Lithos 240–243 (2016) 34–48

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Zircon U–Pb ages and Hf–O isotopic composition of migmatites from theZanjan–Takab complex, NW Iran: Constraints on partial meltingof metasediments

Hadi Shafaii Moghadam a,⁎, Xian-Hua Li a, Robert J. Stern b, Ghasem Ghorbani c, Farzaneh Bakhshizad c

a State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, Chinab Geosciences Dept., University of Texas at Dallas, Richardson TX75083-0688, USAc School of Earth Sciences, Damghan University, Damghan 36716-41167, Iran

⁎ Corresponding author.E-mail address: hadishafaii@mail.iggcas.ac.cn (H.S. Mo

http://dx.doi.org/10.1016/j.lithos.2015.11.0040024-4937/© 2015 Elsevier B.V. All rights reserved.

a b s t r a c t

a r t i c l e i n f o

Article history:Received 29 July 2015Accepted 4 November 2015Available online 10 November 2015

We studymigmatites and other metamorphic rocks in the Zanjan–Takab region of NW Iran and use these resultsto report the first evidence of Oligocene core complex formation in Iran. Four samples of migmatites associatedwith paragneisses, including leucosomes and associated para-amphibolite melanosomes were selected for U–Pbdating andHf–O isotopic analysis. Zircon cores – interpreted as originally detrital zircons–have variable ages thatpeak at ca. 100–110 Ma, but their sedimentation age – indicated by the youngest 206Pb/238U ages – is ca.35–40 Ma. New zircons associated with incipient melting occur as overgrowths around zircon cores and/or asnewly grown grains. Morphologies and internal structures suggest that rim growth and formation of new zirconswere associated with partial melting. All four samples contain zircons with rims that yield 206Pb/238U ages of28–25 Ma, indicating that partial melting occurred in Late Oligocene time. δ18O values for zircon rims vary be-tween 8.2 and 12.3‰, significantly higher than expected formantle inputs (δ18O ~6‰) and consistent with equi-librium with surface materials. Zircon rims yield εHf(t) between 2.2 and 12.4 and two-stage Hf model ages of~448–562Ma, indicating that the region is underlain by Cadomian–Caledonian crust. According to the Hf–O iso-topic values, the main mechanism forming zircon rims was dissolution of pre-existing detrital zircons withreprecipitation of new zircon shortly thereafter. Oligocene ages indicate that partial melting accompanied corecomplex formation in the Zanjan–Takab region. Extension, melting, and core complex formation in south-centralIran are Eocene in age, but younger ages of Oligocene–Miocene in NW Iran and Turkey indicate that extensionwas distributed throughout the region during Cenozoic time.

© 2015 Elsevier B.V. All rights reserved.

Keywords:MigmatiteU–Pb zircon agesHf–O isotopesCore complexZanjan–TakabIran

1. Introduction

Partial melting in the crust is an important feature of orogens(Brown, 2002), and the generation, transport and final fate of crustalmelts are controlled by tectonic forces (Keay et al., 2001; Liu et al.,2014; Whitney et al., 2003; Wu et al., 2007b). The southern margin ofEurasia including Iran is an excellent region where such processes canbe studied. Iran is especially suitable because it has not yet been affectedby “hard” collision with Arabia, in contrast to the “hard-collision”orogens: Alps, Turkey, and India. Unfortunately, the rocks affected byorogenic melting are form deep in the crust. Sometimes these rocksare brought to the surface by regional extension and core-complex for-mation, as was responsible for creating the Zanjan–Takabmetamorphiccomplex of NW Iran, the subject of this study (Fig. 2). In this paper wepresent an integrated study of zircon morphology, U–Pb age and Hf–Oisotope composition from migmatites around Qare–Naz in the Zanjan–

ghadam).

Takab complex and use these results to help understand the relation-ships between partial melting and the evolution of orogens (Brown,2007).

We carried out for zirconU–Pb dating of paragneiss and gneissic am-phibolite of meta-sedimentary protolith, as well as migmatiteleucosomes to unravel the age of partial melting during tectonic uplift,exhumation and core complex formation in central Iran. Zircons can re-tain U–Pb ages andHf isotope signatures through subsequentmetamor-phic and partial melting processes (Corfu, 2007; Liu et al., 2014; Xiaet al., 2014). In-situU–Pb dating of zircons can thus help reveal the com-plex evolution of source rocks and migmatites (Foster et al., 2001; Keayet al., 2001). In addition, zircons have very high Hf contents with verylow 176Lu/177Hf ratios, so their Hf isotopic composition is a sensitive in-dicator of the source of Zr in partially melted rocks (Buick et al., 2008;Foster et al., 2001; Griffin et al., 2002). Moreover, oxygen isotopic com-position of such complex zircons can reveal the involvement of mantle,lower crust and/or upper crust in magma genesis (Kemp et al., 2007;Lackey et al., 2005; Valley et al., 2005). Zircon has high closure temper-ature for O isotopes, resisting resetting even during granulite-facies

35H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

metamorphism (King et al., 1998). Thus, oxygen isotopic composition ofzircon is likely to reflect that of the magma sources. Thus modern anal-yses of zircons can not only provide age information aboutmigmatization-partial melting events, but also constrain the nature ofthe protolith.

2. Geological setting

2.1. General view

The crustal core of Iran and Turkey is dominated by Ediacaran–Cambrian (Cadomian) granitic to tonalitic gneiss and various types ofgranitoids (Abbo et al., 2015; Badr et al., 2013; Balaghi Einalou et al.,2014; Hassanzadeh, 2008; Ramezani and Tucker, 2003; Ustaomeret al., 2012a; Yilmaz Şahin et al., 2014). Cadomian basement is exposedin different parts of Iran (Fig. 1), including the Takab–Zanjan(Hassanzadeh et al., 2008), Soursat (Badr et al., 2013), and Khoy (Aziziet al., 2011)–Salmas (Moghadam et al., 2015a) in NW Iran and inBiarjmand–Torud and near Kashmar in NE Iran (Balaghi Einalou et al.,2014; Moghadam et al., 2015b; Rossetti et al., 2014). Cadomian rocksalso occur in the Alborz (Lahijan pluton) and in the Sanandaj–SirjanZone (Golpayegan–Chadegan region) (Hassanzadeh et al., 2008;Nutman et al., 2014). Cadomian basement is also recognized in theYazd–Tabas blocks, including a pocket of high-grade metamorphicrocks in the Saghand region (Fig. 1), interpreted as a metamorphiccore complex (Ramezani and Tucker, 2003; Verdel et al., 2007).Cadomian exposures are also known in Turkey (Fig. 1), includingBitlis–Poturge massif (Ustaomer et al., 2009, 2012b), Sakarya, Istanbuland Istranca segments (Yilmaz Şahin et al., 2014), Sandikli (Gursuet al., 2004; Kroner and Sengor, 1990), Menderes (Hetzel et al., 1998;

Derik

Black Sea

AzeArmenia

Lesser Caucasus

Mediterranean Sea

ARABIA

Red Sea

B

Pontides

Nigde

30 Eo

40 Eo

Eocene volcanicsEocene granitoidsMesozoic ophiolitesPaleozoic ophiolitesCadomian ophiolitesCadomian granite-gneiss (g)/rhyolite (rgr

?

Scythian Block

Khoy

Z

Ankara

Taurides

BitlismassifPoturge

massifSandikli

Kazdag

Istanbulterranes

Menderesmassif

Salmas

Takab

Fig. 2

Miocene core complex

Oligocene core complex

Eocene core complex

Simav

Fig. 1. Simplified geological map of Iran–Turkey showing the distribution of Eocene magmatismtribution of core complexes is according to Verdel et al. (2007), Stockli et al. (2004), Okay and SWhitney et al. (2007).

Loos and Reischmann, 1999) and Kazdag complexes (Cavazza et al.,2009; Okay and Satir, 2000). Most of these metamorphic complexeshave been exhumed from depth during extension, core complex forma-tion and crustal unroofing related to the younger tectonic movements.

Cadomian basement in central Iran – Saghand – was exhumedduring Eocene extension followed by Early Miocene erosion(Kargaranbafghi et al., 2012). This extension was associated with LateCretaceous and Paleogene igneous activity related to oblique subduc-tion of Neotethys beneath Iran to form the Urumieh–Dokhtarmagmaticbelt (Fig. 1). Eocene extension, magmatism, and core complex forma-tion are also revealed by U–Pb zircon ages of granites and graniticgneisses of the Saghand region (Ramezani and Tucker, 2003).Oligocene–Miocene core complexes are rare in Iran but are known inTurkey. A core complex of Oligocene–Miocene age was reported fromthe Zanjan–Takab region by Stockli et al. (2004). Post-Eocene core com-plexes are also reported from Turkey, in Kazdag (Late Oligocene; Okayand Satir, 2000), Simaw (Miocene; Ring and Collins, 2005), Menderes(Miocene–present; Gessner et al., 2001; Gessner et al., 2013) andNigde (Miocene–present; Whitney and Dilek, 1997; Whitney et al.,2007).

The Zanjan–Takab complex contains a belt of metamorphic rocksthat trend NW–SE, parallel to the larger Zagros orogen. The age of themetamorphic belt is suggested to be Late Neoproterozoic–Early Cambri-an, inferred from U–Pb zircon dating of ca. 548–568 Ma meta-granitesand orthogneisses (Hassanzadeh et al., 2008). The belt is in faultedcontact with Miocene and younger volcano-sedimentary rocks(Fig. 2) – although in some places overlain by unconformity – andmay represent a core complex that formed in response to crustalthickening and orogenic collapse related to initial Arabia–Iran continentalcollision (Gilg et al., 2006; Stockli et al., 2004). (U–Th)/He geochronology

Zarand

EURASIA

Persian Gulf

CaspianSea

rbaijan

Alborz

Kopet Dagh

Sistan suture zone

Makran

Urumieh-Dokhtar Magmatic Belt

itlis-Zagros Suture

Turan Block

50 Eo

60 Eo

30 No

40 No

500 Km

)

Afghanistan

Lahijan

Kahnuj

Birjand

Neyriz

Nain

Mashhad

?anjan

Saghand

Torud

Golpayegan

Paleozoic Suture

Taknar

, Paleozoic–Mesozoic ophiolites, Cadomian basement rocks and core complexes. The dis-atir (2000), Ring and Collins (2005), Gessner et al. (2013), Whitney and Dilek (1997), and

Zaki Kand

Qozlu

Alam Kandi

Qozlu

Almalu

Qazi Kandi

Qare Naz

Poshtuk

Takab

Chah Daka

Moghanlu

Qazi Kandi

Mahneshan

Chomlu

Darvish Rash

Incheh

Khan Qoli

Pichaqchi

Aq Darreh

Gneiss Cadomian granites

Mica schist

Marble

Amphibolite with

Kahar FormationShale, tuff, dolomite

Qaradash FormationRhyolite, tuff, dolomite

Bayandor FormationShale, sandstone

Soltanieh FormationShale, dolomite

Barut FormationShale, dolomite, limestone

Zaigun FormationShale, sandstone

Cenozoic granitoid

Neo

prot

eroz

oic

Late

Ear

lyC

ambr

ian

46º

45'

47º

00'

47º

15'

47º

30'

47º

45'

36º 30'

36º 45'

548.5±13.3

568±44

548.5±13.3

Hassanzadeh et al., 2008

26-28 Mapartial melting

38.4 MaUnpublished data

537±8 Ma

537±8 Ma

543±4 Ma

59±2.7 Ma

56.2±2.1 Ma

Jamshidi et al., 2013

Oligocene melting

26-28 Ma

This study

Neo

prot

eroz

oic

Late

10 km

2000

1000

0 m

1000

2000

SW NE

+ .

A

A B

B

Olig

o-M

ioce

ne

Andesitic lava

Marl, Sandy marllimestone

~491-516 Magneiss

513 Magneiss

~497-568 Maenclave

~23.5-25.9 Mameta-tonalite

28.3 Magneiss38.4 Ma

gneiss

Fig. 2. Simplified geological map of the Zanjan–Takab complex showing sample localities and available U–Pb age data (modified after Takab 1/250,000 map, Geological Survey of Iran).

36 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

(Stockli et al., 2004) as well as Ar–Ar dating (20.07 ± 0.60–20.32 ±0.05Ma) ofmuscovite schist (Gilg et al., 2006) from footwall rocks revealrapid Middle Miocene exhumation of the Zanjan–Takab core complex.

The Zanjan–Takab complex consists of various types of Ediacaran–Cambrian orthogneissic rocks (Hassanzadeh, 2008), metabasite, am-phibolites, calc-silicates, psammitic to pelitic schists andmeta-ultramaf-ic rocks (Fig. 2). Orthogneisses occur mostly around Almalu andMoghanlu villages and show ca. 548–568 Ma ages (Fig. 2; (Hajialioghliet al., 2007; Hassanzadeh et al., 2008; Saki, 2010)). Psammiticparagneiss, with detrital protoliths are intercalated with amphiboliteand schist with rare marbles and calc-silicates and are widespreadaround Qozlu, Alam Kandi and Qare–Naz villages in the centralZanjan–Takab complex. The deposition age of meta-sedimentarygneisses was thought to be Late Neoproterozoic–Early Cambrian(Hassanzadeh et al., 2008), but our results show that some of thesemeta-sedimentswere deposited in Late Eocene time.Metapelitic schistsoccur around the Poshtuk and south of the Qozlu villages (Saki et al.,2012) (Fig. 2). Themineral paragenesis of themetapelites (garnet, stau-rolite, chloritoid, chlorite, muscovite and quartz) reflects peak

metamorphic P–T conditions of 580 °C and ~3–4 kbar, indicating forma-tion ~10–15 km deep in the crust (Saki et al., 2012). Supra-crustal unitsincluding Late Neoproterozoic dolomites, shales, sandstones, tuffs andrhyolites of the Qaradash, Bayandor and Soltanieh Formations are infaulted contact with high T/Pmetamorphic rocks and associated granit-oids (Fig. 2).

High T/P metamorphic rocks exposed in the Zanjan–Takab complexincludemarble, pelitic schist, para-amphibolite and paragneisses,whichshow evidence of partial melting in the form of migmatites (Fig. 3).These migmatites occur among psammitic gneiss and amphibolite andare common NE of Takab, especially near the Qare–Naz village. Howev-er, only a few zircon U–Pb ages from one leucosome sample of thesemigmatites are available (Hajialioghli et al., 2010). Sporadic zircon U–Pb ages on the migmatites reveal Oligocene partial melting and also afew ca. 2960–2770 Ma detrital zircons (Hajialioghli et al., 2010, 2011).Cadomian (ca. 537–543 Ma) granitic gneisses are also common NW ofTakab (Fig. 2). Adakitic, 52–59 Ma old granites crosscut the old gneissicrocks (Badr et al., 2013). The relationship between Oligocenemigmatites and Eocene intrusions is unclear; these may reflect

Fig. 3. Field photographs of theQare–Nazmigmatites. A–E, Schlieren, nebulitic, stromatic, raft- and vein-likemigmatites. F— Late-stage granitic dikes crosscut themigmatitic amphibolites.

37H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

Urumieh–Dokhtar Magmatic Belt activity and core-complex formation.In this study we focus on the Qare–Naz paragneisses, gneissic amphibo-lites and migmatites.

2.2. Geology of the Qare–Naz region

There are six principal lithologies in the study area: amphibolite andgneissic amphibolite; migmatite leucosomes; paragneiss; orthogneiss;clinopyroxene-bearing marble; and granitic dikes. The petrography ofthese six lithologies is described in greater detail in the next section.Migmatites occur mainly around Qare–Naz and Qozlu villages (Fig. 2).Amphibolites and paragneisses with minor metapelitic schists andmigmatitic gneisses are common in the Qare–Naz region (Fig. 3). Am-phibolites can be divided into three types: 1) fine-grained biotite-bear-ing amphibolites; 2) garnet amphibolites and 3) gneissic amphiboliteswith thin (1–2 cm) to thick (20–30 cm) migmatite leucosomes.

Paragneisses and migmatitic gneisses are variably deformed.Orthogneiss is less abundant than paragneiss. Paragneiss is interlayeredwith amphibolite, marble and calc-schist. Leucosomes are millimetersto tens of centimeters thick and vary from weakly to highly foliated(Fig. 3). Leucosome foliation is parallel to that of melanosomes or trun-cates host foliation. Leucosomes have gradational contacts with mela-nosomes and in some cases are folded with axial planes parallel tomelanosomes. Melanosomes are restites composed of massive andgneissic amphibolite, rich in amphibole and plagioclase (and epidote)with minor quartz. Migmatites occur as schlieren, nebulitic, ptygmatic,stromatic, raft- and vein-like structures (Fig. 3). Gneissic graniticleucosomes with large grains of K-feldspar and amphibole also occuramong themigmatites. Some leucosomes are products of in-situ partialmelting but others seem to have invaded from farther away. Unfoliatedgranitoid dikes and small lenses crosscut the amphibolites and gneissicamphibolite. These show sharp contacts with host rocks and seem tohave been injected (Fig. 3F).

38 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

3. Petrography

3.1. Amphibolite and gneissic amphibolite

These foliated rocks consist of hornblende+plagioclase+ quartz±resorbed fine-grained garnet ± titanite ± zircon (Fig. 4A). Biotite-richamphibolites have hornblende + quartz + fine-grained plagioclase.Fine-grained biotite occurs between the amphiboles. One gneissic am-phibolite (QN14-10) was selected for U–Pb zircon dating and in-situHf–O isotope analyses.

3.2. Migmatite leucosomes

These rocks show strong to weak foliation consisting of stretched,coarse-grained quartz withminor kink-bands, serrated margins and re-crystallization, alongwith large grains of K-feldspar+ plagioclase ± bi-otite (2–5% vol.) (Fig. 4B), with slight alteration into chlorite ± titanite.Myrmekitic and graphic texture are common. Augen gnessicleucosomes include highly kinked quartz+ augen K-feldspar± amphi-bole ± epidote with mylonitic foliation. Weakly foliated leucosomescontain plagioclase and K-feldspar (slightly altered into clay min-erals) + highly deformed and stretched quartz grains + amphibole(coarse-grained)± euhedral zircon± titanite (Fig. 4C). Two leucosomesamples (QN14-12 and QN14-13) were selected for U–Pb zircon datingand in-situ Hf–O isotope analyses.

3.3. Paragneiss

Large grains of K-feldspar and plagioclase (4 mm or greater inlength) + biotite + stretched quartz ± clinozoisite ± green amphiboleare the main rock-forming minerals of paragneiss (Fig. 4D). Augenparagneisses are characterized by large grains of K-feldspar and plagio-clase porphyroclasts (av. 2 mm but max. 4 mm) with undulose extinc-tion and deformation twinning. Chloritized fine-grained biotites arewidespread among the quartz ribbons. Mylonitic foliation is common

A

C

QN14-10 Q

QN14-13 Q

amph

amph

plag

K-feld

quartz

quartz

Fig. 4.Microphotographs ofQare–Nazmigmatites. A—Gneissic amphibolites include quartz, plafeldspar, biotite and minor plagioclase. C— Amphibole-rich leucosomes show coarse-grained aK-feldspar, biotite and quartz ribbons in paragneisses.

in augen gneiss. One paragneiss (QN14-14) was selected for U–Pb zir-con dating and in-situ Hf–O isotope analyses.

3.4. Orthogneiss

These rocks are rare and contain small grains of plagioclase and K-feldspar + quartz ribbons + amphibole + allanite + zircon. Most gra-nitic orthogneisses show mylonitic texture, while trondhjemiticgneisses have cataclastic texture and are weakly foliated. Trondhjemiticgneisses have plagioclase with both magmatic and deformationtwining + deformed quartz + chloritized biotite.

3.5. Clinopyroxene-bearing marbles

These rocks contains N70% vol calcite with pale to deep green crys-tals of Cpx + plagioclase + epidote ± fine-grained amphibole ±titanite. They occur as interlayers within psammitic gneiss andamphibolite.

4. Analytical methods

4.1. Mineral major elements

Major-element compositions of plagioclase, amphibole, biotite,clinopyroxene, epidote, and garnet from Qare–Naz metamorphic rockswere determined by JEOL wavelength dispersive electron probe X-raymicro-analyzer (JXA 8100) at the Institute of Geology and Geophysics,Chinese Academy of Sciences (IGG-CAS) (Beijing, China). Acceleratingvoltage, beam current, and beam diameter for the analyses were20 kV, 20 nA, and 3 μm, respectively. Representative mineral composi-tions are used for thermobarometry and to see if there is any zonationin mineral composition; these data are reported in Supplementarydocument 1.

B

D

N14-12

N14-14

amph

biot

biot

K-feld

K-feld

quartz

plag

gioclase, amphibole and titanite crystals. B– Foliated leucosomes contain quartz ribbons, K-mphiboles associated with quartz, K-feldspar and plagioclase. D— Amphibole, plagioclase,

39H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

4.2. Zircon U–Pb dating

Zircon concentrates were separated from ca. 3 kg of rock samplesusing conventional methods. Zircon unknowns together with stan-dards were mounted in epoxy and then polished to section the crys-tals in half for analysis. U–Pb analyses were performed on a CamecaIMS-1280HR SIMS at IGG-CAS using standard operating conditions(7-scan duty cycle, ~8 nA primary O2

− beam, 20 × 30 μm analyticalspot size, mass resolution ~5400). U–Th–Pb ratios and absoluteabundances were 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 isoto-pic compositions were corrected for common Pb using 204Pb. An av-erage Pb of present-day crustal composition (Stacey and Kramers,1975) was used for the common Pb correction assuming that it islargely due to surface contamination introduced during samplepreparation. A long-term uncertainty of 1.5% (1 RSD) for 206Pb/238Umeasurements of the standard zircon was propagated to the un-knowns (Li et al., 2010a). In order to monitor the external uncer-tainties of SIMS U–Pb measurements, analyses of an in-housezircon standard Qinghu were interspersed with unknowns. Eightanalyses yield a weight mean 206Pb/238U age of 160.1 ± 1.9 Ma, iden-tical within errors to the reported age of 159.5 ± 0.2 Ma (Li et al.,2013). Data reduction was carried out using the Isoplot/Ex v. 2.49program (Ludwig, 2003). Zircon U–Th–Pb isotopic data are present-ed in Supplementary document 2.

4.3. Zircon Hf–O isotopes

In situ oxygen-isotope analyses were conducted on zircons thatwere previously dated at IGG-CAS. The Cs+ primary ion beam wasaccelerated to 10 kV, with an intensity of ~2 nA corresponding to abeam diameter of ~10 μm. A normal-incidence electron flood gunwas used to compensate for sample charging during analysis. Nega-tive secondary ions were extracted with a−10 kV potential. Oxygenisotopes were measured using the multi-collection mode on two off-axis Faraday cups. One analysis consists of 16 cycles, with an internalprecision generally better than 0.4‰ (2SE) on 18O/16O. The detailedanalytical procedures are similar to those reported by Li et al.(2010b, 2013). Oxygen isotope ratios are expressed as δ18O,representing deviation of measured 18O/16O values from that of Vi-enna Standard Mean Ocean Water ((18O/16O)VSMOW = 0.0020052)in parts per thousand. The internal precision for one spot analysiswas typically better than 0.4‰ (2SE). The results were thencorrected for instrumental mass fractionation factor (IMF), followingthe equation: δ18OCorrected = δ18OMeasured − IMF. IMF was monitoredin terms of the difference between measured and recommended oxygenisotopic compositions of Penglai zircon standard with a δ18O valueof 5.31‰ (Li et al., 2010b). Zircon O isotopic data are presented inSupplementary document 3.

In situ zircon Lu–Hf isotopic analysis was carried out on aNeptune multi-collector ICP-MS equipped with a Geolas-193 laser-ablation system at IGG-CAS, and the analytical procedures weresimilar to those described by Wu et al. (2006). Lu–Hf isotopicanalyses were obtained on the same zircon grains previouslyanalyzed for U–Pb and O isotopes, with ablation pits of 63 μm diam-eter, ablation time of 26 s, repetition rate of 10 Hz, and laser beamenergy density of 10 J/cm2. During laser ablation analyses, theisobaric interference of 176Lu on 176Hf was negligible due to extreme-ly low 176Lu/177Hf in zircons. Isobaric interference of 176Yb on 176Hfwas corrected using independent mass bias factors for Hf and Yb.During analysis, an isotopic ratio of 176Yb/172Yb = 0.5887 wasapplied (Wu et al., 2006). Measured 176Hf/177Hf ratios were normal-ized to 179Hf/177Hf = 0.7325. The zircon Lu–Hf isotopic results arelisted in Supplementary document 3.

5. Results

5.1. Mineral chemistry

Clinopyroxene (Cpx) in Cpx-marbles is diopside (Fig. 5) with Mg#andAl2O3 from0.54 to 0.58 and from1.5 to 2.2wt.% respectively. Plagio-clase from the granodioritic gneissic and tonalitic leucosomes,paragneissic rocks and Cpx marbles have homogenous compositionsranging from andesine to oligoclase (Fig. 5). Amphibole in myloniticparagneisses and tonalitic leucosomes is ferropargasite (Fig. 5) withAl2O3 and TiO2 ranging from 11.7 to 14.6 and from 0.7 to 1.7 wt.% re-spectively. FeO contents of these amphiboles are high, ranging between17.7 and 22.2 wt.%. Biotite in paragneissic rocks is characterized by lowMg# (~0.4–0.5) and Al2O3 (15.2–16.2 wt.%) but high TiO2 content(3–3.8 wt.%). In comparison, biotites in gneissic leucosomes havelower TiO2 (1.8–2.5 wt.%) but higher Al2O3 (16.7–18.3 wt.%) and Mg#(~0.6). Garnet in granodioritic gneisses is 51.2–53.3% almandine end–member with high pyrope (19.1–20.7%) and spessartine (16.3–17.2%)components.

5.2. Hornblende–plagioclase thermobarometry

We used hornblende–plagioclase thermobarometry to calculate themetamorphic conditions of Qare–Naz migmatites and paragneisses.Hornblende–plagioclase thermometer B (edenite–richterite reaction)(Holland and Blundy, 1994), calibrated by Anderson and Smith (1995)is used here. Temperatures were obtained using the pressure calculatedby the Anderson and Smith (1995) calibration.Metamorphic P–T condi-tions revealed by the hornblende–plagioclase thermobarometer for theQare–Naz migmatites and gneisses are ~766–789 and 750–765 °C,4.9–6.7 and 5.2–7.6 kbar respectively, corresponding to 16–25 kmdeep in the crust and a elevated thermal gradient of ~35°–40 °C/km.The variable pressure estimates for the migmatites may indicate exhu-mation of metamorphic rocks from different crustal levels (Xia et al.,2014).

5.3. Zircon morphology and U–Pb ages

Two leucosome samples (QN14-12 and QN14-13), one gneissic am-phibolite (QN14-10) and one paragneiss (QN14-14) were selected forU–Pb zircon dating and in-situ Hf–O isotope analyses. These resultsare discussed below.

5.3.1. QN14-12 foliated leucosomeThese zircons are prismatic and ~100–200 μm long; some have oval

shapes. Most zircons show core–rim structures on CL images with coresexhibiting an irregular patchy or weakly oscillatory zoning or areunzoned, while less luminescent zircon rims are unzoned (Fig. 6). Zir-cons show resorbed inherited cores, which are luminescent. A few zir-cons lacking cores are considered to have crystallized from ananatectic melt (Wu et al., 2007b, 2008). Zircon cores show a widerange of 206Pb/238U ages from ca. 253 to 2259 Ma and Th/U ratios arevariable (0.06–1.8) (Fig. 9). Such variable zircon ages, Th/U ratios, andmorphologies indicate that the cores were originally detrital. Zirconrims have Th/U between 0.06 and 0.6 (mostly 0.1 to 0.4). Such lowTh/U suggests that the zircon rims may have crystallized from partialmelt of meta-sediments in the middle crust, in an environment withstrong chemical gradients (Vavra et al., 1999). Thirteen of 24 analysesobtained from leucosome zircons yield a weighted mean age of27.78 ± 0.28 Ma (MSWD = 1.3) for zircon rims on the inverse (Terra-Wasserburg) U–Pb concordia plot (Fig. 7), which we interpret as ap-proximating the time of anatexis.

5.3.2. QN14-13 amphibole-rich leucosomeThese zircons are prismatic, ~100–200 μm long. On CL images, they

display intricate internal structures. Nearly all zircons have core–rim

Pigeonite

Augite

Diopside Hedenbergite

En Fs

Wo

Sanidine

Anorthoclase

Oligoclase Andesine Labradorite Bytownite

Ab An

Or

Pyroxene

Feldspar

Amphibole

QN14-1 CPX marbleQN14-5 Tonalitic leucosomeQN14-18 Mylonitic paragneissQN14-3 Granodioritic gneiss

8 7 6 5 4Si (apfu)

0

0.2

0.4

0.6

0.8

1.0

Edenite

Ferroedenite Sadanagaite

Hastingsite

Ferropargasite

Pargasite

Magnesio-Hastingsite

AlVI>Fe+3

AlVI<Fe+3

AlVI<Fe+3

AlVI>Fe+3

Mg/

(Mg+

Fe

)+2

Fig. 5. Pyroxene (Morimoto et al., 1988), feldspar (Deer et al., 1992) and amphibole (Leake et al., 1997) compositions of Qare–Naz migmatites, gneisses and marbles.

40 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

41H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

structures, with inherited cores exhibiting weak oscillatory or planarzoning or are unzoned (Fig. 6). The weak oscillatory zoning of somecores probably represents recrystallization of older magmatic zircons.27 analyses were conducted on rims and cores. Th/U of cores varies be-tween 0.04 and 0.9 (Fig. 9), with common lead corrected ages rangingfrom ca. 45 to 620 Ma (excluding cores of new zircon grains). Ninespots on zircon rims yield a weighted mean age of 26.16 ± 0.33 Ma(MSWD = 0.45) (Fig. 7), which we interpret as approximating thetime of anatexis. Th/U of zircon rims varies between 0.08 and 0.3. Com-mon lead is variable, with f206 ranging from 0.04 to 2.9% for zircon rims.Some newly formed zircon grains from this leucosome have clear core–rim structure,with cores slightly older than rims that probably show theinitiation (ca. 30 Ma) and final crystallization (ca. 26–25 Ma) of partialmelts.

5.3.3. QN14-10 gneissic amphiboliteThese zircons are prismatic, transparent and colorless. In CL images,

most grains contain a core surrounded by a weakly luminescent rim(Fig. 6). The rims exhibit weak oscillatory or in most cases are unzoned,and the cores are interpreted as originally detrital zircons. Twenty-threeSIMS analyses from this sample were obtained. Old, inheritedmagmaticzircons with oscillatory zoning and unzoned young zircons (~26 Ma)are also common. The cores show ages between ca. 35 and 106 Ma;no ages older than 106 Ma have been detected (Fig. 8). Th/U ratios ofthe zircon cores nearly all are higher than 0.1, which we interpret asinherited magmatic zircons (Fig. 9). Core–rim boundaries are irregularbut sharp, suggesting that the rims are overgrowths (Wu et al.,2007b). Zircon rims are weakly zoned or unzoned and have variablyhigh U (245–1328 ppm) and Th (35–74 ppm) contentswith Th/U ratiosbetween 0.05 and 0.2 (Supplementary document 2) (Fig. 9). Commonlead is variable with f206 ranging from 0.19 to 1.8%. Combined206Pb/238U and 207Pb/206Pb (uncorrected for common Pb) define adiscordia on the inverse (Terra-Wasserburg) U–Pb concordia plot withan intercept age of 26.62± 0.55Ma (MSWD=2.3, Fig. 7), whichwe in-terpret as approximating the time of anatexis.

5.3.4. QN14-14 paragneissThese zircons are subhedral to euhedral with short to long prismatic

and even ovoid shapes. Most zircons show core–rim structures on CLimageswith cores exhibiting irregular patchy orweak oscillatory zoning(Fig. 6). The contacts between cores and rims are sharp. Occasionally,the rims penetrate and truncate zonation patterns of the cores, hintingthat the rims formed by recrystallization of older zircons (Corfu et al.,2003; Geisler et al., 2005). 17 analyses were conducted on rims andcores. Four analyses of cores yield variable 206Pb/238U ages of40–325 Ma (Fig. 8). Th/U ratios of cores vary between 0.3 and 0.6, inthe range ofmagmatic zircons (Fig. 9). In contrast, zircon rims have var-iably high (306–1554 ppm) U and Th (24–357 ppm) contents, with Th/U ratios between 0.06 and 0.5 (Fig. 9 and Supplementary document 2).Common lead (f206) ranges from 0.4 to 1.5%. Combined 206Pb/238U and207Pb/206Pb (uncorrected for common Pb) define a discordia on the in-verse (Terra-Wasserburg) U–Pb concordia plot with an intercept ageof 26.9 ± 0.41 Ma (MSWD = 1.9) (Fig. 7), which we interpret as ap-proximating the time of anatexis.

5.4. Zircon Hf–O isotope compositions

εHf(t) for zircon rims of all rock types are positive (N+ 2 except forone analysis) and cluster between two-stage Hf model ages (TDM2) of400 to 700 Ma (Fig. 10A). In contrast, zircon cores have variableεHf(t) values with different TDM2 (~630–2600 Ma). The similar varia-tion in Hf model ages for zircon rims indicates either that these formedthrough the dissolution of pre-existing zircons and re-precipitation ofnew zircon in a closed system (Liu et al., 2014; Wu et al., 2007a) or re-flect invasion of partial melts from similar protoliths.

Twelve Lu–Hf analyses on zircons from gneissic amphibolite (QN14-10) have 176Hf/177Hf varying from 0.28265 to 0.28301 (Supplementarydocument 3). These yield εHf(t) values of−2.1 to 6.2 for cores and 2.2to 8.9 (at t=~26Ma) for zircon rims. Two-stageHfmodel ages for coresvary from577 to 1063Ma and for rims between 442 and 781Ma (with aweighted mean of ~550 Ma for rims) (Fig. 10A).

Thirteen Lu–Hf analyses performed on zircons from the foliatedleucosome (QN14-12) with low 176Hf/177Hf for cores (0.28133 to0.28260) andhigher ratios for rims (0.28292–0.28311) (Supplementarydocument 3). This corresponds to εHf(t) values of−2.6 to 6.1 for zirconcores and 5.7–12.4 for rims (Fig. 10A). Two-stage Hf model ages forcores vary from 1014 to 2638 Ma and between 254 and 604 Ma (withaweightedmean of ~448Ma) for the rims. Nineteen analyses of zirconsin amphibole-rich leucosome (QN14-13) have 176Hf/177Hf of 0.28243 to0.28306 for cores and 0.28288 to 0.28304 for rims, corresponding toεHf(t) values of −11.7 to 10.7 for cores and 4.4–10 for rims (Supple-mentary document 3). These yield two-stage Hf model ages of472–1490 Ma for cores and 379–571 Ma for rims (with a weightedmean of ~540 Ma).

Fifteen analyseswere performed on zircons fromparagneiss (QN14-14). Four analyses on cores exhibit 176Hf/177Hf of 0.28266 to 0.28294(Supplementary document 3). Their εHf(t) values range from 3.1 to7.2, corresponding to two-stage Hf model ages of 549–977 Ma.The other 11 rim analyses show 176Hf/177Hf of 0.28284 to 0.28302.They yield εHf(t) values of 3.1 to 9.2 (except one point withεHf(t) = −2.2) and two-stage Hf model ages of 421–735 Ma (with aweighted mean of ~562 Ma) (Fig. 10A). The large variation in176Hf/177Hf and Hf model ages implies that Hf isotope heterogeneity ofthe protolith zircons was largely preserved during partial melting (Wuet al., 2007a).

δ18O values for migmatite zircons are N6‰ and mostly N9‰(Fig. 11). δ18O for rims varies from 8.2 to 12.3 (average=10.2)whereaszircon cores are slightly lower, 6.4 to 11.3 (average = 9.5). These δ18Ovalues are mostly higher than that of I-type granitic melt (5–8‰; (Liet al., 2007)) and are similar to those expected for δ18O of supracrustalsources (Valley and Lackey, 2005). The median δ18O values formigmatite zircon cores are similar to those of Cadomian granites fromNW Iran (~6.8–10.1; Moghadam et al., 2015b) whereas the zirconrims show slightly higher values.

6. Discussion

Results reported above suggest that an immature volcaniclasticsection of Paleogene age was metamorphosed to amphibolite faciesin the middle crust and experienced incipient local melting toform migmatites and leucosomes. Modest (~10%?) melting ofmetasediments occurred at 26–28 Ma in hot middle crust, 16–25 kmdeep. The mid-crustal package was exhumed to the surface by~20–10 Ma, as shown by overlying sediments of that age. These pointsare further explored in the following discussion. We first outline whatwe infer about the age, nature, and origin of the original supracrustal se-quence.We then address the implications of our results for understand-ing the age and styles of partialmelting to formmigmatites, particularlymeltmobility. Finallywe discuss the tectonic implications of our results.

6.1. Constraints on the age and nature of the protolith succession

In this section we build on our conclusion that zircon cores in Qare–Naz rocks we studied were originally detrital grains. Qare–Nazmigmatites reflect partial melting of a volcano-sedimentary section ofPaleogene age as indicated by the youngest 206Pb/238U ages for zirconcores of ~45 Ma in amphibole-rich leucosome QN14-13, ~35 Ma ingneissic amphibolite QN14-10, and ~40Ma for paragneiss QN14-14. Zir-con cores from paragneiss and amphibolite give ages from ~35 to325 Ma (Fig. 8), although older ages N 620 Ma are also common for zir-con cores frommigmatite leucosomes. Most zircon cores concentrate at

@9@10

25.9 Ma81.7 Ma

8.8

@1 @2

27 Ma

103.2 Ma

-2.1

20 µm

Hf (t)

U-Pb dating

Sample QN14-10, gneissic amphibolite

@17

2539.7 Ma@16

28.0 Ma

10.36

@23@24

27.3 Ma

385.9 Ma5.7

-2.3

Sample QN14-12, foliated leucosome

@4

25.6 Ma 75.8 Ma

7.65.6

@3 @5@6

282 Ma

25.9 Ma

5.0

4.4

Sample QN14-13, amphibole-rich leucosome

Sample QN14-14, paragneiss

@15 @16

26.3 Ma

58.8 Ma

7.2

9.2

@9

25.6 Ma

8.4

4.5

Fig. 6. Representative CL images of zircons from the Qare–Naz migmatites. Analytical spots, measured ages and εHf(t) are marked.

42 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

100–110 Ma (Fig. 8), suggesting that the protoliths of the meta-sedi-mentary rocks were sourced mainly from Late Cretaceous igneousrocks of the early Urumieh–Dokhtar magmatic belt. We infer fromrock associations observed in the field and metamorphic mineral

assemblages that these Late Eocene to Early Oligocene sediments wereoriginally composed of a marine-influenced succession of immaturevolcaniclastic (and minor carbonate) sediments along with lavas. Thissuccession also contained significant arkose, as inferred from the

0.14 0.10 0.000

0.004

0.008

0.012

20

40

60

80

100

120

0

0.016

0.020

0.00 0.02 0.04 0.06 0.08 0.12

206 P

b/23

8 U

207Pb/235U

QN14-10 core & rim

24 25 26 27 28 29

0.042

0.046

0.050

0.054

0.058

0.062

215 225 235 245 255 265

207 P

b/20

6 Pb

238U/206Pb

QN14-10 rim

23 24 25 26 27 28 29 30

0.040

0.044

0.048

0.052

0.056

0.060

210 230 250 270

207 P

b/20

6 Pb

238U/206Pb

QN14-14 rim

Intercept at26.62 ± 0.55MSWD=2.3

Intercept at26.90 ± 0.41MSWD=1.9

Intercept at27.78 ± 0.28MSWD=1.3

Intercept at26.16 ± 0.33MSWD=0.45

25 26 27 28

0.040

0.044

0.048

0.052

0.056

0.060

0.064

0.068

225 235 245 255 265

207 P

b/20

6 Pb

238U/206Pb

QN14-13 rim

0

100

200

300

400

500

600

700

0.00

0.02

0.04

0.06

0.08

0.10

0.12

0.0 0.2 0.4 0.6 0.8 1.0

206 P

b/23

8 U

207Pb/235U

QN14-13 core & rim

238U/206Pb

26 27 28 29 30

0.043

0.045

0.047

0.049

0.051

0.053

0.055

210 220 230 240 250

207 P

b/20

6 Pb

QN14-12 rim

Fig. 7. Concordia and Terra-Wasserberg diagrams for the investigated zircons from the Qare–Naz migmatite leucosomes and melanosomes.

43H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

abundance of Ediacaran–Early Cambrian-aged zircons found as cores.Such a succession is consistent with deposition in a basin as part ofthe (Cretaceous to Eocene) Urumieh–Dokhtar magmatic belt.

6.2. Constraints from zircon rims on age and style of partial melting

The Qare–Naz succession was somehow quickly taken to mid-crust-al depths, perhaps by overthrusting, where it experienced new zircongrowth bracketed between 27.78 ± 0.28 Ma and 26.16 ± 0.33 Ma. Weinterpret that new zircon growth occurred in Late Oligocene time,when crust melted to form migmatites. Both zircon rims and most

newly formed zircons are euhedral, consistent with growth associatedwith cooling of partial melts (Andersson et al., 2002; Keay et al., 2001;Rubatto et al., 2001). Such zircon overgrowths and newly formed zir-cons can form by three processes; 1) dissolution of pre-existing zirconsand reprecipitation of rims and new zircons (Andersson et al., 2002;Flowerdew et al., 2006; Keay et al., 2001; Rubatto et al., 2001); 2) break-down of Zr-bearing phases other than zircon and reprecipitation of rimsand new zircons in a closed system (Flowerdew et al., 2006; Fraser et al.,1997); and 3) crystallization of rims and new zircons from an externallyderived Zr-bearing melt (Andersen et al., 2002; Flowerdew et al., 2006;Foster et al., 2001). The first explanation best fits the case of Qare–Naz

0 100 200 300

1

2

3

4

5

6

0

Abu

ndan

ce

Detrital coreparagneiss &amphibolite

Age (Ma)

Fig. 8.Histogram showing the distribution of ages for zircon cores in gneissic amphibolitesand paragneisses, interpreted as detrital zircon cores.

44 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

leucosome zircons, because Hf isotopic compositions and Th/U of over-growths are indistinguishable from those of older cores (Fig. 10). Theseyield Ediacaran–Cambrian TDM2 ages of 448 to 562Ma, suggesting thatthis Hf largely came from dissolution of Cadomian zircons. This is bestexplained by dissolution of older detrital zircons in silicate melt, limiteddiffusion of dissolved components into a slow-moving felsic melt, andreprecipitation of Zr, SiO2, Hf, U, and Th not far (few cm to m) from

Th/U=1

Th/U=0

.4

Th/U=0

.1

Th/U=0

.01

103

100

101

102

Th

(p

pm

)

101

102

103

104

U (ppm)

Th/U=1

Th/U=0

.4

Th/U=0

.1

Th/U=0

.01

103

100

101

102

Th

(p

pm

)

101

102

103

104

U (ppm)

QN14-10

Core

Rim

Core

Rim

QN14-13

Fig. 9. Th and U concentrations in migmatite zircons di

the site of dissolution. In this interpretation, high εHf values of therims and new zircons can be interpreted as reflecting average Hfcompositions of the anatectic melts of the meta-sedimentary rocks,which will be dominated by contributions from partially-dissolved zir-cons. The εHf value is the “mean crustal residence age” of themetasedimentary rocks, i.e., depending on the timing of their crust–mantle differentiation. Hf isotope compositions of zircon rims thus indi-cate local “closed-system”melting and limited flow of migmatite melts.We see no evidence for addition of mass from a more juvenile source.

Supporting evidence for this interpretation comes from zircon O iso-topes — which in contrast to the situation for Hf–O during migmatiteand new zircon formation could have been derived frommany differentminerals. Mantle-derived magmas (δ18O = 5.3‰ ± 0.3‰; (King et al.,2008; Valley et al., 1998)), lower crustal (δ18O = 7‰), and melts ofsupracrustal materials such as sediments (δ18O = 10‰–30‰; Valleyand Lackey, 2005) have distinct O isotopes. The high δ18O (8.2–12.3‰)for the zircon rims confirm involvement of sediments in themigmatitesfromwhich the zircons crystallized. These supracrustal materials mightbe Eocene-deposited detrital sediments as inferred by youngest detritalzircon cores from paragneiss and para-amphibolite (Fig. 8).

Thermal overprint by pooling magmas at depth might be responsi-ble for re-melting of detrital metamorphic rocks and growth of zirconrims, but we have no direct evidence for this. This process is suggested

Th/U=1

Th/U=0

.4

Th/U=0

.1

Th/U=0

.01

103

100

101

102

Th

(p

pm

)

101

102

103

104

U (ppm)

Th/U=1

Th/U=0

.4

Th/U=0

.1

Th/U=0

.01

103

100

101

102

Th

(p

pm

)

101

102

103

104

U (ppm)

Rim

QN14-14

Core

Core

Rim

QN14-12

scriminated according to their internal structures.

New grains (core)

0 100 200 300 400 500 600 700 800 900 1000

Age (Ma)

0

0.4

0.8

1.2

1.6

2.0

Th/

U

Inherited cores

Rim

-15

-10

-5

0

5

10

15

0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0

Th/U

εHf(

t)

B

C

-15

-10

-5

0

5

10

15

εHf(

t )20

0 200 400 600 800 1000 1200 1400

Age (Ma)

CHUR

“Depleted mantle”0.4 Ga

0.7 Ga

A

Fig. 10. εHf(t) vs. age (A), Th/U ratio vs. age (B) and εHf(t) vs. Th/U (C) for the Qare–Nazmigmatites.

6 7 8 9 10 11 12 130

4

8

12

16

18O (zircon)

Abu

ndan

ce

New grain

Zircon coreZircon rim

Man

tle Sed

imen

ts

Fig. 11. Histogram of zircon O isotopes from the Qare–Naz migmatites.

45H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

by occurrence of high T/low Pmetamorphic events in the region, whichcould reflect high heat flow accompanying Oligocene magmatic intru-sions (Saki et al., 2012).

In addition, themelanosomes preserve an amphibolite facies miner-al assemblage of plagioclase + hornblende + quartz + biotite ± re-sorbed garnet ± titanite; this assemblage is inconsistent with beingresidues after melt-producing reactions in granulite facies (Ge et al.,2013). Although it is possible for melts to be generated by fluid-presentmelting reactions, the abundance of amphibole suggests hydrous fluid-mediated melting, with fluids derived from local metasediments. Thus,the leucocratic veins were probably generated by anatexis ofmetasediments leaving melanosome as melt-depleted residue.

In contrast to the situation for Hf, δ18O values for migmatite zirconrims (8.2–12.3‰) differ from that of zircon cores in the analyzed sam-ples (6.4–11.3‰), which resemble those of Cadomian granites fromNW Iran (~6.8–10.1‰; (Moghadam et al., 2015b)). The different behav-iors of Hf andO isotopes in new rims are becauseO isotopic composition

of the anatectic melt was controlled by breakdown of many differentminerals, some of which (clays, carbonates) exchanged with high δ18Osurface waters in contrast to the situation for Hf in the melt whichwas dominated by contributions from zircon breakdown.

6.3. Tectonic implications

A key aspect of the Cenozoic tectonic evolution of Iran and Turkey isthe formation of metamorphic core complexes, but we have much tolearn about how and when these formed. Our study indicates thatmid-crustal weakening in the Qare–Naz region leading to formation ofa metamorphic core complex began in Late Oligocene time, althoughit took much longer for exhumation to be completed. (U–Th)/He and40Ar–39Ar geochronology indicate Oligocene–Miocene ages for the ex-humation of the Zanjan–Takab core complex (Gilg et al., 2006; Stockliet al., 2004). Stockli et al. (2004) constrained the timing of exhumationwith a range of thermochronologic studies. NearMahneshan (Fig. 2), anextensional low-angle normal fault juxtaposes mylonitized ~25 Magranite in the footwall against ~560Ma granite. Stockli et al. (2004) re-port ~20–22 Ma apatite (U–Th)/He ages from the young granite andinfer rapid exhumation shortly after intrusion. The hanging wall alsocontains a thick sequence of Oligo-Miocene red beds (~30–20 Ma),evaporites, and carbonates (Fig. 2), further constraining core complexexhumation to this time. The Oligo-Miocene clastic rocks are fine-grained and without metamorphic detritus derived from the Zanjan–Takab complex, implying that the Zanjan–Takab complex was still cov-ered at somedepth. Themain detrital constituents of theOligo-Mioceneclastic rocks are Eocene volcanic rocks and younger (Oligocene–Mio-cene) sedimentary clasts. Gneissic clasts indicating final exhumationfirst appeared in Plio-Pleistocene coarse-clastic growth strata, andthese show 3–4 Ma apatite (U–Th)/He ages.

Older core complexes and associated partial melt products(migmatites and anatexites) have also been documented for theSaghand core complex in central Iran (Nadimi, 2007; Ramezani andTucker, 2003; Verdel et al., 2007). U–Pb zircon and 40Ar–39Ar age dataindicate that migmatization, mylonitic deformation, magmatism, andsedimentation all occurred in Middle Eocene time, between ca. 49 and41 Ma (Verdel et al., 2007). Development of the Eocene core complexin central Iran may have been due to lithospheric removal (delamina-tion) accompanied by upwelling of hot asthenosphere and extension,leading to exhumation at a rate of ~0.75–1.3 km/m.y. between ca. 49and 30 Ma (Kargaranbafghi and Neubauer, 2015).

Late Oligocene extension and formation of the Kazdag core complexin western Turkey near the Aegean Sea formed about the same time asthe Zanjan–Takab core complex. Apatite fission-track ages from thiscomplex range from 20.4 ± 2.4 to 10.2 ± 2.5 Ma (Cavazza et al.,2009), showing exhumation to shallow levels around Middle Miocenetime. The relationship of the Kazdag core complex to the Zanjan–

46 H.S. Moghadam et al. / Lithos 240–243 (2016) 34–48

Takab complex is unclear. Whereas the Kazdag complex might reflectrollback of the African slab (Buick, 1991, Gautier and Brun, 1994; Okayand Satir, 2000), the Zanjan–Takab core complex cannot be explainedthis way.

Late Miocene core complexes are also present in NW Turkey and in-clude the Simav (Ring and Collins, 2005), Menderes (Gessner et al.,2001, 2013) and Nigde (Whitney and Dilek, 1997; Whitney et al., 2007)core complexes (Fig. 1). The cause of the extension that formed thesemight differ from those forming core complexes in Iran, which are morelikely related to orogenic collapse. Certainly Urumieh–Dokhtar arc igne-ous activity heated and weakened the crust of this region and mighthave reflected a number of processes including slab break-off and upperplate delamination, perhaps related to magmatic “flare-up” and “lull”(Ducea et al., 2015). Iran was the focus of extensional phases duringEocene–Oligocene, perhaps related to the roll-back of subducted Tethyanslab (Verdel et al., 2007; 2011). Paleogene uplift of Iran has been ascribedto slab break-off and heating of the base of the lithosphere to thin it, amodel based on seismic observations and theoretical considerations(e.g., Hafkenscheid et al., 2006; Kaviani et al., 2007, 2009, 2015; Paulet al., 2006). Extension and continental thinning is the most plausiblemodel for crustal melting within the Iranian block. If we consider thatbreak-off is the main factor, then it is difficult to relate break-off – nowconsidered as ca. 10 Ma (Late Miocene; Agard et al., 2011; Omrani et al.,2008) – to Paleogene core complex exhumation.

Timing of oceanic closure and initiation of collision between Iran andArabia has been suggested as Late Eocene–Oligocene (30±5Ma; Agardet al., 2005; McQuarrie and Van Hinsbergen, 2013), which is consistentwith extension, core complex formation and rapid exhumation of theIranian plateau at this time (Francois et al., 2014). A major exhumationstage in Iran is reported also during ~25–17Ma in central Iran using ap-atite (U–Th/He) and fission track results (Francois et al., 2014).

Formation of Oligocene core complex in NW Iran and partialmeltingevents could showonset of collision and henceforwardwidespread syn-collisional magmatism, indicating delamination and the gravitationalcollapse of the lithospheric keel of Iran, possibly triggered by the upris-ing of the mantle asthenosphere. Partial delamination and/or litho-spheric mantle removal (Hatzfeld and Molnar, 2010) are suggestedhaving major roles during Eocene–Oligocene extension and magmatic“flare-up” in Iran.

7. Conclusions

The following conclusions result from this study:

(a) Cores of zircons in paragneiss and amphibolite melanosomes inthe Zanjan–Takab core complex show variable ages of ca.35–325 Ma (but older ages N625 Ma are also common), indicat-ing that the sedimentary protolithwas deposited about the sametime as subduction-relatedmagmatism of the Urumieh–Dokhtarmagmatic belt.

(b) Zircons in migmatitic leucosomes from the Zanjan–Takab corecomplex show evidence for partial melting of metasediments at28–25 Ma (Oligocene). Core complex exhumation began at thattime and continued for ~20 M.y.

(c) Th/U ratio and Hf isotopic compositions of zircon rims are indis-tinguishable from those of older cores. This indicates local closed-system melting. In contrast, the δ18O values for zircon over-growths (8.2–12.3‰) are generally higher than those of zirconcores (6.4–11.3‰), indicating that rims grew in the presence ofa high δ18O fluid derived from clays and other minerals thatequilibrated with surface materials.

(d) Considering Hf–O isotopic evidence, it appears that rim over-growths reflect dissolution of detrital zircons followed byreprecipitation and new zircon growth.

(e) Anatexis in themiddle crust to formOligocenemigmatites of theZanjan–Takab complex preceded extension and core complex

exhumation and perhaps was associated with mafic under-plating and the addition of hot, mantle-derived material to thebase of continental crust in Oligocene time.

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.lithos.2015.11.004.

Acknowledgments

This work was funded by the Chinese Academy of SciencesPresident's International Fellowship Initiative (PIFI) grant no.2015VEC063 to the first author. We are very grateful to Orhan Karsliand Michael Flowerdew for their constructive reviews of the manu-script. Editorial suggestions by Nelson Eby are appreciated. All logisticalsupport for field studies came from Damghan University. This is UTDGeosciences contribution number 1276.

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