Ar-Ar dating and petrogenesis of the Early Miocene Taşkapı-Mecitli (Erciş-Van) granitoid, Eastern Anatolia Collisional Zone, Turkey

Accepted Manuscript Ar-Ar dating and petrogenesis of the Early Miocene Taşkapı-Mecitli (ErcişVan) granitoid, Eastern Anatolia Collisional Zone, Turkey Vural Oyan PII: DOI: Reference:

S1367-9120(18)30083-X https://doi.org/10.1016/j.jseaes.2018.03.002 JAES 3431

To appear in:

Journal of Asian Earth Sciences

Received Date: Revised Date: Accepted Date:

12 November 2017 22 February 2018 4 March 2018

Please cite this article as: Oyan, V., Ar-Ar dating and petrogenesis of the Early Miocene Taşkapı-Mecitli (ErcişVan) granitoid, Eastern Anatolia Collisional Zone, Turkey, Journal of Asian Earth Sciences (2018), doi: https:// doi.org/10.1016/j.jseaes.2018.03.002

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Ar-Ar dating and petrogenesis of the Early Miocene Taşkapı-Mecitli (Erciş-Van) granitoid, Eastern Anatolia Collisional Zone, Turkey

Vural Oyan1*

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Van Yüzüncü Yıl University, Faculty of Architecture and Engineering, Department of

Mining Engineering, Zeve Campus, 65080 Van, Turkey * Corresponding author: email: [email protected], tel: +90 432 2250128/1138 fax: +90 432 2251730

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Abstract The Early Miocene Taşkapı-Mecitli granitoid that is located in the northern section of the Eastern Anatolia Collision Zone has typical I-type, metaluminous and calk-alkaline characteristics. It also contains mafic microgranular / magmatic enclaves (MMEs). New ArAr dating results show that the age of the Taşkapı-Mecitli granitoid is ~23 Ma and it crystallised in the Early Miocene, in contrast to its previously known Cretaceous age. Identical crystallisation ages (23 Ma), similar mineral assemblages and geochemical compositions, and indistinguishable isotopic compositions of MMEs and host rocks imply that the MMEs are most consistent with a cumulate origin formed at earlier stages of the same magmatic system that produced the Taşkapı-Mecitli granitoid. MELTS modelling suggests that magma of the Taşkapı-Mecitli granitoid was the result of fractionation under a crustal pressure of 4 kbar, with a H2O content of 1.5%. ECAFC model calculation reveals that the Taşkapı-Mecitli granitoid includes from 0.5% to 2% crustal assimilation rates. These rates indicate that crustal contamination can be negligible when compared to fractional crystallisation in the evolution of the magma beneath the Taşkapı-Mecitli granitoid. The partial melting model calculations and MORB-normalised trace element concentrations of the least evolved samples of the Taşkapı-Mecitli granitoid are consistent with those of mafic melts obtained from partial melting of interacting mantlelower crust with a melting degree of 18%. The age (23 Ma) of the post- or syn-collisional Taşkapı-Mecitli granitoid suggests that the collision between Arabian and Eurasian plates could be before/around ~23 Ma (Late Oligocene to Early Miocene). Keywords: Taşkapı-Mecitli granitoid, mafic crust-mantle interaction, partial melting, timing of collision, Eastern Anatolia

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1. Introduction Granitoids are the main component of the continental crust, and they play an important role in the growth and development of the Earth's crust. Therefore, revealing the origin of granitoids can give key hints for understanding the geodynamic processes and lithospheric structures. Granitoids are classified as (1) mantle origin (Turner et al. 1992), (2) crustal origin (Chappell and White, 1992; Chappell, 1999) and (3) mixed origin that can be related to mixing of melts derived from both crust and mantle (Barbarin and Didier, 1992). Turkey is settled on the collision zone between Gondwanan and Eurasian plates, and it lies in important geodynamic zones: arc and syn-post collision tectonic settings. Hence, Turkey can be seen as a natural laboratory in geological terms. The Eastern Anatolia Collision Zone (EACZ), which is located on the eastern part of the Anatolide-Tauride block, is characterised by magmatic activities that continued from Early-Middle Miocene to historical times, following the continent-continent collision between the Eurasian and Arabian lithospheric plates (Şengör and Kidd, 1979). Şengör and Kidd (1979) mentioned that the collision between the Arabian and Eurasian plates started in the Middle Miocene; on the other hand, Okay et al. (2010) have suggested that the last oceanic lithosphere in the Eastern Anatolia region was exhausted 20 Ma ago. Lebedev et al. (2010) reported that volcanic activity in the EACZ began 15 Ma ago in the neovolcanic areas to the north of Lake Van. The products of the volcanism have covered the majority of the EACZ and they have generally masked older units such as ophiolitic mélange, sediments and granitoids. Consequently, the granitoid bodies are observed in limited areas in the EACZ, especially to the north of Lake Van. Most of the previous studies of the Eastern Anatolian magmatism have examined the petrology of the collision-related volcanics (Innocenti et al. 1982; Pearce et al. 1990; Keskin et al. 1998; Keskin, 2003; Oyan, 2011; Özdemir and Güleç, 2014; Lebedev et al. 2016; Oyan

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et al. 2016). However, the granitic intrusions within the EACZ have not been investigated in detail. The Taşkapı-Mecitli granitoid (Erciş, Van) to the north of Lake Van (Eastern Turkey) (Fig. 1) intruded Eocene sedimentary units, ophiolitic mélange and metamorphic basement known as the Aladağ crystalline complex (Milanovskii and Koronovskii, 1973). In turn, it was covered by Miocene and Quaternary volcanics. However, studies carried out on the Taşkapı-Mecitli granitoid are very rare and they are generally related to petrography, geological mapping and geothermal resources (Karamanderesi et al. 1984). In particular, there are no geochemical, petrological or geochronological data available in the literature on the Taşkapı-Mecitli and other granitoids to the north of Lake Van. Although Karamanderesi et al. (1994) stated that the age of the Taşkapı-Mecitli granitoid is Eocene according to field relationships, the new Ar-Ar date obtained from this study indicates that this granitoid is Early Miocene in age (~23 Ma). This study aims to: (1) present new Ar-Ar age data, petrologic, geochemical and NdSr-Pb isotopic data of the Taşkapı-Mecitli granitoid; (2) exhibit the magma chamber processes, such as fractional crystallisation and crustal contamination, obtained from petrological models such as MELTS and EC-AFC; (3) discuss the mineralogical and chemical characteristics of the source region; and, (4) discuss the importance of these findings for the generation of this granitoid and for the geodynamic processes of the EACZ.

2. Geological setting The EACZ is located on the Alpine-Himalayan orogeny. The basement of the EACZ consists of micro-continents that were accreted to one other during a period from Late Cretaceous to Paleogene (Şengör et al. 2008). Ophiolite belts and accretionary complexes

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separate these micro-continents from each other. The tectonic blocks of the region are explained thoroughly in Şengör et al. (2008) and hence will not be reiterated here. The EACZ is a high plateau that has been uplifted by the collision of the Arabian and Eurasian plates. It has an average elevation of 2 km above sea level (Şengör et al., 2008). After this collision, intense tectonic movements occurred within the EACZ, including seismic activity, folding, reverse faulting and strike-slip faulting. The origin of the tectonic activity in the EACZ has changed from post-collisional convergence in the Early Miocene to tectonic escape in the Early Pliocene (Koçyiğit et al., 2001). The starting age of this collision is still controversial. The age ranges from the Late Cretaceous (Hall, 1976), through the Late Eocene (Allen and Armstrong, 2008) and to the Miocene (Şengör and Kidd, 1979; Robertson et al., 2007). Okay et al.’s (2010) recent study was based on apatite fission track ages of the Eocene sandstones from the Bitlis Zagros suture zone. Their findings suggest that the last oceanic lithosphere between the Arabian and Eurasian plates was consumed in the Early Miocene (20 Ma). However, Karaoğlan et al. (2016) have reported that the timing of this collision needs to be placed in the Oligocene to Early Miocene on the basis of the fission track ages of the granitoids. Recently, geophysical studies in the EACZ have revealed that the crust of Eastern Anatolia has an average thickness of 45 km (Zor et al., 2003). In light of this data, Şengör et al. (2008) reported that the lithospheric mantle under the EACZ is absent and that hot asthenospheric melts settled directly below the crust. Keskin (2003) proposed that the Late Cenozoic magmatism that can be observed across the region could be produced with the slab stepping and breakoff model. Meanwhile, Özacar et al. (2006) and Angus et al. (2006) indicated that the lithosphere under the region has an average thickness of 70 to 75 km. In that case, if the lithospheric crust is thought to have a thickness of 45 km, then the lithospheric mantle under the EACZ has a thickness of around 20 to 25 km. More recently, studies on the Late Cenozoic magmatic activity in the region have proposed a model that

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contains a mixture of the melts derived from both the lithospheric and asthenospheric mantle (Özdemir and Güleç, 2014; Oyan et al., 2016 and 2017). This study presents the first geochemical, petrologic and geochronological evaluation of the Taşkapı-Mecitli granitoid that is situated on the culmination of the Lake Van regional dome structure that was identified by Şengör et al. (2008). The study area is located to the north of Lake Van in the EACZ and the basement rocks of the area mainly consist of Palaeozoic metamorphics, Eocene sandstone, and Upper Cretaceous ophiolitic mélange. These basement lithologies have been cut by granitoids and they are overlain by Middle-Late Miocene to Pliocene-Quaternary volcanic and volcanoclastic rocks.

3. Analytical techniques A total of 55 samples were collected from the Taşkapı-Mecitli granitoid, and 39 fresh samples were selected for major, trace and rare earth element (REE) geochemical analysis from mafic microgranular/magmatic enclaves (MMEs) and host rocks based on petrographic descriptions such as alteration, mineral assemblages and rock classifications. First, the samples were cut away the weathered surfaces using a diamond saw and then was crushed the fresh parts with a jaw crusher. Next, the samples were ground in an agate ball mill to produce powders of each sample. Whole rock analyses were carried out by ACME analytical laboratories, Vancouver, Canada. The major elements and trace-rare earth elements were analysed by inductively coupled plasma atomic emission spectroscopy (ICP-AES) and inductively coupled plasma mass spectrometer (ICP-MS), respectively, using lithium metaborate/tetraborate fusion and nitric acid digestion methods. For data quality, the major intra-lab standard (Reference Material STD SO-18) was analysed together with the samples. Standards STD-SO-18 gave the first relative standard deviations of 6% or better for all trace and rare earth elements.

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Two granitoid host rocks and one MME analysed by the

40

Ar/39Ar method at the

University of Nevada Las Vegas were wrapped in aluminium foil and stacked in 6 mm inside-diameter sealed fused silica tubes. Individual packets averaged 3 mm thick and neutron fluence monitors (FC-2, Fish Canyon Tuff sanidine) were placed every 5-10 mm along the tube. Irradiated FC-2 sanidine standards together with CaF2 and K-glass fragments were placed in a copper sample tray in a high vacuum extraction line and were fused using a 20 W CO2 laser. The samples were viewed during laser fusion with a video camera system and they were positioned by a motorised sample stage. The samples were analysed with the furnace step-heating method, which used a double vacuum resistance furnace similar to Staudacher et al.’s (1978) design. For

40Ar/39Ar

analyses, the plateau segment consisted of

three or more contiguous gas fractions that have analytically indistinguishable ages (i.e. all of the plateau steps overlap in age at  2 analytical error) and which comprise a significant portion of the total gas released (typically >50%). The total gas (integrated) ages are calculated by weighting the amount of

39Ar

released, whereas the plateau ages are weighted

by the inverse of the variance. For each sample, inverse isochron diagrams are examined to check for the effects of excess argon. Reliable isochrons are based on the MSWD criteria that were developed by Wendt and Carl (1991) and, as for the plateaus, they must comprise contiguous steps and a significant fraction of the total gas released. All of the analytical data are reported at a confidence level of 1 (standard deviation). Nine representative samples from the MME and host rocks were selected for Nd-SrPb isotopic analysis, which was carried out at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia (UBC), Canada. Measurements of the Sr and Nd isotope compositions in the samples were conducted by thermal ionisation mass spectrometry (TIMS) on a Thermo Finnigan Triton system. The Sr and Nd isotopic compositions were corrected for mass fractionation using

86

Sr/88Sr=0.1194

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and

146

Nd/144Nd=0.7219. The precision of the results was controlled by the international Nd

standard La Jolla (an average value of 0·511858±0·000021 (2α, n=4)) and international Sr standard SRM 987 (an average value of 0·710248±0·000020 (2α, n=5)), which were analysed together with the samples. The US Geological Survey standard rock sample G2 was included with the samples and it gave Sr and Nd isotopic ratios of 0.709781±0.000008 and 0.512233±0.000006, respectively. These values are in agreement with the published values of 0·709770±14 and 0·512228±6, respectively (Weis et al., 2006). Analyses of the Pb isotope compositions were obtained by multiple collectors inductively coupled plasma mass spectrometry (MC-ICP-MS) on a Nu Plasma (Nu Instruments) system. The detailed analytical procedure for the Pb isotopic analyses on the Nu Plasma system at the PCIGR has been described by Weis et al. (2005). During the carrying out of the analysis, analysis of international standard NBS 981 was repeated, giving 206Pb/204Pb=16.9412±0.0006 (2α= n20), 207

Pb/204Pb=15.4983±0.0006 (2α= n20), and

208

Pb/204Pb=36.7183±0.0018 (2α= n20). The

measurement precision and the accuracy of the obtained results were estimated by using the results of systematic parallel analyses of US Geological Survey standard rock sample G2. The

206

Pb/204Pb,

207

Pb/204Pb and

208

Pb/204Pb isotopic ratios of the USGS standard samples

gave 18.4305±0.0008, 15.6387±0.0007, and 38.9094±0.0020, respectively. These values are in agreement with those for leached residues of USGS-G2 that were published by Weis et al. (2006).

4. Results 4.1. Field relationships and petrography The Taşkapı-Mecitli granitoid, which is located north of the town of Erciş (Van, East Turkey), covers an area of approximately 100 km2 at the earth surface (Fig.1c). It mainly consists of host rocks, including diorite-granodiorite-granite with allotriomorphic texture and

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gabbroic-dioritic MMEs with porphyritic texture (Appendix 1). This granitoid is unconformably overlain by the Quaternary Karakaya andesitic lava flow and by Miocene volcanic and volcanoclastic rocks (Fig. c in Appendix 1). The Taşkapı-Mecitli granitoid cuts the Palaeozoic metamorphic rocks, Eocene sandstones and Upper Cretaceous ophiolitic mélange units (Fig. d in Appendix 1). The contacts between the granitoid and these basement rocks are predominantly sharp and discordant. The granitoid body was intensely deformed and it displays many fractures, cracks, and joint planes (Fig. a in Appendix 1). The rocks have been exposed to alterations such as epidotisation and kaolinisation, especially along the fault zones and joint systems. Under the influence of this alteration, mineralisation in the form of pyrite can be observed on the coarse-grained granites and granodiorites (Fig. b in Appendix 1). Although most rocks of the Taşkapı-Mecitli granitoid are medium to coarsegrained, it has been observed to be more fine-grained towards its margins. The granites and granodiorites, which are green to dark grey, constitute the main body and the bulk of the Taşkapı-Mecitli granitoid. They are also coarse-grained and they are altered and weathered due to geothermal fluids. The diorite and monzonites, which are light green and grey in colour, occur at the margin of the Taşkapı- Mecitli granitoid. They are finer-grained and less altered than the granites and granodiorites. The MMEs usually have rounded ellipsoids with diameters ranging from 1 to 15 cm. The contact relations between the granitoid body and the MMEs are commonly sharp (Figs. e, f in Appendix 1). The MMEs, which are blackish and dark grey in colour, are usually found in the coarse-grained granite and granodiorite rather than in the fine-grained diorite and monzonite. The MMEs are fine-grained, like the diorite and monzonite of the host rock, and the observable crystals are a maximum of 0.2 cm in size. The MMEs are comprised of monzogabbro-diorite, monzonite, diorite and quartz-monzonite. Petrographically, the host rocks of the Taşkapı-Mecitli granitoid are generally coarse and fine-grained (Appendix 1). Although coarse-grained granite-granodiorites and fine-

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grained diorite-monzonites are similar in terms of their mineral contents, they display different textures. While the granite and granodiorites display a holocrystalline texture (Fig. g, h in Appendix 1), the diorite and monzonites have holocrystalline, porphyritic, and graphic textures (Figs. i, j in Appendix 1). The major mineral phases of the host rocks of the TaşkapıMecitli granitoid mainly consist of plagioclase (25% to 55%), K-feldspar (5% to 34%), quartz (10% to 33%), hornblende (2% to 15%) and biotite (0% to 8%). The accessory minerals are apatite, zircon and opaque minerals. The secondary minerals in the host rocks of the Taşkapı-Mecitli granitoid are comprised of chlorite and sericite. K-feldspar (orthoclase) and plagioclase form subhedral to anhedral crystals, and some of their crystals are altered into sericite and clay minerals. The plagioclase generally shows prismatic-shaped and polysynthetic-albite twining. The quartz is anhedral and it has irregular cracks. The hornblendes, which are green to brownish green, occur as anhedral to euhedral prismatic crystals and chloritisation is observed in some of their crystals. Biotite has anhedral to euhedral shapes, it is reddish brown in colour, and it forms prismatic crystals and lamellae. A few microstructural textures, such as acicular apatites and blade-shaped biotites, can be observed in the coarse-grained host rocks but they have not been seen in the fine-grained host rocks. The MMEs are finer-grained relative to the host rocks of the Taşkapı-Mecitli granitoid. In terms of mineral types, they are similar to the fine-grained dioritic and monzonitic host rocks. They mainly consist of plagioclase (46% to 66%), K-feldspar (10% to 17%), and rarely quartz (0% to 8%), biotite (6% to 12%) and hornblende (11% to 18%) with accessory zircon, apatite and opaque minerals. They display ophitic, holocrystalline and porphyritic textures, and alteration and secondary minerals such as epidote, chlorite and so on are not observed (Fig. k, l in Appendix 1). The plagioclase crystals, which are observed more frequently than orthoclase as feldspar species, are prismatic and lath-shaped. Quartz is

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anhedral and it is rarely present. Hornblende which is found more frequently than biotite as a mafic phase in the MMEs is greenish and it is forms prismatic crystals. The biotites are reddish brown to yellow, and they form lamellae. Textural features pointing to magma mixing were not observed.

4.2. Ar-Ar dating In this study, new geochronological age data have been obtained by the

40

Ar-39Ar

method. Samples were run as conventional furnace step-heating analyses on the hornblende and biotite separates and the results are given in Table 1 and Fig. 2. The New Ar-Ar dating results are presented for two the host rock granitoids and one is given for MME. The two biotite separates from the host rocks of the Taşkapı-Mecitli granitoid (Sample no M-23 and Z-9A) gave an Early Miocene age (Fig. 2 and Table 1). Sample M-23 produced an ideal, flat and concordant age spectrum, which is characteristic of a highly reliable sample. The total gas age is 22.77  0.14 Ma. Steps 1 to 9 (94% of the

39

Ar released) define an

indistinguishable plateau age of 22.75  0.16 Ma. No isochron is defined by these data. Sample Z-9A produced a generally flat and concordant age spectrum. The total gas age is 22.57  0.05 Ma (Table 1). No plateau or isochron ages are defined by these data. The Ca/K ratios are consistent with the analysis of a homogeneous biotite mineral separate. Meanwhile, the radiogenic yields are high and do not suggest recent alteration. Therefore, the total gas age is the most reasonable to use for this sample. The hornblende separate from the MMEs (sample no Z-9B) presented a similar age (Early Miocene) as the host rocks. Sample Z-9B yielded a good plateau of over 89% of the 39

Ar released between the 5th and 13th steps. No isochron is defined by these data. The Ca/K

ratios are consistent with the analysis of a homogeneous hornblende mineral separate, and the radiogenic yields are reasonably high (for this mineral) and thus do not suggest a recent

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alteration. The

40

Ar/39Ar dating of a hornblende separate from the sample Z-9B yielded a

plateau age of 22.88  0.21 Ma (Fig. 2 and Table 1). The obtained age data indicate that the cooling and emplacement ages for the host rocks and the MMEs are similar.

4.3. Geochemistry The results of the whole-rock major, trace and rare earth element analysis of the host rocks (28 samples) and the MMEs (11 samples) of the Taşkapı-Mecitli granitoid are given in supplementary tables as an MS Excel workbook. Nine samples (three MMEs and six host rocks) were analysed for the Sr, Nd and Pb isotopes (supplementary tables) and three samples (one MME and two host rocks) have been dated by the Ar-Ar geochronological method (Table 1 and Fig. 2). The major element data obtained from these analyses indicate that the loss-on-ignition (LOI) values of the studied host rocks and the MMEs are generally < 2.5 wt. % (see the supplementary tables), except for three of the granodiorites of the host rocks (LOI; 4-4.8 wt. %) and one MME sample (LOI; 4 wt. %) which could be affected by hydrothermal and, or meteoric alteration, as well as low-grade metamorphism.

4.3.1. Major element geochemistry The major element compositions of the host rocks cover a wide range from 58.31 to 72.11 wt. % for SiO2 and 0.39 to 0.53 for Mg# (Mg# = 100 Mg2+/(Mg2+ + Fe2+); Mg# values were calculated on an anhydrous basis adjusted to 100%, using the Fe2O3/FeO ratio after Middlemost, 1989). In the total alkali vs. silica classification diagrams, the majority of the samples plot in the diorite, granodiorite and granite (Fig. 3a). All of the host rocks are subalkaline in character and they are classified as high calc-alkaline series on the SiO2 versus K2O diagram (Fig. 3b). On the AFM diagram (total alkali-MgO-FeO), all of the sub-alkaline samples of the host rocks are classified as calc-alkaline (Fig. 3c). All of the granite-

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granodiorites and diorites of the host rock are metaluminous to weakly peraluminous, whose ASI (molar Al2O3/CaO+Na2O+K2O) values between 1.1 and 0.87 (Fig. 3d). The MMEs of the Taşkapı-Mecitli granitoid lie in a narrow range from 54.7 to 63.7 wt. % for SiO2 and 0.48 to 0.66 for Mg#. The MMEs are classified as monzodiorite, monzonite, gabbroic-diorite and diorite (Fig. 3a) based on the total alkaline versus silica diagram (TAS). In the SiO2 versus K2O diagrams, the MME samples are calc-alkaline to high calc-alkaline in character and they are classified as calc-alkaline on the AFM diagram (Fig. 3b, c). All of the MME samples are metaluminous with ASI index values ranging from 0.69 to 0.97 (Fig. 3d). The major element Harker diagrams of both the host rocks and the MMEs are shown in Fig. 4. The TiO2 (R2=0.71), Fe2O3TOT (R2=0.92), CaO (R2=0.804) and MgO (R2=0.825) contents of the host rocks and the MMEs decrease with increasing SiO2. In contrast, K2O (R2=0.852) displays a positive correlation with SiO2 and Na2O remains nearly constant with increasing SiO2.

4.3.2. Trace element geochemistry The selected trace element compositions of the host rocks and the MMEs are plotted against SiO2 contents in Fig. 5. The transition elements Sc (R2=0.83), Co (R2=0.84) and V (R2=0.68; not shown) decrease with increasing SiO2. The trace elements Sr (R2=0.65), Y (R2=0.71), Nb and Eu display flat and/or weakly negative trends up to 63.78 wt. % for SiO 2. They then suddenly decrease with increasing silica. Ba and Rb trace elements exhibit positive and, or flat trends against SiO2%. Chondrite–normalised REE patterns of the host rocks and the MMEs are presented in Figs. 6a and b. The REE pattern of the host rocks and the MMEs display strong light rare earth elements (i.e. LREE) enrichment relative to the medium and heavy rare earth elements (i.e. MREEs and HREE, respectively). The MMEs exhibit less enriched LREE (e.g. ranging

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from 90 to 180 times the chondritic values for La) compared to those of the host rocks (ranging from ~80 to ~230 times the chondritic values for La). In the REE patterns of the host rocks (Fig. 6a), there is a strong depletion in the MREEs relative to the LREEs, forming a concave upward pattern. The REE patterns of both the host rocks and the MMEs have a negative Eu anomaly (ranging between Eu/Eu* = 0.64 to 0.85 and 0.79 to 0.92, respectively). The N-type MORB-normalized multi-element patterns of the host rocks and the MMEs are shown in Figs. 6c and d. All of the samples from the host rocks and the MMEs display enrichments in large-ion lithophile elements (i.e. LILs; from Sr to Th) relative to HFSE (i.e. high field strength elements; Nb, Ta, Zr, Hf, P and Ti) and HREEs (Yb and Lu). The patterns exhibit a significant depletion in Nb and Ta relative to the adjacent LILs and LREEs (La and Ce), indicating the presence of a subduction signature and/or crustal contamination. Another noteworthy trend in the spider diagrams is a depletion in Ba and P relative to the adjacent elements. It should be noted that the high LILs/HFSEs and low LREEs/HREEs ratios of the host rocks and the MMEs are generally similar, suggesting that they are derived from the same source region.

4.3.3. Isotope geochemistry Sr, Nd and Pb isotopic data obtained from studied samples are presented in supplementary tables as an MS Excel workbook. The initial isotopic ratios for the Sr, Nd and Pb isotopes were calculated at an age of 23 Ma obtained by the dating method. The

87

Sr/86Sr(23Ma) and

143

40

Ar-39Ar geochronological

Nd/144Nd(23Ma) ratios of the host rocks vary in a

narrow range from 0.70640 to 0.70719 and from 0.512551 to 0.512559, respectively. Sr (23Ma) and Nd (23 Ma) values of the MMEs are in a narrow range from 0.70620 to 0.70640 and from 0.512551 to 0.512560, respectively. Lead isotopic ratios of the host rocks have a small range with 206Pb/204Pb(23Ma),

207

Pb/204Pb(23 Ma) and

208

Pb/204Pb(23Ma) varying over 18.939

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to 18.971, 15.656 to 15.690, and 39.001 to 39.078, respectively. The MMEs show limited changes, with ranges for

206

Pb/204Pb(23Ma),

207

Pb/204Pb(23

Ma)

and

208

Pb/204Pb(23Ma) isotopic

values between 18.863 to 18.895,15.678 to 15.682 and 39.021 to 39.046, respectively. Note that the Sr, Nd and Pb isotopic compositions of the host rocks are similar to those of the MMEs. The Sr-Nd isotope diagram is presented in Fig. 7. The Sr and Nd isotopic composition of global subducting sediment (GLOSS; Plank and Langmuir, 1998), Mariana Trough (Pearce et al., 2005), lithospheric mantle melting array (Davis and Von Blanckenburg, 1995) and previously reported isotopic data from volcanic and magmatic areas of the WesternEastern Anatolian and Arabian plates are plotted on the diagram for comparison (Fig. 7). The host rocks and the MMEs exhibit a nearly constant radiogenic Nd with increasing radiogenic Sr. All of the samples plot in the enriched quadrant relative to bulk earth, within the lithospheric mantle melting array and the GLOSS field in Fig. 7a. The lead isotopic data are illustrated in a plot of

206

Pb/204Pb -

207

Pb/204Pb, and

206

Pb/204Pb -

208

Pb/204Pb diagrams (Figs.

7b, c) with EM-I, EM-II mantle, MORB, GLOSS and data from the Arabian plate and magmatic areas of Eastern Anatolia plotted on the diagrams for comparison. All of the samples are situated above the depleted MORB mantle (DMM; Workman and Hart, 1995) and the Northern Hemisphere Reference Line (NHRL; Hart, 1984). On the other hand, all of the data points fall into and, or near the GLOSS field on the Pb-Pb isotope projections and they also plot in the EM-II type mantle field in the 207Pb/204Pb, and 206Pb/204Pb diagram (Fig. 7b). The lead isotopic compositions of the studied host rocks and the MMEs are quite similar to those of the magmatic of Eastern Anatolia.

5. Discussion 5.1. Origin of MMEs

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The question of whether or not the source of both the host rock and the MMEs could be derived from identical magma can be examined through the behaviour of the trace element compositions and the petrographical-mineralogical features, in addition to the isotopic ratios. The MMEs are generally observed in the granitoids, ranging from intermediate to felsic compositions within orogenic belts and continental arcs. The three favourite models suggested by previous studies for the origin of the MMEs are described in the following; (1) The restite model suggests that the MMEs are refractory or restitic portions of the granitoid source (Chappell and White, 1992; Chappel et al. 1987). According to this model, the MEEs would contain metamorphic and, or sedimentary samples of un-melted rock. This indicates that the refractory materials are derived from the source region. Taking into account the fact that these fabrics are not observed in the studied enclaves, it might consider that this model is not consistent with the studied MMEs. In addition, the Ar-Ar isotopic ages of the MMEs and the host rocks are similar which is incompatible with the restite model. (2) MMEs are formed by inclusion of mafic magmas into partially crystallised felsic magma, which is known as the mixing/mingling process (Didier and Barbarin, 1991; Barbarin, 2005; Chen et al., 2009). This model is characterised by resorption and microstructural textures, finer-grained enclaves compared to the host rocks and a strong linear correlation between the host rocks and the MMEs on the Harker diagrams. The studied MMEs are fine-grained and they also geochemically exhibit a linear correlation with the host rocks (Figs. 4 and 5). Even though the linear correlations and some resorption textures such as acicular apatite and blade-shaped biotite that point to magma mixing/mingling are observed in the samples of the Taşkapı-Mecitli granitoid, it should be noted that these features could have been caused by magma replenishment, decompression and fractional crystallisation. For instance, the MELTS thermodynamic modelling results that are discussed in Section 5.3 and presented in Fig. 11 indicate that these linear trends can be obtained by

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fractional crystallisation. The resorption textures can also be formed by magma replenishment occurring by injection of mafic magma into a more intermediate or felsic magma chamber (Özdemir and Güleç, 2014; Oyan et al., 2016; Chen et al., 2015; 2016) or by rapid magma decompression during magma ascent (Nelson and Montana, 1992; WeiFeng et al.2007). Furthermore, Chen et al. (2016) have reported that abundances of incompatible elements Zr and P in the MMEs cannot be explained by magma mixing or mingling. Concentrations of these elements in basaltic magmas derived from the mantle are very low (5 ppm for Zr and 0.019 wt. % for P2O5; Lee and Bachmann, 2014). The Zr and P2O5 values of the studied MMEs vary in range from 92 to 222 ppm and from 0.15 to 0.72 wt. %, respectively and these concentrations are much higher than in the mantle. These findings and data suggest that magma mixing processes cannot be responsible for the formation of the studied MMEs. (3) The autolith model, which is known as a cognate origin for MMEs, can be defined by chilled materials or cumulates of early-formed co-genetic crystals (Donaire et al., 2005; Niu et al., 2013; Chen et al., 2015; 2016). This hypothesis is characterised by similar mineral assemblages, indistinguishable isotopic compositions and similarities in geochemical compositions between the MMEs and the host rocks. However, some researchers (Barbarin, 2005; Chen et al. 2009) have interpreted that isotopic and mineral similarities between the MMEs and the host rocks can occur by magma mixing because of chemical and isotopic equilibration during mixing processes, which is based on some experimental studies stating that the isotopic equilibrium is reached faster than the chemical equilibrium. Nevertheless, Chen et al. (2016) have reported that if major and trace elements cannot reach chemical equilibrium, then it is unlikely that the isotopes will be homogenised because isotopes are carried with the related chemical elements and diffusion of the “carrying” elements can cause isotopic diffusion (Chen et al., 2015; 2016).

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The studied MMEs and the host rocks have similar mineral compositions and isotopic values. Although some representative textures such as acicular apatite and blade-shaped biotite were observed in the coarse-grained host rocks, crystal resorption or reactive overgrowth indicating magma mixing in the MMEs were not determined. The strongly linear or weakly parabolic correlated variations between the host rocks and the MMEs in the Harker diagrams (Figs. 4 and 5) are compatible with modal mineralogical compositions on the basis of magmatic evolutionary processes. This suggests that the MMEs are cumulate and the host rocks are residual melts. The mixing processes of the two different magmas are in the complex and multi-stage and, therefore, the formation of linear trends in the Harker diagrams can be difficult during this magmatic process (Donaire et al., 2005). The MELTS (Fig. 11) and EC-AFC (energy-constrained assimilation and fractional crystallisation; Figs. 12b, c) models obtained in this study also indicate that fractional crystallisation (FC) and crustal contamination evaluation processes can be responsible for the linear trends. In the light of these data, it can be concluded that the studied MMEs are earlier crystallised cumulates that were partly modified by subsequent magma replenishment. The fine-grained textures of the MMEs may have been caused by rapid quenching of the primary mafic magma. Chen et al. (2015) implied that quenching of mafic magma could account for the formation of fine-grained cumulates. In addition, some of the absorption textures observed in the coarse-grained granodiorite of the host rocks could have been caused by magma replenishment that was formed by injection of the primary melts into granitic magmas. Furthermore, hyperbolic trends suggesting magma mixing were not observed in Zr/Y-La, La/Y-La, Ti/Zr-Sr/Zr and Ti/Zr versus Rb/Zr variations diagrams (not shown).

5.2. Magma generation and partial melting

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High-K and metaluminous I-type granitoids are common in the post-collisional and active convergent margins (Pearce et al. 1990; Pearce and Peate, 1995). These types of granitoids are accepted as products that could be derived from either continental crust or subcontinental mantle via FC with or without crustal contamination and also magma mixing (Chappell and White, 1992; Barbarin 2005; Karslı et al., 2010). The MMEs from the TaşkapıMecitli granitoid display metaluminous, I-type characteristics, whereas the host rocks are represented by metaluminous to slightly peraluminous, I-type granitoids. The MMEs and the host rocks have Mg# values ranging from 0.48 to 0.66 and from 0.39 to 0.53, respectively, and the Ni contents vary between 3-15 and 34-64 ppm, respectively. These data indicate that the host rocks and the MMEs were not in equilibrium with primary melts and also that they must have been evolved by the crystallisation processes, with or without crustal contamination. However, the relatively high Mg# values (> 0.53) in the monzonite and diorite samples of the host rocks may indicate that they could have been slightly affected by the evolutionary processes and, therefore, they can be considered as the least evolved samples. The host rocks from the Taşkapı-Mecitli granitoid exhibit an enrichment of LILs and LREEs relative to HFSEs in the spider diagrams (Fig. 6). In addition, Nb, Ta and Ti incompatible elements relative to the adjacent LILs and REEs have negative anomalies, which indicate a crustal source. However, these trends can point to partial melting of a metasomatised mantle source that was enriched by a subduction component. Experimental studies have reported that partial melting of amphibolites or mafic materials in the lower crust could produce metaluminous granitic compositions (Rapp and Watson, 1995) and also that products derived from the melt of the mafic lower crust materials have lower MgO, Mg# and Ni contents. The MgO and Mg# contents of the host rocks vary within a wide range, from 1.76 to 4.07% and 0.43 to 0.64, respectively. While the least evolved samples of the host

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rocks have relatively high MgO (4.07%) and Mg# (0.64) contents, the evolved samples are characterised by lower MgO (2.67%) and Mg# (0.43) values than those of the least evolved samples. If the source of the Taşkapı-Mecitli granitoid were derived from direct partial melting of the mafic crust, then they would have low MgO% and Mg# and relatively high SiO2% contents similar to those in the experimental studies obtained by Rap and Watson (1995). However, this is not the case for the least evolved samples. In addition, slightly radiogenic Sr and Nd compositions (87Sr/86Sr(23Ma)=0.706404 and 143Nd/144Nd(23Ma)=0.512564 for the least evolved sample) of the host rocks are indicative of significant mantle input or juvenile mafic lower crust. Besides, the host rocks have low Al2O3/(Fe2O3+MgO+TiO2) and high Al2O3+Fe2O3+MgO+TiO2 contents, indicating that they could be produced by melting of amphibolitic lower crust materials (Fig. 8). To illustrate what has been discussed above, a Sr-Nd isotope diagram was produced (Fig. 9). The different primitive source compositions are plotted in this diagram because there are three different compositions for the origin of granitoid bodies in the world: pure crust, pure mantle and mixing between mantle and crust materials. Miocene Çökek basalts (Oyan 2011) were chosen to represent the mantle origin because they are the most primitive basaltic samples in the studied field. They also have high MgO (< 9 wt. %) and Mg# (0.68), and low SiO2 (44 wt. %) contents and they are unaffected by the crustal contamination and fractionation (Oyan 2011). It should be noted that Çökek basalts are derived from a mantle source that is enriched by a subduction component (Oyan 2011). Additionally, primitive mantle (Sun and McDonough 1989) and depleted MORB mantle (DMM; Workman and Hart, 2005) compositions were used as mantle origins for comparison. In addition, the mafic lower crust composition of Rudnick and Gao (2003) was taken as the pure mafic crust materials because chemical and isotopic compositions of the mafic lower crust in Eastern Anatolia are unavailable in the literature or in this study. The subducted sediments (SM) were obtained

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from GLOSS (Plank and Langmuir, 1998). A mixture of the mantle and the mafic lower crust was used as the mixing origin. As can be seen in Fig. 9, neither the mantle origin (primitive mantle, DMM and Çökek basalts) nor the mafic lower crustal material source compositions alone could produce rock compositions similar to those of the least evolved samples from the Taşkapı-Mecitli granitoid, because the data points of the least evolved samples fall into an area between these modelling curves or away from them. The position of the data point of the least evolved sample (sample M-29) in this diagram indicates that it could be formed by a mixture between the Çökek basalts (38%) and the mafic lower crust (62%), rather than any of the primitive mantle, depleted mantle or lower crust compositions. To reveal the source region of the Taşkapı-Mecitli granitoid, a partial melting model adopting non-modal batch melting following the approach of Oyan et al. (2016) was constructed. The MORB-normalised multi-element patterns of the least evolved sample of the host rocks were compared with the patterns of the trace element concentrations that obtained from the partial melting calculations. Based on the residual mineral phases obtained from MELTS modelling (Fig. 11), and from experimental studies of the amphibolites and metabasaltic rocks (Rapp and Watsonn, 1995) and petrogenetic simulations (Sajona et al., 2000), it was assumed that modal mineralogy (mineral source mode for the non-modal partial meting model) in the mixing origin (i.e. for processes in the Çökek basalts-mafic lower crust interaction) can be orthopyroxene + clinopyroxene + amphibole + garnet + rutile + apatite. Owing to the fact that the proportions of these mineral phases are unknown, a series of iterative calculations that gradually changing proportions of each mineral from 0% to 60% was performed until obtained the best fit curve to the least evolved sample. This approach has previously been applied to adakites south of Tibet (Guo et al., 2007) and to potassic magmatism of SE Tibet (Guo et al., 2005; 2006). The mineral melting mode that used in the non-modal batch melting model for the mafic crust-mantle interaction source was assumed

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orthopyroxene (3%), clinopyroxene (30 %), garnet (-1%), amphibole (51%), rutile (2%) and apatite (15%), based on experimental studies (Şen and Dunn, 1994; Rapp and Watson, 1995) and petrogenetic modelling (Guo et al. 2007). The partition coefficients for elements using the MORB-normalised melting model are given in supplementary tables in MS Excel workbook. In addition, the curves in the different melting degree of the mafic lower crust and the Çökek basalts are plotted on the diagrams for comparison (Fig. 10). The melting patterns obtained from both the mafic lower crust and the Çökek basalts are incompatible with those of the least evolved sample (Fig. 10), whereas the melting pattern of the mixing origin (38% Çökek basalts and 62% mafic lower crust) indicates that melt with trace element concentrations similar to those of the least evolved samples (M-29) could be generated by modal mineralogy involving orthopyroxene (5%), clinopyroxene (30%), garnet (7%), amphibole (50.79%), rutile (2%) and apatite (5.21%) and non-modal batch melting of the mantle-mafic crust interaction with partial melting degree of about 18% (Fig. 10). These findings reveal that mafic melts derived from partial melting of the lower crust-mantle interaction might be responsible for the formation of the least evolved sample of the host rocks that formed the Taşkapı-Mecitli granitoid to the north of Lake Van.

5.3. Fractional crystallisation The major and trace elements of the studied samples in the Taşkapı-Mecitli granitoid exhibit positive and negative trends with SiO2 content, revealing the importance of fractional crystallisation in evaluating the degree to which the magma chamber(s) depend on the crystallisation of mineral phases. The Al2O3 contents of the host rocks and the MMEs decrease with increasing silica and this trend can be attributed to the fractionation of feldspar. Decreasing Sr and Eu, and nearly constant or increasing Ba with decreasing silica suggest fractionation of plagioclase rather than K-feldspar at shallow crustal levels because Sr and Eu

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have higher partition coefficients for plagioclase than K-feldspar, whereas K-feldspar shows higher Kd values for Ba than Sr and Eu elements (Furham and Graham, 1999). Negative Eu anomalies (average Eu/Eu* = 0.7; Eu/Eu*=EuNM/(SmNMxGdNM)1/2; where NM stands for values normalised to N-type MORB) can be observed in the incompatible trace element spider diagrams (Figs. 6c, d) and are consistent with the variations for plagioclase fractionation. A minor fractionation of K-feldspar will also result in the depletion of Ba relative to the adjacent elements in the incompatible trace element pattern of the samples (Figs. 6c and d). The presence of a notable decrease in CaO, MgO and the transition elements Co, Sc and Ni (not shown) with increasing SiO2 contents are consistent with calcicplagioclase and pyroxene fractionations in the low SiO2 values, and sodic-plagioclase and amphibole fractionations in the high SiO2 contents. The negative correlations in Ni and TiO2 against silica in the samples may reflect the fractionation of ferromagnesian minerals (e.g. pyroxene and amphibole) and oxides. The Nb variations against silica, which present constant and negative correlations for the samples of the Taşkapı-Mecitli granitoid, suggest that biotite fractionation may be important for evolutionary processes in the magma chamber. The decreasing Y with increasing SiO2 may be related to amphibole fractionation. This interpretation is consistent with the presence of a noteworthy depletion in the MREEs relative to the HREEs on chondrite-normalised multi-element patterns (Figs. 6a and b) because the MREEs (i.e. Eu to Ho) are more strongly partitioned into amphibole relative to the HREEs and LREEs. In addition, the decrease in P2O5 (not shown) and TiO2 with increasing silica can be linked to the FC of apatite and titanite (may be magnetite), respectively. These are supported by the depletion of P and Ti elements in the MORB-normalised incompatible trace element patterns of the studied samples. To examine the FC during the evolution of the Taşkapı-Mecitli granitoid at the crustal levels, MELTS models using the major oxide contents were produced. The MELTS code of

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Ghiorso and Sack (1995) allows us to take thermal parameters such as temperature, pressure, oxygen fugacity and buffers of primitive and resultant magmas into consideration. It also provides important clues about the physicochemical conditions of the magma chambers. To this end, the composition of the residual liquid and the cumulate minerals was produced (Fig. 10) based on the major oxides of the Taşkapı-Mecitli granitoid using the MELTS code. The least evolved sample (M-29) of the host rocks was taken as the starting composition for the modelling. Furthermore, different experimental conditions have been tested in the MELTS algorithm [P= 2-4 kbar; H2O=0 to 1.5% and ƒO2=QFM (quartz-fayalite-magnetite system)]. The modelled curves obtained by the MELTS algorithm over temperatures ranging from 12500C to 7500C are given in Fig. 11. The results of the MELTS algorithm indicate that the best match between modelled and observed differentiated trends for the Taşkapı-Mecitli granitoid could have been obtained via FC with 4 kbar pressure, 1.5% H2O and QFM system. The mineral fractionation phases obtained by the MELTS code under these physicochemical conditions are similar to the observed mineralogy in the rocks of the Taşkapı-Mecitli granitoid. The modelled fractionation processes (P=4 kbar, 1.5% H2O, ƒO2=QFM) obtained by the MELTS code from the least evolved sample to the most evolved samples of the host rocks indicate 80% FC and mineral assemblages involving orthopyroxene (~9 wt. %), clinopyroxene (~18 wt. %), plagioclase (~44 wt. %), K-feldspar (~6 wt. %), biotite (~2 wt. %) and apatite (~1 wt. %) would be required (Fig. 9e). In addition, according to the MELTS modelling, orthopyroxene crystallises at high temperature (~12100C) until ~30 wt. % of the liquid mass has fractionated. Clinopyroxene, plagioclase, apatite, K-feldspar and biotite begin to crystallise at 1150, 1090, 990, 840 and 8000C, respectively (Fig. 11f). These mineral assemblages are consistent with existing minerals as determined by petrographic studies, except for pyroxene. We have to note that the pyroxene mineral may be a residual phase or it may be consumed due to the fractionation. It is also crystallised in the fine-grained

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holocrystalline matrix of the fine-grained granite and, or MMEs. Although pyroxene minerals were not observed in either the MMEs or the host rocks, CIPW norms of the rock samples show that orthopyroxene (hypersthene; 0.2 to 5%) and clinopyroxene (diopside; 0.2 to 10%) could crystallise in the magmatic system that formed the Taşkapı-Mecitli granitoid (see the supplementary tables).

5.4. Contamination in the crustal levels Although the major and trace element contents and petrologic models have revealed the importance of FC in the evolution of the Taşkapı-Mecitli granitoid, slightly radiogenic Sr and Nd isotope compositions and slightly high Th (9 to 12 ppm for the MMEs and 12 to 19 ppm for the host rocks) and Pb (1 to 4.5 ppm for the MMEs and 0.8 to 6.6 ppm for the host rocks) values of the samples point to crustal contamination in the evolution of the magma chamber under the Taşkapı-Mecitli granitoid. However,

87

Sr/86Sr(23Ma) and Nd(23Ma) against

SiO2 diagrams (Figs. 12a, b) produce weakly positive or nearly constant trends, which indicate that the crustal contamination is negligible. To reveal the effect of the crustal contamination, EC-AFC models were constructed on the basis of the isotopic compositions and trace element values of the samples, as well as thermal parameters. The EC-AFC code of Bohrson and Spera (2001) and Spera and Bohrson (2001) allow us to use a number of thermal variables (e.g. partial melting of the crust and heat transfer from the magma to the crust) in the formulations. The isotope and trace element compositions of sample M-29 that were used in the MELTS model were taken as starting values for the least evolved magma. The local Palaeozoic schist that was cut by the granitic body and the bulk crust composition of Rudnick and Gao (2003) were taken as the contaminant materials. It can be seen in Fig. 12 that

87

Sr/86Sr vs.

143

Nd/144Nd (Fig. 12c) and Sr vs.87Sr/86Sr

(Fig. 12d) diagrams reveal EC-AFC processes in the Taşkapı-Mecitli granitoid, utilising the

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formulae of Spera and Bohrson (2011). Both the local schists and the bulk crust (Rudnick and Gao, 2003) end members were tested as crustal materials in the EC-AFC models for comparison. The initial temperatures of the crust and magma were selected within the ranges 300 to 600ºC and 1100 to 1400ºC, respectively, because the starting temperatures of the crust and the magma are unknown. EC-AFC simulations with the selected thermal and chemical parameters have been run and the modelling was terminated when the most suitable EC-AFC curve for the samples was obtained. The best fit thermal and chemical-isotopic parameters generated by the EC-AFC modelling are given in Table 2. The modelling results indicate that the Taşkapı-Mecitli granitoid may have been affected by contamination from the local schists rather than the bulk crust composition of Rudnick and Gao (2003), because the data point of the samples from the Taşkapı-Mecitli granitoid align along the modelled curves which were produced by using the values of the local schists. Data on the EC-AFC modelled curves revealed that the Taşkapı-Mecitli granitoid could have been slightly affected by assimilation in the upper crustal level (local metamorphic schist). In addition, the crustal involvement (Ma*) and FC (Mc) during magma genesis are a maximum of 2% and 78%, respectively. These findings imply that FC is a more effective evolutionary process than crustal contamination (which may be negligible) in the magma chamber beneath the Taşkapı-Mecitli granitoid.

5.5. Geodynamic implications The Turkish-Iranian high plateau is a result of the collision between the Eurasian and Arabian plates following the closure of the Neo-Tethys Ocean (Şengör and Kidd, 1979). The timing of the collision is still controversial and most popular ages for the initiation of collision range from 35 Ma to 20 Ma (Şengör and Kidd, 1979; Robertson et al., 2007; Allen and Armstrong, 2008; McQuarrie and Van Hinsbergen, 2013; Okay et al., 2010; Karaoğlan et

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al., 2016). Although these ages indicate a wide range, it can be said that the collision developed in the Late Oligocene to Early Miocene. The Ar-Ar dating results obtained from the MMEs and the host rocks indicate that they formed in the Early Miocene (~23 Ma) and they are post- or syn-collisional granitoids. All of the samples from the Taşkapı-Mecitli granitoid are plotted in the volcanic arc and the post- or syn-collisional granitoid fields (Fig. 13). Karaoğlan et al. (2016) have emphasised that the final continental collision between the Taurids and the Arabian plate should have occurred in the Oligocene, based on the apatite fission track age results that they obtained from the granites of the southeastern Anatolia. In addition, Lebedev et al. (2013) have reported that some of the subvolcanic rocks (~23 Ma) in Armenia and the Caucasus, and in the Taşlıçay granitoid (23 Ma) in Eastern Anatolia are post- or syn-collisional granitoids. These findings indicate that the collision is unlikely to have occurred after 23 Ma and, hence, the starting of the earliest collision should be in the Late Oligocene or beginning of the Early Miocene. The arc-related Cenozoic magmatism that is observed in wide areas on the TurkishIranian plateau began in the Eocene to Early Oligocene. This magmatism that occurred prior to the collision in the plateau is exposed in the southwest and northern Iran (Verdel et al., 2011), Armenia (Lordkipanidze et al., 1988) and also in the Erzurum-Kars Plateau in Northern Anatolia (Keskin et al., 1998). Richards et al. (2012) and Verdel et al. (2011) have reported that there may be a magmatic gap between the Oligocene and Early Miocene in the Erzurum-Kars Plateau due to limited magmatic activity from 30 Ma to 20 Ma, which may be due to thickening of the plateau and the continued continental collision. This age interval, which is defined as a magmatic gap in the plateau, could be characterised by limited plutonic and subvolcanic activities rather than an extensively eruptive style of magmatism. Lebedev et al. (2013) have determined that some of the subvolcanic rocks in Armenia and the Caucasus have ages of 23 to 24 Ma. These researchers have also reported that K-Ar dating of a

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hornblende from the Taşlıçay granitoid located in 10 km north of the Taşkapı-Mecitli granitoid yielded an age of 23 Ma. A geodynamic model proposed for the formation of the Taşkapı-Mecitli granitoid is illustrated in Fig. 14. Based on the whole rock major and trace element contents, isotopic ratios (Nd-Sr-Pb; Fig. 7) in addition to the petrologic modelling of the MMEs and host rocks, it can be proposed that interaction of the mafic lower crust and mantle was responsible for the Taşkapı-Mecitli granitoid. The flare-up of arc magmatism within the Turkish-Iranian Plateau in Eocene to Early Oligocene times (Neill et al. 2013) can be associated with the back-arc extension, flat subduction and/or slab rollback (Berberian and Berberian, 1981; Verdel et al. 2011). If the back-arc extension or slab rollback had caused the flare-up of arc magmatism, then we would expect to see intense magmatic activity associated with the slab instead of a magmatic gap in the Early Oligocene to the beginning of the Early Miocene. However, this is not the case in the plateau. The flat subduction proposed by Berberian and Berberian (1981) can be the most suitable model for the flare-up of arc magmatism because while arc magmatism in the Eastern Pontides, Armenia and the Caucasus was extensively observed in the Late Eocene and Early Oligocene, it is rarely seen in the EACZ where subduction of the southern Neo-Tethys slab was flat (Fig. 14). During the Late Oligocene to Early Miocene, magmatism in Eastern Anatolia probably developed in relation to the slab rollback and crustal thickness-shortening caused by a collision between the Arabian and Eurasian plates. Consequently, the slab rollback of the southern Neo-Tethys must have been caused by asthenospheric upwelling under the EACZ and, hence, re-mobilisation and partial melting in the base of the metasomatised lithospheric mantle (similar to Çökek basalts) could be triggered (Pearce et al. 1990). This situation may have led to the development of the Late Oligocene arc volcanism in the Erzurum-Kars Plateau (Keskin et al. 1998) and Early Miocene granitic intrusions (Lebedev et al. 2013 and this study) north of Lake Van. As

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discussed in Section 5.2, the partial melting of the mantle is insufficient for the formation of the Taşkapı-Mecitli granitoid and additional crustal material is required. Upwelling of asthenospheric melts via slab rollback and dehydration of the subducted slab that enriched the lithospheric mantle would lead to partial melting in the base of the sub-continental lithospheric mantle. The melts produced by this mechanism can settle in the base of the mafic lower crust and they can cause partial melting of the mafic lower crust. In this way, mafic melts derived from both the mafic lower crust and metasomatised lithospheric mantle can mix (homogeneous mixing) in the source region. The mixed origin melt that generates to the Taşkapı-Mecitli granitoid could then settle in the upper crustal levels. This geodynamic scenario is consistent with the formation of the source melts for the MMEs and host rocks, which would be evolved by FC and crustal contamination in a shallow magma chamber.

6. Conclusions The Taşkapı-Mecitli granitoid displays I-type, metaluminous and high-K calk alkaline characteristics. The emplacement of the pluton took place at ~23 Ma (Akitaniyen-Early Miocene). While the MMEs are classified as monzodiorite, monzonite, gabbroic-diorite and diorite, the host rocks have granite, granodiorite and diorite compositions. Ar-Ar dating results of the MMEs and host rocks from the Taşkapı-Mecitli granitoid yield similar ages. They also have the same mineralogy and indistinguishable isotopic compositions as well as similarities in geochemical compositions, all of which imply that the MMEs of the Taşkapı-Mecitli granitoid are earlier crystallised cumulate of the same magmatic system rather than resulting from magma mixing processes. The results of melting model based on the MORB-normalised incompatible trace element compositions suggest that a mixture of the Çökek basalts (38 %) and mafic lower

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crust (62%) with a partial melting degree of 18% might be responsible for the formation of the Taşkapı-Mecitli granitoid. Based on the tectonomagmatic history of Eastern Anatolia and together with age data obtained from the current study and from the previous literature, it can be stated that collision initiation between the Arabian and Eurasian plates should have taken place before/around ~23 Ma (i.e. Late Oligocene to Early Miocene). On the basis of these findings, it was emphasized that upwelling of asthenospheric melts via rollback of the southern Neo-Tethys Slab in the Late Oligocene to Early Miocene caused melting in the base of the subcontinental lithospheric mantle. The mafic melts derived from both the metasomatised lithospheric mantle and lower crust mixed at different rates in the source region. Finally, mafic melts derived from crust-mantle interaction have risen as diapirs. These melts that are located in the upper crustal level could generate the Taşkapı-Mecitli granitoid via FC and slight (negligible) crustal contamination.

Acknowledgments This work has been funded by TUBİTAK, the Scientific and Technical Research Council of Turkey (Project No. ÇAYDAG, 112Y200). I thank Alper Şengül for his help with the fieldwork. I also thank two anonymous reviewers for constructive comments and suggestions and Editor Derek Wyman for handling this manuscript.

References Alıcı, P., Temel, A., Gourgaud, A., 2002. Pb-Nd-Sr isotope and trace element geochemistry of Quaternary extension-related alkaline volcanism: a case study of Kula region (western Anatolia, Turkey). Journal of Volcanology and Geothermal Research 115, 487–510.

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potential: examples from the Tethyan arcs of central and eastern Iran and western Pakistan. Economic Geology 107, 295–332. Robertson, A. H. F., Parlak, O., R|ızaoğlu, T., Ünlgenc, U., Inan, N., Taslı, K., Ustaömer, T., 2007. Tectonic evolution of the South Tethyan ocean: Evidence from the Eastern Taurus Mountains (Elazığ region, SE Turkey). In: Ries, A. C., Butler, R. W. H. & Graham, R. H. (eds) Deformation of Continental Crust. Geological Society, London, Special Publications 272, 231-270. Sajona, F.G., Maury, R.C., Pubellier, M., Leterrier, J., Bellon, H., Cotten, J., 2000. Magmatic source enrichment by slab-derived melts in a young post-collision setting, central Mindanao (Philippines). Lithos 54, 173–206. Şen, C., Dunn, T., 1994. Dehydration melting of a basaltic composition amphibolite at 1.5 and 2.0Gpa: implications for the origin of adakites. Contribution to Mineralogy and Petrology. 117, 394–409. Spera, F.J., Bohrson, W.A., 2001. Energy-constrained open-system magmatic processes I: General model and energy-constrained assimilation and fractional crystallization (ECAFC) formulation. Journal of Petrology 42, 999-1018. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A.D. & Norry, M.J. (ed.) In Magmatism in Ocean Basins, Geological Society of London Special Publications 42, 313-345. Staudacher, Th., Jessberger, E. K., Dorflinger, D., Kiko, J., 1978. A refined ultra-high vacuum furnace for rare gas analysis. Journal of Physics E: Scientific Instruments 11,781-784. Şengör, A.M.C., Kidd, W.S.F., 1979. Post-collisional tectonics of the Turkısh – Iranian plateau and a comparison with Tibet. Tectonophysics 55, 361-376.

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Figure Captions Figure 1. (a, b) location map of the study area and Eastern Anatolia, Turkey, including major structural elements. (c) Geological map of the Taşkapı-Mecitli granitoid and surrounding areas. Figure 2.

39

Ar/40Ar plateau ages of the MMEs and the host rocks from the Taşkapı-Mecitli

granitoid. Figure 3. Classification of the samples from the Taşkapı-Mecitli granitoid. (a) Total alkali vs. silica diagram of Le Bas et al. (1986). Subalkaline-alkaline divisions of TAS diagram are from Irvine and Baragar (IB; 1971) (b) K2O (wt. %) vs. SiO2 (wt. %) plots of Peccerillo and Taylor (1976). (c) Total alkali-FeO-MgO (AFM) triangle diagram of Irvine and Baragar (1971) (d) Al2O3/ (Na2O+K2O) molar vs. Al2O3/ (Na2O+K2O+CaO) molar diagram for the MMEs and host rocks. For a thorough discussion, see the text. Figure 4. Harker diagrams of the representative subset of the major element oxides.

42

Figure 5. Harker diagrams of the representative subset of the trace element oxides. Figure 6. Chondrite-normalized REE and primitive mantle-normalized multi-element spidergrams of the host rocks and MMEs. The chondrite and primitive mantle (PM) values are of Sun and Mc Donough (1989). For an explanation, see the text. Figure 7. (a) 87Sr/86Sr(23Ma) vs. and (c)

206

Pb/204Pb(23Ma) vs.

143

208

Nd/144Nd(23Ma) and (b)

206

Pb/204Pb(23Ma) vs.

207

Pb/204Pb(23Ma)

Pb/204Pb(23Ma) ratio diagrams for the samples. Muş-Nemrut

(i.e. Muş-Nem-Tend), Bingöl, Süphan, Tendürek, Etrüsk and Kars-Ararat fields are from Pearce et al. (1990), Keskin et al. (2006), Özdemir et al. (2006), Özdemir and Güleç (2014), Oyan et al. (2016); Lebedev et al. (2016). GLOSS, Mariana through and lithospheric mantle melting array are from Plank and Langmuir (1998), Pearce et. al. (2005) and Davis and VonBlanckenburg (1996), respectively. Fields for Karacadağ has been produced on the basis of the data from Keskin et al. (2012). Field of the Arabian plate, Israel, Jordan and Afar plume are from Volker et al. (1993) and Krienitz et al. (2006). Kula volcanic field is from Alıcı et al. (2002). For an explanation, see the text. Figure 8. (a) Al2O3+Fe2O3+MgO+TiO2 vs. Al2O3/(Fe2O3+MgO+TiO2) variation diagram for the samples from the Taşkapı-Mecitli granitoid. Compositional fields derived from experimental melts of amphibolites, greywackes, and pelites are from Patiňo Douce (1999). For an explanation, see the text. Figure 9.

87

Sr/86Sr(23Ma) vs.

143

Nd/144Nd(23Ma) was produced to approach source compositions

that can be formed to the Taşkapı-Mecitli granitoid. Çökek volcanics (ÇV), primitive mantle (PM), depleted MORB mantle (DM), lower crust (LC) and sediment (SM; it is compositions of GLOSS obtained from Plank and Langmuir, 1998) were used to source compositions. Data point of the least evolved samples in the host rocks were plotted on the LC-ÇV mixing curve. For an explanation, see the text.

43

Figure 10. Model N-MORB-normalized incompatible trace element patterns of the leastevolved sample (M-29) and of the values obtained from petrologic models. The Pattern of the least-evolved sample is closely overlapped by mixture 38 % mantle (Çökek basalts) melt and 62 % lower crust melt, with the partial melting degree of 18%. Partition coefficients, mineral melting and source modes used in this model curves are presented in supplementary tables. For an explanation, see the text. Figure 11. (a, b, c, d) variations TiO2, Al2O3, CaO, Fe2O3 versus SiO2 for the TaşkapıMecitli granitoid compared with MELTS (Ghiorso & Sack,1995) fractional crystallization modelling curves. Calculation parameters: Crystallization temperature ranging in 1250-750 0

C, moderate pressure (2-4 kbar) and H2O contents between 0.5% and 1.5%. A QFM (quartz-

fayalite-magnetite system) buffer was selected ƒO 2. The least evolved sample (M-29) of the host rocks was taken as starting composition. (e) Fractionating mineral assemblage and liquidus temperature calculated with MELTS for a residual liquid percentage from starting compositions M-29. (f) A proportion of the crystallized mineral phases obtained from MELTS calculations with 4 kbar pressure, 1.5% H2O under QFM conditions. Plagioclase (Plg), ortho (Opx)- and clino (Cpx) pyroxenes, biotite (Bio), K-feldspar (K-feld), Apatite (Ap). Figure 12. Variation of (a) 87Sr/86Sr(23Ma) vs SiO2 (wt. %) and (b) Nd(23Ma) vs. SiO2 (wt. %). EC-AFC models utilizing (c)

87

Sr/86Sr(23Ma) vs.

143

Nd/144Nd(23Ma) and (d) Sr vs

87

Sr/86Sr(23Ma)

plots. The compositional and thermal parameters used in these EC-AFC models are presented in Table 2. Distribution coefficients used in these models are from the Geochemical Earth Reference Model home page (http:/www.earthref.org). Black numbers, italic black number, and grey number denote the percentages of assimilated contaminant, mixing (mantle-mafic lower crust interaction) and fractionation, respectively. For an explanation, see the text.

44

Figure 13. Tectonic discrimination diagrams for the samples. (a) Nb vs. Y diagram (Pearce et. al. 1994), (b) R1 versus R2 diagram after Bowden (1985). Abbreviations: VAG-volcanicarc granites; Syn-COLG-syn collisional granites; WPG-within plate granites; ORG-ocean ridge granites; Post-COLG-post collisional granites. Figure 14. Geodynamic model for the formation of the Taşkapı-Mecitli granitoid and surrounding granitic bodies. EKP-Erzurum-Kars Plateau, MG-Mecitli granitoid, TG-Taşlıçay granitoid. For an explanation, see the text.

45

FIGURE 1

46

FIGURE 2

47

FIGURE 3

48

FIGURE 4

49

FIGURE 5

50

FIGURE 6

51

FIGURE 7

FIGURE 8

52

FIGURE 9

53

FIGURE 10

54

FIGURE 11

55

FIGURE 12

56

FIGURE 13

57

FIGURE 14

58

Table 1. Summary of

40

Ar/39Ar data of samples from the Taşkapı-Mecitli granitoid. All samples were run as

conventional furnace step heating analyses. All data are reported at the 1 uncertainty level. Unit Host rock Host rock MME

Sample No M-23 Z-9A Z-9B

Rock Type Granite Diorite Monzonite

Sample Biotite Biotite Amphibole

Total gas age (Ma) 22.77 22.57 23.04

Plateau age (Ma) 22.75 22.88

59

Table 2. Compositional and thermal parameters used for EC-AFC modeling. Sr and Nd isotopic values of the crust 2 are from Rudnick and Gao (2003). Thermal parameters Crust 1

Crust 2

0

Magma liquidus temperature Tl, m

1280 C

1320 0C

Magma initial temperature Tm o

1280 0C

1320 0C

0

Assimilant liquidus temperature Tl, a

1000 C

1100 0C

Assimilant initial temperature Ta o

400 0C

600 0C

0

Solidus temperature Ts

950 C

950 0C

Equilibration temperature Teq

1000 0C

1000 0C

Crystallization enthalpy, hcry (J/Kg)

396000

396000

Isobaric specific heat of magma Cp, m (J/Kg per K)

1484

1484

Fusion enthalpy hfus (J/Kg)

354000

354000

Isobaric specific heat of assimilant Cp, a (J/Kg per K)

1388

1388

Compositional Parameters Crust 1

Crust 2

Sr

Nd

Sr

Nd

Magma initial concentration (ppm), Cm o

381.6

25.8

381.6

25.8

Magma isotope ratio, £m

0.706404

0.512551

0.706404

0.512551

Magma distribution coefficient, Dm

1.5

0.25

0.5

0.25

Assimilant initial concentrations (ppm) Cm o

1344

28

230

12.7

Assimilant isotope ratio, £a

0.709757

0.512161

0.711

0.5122

Assimilant distribution coefficient, Da

0.5

0.25

0.05

0.25

60

Graphical abstract

61

HIGHLIGHTS 

New whole rock geochemical, Sr-Nd-Pb isotopic and reported.



Mafic micro-granular enclaves are earlier crystallized cumulate of same magma compositions.



Parental melts have been derived from mantle-mafic lower crust interaction.



Collision between Arabian and Eurasian plates must be before/around 23 Ma.

40

Ar/39Ar dating data are