ARTICLE

Age and tectonic setting of the Bavanat Cu–Zn–Ag Besshi-typevolcanogenic massive sulfide deposit, southern Iran

Fardin Mousivand & Ebrahim Rastad &

Sebastien Meffre & Jan M. Peter &

Mohammad Mohajjel & Khin Zaw &

Mohammad Hashem Emami

Received: 7 November 2011 /Accepted: 26 January 2012# Springer-Verlag 2012

Abstract The Bavanat Cu–Zn–Ag Besshi-type volcanogenicmassive sulfide (VMS) deposit occurs within the Surianvolcano-sedimentary complex in the Sanandaj–Sirjan zone(SSZ) of southern Iran. The Surian complex is comprised ofpelite, sandstone, calcareous shale, basalt, gabbro sills, andthin-bedded limestone. Mineralization occurs as stratiformsheet-like and tabular orebodies hosted mainly by greenschistmetamorphosed feldspathic and quartz feldspathic sandstone,basalt, and pelites. The basalts of the Surian complex show

predominantly tholeiitic to transitional affinities, with a fewsamples that are alkalic in composition. Primitive mantle-normalized trace and rare earth element (REE) patterns ofthe Surian basalts display depletions in light REE, negativeanomalies of Nb, Ta, and Ti, and positive anomalies of P.Positive P anomalies are indicative of minor crustal contam-ination. Furthermore, Th enrichments in the mid-ocean ridgebasalt-normalized patterns of the Surian basalts are character-istic of rifted arc basalts emplaced in continental marginsubduction zones. The high MgO content (>6 wt.%) of mostSurian basalts and low TiO2 content of two samples (0.53 and0.62 wt.%) are characteristic of boninites. The aforementionedfeatures of the basalts indicate arc tholeiites emplaced in intra-arc rift environments and continental margin subduction zones.U–Pb dating by laser ablation- inductively coupled plasmamass spectrometry of detrital zircons extracted from the hostfeldspathic and quartz feldspathic sandstone yields variousages that are predominantly Permian and Triassic; however,the youngest zircons give a mean Early Jurassic concordantU–Pb age of 191±12 Ma. This age, together with geologicaland petrochemical data, indicate that VMS mineralizationformed in the Early Jurassic in pull-apart basins within theSSZ. These basins and the VMS mineralization may be tem-porally related to an intra-arc volcano–plutonic event associatedwith Neo-Tethyan oblique subduction.

Keywords Bavanat . Besshi . VMS . U–Pb . SSZ . Iran

Introduction

Iran hosts numerous types of volcanogenic massive sulfide(VMS) deposits that occur within different tectonic assemb-lages and which formed at discrete time periods (Mousivandet al. 2008a). The metamorphic Sanandaj–Sirjan zone

Electronic supplementary material The online version of this article(doi:10.1007/s00126-012-0407-6) contains supplementary material,which is available to authorized users.

F. Mousivand : E. Rastad (*) :M. MohajjelDepartment of Geology, Faculty of Basic Sciences,Tarbiat Modares University,Tehran 14115-175, Irane-mail: rastad@modares.ac.ir

S. Meffre :K. ZawARC Centre of Excellence in Ore Deposits (CODES),University of Tasmania,Private Bag 79,Hobart, TAS 7001, Australia

J. M. PeterCentral Canada Division, Geological Survey of Canada,601 Booth St.,Ottawa, ON, Canada K1A 0E8

M. H. EmamiResearch Institute for Earth Sciences, Geological Survey of Iran,Tehran 13185-1494, Iran

Present Address:F. MousivandSchool of Geosciences, Shahrood University of Technology,Shahrood 3619995161-316, Iran

Miner DepositaDOI 10.1007/s00126-012-0407-6

(SSZ), particularly the southern part, is the most prospectivemetallotect for VMS exploration in Iran (Mousivand et al.2008a). The SSZ hosts four major VMS deposits, includingthe Bavanat Cu–Zn–Ag, Sargaz Cu–Zn, Chahgaz Zn–Pb–Cu,and Barika Au–Ag (Zn–Pb–Cu) deposits (Fig. 1). There arefew published geochronologic age determinations for theVMSdeposits in these terranes (including the SSZ). Age estimatesare mostly based on stratigraphic and geological relationships,but they are generally equivocal because multiple deforma-tional and metamorphic events have obliterated fossil evi-dence (e.g., radiolarians, foraminifera, and conodonts).

The Bavanat Cu–Zn–Ag VMS deposit is located in theBavanat area, 195 km northeast of Shiraz, in the southern partof the SSZ, southern Iran (Fig. 1) and about 170 kmNWof theChahgaz VMS deposit (Figs. 1 and 2). There are numeroussemimassive pyritic Cu–Zn deposits and occurrences in theBavanat area within the upper parts of the Surian volcano–sedimentary complex. The Bavanat deposit is the largest ofthese in the area, and is presently being mined. The JianCorporation Ltd. outlined ore reserves of around 6Mt gradingan average of 3.0% Cu, 0.5% Zn, and up to 68 ppm Ag,0.5 ppm Au and 1,300 ppm Co (Mousivand 2003; Mousivandand Rastad 2005; Mousivand et al. 2007).

The origin of the Bavanat deposit has been the focus ofmuch debate. Several workers proposed a metamorphicorigin for the deposit (e.g., Jannesari 2000a,b; Liaghat andTaghipour 2000; Taghipour and Moore 2002; Taghipour2000; Zarasvandi et al. 2001). Subsequently, Mousivand(2003) and Mousivand et al. (2001, 2004, 2007) suggestedthat it is a metamorphosed Besshi-type or pelitic mafic-type(Mousivand et al. 2008a) VMS deposit.

The deposit is hosted by the greenschist metamor-phosed Surian volcano–sedimentary sequence that waspreviously thought to be Late Devonian–Early Carbonif-erous (Houshmandzadeh and Soheili 1990a,b) or Permo-Triassic (Oveisi 2001) in age. Both age estimates werebased solely on stratigraphic and geological relationships.As part of our study, eight samples of fossiliferous thin-beddedlimestone belonging to the upper parts of the host se-quence were examined by Patrick Quilty of the University ofTasmania, but no diagnostic fossil evidence was found due tothe penetrative and pervasive deformation of the rocks in thearea. Vestigial crinoids and shell fragments were identified,but these provided no reliable biostratigraphic age estimates.Splits of the limestone samples were also dissolved in aceticacid at the Geological Survey of Iran, but no conodonts wereidentified in the residues. Two pelite samples were alsodigested in concentrated hydrofluoric acid at the GeologicalSurvey of Iran and examined for palynomorphs, but nonewere observed. Due to the lack of success of the biostrati-graphic studies, a laser ablation-inductively coupled plasmamass spectrometry (LA-ICPMS) U–Pb geochronologicstudy of zircons extracted from the feldspathic and quartz

feldspathic sandstone host rocks to the deposit was undertaken.Zircon is relatively abundant in feldspathic and quartz feld-spathic sandstone host rocks to the deposit, and this mineral isextremely resistant to resetting of U–Pb ages by deformationand metamorphism (Harley and Kelly 2007).

Previous geological studies (Alric and Virlogeux 1977;Houshmandzadeh and Soheili 1990a) suggested an intra-continental rift setting for the area based on the volcano–sedimentary nature of the Surian sequence, alkaline affinityof some basalts and the probable Paleozoic age. Based onthese data, Mousivand (2003) and Mousivand et al. (2007)also supported such an interpreted setting for deposition ofthe Bavanat VMS deposit.

Herein, we present new U–Pb geochronology data for theBavanat deposit and elucidate the tectonomagmatic setting forthe mineralization. These U–Pb ages should be of interest tomineral explorationists working in the SSZ, but they also havea broader geological significance because the Bavanat districtis part of the Dehbid–Neyriz–Shahre Babak region, an areacritical to understanding the opening, spreading, and closureof the Neo-Tethyan Ocean.

Geotectonic setting

The Tethyan orogen was caused by the collision of Eurasiawith Gondwanaland (Sengör and Natal’in 1996). The Zagrosorogenic belt is part of the Tethyan orogen (e.g., Takin 1972;Stöcklin 1974; Berberian and King 1981; Sengör 1991; Alavi1994; Shahabpour 2005) and is comprised of four paralleltectonic subdivisions from southwest to northeast: (1) theZagros simply folded belt, (2) High Zagros belt, (3) theSanandaj–Sirjan zone (SSZ), and (4) the Urumieh–Dokhtarmagmatic arc (Ghasemi and Talbot 2006; Figs. 1 and 2).

The 150–250 km-wide SSZ extends over a strike distanceof 1,500 km from the southeast to northwest of Iran (Fig. 1).The SSZ is characterized by regionally metamorphosed anddeformed rocks that are spatially associated with abundantdeformed and undeformed plutons, as well as widespreadMesozoic volcanic rocks. The Chahdozdan and Bon-Donointrusions are examples of the intrusions in the SSZ (Fig. 2).Berberian (1983) considered this zone to be a Mesozoicmagmatic arc and a Tertiary fore-arc. Two magmatic arcsof Mesozoic and Tertiary age (the Sanandaj–Sirjan andUrumieh–Dokhtar magmatic arcs, respectively) formeddue to the convergence of Arabia and Eurasia along thestrike of the Zagros Mountains (Omrani et al. 2008; Figs. 1and 2). Magmatism was active during subduction, fromEarly Jurassic to the Late Eocene, and later resumed in theUrumieh–Dokhtar magmatic arc after the onset of collision,from Late Miocene onwards (Omrani et al. 2008). TheSirjan depression, which is located between theSSZ (Mesozoic magmatic arc; Tertiary fore-arc) and the

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Urumieh–Dokhtar volcanic belt (Tertiary magmatic arc;Figs. 1 and 2) is interpreted to be an inter-arc basin(Shahabpour 2005).

Regional geology of the Bavanat area

Detailed descriptions of the geology of the area have beengiven by several authors (Harrison et al. 1936; Ricou 1974;Taraz 1974; Alric and Virlogeux 1977; Pourkermani 1977;Houshmandzadeh and Soheili 1990a; Oveisi 2001; Mousivand2003).

Within the SSZ, the Bavanat area is represented by anarrow linear belt of deformed and metamorphosed volca-no–sedimentary sequences that are collectively 5–10 kmwide and about 100 km long (Fig. 3). The overall strike ofthe belt is broadly NW–SE, parallel to the general regionaltrend of the Zagros belt and the Urumieh–Dokhtar magmaticarc (Figs. 2 and 3).

Structurally, the Bavanat area is bounded by two majorfaults; the Jian listric fault along its southwestern contact withthe High Zagros Belt and the Surian fault along its northeast-ern side (Fig. 3). This trend has been modified by subordinateNE–SW faults. The area is characterized by vertical to sub-vertical isoclinal folds and thrust faults, the main trends ofwhich are NW–SE. A major regional fold (anticline) through-out the Bavanat area strikes NW–SE over a distance of about100 km, from Safa Shahr (Dehbid) to Harat (Fig. 3).

The Bavanat area is comprised of three major deformedand metamorphosed volcano–sedimentary stratigraphicsequences (from base to top): (1) the middle–upper Devo-nian Toutak complex forms the core of the anticline andincludes marble, mica schist, and amphibolite; (2) the EarlyJurassic (this study, see below) Surian complex whichincludes metamorphosed pelite, sandstone, calcareous shale,basalt, gabbro sills, and thin-bedded limestone; and (3) theTriassic–Jurassic Kulikosh complex that consists of meta-morphosed carbonate and clastic sedimentary and volcanic

Urumieh-Dokhtarmagmatic arc

Ophiolite

City/Town

Sanandaj-Sirjan zone

VMS deposit

N

200 km60

35

5550

30

35

30

6055

ArabianPlatform

Caspian Sea

Tehran

4

1

3

2

Nain

Shiraz

50

BoroujerdCentral Iranianmicrocontinent

BaftShahre Babak

Esfandagheh

Hamadan

Fig. 1 Tectonic subdivisions ofthe Zagros Orogen (modifiedafter Ghasemi and Talbot 2006)and location of VMS depositsin the Sanandaj–Sirjan zone:1 Bavanat (Early Jurassic),2 Sargaz (Late Triassic–EarlyJurassic); 3 Chahgaz (MiddleJurassic); and 4 Barika(Middle–Late Cretaceous)

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rocks. Both the Toutak and Surian complexes extend to theHarat area; however, the name of both of these complexeschanges (to Harat complex) in the Anar map area (Soheili1981).

The marbles of the Toutak complex host stratiform magnetitedeposits that have been mined in ancient times at Kan Gohar(e.g., Ebrahimi 1999). The Kulikosh complex also contains

similar iron deposits that are hosted in carbonates and rhyoliticvolcanic rocks (e.g., Kazemi Rad et al. 2010).

Two main phases of metamorphism have affected rocksin the area. Strata of the SSZ have undergone amphibolitefacies dynamothermal metamorphism, followed by a retro-grade greenschist facies regional dynamothermal metamor-phism that has obscured features of the earlier amphibolite

N

HaratShahre Babak

Safa Shahr(Dehbid)

Marvast

Zagros zone

Sanandaj-Sirjan zone

Urumieh-Dokhtar magmatic arc

Ophiolite

Mesozoicintrusions

VMS depositZagros fault, suture

Chahdozdan anorthosite-granitoid intrusions

Chahghand and Robat Gabbro- pyroxenite intrusions

Bon-Dono granitoidintrusion

City/Town

54 5529

30

31

30

3153 54 55

Bon-Dono intrusion

Robat intrusion

Chahghand intrusion

Chahdozdan intrusion170.5 1.9 Ma

Chahgaz Zn-Pb-Cu deposit174.1 1.2 Ma+

+

Bavanat Cu-Zn-Ag deposit191 12 Ma (this study)+

60 km

Neyriz

Fig. 2 Simplified geologicalmap of the Dehbid–Neyriz–Shahre Babak region, basedon the Eqlid, Anar and Neyriz1:250,000 scale maps (e.g.,Houshmandzadeh and Soheili1990b; Soheili 1981; Sabzeheiet al. 1993), showing tectonicsubdivisions and location of themajor Mesozoic intrusionswithin the area and the BavanatCu–Zn–Ag VMS depositwhich is about 170 km NW ofthe Chahgaz Zn–Pb–Cu VMSdeposit. Age data for theChahdozadan intrusion andthe Chahgaz deposit are fromFazlnia et al. (2007) andMousivand et al. (2011),respectively

Major fault

Axis of anticline

Bavanat deposit

Copper occurrence

Triassic-Jurassic Kulikosh complex: mafic volcanics, rhyolite, pelite and marble

Early Jurassic Surian complex: pelite andmafic volcanics and sills

Late Devonian Toutak complex: marble,pelite and mafic volcanics

Jian Surian Chir Surian fault Harat

Late TriassicBon-Dono granitoidintrusion

Town

Village

Mazijan

10 km

Safashahr(Dehbid)

SSZ

High Zagros Belt

Fig. 3 Simplified geological map of the Bavanat area showing loca-tions of the Bavanat deposit and other copper occurrences hosted in theSurian complex. The Toutak complex forms the core of the anticlinal

structure, and the Kulikosh complex outcrops in the extreme northernpart of the Bavanat area

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facies event (Houshmandzadeh and Soheili 1990a). Bothphases are of Barrovian-type (medium pressure) metamor-phism (Alric and Virlogeux 1977; Houshmandzadeh andSoheili 1990a). Alric and Virlogeux (1977) suggested thatboth metamorphic events are related to the Early Cimmerianorogeny. However, based on Houshmandzadeh and Soheili(1990a), the amphibolite facies metamorphism is interpretedto record the early Cimmerian (Middle to Late Triassic)event and the greenschist facies metamorphism the lateCimmerian (upper part of Mid-Jurassic) event. Although theToutak complex displays evidence of both the amphiboliteand greenschist facies metamorphisms, the Surian complexonly displays effects of the later event (Mousivand 2003).Two deformational phases have affected the rocks of theSurian complex, and these have imparted folds, faults andfoliations (NW–SE trending S1 and S2). The S2 foliation isparallel to the axial planes of the folds.

Deposit geology and host rock petrography

The Bavanat VMS deposit occurs on the southern limb of ananticline with a NE-trending axial plane. The deposit strati-graphic succession dips subvertically and youngs consis-tently to the south. Faults in the area trend mainly NE–SWand a few trend NW–SE.

The deposit is hosted by the Surian volcano–sedimentarycomplex which averages about 1,800 m in thickness(Houshmandzadeh and Soheili 1990a) and has been dividedinto four units (from bottom to top; Mousivand et al. 2007;Fig. 4):

Unit 1, metamorphosed conglomerate and associated micaschist and pelite (~500 m thick); unit 2, metamorphosedintermediate to mafic volcanic and subvolcanic rocks, andinterbedded pelitic rocks (~400 m thick); unit 3, a ~300 mthick sequence of gray to black metapelites; unit 4, metamor-phosed dominantlymafic to intermediate subvolcanic sills andvolcanic rocks intercalated with quartz–feldspathic and feld-spathic sandstone and gray to black pelites (~600 m thick).Unit 4 is capped by fossiliferous microsparite and metagrey-wacke (Fig. 4).

The host rock sequence to the Bavanat deposit is com-prised of intercalated feldspathic and quartz feldspathicsandstone, pelites, and basalts. Mineralogically, the feld-spathic and quartz feldspathic sandstone rocks are com-posed predominantly of fine-grained chlorite, plagioclase(up to 2 mm diameter), fine-grained quartz, and minorcalcite together with zircon (Fig. 5a, b), sericite, sphene,rutile, and ilmenite. Up to 1 vol.% zircon is present in theserocks as euhedral to rounded grains ranging from 10 to30 μm in diameter (Fig. 5b, c). The pelites contain fine-grained muscovite, quartz, calcite, chlorite, and organiccarbon. The darker-colored (black) pelites contain higher

total organic carbon or graphite and lower calcite contentsthan the lighter-colored (gray) pelites. The basalts are com-prised of dominantly chlorite and minor tremolite, actinolite,epidote, plagioclase, and calcite.

Mineralization and hydrothermal alteration

Cu–Zn–Ag VMS mineralization in the Surian volcano–sedimentary complex occurs along two discrete, discontin-uous narrow horizons that have a strike length of >100 kmthroughout the Bavanat area (Fig. 3). A lower mineralizedhorizon hosts numerous occurrences in the upper part of unit2, and an upper mineralized horizon hosts the Bavanat depositin the lower–middle part of unit 4 (Fig. 4; Mousivand et al.2007). Mineralization in the Bavanat deposit occurs as severalstratiform, sheet-like and tabular sulfide orebodies, each ofwhich is stratigraphically underlain by a stringer zone (Fig. 6)of quartz–sulfide veins and veinlets. The mine stratigraphy isoverturned (Fig. 7). The orebodies are up to 1 m thick and upto 40 m long (Fig. 7).

Host rocks to the mineralization are metamorphosed basalt,quartz feldspathic and feldspathic sandstone, and gray to blackpelites. The immediate stratigraphic footwall and hangingwall to the mineralization are metamorphosed pelites. Thefootwall pelites are darker colored than the hanging wallpelites due to their higher organic carbon contents. The uppermineralized horizon has a red to brown gossan-like layerdirectly adjacent (above) the orebodies. This layer is domi-nantly comprised of quartz, and minor iron and manganeseoxides, and may be a product of primary weathering on theseafloor.

Ore minerals in the mineralized horizon are dominated bypyrite with subordinate amounts of chalcopyrite, sphalerite,pyrrhotite, marcasite, and rare to minor galena and cubanite.The gangue minerals are mainly quartz, chlorite, plagioclase(close to albite end-member), sericite, and carbonate (mainlysiderite).

Four ore types are recognized in the Bavanat deposit: (1)semimassive to massive, (2) banded–bedded–laminated, (3)disseminated, and (4) stringer vein. Three types of massivesulfide ore are recognized, on the basis of predominance ofsulfide minerals: (1) pyrite-rich, (2) chalcopyrite-rich, and (3)pyrrhotite-rich. The massive pyrite-rich ore is dominated byfine- to coarse-grained (up to 1 cm) anhedral pyrite, and showsbreccia textures in places. The massive chalcopyrite ore maycontain chlorite- and quartz-rich clasts whereas the massivepyrrhotite ore may contain quartz-rich stringer clasts. Thestringer ores consist of deformed quartz-carbonate sulfideveins.

A variety of primary and secondary textural featuresoccurs in these ores. The dominant primary texture withinbedded ores is alternating sulfide- and feldspar-,quartz-,

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carbonate-rich bands within the quartz–feldspathic and fel-spathic sandstones and calcareous shales that are interpretedto be primary layering that have been overprinted by S1 andS2 foliations associated with folding. Minor secondarybanded ores comprised of macroscopic bands of sulfideand silicate minerals also occur but show only S2 foliation.Several other secondary ore textures are observed, includingboudinage and cataclastic textures, micro- and macroscalefolding, triple junctions, elongations, pressure shadows, andopen space fillings. These textures are typical of deformedand metamorphosed sulfide ore deposits (e.g., Vokes andCraig 1993; Vokes 1995, 2000a,b; Craig and Vokes 1993;Marshall et al. 1999, 2000; Marshall and Spry 2000).

Well-developed metal andmineralogical zonations are pres-ent in the Bavanat orebody. Chalcopyrite is more abundant in

the stratigraphic lower part of the sulfide orebodies near thefootwall, whereas sphalerite and/or galena are more abundantin the stratigraphic upper parts of the orebodies near thehanging wall. Concomitantly, Cu contents decrease and Znand/or Pb contents increase from the footwall toward thehanging wall.

Host rocks to the Bavanat deposit have been intenselyhydrothermally altered and deformed, with the alterationzones developed predominantly in the footwall withinand around the stringer zones. Chloritization is the pre-dominant wall rock hydrothermal alteration style with subor-dinate argillic and sericitic alteration. The chloritized hostrocks were converted to chlorite schist and quartz chloriteschist during penetrative deformation (Fig. 5d). As a resultof deformation and metamorphism, stringer veins were

Age Symbol LithologyOre -bearing

horizons

Marble

Meta-pelite

Meta-pelite and quartzite

Amphibolite

Meta-conglomerate andmeta-greywacke

Ore-bearing chlorite schist andmeta-pelite

Mafic to intermediate meta-volcanics

M eta-greywacke

Intermediate meta-volcanics and meta-volcaniclastics

Meta-basic sills and volcanics

Ore-bearing chlorite schist, basic meta-volcanics, graphite schist, meta-calcarousshale, meta-sandstone and meta-pelite

Bavanat deposit,and Didebanki,Surian 3-6 andChir 9occurrences

Chir 1-7, andSurian 1, 2occurrences

200 m

100

0

Unconformity

Fossiliferous thin-bedded limstone(microsparite)

Fig. 4 Stratigraphic section ofthe Surian volcano–sedimentary complex and itsore horizons

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transposed into quartz–sulfide-rich bands developed throughfoliation (Fig. 5d).

Petrochemistry of the volcanic rocks

Major, trace, and rare earth element (REE) contents weremeasured on representative unaltered samples of the Bava-nat host basalts (ESM Table 1). The samples were analyzedusing both X-ray fluorescence at Tarbiat Modares Universi-ty of Iran and ICPMS and inductively coupled plasmaatomic emission spectrophotometry (ICPAES) methods atAcme Laboratories, Vancouver, Canada, and the Universityof Tasmania, Hobart, Australia.

In the immediate vicinity of the Bavanat deposit, thebasalts are intensely altered to an assemblage of chlorite–calcite–pyrite with minor sericite. The alteration intensitydecreases with increasing distance from zones of minerali-zation. To characterize the primary petrochemical character-istics of the basalts, we have relied primarily on theimmobile major element TiO2, the high field strength ele-ments Zr, Nb and Y, and REE. We have not used most of themajor elements or low-field strength elements as these arereadily mobilized during hydrothermal alteration and meta-morphism and substitute for Na, K, and Ca.

The basalts show subalkaline to alkaline affinities, andbasaltic to andesitic–basaltic compositions on the discriminant

plot of Winchester and Floyd (1977; Fig. 8a), and displaytholeiitic to transitional affinities on the Zr versus Y diagramof Barrett and MacLean (1999; Fig. 8b). On the Zr versusZr/Y binary diagram of Pearce and Norry (1979) and Zr/4–2Nb–Y ternary diagram of Meschede (1986), these samplesdominantly plot in the mid-ocean ridge basalt (MORB)and within-plate tholeiitic basalt fields (Fig. 8c, d). On the Zr–Ti/100–3Y ternary diagram of Pearce and Cann (1973) theydominantly show MORB and volcanic arc basalt (VAB) orcalc–alkaline basalt features (Fig. 8e). On the V versusTi/1,000 binary diagram of Shervais (1982), the samples liedominantly in the MORB and BABB fields, but two samplesdisplay island arc tholeiite (IAT) and boninitic compositions(Fig. 8f).

On a primitive mantle-normalized spider diagram (Sunand McDonough 1989) the basalts have overall low traceelement abundances, show no light rare earth element(LREE) enrichment, but display positive P anomalies andnegative Nb, Ta, and Ti anomalies (Fig. 9a). Furthermore,on an N-MORB-normalized spider diagram the Surianbasalts show Th enrichment (Fig. 9b). The Surian basaltsamples (ESM Table 1) have variable SiO2 (43.3–49.0 wt.%,with two high outliers at 55.8 and 61.7 wt.%), TiO2 (0.6–2.8 wt.%, with the exception of one low outlier at 0.5 wt.%),MgO (3.3–13.8 wt.%, with most >6 wt.%), and Al2O3 (12.3–17.7wt.%), Ni (10–144 ppm), Co (19.9–43 ppm), and Cr (65–473 ppm) contents (ESM Table 1).

Fig. 5 a Transmitted light(crossed polars)photomicrograph of hostfeldspathic and quartzfeldspathic sandstone rock,showing feldspar (f), quartz(qz), chlorite (chl), calcite (ca),and pyrite (py). b Transmittedlight photomicrograph of hostfeldspathic and quartzfeldspathic sandstone rockshowing zircon grain (zr) andquartz (qz) within matrix ofchlorite (chl). The quartz grainshave been elongated along withchlorites due to deformation. cCathodoluminescence (CL)image of a single zircon grainshowing internal zonation. dSulfide (mainly pyrite)–quartz-rich stringer veinstransposed parallel to majorfoliation of the host feldspathicand quartz feldspathicsandstone (chlorite schist).The veins show late faulting(dashed lines)

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Zircon morphology and U–Pb geochronology

Four samples of hydrothermally altered feldspathic andquartz feldspathic sandstone host rocks were collected fromdrill cores from the Bavanat deposit for age dating. Thezircons are euhedral to subhedral with prominent growthzoning that is continuous from core to rim (Fig. 5c). Thezircons are free of any inclusions of other minerals such assulfides and silicates, are brownish to colorless transparentand show well-developed oscillatory zonation.

Approximately 100 g of each rock sample was repeatedlysieved and crushed in a Cr-steel ring mill to a grain size of<400 μm. Nonmagnetic heavy minerals were then separatedusing a gold pan and a Fe–B–Nd hand magnet. The zirconswere hand-picked from the heavy mineral concentrate under abinocular microscope in cross-polarized transmitted light. Theselected crystals were placed on double-sided sticky tape andepoxy glue was then poured into a 2.5 cm diameter mold ontop of the zircons. The mount was dried for 12 h and polished

using clean sandpaper and a clean polishing lap. The sampleswere then washed in distilled water in an ultrasonic bath. Afterpolishing to expose the top of the crystals, the zircons wereexamined under a microscope using cathodoluminescence toreveal any internal structures. The zircons were analyzedusing the LA-ICPMS method to measure U, Th, and Pbisotopic ratios (e.g., Fryer et al. 1993; Compston 1999; Blacket al. 2003, 2004; Kosler and Sylvester 2003; Jackson et al.2004; Harley and Kelly 2007).

The analyses were performed on an Agilent 7500a quadru-ple ICPMS with a 193 nm NewWave Laser solid-state laser atthe University of Tasmania in Hobart. The downhole fraction-ation, instrument drift and mass bias correction factors for Pb/Uratios on zircons were calculated using four analyses on theprimary (Temora standard of Black et al. 2003) and twoanalyses on the secondary standard zircons (91500 standardof Wiedenbeck et al. 1995) determined at the beginning of thesession and every 12 unknown zircons (approximately everyhour) using the same spot size and conditions as the unknowns.

DH6DH1

DH2 & 5

DH9 & 10DH7 & 8

DH11 &12

DH13 & 14 & 15

A

B

81

79

77

7578

82

77

80

80

DH7

25 m

Altered host rocks (chlorite schist) includingbasalt and sandstone

Black meta-pelite

Grey meta-pelite

Massive sulfideorebody

Drill hole location

N

Trace of cross-sectionshown in figure 7

Stringer zone

Strike and dip80

Fig. 6 Surface geologicalmap of the Bavanat deposit.Shown are the differentsheet-like massive sulfidelenses and stringer zones thatstratigraphically underlieseveral of the lenses. Themineralization is hosted inchloritized feldspathic andquartz feldspathic sandstoneand basalt intercalatedwith pelites

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Additional secondary standards (The Mud Tank Zircon ofBlack and Gulson 1978) were analyzed after the sample. Thecorrection factors for the 207Pb/206Pb ratio were calculatedusing six large spots (100 μm diameter, 10 Hz laser) ofNIST612 analyzed at the beginning and end of the day usingthe values recommended by Baker et al. (2004).

Each analysis began with a 30 s blank gas measurementfollowed by a further 30 s of analysis time when the laser wasswitched on. Zircons were sampled on 35 μm spots using the

laser at 5 Hz and a density of approximately 1.5 J/cm2. A flowof He carrier gas at a rate of 0.5 l/min carried particles ablatedby the laser out of the chamber to be mixed with Ar gas andcarried to the plasma torch. Elements measured include 49Ti,96Zr, 178Hf, 202Hg, 204Pb, 206Pb, 207Pb, 208Pb, 232Th and 238Uwith each element being measured sequentially every 0.14 s,with longer counting time for the Pb isotopes than for theother elements. The data reduction methods used are basedon the method outlined in detail in Appendix 1 of Meffre

DH7 DH8DH9

DH10DH12 DH11

Old workings

A B

4 m

Zr

Zr

Zr

Zr

Fig. 7 Cross-section A–B (forlocation of trace, see Fig. 6),showing diamond drill holetraces, distribution of sheet-likemassive sulfide mineralizationand the underlying stringerzones (note that the stratigraphyis overturned.). Zr 0 U–Pbsample location. For legend,see Fig. 6

Miner Deposita

Nb/Y

0.1

Trachyte

Phonolite

Andesite

Rhyodacite+ Dacite

And/Bas-And

Rhyolite

Alk-Bas

TrachyAnd

Sub alkaline Basalt

.01 1 10

5

1

0.1

0.01

.002

Com/Pant

Bsn/Nph

Zr ppm

B

B: Island arc basalts

C

A

101

100 1000

C: Mid oceanic ridge basalts

A: Within plate basalts

10

20

AI, AII = Within plate alkaline

AIl, C = Within plate tholeiitic

B = E-MORB

D = N-MORB

C, D = VAB

Nb*2

Zr/4 Y

AI

AII

B

CD

(a)

(c) (d)

A = Volcanic-arc basalts

B = MORB + VAB

C = VAB

D = Within plate basalts

Ti/100

Zr Y*3

AI

DB

C

A

Ti ppm/1000

IATBoninite andLOTI-Associated MORB-Associated

0

100

200

400

300

500

600

0 10 15 20 255

MORB/BABB

Low Ti IATBON ARC <20> OFB

100

50

10

(e) (f)

Zr ppm

Calc-alkaline

Transitional

Tholeiitic

0

10

20

40

30

50

60

0 100 150 200 25050

(b)

Fig. 8 Selected discriminant diagrams for mafic volcanic host rocks tothe Bavanat deposit. a Zr/TiO2 versus Nb/Y diagram of Winchester andFloyd (1977), showing basaltic to andesitic compositions; b Bivariateplot of Yversus Zr (after Barrett andMacLean 1999) showing dominantlytholeiitic to transitional magmatic affinities; c Zr versus Zr/Y binarydiagram of Pearce and Norry (1979) and d Zr/4–2Nb–Y ternary diagram

of Meschede (1986), both showing dominantly mid-ocean ridge basalts(MORB) and within-plate tholeiitic affinities; e Zr–Ti/100–3Y ternarydiagram of Pearce and Cann (1973) illustrating the dominantly MORBand calc-alkaline basalt features; f V versus Ti/1000 binary diagram ofShervais (1982) illustrating dominantlyMORB and back arc basin basalts(BABB) affinities, but two samples plot in IAT and boninitic field

Miner Deposita

et al. (2007, 2008) modified from that of Black et al.(2004) to suit the ICPMS and laser at the University ofTasmania.

Element abundances on zircons were calculated employ-ing the method outlined by Kosler (2001) that uses Zr as theinternal standard element, assuming stoichiometric propor-tions and using the secondary standard 91500 to correct formass bias.

Geochronology results

Zircons from the feldspathic and quartz feldspathic sand-stone host rocks yield U–Pb ages ranging from Precambrianto Mesozoic (Fig. 10 and ESM Table 2). The oldest age is695±12 Ma, and the youngest ages obtained are 181±2,189±6, and 201±15 Ma. The mean age of the three youn-gest zircons is 191±12 Ma. Both rims and cores of somegrains have been analyzed. The cores typically show olderages than rims. For example, in one grain the core and rimyield ages of 360 and 307 Ma, respectively.

Discussion

Petrochemical characteristics of the Surian basalts

Alric and Virlogeux (1977), Houshmandzadeh and Soheili(1990a), Mousivand (2003), and Mousivand et al. (2007)suggested an intracontinental rift setting for the Bavanatarea, based on the alkalic nature of some of the basaltsintercalated with sedimentary rocks and the Late Devonian–Early Carboniferous and Permo-Triassic age estimates. Sincea magmatic arc setting had not been recognized in thePaleozoic rocks of Iran, these age estimates for the Suriancomplex were interpreted by previous workers to reflect de-position of the complex in an intracontinental rift setting.However, our U–Pb age determinations and petrochemicaldata do not support this interpretation. The MORB and VABaffinities of the Surian basalts are typical of rift settings withincontinental margin arcs and subduction zones (e.g., Pearce1996; Murphy 2006). The alkalic nature of a few samples mayindicate that volcanism occurred in a nascent or primitive arcenvironment (e.g., Murphy 2006, 2007). Primitive mantle-normalized trace and REE patterns of the Surian basalts dis-play depletions in LREE, negative Nb, Ta, and Ti anomaliesand positive P anomalies. These features are characteristic ofarc tholeiites and primitive arc environments (e.g., Swinden1996; Barrie and Pattison 1999; Wyman et al. 1999; Wyman2000; Piercey et al. 2001b; Arndt 2003; Murphy 2007). Pos-itive P anomalies reflect minor crustal contamination (e.g.,Kerrich and Wyman 1996). Furthermore, Th enrichment inthe MORB-normalized patterns of the Surian basalts is afeature common to rifted arcs belonging to continental marginsubduction zones (e.g., Pearce 1996). The variable SiO2 andMgO contents of the Surian basalts may reflect element mo-bility during hydrothermal alteration, alteration by seawater,and/or metamorphism. The high MgO content (>6 wt.%) ofmost of the Surian basalt samples, and low TiO2 of two

HfEu Ti

Gd TmDyY Er U

YbLu

HoTa

10000

100

LaCe P

Nd SmZrNbTh

1

0.1

10

1000

Th Nb Ce Zr Ti Y

0.1

1

10

(a)

(b)

Fig. 9 a Primitive mantle-normalized (Sun and McDonough 1989)spider diagram for mafic volcanic host rocks to the Bavanat deposit.The samples lack LREE enrichment and display negative Nb, Ta, andTi anomalies, and positive P anomalies; b MORB-normalized spiderplot of the Surian mafic volcanic host rocks to the Bavanat depositshowing Th enrichment

150 250 350 450 550 650 750 800500 600 7004003002000

1

2

3

Age (Ma)

PrecambrianPaleozoic

Fig. 10 Histogram of all 207Pb corrected 206Pb/238U LA-ICPMS zir-con age data for the Bavanat host feldspathic and quartz feldspathicsandstones

Miner Deposita

samples (0.53 and 0.62 wt.%) are characteristic of boninites(e.g., Crawford et al. 1989; Pearce et al. 1992; ESM Table 1).High SiO2 contents in two samples (55.8 and 61.7 wt.%), highCo in most samples (mean of 33.2 ppm), moderate to high Niand Cr in several samples, and moderate Al2O3 contents(12.3–17.7 wt.%, ESM Table 1) are also features of boninites(e.g., Crawford et al. 1989; Pearce et al. 1992).

Pelitic-mafic or mafic-siliciclastic (or Besshi)-type VMSdeposits are associated with MORB (e.g., Middle Valley,Guaymas, Escanaba Trough), boninite (e.g., Fyre Lake) andalkalic/ocean island basalts (e.g., Windy Craggy; Peter andScott 1999; Piercey 2011). The Surian basalts show all thesefeatures.

The REE and trace elements patterns of the Surian basalts aresimilar to high-Ca boninites (HCB; e.g., Crawford et al.1989;Meffre et al. 1996), but show greater element enrichments(Fig. 11). Furthermore, the moderate Sc contents (30–41 ppm,ESM Table 1) of the basalts are also typical of HCB (e.g.,Crawford et al. 1989; Meffre et al. 1996).

Boninites are thought to have formed in the followingtectonic settings: (1) the initiation stages of subduction(Stern and Bloomer 1992), (2) intra-arc rifting (Crawfordet al. 1981), (3) ridge subduction (Cameron 1989), (4)seafloor spreading within a fore-arc basin (Bedard et al.1998), (5) an active arc undergoing back-arc spreading(e.g., Falloon and Crawford 1991; Monzier et al. 1993),(6) nascent or primitive arcs (Crawford et al. 1989), and(7) continental margin environments (e.g., Piercey et al.2001a; Rogers and Saunders 1989). It is generally acceptedthat the SSZ was the southwestern continental margin of theIranian plate. Neo-Tethyan subduction of the oceanic crustbeneath the Iranian microcontinent likely occurred at varioustimes between the Late Triassic and the Late Cretaceous(Berberian and King 1981; Alavi 1994; Mohajjel et al. 2003;

Shahabpour 2005, 2007). The Late Triassic Siahkuh calc–alkaline granitoid intrusion located in the southern SSZ, closeto the Zagros suture, was emplaced during the initial phase ofsubduction (Arvin et al. 2007). Therefore, the Late Triassicage of the calc–alkaline Bon-Dono granitoid intrusion alsoreflects the time of initiation of subduction in the Bavanatregion. The Bon-Dono intrusion is close (<20 km) to theZagros suture (Fig. 2). The presence of some Surian basaltsof boninitic affinity might indicate a fore-arc, or intra-arcrifting, or an active arc undergoing back-arc spreading, ornascent/primitive arc environment for the Bavanat region.The close proximity of the Bavanat area to the Zagros suture(Fig. 2) supports a fore-arc setting for magmatism in the area.Also, the MORB–BABB-like and HCB-like nature (Figs. 8fand 11) of some Surian basalts may indicate a back-arc envi-ronment (e.g., Hawkins 1995; Crawford et al. 1981, 1989;Piercey et al. 2001a; Piercey 2011). However, the coexistenceof the calc–alkaline Bon-Dono granitoid and the Surian tho-leiitic to transitional basalts in the Bavanat area (Fig. 3) andvolcano–sedimentary nature of the Surian sequence morelikely indicates formation in an intra-arc rift setting associatedwith continental margin subduction.

Timing of the mineralization

The oscillatory growth zonation of the zircons of the Bavanathost feldspathic and quartz feldspathic sandstone indicates amagmatic origin (e.g., Corfu et al. 2003). The origin of growthzoning in zircons is discussed in detail by Mattinson et al.(1996) who conclude that episodic growth results from theinterplay between the stage of crystal growth, the nature of thecrystal–liquid interface, the degree of supersaturation ofthe melt, the rates of diffusion, and the state of oxidationof the magma. The zoning reflects compositional varia-tion (up to an order of magnitude for some elements) ofZr, Si, Hf, P, Y, the REE, U and Th (e.g., Hanchar andRudnick 1995; Fowler et al. 2002). The composition ofthe growth zones tends to vary between two end-members,one of which is very low in trace elements, approaching thecomposition of pure zircon, and the other highly enriched intrace elements with up to several weight percent of the impu-rity element (Speer 1982).

The large spread in zircon ages obtained may be due to lossof lead since time of deposition and/or the detrital nature of thezircons within the host feldspathic and quartz feldspathicsandstone. Additionally, the spread in ages might indicate thatthe zircons were sourced both from felsic volcanic or sedi-mentary rocks that are much older than the Bavanat basaltsand from magmas that were emplaced at the same time as thebasalts. The mean age of the youngest zircons (191±12Ma) isEarly Jurassic, and this is interpreted to be the depositional ageof the feldspathic and quartz feldspathic sandstone host rocksof the Bavanat deposit. Such rocks are deposited essentially

HfEu

TiGd

TbDy

YEr V

YbLu

Al Sc

100

10

0.1La

CePr

NdSm

ZrNbTh

1

Fig. 11 Primitive mantle-normalized (Sun and McDonough 1989)spider diagram for mafic volcanic host rocks to the Bavanat deposit.Also shown for comparison is the (gray) field for high-Ca boninites(HCB; e.g., Crawford et al. 1989; Meffre et al. 1996); the Bavanatsamples show a similar pattern to HCB, and also display Nb and Tianomalies, but they have higher abundances of most of these elements

Miner Deposita

instantaneously, and therefore this age is also taken to be theage of the stratabound and stratiform VMS mineralization.

Implications for regional magmatism and tectonics

The age and geochemical characteristics of the Bavanat hostvolcanic rocks can be attributed to the subduction of Neo-Tethys oceanic crust beneath the Central Iranian microcontinentin the Mesozoic (e.g., Berberian and King 1981; Sengör 1990;Ghasemi and Talbot 2006; Arvin et al. 2007; Omrani et al.2008; Sheikholeslami et al. 2008; Fazlnia et al. 2009). Varioustime periods have been suggested by different authors for theonset of subduction of the Neo-Tethys oceanic crust: Triassic(Arvin et al. 2007), Late Triassic (Sheikholeslami et al. 2008),Early Jurassic (Omrani et al. 2008), Late Jurassic (Fazlnia et al.2009), and Late Jurassic–Early Cretaceous (Sengör 1990;Ghasemi and Talbot 2006).

Geochemical features of the Late Triassic Bon-Dono intru-sion including the calc–alkaline, I-type and post-orogenic af-finities (Noori Khankahdani et al. 2006) can be attributed tosubduction of the Neo-Tethyan oceanic crust beneath the Ira-nian plate. The existence of clasts of this intrusion in the basalconglomerate of the Surian complex indicates the onset ofarc plutonism in the Bavanat area prior to tholeiitic totransitional volcanism. Similarly, Late Triassic–Early Ju-rassic volcano–plutonism has also been documented inthe Esfandagheh region (Babakhani and Alavi Tehrani1992; Sahandi 2001; Badrzadeh 2009; Badrzadeh et al.2011) 400 km SE of the Bavanat area (Fig. 1). Here, theLate Triassic, calc–alkaline to transitional Siahkuh granitoid(Arvin et al. 2007) predates tholeiitic to transitional bimodalvolcanism in the area (Babakhani and Alavi Tehrani 1992;Sahandi 2001). Similar Late Triassic–Early Jurassic andMiddle Jurassic volcano–plutonism also occurs in the Chah-gaz area, about 170 km southeast of the Bavanat area(Sheikholeslami et al. 2008; Fazlnia et al. 2007; Mousivandet al. 2011; Figs. 1, 2). In the Chahgaz area, the Chahghandand Robat gabbro pyroxenite and the calc–alkaline Chahdoz-dan granitoid (Fig. 2; Sabzehei et al. 1993) and anorthositic(Fazlnia et al. 2007) intrusions are coeval with or slightlypredate the bimodal volcanism (Mousivand et al. 2011).

Petrochemical data for the Chahghand intrusion (e.g., Otrodi2006; Otrodi et al. 2006) show transitional to calc–alkaline andwithin-plate signatures. However, an intracontinental rift hasbeen suggested as the emplacement setting for the intrusion bythese authors. Sheikholeslami et al. (2008) suggested thatemplacement of the Chahdozdan and Chahghand intrusionsoccurred within extensional basins due to oblique subductionat 167 and 159Ma, respectively. Fazlnia et al. (2009) suggestedthat opening and expansion of the Neo-Tethys Ocean occurredin the Middle Jurassic (173.0±1.6 Ma) at the time of emplace-ment of the Chahdozdan anorthosite intrusion and, based onage of the trondhjemites, suggested subduction beneath the

southern part of the SSZ commenced in the Late Jurassic(Volgian; 147.4±0.76 Ma).The calc–alkaline nature of theChahdozdan and Chahghand intrusions (e.g., Sheikholeslamiet al. 2003) cannot be attributed to the opening of the Neo-Tethyan Ocean, as proposed by Fazlnia et al. (2009), butmay be related to the same subduction event recorded in theChahgaz volcanic rocks.

Our study suggests that intra-arc-related tholeiitic to tran-sitional volcanism in the Bavanat area occurred at 191.0±12Ma, approximately contemporaneously with Late Triassic–Early Jurassic tholeiitic to transitional, bimodal basalt–andesite to dacite volcanism in the Esfandagheh region insouth SSZ (e.g., Badrzadeh et al. 2011) where the bimodalvolcanism is associated with Cu–Zn–Au and Cu–Zn VMSdeposits (e.g., Badrzadeh et al. 2010). The tholeiitic to transi-tional volcanic rocks are interpreted to have been emplaced ina back-arc setting (e.g., Badrzadeh et al. 2011). However, thepresence of some volcanic rocks of boninitic affinity (e.g.,Badrzadeh 2009) might also indicate one of the followingenvironments for the Esfandagheh region: (1) a fore-arc(e.g., Bedard et al. 1998; Murphy 2006); (2) intra-arc rifting(Crawford et al. 1981); (3) an active arc undergoing back-arcspreading (e.g., Falloon and Crawford 1991; Monzier et al.1993); (4) nascent or primitive arc (Crawford et al. 1989;Murphy 2006).

Both of the Late Triassic Bon-Dono (Noori Khankahdaniet al. 2006) and Siahkuh (Arvin et al. 2007) calc–alkalineintrusions reflect an arc setting which has subsequently beenrifted, resulting in later tholeiitic to transitional volcanism inthe Bavanat and Esfandagheh regions. The emplacementsettings for the Siahkouh and Bon-Dono granitoids, andthe Chahghand and Robat gabbro intrusions (Otrodi 2006)may be similar to that for the Middle Jurassic Chahdozdanintrusion (Sheikholeslami et al. 2008), namely an arc pull-apart basin. This is similar to the model suggested for theSaqqez area in the northwest part of the SSZ (e.g., Azizi et al.2006; Azizi and Moinevaziri 2007, 2009; Azizi and Jahangiri2008). The Saqqez area hosts the Barika gold-rich Zn–Pb–CuVMS deposit (Yarmohammadi 2006; Yarmohammadi et al.2008). A similar suchmodel has previously been proposed forthe Perubar Ba–Pb–Zn VMS deposit in the Peruvian Andes(e.g., Polliand et al. 2005).

Houshmandzadeh and Soheili (1990a) observed that con-tacts between the Toutak and overlying Surian complex aregradational. However, the presence of the basal conglomer-ate of the Surian complex (so-called “Morshedi conglomer-ate”; Ricou 1974; Alric and Virlogeux 1977) indicates thatthis is an erosional unconformity (Oveisi 2001). The con-glomerate is composed of marble and pebble- and gravel-sized clasts of granite. The granitic clasts are sourced fromthe Bon-Dono granitoid intrusion that occurs in the core ofthe Toutak anticline (Figs. 2 and 3). The Early Jurassic ageof the Surian complex is consistent with a Late Triassic

Miner Deposita

emplacement (e.g., Noori Khankahdani et al. 2006) for theBon-Dono intrusion, since pebbles of this intrusion occurin the basal part (unit 1) of the Surian complex (Fig. 4).However, Alizadeh et al. (2010) suggested that the core ofthe Toutak anticline or the Toutak gneiss dome hosts twotypes of granite; an old aplite-granite and a young Bon-Dono (Bendenow) granite-gneiss. Based on 40Ar–39Ardating of biotite and muscovite, these authors suggestedthat the aplite-granite intruded the Toutak complex in thePaleozoic and was deformed at 180 Ma as a result of theinitial subduction of the Neo-Tethyan Oceanic crust,whereas the Bendenow granite-gneiss was emplaced dur-ing continent–continent collision in the Late Cretaceous(i.e., 77 Ma).

The tholeiitic affinity of the Bavanat volcanic rocks mayalso be due to the attenuation of nearby continental crustwhich resulted in rapid rifting and low degrees of crustalcontamination; the Bavanat area is only about 10 km fromthe Zagros Suture (Fig. 2). Similarly, Arvin et al. (2007) notedthe close proximity of the calc–alkaline Siahkuh granitoidintrusion to the Zagros suture, which caused low degrees ofcrustal contamination reflected in tholeiitic and transitionalnature of some of the intrusion samples.

VMS metallogeny

There is growing consensus that VMS deposits preferentiallyform during episodic rifting of oceanic and continental volca-nic arcs, fore-arcs, and in back-arc extensional environments(e.g., Piercey et al. 2001b; Allen et al. 2002; Hart et al. 2004;Hannington et al. 2005; Franklin et al. 2005; Piercey 2011).Approximately 80% of the world’s VMS deposits occur inarc-related successions (Franklin et al. 1999). Many VMSdeposits form on continental margins (e.g., Piercey et al.2008) and in arc or intra-arc settings related to oblique sub-ductions (Franklin et al. 2005). Examples include VMSdeposits in the Skellefte district, Sweden (e.g., Allen et al.1996), Mount Read Volcanic belt, Tasmania, Australia (e.g.,Crawford et al. 1992), and the Perubar deposit which formedwithin an oblique subduction arc-related pull-apart basin (e.g.,Polliand et al. 2005). Furthermore, the Upper Triassic WindyCraggy deposit, the world largest Besshi-type deposit (e.g., Peterand Scott 1999), and the Jurassic Besshi-type deposits (includ-ing the namesake Besshi deposit) in the Sanbagawa metamor-phic belt, Shikoku Island, Japan all were formed in back-arcbasins (e.g., Fox 1984; Slack 1993; Watanabe et al. 1993).

Some authors have previously presented geodynamicmodelsfor the opening, spreading, and closure of the Neo-Tethyanocean basin in southwestern Iran (e.g., Stöcklin 1968; Berberianand King 1981; Sengör 1991; Glennie 2000; Mohajjel et al.2003; Agard et al. 2005; Shahabpour 2005; Ghasemi andTalbot 2006; Sheikholeslami et al. 2008; Omrani et al. 2008;Fazlnia et al. 2009). Our data support the geodynamic model of

Shahabpour (2005) and Sheikholeslami et al. (2008) for theNeyriz–Shahre Babak area. Figure 12 shows the followingrelationships between VMS mineralization and geodynamicevolution of the Bavanat area:

1. Paleozoic opening of the Neo-Tethyan OceanEvidence of an Early Paleozoic episode of crustal

thinning in parts of the SSZ (e.g., Sabzehei 1974; Alricand Virlogeux 1977; Houshmandzadeh and Soheili1990a; Rachid Nejad-Omran et al. 2002) and the Permiancarbonates in the Heneshk area (Houshmandzadeh andSoheili 1990a,b; Shahidi 2001) may reflect an extensionaltectonic regime which is coincident with the separation ofthe central Iranian block from Gondwana (Takin 1972;Berberian and King 1981; Sheikholeslami 2002;Sheikholeslami et al. 2008). Following intracontinen-tal rifting (Fig. 12a) that was initiated during the LatePaleozoic, the accretion of the Neo-Tethys commenced inthe Mid-Permian (Fig. 12b; Sheikholeslami et al. 2008).

2. Early Mesozoic oblique subduction of the Neo-Tethyanoceanic lithosphere

Consumption of the young Tethyan lithosphere thatseparated the northern margin of Gondwana and thesouthern margin of Iran during the Mid-Permian–LateTriassic was initiated by oblique subduction under theIranian plate in the Late Triassic (Fig. 12c; Sheikholeslamiet al. 2008). The Early Cimmerian Orogeny resulted fromcompression related to this subduction and led to thedevelopment of an accretionary prism over an ancientthinned margin, which was subsequently thickened andcrustally shortened, and the Paleozoic rocks metamor-phosed at amphibolite to greenschist facies conditions(Fig. 12c; Sheikholeslami et al. 2008).

3. Formation of arc/pull-apart basins; VMS mineralizationOne of the consequences of the oblique subduc-

tion is the development of post-Early Cimmeriansedimentary and magmatic arc basins in the area(Sheikholeslami et al. 2008; Fig. 12c). Magmas wereemplaced as intrusions in these basins, and lavas eruptedin the Triassic–Jurassic volcano–sedimentary sequences(Sheikholeslami et al. 2008). Collectively, these largegranitoid intrusions in the southern part of the SSZ werelikely emplaced in the Late Triassic, and include theSiahkuh granitoid intrusion 400 km southeast of theBavanat area (Arvin et al. 2007), the Bon-Dono myloniticgranitoid intrusion in the Bavanat area (Noori khankahdaniet al. 2006), and the Chahghand gabbro-pyroxenite andChahdozdan anorthosite-granitoid in the Chahgaz area(~170 km south east of the Bavanat area) that wasemplaced at about 170–173 Ma (Fazlnia et al. 2007). Inthe Chahgaz area, coeval bimodal (dominantly felsic) vol-canism occurred at 174 Ma (Mousivand et al. 2011). Thecalc–alkaline geochemical affinity of the intrusions and

Miner Deposita

volcanic rocks there indicates that they were related to anortheast-directed subduction (e.g., Sheikholeslami 2002;Shahabpour 2005; Arvin et al. 2007; Noori Khankahdaniet al. 2006; Mousivand et al. 2011). Similar Middle

Jurassic calc–alkaline intrusions occur in the northern partof the SSZ at Boroujerd (Boroujerd granitoid complex,dated between 169 and 172 Ma, Ghaderi et al. 2004;Ahmadi Khalaji et al. 2006, 2007) and at Hamadan

Paleozoicmetamorphic complexesNeo-Tethys Bavanat intra-arc basin

Central Iran blockNorth Gondwanan margin(Zagros)/Arabian plate

Neo-Tethys

Central Iran blockNorth Gondwanan margin

Paleozoicmetamorphic complexes

Neo-Tethys

Bavanat metamorphic rocks

Central Iran blockNorth Gondwanan margin(Zagros)/Arabian plate

Nain-Shahre Babak-Baft back-arc basin

Zagros

Neyrizophiolites

Urumieh-Dokhtar magmatic arc

Bavanat metamorphic rocks

Central Iran block

Nain-Shahre Babak-Baft ophiolites

Paleozoicmetamorphic complexes

(a) Paleozoic

(b) Middle Permian-Late Triassic

(c) Late Triassic-Late Jurassic

(d) Late Jurassic-Late Cretaceous

(e) Cenozoic

N

Fig. 12 Relationship between VMS mineralization and tectono–mag-matic evolution of the SSZ in the Bavanat area based on the geodynamicmodel proposed by Sheikholeslami et al. (2008) and Mousivand et al.(2011) for reconstruction of the SW margin of central Iran and northGondwanan margin from the Paleozoic to Cenozoic: a lithospheric thin-ning in middle Paleozoic and development of an intracontinental rift basin;b development of a Neo-Tethys oceanic basin from Middle Permian toLate Triassic; c oblique subduction at Late Triassic; amphibolite to green-schist facies metamorphism of the Paleozoic rocks (Early CimmerianOrogeny); generation of an accretionary prism and formation of a primi-tive magmatic arc as a pull-apart basin over the subduction zone from LateTriassic to Early Jurassic time: emplacement of the Bon-Dono intrusion

(Late Triassic) and Bavanat volcanism (Early Jurassic) and formation ofthe Bavanat VMS deposit within this volcano-sedimentary arc basin; dGreenschist facies metamorphism and associated S1 foliation of the oreand host rocks of the Bavanat deposit; and incipient formation of theNain–Shahre Babak–Baft back-arc basin at Late Jurassic time (Late CimmerianOrogeny); and e Closure of the Neo-Tethys and Nain–Shahre Babak–Baftback-arc basin; development of volcanic arc of Urumieh-Dokhtar zone andcollision and generation of an orogenic prism; formation of S2 foliation inthe Bavanat deposit host rocks; and finally, development of thrusts andstrike-slip faulting in the Bavanat area during the Late Mesozoic andCenozoic (Laramide Orogeny and later phases)

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(Alvand plutonic complex, dated between 153 and166 Ma; Shahbazi et al. 2010). In addition, Mousivand etal. (2011) suggest that in the Middle Jurassic (174.0±1.2 Ma), calc–alkaline bimodal volcanism and spatiallyand temporally associated VMS mineralization formed inthe Chahgaz area, probably coevally with the Chahdozdanintrusion or slightly later than it.

VMS mineralization in the Bavanat area formed withinthese arc pull-apart basins (Fig. 12c) contemporaneouslywith the tholeiitic volcanism and sedimentation in the LateTriassic–Jurassic. High heat flow induced hydrothermalfluid circulation, upflow of fluids along syn-sedimentaryfaults, and precipitation of mineralization at and nearthe paleoseafloor. Several other similar VMS andvolcanic-sediment-hosted stratabound and stratiformdeposits formed in the SSZ during the Triassic–Jurassic, including: (1) the Late Triassic–Early JurassicSargaz Cu–Zn VMS deposit in the Esfandagheh area(Badrzadeh et al. 2011); (2) Late Triassic–Early Jurassicvolcanic- and carbonate-hosted stratiform magnetitedeposits in the Hamedan–Kordestan region (e.g.,Tavakkoli 2004; Motevali 2005); (3) Late Triassic–EarlyJurassic volcanic- and carbonate-hosted stratiform magne-tite mineralization in the Heneshk–Gushti area of theKulikosh region, adjacent to the Bavanat area (KazemiRad et al. 2010); (4) Middle Jurassic stratabound scheelitedeposits hosted in volcano–sedimentary rocks in theShahzand–Bamsar–Nezamabad region (AzizpourMaghvan 2000; Fardindoust 2004; Abdi 2007); and (5)the Middle Jurassic Chahgaz Zn–Pb–Cu VMS depositin the Chahgaz area, southeast of the Bavanat area(Mousivand et al. 2008b, 2011; Mousivand 2010).

4. Late deformation and metamorphism of the VMSmineralization

The Late Cimmerian Orogeny occurred in LateJurassic time (e.g., Berberian and King 1981) and inthe Bavanat area it is evidenced by a penetrativecleavage and folding, both of which are similar inorientation to those of the Early Cimmerian phase(Sheikholeslami et al. 2008). Since there is no evi-dence of a change in stress field between the Earlyand Late Cimmerian phases, Sheikholeslami et al.(2008) suggested that the so-called Late Cimmerianphase corresponds to the continuation, in the samemode of deformation, of the compression that wasresponsible for the Early Cimmerian phase. There-fore, the south-Iranian accretionary prism was stillactive during the Jurassic (and probably the Cretaceous),and the Late Cimmerian Phase is simply part of thecontinuous deformation of the south-Iranian margin(Sheikholeslami et al. 2008). This orogeny, therefore, isresponsible for greenschist facies metamorphism, D1 de-formation, and associated S1 foliation in the Bavanat

VMS deposit and its host rocks (Fig. 12d). The modelof Sheikholeslami et al. (2008) does not account for thepresence of the Nain–Shahre Babak–Baft ophiolites.However, in our modified model, rifting in the nascentNain–Shahre Babak–Baft back-arc basin (e.g., Arvin andRobinson 1994; Shahabpour 2005) was initiated duringLate Jurassic–Cretaceous time as a consequence of theLate Cimmerian Orogeny (Fig. 12d).

5. Closure of the Neo-TethysClosure of the Neo-Tethys and collision of Arabia with

central Iran took place during the Neogene (Berberian andBerberian 1981; Berberian et al. 1982). This collisionresulted in thrusting of the metamorphic rocks over thefolded Zagros unit (the old Arabian margin of Gondwana)(Sheikholeslami et al. 2008) and was followed by dextralstrike-slip faulting described by Talebian and Jackson(2002). Some brittle thrust faults observed in the areamay be related to these later deformational events(Sheikholeslami et al. 2008). The Nain–Shahre Babak–Baft ophiolite mélanges formed due to closure of theback-arc basin, perhaps during the Laramide Orogeny inthe Late Cretaceous (Fig. 12e). All of the Paleozoic andMesozoic metamorphic rocks formed an accretionaryprism. Continued subduction resulted in formation of theEocene–Oligocene Urumieh–Dokhtar magmatic arc with-in the Central Iran zone (Fig. 12e). The Late Cretaceousorogeny likely caused D2 deformation and associated S2foliation and thrusting, faulting and shearing in the Bavanatarea.

Conclusions

Lithogeochemical data show that the mafic volcanic rocks ofthe Surian volcano-sedimentary complex have tholeiitic totransitional affinities, and indicate that magmatism issubduction-related and occurred in an intra-arc rift setting. This,together with LA-ICPMS U–Pb dates of zircons from feld-spathic and quartz feldspathic sandstone host rocks of theBavanat Cu–Zn–Ag VMS deposit that yield a mean age of191±12 Ma indicates that basin subsidence, submarine volca-nism and plutonic activity occurred in close spatial and tempo-ral relationship within the SSZ magmatic arc during the EarlyJurassic. During the Late Triassic to Middle Jurassic, the SWIranian coastal margin underwent rapid Neo-Tethyan obliquesubduction and arc plutonism (evidenced by the Siahkuh,Bon-Dono, Chahghand, Robat and Chahdozdan, BoroujerdandAlvand intrusions) and formation of volcano–sedimentarypull-apart basins in the SW part of the SSZ (including theBavanat area) at an active continental margin. Our studyindicates that Early Jurassic time was a period of formationof VMS mineralization within the Mesozoic magmatic arc ofthe SSZ. Therefore, the Jurassic volcano–sedimentary

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sequences within the SSZ are highly prospective for VMSdeposit exploration.

Acknowledgments The present study constitutes a part of the firstauthor’s Ph.D. thesis at Tarbiat Modares University, Tehran, Iran.ICPMS/ICPAES petrochemical analyses were funded from ioStipendstudent Grant no. VAN08010689 supported by ioGlobal Co., Australia,and Acme Analytical Laboratories, Vancouver, Canada. The seniorauthor gratefully thanks Dave Lawie of ioGlobal and John Gravel ofAcme labs for their support. Costs of U-Pb analyses were defrayed bythe Hugh E. McKinstry Student Research Grant of the Society ofEconomic Geologists (SEG), USA, and the ARC Centre of Excellencein Ore Deposits (CODES), University of Tasmania, Australia. Fieldstudies and lapidary services were funded by the Tarbiat ModaresUniversity Research Grant Council. The first author thanks the Ministryof Science, Research, and Technology of Iran for financial support of hisresearch stay at University of Tasmania, Australia, under supervision ofMike Solomon. Patrick Quilty, University of Tasmania is kindly acknowl-edged for studying the paleontology samples. N. Rachid Nejad-Omran isthanked for his constructive comments on the lithogeochemical data. TheInternational Geoscience Programme (IGCP) project number 502 (Globalstudy of VMS deposits) provided financial support for the senior author toattend the 33rd International Geological Congress at Oslo, and participatein the post congress field trip to the VMS deposits of the Bergslagenregion, Sweden. We thank Wayne Goodfellow, Geological Survey ofCanada for reviewing the manuscript, and B. Lehmann for a carefuleditorial handling. This is GSC contribution number 20110192.

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