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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: [email protected]

International Geology Review

iFirst article, 2009, 1–43

Downloaded By: [Dilek, Yildirim] At: 18:11 12 November 2009

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.

Y. Dilek et al.2


International Geology Review 3


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.

R

Y. Dilek et al.4


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).



Y. Dilek et al.6


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).

R



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.

Y. Dilek et al.8


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



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

Y. Dilek et al.10


(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



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


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



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


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);



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


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



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


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



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

16

16

24

15

21

27

25

27

24

32

32

16

19

15

10

n.d

.n.d

.19

Y. Dilek et al.20


00

022

244

466

688

81010

101212

121414

1416A

B

CD

16

16

3540

4550

5560

6570

7580

SiO

2 (w

t. %

)

SiO

2 (w

t. %

)

SiO

2 (w

t. %

)

SiO

2 (w

t. %

)

PB

BB

AA

D

RT

PTA

TP

TP

P

TP

B

PT

TP

B

PT

I&B

I&B

TD

3540

4550

5560

6570

7580

Na2O + K2O (wt.%) Na2O + K2O (wt.%)

Na2O + K2O (wt.%) Na2O + K2O (wt. %)

PB

BB

A

AD

RT

BT

A

I&B

0246810121416

3535

4040

4545

5050

5555

6060

6565

7070

7575

8080

PB

BB

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BT

A

I&B

Nem

rut

Ara

rat

Ten

dure

k

Mus

Sup

han

AZ

ER

BA

IJA

NE

RZ

UR

UM

-KA

RS

PLA

TE

AU

IRA

NE

AS

TE

RN

AN

AT

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Late

Pha

se

Ear

ly P

hase

Eoc

ene

Lava

sG

abbr

o no

dule

s

Aze

rbai

jan

Cen

tral

Iran

Nor

ther

n Ir

an

Eas

tern

Iran

Kis

lako

y

Late

Sta

ge

Ear

ly S

tage

Mid

dle

Sta

ge

TA

TB

TP

B

PT

TP

P

P

TP

B

PT

TP

BB

AA

D

TD

RT

TA

TB

TA

TB

TD



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


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.



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

45

50

55

60

65

70

75

SiO

2 (w

t %)

10

12

14

16

18

20

22

Al 2O

3 (w

t %)

00.20.40.60.81.01.21.41.61.8

TiO

2 (w

t %)

0

2

4

6

8

10

Fe 2O

3 (w

t %)

0

2

4

6

8

10

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14

0 105 15

MgO (wt %)0 105 15

MgO (wt %)

CaO

(w

t %)

Eocene lavas

Gabbro nodules

020406080

100120140160180

Rb

(ppm

)

0

500

1000

1500

2000

2500S

r (p

pm)

20406080

100120140160180200

0

Ni (

ppm

)

00

0200400600800

10001200140016001800

Ba

(ppm

)

0102030405060708090

Nb

(ppm

)

Early phase

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


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

0.1

1

10 10

100 100

1000

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

k/N

-MO

RB

Roc

k/N

-MO

RB

Roc

k/N

-MO

RB

Roc

k/N

-MO

RB

0.1

1000

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.



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


1

10

100

1000


Roc

k/C

hond

rite

Roc

k/C

hond

rite

Roc

k/C

hond

rite

1

10

100

1000


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


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.



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


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.



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


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


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);



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


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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