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International Geology Review

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Age and nature of 560–520 Ma calc-alkalinegranitoids of Biarjmand, northeast Iran: insightsinto Cadomian arc magmatism in northernGondwana

Hadi S. Moghadam, Xian–Hua Li, Robert J. Stern, Jose F. Santos, GhasemGhorbani & Mehrdad Pourmohsen

To cite this article: Hadi S. Moghadam, Xian–Hua Li, Robert J. Stern, Jose F. Santos, GhasemGhorbani & Mehrdad Pourmohsen (2016) Age and nature of 560–520 Ma calc-alkaline granitoidsof Biarjmand, northeast Iran: insights into Cadomian arc magmatism in northern Gondwana,International Geology Review, 58:12, 1492-1509, DOI: 10.1080/00206814.2016.1166461

To link to this article: http://dx.doi.org/10.1080/00206814.2016.1166461

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Age and nature of 560–520 Ma calc-alkaline granitoids of Biarjmand, northeastIran: insights into Cadomian arc magmatism in northern GondwanaHadi S. Moghadama*, Xian–Hua Lia, Robert J. Sternb, Jose F. Santosc, Ghasem Ghorbanid

and Mehrdad Pourmohsene

aState Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China;bGeosciences Dept., University of Texas at Dallas, Richardson, TX, USA; cGeobiotec, Departamento de Geociências, Universidade de Aveiro,Aveiro, Portugal; dSchool of Earth Sciences, Damghan University, Damghan, Iran; eDepartment of Geology, Payame Noor University (PNU),Tehran, Iran

ABSTRACTThe Biarjmand granitoids and granitic gneisses in northeast Iran are part of the Torud–Biarjmandmetamorphic complex, where previous zircon U–Pb geochronology show ages of ca. 554–530 Mafor orthogneissic rocks. Our new U–Pb zircon ages confirm a Cadomian age and show that thegranitic gneiss is ~30 million years older (561.3 ± 4.7 Ma) than intruding granitoids(522.3 ± 4.2 Ma; 537.7 ± 4.7 Ma). Cadomian magmatism in Iran was part of an approximately100-million-year-long episode of subduction-related arc and back-arc magmatism, which domi-nated the whole northern Gondwana margin, from Iberia to Turkey and Iran. Major REE and traceelement data show that these granitoids have calc-alkaline signatures. Their zircon O (δ18O = 6.2–8.9‰) and Hf (–7.9 to +5.5; one point with εHf ~ –17.4) as well as bulk rock Nd isotopes (εNd(t) = –3 to –6.2) show that these magmas were generated via mixing of juvenile magmas with an oldercrust and/or melting of middle continental crust. Whole-rock Nd and zircon Hf model ages (1.3–1.6 Ga) suggest that this older continental crust was likely to have been Mesoproterozoic or evenolder. Our results, including variable zircon εHf(t) values, inheritance of old zircons and lack ofevidence for juvenile Cadomian igneous rocks anywhere in Iran, suggest that the geotectonicsetting during late Ediacaran and early Cambrian time was a continental magmatic arc rather thanback-arc for the evolution of northeast Iran Cadomian igneous rocks.

ARTICLE HISTORYReceived 10 January 2016Accepted 12 March 2016

KEYWORDSCadomian; U–Pb zircongeochronology; zircon O–Hfisotopes; active continentalmagmatism; Iran

1. Introduction

Peri-Gondwana terranes in the Alpine–Himalayan oro-genic belt of Europe and Asia are characterized byEdiacaran–Cambrian ‘Cadomian’ arc-type igneous rocks(Fernandez-Suarez et al. 2002, 2003, 2014; Ustaomeret al. 2009, 2011). The Cadomian belt stretches fromIberia through central and southeast Europe intoTurkey and Iran (Shafaii Moghadam et al. 2015; Avigadet al. 2016) and may continue into the Qingtang terraneof Tibet (Wang et al. 2016). Formation of the Cadomianarc followed earlier collisional events that produced thesupercontinent Gondwana (Powell et al. 1993; Dalziel1997) and reflects subduction along its entire northernperiphery to form the several-thousand-kilometre-longCadomian–Avalonian arc (Kroner and Romer 2013;Garfunkel 2015). The Cadomian orogeny centred a mar-ginal tectono-magmatic system of the Western Pacific

style (Drost et al. 2011). Nearly all Cadomian crust riftedaway from Gondwana during Permian–Triassic time,including that of Iran (Berberian and King 1981;Robertson et al. 1991; Garfunkel 2004) and accreted tothe southern flank of Eurasia.

Cadomian igneous and metamorphic rocks comprisemost of the basement of Iran. Late Neoproterozoic sedi-mentary rocks comprise a subordinate part of the base-ment. Exposures are documented from the west(Golpayegan), northwest (Khoy–Salmas, Zanjan–Takab),northeast (Torud–Biarjmand, Taknar), north (Lahijangranites), and central Iran (Saghand and Zarand)(Figure 1). This old crust comprises the core of thePalaeozoic ribbon continent ‘Cimmeria’ (Sengor 1991).Understanding the age, nature, and distribution ofCadomian crust is key to reconstructing the tectonicevolution of Iran. Our understanding of Cimmerian

CONTACT Hadi S. Moghadam hadi.shafaeimoghaddam@mq.edu.au State Key Laboratory of Lithospheric Evolution, Institute of Geology andGeophysics, Chinese Academy of Sciences, Beijing, China

*Present address: ARC Centre of Excellence for Core to Crust Fluid Systems and GEMOC ARC National Key Centre, Department of Earth and Planetary Sciences,Faculty of Science and Engineering, Macquarie University, NSW 2109, Australia. Email: hadi.shafaeimoghaddam@mq.edu.au

Supplemental data for this article can be accessed here.

INTERNATIONAL GEOLOGY REVIEW, 2016VOL. 58, NO. 12, 1492–1509http://dx.doi.org/10.1080/00206814.2016.1166461

© 2016 Informa UK Limited, trading as Taylor & Francis Group

blocks and their Cadomian basement is advancingrapidly but remains incomplete, especially in Turkeyand Iran. Reliable geochronological data for meta-igneous rocks of Cadomian terranes in Iran have onlyrecently begun to be reported (Badr et al. 2013; BalaghiEinalou et al. 2014; Rossetti et al. 2015; ShafaiiMoghadam et al. 2015).

In this study, we advance our understanding of theCadomian basement of Iran by presenting and discuss-ing the significance of new in situ LA-ICP-MS and SIMSzircon U–Pb ages as well as zircon Hf–O and bulk rockSr–Nd isotope data from the Biarjmand granitoids ofnortheast Iran. Our previous study focused on the gneis-sic rocks (ortho-and paragneissic) from north of Sahadand north and south of Delbar. In this study we focus on

poorly known granitic rocks and associated gneissesnorth of Dochah (Figure 2). To provide context for thisstudy, we also compile all zircon U–Pb and Hf–O isotopedata from Iranian Cadomian terranes and use thesecombined data to better understand lateNeoproterozoic arc magma formation and evolution.

2. Geological setting

The Cadomian basement of Iran is characterized byboth felsic igneous rocks – and their metamorphicequivalents – and detrital sedimentary successions.The oldest igneous rocks have Ediacaran–earlyCambrian ages (520–600 Ma) (Ramezani and Tucker

200 km

46 48 54 56 58

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568-637

Figure 1. Simplified geological map of Iran showing the distribution of Eocene magmatism, Mesozoic ophiolites, and lateNeoproterozoic–early Cambrian basement rocks. Numbers show U–Pb zircon ages (the age of the Soursat complex is from Badret al. (2013); Khoy is from Azizi et al. (2011); Golpayegan and Lahijan ages are from Hassanzadeh et al. (2008); south of Taknar is fromRossetti et al. (2015) and other ages are from Moghadam et al. (submitted).

INTERNATIONAL GEOLOGY REVIEW 1493

2003). Cadomian basement was deeply buried byyounger sediments and exhumed during Cenozoicextension (Stockli 2004; Kargaranbafghi et al. 2015;Moghadam et al. 2016). Most Cadomian gneissic expo-sures show evidence of Eocene–Oligocene partial melt-ing to form anatexite-diatexite and melt invasion andare often intruded by Eocene–Miocene plutonic rocks(Ramezani and Tucker 2003). Petrographic and geo-chemical data (bulk rock major and trace elements)show that most Cadomian igneous rocks in Iran arefelsic plutons – with minor rhyolitic extrusive rocks –and on a basis of trace element considerations belongto ‘volcanic arc granite’ suites (VAG) (Badr et al. 2013;Balaghi Einalou et al. 2014; Hosseini et al. 2015; Rossettiet al. 2015; Moghadam et al. 2015a). Moreover, bulk rockNd and zircon Hf isotopic compositions of these rocksindicate that they formed by partial melting of oldercontinental crust or by mixing between juvenile mantle-derived melts and continental crust, although exposuresof this older crust are so far unknown (see Section 6).Interaction between juvenile melt and older crust is alsosuggested by Abbo et al. (2015) for isotopically evolvedCadomian igneous rocks. Remelting of Gondwana-derived sediments deposited in the back-arc basin isalso suggested as an alternative process for theobserved Hf isotope trends and Hf model ages inCadomian igneous rocks of Turkey (Zlatkin et al. 2013).However, remelting of clay-and sand-rich graywackes isnot confirmed by O and Hf isotopic composition ofCadomian zircons (see the next section).

In this article we contribute to our understanding ofCadomian crust formation in Iran by presenting newdata for a representative example from the Torud–Biarjmand area of northeast Iran (Figure 1).

The study area is situated between the AlborzMountains in the north and the Lut–Tabas block to thesouth. Torud–Biarjmand basement rocks are exposedover about 5000 km2; we refer to combined Cadomianbasement exposures here as the Torud–Biarjmand com-plex. To the north-northwest, the Torud–Biarjmand com-plex is surrounded by Eocene magmatic rocks and byPalaeozoic phyllite, schist, and marble. The contactsbetween Paleozoic rocks and Cadomian rocks are faultedand sometimes covered by Quaternary alluvium. TheTorud–Biarjmand complex contains greenschist toamphibolite-facies meta-igneous and meta-sedimentaryrocks and is overlain by un-metamorphosed Jurassic andCretaceous sedimentary rocks. The age of metamorphismof the Biarjmand rocks is unclear, as metamorphic rimsare absent in our zircons. However, Ar–Ar and K–Ar ageson muscovite and biotite grains from Torud orthog-neisses yield ages of ca. 160 and 171 Ma, respectively(Rahmati-Ilkhchi et al. 2010). The region was affected byfour major deformation phases during Cadomian, Mid-Jurassic, Cretaceous and Neogene (Rahmati-Ilkhchi et al.2010). The Torud–Biarjmand complex contains bothpsammitic to volcanogenic metasediments (paragneiss,mica schist, garnet-mica schist and amphibolite) andigneous felsic to mafic plutonic rocks (granitoids, gabbrosand their metamorphic equivalent (orthogneiss)). Less

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55º

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Sahad

ChahJam Dochah

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Cretaceous limestone

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(552.3±3.6 Ma)

(546.8±9.1)

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(529.9±3.8 Ma)

(547.4±6.8 Ma)

(546.4±3.7 Ma) (543.1±9.9 Ma)

(543.0±6.1 Ma)

(541.8±4.7 Ma)

(561.3±4.7 Ma)

+ +

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(537.7±4.7 Ma)

(522.3±4.2 Ma)

This study

Moghadam et al., 2015

Maryam et al., 2014

+ + + + + + =

== = == =

Gabbroic rocks=

=

==

Figure 2. Simplified geological map of the Torud–Biarjmand metamorphic complex. Modified after 1/250,000 maps of Kharturan andTorud, Geological Survey of Iran. For comparison, we show the zircon U–Pb data for the Torud–Biarjmand orthogneissic rocks(Maryam et al. 2014; Moghadam et al. 2015b).

1494 H. S. MOGHADAM ET AL.

deformed granitic bodies intrude more deformedorthogneiss of broadly similar composition. Outcropsnear Sahad in the south and Delbar in the east(Figure 2) have been studied by Maryam et al. (2014)and Moghadam et al. (2015b) (Figure 2). Here, U–Pbzircon dating of orthogneissic rocks yield lateNeoproterozoic–early Cambrian (530–550 Ma) crystalliza-tion ages. Metamorphosed felsic and mafic (amphibolite)dikes and sills yield zircon U–Pb ages of ca. 534–554 Ma(Hosseini et al. 2015). Paragneissic rocks mostly consist ofalternating quartzo–feldspathic and foliated biotite-richlayers. Garnet mica schists are similar to mica-rich para-gneisses. Zircons from Biarjmand mica-schists show206Pb/238U ages of 549–551 Ma (Maryam et al. 2014). Inthe Biarjmand region, fine-grained schists with marbleintercalations are also common. Coarse-grained, non-metamorphosed granitic to granodioritic intrusions areabundant in the Biarjmand region (Figures 2 and 3A). Thegranitoids are either enriched in K-feldspar with minorferromagnesian minerals or contain abundant amphibole

and biotite. They show both coarse-grained proto-gran-ular and/or mylonitic textures. Aplitic and dioritic dikesintrude the Biarjmand granitic pluton (Figure 3B). Garnet-rich granitic mylonites with quartz, feldspar, and biotitestretching lineations are abundant in the Torud–Biarjmand region. Mylonite is especially obvious at thegranite contacts with metasedimentary host rocks.Granitic gneiss is common and characterized by stretch-ing foliation of quartz, amphiboles and micas. Gabbrosare a minor component of the Torud–Biarjmand complexand are petrographically classified as olivine gabbros(olivine + Cpx + plagioclase), mylonitic gabbros (Cpx +plagioclase ± K-feldspar augens) and garnet bearing gab-bros (Cpx + plagioclase + garnet) (Figure 3D). Gabbroicrocks are cross-cut by younger granitic dikes (Figure 3C).

We studied orthogneiss and granitoids in a largeexposure north of Dochah (Figure 2) in terms of majorand trace element compositions, U–Pb geochronology,whole rock Sr–Nd and zircon Hf–O isotopesgeochemistry.

Figure 3. (A) Hillside exposure of Biarjmand Cadomian granites. (B) Centimetric aplitic dikes injected into orthogneiss. (C) Close-up ofgabbros with intruding granitic dikes in the Torud area. (D) Olivine + clinopyroxene and plagioclase in Torud gabbros. (E) and (F)Quartz + K-feldspar + biotite + plagioclase in the Biarjmand granodiorites.

INTERNATIONAL GEOLOGY REVIEW 1495

3. Petrography and mineral assemblages

Because our area contains both metamorphic rocks withigneous and sedimentary protoliths and also non-meta-morphosed igneous rocks, for clarity we modify IUGSQAP modal nomenclature for non-metamorphosed fel-sic igneous rocks (e.g. granite) and igneous prefix plusmetamorphic terms (e.g. granitic gneiss) for our meta-morphic rocks.

Studied granitoids and granitic gneisses show proto-granular to mylonitic textures and contain alkali feld-spar, plagioclase, quartz, and biotite with rare amphi-bole. Samples with more alkali feldspar aremonzogranite (Figure 4), whereas granodiorite, tonaliteand quartz monzodiorite contain more plagioclase andamphibole. Chlorite, epidote, clinozoisite, sericite, tita-nite and iron oxides are present as secondary minerals.Granitic gneisses contain large (perthitic) K-feldspargrains, highly deformed and recrystallized quartz,zoned plagioclase, biotite and allanite. Titanite + clino-zoisite + epidote are secondary minerals (Figures. 3E–F).Aplitic dikes have perthitic K-feldspar, quartz and minorbiotite and plagioclase, whereas dioritic dikes containplagioclase and hornblende with minor quartz, biotiteand K-feldspar.

In the QAP diagram (Streckeisen 1976), the Biarjmandgranitoids and granitic gneisses plot predominantly inthe field of monzogranite with four exceptions: twogranodiorites (BJ11–26 and BJ11–27), a tonalite (BJ11–22) and a quartz monzodiorite (BJ11–12) (Figure 4).

4. Analytical procedures

Major and trace elements were analysed at Actlabs, Canada(www.actlabs.com) using ICP-AES and ICP-MS. The uncer-tainty (1σ) is ~2% for major elements and 5–10% for traceelements (depending on concentration). Major and traceelements are presented in Supplementary Document 1 (seehttp://dx.doi.org/10.1080/00206814.2016.1166461 for sup-plementary documents).

Sr and Nd isotopic composition were determined at theLaboratório de Geologia Isotópica da Universidade deAveiro, Portugal. The selected powdered samples weredissolved with HF/HNO3 in Teflon Parr acid digestionbombs at 200°C. After evaporation of the final solution,the samples were dissolved with 6 N HCl and evaporatedto dry cake, then redissolved for chromatographic separa-tion. The elements for analysis were purified using a con-ventional two-stage ion chromatography technique: (i)separation of Sr and REE elements in ion exchange columnwith AG8 50 W Bio-Rad cation exchange resin; (ii) purifica-tion of Nd from other lanthanide elements in columns withLn resin (ElChroM Technologies) cation exchange resin. Allreagents used in sample preparation were sub-boiling dis-tilled, and pure water was produced by a Milli-Q Element(Millipore) apparatus. Sr was loaded with H3PO4 on a singleTa filament, whereas Nd was loaded with HCl on a Ta sidefilament in a triple filament arrangement. 87Sr/86Sr and143Nd/144Nd isotopic ratios were determined using amulti-collector thermal ionization mass spectrometer(Timsic and Patterson 2014) VG Sector 54. Data wereobtained in dynamic mode with peak measurements at1–2 V for 88Sr and 0.5–1 V for 144Nd. Sr and Nd isotopicratios were corrected for mass fractionation relative to88Sr/86Sr = 0.1194 and 146Nd/144Nd = 0.7219. During thisstudy, the SRM–987 standard gave a mean value of87Sr/86Sr = 0.710255 ± 23 (N = 10; 95% CI) and the jNdi – 1

standard yielded 143Nd/144Nd = 0.5121009 ± 66 (N = 12;95% CI). Nd model ages were calculated according to theprocedure of Depaolo (1981). The one-stage model age(TDM) was calculated assuming Nd isotopic growth of thedepleted mantle reservoir from εNd(t) = 0 at 4.56 Ga to εNd(t) = +10 at present, according to the following equation:

TDM = 1/λ ln{1 + [(143Nd/144Nd)s – 0.51315]/[(-147Sm/144Nd)s – 0.2137]}, where λ = 6.54 × 10–12 a–1 and(143Nd/144Nd)s and (147Sm/144Nd)s are the measured ratiosof samples. Bulk rock Sr–Nd isotopic data are presented inSupplementary Document 2.

Zircon U–Pb dating analysis was carried out using bothLA-ICP-MS and SIMS. LA-ICP-MS analyses were done on anAgilent 7500a ICP-MS equipped with a 193 nm laser at theInstitute of Geology and Geophysics, Chinese Academy ofSciences (IGG–CAS) using the analytical proceduresdescribed by Xie et al. (2008). Zircon 91,500 and GJ–1were used as external calibration standards and the stan-dard silicate glass NIST 610 was used to optimize the

Maryam et al., 2014Moghadam et al., 2015

Q-rich granitoid

AF g

rani

te

syeno-granite

monzo-granite

grano-diorite

tonalite

Q-syeniteQ-monzonite Q-monzodi

syenite monzonite

Q

A P

granitic gneissgranitoid

Figure 4. Quartz–alkali feldspar-plagioclase (QAP) normativeclassification diagram for the Biarjmand gneiss and granitoids.For comparison, we show the normative composition of pre-viously Torud–Biarjmand orthogneiss and granitoids (Maryamet al. 2014; Moghadam et al. 2015b).

1496 H. S. MOGHADAM ET AL.

machine. Common Pb correction followed the methoddescribed by Andersen (2002), since the signal intensityof 204Pb was much lower than the other Pb isotopes andthere is a large isobaric interference from Hg. Raw countrates for 29Si, 204Pb, 206Pb, 207Pb, 208Pb, 232Th and 238Uwere collected for age determination. U, Th and Pb con-centrations were calibrated using 29Si as an internal cali-brate and NIST 610 as reference material. 207Pb/206Pb and206Pb/238U ratios were calculated using the GLITTER pro-gram. Average 207Pb/206Pb, 206Pb/238U and 208Pb/232Thratios in zircon 91,500 obtained during our analytical ses-sion were used to calculate correction factors. These cor-rection factors were then applied to each sample tocorrect for both instrumental mass bias and depth-depen-dent elemental and isotopic fractionation. Ages weredetermined and concordia plots were constructed usingISOPLOT (Ludwig 2003). The zircon U–Pb age results arelisted in Supplementary Document 3.

SIMS U–Pb analyses were performed on a CamecaIMS–1280HR at IGG–CAS using standard operating con-ditions (7-scan duty cycle, ~8 nA primary O2

− beam,20 × 30 μm analytical spot size, mass resolution~5400). U–Th–Pb ratios and absolute abundanceswere determined relative to the standard zirconPlešovice (337 Ma, Sláma et al. (2008)) and M257(U = 840 ppm, Th/U = 0.27, Nasdala et al. (2008)),respectively. Measured Pb isotopic compositions werecorrected for common Pb using 204Pb. Average Pb ofpresent-day crustal composition (Stacey and Kramers1975) was used for the common Pb assuming that thiswas due to surface contamination during sample pre-paration. A long-term uncertainty of 1.5% (1 RSD) for206Pb/238U measurements of the standard zircon waspropagated to the unknowns (Li et al. 2010). In orderto monitor the external uncertainties of SIMS U–Pbmeasurements, analyses of an in-house zircon standardQinghu were interspersed with unknowns. Nine ana-lyses yield a weighted mean 206Pb/238U age of160.7 ± 1.8 Ma, identical within errors to the reportedage of 159.5 ± 0.2 Ma (Li et al. 2013). Weight mean206Pb/238U age of 337.6 ± 2.3 Ma was also obtained forstandard Plešovice. Uncertainties on individual analysesin the data table are reported at a 1σ level; mean agesfor pooled U/Pb and Pb/Pb analyses are quoted with95% confidence interval. Data reduction was carriedout using an Isoplot/Ex v. 2.49 program (Ludwig2003). Zircon U–Th–Pb isotopic data are presented inSupplementary Document 4.

In situ oxygen isotope analyses were conducted onpreviously dated zircons using the same Cameca IMS-1280 SIMS at IGG–CAS. After U–Pb dating, the samplemount was reground to ensure that any oxygenimplanted in the surface from the O2

– beam used for

U–Pb analysis was removed. The Cs+ primary ion beamwas accelerated to 10 kV, with an intensity of ~2 nAcorresponding to a beam size of ~10 μm in diameter. Anormal-incidence electron flood gun was used to com-pensate for sample charging during analysis. Negativesecondary ions were extracted with a –10 kV potential.Oxygen isotopes were measured using the multi-collec-tion mode on two off-axis Faraday cups. One analysisconsists of 16 cycles, with an internal precision generallybetter than 0.4‰ (2SE) on the 18O/16O ratio. Thedetailed analytical procedures are similar to thosereported by Li et al. (2010, 2013). Oxygen isotope ratiosare expressed as delta notation δ18O, representingdeviation of measured 18O/16O values from that of theVienna Standard Mean Ocean Water ((18O/16O)

VSMOW = 0.0020052) in parts per thousand. The internalprecision from one spot analysis is typically better than0.4‰ for 18O/16O ratio (2SE). The results were correctedfor instrumental mass fractionation factor (IMF), follow-ing the equation: δ18OCorrected = δ18OMeasured – IMF. IMFis monitored in terms of the difference between mea-sured and recommended oxygen isotopic compositionsof Penglai zircon standard with a δ18O value of 5.31‰(Li et al. 2010). Zircon O isotopic data are presented inSupplementary Document 5.

In situ zircon Lu–Hf isotopic analysis was carried outon a Neptune multi-collector ICP-MS equipped with aGeolas–193 laser-ablation system at IGG–CAS, and theanalytical procedures were similar to those described byWu et al. (2006). Lu–Hf isotopic analyses were obtainedon the same zircon grains that were previously analysedfor U–Pb and O isotopes, with ablation pits of 63 μm indiameter, ablation time of 26 seconds, repetition rate of10 Hz and laser beam energy density of 10 J/cm2.During laser ablation analyses, the isobaric interferenceof 176Lu on 176Hf was negligible due to extremely low176Lu/177Hf in zircons. Isobaric interference of 176Yb on176Hf is corrected using independent mass bias factorsfor Hf and Yb for correction. During analysis, an isotopicratio of 176Yb/172Yb = 0.5887 was applied (Wu et al.2006). Measured 176Hf/177Hf ratios were normalized to179Hf/177Hf = 0.7325. The measured 176Lu/177Hf ratiosand the 176Lu decay constant of 1.865 × 10−11 yr−1

reported by Scherer et al. (2000) were used to calculateinitial 176Hf/177Hf ratios. The chondritic values of176Hf/177Hf = 0.0332 and 176Lu/177Hf = 0.282772reported by Blichert-Toft and Albarede (1999) wereused for the calculation of εHf values. The present-day176Hf/177Hf = 0.28325 and 176Lu/177Hf = 0.0384 values(Griffin et al. 2004) were used to calculate the depletedmantle model age (TDM). Because zircons were formedin a granitic magma derived from crustal sources, weused the crustal residence model ages (T2DM). The mean

INTERNATIONAL GEOLOGY REVIEW 1497

176Lu/177Hf ratio of 0.0093 for the upper continentalcrust was used to calculate the crustal residencemodel age (T2DM). The zircon Lu–Hf isotopic results arelisted in Supplementary Document 5.

5. Results

5.1. Bulk rock major and trace elementsgeochemistry

Fifteen samples including granitoids (11 samples) andgranitic gneisses (four samples) were selected forwhole rock geochemistry (Supplementary Document1). The rocks are similarly strongly felsic, with SiO2

contents of 74.2–78.1 wt.% along with low MgO abun-dances (0.08–0.8 wt.%), except sample BJ11–12 (quartzmonzodiorite) with SiO2 = 53.9 wt.%, MgO = 5.4 wt.%and high LOI content (1.5 wt.%). K2O content of theserocks is variable but mostly high; 0.5–5.1 wt.%. Allsamples are peraluminous A/CNK (molar Al2O3/(CaO+Na2O+K2O) > 1 (Figure 5A). In the Rb vs. Nb + Ytectonic discriminant diagram (Pearce et al. 1984), thesamples plot as volcanic arc granites (VAGs)(Figure 5B), except sample BJ11–24 (tonalite) withhigher Nb and Y, perhaps reflecting abundant titaniteand amphibole. Biarjmand granitoids are variablyenriched in light rare earth elements (LREEs) relative

to heavy rare earth elements (HREEs) – La(n)/Yb(n)~ 0.7–17.3 – in the chondrite-normalized diagram (Figure 6).Granitoids show conspicuous depletion in Eu, indicat-ing equilibrium with feldspar in residue or cumulate.These rocks have positive Th, U, Rb anomalies andnegative Nb, Ta and Ti anomalies compared toN-mid-oceanic ridge basalts (MORB) (Figure 6), furthersuggesting formation of these melts at a convergentplate margin.

Granitic gneisses share most trace element character-istics with granitic rocks, except that they show traceelement affinities with adakites. They are more enrichedin LREEs relative to HREEs (La(n)/Yb(n) ~ 3.9–20.6) andlack Eu depletion. Granitic gneisses and granitoids areeasily distinguished using Sr/Y vs. Y and La(n)/Yb(n) vs.Yb(n) diagrams (Figure 5C and D). Gneissic rocks havelow Y and Yb contents and overlap with adakites intheir Yb(n) and La(n)/Yb(n) relationships. They also showgreater enrichments in Ba and Sr along with greaterdepletions in Nb–Ta–Ti, but both have trace elementpatterns that are consistent with formation at a conver-gent plate margin.

5.2. Bulk rock Sr–Nd isotopes

Sr–Nd isotopic analyses of the six Biarjmand samples –including three monzogranites, two granodiorites, and

Metaluminous

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Maryam et al., 2014Moghadam et al., 2015

Figure 5. ANK (molar Al2O3/Na2O+K2O) vs. A/CNK (molar Al2O3/CaO+ Na2O+K2O) and Rb vs. Y + Nb diagrams (Pearce et al. 1984) forBiarjmand granitoids and granitic gneisses. For comparison, we show the composition of previously studied Torud–Biarjmandorthogneisses and granitoids (Maryam et al. 2014; Moghadam et al. 2015b).

1498 H. S. MOGHADAM ET AL.

an orthogneiss – are presented in SupplementaryDocument 2 and plotted in Figure 7.

The initial 87Sr/86Sr (t = 520 Myr) for Biarjmand gneissand granitoids ranges between 0.703 and 0.714(Supplementary Document 2), but this scatter mayreflect the high Rb/Sr of these rocks, their low Sr con-tents (20–163 ppm), the choice of 520 Ma as the time offormation for calculating initial Sr isotope composition,and their subsequent deformation and alteration. Therocks yield a whole rock Rb–Sr errorchron age of511 ± 14 Ma (MSWD = 2.0 and initial87Sr/86Sr = 0.70412 ± 0.00077).

Nd isotopes are more resistant to alteration andother concerns for Rb/Sr. The εNd of our samples variesbetween –3 and –6.2. Cadomian granitoids and gneissof the Torud–Biarjmand complex have low 147Sm/144Nd(0.09–0.147) that are suitable for calculating Nd modelages, except for monzogranite sample BJ11–8. Exceptfor BJ11–8, they give Nd model ages (TDM) that rangefrom 1.28 to 1.62 Ga (Supplementary Table 2). On a143Nd/144Nd vs. 87Sr/86Sr plot (Figure 7), the Biarjmandgranitoids are isotopically similar to other Cadomiangneissic rocks from both the Torud–Biarjmand complex(northeast Iran) and the Zanjan–Takabmeta-igneous rocks(northwest Iran). The εNd (t = 520) values of Cadomian

Figure 6. Chondrite-normalized rare earth element patternsand N-MORB-normalized trace element patterns for studiedBiarjmand granitoids and granitic gneisses.

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in NE Iran

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Biarjmand gneissic rocks

Cadomian ophiolite

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Nain-BaftLate K ophiolites

Figure 7. Initial εNd vs. 87Sr/86Sr for the Biarjmand granitoids compared with other Cadomian rocks of Iran. Data for Iran Cadomiangneisses-granites and Kashmar granites are from (Moghadam et al. 2015b) and (Moghadam et al. 2015c), respectively. Data for thePosht-e-Badam Cadomian ophiolitic lavas, Nain–Baft Late Cretaceous ophiolites and Zanjan Cadomian rocks are from Moghadamet al., (submitted).

INTERNATIONAL GEOLOGY REVIEW 1499

granites and gneissic rocks of the Torud–Biarjmand com-plex (−2.2 to −5.8) are also similar to Cadomian granitesfrom the Bitlis massif (Turkey) (εNd = − 1.2 to −2.9)(Ustaomer et al. 2009). Cadomian gneissic rocks are iso-topically different from Cadomian ophiolitic lavas fromCentral Iran, which show highly radiogenic Nd values(Figure 7). They are also different from Cretaceous ophio-lites and Eocene granitic rocks (Figure 7).

5.3. Zircon U–Pb geochronology

Three samples were dated: one granitic gneiss, grano-diorite and monzogranite. These results are discussedbelow.

5.3.1. Sample BJ11–8 (monzogranite)Zircons are euhedral to subhedral with long to short pris-matic and even stubby shapes (av. ~70–150 μm). Zirconsfrom this sample show evidence of inheritance in theircores. In cathodoluminescence images both new zirconsand inherited cores have magmatic oscillatory zoning.Some inherited cores are unzoned. Results for 21 LA-ICP-MS analyses are listed in Supplementary Document 3.Zircons are characterized by moderately high Th/U ~0.2–0.6. These characteristics are similar to those of magmaticzircons (Corfu et al. 2003; Corfu 2004). The 207Pb/206Pb vs.238U/206Pb (uncorrected for common 206Pb), define a dis-cordia on the inverse (Terra–Wasserburg) U–Pb concordiaplot (Figure 8) with an intercept age of 519.9 ± 4.4 Ma(MSWD = 2.1). This is interpreted as the age of granitecrystallization. Twelve out of 21 analyses give 207Pb/206Pbages older than 520Ma, as old as 930Ma are also commonand reflect inherited cores.

5.3.2. Sample BJ11–28 (granodiorite)Zircons from sample BJ11–28 are long to short pris-matic, ~100–200 μm long. In CI images, they showmagmatic concentric zoning. Core–rim structure is com-mon. The cores show either complex zoning or areunzoned. Some zircon grains have dark discontinuousunzoned core. Data for 21 SIMS analyses on zircons arelisted in Supplementary Document 4. These show mod-erately high Th/U ratios of 0.2–0.9, typical for magmaticzircons (Corfu et al. 2003; Corfu 2004). Except for threespots (points 10, 21 and 22) that show Pb loss, 13analyses on zircons from coherent groups that are con-cordant and are close to the lower intercept with an ageof 537.7 ± 4.7 Ma (MSWD = 1.12) (Figure 9). 207Pb/206Pbages as old as 988 Ma reflect inherited cores.

5.3.3. Sample BJ11–19 (granitic gneiss)Zircons from the Biarjmand granitic gneiss are short tolong prismatic with variable length from 50 to >200 μmand aspect ratios of 1:2 to >1:4. Stubby-form zircons arealso common. Zircons show oscillatory zoning, consis-tent with their magmatic origin (Corfu et al. 2003; Corfu2004). Rare inherited cores are present, and show over-growth by wide magmatic rims. Data for 24 SIMS ana-lyses of zircons are listed in Supplementary Document3. These display high Th/U ratios of 0.4–0.9 for bothcores and rims. In Terra–Wasserburg diagram (uncor-rected for common 206Pb), zircons (excluding minorinherited cores with 207Pb/206Pb ages ~572, 589 and627 Ma) are concordant with lower intercept at561.3 ± 4.7 Ma (Figure 9).

5.4. Zircon Hf and O isotopes

δ18O values for granitoid and granitic gneiss zircons (gran-odiorite BJ11–28 and BJ11–19) are moderately high andvary between 6.2‰ and 8.9‰. These δ18O values aresimilar to that of I-type granitic melt (5–8‰; (Li et al.2007) and are lower than expected for supracrustal sources(Valley and Lackey 2005). The median δ18O values(7.5 ± 0.2‰) for Biarjmand zircons are similar toCadomian granites from northwest Iran (~6.8–10.1;Moghadam et al. 2015d), although the latter may showmore addition of crust into magma.

Thirty-six Lu–Hf analyses were obtained from twosamples of Cadomian (520–560 Ma) zircons(Supplementary Document 5), one granitic gneiss(BJ11–19) and one granodiorite (BJ11–28). The initialHf isotope ratios were calculated using measured206Pb/238U ages of each analysis point. The Cadomianzircons (excluding inherited cores) show variable176Hf/177Hf and hence εHf(t) values (Figure 10). TheirεHf(t) varies from ~ –7.9 to +5.5 (average ~ –2.5)

500 510 520 530 540 550

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505

515

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535

545

BJ11-8monzogranite

Figure 8. LA-ICP-MS zircon U–Pb data for zircons from sampleBJ11–8 (monzogranite).

1500 H. S. MOGHADAM ET AL.

excluding one point with εHf ~ –17.4 (point BJ11–28–17). These 176Hf/177Hf values for Cadomian zircons cor-respond to a mean T2DM age of ~ 1.4 Ga.

6. Discussion

In the following three subsections, our results providenew insights for three questions: (1) what was the tim-ing of Cadomian magmatism in northeast Iran?; (2) whatdo Nd bulk rock and zircon Hf and O isotopic datareveal about crustal evolution in this region?; and (3)What are the tectonic implications of these data?

6.1. Timing of Cadomian magmatism in northeastIran

The Cadomian orogen as preserved in rocks of southernEurasia encompassed magmatic, tectonic and meta-morphic events and sedimentary responses that spannedthe period from mid-Neoproterozoic (~750 Ma,Cryogenian–Ediacaran) to the early Cambrian (~540–

520 Ma) that happened along the periphery of the super-continent Gondwana (peri-Gondwana) (Linnemann et al.2004, 2014) (Linnemann et al. 2008). However, it seemsthat Cadomian events encompassed a shorter time in east-ern peri-Gondwana (Turkey and Iran), where it appears tohave lasted from Ediacaran (~600 Ma) to early-middleCambrian (~500 Ma) time, with the most intense activityoccurring around 550 Ma (Figure 12A) (Ustaomer et al.2009; Badr et al. 2013; Balaghi Einalou et al. 2014; Abboet al. 2015; Rossetti et al. 2015; Avigad et al. 2016). The U–Pbages presented in this study as well as previous publisheddata (Maryam et al. 2014; Moghadam et al. 2015b) on theTorud–Biarjmand region support the conclusion thatCadomian magmatism lasted a few tens of millions ofyears (~40 million years) in northeast Iran.

Our compiled data (Figure 1, Figure 12A) shows that theCadomianmagmatism in Iran lasted from~600 to <500Ma.The only known exceptions are the Chadegan rocks fromwest Iran, which show U–Pb zircon ages of ~637 Ma(Nutman et al. 2014). Interestingly, zircons with older agesthan >600 Ma occur as inherited cores in Iran Cadomianrocks. It seems that igneous rocks that are older than

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Figure 9. SIMS zircon U–Pb data for zircons from samples BJ11–28 (granodiorite) and BJ11–19 (granitic gneiss).

INTERNATIONAL GEOLOGY REVIEW 1501

>600 Ma should be present somewhere in Iran and Turkey,but no one has found these yet. However, as the Cadomiangranites intruded metasediments with a variety ofNeoproterozoic and older detrital zircons, it may provethat the older zircons are inherited from the metasedi-ments. On the other hand, this suspicion is challenged bythe distinctive zircon Hf and bulk rock Nd model ages ofIran and Turkey Cadomian rocks, which show involvementof Mesoproterozoic igneous crust during Cadomianmagma generation (see also Zlatkin et al. 2013; Abboet al. 2015; Avigad et al. 2016). Extensive melt–sedimentmixing and partial melting of clay-rich sediments canincrease O isotope but decrease Hf isotope values (Zhenget al. 2007; Zhao et al. 2013). Furthermore, detrital zirconsfrom the Torud–Biarjmand mica-schists show a concentra-tion of ages around 551 ± 5.1 to 549 ± 5.1 Ma, confirming alocal source (Maryam et al. 2014).

6.2. Isotope perspective on the crustal evolution ofnortheast Iran

Geochemical evidence discussed in the next sectionindicates that the rocks we studied formed in an arc.

The involvement of older crust is strongly suggested bythe distinctive zircon Hf model ages (1.0–2.2 Ga, mostly~1.4 Ga) and bulk rock Nd model ages (1.3–1.6 Ga) ofBiarjmand orthogneiss and granitoids. This could indi-cate anatexis of Mesoproterozoic crust or mixing ofjuvenile Cadomian melts with Palaeoproterozoic orArchaean crust. The lack of known crust of pre-Cadomian crust in Iran and dearth of inherited zirconsof either age makes this question unresolvable at pre-sent. Nevertheless, the isotopic information indicatesthat the Cadomian arc in Iran was a continental arc,similar to the modern Andes. Moderately high zirconO (6.2–8.9‰) but variably low εHf values (–7.9 to +5.5)and bulk rock εNd values (–2.2 to –5.8) are also consis-tent with a continental arc setting (Figure 10). On theδ18O against εHf(t) diagram (Figure 11), Torud–Biarjmand complex rocks differ from Cadomian juvenilegranites and gabbros from east Iran and also fromSalmas (northwest Iran) rocks, which show high δ18Oand constant εHf(t) values.

U–Pb dates and Hf isotope compositions indicate animportant episode of arc granitoid production and recy-cling of older continental crust during Cadomian timealong the whole of northern Gondwana. Regional con-siderations suggest a geotectonic setting starting withan Andean-type marginal continental arc that changedto a Western Pacific style continental arc and a back-arcbasin developed on a thinned and stretched continentalcrust (Linnemann et al. 2014).

Cadomian crustal evolution (ca. 600–500 Ma) of Iranwas dominated by recycling of continental crust assuggested by the abundance of granitoids with zirconshaving negative or low values of εHf(t) (Figures 10 and12). During the ~100-million-year long Cadomian mag-matic arc activity, juvenile arc magmas interacted witholder continental crust or sediments derived from these.None of the εHf(t) values for the zircons sits on thedepleted mantle array, suggesting that these havebeen affected by crustal contamination. However, ourcompiled Hf isotope data (Figures 10 and 12C) revealthat Cadomian rocks from the Taknar and Zanjan–Takabcomplexes (east and northwest Iran, respectively) aremore juvenile than other Cadomian rocks of Iran, sug-gesting the presence of juvenile Neoproterozoic arccrust (new crust in Figure 10) in eastern Iran.Cadomian (630–540 Ma) juvenile rocks are inferred else-where in West Gondwana (Anti–Atlas belt, Morocco)from detrital zircon survey (Avigad et al. 2012). Also,Cadomian ophiolitic basalts with high bulk rock Nd iso-topic compositions are known from central Iran (Posht-e-Badam = Saghand) (Figure 7). Crustal contaminationcould be related to the crustal thickening along theCadomian arc. In this scenario, long-lasting magmatic

TD

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Biarjmand orthogneissBiarjmand granites

New crust

Figure 10. Diagram showing zircon εHf(t) values against U–Pbages. The depleted mantle array was calculated using data formodern mid-oceanic ridge basalts (Chauvel and Blichert-Toft2001). The curve for new continental crust evolution is fromDhuime et al. (2011). Zircon Hf isotope data for Salmas, Zanjan–Takab and northeast Iran Cadomian rocks are from Moghadamet al., (submitted); data for Torud–Biarjmand and Ghushchi orthog-neisses are from Moghadam et al. (2015b, 2015d).

1502 H. S. MOGHADAM ET AL.

pulses (approximately 100 million years) led to crustalthickening, therefore enhancing MASH processes(Melting, Assimilation, Storage and Homogenization).This crustal thickening can be responsible for adakiticsignature of the approximately 560-million-year-oldgranitic gneisses. However, isotopic signatures of graniticgneisses and granitoids are quite similar (althoughorthogneiss has lower mean εHf(t) values), and the La

(n)/Yb(n) ratio for gneissic rocks is not too high. This mayshow that the adakitic signature could be inherited bymore crustal addition to the parent melt, or even byamphibole fractionation in the source area.

6.3. Tectonic implications

The geochemical data summarized in Figures 5 and 6suggest that Cadomian rocks from the Torud–Biarjmandcomplex of northeast Iran formed in a magmatic arcabove a subduction zone. The isotopic data discussed inthe previous section indicate that this was a continentalarc. This magmatic system evolved quickly and we seeevidence in the Biarjmand gneiss that early-formed arc

Figure 11. Plot of δ18O versus εHf(t) for Cadomian rocks of Iran(modified after Kemp et al. 2007), showing curves correspond-ing to magma evolution by crustal assimilation–crystallization(AFC). Hfpm/Hfc = ratio of Hf concentration in the parentalmagma (pm) to Hf concentration in crustal (c) rocks. ZirconHf–O isotope data for Iran Cadomian rocks, Eocene Kurdistanjuvenile gabbros and Nain ophiolite are from Moghadam et al.(submitted). The Kurdistan and Nain ophiolite zircons presentmantle-derived juvenile melts and are for comparison to revealcrustal input in Cadomian rocks.

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Figure 12. Histograms of zircon U–Pb ages (A), T2DM Hf model ages (B), εHf(t) (C) and zircon δ18O (D) isotopes for all Cadomian rocksof Iran. Zircon U–Pb ages and Hf–O isotopes data for Iran Cadomian rocks are from Moghadam et al., (submitted) excluding those forTorud–Biarjmand, which are from (Maryam et al. 2014; Moghadam et al. 2015b) .

INTERNATIONAL GEOLOGY REVIEW 1503

rocks (represented by 561 ± 4 Ma gneiss) weredeformed and uplifted before late-formed arc magmas(represented by 538 ± 5 and 522 ± 4 Ma granitoids)intruded (Figure 13). This observation is consistent withconclusions from farther west, where the differentpulses of magmas intrude the older gneissic rocks andmetasediments. Following the assembly of theGondwana supercontinent, which started in theCryogenian (Collins and Pisarevsky 2005), Cadomian

arc-type magmatism formed thickened continental arcin the peri-Gondwana region including Iran, Turkey andBohemian, Armorican, and Iberian Massifs (Linnemannet al. 2008, 2014; Ustaomer et al. 2009; Pereira et al.2011; Yilmaz Şahin et al. 2014; Orejana et al. 2015;Moghadam et al. 2015b). Cadomian arc magmatismevolved to a marginal Western Pacific style orogenicsystem at the end of the Ediacaran (e.g. Drost et al.2004, 2011; Linnemann et al. 2008). In fact there are

Figure 13. Summary of the tectonic evolution of Cadomian terranes within Iran (and North Gondwana) is shown in a series of panels(see text for explanations).

1504 H. S. MOGHADAM ET AL.

two alternatives for the evolution of the Cadomian arcrocks. For the first scenario, Linnemann et al. (2014)considered an arc–back-arc basin model for the evolu-tion of peri-Gondwana during Ediacaran–Cadomiantime. In this scenario, the back-arc basin was flooredby thinned continental crust and flanked by a magmaticarc to the north and by intact Gondwana to the south(Figure 13). According to Linnemann et al. (2014), theCadomian back-arc spreading started at ca. 570 Ma andproduced pillowed E–MORB-like lavas, andesites, andcalc-alkaline meta-basalts, all associated with blackcherts (Buschmann et al. 2001). Nance and Murphy(1994) proposed a model in which oblique subductionbeneath the arc led to strike–slip shearing and back-arcbasin opening.

The distribution of Cadomian ophiolites is taken asevidence of one or more back-arc basins (Kounov et al.2012; Von Raumer et al. 2015). Suspected Cadomianophiolites in Iran are identified in two regions, Posht-e-Badam (Saghand, central Iran) and Zanjan–Takab (north-west Iran; Figure 1). There are not yet any detailed studiesor U–Pb zircon ages for these ophiolites but their asso-ciation with known Cadomian metamorphic rocks istaken as evidence that they are also Cadomian. InPosht-e-Badam, suspected Cadomian ophiolitic rocksinclude metamorphosed peridotites, MORB-like metavol-canic rocks, marble, and rodingites. The whole rock geo-chemical compositions of these lavas are similar to thosefound in back-arc basins (Moghadam et al., submitted)and isotopically these lavas are juvenile (Figure 7). InZanjan–Takab, the ophiolite contains metamorphosedperidotites and metamorphosed magmatic rocks (tremo-lite-actinolite schists). Further studies are needed todetermine the relationship between these inferredCadomian ophiolites and Cadomian continental arcigneous rocks like those of Biarjmand. In the secondscenario, which was recently proposed by Avigad et al.(2016), the back-arc basin was opened as early as 570 Ma,as graywacke sediments were deposited, starting ca.570 Ma. This is confirmed by the youngest detrital zirconsin Eastern Mediterranean (Tauride block) Neoproterozoicsediments (Abbo et al. 2015). Avigad et al. (2016) consid-ered an approximately 50-million-year time interval forthe cycle of back-arc sedimentation, metamorphism andlater granitic magmatism, which were responsible for theconstruction of the Tauride block. They concluded thatTauride (and other Cadomian terranes) were finallyaccreted to Afro-Arabia during the late Ediacaran–earlyCambrian. This model best explains the existence ofyounger granitic intrusions (~540 Ma) intruding theTauride metasediments which were deposited in aback-arc basin subsequent to ~580 Ma. However, thetrace of the older Cadomian arc north of the back-arc

basin isn’t clear. Have the older Cadomian rocks beentotally uplifted and eroded and therefore supplied theback-arc sediments? Furthermore, the presence of ca.580 Ma granulites shows crustal thickening beneath anevolved arc during that time (Koralay 2015), whereasexhumation of eclogites at 535 Ma shows that subduc-tion beneath north Gondwana was active during theearly Cambrian (Candan et al. 2015).

The presence of gabbros with ages of ~556–585 Main northeast Iran (Moghadam et al., submitted), whichdo not intrude metasediments, are another matter ofdebate. These are clearly older than the sedimentationage of ca. 570 Ma. We suggest that the oldest metasedi-ments were deposited in the hinterland near the Afro-Arabian margin, and the back-arc basins opened in thelatest Ediacaran. These back-arc basins continued toopen to form the Rheic Ocean beginning ~530 Ma,allowing the western arc fragments (including Iberia)to migrate north where they accreted to the southernmargin of Laurussia.

7. Conclusions

Bulk rock geochemistry reveals that granitoids and gneissof the Torud–Biarjmand complex in northeast Iran are arc-related calc-alkaline rocks that formed in a continental arc.U–Pb zircon geochronology shows that these rocks haveages of ca. 560–520 Ma and are part of Cadomian igneousactivity in peri-Gondwana. Zircon Hf–O and bulk rock Ndisotope compositions indicate the involvement ofMesoproterozoic or older continental crust. Our geochem-ical and geochronological data are consistent with lateEdiacaran–early Cambrian arc magmatism along theactive continental margin of north Gondwana, and wesuggest that evidence of a Cadomian back-arc basin mayexist in central and northwest Iran.

Acknowledgements

We are very grateful to G. Topuz and an anonymous refereefor their constructive review of the manuscript.

Disclosure statement

No potential conflict of interest was reported by the authors.

Funding

This work was funded by the Chinese Academy of SciencesPresident’s International Fellowship Initiative (PIFI) [grantnumber 2015VEC063] to the first author.

INTERNATIONAL GEOLOGY REVIEW 1505

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