 0.41 −0.28 68.37 99.32

48.81 15.79 8.97 10.18 10.43 3.43 0.80 1.10 0.15 0.37 −0.32 72.62 99.75

48.88 15.30 8.60 10.92 10.86 3.00 0.88 1.06 0.14 0.25 −0.23 74.80 99.83

49.39 16.23 9.39 8.81 9.47 3.73 1.15 1.49 0.15 0.41 0.22 68.68 100.42

56.24 17.08 7.31 4.07 7.57 4.19 1.86 1.25 0.13 0.31 0.24 59.19 100.23

49.75 16.48 8.53 8.12 11.02 3.36 0.95 1.03 0.14 0.30 −0.32 68.99 99.35

50.87 17.65 9.17 5.69 9.52 3.96 1.13 1.27 0.15 0.35 −0.38 60.97 99.38

52.40 15.99 8.20 8.54 9.11 3.50 1.12 1.00 0.14 0.33 0.12 72.38 100.43

4.46 2.57 1.54 4.63 3.13 0.95 28.31 0.41 10.76 25.20

6.06 6.29 23.06 2.62

142

3.84 2.19 1.31 4.04 2.82 0.79 29.92 0.36 9.96 22.97

4.17 6.06 20.85 2.23

134

370 56.03 33.44 288 3.50 2.01 1.29 3.76 2.97 0.70 29.72 0.32 8.41 23.48 135 6.91 6.39 27.01 2.02 66.02 117 346 56.15

97

4.63 4.09 13.20 1.85 4.40 5.83 17.10 2.18

166

5.10 7.73 23.40 2.32

186

4.86 5.47 12.20 2.19

3.39 1.91 1.15 3.32 2.21 0.69 18.07 0.28 7.78 16.96 3.74 2.17 1.36 3.82 2.65 0.79 25.94 0.35 11.72 21.37 4.14 2.23 1.61 4.44 3.43 0.83 24.20 0.35 14.88 23.73 4.47 2.38 1.83 5.13 3.84 0.90 31.46 0.36 13.53 31.85

126 'na' not applicable.

KA19

343 56.34

433 50.64 23.25 44 4.11 2.36 1.42 4.43 3.87 0.82 25.75 0.38 12.16 22.93 27 8.72 5.41 44.02 2.41 67.66 162 253 48.21 38.87 296 4.35 2.44 1.73 5.03 3.73 0.85 22.50 0.37 12.44 25.64 156 5.02 5.64 17.24 2.35 93.15 164 260 35.80

KA18 HA27 HA26 HA17 HA16

290 49.87 301 50.26

HA15 HA14

360 66.02

350 53.74 32.99 20 4.42 2.49 1.64 4.79 3.15 0.87 28.14 0.38 12.70 24.80 49 5.58 6.41 19.51 2.45 92.22 135 349 54.06 35.42 275 4.41 2.45 1.62 4.65 3.64 0.85 27.73 0.38 16.09 25.60 145 5.52 6.54 21.06 2.42 87.10 173 376 53.92 33.36 206 4.27 2.33 1.63 4.64 3.56 0.82 28.30 0.36 16.25 25.32 128 5.72 6.43 26.91 2.32 80.75 168

HA19 ER32

252 38.13 23.98 115 3.30 1.81 1.21 3.57 3.40 0.65 20.00 0.28 9.44 17.92 97 6.83 4.54 20.21 1.81 69.01 154 413 45.77 24.48 86 4.21 2.37 1.60 4.59 3.79 0.83 22.86 0.37 14.92 22.47 57 6.20 5.61 26.18 2.39 86.46 173

ER30 ER29

520 70.27 28.12 56 4.73 2.54 1.92 5.41 4.34 0.92 36.35 0.39 23.95 31.36 46 8.31 8.30 29.30 2.55 94.67 213 247 39.43 26.48 133 3.13 1.70 1.14 3.51 3.10 0.62 19.70 0.27 10.06 17.63 85 6.06 4.63 25.37 1.71 63.33 128

ER27 ER26

292 46.57 21.51 77 3.53 1.94 1.29 3.81 3.49 0.70 24.50 0.31 12.83 20.48 46 7.34 5.39 29.62 1.98 73.38 161 296 47.91 26.73 144 3.71 2.07 1.39 3.99 3.79 0.73 24.91 0.33 13.76 21.68 97 6.57 5.63 23.52 2.11 75.45 181

ER20 ER17

348 52.51 24.12 106 4.14 2.33 1.46 4.50 4.45 0.83 27.11 0.37 14.76 23.47 72 7.72 6.12 27.28 2.39 78.17 214 228 48.59 37.90 27 6.04 3.47 2.13 6.29 5.36 1.23 21.92 0.52 13.07 26.59 38 3.88 6.10 8.51 3.38 116 265

ER13 ER11

356 54.56 20.70 61 4.13 2.28 1.54 4.59 4.13 0.81 28.05 0.36 15.06 24.53 49 8.23 6.12 24.29 2.31 84.68 194 226 45.35 32.14 187 4.57 2.54 1.57 4.76 4.09 0.90 22.49 0.39 11.39 22.32 123 5.68 5.30 18.49 2.52 85.69 194 261 46.09 na na 4.37 2.46 1.34 4.47 3.72 0.92 21.96 0.38 12.12 21.79 na 7.70 5.43 18.15 2.33 na 192

ER10

The basaltic volcanic rocks of ES, HS, OZ and KA exhibit basic to intermediate compositions (48 b SiO2 b 56). Coeval basaltic rocks exhibit a calc-alkaline affinity in ES and alkaline compositions in HS, OZ and KA complexes. The majority of the contemporaneous alkaline and calcalkaline basaltic rocks have been dated to the period of post-collisional volcanism in Central Anatolia and are younger than 600 ka. The mineral assemblage of ES, HS, OZ and KA samples are plagioclase + olivine + clinopyroxene ± orthopyroxene + oxide minerals. The pressuretemperature and depth calculations reveal multistage crystallization in magma chamber(s) located at different depths, with the depth of ES magma reservoir shallower than those of HS, OZ and KA. The average value of calculated magma storage depths increase from NE to SW in Central Anatolia. Additionally, the oldest alkaline basalts can be found in the OZ complex in SW Central Anatolia, for which the highest cpx-liq temperature, pressure and depth were estimated. This may be explained by the deeper magma chamber(s) that were active at different periods during the Quaternary period. Harker variation diagrams and least-squares mass balance calculations show that fractional crystallization is the most dominant process at ES. All ES, HS, OZ and KA basaltic rocks are enriched in LILE and LREE. Lack of negative anomalies for HFSE (Y, Yb) and the La/Nb N1, favor a shallow lithospheric source. Zr/Ba ratio is a parameter used to distinguish lithospheric sources (Zr/ Ba: 0.3–0.5), and asthenospheric sources (Zr/Ba N0.5), (Menzies et al., 1991; Kürkçüoğlu, 2010). The Zr/Ba ratio ranges between 0.4 and 1.2 for ES, 0.38–0.51 for HS, 0.38–0.65 for OZ and 0.32–0.41 for KA samples implying foremost a lithospheric source for basaltic rocks in Central Anatolia. The negative anomalies of Nb-Ta-Ti and the positive Pb spikes suggest that rocks of these volcanic areas were affected by crustal contamination and/or subduction. Our geochemical data and published mantle tomography argue against the presence of a mantle plume below Central Anatolia. This is because there is no discernible spatio-temporal evolution of the Central Anatolian mafic volcanism, and magma generation appears to have been controlled by similar processes for at least the past 600 ka. Previous data indicate that the effect of subduction and/or crustal contamination in Central Anatolia decreased from Miocene to Quaternary, but our Quaternary samples show evidence for sediment input rather than for slab-derived fluids. The geochemical footprints of this study suggest that ES, HS, OZ and KA basaltic rocks have evolved mainly from a heterogeneous lithospheric mantle source. While the asthenospheric mantle may have also affected the lithospheric mantle; if not directly, through percolations as supported by Reid et al. (2017). According to the geochronological and geochemical data, alkaline and calc-alkaline series basaltic rocks in Central Anatolia are coeval and the magmatic evolution of the four studied volcanic areas are heterogeneous.

383 58.90 38.05 346 4.35 2.38 1.64 4.72 3.49 0.85 31.01 0.37 14.88 26.83 177 6.03 6.89 22.47 2.33 88.27 159 331 49.52 na na 3.68 2.05 1.22 3.88 3.53 0.76 25.04 0.33 14.85 20.96 na 9.01 5.47 27.44 1.98 na 175 344 50.63 na na 3.43 1.89 1.18 3.56 3.15 0.70 27.34 0.30 15.99 19.69 na 9.94 5.31 23.27 1.91 na 153

ER 3

9. Conclusions

Acknowledgements

Ba Ce Co Cr Dy Er Eu Gd Hf Ho La Lu Nb Nd Ni Pb Pr Rb Yb Zn Zr

HA20 HA18 ER 2 ER 1

HA23b

HA31

Obruk-Zengen Hasandag stratovolcano Erciyes stratovolcano Trace element (ppm)

Table 7 Trace element analysis of ES, HS, OZ and KA volcanic rocks.

71

1.72–2.22 (HS), 1.63–2.33 (OZ), and 2.63–3.53 (KA), therefore crustal contamination has occurred within magma for all of the volcanic areas. The Th/Yb vs. Ta/Yb diagram in Fig. 9b reveals that both crustal contamination and subduction-related enrichment were effective in all four volcanic areas, but to varying degrees. For example, KA rocks show signs of both crustal contamination and subduction enrichment. The existence of enrichment in ES samples indicates, however, lithosphere-asthenosphere interaction rather than crustal contamination, which was more dominant in HS, OZ and KA.

308 48.59 35.24 286 4.51 2.55 1.55 4.81 3.61 0.90 23.99 0.41 11.73 23.39 130 7.18 5.39 16.51 2.59 83.90 157

Karapınar

KA21


All analyses were financed by TUBITAK (The Scientific and Technological Research Council of Turkey; Research Project 109Y064). This research is the integrated PhD thesis of the G.Deniz Dogan-Kulahci (G.D.D. K) from Hacettepe and Blaise Pascal Universities. G.D.D.K is grateful to Dr. Sarah B. Cichy from Postdam University, Dr. Kim Altinkaynak from

72


Table 8 Fractional crystallization (FC) test results for ES, HS, OZ and KA rock samples after Byran et al. (1969). Less differentiated samples ES ER3 ER20 ER27 ER17 ER1 HS HA20 HA23b HA13b HA20 HA20 HA19 OZ HA16 HA17 KA KA18

More differentiated samples

Olivine %

Orthopyroxene %

Clinopyroxene %

Plagioclase %

→ → → → →

ER11 ER17 ER32 ER11 ER27

11 0.94 0.25 2.62 1.45

2 – 3.4 1.27 –

7 3 – – 0.87

11 – 4.17 7 1.52

→ → → → → →

HA18 HA18 HA19 HA23b HA19 HA18

3.3 – 2.35 3.01 2.25 1.72

6.34 4.04 – 2.27 – 3.79

4.95 0.67 2.31 2.44 0.95 3.66

→ →

HA26 HA26

4.28 5.84

6.35 11.56

6.89 8.27

→

KA19

4.3

12.05

9.33

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Further reading http://helium.crpg.cnrs-nancy.fr/SARM/pages/roches.html. http://www2.jpl.nasa.gov/srtm/.