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This article was downloaded by: [Dilek, Yildirim]On: 12 November 2009Access details: Access Details: [subscription number 916753623]Publisher Taylor & FrancisInforma Ltd Registered in England and Wales Registered Number: 1072954 Registered office: Mortimer House, 37-41 Mortimer Street, London W1T 3JH, UK
International Geology ReviewPublication details, including instructions for authors and subscription information:http://www.informaworld.com/smpp/title~content=t902953900
Geochemistry and tectonics of Cenozoic volcanism in the Lesser Caucasus(Azerbaijan) and the peri-Arabian region: collision-induced mantledynamics and its magmatic fingerprintYildirim Dilek a; Nazim Imamverdiyev b; Şafak Altunkaynak c
a Department of Geology, Miami University, Oxford, OH, USA b Department of Geology, Baku StateUniversity, Baku, Azerbaijan c Department of Geological Engineering, Istanbul Technical University,Maslak Istanbul, Turkey
First published on: 10 November 2009
To cite this Article Dilek, Yildirim, Imamverdiyev, Nazim and Altunkaynak, Şafak(2009) 'Geochemistry and tectonics ofCenozoic volcanism in the Lesser Caucasus (Azerbaijan) and the peri-Arabian region: collision-induced mantledynamics and its magmatic fingerprint', International Geology Review,, First published on: 10 November 2009 (iFirst)To link to this Article: DOI: 10.1080/00206810903360422URL: http://dx.doi.org/10.1080/00206810903360422
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Geochemistry and tectonics of Cenozoic volcanism in the LesserCaucasus (Azerbaijan) and the peri-Arabian region: collision-induced
mantle dynamics and its magmatic fingerprint
Yildirim Dileka*, Nazim Imamverdiyevb and Safak Altunkaynakc
aDepartment 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
(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 post-
collisional 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.
ISSN 0020-6814 print/ISSN 1938-2839 online
q 2009 Taylor & Francis
DOI: 10.1080/00206810903360422
http://www.informaworld.com
*Corresponding author. Email: dileky@muohio.edu
International Geology Review
iFirst article, 2009, 1–43
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Keywords: Lesser Caucasus (Azerbaijan); peri-Arabian region; Turkish–Iranian highplateau; 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 peri-
Arabian magmatism and the causes of heat supply that triggered melting are particularly
important questions for the geodynamic conditions and mechanisms that result in high-
magmatic 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 within-
plate 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 calc-
alkaline 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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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
Figure 2. Tectonic map of the eastern Mediterranean–Persian Gulf region, showing the main plateboundaries, 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 volcanicunits (i.e. Maden complex, Kislakoy volcanics) discussed in the text are also shown. CACC, CentralAnatolian 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–Akerasuture zone.
Figure 1. Simplified tectonic map of the eastern Mediterranean–Persian Gulf region, showing theactive plate boundaries, plate convergence vectors (in green) with respect to fixed Eurasia, and post-collisional 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–Karsplateau; HT, Hellenic trench; IAESZ, Izmir–Ankara–Erzincan suture zone; ITSZ, Inner-Tauridesuture zone; MTJ, Maras triple junction; NAF, North Anatolian fault; NEAF, Northeast Anatolianfault; PSF, Pampak–Sevan fault; TF, Tabriz fault; TGF, Tuzgolu 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–Puturge (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 melanges, 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
region (Figures 2 and 3; Ozgul and Tursucu 1984). The Tauride microcontinent consists of
late Proterozoic–Palaeozoic and Mesozoic carbonate, siliciclastic, and volcanic rocks
(Ozgul 1976; Demirtasli et al. 1984) and represents a ribbon continent rifted off from the
northwestern edge of Gondwana (Robertson and Dixon 1984; Sengor 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 (Perincek
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 melange
(Figure 4), consisting of blocks of serpentinites, metamorphic rocks, and pelagic
limestones in a fine-grained, phyllitic matrix (Ozgul and Tursucu 1984). Both the Ovacik
melange 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; Perincek 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, melanges, and Upper Cretaceous and younger volcanic and volcaniclastic rocks. The
Puturge 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
(Goncuoglu and Turhan 1984; Helvaci and Griffin 1984), and Palaeozoic to late Mesozoic
Figure 3. Modern topography of the Arabia–Eurasia collision zone and the Turkish–Iranian highplateau, 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 andvelocity in mm/year) with respect to fixed Eurasia based on the GPS data of McClusky et al. (2000).
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granitoids. Oberhansli et al. (2008) reported a regionally distributed high-pressure/low-
temperature 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 Puturge massifs and the overlying volcanic
and ophiolitic rocks are structurally underlain in the south by an Upper Cretaceous–early
Tertiary melange, 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,
Figure 4. Simplified geological map of Eastern Anatolia and the Arabian foreland, showing thedistribution of major tectonic units in the region and the post-collisional volcanic rocks in theTurkish high plateau. Munzur Platform constitutes the eastern extension of the platform carbonatesand basement rocks of the Eastern Tauride ribbon-continent. B–P massif is a rifted off fragment ofthe Arabian plate, analogous to the Eastern Tauride block. The Turkish high plateau is covered byMiocene–Quaternary volcanic rocks; its basement is composed of Tethyan ophiolites and ophioliticmelanges, flysch and molasses deposits, and platform carbonates of the Tauride block. Kackarbatholith and the Jurassic–Upper Cretaceous sandstone, volcanic tuff, and limestone in northernTurkey constitute the Eastern Pontide Arc. EAF, East Anatolian fault; EAFZ, East Anatolian faultzone; EKP, Erzurum–Kars plateau; IAESZ, Izmir–Ankara–Erzincan suture zone; KOTJ, Karliovatriple junction; NAFZ, North Anatolian fault zone; NEAF, Northeast Anatolian fault.
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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; Sengor 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 (Perincek 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 strike-
slip (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
International Geology Review 11
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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 Gumushane area (Y. Eyuboglu, 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 Kislakoy 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 Kackar 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 Kackar 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 Kosedag 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 Corum area along the IAESZ, north of the Central Anatolian crystalline complex
(CACC; Keskin et al. 2008).
Y. Dilek et al.12
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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 Yuksekova 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, Suphan, Tendurek, 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
International Geology Review 13
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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
Figure 5. Geological map of the Palaeogene–Neogene and Quaternary magmatic units andvolcanic centres in the central Lesser Caucasus, Azerbaijan.
Y. Dilek et al.14
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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);
International Geology Review 15
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Tab
le1
.M
ajo
ran
dtr
ace
elem
ent
com
po
siti
on
so
fre
pre
sen
tati
ve
Eo
cen
ev
olc
anic
un
its,
Les
ser
Cau
casu
s(A
zerb
aija
n).
TA
6-1
*la
va
TA
9-1
*la
va
TA
10
-3*
sill
TA
19
-1*
sill
TA
39
-1*
sill
TA
39
-3*
sill
TA
64
-1*
clas
tT
A6
6-1
*cl
ast
TA
70
-1*
sill
TA
77
-1*
sill
TA
82
-1*
sill
TA
87
-1*
lav
a
Maj
or
ox
ides
(wt%
)S
iO2
52
.30
55
.75
51
.09
50
.44
47
.17
47
.18
52
.50
51
.46
51
.77
48
.82
51
.79
50
.85
TiO
20
.86
0.7
41
.14
1.1
31
.16
1.2
01
.09
0.8
31
.14
0.9
81
.25
1.1
6A
l 2O
31
5.9
51
9.5
51
6.2
81
5.7
01
6.3
51
6.9
51
4.1
71
8.4
31
6.5
11
7.6
11
7.0
91
6.9
6F
e 2O
38
.43
5.6
98
.32
9.8
29
.07
9.4
47
.34
7.1
79
.11
8.9
38
.99
8.1
2M
nO
0.1
60
.12
0.1
70
.16
0.2
60
.19
0.2
10
.16
0.1
70
.17
0.1
50
.16
Mg
O4
.00
2.0
14
.92
5.9
25
.18
4.5
97
.62
4.9
85
.17
4.8
74
.82
5.0
8C
aO6
.61
7.0
59
.62
8.4
57
.25
5.8
61
0.3
47
.14
9.0
38
.12
8.3
99
.68
Na 2
O3
.38
3.9
72
.97
2.6
73
.49
3.7
52
.30
2.7
52
.94
3.3
93
.17
3.0
1K
2O
5.3
23
.99
3.0
03
.41
2.8
23
.84
2.5
74
.80
2.9
33
.33
3.5
03
.02
P2
O5
0.5
20
.57
0.4
80
.46
0.4
10
.43
0.3
60
.67
0.4
90
.55
0.5
00
.49
LO
I2
.39
0.5
31
.36
1.3
75
.59
5.8
81
.34
1.7
11
.07
2.8
90
.99
1.5
1T
ota
l9
9.6
71
00
.00
99
.46
99
.63
99
.26
99
.46
99
.91
10
0.1
71
00
.38
99
.76
99
.63
99
.46
Tra
ceel
emen
ts(p
pm
)R
b1
70
11
37
18
37
19
57
61
22
77
69
91
70
Sr
89
07
61
65
66
02
92
51
00
85
28
11
52
68
69
88
59
96
47
Ba
92
38
97
66
96
37
10
09
10
63
48
61
16
77
18
87
37
22
65
2Z
r1
22
16
91
32
12
07
06
48
91
00
14
18
91
74
12
2N
b1
61
81
71
61
51
31
21
61
81
32
31
5T
a0
.97
1.0
01
.09
0.9
10
.98
0.8
30
.69
0.9
21
.05
0.8
01
.38
0.8
8H
f3
.17
4.3
83
.49
3.1
81
.85
1.7
02
.43
2.4
03
.68
2.3
74
.48
3.1
7T
h9
.02
11
.35
7.7
06
.88
3.7
93
.09
3.7
28
.13
8.1
56
.34
11
.14
6.2
0N
i3
71
42
02
11
01
69
37
22
31
81
82
4L
a3
43
63
22
92
31
92
03
43
42
83
72
9C
e6
16
75
95
34
33
63
95
96
35
16
75
4S
m5
.78
5.7
85
.80
5.3
14
.72
3.8
74
.51
4.9
86
.17
5.1
06
.21
5.5
0E
u1
.69
1.6
11
.68
1.5
11
.56
1.2
01
.44
1.5
71
.77
1.6
31
.68
1.6
2T
b0
.77
0.7
50
.83
0.7
70
.72
0.6
10
.72
0.6
50
.90
0.6
90
.89
0.7
9Y
b2
.08
2.2
52
.32
2.1
11
.93
1.6
31
.91
1.7
42
.46
1.8
32
.55
2.1
9L
u0
.31
0.3
40
.34
0.3
10
.27
0.2
30
.28
0.2
60
.35
0.2
70
.38
0.3
2Y
23
24
26
24
23
19
23
20
27
21
28
25
No
tes:
Sam
ple
sw
ith
( *)
sym
bo
lar
efr
om
Vin
centet
al.
(20
05)
and
n.d
.,N
ot
det
ecte
d.
Y. Dilek et al.16
Downloaded By: [Dilek, Yildirim] At: 18:11 12 November 2009
Tab
le2.
Maj
or
and
trac
eel
emen
tco
mposi
tions
of
repre
senta
tive
sam
ple
sof
the
earl
yphas
e(l
ate
Mio
cene
–ea
rly
Pli
oce
ne)
volc
anic
unit
san
dgab
bro
no
du
les,
Les
ser
Cau
casu
s(A
zerb
aija
n).
40
lav
a1
5la
va
10
0la
va
19
0la
va
19
3la
va
19
4la
va
8la
va
96
lav
a1
06
lav
a7
4la
va
20
0la
va
95
3la
va
97
3la
va
19
0/G
Gab
bro
no
du
le
19
4/A
Gab
bro
no
du
le
19
4/B
Gab
bro
no
du
le
Maj
or
ox
ides
(wt%
)S
iO2
61
.87
62
.61
62
.10
61
.75
62
.04
62
.84
63
.80
70
.62
65
.01
64
.97
64
.51
70
.40
74
.21
49
.80
45
.94
51
.42
TiO
20
.59
0.5
80
.60
0.8
10
.79
0.7
50
.49
0.2
70
.60
0.5
20
.55
0.0
10
.32
1.1
51
.58
1.0
1A
l 2O
31
5.7
01
6.9
01
6.6
01
4.8
11
6.2
51
7.1
51
5.4
11
5.7
71
7.0
31
6.4
11
5.9
61
5.1
01
5.6
78
.46
13
.09
17
.82
Fe 2
O3
3.4
73
.91
3.2
83
.91
4.8
14
.94
2.5
01
.69
3.3
83
.59
3.5
51
.36
1.0
05
.62
9.6
66
.03
FeO
1.2
91
.01
1.2
92
.46
0.7
20
.43
0.9
40
.43
0.7
30
.28
1.0
11
.48
0.4
33
.33
1.7
42
.17
Mn
O0
.06
0.0
40
.09
0.1
00
.09
0.0
90
.06
0.0
40
.03
0.0
90
.08
0.0
90
.03
0.1
80
.19
0.1
6M
gO
1.8
51
.95
1.9
03
.18
2.0
21
.86
1.7
70
.05
1.4
31
.31
1.1
31
.14
1.0
51
2.4
38
.13
5.8
7C
aO4
.85
4.2
44
.32
6.1
35
.04
5.2
55
.34
1.3
23
.97
3.1
93
.30
0.9
70
.54
13
.74
13
.47
8.8
3N
a 2O
4.1
94
.07
4.0
83
.37
3.1
83
.30
3.9
34
.57
4.2
74
.05
4.0
02
.94
2.0
62
.00
2.9
93
.38
K2
O3
.54
2.9
53
.08
2.3
72
.57
1.8
72
.73
4.1
43
.47
2.5
53
.47
3.2
53
.14
0.9
20
.89
1.2
7P
2O
50
.41
0.2
80
.30
0.2
80
.40
0.3
50
.38
0.0
60
.33
0.2
30
.22
0.1
60
.07
0.1
31
.70
0.5
2L
OI
0.8
10
.54
0.4
60
.13
0.3
60
.38
1.9
60
.27
0.4
70
.96
0.8
52
.88
1.5
60
.39
0.4
90
.46
To
tal
98
.63
99
.08
98
.10
98
.39
8.2
79
9.2
19
9.3
19
9.2
31
00
.72
98
.15
98
.43
99
.78
10
0.0
89
8.1
59
9.8
79
8.9
4T
race
elem
ents
(pp
m)
Rb
90
63
55
45
66
79
56
97
86
72
90
12
81
18
11
11
22
Li
7.2
19
17
16
15
13
14
13
12
13
22
n.d
.n
.d.
23
22
14
Sr
89
05
90
60
08
50
94
09
35
68
04
20
93
07
90
71
03
00
15
04
60
11
00
13
00
Ba
12
40
79
07
30
90
06
50
11
20
92
08
30
74
06
60
10
70
35
05
00
38
04
40
65
0Z
n6
56
55
47
07
05
26
37
05
75
94
61
00
30
01
00
10
01
10
Cu
20
37
83
41
25
37
20
13
22
26
31
20
85
41
41
71
Zr
15
01
70
15
0n
.d.
15
0n
.d.
18
02
40
17
01
50
20
01
50
20
01
10
85
78
Nb
10
10
10
n.d
.1
0n
.d.
10
17
14
14
14
n.d
.n
.d.
10
19
10
Ta
0.3
1n
.d.
0.9
4n
.d.
0.8
2n
.d.
0.8
21
.20
1.4
01
.00
n.d
.n
.d.
n.d
.0
.40
0.6
80
.47
Hf
2.6
0n
.d.
3.3
0n
.d.
4.0
0n
.d.
4.0
06
.00
4.7
04
.00
n.d
.n
.d.
n.d
.2
.80
3.1
02
.10
U2
.70
n.d
.4
.50
n.d
.4
.70
n.d
.4
.70
5.2
05
.40
3.4
0n
.d.
4.7
07
.00
2.0
05
.30
4.0
0T
h1
1.0
0n
.d.
10
.00
n.d
.1
1.0
0n
.d.
11
.00
14
.00
18
.00
15
.00
n.d
.1
0.0
02
1.0
04
.00
5.1
03
.80
Cr
12
03
10
18
0n
.d.
18
01
40
18
0n
.d.
18
01
00
n.d
.3
03
07
10
n.d
.n
.d.
V1
70
80
65
90
10
01
00
70
40
10
01
00
12
08
51
51
70
21
02
30
International Geology Review 17
Downloaded By: [Dilek, Yildirim] At: 18:11 12 November 2009
Tab
le2
–continued
40
lav
a1
5la
va
10
0la
va
19
0la
va
19
3la
va
19
4la
va
8la
va
96
lav
a1
06
lav
a7
4la
va
20
0la
va
95
3la
va
97
3la
va
19
0/G
Gab
bro
no
du
le
19
4/A
Gab
bro
no
du
le
19
4/B
Gab
bro
no
du
le
Ni
24
30
30
40
22
32
22
15
32
25
20
25
10
70
45
56
Co
20
35
15
35
30
25
30
93
01
53
05
08
51
38
24
Sc
27
20
15
10
81
07
38
10
10
10
38
44
22
8L
a4
54
33
62
3n
.d.
47
37
47
47
38
53
n.d
.n
.d.
23
68
43
Ce
88
77
76
57
n.d
.9
17
37
88
77
47
9n
.d.
n.d
.4
61
40
74
Sm
4.2
03
.90
4.2
07
.50
n.d
.5
.10
3.6
05
.00
3.6
04
.40
6.3
0n
.d.
n.d
.7
.90
14
.00
6.7
0E
u1
.20
1.2
01
.00
1.6
0n
.d.
1.6
01
.00
0.7
91
.10
0.9
51
.20
n.d
.n
.d.
2.0
03
.20
1.8
0T
b0
.67
0.5
60
.58
1.1
0n
.d.
0.9
00
.43
0.5
70
.44
0.4
20
.99
n.d
.n
.d.
1.6
01
.50
0.8
1Y
b1
.20
1.4
01
.50
3.6
0n
.d.
1.8
01
.30
1.4
01
.30
1.3
01
.70
n.d
.n
.d.
3.7
03
.00
1.5
0L
u0
.19
0.2
00
.20
0.6
9n
.d.
0.2
30
.18
0.1
80
.17
0.1
70
.21
n.d
.n
.d.
0.6
40
.52
0.2
5Y
11
16
14
29
n.d
.n
.d.
16
16
10
97
n.d
.n
.d.
25
27
21
Y. Dilek et al.18
Downloaded By: [Dilek, Yildirim] At: 18:11 12 November 2009
Tab
le3.
Maj
or
and
trac
eel
emen
tco
mposi
tions
of
repre
senta
tive
sam
ple
sof
the
late
phas
e(l
ate
Pli
oce
ne
–Q
uat
ernar
y)
volc
anic
unit
san
dgab
bro
nodu
les,
Les
ser
Cau
casu
s(A
zerb
aija
n).
10
5la
va
129
lava
132
lava
13
4la
va
21
lava
57
lava
20
8la
va
19/P
lava
53
lava
87
lava
109
lava
36/P
lava
12
0la
va
16
7la
va
174
Lav
a180
Lav
a13
Lav
a25
Lav
a33
Lav
a1
43
Lav
a1
60
Lav
a1
85
Lav
a73/P
Lav
a12
Lav
a6-1
74
Lav
a
13/3
-VG
abbro
nodule
Maj
or
oxid
es(w
t%)
SiO
251.2
348.3
548.8
84
8.0
551.8
449.4
252.9
750.5
053.3
253.0
554.9
254.9
055.6
754.3
154.0
155.2
157.6
658.5
259.8
557.0
85
9.2
857.8
567.8
073.9
975.5
151
.41
TiO
21.3
91.2
01.5
71.4
51.3
61.4
41.3
01.1
80.9
71.1
41.1
40.9
21.0
81.1
81.5
01.5
20.7
90.8
20.8
01.2
41.2
40
.75
0.4
80.0
10.0
11.4
5A
l 2O
316.4
915.7
715.8
61
5.5
316.6
416.2
716.4
617.7
017.3
917.4
616.3
817.6
017.1
316.8
217.4
916.9
916.4
116.2
316.6
717.2
51
6.5
517.7
015.7
013.4
813.7
918
.73
Fe 2
O3
7.7
46.3
85.6
13.5
56.1
17.1
67.0
47.0
06.1
15.6
64.5
47.0
06.5
95.0
25.7
93.6
94.0
94.8
04.8
84.6
24.9
53
.79
4.0
01.2
00.5
55.9
7F
eO0.8
62.1
62.7
34.4
61.0
10.7
20.3
00.8
00.5
71.6
52.5
90.3
00.4
32.1
72.4
63.9
01.8
70.8
70.5
03.0
91.3
01
.88
3.0
01.7
80.7
11.5
9M
nO
0.1
30.1
50.1
40.1
30.1
10.1
20.1
20.1
50.1
00.1
30.1
00.1
30.1
20.1
20.1
20.1
20.0
50.0
90.1
10.1
10.1
00
.13
0.0
50.0
10.0
10.1
2M
gO
6.0
46.7
46.2
96.8
14.4
25.2
73.6
55.3
03.8
14.1
23.7
63.9
04.6
63.8
43.3
72.5
03.1
83.2
32.6
72.2
92.7
92
.77
1.1
00.1
40.3
64.8
9C
aO8.3
39.8
09.0
99.1
98.5
89.1
07.0
09.2
07.1
76.7
16.8
87.1
06.2
46.6
66.8
05.9
66.2
56.2
45.6
16.0
95.8
26
.12
2.2
00.5
31.9
09.5
8N
a 2O
4.2
23.6
14.0
04.1
84.1
43.2
24.3
94.5
05.0
34.2
73.7
04.6
04.2
24.7
84.5
35.0
43.8
54.0
04.3
84.5
34.6
54
.53
5.5
03.2
72.9
24.1
1K
2O
1.4
21.9
61.9
21.7
32.9
22.4
83.1
62.9
02.8
02.7
72.1
73.0
02.6
02.9
63.2
53.1
13.0
12.8
03.1
12.8
73.4
62
.89
4.0
04.8
73.9
61.6
1P
2O
50.6
51.0
31.1
81.1
31.3
11.0
40.9
30.8
90.8
20.8
30.9
40.7
80.5
80.7
50.9
40.9
10.5
70.6
80.7
90.6
80.7
60
.44
0.3
50.0
10.0
10.4
0L
OI
0.7
01.5
00.9
31.7
90.6
11.9
01.1
01.0
00.1
40.3
50.8
51.0
00.4
10.1
90.4
40.0
20.6
40.4
00.3
50.2
70.2
01
.15
1.0
00.3
80.5
40.3
9T
ota
l99.2
098.6
598.2
09
8.0
099.0
598.1
498.4
2100.1
298.2
398.1
498.4
7100.2
399.7
398.8
0100.7
098.9
798.3
798.6
899.7
2100.1
2101.1
0100.0
0100.1
899.6
7100.2
7100.2
5T
race
elem
ents
(ppm
)R
b16
30
32
34
37
27
53
43
34
36
39
42
55
37
43
43
55
49
66
40
56
48
70
160
180
33
Li
10
98
98
913
12
13
12
13
13
17
14
13
15
10
12
16
14
17
15
20
67
70
11
Sr
910
1310
1360
1490
2400
2600
1900
1780
1420
1615
1130
1433
730
1700
1700
1190
1360
1275
1615
1647
1360
790
1356
150
100
1400
Ba
600
1040
1020
990
1300
1170
1170
1267
980
1000
1240
1054
800
900
900
840
830
1060
900
900
1016
930
1100
100
100
500
Zn
49
64
100
56
100
120
110
78
100
95
150
66
80
100
140
100
91
70
80
100
100
100
55
100
30
95
Cu
75
71
70
67
90
90
66
58
50
21
45
35
46
46
21
32
63
37
100
50
28
35
41
30
2112
Zr
178
229
259
244
200
338
250
240
222
210
252
235
222
250
244
222
190
180
220
207
200
160
303
100
80
260
Nb
35
35
28
35
35
42
23
20
42
21
28
25
19
23
35
21
18
13
18
21
23
15
33
15
10
28
Ta
0.9
20.9
20.9
20.9
61.2
01.7
01.5
01.8
00.8
00.9
9n.d
.1.2
01.2
01.4
01.3
01.3
00.8
10.8
71.0
00.9
81.4
00.8
81.4
3n
.d.
n.d
.n
.d.
Hf
4.6
04.7
05.2
05.1
04.5
04.6
05.2
04.7
54.0
04.7
0n.d
.4.9
04.4
04.8
05.0
05.1
04.8
04.5
05.3
04.7
04.7
04.3
06.6
0n
.d.
n.d
.n
.d.
U3.0
03.0
03.0
03.0
05.2
07.4
03.0
03.0
04.0
05.3
0n.d
.4.6
04.0
02.8
03.8
04.0
06.3
06.5
08.8
05.6
09.5
09.7
03.2
09
.30
12.0
0n.d
.T
h2.6
03.2
02.6
04.9
04.0
03.8
08.1
05.7
06.0
04.0
0n.d
.7.2
05.6
06.4
06.5
07.5
03.6
06.3
04.0
04.0
04.0
04.0
012.2
025.0
031.0
0n.d
.C
r310
412
280
450
170
220
n.d
.28
141
200
261
27
160
n.d
.n.d
.n.d
.160
188
100
n.d
.n
.d.
n.d
.140
30
n.d
.3
22
V1
65
170
210
260
140
220
150
96
200
200
150
129
130
240
150
170
80
130
100
140
140
110
70
n.d
.20
190
Ni
100
93
110
10
043
64
45
25
38
48
25
25
30
44
19
15
50
54
50
33
29
31
13
.520
31
51
Co
30
26
60
24
26
50
45
21
18
50
27
21
14
21
18
17
45
16
20
40
19
13
11
53
24
Sc
15
18
21
26
20
20
20
14.5
10
20
10
4.9
20
14
16
20
20
14
11
18
10
10
6.7
n.d
.n.d
.35
La
40
65
63
62
76
77
77
73.5
59
66
69
72
52
69
80
69
60
60
70
59
67
48
72
n.d
.n.d
.29
Ce
81
130
130
120
150
160
160
135
120
130
130
130
98
120
160
140
120
120
120
120
140
88
115
n.d
.n.d
.72
Sm
5.3
09.5
09.8
09.1
010.0
011.0
09.5
08.4
06.3
07.4
07.4
08.1
05.9
07.4
09.8
08.0
05.7
05.3
05.8
07.2
08.6
05.7
06.0
0n
.d.
n.d
.5.9
0
International Geology Review 19
Downloaded By: [Dilek, Yildirim] At: 18:11 12 November 2009
Tab
le3
–continued
10
5la
va
129
lava
132
lava
13
4la
va
21
lava
57
lava
20
8la
va
19/P
lava
53
lava
87
lava
109
lava
36/P
lava
12
0la
va
16
7la
va
174
Lav
a180
Lav
a13
Lav
a25
Lav
a33
Lav
a1
43
Lav
a1
60
Lav
a1
85
Lav
a73/P
Lav
a12
Lav
a6-1
74
Lav
a
13/3
-VG
abbro
nodule
Eu
1.7
02.5
02.5
02.4
02.5
02.8
02.5
02.1
51.6
01.8
02.0
01.9
51.7
02.2
02.7
02.3
01.6
01.7
01.7
02.0
02.0
01.4
01.5
0n
.d.
n.d
.1.8
0T
b0.8
81.5
01.3
01.1
01.0
01.3
01.3
01.3
51.0
01.4
01.1
01.0
50.9
01.1
00.9
51.4
01.1
00.9
40.8
51.8
01.2
00.5
91.1
2n
.d.
n.d
.1.7
0Y
b2.4
02.7
02.4
02.2
01.8
01.9
02.3
02.3
51.8
02.1
02.0
02.3
52.0
02.2
02.0
02.2
01.8
01.9
02.0
02.2
02.1
01.3
02.1
0n
.d.
n.d
.2.3
0L
u0.4
20.3
90.3
30.3
10.2
20.3
40.3
40.2
70.2
50.2
80.2
20.3
30.3
90.3
10.2
70.3
50.3
10.3
00.2
60.2
50.2
40.2
40.2
5n
.d.
n.d
.0.3
7Y
31
30
34
29
16
23
23
15
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022
244
466
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81010
101212
121414
1416A
B
CD
16
16
3540
4550
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6570
7580
SiO
2 (w
t. %
)
SiO
2 (w
t. %
)
SiO
2 (w
t. %
)
SiO
2 (w
t. %
)
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BB
AA
D
RT
PTA
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TP
P
TP
B
PT
TP
B
PT
I&B
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TD
3540
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7580
Na2O + K2O (wt.%) Na2O + K2O (wt.%)
Na2O + K2O (wt.%) Na2O + K2O (wt. %)
PB
BB
A
AD
RT
BT
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I&B
0246810121416
3535
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dure
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se
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dule
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lako
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International Geology Review 21
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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
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 latephases 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 EasternAnatolia (Pearce et al. 1990; Yilmaz et al. 1998); early, middle and late stages of volcanism andKislakoy 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; Aftabiand Atapour 2000).
R
Y. Dilek et al.22
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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)).
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 datasources.
International Geology Review 23
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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
40
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2 (w
t %)
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0102030405060708090
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)
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Late phase
Figure 8. Selected major and trace element vs. MgO variation diagrams for the Cenozoic volcanicsequences in the Lesser Caucasus of Azerbaijan. See Figure 6 for the data sources.
Y. Dilek et al.24
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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,
0.1 0.1
10 10
100 100
1000 1000
1 1
1
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10 10
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Sr K Rb Ba Th Ta Nb La Ce P Zr Sm Tb Ti Y Yb Sr K Rb Ba Th Ta Nb La Ce P Zr Sm Tb Ti Y Yb
Sr K Rb Ba Th Ta Nb La Ce P Zr Sm Tb Ti Y YbSr K Rb Ba Th Ta Nb La Ce P Zr Sm Tb Ti Y Yb
Roc
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-MO
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-MO
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Intermediate-mafic lavas
Felsic Lavas
EARLY PHASE Eocene LavasLate Phase
Middle StageLate Stage Early Stage
Kislakoy (Eocene)
ERZURUM-KARS PLATEAU EASTERN ANATOLIA
NemrutTendurek
AraratSuphan
AZERBAIJANAZERBAIJAN
Gabbro nodulesKislakoy (Eocene)
Figure 9. N-MORB-normalized multi-element patterns for the Cenozoic volcanic sequences in theLesser Caucasus of Azerbaijan (top two panels) and in the Erzurum–Kars plateau and EasternAnatolia (lower two panels). N-MORB normalizing values are from Sun and McDonough (1989).See Figure 6 for the data sources.
International Geology Review 25
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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; Koprubasi 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
1
10
100
1000A B
C D
La Ce Sm Eu Tb Yb Lu
1
10
100
1000
La Ce Sm Eu Tb Yb Lu
1
10
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La Ce Sm Eu Tb Yb Lu
Roc
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hond
rite
Roc
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hond
rite
Roc
k/C
hond
rite
1
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1000
La Ce Sm Eu Tb Yb Lu
Late PhaseEarly PhaseEocene Lavas
Intermediate-mafic lavas
Felsic Lavas
Nemrut
Ararat
TendurekMu SuphanKislakoy (Eocene)
Late Stage Early StageMiddle Stage
Roc
k/C
hond
rite
ERZURUM-KARS PLATEAU EASTERN ANATOLIA
AZERBAIJAN AZERBAIJAN
Gabbro nodules
Figure 10. Chondrite-normalized REE element patterns for the Cenozoic volcanic sequences in theLesser Caucasus of Azerbaijan (top two panels) and in the Erzurum–Kars plateau and EasternAnatolia (lower two panels). Chondrite normalizing values are from Boynton (1984). See Figure 6for the data sources.
Y. Dilek et al.26
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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 collision-
related 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
0.01
0.1
1
0.01 0.1 1 10
Early phase Late phaseEocene Lavas
Th/
Yb
Ta/Yb
AverageN-type MORB
UC
MM:mantle metasomatism trendSZE:subduction zone enrichmentUC:average upper crust
MM
AFC
10
GabbroNodules
SZE
Figure 11. Th/Yb vs. Ta/Yb diagram (after Pearce 1982) for mafic to intermediate lavas from theEocene and Miocene–Quaternary volcanic sequences in the Lesser Caucasus of Azerbaijan.
International Geology Review 27
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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
Intraplate
Collision-related
Subduction-related
10
0.110 100 1000
Zr (ppm)
Nb/
Zr
1
Eocene Lavas
Late phase(mafic)Late phase (felsic)
Early phase
Figure 12. Nb/Zr(n) vs. Zr diagram (Thieblemont and Tegyey 1994) for the Cenozoic volcanicsequences in the Lesser Caucasus of Azerbaijan. (n), N-MORB normalized values (Sun andMcDonough 1989).
1
10
100
0.1 1 10
La/Nb
Ba/
Nb
Early phase
Gabbro nodules
Late phase
Eocene lavas
MORB
PMOIB
CC
Figure 13. Ba/Nb vs. La/Nb diagram for the Cenozoic volcanic sequences in the Lesser Caucasusof Azerbaijan. PM, primary mantle; OIB, ocean island basalt, MORB values are from Sun andMcDonough (1989); CC, Continental crust from Rudnick and Gao (2003).
Y. Dilek et al.28
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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 subduction-
enriched, 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.
0
10
20
30
40
50
60
500 100 150
La
La/Y
bEarly Phase
Late Phase
Eocene LavasIn
crea
sing
parti
al m
eltin
gFigure 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.
International Geology Review 29
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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 Tendurek 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 chondrite-
normalized 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; Koprubasi 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
Y. Dilek et al.30
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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 Kackar 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; Kocyigit et al. 1988; Tuysuz 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 subduction-
metasomatized 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 post-
collisional 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 melange and flysch deposits along the Bitlis–
Zagros suture zone (Figure 15(b)). The occurrence of relatively undeformed Oligo-
Miocene 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.
Figure 15. Sequential geodynamic diagram depicting the tectonic evolution of the Cenozoicvolcanism within a Tethyan realm in the peri-Arabian region. See text for discussion.
R
Y. Dilek et al.32
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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 subduction-
metasomatized 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 southward-
migrating magmatism in the Eastern Pontide arc during the Eocene–Oligocene, while the
subduction–accretion complex widened towards the south (Figure 15(d); Sengor 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; Sengor 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);
International Geology Review 33
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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; Sengor 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 (Kocyigit 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 strike-
slip fault systems appear to have played a significant role in the development of post-
collisional fissure eruptions and major eruptive centres (i.e. Ararat, Tendurek 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 (Sengor 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; Gok 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; Gok 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.
Y. Dilek et al.34
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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) Plio-
Quaternary, 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 ofthe Cenozoic magmatism in the Lesser Caucasus, eastern Anatolia, and northern Iran. We thankFarahnaz Daliran (Germany), Manuel Pubellier (France), and Paul Robinson (Canada) for theirconstructive and insightful comments on the manuscript.
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