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International Geology Review Vol. 52, Nos. 4 – 6, April– June 2010, 536–578

Geochemistry and tectonics of Cenozoic volcanism in the Lesser Caucasus (Azerbaijan) and the peri-Arabian region: collision-induced mantle dynamics and its magmatic fingerprint Yildirim Dileka*, Nazim Imamverdiyevb and S¸afak Altunkaynakc a

Department of Geology, Miami University, Oxford, OH 45056, USA; bDepartment of Geology, Baku State University, Baku AZ1148, Azerbaijan; cDepartment of Geological Engineering, Istanbul Technical University, Maslak Istanbul 34469, Turkey

Downloaded By: [Dilek, Yildirim] At: 17:09 15 February 2010

(Accepted 22 September 2009) The Lesser Caucasus occurs in the hinterland of the Arabia – Eurasia collision zone in the broad Alpine– Himalayan orogenic belt and includes Cenozoic plutonic and volcanic sequences that provide important clues for collision-driven continental magmatism and mantle dynamics. Two main magmatic episodes (Eocene and late Miocene –Quaternary) formed the volcanic landscape and the igneous assemblages in the Lesser Caucasus of Azerbaijan. (1) The Eocene sequence consists of trachybasalt and basaltic trachyandesite with subordinate tephrite-basanite, basaltic andesite, and trachyandesite, showing shoshonitic and mildly alkaline compositions. The Miocene – Quaternary magmatic episode is represented by (2a) an early phase of upper Miocene – lower Pliocene andesite, trachyandesite, trachydacite, dacite and rhyolite lavas, and by (2b) a late phase of upper Pliocene– Quaternary trachybasalt, basaltic trachyandesite, basaltic andesite, trachyandesite, trachyte, and rhyolite flows. The rocks of the early phase have high-K calc-alkaline compositions, whereas those of the late phase show high-K shoshonitic compositions, defining an alkaline trend and a K2O-enriched melt source. All three volcanic associations show variant troughs in Nb, Ta, Hf, and Zr, strong enrichment in Rb, Ba, Th, La, and depletion in Ti, Yb, Y relative to mid-ocean ridge basalt N-(MORB) in their multi-element patterns. The enrichment of incompatible elements and K suggests derivation from a metasomatized mantle source, whereas the troughs in Nb and Ta indicate a subduction influence in the mantle melt sources. Mantle-derived magmas were modified by AFC/FC processes for all three volcanic sequences. These geochemical features are similar to those of coeval volcanic associations in the peri-Arabian region, and indicate the existence of subduction-metasomatized lithospheric mantle beneath the Lesser Caucasus during the Cenozoic. Partial melting of this subduction-modified subcontinental lithospheric mantle in the peri-Arabian region was triggered initially by slab breakoff following discrete continental collision events in the early Eocene. The heat source for the later Miocene –Quaternary volcanism in the entire peri-Arabian region was provided by asthenospheric upwelling, which itself was caused by delamination of the mantle lithosphere following the final Arabia – Eurasia collision at ,13 Ma. Increased alkalinity of successively younger units in the Plio-Quaternary volcanic associations resulted from the input of enriched asthenospheric melt during the last stages of postcollisional magmatism. Active, crustal-scale and orogen-parallel, transtensional fault systems in the peri-Arabian region facilitated the formation of fissure eruptions and stratovolcanoes in the latest Cenozoic.

*Corresponding author. Email: [email protected] ISSN 0020-6814 print/ISSN 1938-2839 online q 2010 Taylor & Francis DOI: 10.1080/00206810903360422 http://www.informaworld.com

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Keywords: Lesser Caucasus (Azerbaijan); peri-Arabian region; Turkish– Iranian high plateau; post-collisional magmatism; slab breakoff; lithospheric delamination

Introduction Cenozoic magmatic rocks occur extensively in the peri-Arabian region north of the Bitlis – Zagros suture zone (Figure 1), and they constitute a significant component of the continental crust in this segment of the Alpine –Himalayan orogenic belt. Although they range in age from Eocene to Quaternary, their temporal distribution reflects significant pulses of magmatism in the late Eocene, late Miocene – Pliocene, and Plio-Quaternary. The timing of their formation mostly coincides with and postdates a series of continental collision events in the region (Dilek and Whitney 2000, and references therein). Together with the nearly coeval volcanic-plutonic units in central and western Anatolia and in the Aegean region to the west (Yılmaz 1989; Altunkaynak and Yılmaz 1998; Dilek et al. 1999b; Aldanmaz et al. 2000; Pe-Piper and Piper 2001, 2006; Yilmaz et al. 2001; Agostini et al. 2007; Altunkaynak 2007; Dilek and Altunkaynak 2007, 2009; Kadioglu and Dilek in press), the Cenozoic peri-Arabian magmatic belt is part of a much larger igneous province, which developed in a broad zone of convergence between Afro-Arabia and Eurasia (Figure 1; Jackson and McKenzie 1984; Dewey et al. 1986; McClusky et al. 2000, 2003; Allen et al. 2004; Dilek and Sandvol 2009). The melt sources of the Cenozoic periArabian magmatism and the causes of heat supply that triggered melting are particularly important questions for the geodynamic conditions and mechanisms that result in highmagmatic productivity in post-collisional orogenic belts. The Eocene magmatic units in the peri-Arabian region are exposed in mainly narrow, E –W-trending, curvilinear belts that straddle the suture zones between the continental blocks (Figure 2). These magmatic units include granitoid – syenitoid plutons and coeval volcanic sequences intercalated with clastic – volcaniclastic rocks. Volcanic units have mildly alkaline, shoshonitic affinities and are overlain by late Eocene flysch deposits and/or late Miocene volcanic sequences. The next magmatic pulse in the region is represented by upper Miocene – Pliocene volcanic sequences, occurring in the northern part of the Turkish – Iranian high plateau and the Lesser Caucasus, which are characterized by calc-alkaline affinities reminiscent of extrusive rocks forming at active convergent margins (Pearce 1982; Wilson 1989; Thirlwall et al. 1994). The latest magmatic pulse in the Plio-Quaternary is represented by alkaline rocks that occupy much of the southern part of the Turkish –Iranian plateau and the western Lesser Caucasus, and that show withinplate basalt geochemical characteristics (Pearce et al. 1990; Yilmaz et al. 1998; Keskin 2003; Kheirkhah et al. 2009). These variations in the lava chemistry of the late Cenozoic volcanic rocks (Miocene to Quaternary) indicate a geochemical progression from calcalkaline to more alkaline compositions in time and a spatial shift from north to south towards the Arabian plate. The geological factors that controlled the temporal and spatial distribution of the Cenozoic magmatic rocks in the hinterland of the Arabia – Eurasia collision zone and the melt regimes and tectonic settings of their formation are outstanding questions both in the geodynamics of the eastern Mediterranean region and in continental magmatism in young orogenic belts. In this paper, we present new geochemical data from representative Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan, filling a major gap in our knowledge of the post-collisional magmatism in the peri-Arabian region, and we use these data to infer the petrogenesis of these rocks in order to interpret their melt sources and magmatic evolution.


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Figure 2. Tectonic map of the eastern Mediterranean – Persian Gulf region, showing the main plate boundaries, collision zones, distribution of Neotethyan ophiolites and Eocene volcanic sequences, microcontinental fragments with Arabian affinity, and Tauride ribbon-continent with Gondwana (Afro-Arabian) origin. Major magmatic belts (i.e. Ahar –Arasbaran, Urumieh-Dokhtar) and volcanic units (i.e. Maden complex, Kislakoÿ volcanics) discussed in the text are also shown. CACC, Central Anatolian crystalline complex; DSF, Dead Sea fault; EAAC, East Anatolian accretionary complex; EAF, East Anatolian fault; EF, Ecemis fault; IAESZ, Izmir – Ankara – Erzincan suture zone; ITSZ, Inner-Tauride suture zone; KOTJ, Karliova triple junction; MTJ, Maras¸ triple junction; MZMM, Mishkana – Zangezur metamorphic massifs; NAF, North Anatolian fault; SASZ, Sevan – Akera suture zone.

We also describe the spatial and temporal distribution of the Cenozoic volcanic rocks in nearby Iran, Armenia, and eastern Turkey, and compare their geochemical features to those of the coeval volcanic units in Azerbaijan. Finally, we evaluate the petrogenetic and tectonomagmatic evolution of the Cenozoic magmatism in the Lesser Caucasus and in the R Figure 1. Simplified tectonic map of the eastern Mediterranean – Persian Gulf region, showing the active plate boundaries, plate convergence vectors (in green) with respect to fixed Eurasia, and postcollisional volcanic rocks in the peri-Arabian region. Continental blocks with Afro-Arabian (Gondwana) affinity are shaded in light yellow. AF, Aksu fault; ASF, Aras fault; BF, Burdur fault; DSFZ, Dead Sea fault zone; EAF, East Anatolian fault; EF, Ecemis fault; EKP, Erzurum – Kars plateau; HT, Hellenic trench; IAESZ, Izmir – Ankara – Erzincan suture zone; ITSZ, Inner-Tauride suture zone; MTJ, Maras¸ triple junction; NAF, North Anatolian fault; NEAF, Northeast Anatolian fault; PSF, Pampak – Sevan fault; TF, Tabriz fault; TGF, Tuzgo¨lu¨ fault.

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peri-Arabian region in a simple geodynamic model, which we present here as a working hypothesis to be further tested with future studies particularly in Iran and Azerbaijan. Regional geology


Much of the peri-Arabian region north of the Bitlis – Zagros suture zone is occupied by the Turkish –Iranian High Plateau, where the mean surface elevation is about 2 –2.5 km above sea level with scattered Plio-Quaternary volcanic cones over 5 km high (e.g. Mt Ararat; Figure 3; Dhont and Chorowicz 2006). The plateau is bounded on the north by the Eastern Pontide arc and the Lesser Caucasus, and on the south by a series of continental blocks including the Bitlis– Pu¨tu¨rge (B –P) massifs in Turkey and the Sanandaj – Sirjan (S –S) massif in Iran (Figures 2 and 3). The basement geology of the plateau is composed of ophiolites and ophiolitic me´langes, latest Cretaceous and Cenozoic flysch and molasse deposits, and the eastward extension of the Tauride microcontinent in the Munzur carbonate platform and in the South Armenian Block (Figure 2). Eastern Pontide block The Eastern Pontide block north of the Turkish – Iranian plateau mainly consists of a south-facing Jurassic – Late Cretaceous volcano-plutonic arc that developed over a subduction zone dipping northwards (Yilmaz et al. 1997), and post-collisional Eocene volcano-sedimentary units and plutons. The collision of the Eastern Pontide arc with the Eastern Tauride – South Armenian continental block in the early Eocene terminated the subduction zone magmatism in the Pontides and produced extensive flysch deposits with intense folding in and across the collision zone (Dewey et al. 1986). Lesser Caucasus The Lesser Caucasus includes the Transcaucasian Massif in the north, the Sevan –Akera suture zone (SASZ) with ophiolite exposures in the centre, and the Miskhana – Zangezur metamorphic massifs (MZMM) in the south (Figure 2), which represent a continental fragment (Khain and Kornousky 1997; Golonka 2004). A Cretaceous island arc complex with calc-alkaline to alkaline extrusive rocks, and pyroclastic deposits, flysch units, and marl-limestone rocks occurs north of the suture zone. Eocene and Plio-Quaternary volcanic and plutonic rocks are widespread in the Lesser Caucasus and are described in the next section. The Transcaucasian Massif includes Pan-African orogenic crust intruded by latest Proterozoic to Palaeozoic granitoids, which are multiply deformed and migmatized, and by Jurassic to Early Cretaceous plutons representing a magmatic arc (Zakariadze et al. 2007). This arc continues into the Eastern Pontide block in the west. The Transcaucasian Massif was already accreted to the southern continental margin of Eurasia by 350 Ma. The SASZ includes Late Jurassic – Early Cretaceous suprasubduction zone ophiolites, which were emplaced southwestwards onto the MZMM by the Late Cretaceous (Khain and Kornousky 1997). This suture zone and the ophiolites continue northwestwards into Armenia, and then into northeastern Turkey, where they connect with the Izmir – Ankara – Erzincan suture zone (IAESZ) and the Northern Neotethyan ophiolites (Dilek and Thy 2006). The Miskhana – Zangezur massifs consist of late Proterozoic to early Palaeozoic schist, amphibolite, and marble units, unconformably overlain by Devonian and younger metasedimentary rocks (Khain and Kornousky 1997; Rolland et al. 2009a, 2009b). This continental fragment is a likely counterpart of the South Armenian Block to the northwest (Figure 1).


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Eastern Tauride and South Armenian blocks The Eastern Tauride block, part of the Tauride microcontinent occupying much of southern Turkey, is represented by the Upper Triassic– Cretaceous Munzur platform in the ¨ zgu¨l and Turs¸ucu 1984). The Tauride microcontinent consists of region (Figures 2 and 3; O late Proterozoic – Palaeozoic and Mesozoic carbonate, siliciclastic, and volcanic rocks ¨ zgu¨l 1976; Demirtas¸li et al. 1984) and represents a ribbon continent rifted off from the (O northwestern edge of Gondwana (Robertson and Dixon 1984; Sengo¨r et al. 1984; Dilek and Moores 1990; Garfunkel 1998). The Palaeozoic –Jurassic tectonostratigraphic units in the Tauride microcontinent are tightly folded and imbricated along major thrust faults. Southwest of the Munzur platform, the Eastern Tauride block includes Lower Cretaceous carbonates, overlain by Maastrichtian –lower Eocene pelagic and hemipelagic limestones (Akdere Formation; Robertson et al. 2006). These units are unconformably overlain by middle Eocene conglomerate, sandstone, and shale with no major tectonic break (Perinçek and Kozlu 1984), indicating that sedimentation was nearly continuous throughout the Mesozoic and early Palaeogene. The Munzur platform carbonates are tectonically overlain by the Ovacik me´lange (Figure 4), consisting of blocks of serpentinites, metamorphic rocks, and pelagic ¨ zgu¨l and Tursucu 1984). Both the Ovacik limestones in a fine-grained, phyllitic matrix (O me´lange and Munzur carbonates are thrust to the south over the Keban – Malatya metamorphic rocks (Figure 4) that consist of Permian to Cretaceous metacarbonate rocks, micaschist, phyllite, meta-clastic rocks, and meta-chert (Michard et al. 1984; Perinçek and Kozlu 1984). The Keban – Malatya metamorphic units likely represent the metamorphosed (greenschist facies) passive margin sequence of the northern edge of the B –P continental block, facing a Neotethyan seaway to the north (Robertson et al. 2006; Dilek and Sandvol 2009). The South Armenian Block constitutes the northeastern extension of the Tauride microcontinent. It includes a Proterozoic crystalline basement, overlain by Palaeozoic – Mesozoic sedimentary sequences (Rolland et al. 2009a; Sosson et al. 2009), reminiscent of the Eastern Tauride block. It was accreted to the Eurasian margin in the latest Cretaceous– early Palaeogene as the marginal basin south of the Eurasian continental margin collapsed and closed (Rolland et al. 2009a).

B –P massif and S– S zone The B – P massif to the south is an approximately E – W-trending continental block (Figures 2 and 4) that was rifted from Arabia in the Permo-Triassic. It is bounded by ophiolitic thrust sheets, me´langes, and Upper Cretaceous and younger volcanic and volcaniclastic rocks. The Pu¨tu¨rge massif is composed of pre-Triassic gneisses and micaschists, and granitoids (Michard et al. 1984; Aktas¸ and Robertson 1990). The Bitlis massif consists of a Precambrian crystalline basement, metamorphosed Palaeozoic – Triassic carbonate rocks (Go¨ncuöglu and Turhan 1984; Helvaci and Griffin 1984), and Palaeozoic to late Mesozoic R Figure 3. Modern topography of the Arabia – Eurasia collision zone and the Turkish –Iranian high plateau, bounded to the north by the Eastern Pontide arc (Turkey), Greater Caucasus Mountains (Russia), and Elborz Mountains (Iran). Major active faults, regional tectonic entities, stratovolcanoes (marked in red) and lakes are shown. White arrows show relative plate motions (direction and velocity in mm/year) with respect to fixed Eurasia based on the GPS data of McClusky et al. (2000).



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Figure 4. Simplified geological map of Eastern Anatolia and the Arabian foreland, showing the distribution of major tectonic units in the region and the post-collisional volcanic rocks in the Turkish high plateau. Munzur Platform constitutes the eastern extension of the platform carbonates and basement rocks of the Eastern Tauride ribbon-continent. B– P massif is a rifted off fragment of the Arabian plate, analogous to the Eastern Tauride block. The Turkish high plateau is covered by Miocene – Quaternary volcanic rocks; its basement is composed of Tethyan ophiolites and ophiolitic me´langes, flysch and molasses deposits, and platform carbonates of the Tauride block. Kaçkar batholith and the Jurassic –Upper Cretaceous sandstone, volcanic tuff, and limestone in northern Turkey constitute the Eastern Pontide Arc. EAF, East Anatolian fault; EAFZ, East Anatolian fault zone; EKP, Erzurum – Kars plateau; IAESZ, Izmir – Ankara – Erzincan suture zone; KOTJ, Karliova triple junction; NAFZ, North Anatolian fault zone; NEAF, Northeast Anatolian fault.

granitoids. Oberha¨nsli et al. (2008) reported a regionally distributed high-pressure/lowtemperature overprint in its metamorphic evolution. The entire Bitlis massif displays a doubly plunging, multiply folded anticlinorium with overturned limbs both to the north and south (Dilek and Moores 1990). The relatively youngest thrust faults are south vergent and synthetic to the Bitlis suture. Both the Bitlis and Pu¨tu¨rge massifs and the overlying volcanic and ophiolitic rocks are structurally underlain in the south by an Upper Cretaceous –early Tertiary me´lange, which is underthrust to the south by the foreland sedimentary sequences of the Arabian plate (Figure 4). The eastward extension of the B – P continental block is represented by the S –S zone, which extends for , 1500 km along strike from northwest (Sanandaj) to southeast (Sirjan) in western Iran (Figures 2 and 3; Emami et al. 1993; Mohajjel and Fergusson 2000). It is , 150 –200 km wide and consists mainly of late Proterozoic – Mesozoic meta-carbonates,


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schist, gneiss, and amphibolite that are intruded by deformed to undeformed granitoid plutons (Barberian and King 1981; Mohajjel and Fergusson 2000; Moritz et al. 2006; Mazhari et al. 2009). Metamorphosed Triassic –Lower Jurassic volcano –sedimentary sequences within the S – S zone are interpreted to represent rift-drift units associated with the early-stage evolution of the Southern Neotethys (Alavi and Mahdavi 1994; Mohajjel et al. 2003). Middle Jurassic to Late Cretaceous, medium- to high-pressure metamorphism and deformation recorded in the S – S rocks were related to the subduction of the Southern Neotethyan seafloor northeastwards beneath this continental block (Barberian and King 1981; Moritz et al. 2006). The general structural fabric is defined by NW-trending and SW-overturned folds, SW-vergent thrust faults, and NW-trending reverse faults that collectively resulted in crustal thickening in the S – S zone. This contractional fabric was overprinted by regional-scale, right-lateral transpressional deformation as evidenced by a pervasive sub-horizontal stretching lineation and dextral shearing (Mohajjel and Fergusson 2000). Major magmatic episodes in the tectonic evolution of the S –S zone are represented by widespread Late Jurassic –Cretaceous, calc-alkaline plutons intruded into the crystalline basement, and by Eocene shoshonitic granitoids crosscutting all its structural fabric elements (Ghasemi and Talbot 2006; Mazhari et al. 2009). This Eocene magmatic pulse is coeval with the magmatism in the Urumieh – Dokhtar arc (or the Central Iranian Volcanic Belt) to the NE (Figure 2). Tethyan ophiolites The Jurassic(?) – Cretaceous ophiolites underlying the molasse deposits and the Tertiary volcanic cover in the Turkish– Iranian High Plateau and in the Lesser Caucasus represent the remnants of a Mesozoic Tethyan ocean and are commonly displaced southwards onto the margins of the Eastern Tauride platform (Munzur platform), South Armenian Block, B –P massif, and S –S Zone (Figures 2 and 4; Dilek and Moores 1990; Ghasemi and Talbot 2006; Mazhari et al. 2009; Rolland et al. 2009a). The ophiolites resting tectonically on the Eastern Tauride and South Armenian Blocks were derived from the IAESZ (Figures 2 and 4) between the Eastern Pontide block and the Tauride microcontinent. The coeval ophiolites resting tectonically on the B –P massif and the S– S Zone farther south (Figures 2 and 4) were derived, on the other hand, from a separate Neotethyan basin that had evolved along the northern periphery of Arabia throughout the Mesozoic (Robertson and Dixon 1984; Sengo¨r et al. 1984; Dilek and Moores 1990; Dilek et al. 1999a). Eastern Anatolian and Urumieh –Dohktar magmatic arcs A regional, late Mesozoic to Eocene magmatic arc system extends along the northern edge of the B – P and S –S continental blocks immediately north of the Arabian plate (Figure 2). The Late Cretaceous Neotethyan ophiolites and the B – P and Keban – Malatya metamorphic units in southeastern Turkey are crosscut by kilometre-scale granitoid plutons (Perinçek and Kozlu 1984; Yazgan and Chessex 1991; Parlak 2006), which have I-type, calc-alkaline geochemical affinities (Parlak 2006). The Baskil magmatic sequence (Figure 4; in the Elazig –Palu nappe of Yazgan 1984) north of the B –P massif consists of calc-alkaline intrusive and extrusive rocks, with overlying Campanian– Maastrichtian volcaniclastic and flysch deposits (Michard et al. 1984; Yazgan 1984). The Santonian – Campanian (85 – 77 Ma) granodiorite, tonalite, quartz monzonite, monzodiorite, diorite, and gabbro rocks of the Baskil igneous sequence represent a magmatic arc constructed


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over and across a tectonic assemblage of Neotethyan oceanic crust and microcontinental rocks (Michard et al. 1984; Yazgan 1984). Therefore, the construction of this magmatic arc largely postdated the tectonic imbrication of the Cretaceous ophiolites and microcontinental units. An Eocene magmatic episode overprinted the previously formed latest Cretaceous magmatic arc system along the B – P and S – S continental blocks and developed extensively in the southeastern Anatolian orogenic belt (Turkey) and the Urumieh – Dokhtar magmatic zone (and the Central Iranian Volcanic Belt) in Iran (Figure 3; Emami et al. 1993). These Eocene magmatic occurrences are covered in detail in the next section. Peri-Arabian Cenozoic volcanism Cenozoic volcanic rocks occur extensively in Iran, Armenia, Georgia, and eastern Turkey around the northern and eastern periphery of the Arabian plate (Figures 1 and 3). In this section, we describe these units in order to compare their geology and main geochemical features with those of the main volcanic sequences we have studied in the Lesser Caucasus in Azerbaijan.

Iran Cenozoic volcanic rocks in Iran include Eocene, upper Miocene, and Pliocene – Quaternary sequences, and occur mainly around the southern periphery of the Caspian Sea, in several major eruptive centres and volcanoes along the eastern part of Lake Urmiyeh, and in the Ahar– Arasbaran and Central Iranian Volcanic Belts (Figures 2 and 3; Emami et al. 1993). Eocene (50 – 39 Ma) trachyandesite, trachyte (locally sanidine and analcime bearing), and basanite rocks of mainly shoshonitic affinity crop out in the Azerbaijan – Alborz – Sabzevar Zone (specifically in the Ahar–Arasbaran volcanic belt in the Azerbaijan province of northern Iran) and SW of Tabriz city in northern Iran (Lotfi 1975; Lescuyer and Riou 1976; Comin-Chiaramonti et al. 1979; Alberti et al. 1981; Haghipour and Aghanabati 1985; Aftabi and Atapour 2000). The Sahand volcano (Sh in Figure 3) south of Tabriz also includes shoshonitic lower Eocene breccia tuffs, porphyritic trachyandesites, and analcime-bearing trachytes. Volcanism here appears to have continued intermittently during the early Eocene, Miocene, and then in the Quaternary (Didon and Germain 1976). The Eocene shoshonitic volcanism in northern Iran extends into the Central Iranian Volcanic Belt along a NW – SE-trending linear zone (Figure 3). This volcanic belt contains lower Eocene basalts, trachybasalts, trachytes, and trachyandesites in the Qom-Aran area in its northern segment (Emami 1981; Amidi et al. 1984), and slightly more evolved shoshonites composed of middle to upper Eocene absarokites and basaltic lavas, tephrites, phonolites, and tephritephonolites in the Natanz – Nain and Shahrebabak areas in its central parts (Moradian 1990; Hassanzadeh 1993). Upper Eocene trachybasalt and trachyandesite occur in the Rafsanjan area (Aftabi and Atapour 1997) and absarokite, shoshonite, latite, and analcime-rich pyroclastic rocks crop out in the Bardsir area (Atapour 1994) in the southern end of the belt. Miocene and younger volcanic rocks in Iran occur mainly in the north, near the Turkish and Azerbaijan borders (Emami et al. 1993). The Saray volcano east of Lake Urmiyeh, one of the major eruptive centres in northern Iran, is composed of upper Miocene basanite, leucite tephrite, and associated pyroclastic rocks in the lower volcanic units, and phonolite, trachyte, and analcime basanite in the upper volcanic units

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(Moine-Vaziri et al. 1991). The Sahand volcano contains upper Miocene banakites, and the Takab – Qorveh area to the south includes shoshonitic lavas composed of absarokite, banakite, and quartz latite (Atapour 1994). The Bijar area near the Zagros fold-thrust belt consists mainly of potassic ignimbrites (Innocenti et al. 1982). In the Alborz region, shoshonitic rocks of Quaternary age occur at Damavand (Brousse et al. 1977; Darvishzadeh 1983) and similar rocks of Miocene – Pliocene age crop out in the Sabzevar area (Spies et al. 1984). Armenia Late Cenozoic (late Pliocene – Holocene) volcanism in Armenia occurred mainly in the Western and Eastern volcanic belts (Karapetian and Adamian 1973; Shirinian 1975; Mitchell and Westaway 1999; Badalyan 2000), in the southern part of the country. There is no evidence of an early Cenozoic magmatic episode. The Western volcanic belt extends north into the Greater Caucasus in Georgia and Russia (i.e. Kazbeg and Elbruz Mountains; Figure 1) and continues west into the Erzurum – Kars plateau (Erzurum – Kars volcanic plateau, EKP; in Figure 1) in NE Turkey. Large cinder cones and domes occur along major strike-slip fault systems in this belt. The Eastern volcanic belt, situated W –SW of Lake Sevan, trends in a NW – SE direction and forms the eastern extension of the Turkish high plateau (Figures 1 and 3). The Aragats volcano (At in Figure 3) in this belt is the northern extension of Mount Ararat (Ar in Figure 3) in eastern Turkey. The Gegham and Javakhet plateaus (Figure 3), with elevations generally . 3000 m, occur in the Eastern volcanic belt, and continue southeastwards into the Lesser Caucasus in Azerbaijan (Talysh region; Figure 3). Volcanism here also appears to be spatially associated with major dextral strikeslip (i.e. Garni – GF and Pampak – Sevan – PSF faults; Figure 3) and oblique-normal faults (Karakhanian et al. 2002). The early stages of the late Pliocene volcanism (3.5 Ma) were characterized by fissure eruptions of olivine basalts along fault systems mainly within the Western volcanic belt. As volcanism evolved from fissure eruptions to central eruptive centres, its character changed from mafic to silicic. Quaternary volcanism was more widespread in the Eastern belt than in the Western belt and produced more than 600 well-preserved monogenetic volcanic centres (i.e. cinder cones, domes, and lava flow fields; Karapetian and Adamian 1973). The main rock types of this phase include andesitic basalts, andesites, dacites, rhyolites and associated pyroclastic rocks (Karapetian 1963; Shirinian 1975; Karapetian et al. 2001). Eastern Anatolia, Turkey Cenozoic volcanism in eastern Anatolia (Turkey) occurred in spatially and temporally discrete zones. Early Cenozoic magmatism was limited to the Eocene in the Eastern Pontide block in the north and the southeastern Anatolian orogenic belt in the south (Figure 2). Late Cenozoic volcanism, on the other hand, affected much of eastern Anatolia occurring in discrete pulses in the late Miocene, Plio-Pleistocene and Quaternary (Figure 4). Eastern Pontide block Eocene volcanism was extensive throughout the Eastern Pontides (Robinson et al. 1995). It is represented by dominantly basalt, tephrite, andesite, dacite, and associated pyroclastic



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rocks, which disconformably overlie Upper Cretaceous basement rocks in the northern part of the Eastern Pontides (Figure 4). These volcanic rocks are alkaline in composition and commonly include phenocrysts of augite, olivine, plagioclase, phlogopite, nepheline, sanidine, cancrinite, and Fe – Ti oxides (Sen et al. 1998; Arslan et al. 2000). U – Pb isotope dating of zircon and titanite indicates that these rocks have latest Palaeocene –early Eocene ages (Hoskin et al. 1998; Arslan et al. 2000). Strong LREE-enrichment and enrichment in the LILEs of the alkaline series suggest that the mantle source region beneath the Eastern Pontides was heterogeneously enriched by subduction-related metasomatism prior to the Eocene magmatism (Arslan et al. 1997; Sen et al. 1998). Adakitic andesite and dacite rocks are also extensive in the Eastern Pontides, particularly in the Gu¨mu¨shane area (Y. Eyu¨boglu, personal communication 2008; Karsli et al. 2009). These high-K calc-alkaline rocks show enrichment in LILEs, depletions in Nb, Ta, and Ti, and high La/Yb and Sr/Y ratios (Karsli et al. 2009). They have been dated at 48– 50 Ma (40Ar/39Ar), giving a narrow age range span in the Ypresian –Lutetian (Karsli et al. 2009). In the southern part of the Eastern Pontides, the Eocene volcanic sequence consists mainly of basalt, andesite, and associated pyroclastic rocks that contain plagioclase, augite, hornblende, biotite, and lesser Fe – Ti oxide and quartz. These rocks are mainly calc-alkaline and low- to medium-K in composition, and are intercalated with clastic sedimentary rocks. Along the boundary between the Eastern Pontide block and the Erzurum – Kars plateau (EKP) to the southeast, the Eocene volcanic rocks of the Narman group rest unconformably on deformed flysch units and ophiolites. Known as the Kislakoÿ volcanic rocks, these andesitic lavas and pyroclastic rocks are exposed beneath the dacitic tuff and epiclastic rocks of the earliest volcanic associations (late Miocene) of the Erzurum –Kars plateau (Keskin et al. 1998). These rocks have a K/Ar age of 38.5 ^ 0.7 Ma (Keskin et al. 1998), confirming their eruption in the middle– late Eocene. In the southwestern part of the Eastern Pontide block, the Eocene magmatism is represented by E– W- to NE – SW-trending and fault-bounded volcano-sedimentary units, which are intruded by granitoid-syenitoid plutons. These plutons are part of the much larger Kaçkar batholith (Figures 2 and 4) that makes up the backbone of the Eastern Pontide block. Recent petrological, geochemical, and geochronological studies have shown that the composite Kaçkar batholith consists of Early Cretaceous (112 Ma) to late Palaeocene (52 Ma) granitoid plutons of a mature volcanic arc and late Palaeocene – Eocene monzonitic to syenitic post-collisional plutons emplaced into this arc and into their own volcanic carapace (Boztug et al. 2006, 2007). The 52.1 ^ 1.6-Ma Ko¨sedag syenitic pluton, exposed south of the North Anatolian fault zone (Figure 4), is the westernmost member of this batholith. Geochemical data from the Eocene volcanic rocks are lacking, but the geochemistry of the monzonitic to syenitic post-collisional plutons indicate that they are high-K, alkaline, and metaluminous to slightly peraluminous rocks, whose magmas were produced by mingling and mixing of coeval mantle- and crustal-derived melts (Boztug et al. 2007; Boztug 2008). Trace element geochemistry of these plutonic rocks suggests a subduction-metasomatized mantle as their melt source (Boztug et al. 2006). The Eocene volcanic units in the Eastern Pontide block appear to extend westwards in to the C ¸ orum area along the IAESZ, north of the Central Anatolian crystalline complex (CACC; Keskin et al. 2008).


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Y. Dilek et al.

Southeastern Anatolian orogenic belt Eocene volcanism in the B –P continental block to the north and the Arabian platform to the south is represented by the middle to upper Eocene volcanic sequences of the Maden complex (Figures 2 and 4; Yigitbas¸ and Yilmaz 1996; Elmas and Yilmaz 2003). The Maden complex consists mainly of basal conglomerate, sandstone, siltstone-claystone, pelagic limestone and basaltic lava flows and diabasic intrusions, which collectively lie on a metamorphic crystalline and ophiolitic basement. These relationships suggest that tectonic imbrication of the B –P metamorphic rocks and the Late Cretaceous ophiolitic units (i.e. Guleman; Figure 4) in south-directed nappe systems must have occurred prior to the formation of the Maden complex. In the upper part of the Maden complex, lava flows and pelagic deposits, which are collectively , 400 m thick, are stratigraphically intercalated. Upper Cretaceous –lower Eocene andesitic lavas and associated pyroclastic rocks of the Yu¨ksekova formation and upper Eocene andesitic-dacitic rocks of the Helete Formation occur to the north and south of the Maden complex, respectively, and form two separate calc-alkaline sequences. Andesitic volcanism in this region appears to have waned by the latest Eocene and the rocks grade upwards into upper Eocene – Oligocene flysch deposits (Yigitbas and Yilmaz 1996).

Erzurum –Kars and Turkish high plateaus Late Cenozoic volcanism in eastern Anatolia is represented by stratovolcanoes with significant relief (i.e. Nemrut, Su¨phan, Tendu¨rek, Ararat; Figures 3 and 4) in the southern part of the Turkish high plateau, and by an extensive (over 5000 km2) and relatively flat volcanic field (Erzurum– Kars plateau; Figure 4) with an average elevation of , 1.5 km in its northern part. The Erzurum – Kars plateau consists mainly of lava flows intercalated with subordinate ignimbrite units and sedimentary layers with ages ranging from 6.9 ^ 0.9 to 1.3 ^ 0.3 Ma (Innocenti et al. 1982; Keskin et al. 1998). Pleistocene scoriaceous spatter cones locally overlie this lava-ignimbrite sequence. The initial eruptive phase of the late Cenozoic volcanism in the Turkish high plateau is characterized by mafic and intermediate alkaline rocks and was followed by widespread eruptions of andesitic to dacitic calc-alkaline lavas during the Pliocene; the last volcanic phase involved the eruption of alkaline and transitional lavas throughout the Plio-Pleistocene and Quaternary (Yilmaz et al. 1987, 1998; Pearce et al. 1990; Kheirkhah et al. 2009). Most of the major stratovolcanoes in the Turkish high plateau were built during this last phase of volcanism, which continued until historical times.

Geology of Cenozoic volcanism in the Lesser Caucasus (Azerbaijan) In the Azerbaijan part of the Lesser Caucasus, the Cenozoic volcanic rocks occur in a broadly NW –SE-trending zone, which includes a series of fault-bounded troughs that are separated by structural and topographic highs. The Kelbajar trough in the northeastern part of the Lesser Caucasus in Azerbaijan contains strongly faulted, , 3 km-thick Eocene volcanogenic and sedimentary formations that are uncomformably overlain by nearly 1.5 km of upper Miocene –lower Pliocene lavas and pyroclastic rocks (Figure 5; Imamverdiyev 2001a). These volcanic formations and the NW – SE-trending oblique-slip faults are crosscut by NE – SW-orientated, high-angle normal faults that form well-defined structural grabens (Figure 5). Numerous vertical to steeply dipping and NE-striking rhyolite and dacite dikes occur within these NE-trending graben systems. These spatial



549

Figure 5. Geological map of the Palaeogene – Neogene and Quaternary magmatic units and volcanic centres in the central Lesser Caucasus, Azerbaijan.

relations suggest that felsic magmatism and NW –SE-directed extension were mostly synchronous events during the Plio-Pleistocene (Imamverdiyev and Mamedov 1996). The Kelbajar Trough is bounded to the S – SW by the Murovdag and Dalidag topographic highs that are occupied by upper Oligocene – lower Miocene granite, granodiorite, monzonite, and quartz syenite plutons (Figure 5). The NW –SE-trending Gochass synclinorium to the S –SW of the Murovdag – Dalidag high includes Upper Cretaceous basement units in the SE and upper Pliocene – Quaternary lavas and volcaniclastic rocks, Quaternary volcanoes, and their volcanic products to the NW (Figure 5). The Upper Cretaceous units consist of calc-alkaline to alkaline lavas and pyroclastic rocks, a flysch series, and marl-limestone deposits. Collectively, these units constitute a Late Cretaceous island arc complex in the Lesser Caucasus that is likely to be the eastern continuation of the Late Cretaceous arc system in the Eastern Pontide block in Turkey. The upper Pliocene –lower Quaternary volcanic units within the Gochass synclinorium are composed mostly of trachyandesite, basaltic trachyandesite, and trachybasalt (Imamverdiyev 2001b). Felsic units of the same sequence include rhyolite (mostly as domes), trachyrhyolite, perlite, and obsidian. The late Quaternary trachybasalt, basaltic trachyandesite, and trachyandesite rocks are widespread, forming a young volcanic plateau described in the literature under various names (i.e. Yaylag, Alagellar, Zar; Imamverdiyev 2000, 2001a, 2001b). This vast Quaternary plateau is dotted with numerous volcanoes, including Galingaya, Karagel, Sagliyali, Ayichingilli, Sarchali, and Sarimsagli (Figure 5). Farther southeast within the Gochass synclinorium, the eruptive centres and

550

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individual volcanoes (i.e. Boyuk and Kocuk Ishikhli, Kagramanbektepe, Lyulpar, Uchtape) become a little older, late Pliocene –early Quaternary in age. Geochemistry of Cenozoic volcanism in Azerbaijan Major and trace element analyses of selected volcanic rocks from the Lesser Caucasus in Azerbaijan are listed in Tables 1 – 3. Based on the field occurrences and the available ages, we distinguish two magmatic episodes in this region, the Eocene and the late Miocene – Quaternary. The Eocene episode is represented by a mafic to intermediate, shoshonitic rocks in the Kelbajar trough. The Miocene – Quaternary magmatic episode can be subdivided into early and late phases. The volcanic sequence of the early phase is represented by an intermediate to felsic, calc-alkaline association, currently exposed in the Kelbajar trough and formed during the late Miocene – early Pliocene. The volcanic sequence of the late phase consists of a mafic to felsic, mildly alkaline – shoshonitic association, exposed in the Gochass synclinorium and formed during the late Pliocene – Quaternary. Both mafic rock groups belonging to the early and late phases include gabbroid nodules. Chemical compositions of these nodules are also listed in Tables 2 and 3. Eocene volcanic sequence The Eocene volcanic sequence consists mainly of trachybasalt and basaltic trachyandesite with subordinate tephrite-basanite, basaltic andesite, and trachyandesite (Figure 6(a)). These units all have , 53 wt % SiO2 and are characterized by moderate TiO2 (0.83 – 1.16 wt %), medium to high Al2O3 (14.1 –19.55 wt %) and low to moderate MgO (2 –7.6 wt%). Most of the rocks have K2O . Na2O and are mildly alkaline with the exception of one subalkaline sample (Figure 6(a)). The analysed samples of this sequence plot in the shoshonitic field on a K2O vs. SiO2 diagram (Figure 7(a)). Selected major and trace elements vs. MgO variation diagrams are shown in Figure 8. In general, TiO2, Fe2O3*, CaO correlate positively with MgO whereas SiO2 and Al2O3 correlate negatively. The Eocene volcanic sequence exhibits a wide range of trace element contents (Table 1; Figure 8), with the Ba, Rb, Sr, and Ni contents being the most variable among them. Ni increases whereas Rb decreases with increasing MgO, and Nb remains almost constant with varying MgO contents. It is noteworthy that the major and trace element compositions of the Eocene sequence (with the exception of Rb, which is higher in the Eocene sequence) overlap with those of mafic lavas belonging to the late Pliocene –Quaternary sequence. Plots of MgO against major oxides and trace elements display variations similar to those of the mafic lavas of the late phase, as well (Figure 8). This observation is also supported by N-mid-ocean ridge basalt (MORB)-normalized (Sun and McDonough 1989) and chondrite-normalized (Boynton 1984) multi-element diagrams (Figures 9 and 10). In Figure 9, all units of the Eocene volcanic sequence display similar patterns to those of the mafic lavas belonging to late Pliocene – Quaternary sequence. They all show enrichment in the most incompatible elements (Ba, Rb, Th, K, La, Ce), troughs in Nb, Ta, Zr, and a nearly flat trend in Ti, Y, Yb. Miocene –Quaternary volcanic sequence All the upper Miocene – lower Pliocene lavas of the early phase in this sequence are sub-alkaline andesite, trachyandesite, trachydacite, dacite, and rhyolite (Figure 6(a);

51.09 1.14 16.28 8.32 0.17 4.92 9.62 2.97 3.00 0.48 1.36 99.46

71 656 669 132 17 1.09 3.49 7.70 20 32 59 5.80 1.68 0.83 2.32 0.34 26

113 761 897 169 18 1.00 4.38 11.35 14 36 67 5.78 1.61 0.75 2.25 0.34 24

TA10-3* sill

55.75 0.74 19.55 5.69 0.12 2.01 7.05 3.97 3.99 0.57 0.53 100.00

TA9-1* lava

83 602 637 120 16 0.91 3.18 6.88 21 29 53 5.31 1.51 0.77 2.11 0.31 24

50.44 1.13 15.70 9.82 0.16 5.92 8.45 2.67 3.41 0.46 1.37 99.63

TA19-1* sill

71 925 1009 70 15 0.98 1.85 3.79 10 23 43 4.72 1.56 0.72 1.93 0.27 23

47.17 1.16 16.35 9.07 0.26 5.18 7.25 3.49 2.82 0.41 5.59 99.26

TA39-1* sill

95 1008 1063 64 13 0.83 1.70 3.09 16 19 36 3.87 1.20 0.61 1.63 0.23 19

47.18 1.20 16.95 9.44 0.19 4.59 5.86 3.75 3.84 0.43 5.88 99.46

TA39-3* sill

Notes: Samples with (*) symbol are from Vincent et al. (2005) and n.d., Not detected.

Major oxides (wt %) SiO2 52.30 TiO2 0.86 Al2 O3 15.95 Fe2 O3 8.43 MnO 0.16 MgO 4.00 CaO 6.61 Na2 O 3.38 K2 O 5.32 P2 O5 0.52 LOI 2.39 Total 99.67 Trace elements (ppm) Rb 170 Sr 890 Ba 923 Zr 122 Nb 16 Ta 0.97 Hf 3.17 Th 9.02 Ni 37 La 34 Ce 61 Sm 5.78 Eu 1.69 Tb 0.77 Yb 2.08 Lu 0.31 Y 23

TA6-1* lava

76 528 486 89 12 0.69 2.43 3.72 93 20 39 4.51 1.44 0.72 1.91 0.28 23

52.50 1.09 14.17 7.34 0.21 7.62 10.34 2.30 2.57 0.36 1.34 99.91

TA64-1* clast

122 1152 1167 100 16 0.92 2.40 8.13 72 34 59 4.98 1.57 0.65 1.74 0.26 20

51.46 0.83 18.43 7.17 0.16 4.98 7.14 2.75 4.80 0.67 1.71 100.17

TA66-1* clast

77 686 718 141 18 1.05 3.68 8.15 23 34 63 6.17 1.77 0.90 2.46 0.35 27

51.77 1.14 16.51 9.11 0.17 5.17 9.03 2.94 2.93 0.49 1.07 100.38

TA70-1* sill

Table 1. Major and trace element compositions of representative Eocene volcanic units, Lesser Caucasus (Azerbaijan).


69 988 873 89 13 0.80 2.37 6.34 18 28 51 5.10 1.63 0.69 1.83 0.27 21

48.82 0.98 17.61 8.93 0.17 4.87 8.12 3.39 3.33 0.55 2.89 99.76

TA77-1* sill

91 599 722 174 23 1.38 4.48 11.14 18 37 67 6.21 1.68 0.89 2.55 0.38 28

51.79 1.25 17.09 8.99 0.15 4.82 8.39 3.17 3.50 0.50 0.99 99.63

TA82-1* sill

70 647 652 122 15 0.88 3.17 6.20 24 29 54 5.50 1.62 0.79 2.19 0.32 25

50.85 1.16 16.96 8.12 0.16 5.08 9.68 3.01 3.02 0.49 1.51 99.46

TA87-1* lava


15 lava

Major oxides (wt %) SiO2 61.87 62.61 TiO2 0.59 0.58 Al2 O3 15.70 16.90 3.47 3.91 Fe2 O3 FeO 1.29 1.01 MnO 0.06 0.04 MgO 1.85 1.95 CaO 4.85 4.24 Na2 O 4.19 4.07 K2 O 3.54 2.95 P2 O5 0.41 0.28 LOI 0.81 0.54 Total 98.63 99.08 Trace elements (ppm) Rb 90 63 Li 7.2 19 Sr 890 590 Ba 1240 790 Zn 65 65 Cu 20 37 Zr 150 170 Nb 10 10 Ta 0.31 n.d. Hf 2.60 n.d. U 2.70 n.d. Th 11.00 n.d. Cr 120 310 V 170 80

40 lava

61.75 0.81 14.81 3.91 2.46 0.10 3.18 6.13 3.37 2.37 0.28 0.13 98.3

45 16 850 900 70 41 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 90

55 17 600 730 54 83 150 10 0.94 3.30 4.50 10.00 180 65

190 lava

62.10 0.60 16.60 3.28 1.29 0.09 1.90 4.32 4.08 3.08 0.30 0.46 98.10

100 lava

66 15 940 650 70 25 150 10 0.82 4.00 4.70 11.00 180 100

62.04 0.79 16.25 4.81 0.72 0.09 2.02 5.04 3.18 2.57 0.40 0.36 98.27

193 lava

79 13 935 1120 52 37 n.d. n.d. n.d. n.d. n.d. n.d. 140 100

62.84 0.75 17.15 4.94 0.43 0.09 1.86 5.25 3.30 1.87 0.35 0.38 99.21

194 lava

56 14 680 920 63 20 180 10 0.82 4.00 4.70 11.00 180 70

63.80 0.49 15.41 2.50 0.94 0.06 1.77 5.34 3.93 2.73 0.38 1.96 99.31

8 lava

97 13 420 830 70 13 240 17 1.20 6.00 5.20 14.00 n.d. 40

70.62 0.27 15.77 1.69 0.43 0.04 0.05 1.32 4.57 4.14 0.06 0.27 99.23

96 lava

86 12 930 740 57 22 170 14 1.40 4.70 5.40 18.00 180 100

65.01 0.60 17.03 3.38 0.73 0.03 1.43 3.97 4.27 3.47 0.33 0.47 100.72

106 lava

72 13 790 660 59 26 150 14 1.00 4.00 3.40 15.00 100 100

64.97 0.52 16.41 3.59 0.28 0.09 1.31 3.19 4.05 2.55 0.23 0.96 98.15

74 lava

90 22 710 1070 46 31 200 14 n.d. n.d. n.d. n.d. n.d. 120

64.51 0.55 15.96 3.55 1.01 0.08 1.13 3.30 4.00 3.47 0.22 0.85 98.43

200 lava

128 n.d. 300 350 100 20 150 n.d. n.d. n.d. 4.70 10.00 30 85

70.40 0.01 15.10 1.36 1.48 0.09 1.14 0.97 2.94 3.25 0.16 2.88 99.78

953 lava

118 n.d. 150 500 300 85 200 n.d. n.d. n.d. 7.00 21.00 30 15

74.21 0.32 15.67 1.00 0.43 0.03 1.05 0.54 2.06 3.14 0.07 1.56 100.08

973 lava

11 23 460 380 100 41 110 10 0.40 2.80 2.00 4.00 710 170

49.80 1.15 8.46 5.62 3.33 0.18 12.43 13.74 2.00 0.92 0.13 0.39 98.15

190/G Gabbro nodule

11 22 1100 440 100 41 85 19 0.68 3.10 5.30 5.10 n.d. 210

45.94 1.58 13.09 9.66 1.74 0.19 8.13 13.47 2.99 0.89 1.70 0.49 99.87

194/A Gabbro nodule

22 14 1300 650 110 71 78 10 0.47 2.10 4.00 3.80 n.d. 230

51.42 1.01 17.82 6.03 2.17 0.16 5.87 8.83 3.38 1.27 0.52 0.46 98.94

194/B Gabbro nodule

Table 2. Major and trace element compositions of representative samples of the early phase (late Miocene – early Pliocene) volcanic units and gabbro nodules, Lesser Caucasus (Azerbaijan).


552 Y. Dilek et al.

Ni Co Sc La Ce Sm Eu Tb Yb Lu Y

24 20 27 45 88 4.20 1.20 0.67 1.20 0.19 11

30 35 20 43 77 3.90 1.20 0.56 1.40 0.20 16

30 15 15 36 76 4.20 1.00 0.58 1.50 0.20 14

40 35 10 23 57 7.50 1.60 1.10 3.60 0.69 29

22 30 8 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.

32 25 10 47 91 5.10 1.60 0.90 1.80 0.23 n.d.

22 30 7 37 73 3.60 1.00 0.43 1.30 0.18 16

15 9 3 47 78 5.00 0.79 0.57 1.40 0.18 16

32 30 8 47 87 3.60 1.10 0.44 1.30 0.17 10

25 15 10 38 74 4.40 0.95 0.42 1.30 0.17 9

20 30 10 53 79 6.30 1.20 0.99 1.70 0.21 7




70 51 84 23 46 7.90 2.00 1.60 3.70 0.64 25

45 38 42 68 140 14.00 3.20 1.50 3.00 0.52 27

56 24 28 43 74 6.70 1.80 0.81 1.50 0.25 21


105 lava

129 lava

134 lava

21 lava

57 lava

208 lava

19/P lava

53 lava

87 lava

109 lava

36/P lava

120 lava

167 lava

174 Lava

180 13 25 33 Lava Lava Lava Lava

143 Lava

160 Lava

185 Lava

73/P Lava

13/3-V 12 6-174 Gabbro Lava Lava nodule

32 34 37 27 8 9 8 9 1360 1490 2400 2600 1020 990 1300 1170 100 56 100 120 70 67 90 90 259 244 200 338 28 35 35 42 0.92 0.96 1.20 1.70 5.20 5.10 4.50 4.60 3.00 3.00 5.20 7.40 2.60 4.90 4.00 3.80 280 450 170 220 210 260 140 220 110 100 43 64 60 24 26 50 21 26 20 20 63 62 76 77 130 120 150 160 9.80 9.10 10.00 11.00

53 13 1900 1170 110 66 250 23 1.50 5.20 3.00 8.10 n.d. 150 45 45 20 77 160 9.50

43 12 1780 1267 78 58 240 20 1.80 4.75 3.00 5.70 28 96 25 21 14.5 73.5 135 8.40

34 36 39 13 12 13 1420 1615 1130 980 1000 1240 100 95 150 50 21 45 222 210 252 42 21 28 0.80 0.99 n.d. 4.00 4.70 n.d. 4.00 5.30 n.d. 6.00 4.00 n.d. 141 200 261 200 200 150 38 48 25 18 50 27 10 20 10 59 66 69 120 130 130 6.30 7.40 7.40

42 13 1433 1054 66 35 235 25 1.20 4.90 4.60 7.20 27 129 25 21 4.9 72 130 8.10

55 17 730 800 80 46 222 19 1.20 4.40 4.00 5.60 160 130 30 14 20 52 98 5.90

37 14 1700 900 100 46 250 23 1.40 4.80 2.80 6.40 n.d. 240 44 21 14 69 120 7.40

43 13 1700 900 140 21 244 35 1.30 5.00 3.80 6.50 n.d. 150 19 18 16 80 160 9.80

43 55 49 66 15 10 12 16 1190 1360 1275 1615 840 830 1060 900 100 91 70 80 32 63 37 100 222 190 180 220 21 18 13 18 1.30 0.81 0.87 1.00 5.10 4.80 4.50 5.30 4.00 6.30 6.50 8.80 7.50 3.60 6.30 4.00 n.d. 160 188 100 170 80 130 100 15 50 54 50 17 45 16 20 20 20 14 11 69 60 60 70 140 120 120 120 8.00 5.70 5.30 5.80

40 14 1647 900 100 50 207 21 0.98 4.70 5.60 4.00 n.d. 140 33 40 18 59 120 7.20

56 17 1360 1016 100 28 200 23 1.40 4.70 9.50 4.00 n.d. 140 29 19 10 67 140 8.60

48 15 790 930 100 35 160 15 0.88 4.30 9.70 4.00 n.d. 110 31 13 10 48 88 5.70

70 160 180 20 67 70 1356 150 100 1100 100 100 55 100 30 41 30 2 303 100 80 33 15 10 1.43 n.d. n.d. 6.60 n.d. n.d. 3.20 9.30 12.00 12.20 25.00 31.00 140 30 n.d. 70 n.d. 20 13.5 20 3 11 5 3 6.7 n.d. n.d. 72 n.d. n.d. 115 n.d. n.d. 6.00 n.d. n.d.

33 11 1400 500 95 112 260 28 n.d. n.d. n.d. n.d. 322 190 151 24 35 29 72 5.90

48.88 48.05 51.84 49.42 52.97 50.50 53.32 53.05 54.92 54.90 55.67 54.31 54.01 55.21 57.66 58.52 59.85 57.08 59.28 57.85 67.80 73.99 75.51 51.41 1.57 1.45 1.36 1.44 1.30 1.18 0.97 1.14 1.14 0.92 1.08 1.18 1.50 1.52 0.79 0.82 0.80 1.24 1.24 0.75 0.48 0.01 0.01 1.45 15.86 15.53 16.64 16.27 16.46 17.70 17.39 17.46 16.38 17.60 17.13 16.82 17.49 16.99 16.41 16.23 16.67 17.25 16.55 17.70 15.70 13.48 13.79 18.73 5.61 3.55 6.11 7.16 7.04 7.00 6.11 5.66 4.54 7.00 6.59 5.02 5.79 3.69 4.09 4.80 4.88 4.62 4.95 3.79 4.00 1.20 0.55 5.97 2.73 4.46 1.01 0.72 0.30 0.80 0.57 1.65 2.59 0.30 0.43 2.17 2.46 3.90 1.87 0.87 0.50 3.09 1.30 1.88 3.00 1.78 0.71 1.59 0.14 0.13 0.11 0.12 0.12 0.15 0.10 0.13 0.10 0.13 0.12 0.12 0.12 0.12 0.05 0.09 0.11 0.11 0.10 0.13 0.05 0.01 0.01 0.12 6.29 6.81 4.42 5.27 3.65 5.30 3.81 4.12 3.76 3.90 4.66 3.84 3.37 2.50 3.18 3.23 2.67 2.29 2.79 2.77 1.10 0.14 0.36 4.89 9.09 9.19 8.58 9.10 7.00 9.20 7.17 6.71 6.88 7.10 6.24 6.66 6.80 5.96 6.25 6.24 5.61 6.09 5.82 6.12 2.20 0.53 1.90 9.58 4.00 4.18 4.14 3.22 4.39 4.50 5.03 4.27 3.70 4.60 4.22 4.78 4.53 5.04 3.85 4.00 4.38 4.53 4.65 4.53 5.50 3.27 2.92 4.11 1.92 1.73 2.92 2.48 3.16 2.90 2.80 2.77 2.17 3.00 2.60 2.96 3.25 3.11 3.01 2.80 3.11 2.87 3.46 2.89 4.00 4.87 3.96 1.61 1.18 1.13 1.31 1.04 0.93 0.89 0.82 0.83 0.94 0.78 0.58 0.75 0.94 0.91 0.57 0.68 0.79 0.68 0.76 0.44 0.35 0.01 0.01 0.40 0.93 1.79 0.61 1.90 1.10 1.00 0.14 0.35 0.85 1.00 0.41 0.19 0.44 0.02 0.64 0.40 0.35 0.27 0.20 1.15 1.00 0.38 0.54 0.39 98.20 98.00 99.05 98.14 98.42 100.12 98.23 98.14 98.47 100.23 99.73 98.80 100.70 98.97 98.37 98.68 99.72 100.12 101.10 100.00 100.18 99.67 100.27 100.25

132 lava

Major and trace element compositions of representative samples of the late phase (late Pliocene – Quaternary) volcanic units and gabbro nodules, Lesser Caucasus (Azerbaijan).

Major oxides (wt %) SiO2 51.23 48.35 TiO2 1.39 1.20 Al2 O3 16.49 15.77 Fe2 O3 7.74 6.38 FeO 0.86 2.16 MnO 0.13 0.15 MgO 6.04 6.74 CaO 8.33 9.80 Na2 O 4.22 3.61 K2 O 1.42 1.96 P2 O5 0.65 1.03 LOI 0.70 1.50 Total 99.20 98.65 Trace elements (ppm) Rb 16 30 Li 10 9 Sr 910 1310 Ba 600 1040 Zn 49 64 Cu 75 71 Zr 178 229 Nb 35 35 Ta 0.92 0.92 Hf 4.60 4.70 U 3.00 3.00 Th 2.60 3.20 Cr 310 412 V 165 170 Ni 100 93 Co 30 26 Sc 15 18 La 40 65 Ce 81 130 Sm 5.30 9.50

Table 3.


554 Y. Dilek et al.

Eu Tb Yb Lu Y

1.70 0.88 2.40 0.42 31

2.50 1.50 2.70 0.39 30

2.50 1.30 2.40 0.33 34

2.40 1.10 2.20 0.31 29

2.50 1.00 1.80 0.22 16

2.80 1.30 1.90 0.34 23

2.50 1.30 2.30 0.34 23

2.15 1.35 2.35 0.27 15

1.60 1.00 1.80 0.25 16

1.80 1.40 2.10 0.28 16

2.00 1.10 2.00 0.22 24

1.95 1.05 2.35 0.33 15

1.70 0.90 2.00 0.39 21

2.20 1.10 2.20 0.31 27

2.70 0.95 2.00 0.27 25

2.30 1.40 2.20 0.35 27

1.60 1.10 1.80 0.31 24

1.70 0.94 1.90 0.30 32


1.70 0.85 2.00 0.26 32

2.00 1.80 2.20 0.25 16

2.00 1.20 2.10 0.24 19

1.40 0.59 1.30 0.24 15

1.50 1.12 2.10 0.25 10

n.d. n.d. n.d. n.d. n.d.

n.d. n.d. n.d. n.d. n.d.

1.80 1.70 2.30 0.37 19



556 Y. Dilek et al.



557

Imamverdiyev 2001b). On a K2O vs. SiO2 diagram, the volcanic units of the early phase plot in the high-K calc-alkaline field (Figure 7(a)). The rhyolites of this phase are rich in K with 3 – 4 wt % K2O. The calc-alkaline nature of the early phase lavas is also evident from major and trace element compositions (Table 1). These lavas are strongly depleted in highly compatible elements but moderately to strongly enriched in highly incompatible elements (Ba, Th, La; Figures 8– 10), yielding high Th/Yb and Zr/Y ratios. Volcanic units of the late phase define a bimodal series with a silica compositional gap between the felsic lavas (68 –75.5 wt % SiO2) and mafic ones (48 –59 wt % SiO2; Figure 6(a)). Majority of the lavas in the latter group are composed predominantly of mildly alkaline lavas, including trachybasalt, basaltic trachyandesite, basaltic andesite, trachyandesite, and rhyolite (Figure 6(a)). Rhyolites (and subordinate trachyte) of this phase have higher K2O contents than those rhyolitic rocks of the early phase. Volcanic units of the late phase plot in the fields of the high-K calc-alkaline and shoshonitic series (Figure 7(a)). For both the early and late phase volcanic units, Fe2O3*, CaO, and TiO2 correlate positively with MgO (Figure 8), although rocks of the early phase have lower contents of Fe, Ti, Ca with lower MgO wt %. The rocks of the late phase show a much wider range in MgO contents relative to the lavas of the early phase. The felsic and intermediate rocks of the early phase display a positive correlation (with a steep slope) between Sr and MgO (Figure 8). In MORB-normalized trace element diagrams, mafic to intermediate rocks of both the early and late phases are enriched in the LILE, LREE, and HFSE relative to MORB, and both have high LILE/HFSE ratios (e.g. Ba/Nb; Figure 9). By contrast, the Ti, Y, and HREE abundances are lower than those of the MORB. There is also a slight depletion in Ti in the calc-alkaline intermediate lavas of the early phase that is absent in the alkaline rocks of the late phase. The Ba/Nb ratio in the alkaline rocks is also slightly lower. Similar trace element patterns are observed in intermediate to mafic lavas from the Erzurum – Kars plateau and Suphan, Ararat, Tendurek stratovolcanoes (Figures 3 and 4) of the Turkish high plateau (and Nemrut with more pronounced troughs in P and Ti). The rhyolites that belong to the early phase have broadly similar trace element patterns to the intermediate lavas of this phase, although troughs in Sr, Ba, P, and Ti are significantly more pronounced (Figure 9). By contrast, their Nb –Ta depletion relative to the LREE is much less pronounced than in the intermediate lavas. The abundances of the REE in the mafic to intermediate lavas from both the alkaline and calc-alkaline series of the Miocene –Quaternary volcanic sequence are very similar, with no Eu anomalies (Figure 10). The rhyolites of the early phase (late Miocene – early Pliocene) have similar or slightly lower REE abundances relative to coeval R Figure 6. Total alkali vs. SiO2 classification diagrams of Cenozoic volcanic units from (a) Azerbaijan, (b) Erzurum –Kars Plateau, (c) Eastern Anatolia, and (d) Iran (Le Bas et al. 1986). I & B – Alkali – subalkali subdivision is from Irvine and Baragar (1971). Data sources: early and late phases of the Miocene– Quaternary volcanism in Azerbaijan (this study); Eocene lavas in Azerbaijan (Vincent et al. 2005); volcanic units of Ararat, Nemrut, Suphan, Tendurek, and Mus in Eastern Anatolia (Pearce et al. 1990; Yilmaz et al. 1998); early, middle and late stages of volcanism and Kislakoy volcanic rocks of the Erzurum – Kars Plateau (Keskin et al. 1998, 2003); Northern, Eastern, and Central Iran, and Azerbaijan volcanic rocks (Didon and Germain 1976; Atapour 1994; Aftabi and Atapour 2000).


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Figure 7. K2O vs. SiO2 diagrams (Peccarillo and Taylor 1976) of Cenozoic volcanic units in: (a) Azerbaijan, (b) Erzurum– Kars Plateau, (c) Eastern Anatolia, and (d) Iran. See Figure 6 for the data sources.

intermediate-mafic lavas (Figure 10(a)). Compared to the intermediate-mafic lavas, they have higher La/Sm ratios, a slight negative Eu anomaly, and depletion in the HREE, Yb, and Lu. By contrast, the rhyolites associated with alkali basalts of the late phase (late Pliocene – Quaternary) have significantly higher REE concentrations than the alkali basalt lavas and a more pronounced negative Eu anomaly. The (La/Yb)n ratios of these volcanic rocks range from 10 to 35. The early and late phase mafic volcanic sequences include gabbroid nodules with higher contents of Cr (320 –710 ppm), Ni (70 –350 ppm), and MgO (8 – 13 wt %), and lower silica content (42 – 51 wt % SiO2) than the host mafic lavas. They are more enriched in Ba, Rb, Th, K, La, Ce, and more depleted in Ta, Zr than their host basalts (Figures 8 and 9(a)). These values are also lower than expected values for primary magmas. The samples have steeply sloping chondrite-normalized REE patterns characterized by strong enrichment in LREE and slight enrichment in Tb and Lu (Figure 10(b)).

559



Figure 8. Selected major and trace element vs. MgO variation diagrams for the Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan. See Figure 6 for the data sources.

Petrogenesis of Cenozoic volcanism in the Lesser Caucasus (Azerbaijan) Both the Eocene and the Miocene – Quaternary volcanic sequences in the Lesser Caucasus of Azerbaijan show broad geochemical similarities (variations with increasing MgO and trace element patterns; Figures 8, 9(a,b) and 10(a,b)), suggesting that they were derived from similar magma source(s). These Cenozoic volcanic rocks have low contents of Cr and Ni (up to 450 and 110 ppm, respectively, for the least evolved basaltic lavas) relative to primary magmas. The Cr (up to 710 ppm), Ni (up to 350 ppm), and MgO (8 – 13 wt %) contents are higher in gabbroid nodules than in their host basalts. These gabbroid rocks


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Figure 9. N-MORB-normalized multi-element patterns for the Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan (top two panels) and in the Erzurum – Kars plateau and Eastern Anatolia (lower two panels). N-MORB normalizing values are from Sun and McDonough (1989). See Figure 6 for the data sources.

display less evolved major and trace element concentrations than the lavas, and therefore they may be closer in composition to the parental magmas. However, even in the gabbroid nodules, the MgO, Ni, and Cr contents are lower than the expected for primary melts. Generally, it is assumed that primary magmas are represented by upper mantle mineralogies having high Mg# values (. 0.7), high Ni (. 400 – 500 ppm), high Cr (. 1000 ppm), and , 50 wt % SiO2 (Taylor and McLennan 1985; Wilson 1989; Condie 2001). Therefore, the majority of the volcanic samples from the Lesser Caucasus display a broad range from slightly to highly evolved compositions, as evidenced by their variable MgO contents (1.9 –8 wt %). It is important to note that the three volcanic sequences (Eocene, late Miocene –early Pliocene, and late Pliocene –Quaternary volcanic associations) have similar trace and REE patterns. N-MORB-normalized spider diagrams for all mafic to intermediate rocks of the three volcanic sequences are characterized by troughs in Nb, Ta, Hf, and/or Zr that are stronger in felsic rocks of the early and late phases, strong enrichment in Rb, Ba, Th, La, and depletion in Ti, Yb, Y relative to N-MORB (Figure 9(a,b)). This enrichment in incompatible elements implies that the melt source from which the magmas were derived was a metasomatized lithospheric mantle, enriched in K and incompatible elements. The troughs in Nb – Ta are commonly considered as typical features of subduction-related magmatism. In subduction zones, K, Rb, Th, La are transferred into the melt in the overlying mantle wedge, whereas Nb and Ta remain behind in the solid peridotite causing depletion in Nb and Ta in the mantle wedge-generated magmas (Condie 2001). However,



561

Figure 10. Chondrite-normalized REE element patterns for the Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan (top two panels) and in the Erzurum – Kars plateau and Eastern Anatolia (lower two panels). Chondrite normalizing values are from Boynton (1984). See Figure 6 for the data sources.

the LILE enrichment in our samples appears to be high with respect to the arc basalts. The high La, Th, Ce, and Pb contents of the analysed samples are also consistent with crustal contamination. Therefore, the trace element and REE patterns of the Eocene and the Miocene – Quaternary volcanic sequences are compatible with the patterns of magmas formed in other post-collisional settings (Turner et al. 1996; Nemcock et al. 1998; Maury et al. 2000; Pe-Piper and Piper 2001; Williams et al. 2004; Zhao et al. 2009), as in the case of the Cenozoic volcanic assemblages in Eastern and Western Anatolia (Yilmaz et al. 1987; Pearce et al. 1990; Yilmaz 1990; Altunkaynak and Yılmaz 1998; Keskin et al. 1998; Aldanmaz et al. 2000; Ko¨pru¨basi and Aldanmaz 2004; Dilek and Altunkaynak 2007, 2009). The geochemical data, particularly the high Th/Nb, Ba/Nb, K/Ti ratios, and low Nb/Y and Ti/Y ratios, combined with the regional geological constraints, indicate that the mantle sources beneath the Lesser Caucasus were metasomatized by ancient subduction events, which provided K-rich and HFSE-depleted aqueous fluids. The gabbroid nodules and least-evolved basaltic lavas of both the Eocene and Miocene – Quaternary volcanic sequences have similar compositions, indicating derivation from enriched lithospheric mantle source(s). The general slope (from left to right) of the multi-element patterns is also typical of basic igneous rocks generated by small degrees of partial melting (Figure 9(a,b)). The abundances of the REE in the mafic to intermediate lavas from both the alkaline and calc-alkaline series of the Miocene – Quaternary volcanic sequences are very similar, with no Eu anomalies, indicating that the source of their magmas was


562

Y. Dilek et al.

Figure 11. Th/Yb vs. Ta/Yb diagram (after Pearce 1982) for mafic to intermediate lavas from the Eocene and Miocene– Quaternary volcanic sequences in the Lesser Caucasus of Azerbaijan.

plagioclase-free, or that plagioclase fractionation was not important during the evolution of their magma(s) (Figure 10). Subduction enrichment in the melt source region of the Eocene and Miocene – Quaternary volcanic sequences can also be detected in a Th/Yb vs. Ta/Yb diagram (Figure 11), which displays source variations and crustal contamination effects (Pearce 1982). Both the Eocene and late Miocene –Quaternary lavas show a trend that is subparallel to the mantle array but shifted towards higher Th/Yb ratios. This feature indicates a lithospheric mantle source enriched by a subduction component. There is some evidence, however, indicating that this subduction signature decreased as the effects of an asthenospheric input increased through time during the evolution of the Eocene and Miocene –Quaternary volcanic sequences. In Figure 12 (Thieblemont and Tegyey 1994), the samples from the Eocene sequence straddle the boundary between subduction-related and collision-related settings, whereas all samples from the early phase and felsic products of the late phase fall into the field of collision-related magmatic rocks. By contrast, alkaline mafic lavas of the late phase show transitional compositions between collisionrelated and intraplate lavas. These data indicate a decreasing subduction signature and an increasing asthenospheric mantle component for the rocks all the way from the middle Eocene sequence to the Miocene –Quaternary sequences. An asthenospheric upwelling overprint might have masked the subduction signature in time. This inference is also supported by the Ba/Nb vs. La/Nb relationships of these volcanic associations (Figure 13). On this diagram, lavas from the early and late phases define a linear trend between the crustal values and PM, indicating a compositional shift from the lithospheric array towards



563

Figure 12. Nb/Zr(n) vs. Zr diagram (Thieblemont and Tegyey 1994) for the Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan. (n), N-MORB normalized values (Sun and McDonough 1989).

the primitive mantle array as a result of interactions of the continental crust and old lithospheric mantle material with asthenosphere-derived magmas. The steep La/Yb trend in Figure 14 indicates that the effects of different degrees of partial melting were important for the generation of the compositional variations in magmas of the Cenozoic volcanic sequences in the Lesser Caucasus (Thirlwall et al. 1994). In this diagram, alkaline lavas of the late phase reflect small degrees of partial melting, whereas we

Figure 13. Ba/Nb vs. La/Nb diagram for the Cenozoic volcanic sequences in the Lesser Caucasus of Azerbaijan. PM, primary mantle; OIB, ocean island basalt, MORB values are from Sun and McDonough (1989); CC, Continental crust from Rudnick and Gao (2003).


564

Y. Dilek et al.

Figure 14. La/Yb vs. La (ppm) diagram illustrating the partial melting and fractionation effects. Vectors for FC and PM are from Thirlwall et al. (1994). See Figure 6 for the data sources.

see an increasing amount of partial melting effect from the early phase lavas to the Eocene volcanic sequences (Figure 14). On the other hand, the Th/Nb and Ta/Yb relationships (Pearce 1982; Figure 11) show the effects of fractional crystallization (FC) and assimilation – fractional crystallization (AFC) processes on magma evolution. In Figure 11, mafic lavas of the Eocene and the late phase of the Pliocene –Quaternary sequences define different AFC trends among the mantle, gabbroid nodules, and upper continental crust values, indicating different AFC paths from a common parental magma source. It is also apparent in Figure 11 that volcanic rocks of both the early and late phases followed different AFC paths from a common parental magma source. The bimodal nature of the volcanic units of the late phase is defined by a large silica compositional gap between the felsic (68 – 75.5 wt % SiO2) and mafic lavas (48 – 59 wt % SiO2). The major and trace element features (Figures 8 and 9) probably reflect the effects of FC during the evolution of bimodal rocks of the Pliocene – Quaternary sequence. Compatible trace elements such as Cr and Ni decrease with MgO (Figure 8), and these variations are consistent with the fractionation of a phenocryst assemblage of clinopyroxene, magnetite, and olivine. The rhyolites display broadly similar trace element patterns to the intermediate lavas of this phase, although depletions of Sr, Ba, P, and Ti are significantly more pronounced, probably reflecting fractionation of feldspar, apatite, and Fe – Ti oxides (Figure 9). Therefore, when we evaluate these features together with Th/Nb and Ta/Yb relationships (Figure 11), we infer that the compositional variations may have resulted from FC; in addition, AFC appears to have played an important role during the formation of bimodal rocks of the Pliocene – Quaternary sequence. We realize, however, that it is necessary to test this interpretation with isotopic compositions, the data for which are currently lacking. In conclusion, the major and trace element characteristics suggest that the magmas that produced the Eocene and Miocene – Quaternary volcanic sequences in the Lesser Caucasus were derived by different degrees of partial melting of a variously subductionenriched, subcontinental lithospheric mantle. The subduction signature in the melt evolution of these volcanic sequences appears to have diminished through time because of an increased asthenospheric component from the Eocene to the Quaternary. FC and/or AFC processes were also important during the evolution of these magmas.



565

Comparative petrogenesis of peri-Arabian Cenozoic volcanism In general, the geochemical characteristics of the Eocene volcanic sequence in the Lesser Caucasus are similar to those of the Eocene volcanic associations in the peri-Caspian and Ahar –Arasban and Central Iranian Volcanic Belts in Iran (Lotfi 1975; Lescuyer and Riou 1976), and in the Eastern Pontide and the southeastern Anatolian orogenic belts in eastern Turkey (Yigitbas and Yilmaz 1996; Keskin et al. 1998; Elmas and Yilmaz 2003). The late Miocene –Quaternary lavas in the Lesser Caucasus also show similar geochemical characteristics to those of the Sabalan – Sahand and Saray volcanoes in NW Iran (Didon and Germain 1976; Atapour 1994; Aftabi and Atapour 2000), the Plio-Pleistocene volcanic assemblages of the Nemrut and Tendu¨rek volcanoes and the Mus-Solhan volcanic field in the Turkish high plateau (Yilmaz et al. 1987; Pearce et al. 1990), and the middle to late volcanic units of the Erzurum –Kars Plateau (Figures 6, 7, 9 and 10; Keskin et al. 1998, 2006; Keskin 2003). Comparison of the source compositions presented on N-MORB- and chondritenormalized spider diagrams (Figures 9 and 10) indicates that the source regions of all these volcanic domains are similar in terms of their incompatible element signatures. The majority of the volcanic domains displayed on these diagrams (Erzurum – Kars Plateau, Turkish high plateau, and Azerbaijan-Lesser Caucasus) are enriched in the LILE and LREE – MREE relative to MORB, and show similar depletions in HREE. These features collectively suggest that the post-collisional Cenozoic magmas in this region were derived from small degrees of melting of subduction-metasomatized, depleted peridotite sources within the sub-continental lithospheric mantle (SCLM; Pearce et al. 1990; Keskin et al. 1998; Yilmaz et al. 1998). The subduction component was likely inherited from earlier subduction events in the region, for no active oceanic lithospheric subduction was in operation here during the late Cenozoic (after middle Miocene). Slab breakoff and/or delamination of all, or part of, the mantle lithosphere were likely processes, which triggered partial melting of the subduction-metasomatized continental lithospheric mantle, reminiscent of the late Cenozoic, post-collisional volcanism in the Maghrebian orogenic belt in NW Africa (Maury et al. 2000; Coulon et al. 2002), the Carpathian – Pannonian region (Nemcok et al. 1998; Seghedi et al. 2004), the Tibetan plateau (Turner et al. 1996; Williams et al. 2004; Zhao et al. 2009), and western Anatolia (Altunkaynak and Yılmaz 1998; Aldanmaz et al. 2000; Yilmaz et al. 2001; Ko¨pru¨basi and Aldanmaz 2004; Dilek and Altunkaynak 2007, 2009). Tectonic model for peri-Arabian Cenozoic volcanism The Cenozoic magmatism in the peri-Arabian region was directly associated, both spatially and temporally, with a series of collisional events and related mantle dynamics. The early Eocene was a time of regional contraction within the Tethyan realm in the eastern Mediterranean region, and the Gondwana-derived microcontinents were accreted along north-dipping subduction zones. The main collisions occurred in the northern and southern segments of the Tethyan realm, near the Eurasia and Arabia continental plates, respectively. The existence of three coeval Cretaceous arc systems, the Eurasian magmatic arc, the Eastern Pontide island arc, and the Baskil –S – S continental arc (from north to south, respectively), indicates the operation of at least two different, north-dipping (in present coordinate system) subduction zones within the Tethyan system by the Late Cretaceous (Figure 15). A north-dipping subduction zone within the Northern Neotethys was responsible for the evolution of the Eastern Pontide arc and its eastward continuation in the Lesser


566 Y. Dilek et al.



567

Caucasus (Armenia and Azerbaijan) and a backarc basin (Black Sea) during the Late Jurassic through Late Cretaceous (Figure 15(a); Okay et al. 1994; Shillington et al. 2008). The backarc basin behind (north of) this island arc was closing at a subduction zone dipping northwards beneath the Eurasian active margin by 65 Ma (Robinson et al. 1995; Yilmaz et al. 1997; Rolland et al. 2009a, 2009b). These subduction-accretion systems in the northern part of the Neotethys collapsed into the southern Eurasian margin by the early Eocene (Figure 15(c); Robinson et al. 1995). The arrival of the Eastern Tauride block and its eastward continuation in the Lesser Caucasus (South Armenian Block, SAB) at the Eastern Pontide trench resulted in the onset of an arc –continent collision along the IAESZ by the late Palaeocene– early Eocene (Figure 15(c); Dilek and Moores 1990; Boztug et al. 2006; Keskin et al. 2008). This collision followed the emplacement of the Cretaceous and older Neotethyan ophiolites onto the northern edge of the Eastern Tauride – South Armenian Block in the latest Cretaceous (Dilek and Sandvol 2009; Rolland et al. 2009b). Continued arc – continent collision and the underplating of the Eastern Tauride – South Armenian Block beneath the Eastern Pontide arc caused rapid uplift of the Kaçkar batholith and the associated plutons in the arc (Boztug et al. 2004, 2007) and widespread flysch deposition along and across the IAESZ (Dewey et al. 1986; Koçyigit et al. 1988; Tuÿsu¨z et al. 1995; Yilmaz et al. 1997). The partial subduction of the Eastern Tauride – South Armenian microcontinent led to slab breakoff and opening of an asthenospheric window beneath the arc mantle wedge and the collision zone (Figure 15(c)). This heat source triggered partial melting of the subductionmetasomatized lithospheric mantle and development of mid to late Eocene calc-alkaline to alkaline volcanism in a curvilinear belt from the Eastern Pontides to the Lesser Caucasus and the peri-Caspian Sea region in northern Iran. A slab breakoff origin for the Eocene volcanic rocks in the Eastern Pontides and in the northern edge of the Erzurum –Kars Plateau has been proposed by other researchers as well (Arslan et al. 1997; Sen et al. 1998; Keskin et al. 2006; Boztug et al. 2007). The coeval (Eocene) shoshonitic and calc-alkaline volcanic and plutonic sequences along the IAESZ farther west in north-central Turkey (Keskin et al. 2008) and in western Turkey (Altunkaynak and Dilek 2006; Dilek and Altunkaynak 2007) have also been interpreted as products of slab breakoff-induced postcollisional magmatism. The collision of the Arabian plate with the B – P and S – S continental blocks and their magmatic arcs occurred in the early Eocene (Yilmaz 1993; Ghasemi and Talbot 2006; Mazhari et al. 2009) and produced the me´lange and flysch deposits along the Bitlis – Zagros suture zone (Figure 15(b)). The occurrence of relatively undeformed OligoMiocene sedimentary units (i.e. Lower Red and Qom formations in the northern S– S zone) unconformably overlying the suture zone rocks suggests that much of the collisional deformation had ceased by the latest Eocene (Alavi 1994; Ghasemi and Talbot 2006). The collision of the Arabian plate with the fringing continental blocks to the N – NE was a diachronous event such that the accretion of the S –S continental block to the northeastern edge of Arabia along a dextral transpressional zone (Mohajjel and Fergusson 2000) may have preceded the head-on collision of the B – P continental block to the north– northwest by several millions of years. R Figure 15. Sequential geodynamic diagram depicting the tectonic evolution of the Cenozoic volcanism within a Tethyan realm in the peri-Arabian region. See text for discussion.


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This continental collision in the early Palaeogene led to slab breakoff and development of an asthenospheric window (Figure 15(c); Agard et al. 2005; Ghasemi and Talbot 2006; Dilek and Sandvol 2009), which in turn facilitated partial melting of the subductionmetasomatized lithospheric mantle beneath the newly accreted B – P and S– S continental blocks. This event resulted in the formation of shoshonitic magmatism in the hinterland of the collision zone in the upper plate. The middle– upper Eocene volcanic sequences of the Maden Complex in southeastern Anatolia (Yigitbas and Yilmaz 1996; Elmas and Yilmaz 2003), the 54 –38 Ma granitoid-gabbroic intrusions in the S –S continental block, and the shoshonitic volcanic-plutonic sequences in the Urumieh – Dokhtar magmatic belt and in the Ahar–Arasbaran and Central Iranian Volcanic Belts are the products of this Eocene magmatism in southeast Anatolia and the Zagros region in west – northwest Iran that was induced by slab breakoff. Following the detachment of the subducting oceanic lithosphere, the negative buoyancy of the underplated Arabian crust and the asthenospheric upwelling triggered rapid post-collisional uplift of the accreted microcontinents (B –P and S –S). This uplift in turn resulted in crustal exhumation, tectonic extension and core complex formation in the crystalline basement rocks in the region (Hassanzadeh et al. 2005; Moritz et al. 2006; Verdel et al. 2007; Dilek and Sandvol 2009). Continued subduction of the Tethyan seafloor beneath Eurasia farther to the north and steepening of the subducting slab associated with slab rollback produced southwardmigrating magmatism in the Eastern Pontide arc during the Eocene – Oligocene, while the subduction– accretion complex widened towards the south (Figure 15(d); Sengo¨r et al. 2003). As the Neotethyan lithosphere continued to subduct beneath the Pontide arc, the East Anatolian accretionary complex shortened and thickened within the closing basin. North – south contraction across the Neo-Tethyan realm between the converging Arabia composite plate and Eurasia caused vertical thickening of the East Anatolian subduction – accretion complex to an average crustal thickness of , 40 km by the late Oligocene –early Miocene (, 24 Ma; Sengo¨r et al. 2003). Southward retreat of the subducting Tethyan lithosphere may have peeled off the base of the subcontinental lithosphere, triggering partial lithospheric delamination beneath the southern margin of the Eastern Pontide arc and the northern part of the Turkish –Iranian high plateau (Figure 15(e)). Asthenospheric upwelling to replace the sinking lithospheric material resulted in remobilization and partial melting of the subduction-metasomatized mantle lithosphere (Pearce et al. 1990; Dilek and Sandvol 2009). This event produced the initial stages of calc-alkaline magmatism in the Erzurum – Kars Plateau by the middle Miocene (Keskin et al. 2006) and the early late Miocene magmatism in the western volcanic belt in Armenia (Karapetian and Adamian 1973; Badalyan 2000) and in the Lesser Caucasus of Azerbaijan (Imamverdiyev and Mamedov 1996; Imamverdiyev 2001a; this study). The arrival of the Arabian plate with the accreted microcontinents along its northern edge at the trench and the ensuing continent – trench collision by , 13 Ma resulted in widespread deformation and metamorphism in the collision zone as manifested in the formation of south-directed thrust sheets and nappes and south-vergent folding in the B – P (Figure 15(e); Michard et al. 1984; Yazgan 1984; Robertson et al. 2006) and S –S (Alavi 1994; Ghasemi and Talbot 2006). Oblique collision along the eastern edge of the Arabian promontory caused dextral transpression and related strike-slip deformation across the Zagros orogenic belt (Mohajjel and Fergusson 2000; Talebian and Jackson 2002). This continental collision slowed down and temporarily arrested the northward subduction beneath the East Anatolian subduction –accretionary complex. However, the continued sinking of the Neotethyan oceanic lithosphere in this subduction zone caused detachment of the subducting slab and development of an asthenospheric window (Figure 15(e);



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Molinaro et al. 2005; Lei and Zhao 2007; Omrani et al. 2008; Dilek and Sandvol 2009). Rising hot asthenosphere beneath the subduction –accretion complex resulted in thermal uplift and widespread partial melting both in the upwelling and convecting asthenosphere and in the overlying crust (Keskin 2003; Sengo¨r et al. 2003; Dilek and Sandvol 2009; Kheirkhah et al. 2009), which produced bimodal volcanism throughout the uplifted Turkish – Iranian plateau and the Lesser Caucasus. Extensive NW – SE-oriented transtensional and NNE – SSW-oriented extensional normal faulting in the Turkish – Iranian high plateau and the Lesser Caucasus (Koçyigit et al. 2001; Allen et al. 2004; Copley and Jackson 2006; Dhont and Chorowicz 2006) facilitated the rise and eruption of asthenosphere-derived alkaline olivine basalts with minimal continental contamination in the late Miocene –Pliocene (Figure 15(f)). Thus, orogen-parallel and crustal-scale strikeslip fault systems appear to have played a significant role in the development of postcollisional fissure eruptions and major eruptive centres (i.e. Ararat, Tendu¨rek and Sahand stratovolcanoes) in the peri-Arabian region. Widespread volcanism across the entire Turkish –Iranian high plateau (. 250 km wide), the Lesser Caucasus, and the peri-Arabian region throughout the late Cenozoic and until historic times indicates the presence of a significant heat source beneath the region, which produced extensive melting (Sengo¨r et al. 2003; Dilek and Sandvol 2009; Kheirkhah et al. 2009). The findings of the recent Eastern Turkey Seismic Experiment (ETSE) and tomographic models suggest an average continental crustal thickness , 40 –45 km, a lack of mantle lithosphere, a lack of earthquakes deeper than , 30 km, and very low Pn velocity zones indicating the presence of partially molten material beneath the region (Al-Lazki et al. 2003; Go¨k et al. 2003; Sandvol et al. 2003; Zor et al. 2003; Angus et al. 2006). These observations collectively suggest that the Turkish – Iranian high plateau is supported in part by hot asthenospheric mantle (Maggi and Priestley 2005), not by overthickened crust (Dewey et al. 1986) or subducted Arabian continental lithosphere (Rotstein and Kafka 1982). The Plio-Pleistocene and Quaternary volcanism in the peri-Arabian region becomes compositionally more alkaline in time and towards the south (Keskin 2003; Keskin et al. 2006; Kheirkhah et al. 2009; this study). However, all volcanic units still show a subduction zone fingerprint (high La/Nb ratios and LILE enrichment) despite the lack of a subducting Neotethyan oceanic lithosphere in the eastern Mediterranean region since , 13 Ma. These observations combined with trace element and available isotope characteristics of these volcanic sequences suggest that their magmas were derived from partial melting of subduction-metasomatized continental lithospheric mantle in the spinel lherzolite field (, 80 km) beneath the Turkish – Iranian plateau and the Lesser Caucasus (Kheirkhah et al. 2009; this study). The progressively more alkaline nature of the younger volcanic units indicates the stronger influence and an increased input of melts derived from the upwelling, enriched asthenospheric mantle through time. This geochemical boundary condition requires the existence of at least , 30 km of lithospheric mantle beneath the continental crust here (Figure 15(f)), rather than the lack of a conventional lithosphere as inferred from the findings of the ETSE (Al-Lazki et al. 2003; Go¨k et al. 2003; Zor et al. 2003). Recent S-wave receiver function analysis of the lithospheric structure of the Arabia– Eurasia collision zone in eastern Turkey (Angus et al. 2006) predicts the lithospheric thickness to be , 60– 80 km there, consistent with our geochemical inferences and modelling of the latest Cenozoic volcanism in the peri-Arabian region.


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Conclusions The Cenozoic plutonic and volcanic sequences in the Lesser Caucasus of Azerbaijan are part of the broader peri-Arabian post-collisional igneous province. This magmatism evolved in three main pulses in the (1) Eocene, (2a) late Miocene –Pliocene, and (2b) PlioQuaternary, and progressed through time from shoshonitic, calc-alkaline to more alkaline compositions towards the south. All volcanic sequences show similar trace element and REE patterns, with troughs in Nb, Ta, Hf, and Zr, strong enrichments in Rb, Ba, Th, La, and depletions in Ti, Yb, Y, relative to N-MORB, indicating a subduction-metasomatized lithospheric mantle as their melt source(s). Middle to upper Eocene magmatic units in the peri-Arabian region occur in the Eastern Pontide, Lesser Caucasus, and peri-Caspian areas to the north, and in the B –P and S– S continental blocks to the south. Two coeval but separate collisional events within the Tethyan realm in the early Eocene were responsible for slab detachment and asthenospheric heat input: (1) collision of the Eastern Tauride –South Armenian microcontinent with the Eastern Pontide arc at a north-dipping subduction zone in the Northern Neotethys, and (2) collision of the Arabian plate with the B –P and S– S continental blocks at another north-dipping subduction zone in the Southern Neotethys. Partial melting of the subcontinental lithospheric mantle and assimilation/FC processes produced evolved magmas that developed the post-collisional magmatic units in discrete, , EW-trending belts, straddling the early Eocene suture zones. The Miocene through Plio-Quaternary volcanic sequences occupy much of the Turkish –Iranian high plateau, Lesser Caucasus, peri-Caspian area, and Central Iranian Volcanic Belt, and occur as fissure eruptions and stratovolcanic centres mainly along NW – SE-trending transtensional, dextral strike-slip fault systems. Although these volcanic sequences display increased alkalinity in successively younger units, their high La/Nb ratios and LILE enrichments hint at a subduction zone influence in their mantle melt source. This inherited subduction fingerprint in the Plio-Quaternary volcanic units points to the existence of some mantle lithosphere beneath the modern Turkish –Iranian plateau. Partial melting of an upwelling asthenosphere in the hinterland of the Arabia – Eurasia collision zone contributed a greater enrichment in alkali content to the younger magmas, and it was triggered by the regional delamination of the mantle lithosphere. Acknowledgements This study was supported in part by research grants from the Havighurst Center at Miami University (USA) and Baku State University (Azerbaijan), and constitutes part of our ongoing investigation of the Cenozoic magmatism in the Lesser Caucasus, eastern Anatolia, and northern Iran. We thank Farahnaz Daliran (Germany), Manuel Pubellier (France), and Paul Robinson (Canada) for their constructive and insightful comments on the manuscript.

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