Zircon trace element geochemical constraints on the evolution of the Ediacaran (600–614Ma) post-collisional Dokhan Volcanics and Younger Granites of SE Sinai, NE Arabian–Nubian

Chemical Geology 360–361 (2013) 54–73

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Chemical Geology

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Zircon trace element geochemical constraints on the evolution of theEdiacaran (600–614Ma) post-collisional Dokhan Volcanics and YoungerGranites of SE Sinai, NE Arabian–Nubian Shield

Mohammed Z. El-Bialy a,⁎, Kamal A. Ali b,c

a Geology Department, Faculty of Science, Port Said University, Port Said 42522, Egyptb Department of Mineral Resources and Rocks, Faculty of Earth Sciences, King Abdulaziz University, P.O. Box 80206, Jeddah 21589, Saudi Arabiac Geosciences Department, University of Texas at Dallas, 800W Campbell Rd., Richardson, TX 75080, USA

⁎ Corresponding author. Tel.: +20 1223282650; fax: +E-mail address: [email protected] (M.Z. El-Bialy).

0009-2541/$ – see front matter © 2013 Elsevier B.V. All rihttp://dx.doi.org/10.1016/j.chemgeo.2013.10.009

a b s t r a c t
a r t i c l e i n f o
Article history:Received 15 May 2013Received in revised form 21 September 2013Accepted 8 October 2013Available online 18 October 2013

Editor: K. Mezger

Keywords:ZirconPost-collisional magmatismDokhan volcanicsYounger granitesArabian–Nubian shieldSinai

The post-collisional stage (620–590Ma) of the Egyptian part of the Arabian–Nubian Shield was characterized bythe eruption and emplacement of the Dokhan Volcanics and Younger Granites. This study presents LA-ICP-MSanalyses of the trace element abundances of zircons separated from Ediacaran (600–614Ma.) Dokhan Volcanic(rhyolite) and Younger Granite (syenogranite) samples from SE Sinai. Whole-rock geochemical data fromthese two rock units indicate both are peraluminous, calc-alkaline and with A-type characteristics and indicatetheir magmas were generated by partial melting of continental crust or underplated crust in a post-collisionalregime. Zircons separated from the Dokhan sample contain higher abundances of Hf, ΣREE, Pb, Th and U thanthose from the Younger Granite sample, which suggest a higher degree of magmatic evolution in the former.The chondrite-normalizedREE patterns of zircons from the two rock types are characterized byHREE enrichmentrelative to LREE andMREEwith positive Ce and negative Eu anomalies, typical of magmatic zircons. Compared tounaltered magmatic zircons, most of studied zircons display an evident LREE overabundance, whereas the vastmajority of the analyzed zircons have Th/U ratios≥ 0.5 common in igneous zircons. The Ce and Eu anomaliesof both zircon populations indicate that crystallization of the zircon grains from the Dokhan Volcanics in morevariable and higher oxygen fugacity conditions than those of the Younger Granites. The relation between theU/Yb ratio vs. Y and Hf contents suggest crystallization of the analyzed zircons in continental crust. Theapplication of the Ti-in-zircon thermometer returns high temperatures of 790–986°C and 781–934°C for zirconsextracted from the Dokhan Volcanic and Younger Granite samples, respectively, indicative of deep crustalmelting within the lower crust. The zircon saturation temperatures for both rock types are slightly lower thanthe temperature of zircon crystallization (≈40–55 °C). The two zircon populations formed by equilibriumcrystallization in a late-magmatic closed system that was progressively enriched with late orthomagmaticLREE-enriched fluids.

© 2013 Elsevier B.V. All rights reserved.

1. Introduction

The Neoproterozoic basement of Sinai crops out over approximately14,000 km2. It constitutes, along with the basement of the EasternDesert of Egypt, thenortheastern segment of theArabian–Nubian Shield(ANS) and is bound by the two branches of the East African Rift System;the Gulf of Suez and Gulf of Aqaba (Fig. 1).

The ANS represents well-preserved juvenile continental crustthat has evolved during the Neoproterozoic Pan-African orogeny(900–530 Ma; Stern, 1994) by accretion and assembly of oceanic andcontinental magmatic arcs through subduction and obduction of

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oceanic crust during closure of the Mozambique Ocean and con-solidation of east and west Gondwana (Stern, 1994, 2002; Johnson,2003; Eyal et al., 2010; Stern et al., 2010; Johnson et al., 2011; Be'eri-Shlevin et al., 2012, 2009a, 2011). Arc accretion is considered to haveterminated in the ANS at ca. 700Ma and was followed by continentalcollision at 640–650Ma (Stern, 1994, 2002).

The Neoproterozoic crustal evolution of Sinai (870–550Ma; Bentor,1985; Stern and Manton, 1987; Be'eri-Shlevin et al., 2012, 2009a,2009b; Eyal et al., 2010;Moghazi et al., 2012) involved the developmentof basins filled with volcano-sedimentary successions that were invadedby post-collisional and anorogenic granitoid intrusions and felsic volcanicrocks (El-Bialy and Streck, 2009; El-Bialy, 2010; Be'eri-Shlevin et al., 2011;Samuel et al., 2011; El-Bialy and Hassen, 2012).

The Late Neoproterozoic post-collisional stage of the ANS evolutionstarted at ~620Ma. Transition from convergence to extension occurred

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Fig. 1. (a) Basic map showing the location and extension of the Arabian–Nubian Shield. (b) Generalized geological map of Sinai, showing the distribution of the main Precambrian rockunits, the main metamorphic complexes (MC) and the location of the mapped area of Fig. 2 (After Be'eri-Shlevin et al., 2009b; Moghazi et al., 2012).

55M.Z. El-Bialy, K.A. Ali / Chemical Geology 360–361 (2013) 54–73

at ~ 590–600Ma (Stern, 1994; Garfunkel, 1999; Jarrar et al., 2003; El-Bialy, 2010; Eyal et al., 2010) and resulted in a stable craton andplatform setting (Bentor, 1985; Garfunkel, 1999; Avigad et al., 2003).The post-collisional stage in the crustal evolution of the Egyptian partof the ANS was characterized by the eruption and emplacementof calc-alkaline volcanic rocks and granitoid intrusions named theDokhan Volcanics and Younger Granites, respectively (Be'eri-Shlevinet al., 2009b; El-Bialy and Streck, 2009; Eyal et al., 2010; Farahatand Azer, 2011; Johnson et al., 2011; El-Bialy and Hassen, 2012). TheDokhan Volcanics are thick sequences of high-K calc-alkaline lavaflows and pyroclastics of predominantly andesitic to rhyolitic com-position in association with ignimbrites that were erupted during aa tectono-magmatic transition between the older calc-alkaline arc-related magmatism and the subsequent alkaline anorogenic mag-matism (590–620 Ma) in the northern part of the Arabian–NubianShield (Stern and Hedge, 1985; Wilde and Youssef, 2000; Breitkreuzet al., 2010; El-Bialy, 2010; Gharib and Ahmed, 2012; Moghazi et al.,

2012). Younger granites are undeformed and highly fractionatedperaluminous calc-alkaline I-type to A-type granitoids (syeno- andmonzogranites and rare granodiorites) that were emplaced at shallowlevels between 610 and 590Ma (Stern and Hedge, 1985; Beyth et al.,1994; Katzir et al., 2007; Jarrar et al., 2008; Be'eri-Shlevin et al., 2009a,2009b; Eyal et al., 2010).

Zircon is a ubiquitousmineral because of its occurrence in almost alligneous rocks, existing not only in crustal rocks but also in mantlexenoliths (e.g. Zheng et al., 2006; Page et al., 2007; Liu et al., 2010;Orejana et al., 2011; Nikitina et al., 2012). It is chemically resistant,relatively insoluble and refractory, and can withstand weathering andrecycling, plus high temperature metamorphism and anatexis. Anadditional reason for its importance stems from its tendency toincorporate many geochemically important trace elements (e.g. Sc, Y,Ti, Hf, Th, U, Nb, Ta, V, P, and REE). Because of this zircon can contain arich and varied record that can be related to geological processes.Previous studies on trace element geochemistry of igneous zircons

56 M.Z. El-Bialy, K.A. Ali / Chemical Geology 360–361 (2013) 54–73

have provided clues to the composition of parental melts (Hinton andUpton, 1991; Rubatto, 2002), the differences between ocean- andcontinent-derived zircons (Peck et al., 2001; Iizuka et al., 2006; Grimeset al., 2007, 2009), source-rock type and crystallization environment(Hoskin and Ireland, 2000; Hoskin et al., 2000; Belousova et al., 2002;Hoskin and Schaltegger, 2003; Hanchar and van Westrenen, 2007;Barth and Wooden, 2010; Nardi et al., 2012; Trail et al., 2012), thediscrimination between different genetic types (i.e., I-, S-, or A-typegranites) of granitoids (Wang et al., 2012) and temperature of forma-tion using Ti-in-zircon thermometry (Watson and Harrison, 2005;Watson et al., 2006; Ferry and Watson, 2007; Harrison et al., 2007;Page et al., 2007; Fu et al., 2008).

Laser ablation inductively-coupled plasma mass spectrometry(LA-ICP-MS) has enabled highly-sensitive elemental analysis to beperformed in-situ on zircon and other accessory minerals. This studypresents LA-ICP-MS analyses of major and trace (including REE)elements of zircons from samples from the Dokhan Volcanics andYounger Granite in Sinai (Fig. 2), which were previously analyzed forU\Pb geochronology and Sr\Nd isotopes by Moghazi et al. (2012). Inorder to contribute to the application of trace element geochemistryof zircon to igneous petrogenesis, we examine here the trace elementcomposition of zircons from the post-collisional calc-alkaline DokhanVolcanics and Younger Granites of the Wadi Kid area, Sinai (Figs. 1, 2).The Ce and Eu contents in the analyzed zircons have been employedas oxybarometers since they are sensitive to magmatic oxidationstate. The possible hydrothermal contribution versus late-magmaticorigin for the zircons contained in these post-collisional igneous rocksis also investigated. Further, many petrogenetic implications havebeen inferred from their variation in some trace element abundances.Additionally, we have used the Ti-in-zircon thermometer to relate

Fig. 2. Geological map of the central-northern Kid Metamorphic Complex and its environs insamples and their age, previously determined by Moghazi et al. (2012), is further indicated.

zircon crystallization to magma temperature (Watson and Harrison,2005; Ferry and Watson, 2007).

2. Geological setting

The Precambrian basement rocks exposed in theWadi Kid area con-sist of low- to medium-grade metamorphosed, thick, folded volcano-sedimentary successions and meta-plutonites (Kid metamorphic com-plex; KMC), extruded by non-metamorphosed felsic to intermediatevolcanic rocks (Dokhan Volcanics) and intruded by plutonic rocksranging in composition from quartz diorite to granite (the Older andYounger Granites) (Fig. 2).

The KMC volcano-sedimentary sequence has been divided into fourformations; the Malhaq and Um Zariq formations (northern Kid area),and the Heib and Tarr formations (southern Kid area) (Furnes et al.,1985).

The Heib Formation is a mixed sequence of weakly metamorposedfelsic to intermediate volcanic rocks with subordinate sediments,whereas the Tarr Formation comprises variable lithologies includinglow-grade metamorphosed lavas, ignimbrites, volcanic breccias andtuffs of dacitic to andesitic composition, mudstones and pebblyvolcanogenic greywackes, and calcareous pelites (Khalifa et al., 2011).The Malhaq Formation comprises a series of dark gray stuctureless(massive) to schistose meta-volcanic rocks interbedded and inter-calated with gray fine- to medium-grained foliated metasediments.The Um Zariq Formation, located in the western central part of theKMC (Fig. 2), is dominantly a metasedimentary sequence, representedmainly by well-bedded metapelitic schists with scarce graphiticphyllites (Abu El-Enen, 2008; El-Bialy, 2013).

Sinai, Egypt (compiled and modified after El-Bialy, 2010, 2013). Location of the studied


Beside these metamorphosed volcano-sedimentary formations, theKMC includes the Quneia Formation and the Shahira metagabbro-diorite complex (meta-plutonites). The rocks of Quneia Formation(Bentor and Eyal, 1987) are exposed as irregular NE- to NW-trendingmasses at the western and eastern margins of the KMC (Fig. 2). Theyconsist of strongly foliated metadiorites, metatonalites and meta-granodiorites enclosing intermittent mafic-rich xenoliths and locallyshowing augen structure. The Shahira metagabbro-diorite complexoccurs as a large intrusive body (about 70–80 km2) at the northernmargin of the KMC (Fig. 2). It intrudes the Malhaq Formation volcano-sedimentary succession and in turn is intruded by the post-collisionYounger Granites along its northern contact at Wadi Malhaq. Recently,the Shahira complex has yielded a U\Pb zircon age of 632 ± 4 Ma(Be'eri-Shlevin et al., 2009a).

The abovementioned KMC rocks are bordered to the north, west andsouth and commonly pierced by unmetamorphosed late Pan-Africanmagmatic rocks. The latter include the syn-orogenic Older Granitesand the post-collisional Younger Granites and the Dokhan Volcanics(Fig. A1a — supplementary materials). Older Granites (650–630 Ma:Stern and Manton, 1987; Moussa et al., 2008) are exposed in thenorthwestern part of themapped area alongWadi Lig andWadi Edakar.They are heterogeneous, exhibiting several different compositions and/or textures. These rocks are gray colored, ranging in composition fromquartz diorite through quartz monzodiorite to tonalite, and in texturefrom granular to porphyritic and gneissose-textured where they aredeformed. The Dokhan Volcanics are mainly exposed along Wadi Kid,Wadi Madsus, Wadi Al Ghuraby and Wadi Qabiliya (Fig. 2). Thesevolcanic rocks form moderate relief and cover a considerable partof the mapped area (Fig. 2). The Dokhan Volcanics consist of non-metamorphosed, brown to dark purple alternating successions ofporphyritic lava flows of mostly felsic composition (rhyolite–dacite),interlayeredwith compositionally equivalent pyroclastic units. The latterare dominated by welded ash tuffs (ignimbrites), although coarser-grained pyroclastic rocks, including lapillistone and agglomerate, arequite common (Fig. A1b — supplementary materials). The pyroclasticrocks have markedly more felsic composition than the lavas. Individuallava flows average less than 10m thick. Rhyolitic flows locally enclosecognate xenoliths, up to 1 meter in diameter, of darker aphyric dacites,indicating later eruption of the former (Fig. A1c — supplementarymaterials). The Dokhan Volcanics extrude and overlie other oldermetamorphosed KMC units including the Quneia Formation gneissesand the volcanosedimentary successions of the Um Zariq and Malhaqformations, and are intruded by the Younger Granites at Wadi Qabiliya.The investigated rhyolite sample (BY-1; Fig. 2) has yielded a U\Pbzircon eruption age of 609± 5 (Moghazi et al., 2012). The eruption ofthe post-collisional Dokhan volcanics defines a tectono-magmatic tran-sition between the older calc-alkaline arc-related and the subsequentalkaline magmatism in the northern part of the Arabian–Nubian Shield(El-Bialy, 2010; Farahat and Azer, 2011).Younger granites are the lastmajor magmatic activity in the region. They form huge expanses oflarge plutons and are easily discriminated in the field into biotitemonzogranites and leuco-syenograniteswith diffuse to quite gradationalcontacts. Additionally, they vary in texture from fine- to coarse-grainedgranular to porphyritic with pink alkali feldspar megacrysts. Theypostdate and intrude the Older Granites, as well as all the KMC rockunits, with sharp contacts (Fig. A1d — supplementary materials), but inturn are cross-cut by numerous dykes of variable composition andorientation. Younger Granites in this region occasionally enclosemicrogranular mafic enclaves. The enclaves are rounded to sub-rounded and tend to have sharp contacts with their granitic host(Fig. A1e — supplementary materials). Recently, ion-probe U–Pb zircondating of some neibouring Younger Granite plutons (e.g. Girgar andRahaba plutons) has revealed that they were emplaced at 604–610Ma(Be'eri-Shlevin et al., 2009a). A comparable U\Pb zircon crystallizationage to that given by Be'eri-Shlevin et al. (2009a) was obtained byMoghazi et al. (2012) from sample KI-1(604 ± 5Ma; Fig. 2) which its

zircons will be investigated in this study. The Younger Granites of Sinai,including the studied sample, have been referred to as the “post-collisional calc-alkaline suite” and are demonstrated to be emplaced ca.635 to 590Ma (Be'eri-Shlevin et al., 2009a, 2009b; Eyal et al., 2010).

3. Petrography of the zircon host-rocks

3.1. Younger granites

The Younger Granites show a subsolvus to transsolvus, allotri-omorphic, medium- to coarse-grained granular to slightly pophyritictexture, and are composed essentially of variable amounts of alkalifeldspar, quartz and plagioclase. The relative proportions of thesemineral constituents vary between the compositions of monzo- andsyenogranite and occasionallygranodiorite. The syenogranites andmonzogranites are biotite, two-mica, and muscovite granites, whereasthe less evolved granodiorites contain hornblende as well as biotite.

Alkali feldspar in the transsolvous granitoids is represented byperthitic varities of microcline and orthoclase, while the subsolvousgranitoids (granodiorites and somemonzogranites) commonly containsingle K-feldpar phases (i.e. microcline). Intergranular swapped rims ofof exsolved albite are widely developed along themutual boundaries ofadjacent microperthite grains (Fig. A2a). Some large microperthitecrystals may enclose relatively abundant intact and variably resorbedsmall plagioclase inclusions (monzonitic texture) (Fig. A2b). Opticalmeasurement of anorthite content (An%) of plagioclase crystals revealswide compositional variation but within the oligoclase-andesinecomposition (An 17–36), with the highest values being obtained fromgranodiorites. In granodiorite and biotite monzogranite, myrmekiticintergrowths are common (Fig. A2c). Quartz forms either large inter-locking grains, up to 6mm diameter, with sutured mutual boundarieswhich may enclose undigested inclusions of early-formed minerals, orsmaller interstitial grains to feldspars. Apatite, zircon and opaque Fe-oxides are the common accessory minerals, with occasional titanite inthe less evolved granodiorites.

3.2. Dokhan Volcanics

Based on their mineralogical composition and textural char-acteristics, the Dokhan Volcanics in the Wadi Kid area can be dividedinto two units, namely lava flows and ignimbrites (El-Bialy, 2010).

Lava flows predominate, comprising chiefly rhyolites with subor-dinate dacites and trachydacites (Fig. 3a; see also El-Bialy, 2010;Moghazi et al., 2012). Lavas are strongly porphyritic, with abundantphenocrysts set in an originally partly glassy, now devitrified, ground-mass. Rhyolite contains quartz, sodic plagioclase and alkali feldsparphenocrysts with occasional biotite and/or hornblende micro-phenocrysts (b0.5 mm). Quartz phenocrysts occur as highly corrodedand embayed crystals 0.5–3 mm across (Fig. A2d). Plagioclase pheno-crysts (An10–19) are lath- or tabular-shaped, euhedral to subhedral, andrange up to 1cm in size, tending sometimes to cluster in glomerocrysts.Some rhyolite samples lack plagioclase as a phenocryst phase. Alkalifeldspar phenocrysts are common in most rhyolites, represented byintact euhedral to resorbed perthitic and normal sanidine and fine-cross-hatch twinned anorthoclase (Fig. A2e). Coexistence of phenocrystsof these two alkali feldspar polymorphs (Fig. A2e), indicates theircontinuous crystallization over a wide range of magmatic temperatures.The groundmass is fine-grained anhedral granular displaying flowbanding, and is composed of quartz, alkali feldspar and plagioclasewith intergranular biotite, opaques, chlorite and accessory apatite andzircon. Dacites and trachydacites consist of abundant plagioclase, quartzand mafic phenocrysts that are enclosed in a felty textured, micro-crystalline groundmass of quartz, feldspars, biotite and opaques.Plagioclase phenocrysts have albite–oligoclase composition (An8–22)and are either independent or as glomerocrysts. Devitrification textures

Fig. 3. Plots of theW. Kid Younger Granites and Dokhan Volcanics on: (a) total alkalis vs. silica (TAS) diagram (Le Bas et al., 1986), with the discriminating boundary between alkaline andsubalkaline fields after Irvine and Baragar (1971). (b) MALI (modified alkali lime index) vs. SiO2 (after Frost et al., 2001). (c) A (Na2O+K2Owt.%)-F (total Fe as FeOwt.%)-M (MgOwt.%)classification diagram of Irvine and Baragar (1971). (d) Yb+Nb vs. Rb tectonic discrimination diagram (Pearce et al., 1984), with the A-type granites field ofWhalen et al. (1987).Fields ofSinai Dokhan Volcanics (gray-filled) are based on the data of Azzaz et al. (2000), Hassen et al. (2001), El-Bialy (2010), Azer and Farahat (2011), and Be'eri-Shlevin et al. (2011), while thoseof Sinai YoungerGranites (bounded bydashed line) are constructedusing thedata ofHassen (1997),Moghazi et al. (1998), El-Bialy (1999), Katzir et al. (2007), Be'eri-Shlevin et al. (2009b)and El-Tokhi et al. (2009).


in the mesostasis of all lava types include granophyric intergrowths,micro-poikilitic quartz and bow-tie spherulites.

Ignimbrites are of rhyolitic composition, with sparse miniaturecrystals and lithic fragments (b10%). They are characterized by theprevalence of glass shards and pumaceous fragments and displayingfeatures indicative of pyroclastic flow origin. The degree of weldingand nature of the principle juvenilematerials (pumice and glass shards)can readily distinguish them into thin-laminated vitric and eutaxitic-textured tuffs (Fig. A2f). Crystals and crystal fragments found includeinteratelluric quartz, sanidine, anorthoclase and plagioclase, while lithicfragments are predominated by cognate volcanic rocks.

4. Whole-rock geochemical affinities

The whole-rock chemical data of samples BY-1 (Dokhan Volcanics)and KI-1 (Younger Granites) from which the zircon grains wereextracted for this work, in conjunction with other 23 samples fromWadi Kid area, previously analyzed byMoghazi et al. (2012), are plottedin selected geochemical diagrams (Fig. 3). These data, in addition tocalculated CIPW normative compositions and several geochemical

indices, are listed in Table A1 (Supplementary materials). For thepurpose of comparison, a compilation of published whole-rock datawas used to construct compositional fields for the Younger Granitesand Dokhan Volcanics of Sinai on the diagrams of Fig. 3.

The DokhanVolcanics and Younger Granites ofWadi Kid area showasubalkaline trend and are of rhyolite–trachydacite composition on thetotal alkali vs. silica (TAS) diagram (Le Bas et al., 1986) (Fig. 3a; Fig. 3in El-Bialy, 2010). The calc-alkaline character of both rock units isverified on the modified alkali-lime index diagram (Frost et al., 2001)and the AFM diagram of Irvine and Baragar (1971) (Fig. 3b, c). Theydisplay a peraluminous to slightly metaluminous character since bothtend to be corundum-normative rather than diopside-normative andmost have A/CNK ratios N1 (Table A1).

On the Y+Nb vs. Rb tectonic discrimination diagram (Pearce et al.,1984), the Wadi Kid Younger Granite and Dokhan Volcanic samplesplot in the volcanic arc and within-plate granite fields (Fig. 5d),reflecting combined geochemical characteristics of subduction-relatedand within-plate settings. It is therefore not surprising that the vastmajority of samples in this diagram cluster within the field of post-collision granites Pearce (1996).


All of the 25 samples fall in the A-type granite field on theMALI-SiO2

diagram of Frost et al. (2001), and further most of them plot in theA-type granite field defined byWhalen et al. (1987) on the Rb\Y+Nbtectonic discrimination diagram (Fig. 3b, d). Additionally, the vastmajority of samples plot in the A-type granite field on all the diagramsof Whalen et al. (1987) and are classified as A2-granites on the Rb/Nb\Y/Nb, Nb\Y\Ga*3 and Nb\Y\Ce diagrams of Eby (1992) (notshown here). The A-type character of these two igneous suites doesnot necessarily imply anorogenic rift-related setting (A1 granites ofEby, 1990, 1992). The A2 sub-type affinity (post-orogenic or post-collisional) refers to generation of their magmas from continentalcrust or underplated crust that has been through a cycle of continent–continent collision or island–arc magmatism (Eby, 1992).

5. Analytical methods

Two samples, one from the Dokhan Volcanics (BY-1) and the otherfrom the Younger Granites (KI-1) from theWadi Kid area were selectedfor studying zircon geochemistry. Zircons were separated at theUniversity of Texas at Dallas (UTD) using standard crushing, heavyliquid and magnetic separation techniques. Grains from the non-magnetic fractions were hand-picked under a binocular microscope,mounted on double-sided adhesive tape, and set in Epirez™ resin.Mounted zircons were ground and polished to effectively cut them inhalf and then they were imaged by cathodoluminescence (CL) using ascanning electronmicroscope prior to gold coating. In situ trace elementmeasurements were performed by LA-ICP-MS using an Agilent 7500aQ-ICPMS connected to a 193 nm Excimer laser ablation system at theInstitute of Geology and Geophysics, Chinese Academy of Sciences,Beijing, China, following techniques described by Xie et al. (2008).One spot was measured on each of fourty-six zircons (20 from sampleBY-1 and 26 from sample KI-1). The isotopes free from isobaricinterference were selected for measurements of trace elements.Standard silicate glass NIST SRM610 was used as external standard for

Fig. 4. Cathodoluminescence (CL) images of zircons from samples analyzed in this study. (a) YoICP-MS spot (30 μm) shown by white dotted circles and zircon sample numbers (S#) are also

the concentration of trace elements in conjunction with the internalstandardization using 29Si (32.8 SiO2 in zircon). Oxide and molecularinterference were assumed to be negligible because of low oxides(e.g. BaO+/Ba+ b 0.3 %, SmO+ b 0.5 %, and ThO+/Th+ b 0.3 %). Highgrade-argon gas, carrying away the ablated materials into the massspectrometer, was measured twice to establish a blank prior to startingeach analysis. Limits of detection (LOD) are typically better than 10 to30 ppb for REE, Nb, Ta, Ba, Hf, Th, and U; 0.1 to 0.3 ppm for Rb, Sr, andY; and 5 to 10ppm for P.

6. Results

6.1. Cathodoluminescence Imaging

Herein, Cathodoluminescence (CL) images were used for spotselection during trace element analyses on the La-ICP-MS,which ablatesa ~30μm spot on each zircon grain (Fig. 4). Representative CL images ofsome of the analyzed zircon grains from samples KI-1 and BY-1 arepresented in Fig. 4a and b, respectively.

Zircons extracted from granite sample KI-1 (28° 19′ 49.83″ N; 34°12′ 33.40″ E) are euhedral to subhedral, displaying long prismaticforms, and are pale brown in color. Generally, the size ranges from160 to 240 μm in length, with length to width ratios ranging from 1:1to 5:1. CL images show faint zoning for some zircons, but others aredominantly homogeneous unzoned grains (Fig. 4a). In a few grains,the internal parts (cores) have many inclusions (e.g. apatite inclusionsin zircon sample KI-1 #15 in Fig. 4a), which were avoided duringanalysis.

Zircons separated from the rhyolite sample BY-1 (28° 12′ 28.1″ N;34° 17′ 53.5″ E) are mostly euhedral to subhedral, stubby prismatic toequant and pale brown in color (Fig. 4b). The zircon grains of thissample are quite smaller and tend to be more resorbed andequidimensional, compared with those of KI-1, ranging from 80 to170 μm in length, with an aspect ratio of 1:1–2:1. Further, CL images

unger Granite sample # KI-1. (b) Dokhan Volcanic rhyolite sample # BY-1. Location of LA-indicated.


(Fig. 4b) show that zircons of this sample displaymore complex internalstructures relative to those separated from sample KI-1, with well-developed oscillatory zoning evident in many crystals, indicatingmultiple stages of zircon crystallization and its primarymagmatic origin(Corfu et al., 2003; Hoskin and Schaltegger, 2003; Cavosie et al., 2006;Barth and Wooden, 2010).

6.2. Zircon geochemistry

Forty six zircon grains were analyzed from samples BY-1 (n= 20)and KI-1 (n= 26). Major (Si, Zr and Hf) and trace element (includingREE) contents were in-situ determined for only one spot in each grain,and the whole data are presented in Tables A2 and A3 (supplementarymaterials). Chondrite-normalized (McDonough and Sun, 1995) spiderdiagrams andREE patterns of the analyzed zircon samples are presentedin Figs. 5 and 6 respectively.

Fig. 5. Chondrite-normalized multi-element spider diagrams of zircons from (a) sample B

Zircon is the primary reservoir for both Zr and Hf, and accordinglythe abundances of both elements are obviously correlated in all of theanalyzed zircons, ranging together between tens to hundered thousandtimes of their concentration in chondrites (Fig. 5). In spite of that crustalmaterials maintain near chondritic Zr/Hf ratios of ~35–40 (Weyer et al.,2002),most of the studied zircons have relative abundances of Zr andHfthat are quite different from that of chondrites. Tables A2 and A3 revealthat the Zr/Hf ratios of most zircons from samples BY-1 (average =52.61) and KI-1 (average=42.63) are higher than the chondritic valuesbut on the other side comparable to igneous zircons from felsic igneousrocks (cf. Claiborne et al., 2006;Wang et al., 2010). Hafnium contents inzircons separated from samples BY-1 and KI-1 vary from 6708 to15444 ppm (11146 ppm in average) and from 8168 to 12517 ppm(10294 ppm in average) respectively. This demonstrates the higherenrichment of Hf in zircons of the Dokhan Volcanics relative to thoseof the Younger Granites and consequently mirrors their crystallization

Y-1 and (b) sample KI-1. C1 chondrite values are after McDonough and Sun (1995).

Fig. 6. Plots of chondrite-normalized (McDonough and Sun, 1995) REE contents in zirconsfrom (a) sample BY-1 and (b) sample KI-1. Plot (c) showswhole-rock REE data from thesetwo samples.


at higher degree of magmatic evolution, since the Hf content of igneouszircon increases with magmatic differentiation (Hoskin and Schaltegger,2003 and references therein).

The chondrite-normalized REE patterns of all zircons in the presentstudy are characterized by a rather steeply-rising slope due to seriousHREE enrichment relative to LREE and MREE with distinctive positiveCe and negative Eu anomalies, which is typical of unaltered magmaticzircons (e.g., Hinton and Upton, 1991; Hoskin and Ireland, 2000;

Belousova et al., 2002; Hoskin and Schaltegger, 2003) (Fig. 6). Theindividual HREE patterns of zircons from each rock sample are almostparallel with restricted ranges of high (Yb/Gd)N ratios (BY-1 = 6.23–40.05; KI-1=8.23–27.43), while the LREE and MREE segments exhibitsomewhat dissimilar patterns as indicated by the wide variation intheir (Sm/La)N ratios (BY-1 = 0.32–305; KI-1 = 0.6–276) (Tables A2and A3). The sum of REE contents in the zircon grains from DokhanVolcanics exhibit broader variation (736–5921 ppm) and higheraverage (1883 ppm) compared with the Younger Granites (631–2481ppm; average=1302ppm). The steepness of the REE patterns ofzircons from Dokhan Volcanics is greater than that of the Youngergranites as expressed by the higher (Yb/La)N ratios of the former(averages; 2129 and 1151 respectively) (Tables A2 and A3).

Compared to chondritic abundance, many of the studied zirconsfrom both samples display an evident overabundance in LREE. Zirconsfrom the samples BY-1 and KI-1 have normalized La values (LaN)above the normal range of igneous zircons (≤10x; Hoskin andSchaltegger, 2003), with values up to 350 and 283 and averages of 34and 65 respectively (Tables A2 and A3; Figs. 5 and 6). This Laoverabundance is more common in the zircons of sample KI-1 (19zircons out of 26). Further, Cerium which is characteristically high inigneous zircons leading to its universal positive anomaly recordoverwhelming normalized values surpassing the normal limit of 100chondrite abundance (Hoskin and Schaltegger, 2003) as indicated inTables A2 and A3 and Fig. 6. Cerium contents in these igneous zirconsare relatively high, implying crystallization from an oxidized magmawhere tetravalent Ce was abundant. Nonetheless, these excessiveLREE abundances will be discussed in depth (see section 7.2.). On theother hand, MREE occasionally display slight enrichment (N100×chondrite), whereas HREE in all zircons fluctuate within the typicalrange of 103 and 104chondrite (Hoskin and Ireland, 2000; Poller et al.,2001; Hoskin and Schaltegger, 2003; Barth and Wooden, 2010)(Fig. 6; Tables A2 and A3). Yttrium, which behaves like the HREE, haschondrite-normalized abundances similar to them (Fig. 5). Excludingzircon grain BY-1-07 (Y=7947ppm), all the Y concentrations obtainedin this study (722–4584ppm) resemble those of zircons from granitoidsin the literature (Hoskin and Ireland, 2000; Poller et al., 2001; Hoskinand Schaltegger, 2003; Belousova et al., 2006; Barth and Wooden,2010; Orejana et al., 2012; Wang et al., 2012).

With the exception of zircon grains BY-105 and KI-1-09, displayingslightly negative Ce anomalies (Ce/Ce* = 0.54 and 0.78 respectively),the studied zircons exhibit pronounced positive Ce anomalies ofvariable magnitudes. Sample BY-1 zircons display wider variationsand higher positive anomaly values relative to sample KI-1 (2.03–170.79 and average = 44.98 versus 1.12–23.23 and average = 4.6respectively). In both samples, it is remarkable that the size of thepositive Ce-anomaly correlates with the degree of LREE depletion(Fig. 6) (cf. Pettke et al., 2005). All the zircons analyzed possess negativeEu anomalies of variable sizes. Zircons from sample KI-1 showconstrained range of small Eu/Eu*values (0.07–0.33, average = 0.17)leading to seriously deep negative anomlies (Fig. 6b), while zircons ofBY-1exhibit shallower negative Eu anomalies (0.2–0.62, average =0.32; Fig. 6a). The markedly deeper negative Eu anomalies in sampleKI-1 zircons imply the major role of feldspar fractionation prior to (orduring) zircon crystallization. This combination of both positive Ceand negative Eu anomalies is a feature that may be unique to zircon(Hinton and Upton, 1991; Hoskin and Schaltegger, 2003; Cavosieet al., 2006; Hanchar and van Westrenen, 2007; Claiborne et al., 2010;Reid et al., 2011; Trail et al., 2012). Positive Ce and negative Euanomalies may be respectively linked to favorable incorporation ofsmall quantities of Ce4+ from a relatively oxidized melt and to a Eu-depleted magma composition due to plagioclase fractionation (e.g.Hinton and Upton, 1991; Hoskin et al., 2000; Pettke et al., 2005;Claiborne et al., 2010; Trail et al., 2011, 2012; Yao et al., 2012). Adistinctive feature of sample BY-1 zircons is the evident positiveSm anomalies in their normalized REE patterns (cf. Nardi et al.,


2012).There is almost no correlation between the sum of REE and Ycontent in zirconswith progress ofmagmatic differentiation (expressedby Hf as differentiation index). The largest variation in both may occurat the same level of Hf concentration (Fig. 7).

Zircons from the Dokhan Volcanics (Sample BY-1) containmarkedlyhigh abundances in Pb, Th and U (averages; 38.29, 274 and 258 ppmrespectively) relative to zircons from the Younger Granites (averages;19.38, 86.43 and 172.62 ppm respectively) (Tables A2 and A3).Conversely, U is relatively enriched compared with Th in zircons ofsample KI-1 (average Th/U= 0.55) while both are present in roughlysubequal concentrations in sample BY-1 zircons (average Th/U= 1.0)(Tables A2 and A3). Most of the analyzed zircons have Th/U ratios≥0.5that is typical of igneous zircons (Hoskin and Schaltegger, 2003).Noteworthy, there exists a considerable large inter-grain compositionalvariation in U and Th, which is a characteristic feature of most zircon(e.g. Hidaka et al., 2002; Pettke et al., 2005; Kaczmarek et al., 2008;Orejana et al., 2011; Ye et al., 2011; Li et al., 2012). Such inter-grainvariation may exceed ten order-of-magnitude, such as in Th of BY-1zircons (93–1596 ppm; Table A2) and U of KI-1 grains (40–487 ppm;Table A3). The covariation of Th and U against Hf rise in zircons fromthe two rock types is interesting. While both compatible elementsshow gradual increase along quite broad trends with Hf elevation insample KI-1 zircons, the Hf concentrations of analyzed grains fromsample BY-1 are positively and negatively correlated with U and Threspectively (Fig. 7). The Th/U ratios of zircons from the two rocksamples are well correlated with magma differentiation and cooling(Hf rise) along two discernable parallel trends of regression (Fig. 8a).This negative correlation between Hf and Th/U (both are commonlyused as differentiation indices) is consistent with the sensitivity of theTh/U ratio to variation of the melt temperature (inverse proportion)(Bolhar et al., 2008; Gagnevin et al., 2010).

Fig. 7. Trace element variation diagrams for the studied zircons using Hf (abscissa) as a differentiation index. Abundances of all variables are in ppm.

Fig. 8. Plots of the studied zircons on the: (a) Hf\Th/U and (b) Ca\P variation diagrams.


The high-field strength elements (HFSE) Nb, Ta and Ti, substitutingfor Zr (Hoskin and Schaltegger, 2003; Van Lichtervelde et al., 2011),are found in significant detectable amounts in all of the analyzed zircons(Tables 1 and 2). The most abundant of them is titanium, with averagecontents of 52.92 ppm and 31.33 ppm in samples BY-1 and KI-1respectively, which is in conformity with the normal abundance of Tiin zircon (≤75 ppm; Hoskin and Schaltegger, 2003). Nonetheless, twograins out of all the analyzed zircons have yielded exceedingly high Tivalues (BY-1 15=838ppm, KI-1 15=236ppm). Such surplusmeasuredvalues cannot represent primary igneous zircon and most likely reflectanalysis of altered zircon or zircon+ inclusions (e.g. rutile). Nb and Tacontents of zircons from sample KI-1 (average = 11.26 and 3.03 ppmrespectively) are noticeably higher (by one order-of-magnitude) thantheir abundances in zircons of BY-1 (average = 5.94 and 1.08 ppmrespectively). The abundances of Nb and Ta in BY-1 zircons and Nb inKI-1 zircons fall within the normal limit of unaltered magmatic zircon(Nb ≤ 62 ppm, Ta b3 ppm; Hoskin and Schaltegger, 2003). On thecontrary, several zircons from sample KI-1 (ten out of twenty six) containhigh Ta concentrations (up to 7.22 ppm; Table A3). These Ta-enrichedzircons have been assumed to be crystallized from a Ta-rich graniticmelt before the precipitation of tantalite or other Ta oxides (VanLichtervelde et al., 2009, 2011). The enrichment of Ta over Nb in all theanalyzed zircons relative to chondritic abundance is evident in Fig. 5.

Phosphorous contents in zircons of sample BY-1 are tremendouslyhigher relative to those from sample KI-1 (averages: 2943 and733 ppm respectively). Recent experimental work of Van Lichterveldeet al. (2011) has evidenced that P can be incorporated by coupledsubstitution in zircon via two different substitution mechanism (plusAl+3 or Mn+2) to maintain charge balance. Although, many of the Pcontents in zircons analyzed in this work are somewhat higher thanthose reported for natural zircons in the literature (e.g. Hoskin et al.,2000; Ballard et al., 2002; Pettke et al., 2005), they fall within thephosphorous values of the synthetic zircons of Van Lichtervelde et al.(2011). Calcium is the next element in abundance to Zr and Si amongthe analyzed elements, reaching percent level in many zircons(Tables A2 and A3). These appreciable calcium abundances in zirconcombined with the reasonably good positive correlation between Caand P in the present zircons (Fig. 8b) suggest that not all of themeasured phosphorous is contained in the zircon lattice, but may beattributed to micro-inclusions of apatite.

Regarding the large ion lithophile elements Ba, Sr and Rb in thestudied zircons, they are found in very low concentrations, rangingfrom decimals to few ppm's (cf. Hoskin and Schaltegger, 2003; Xiaet al., 2010). Ba and Sr display subchondritic values, while Rb surpasseswith contents up to 10x chondrite (Fig. 5).

7. Discussion

7.1. Ce and Eu anomalies of zircons and magma oxygen fugacity

Zircon possess distinctive characteristics in its REE pattern, and itsvariations in Ce and Eu anomalies can be interpreted to reflect zirconcrystallization under physiochemical conditions (Ballard et al., 2002;Pettke et al., 2005; Barth and Wooden, 2010; Claiborne et al., 2010;Trail et al., 2011, 2012; Li et al., 2012; Burnham and Berry, 2012;). TheCe and Eu contents in zircons have been recently employed inoxybarometry since they are sensitive to magmatic oxidation statedue to multiple ionic states.

Themagnitude of thenegative Eu anomaly depends on the oxidationstate during mineral crystallization from fluid/melt (Hoskin andSchaltegger, 2003; Xia et al., 2010; Trail et al., 2012). On the otherside, once Ce3+ (ionic radius r = 1.143 Å) is oxidized to Ce4+ (r =0.970 Å), it behaves similar to Zr or Hf, and accordingly is preferablysubstituted for them in comparison to the other LREE in zircon.Therefore, the positive Ce anomaly occurs because Ce4+ (vs. La3+ andPr3+) is more compatible in zircon (substitutes for Zr4+, Hf4+, U4+,

Th4+) than in the other minerals (Ballard et al., 2002; Hidaka et al.,2002; Hoskin and Schaltegger, 2003; Xia et al., 2010; Trail et al., 2011,2012).

Chondrite normalized REE patterns of zircons commonly haveenriched Ce contents relative to La and Pr, and depleted Eu abundancessrelative to Sm and Gd. High Ce concentrations in zircon may reflectoxidizing conditions (Ce4+ is more compatible than Ce3+), whereasdepleted Eu contents may mirror reducing conditions as Eu2+ doesnot substitute into the zircon lattice (Trail et al., 2012). The coexistenceof these two anomalies in zircon results in a contradiction since thepresence of a positive Ce anomaly indicates oxidizedmagma conditionsand a coexisting negative Eu-anomaly implies reducing conditions. Thispuzzle may be rationalized by plagioclase fractionation which causesdepletion of Eu from the magma before or during zircon crystallization(Hoskin et al., 2000; Hoskin and Schaltegger, 2003; Kaczmarek et al.,2008; Burnham and Berry, 2012). Therefore, the Eu-anomaly inzircon is inherited from the Eu-depleted melt (as exampled in Fig. 6a,b and c) as well as being influenced by ƒO2. Essentially, at all terrestrialoxygen fugacities, Ce and Eu anomalies will coexist in zircon and foundto vary systematically with fO2 (Trail et al., 2011, 2012). Consequently,both Ce and Eu anomalies may be produced exclusively by zircon/meltpartitioning under terrestrial conditions (Burnham and Berry, 2012;Trail et al., 2012).

Herein, for evaluating the oxygen fugacity (fO2) using zircon Ce andEu anomalies, both of these ratios are calculated according to theexpressions given in Hoskin and Schaltegger (2003) and presented inTables A2 and A3.

Zircons from sample KI-1 exhibit deeper negative anomaliesexpressed by lower Eu/Eu* values (average = 0.17) compared withthe shallower Eu anomalies of sample BY-1 zircons (average Eu/Eu*=0.32) (Fig. 6a, b). This indicates that the crystallization of sample BY-1zircons from awetter oxidizedmelt (Higher fO2). Similarly, Ce/Ce* ratiosconfirm the same finding as zircon of sample BY-1 havemarkedly higherCe contents and more spiky positive anomalies (Tables A2 and A3;Fig. 6a, b), inferring higher oxygen fugacity. However, a single zircongrain from each sample yields a bizarre negative Ce anomaly (BY-1 05and KI-1 09), which may be related to apatite inclusions that can biasthe LREE concentrations in zircons.

A Ce/Ce* vs. Eu/Eu* diagram for all zircons from the two rocksamples is shown in Fig. 9a. Zirons from the Younger Granite sample(KI-1) show a constrained variation in both ratios within quitemoderate values, while zircons of sample BY-1 spread over a widerranges of both values in general and for Ce/Ce* values in particular.This gives clues to the crystallization of the zircon from the DokhanVolcanics (sample BY-1), perhaps in more variable and higher oxygenfugacity conditions relative to those from sample KI-1. A positive linearto curvilinear correlation between Ce and Eu anomalies would beexpected if both anomalies were simply controlled by the oxidationcondition during zircon crystallization (cf. Hidaka et al., 2002; Orejanaet al., 2011; Ye et al., 2011; Wang et al., 2013). Therefore, the broadanti-correlation observed between the two anomalies (Fig. 9a), suggeststhat oxygen fugacity is not the only factor that controls the Eu and Ceoxidation states in both rock types hosting zircon. Further, this leadsto conclude that the Ce/Ce* is more sensitive than Eu/Eu* to variationin the oxidation state, and thusmay bemost readily applied to constrainthe oxidation state of natural melts (Trail et al., 2012).

Monitoring the variation of the oxidation state with magmadifferentiation is acheived through plotting of the Hf versus Ce/Ce*(Fig. 9b). With slight scatter of few data points, sample KI-1 zirconsappears to be crystallized in a limited range of fO2 over a rathershort time interval (narrow ranges of Hf and Ce/Ce* values). Althoughthe plots of sample BY-1 zircons are not perfectly correlated anddisplaying some kind of scatter, it is obvious that they werecrystallized from more oxidized magma that had gradually changedto more reduced (lower Ce/Ce* and fO2) with progressive magmaevolution (Hf elevation).

Fig. 9. Plots revealing the oxidation state of the magma from which studied zircons werecrystallized. (a) Ce/Ce* vs. Eu/Eu* diagram. (b) Hf vs. Ce/Ce* diagram.

Fig. 10. The overabundance of the LREE in the studied zircons illustrated by plotting Hfagainst LaN (a), CeN (b) and (LREE)N (c). (LREE)N=LaN+CeN+PrN. Igneous zircon fieldsdata are from Hoskin and Schaltegger (2003).


7.2. Zircon LREE overabundance and its source

In zircons, HREE and MREE readily substitute for Zr, whereas LREEare considered incompatible during zircon crystallization from felsicmelts because of their relatively large ionic radii and high charge, andhence rather prefer to reside in coexisting melts (Hanchar et al., 2001;Hoskin and Schaltegger, 2003; Hanchar and van Westrenen, 2007).

The normal absolute abundances of the LREE La and Pr in igneouszircon are sub-ppm to ppm-level, whereas Ce abundances are ex-ceptional, as it is much more compatible in zircon, reaching up toabout 50 ppm (Hoskin and Schaltegger, 2003 and references therein).The LREE (La-Pr) absolute and chondrite-normalized values of theanalyzed zircons from samples BY-1 and KI-1 show obvious enrich-ment, which may reach into overall radical overabundance (Tables A2and A3). Compared with chondritic abundances, lots of the analyzedzircons from the two studied samples have overenriched La and Cecontents exceeding 1 and 100 times chondrite value respectively(Fig. 10a, b). Further, the sum of the LREE (La, Ce and Pr) of all zirconsamples surpasses the normal content in igneous zircon (N10chondrite), confirming the inclusive overabundance in LREE for thestudied zircon samples (Fig. 10c). This extreme overenrichment inLREE (average (Sm/La)N = 45.61 and 27.19 for BY-1 and KI-1respectively), leads to considerably flat LREE patterns for most of thestudied zircons in contrast to the negative steep slope LREE patterns ofcontinental crust zircon ((Sm/La)N = 57–547), which are furtherapproaching LREE patterns of zircons of mantle-affinity ((Sm/La)N =

11.3–38; Hoskin and Schaltegger, 2003). Such unsystematic andexcessive incorporation of LREE in zircons were discerned anddocumented recently by many studies (e.g. Whitehouse and Kamber,2002; Geisler et al., 2003; Cavosie et al., 2006; Xia et al., 2010; Nardiet al., 2012; Orejana et al., 2012). Several explanations on the sourceof the observed LREE overenrichment in the analyzed zircons arediscussed below.

7.2.1. Accidental analysis of LREE-bearing inclusionsGiven the 6–8μmdepth of the REE analysis ablation pit (Fig. 4), LREE

enrichment of zircon could be an analytical artifact of analyzing micro-scale LREE-enriched phosphate mineral inclusions (i.e. monazite and/orapatite) that may be present in zircon grains, but were undetectable at


the scale of detailed pre-analytical optical and CL imaging investigation.Throughout the LA-ICP-MS analyses, these LREE-bearing mineralinclusions may possibly be incorporated into the analyzed volumes,giving rise to outstandingly high LREE concentrations (e.g., Cavosieet al., 2006; Xia et al., 2010).

The presence of apatite and monazite inclusions can be preliminaryevaluated by testing the correlation between degree of LREE enrich-ment ((Sm/La)N) and phosphorous concentration. As shown inFig. 11a, there is no correlation between P content and degree of LREEenrichment, but instead the (Sm/La)N ratio tend to be almost constantover wide range of P concentrations, indicating that apatite or monazitelikely does not exist pervasively in the studied zircon grains. Morespecifically, monazite ((Ce,La,Th,Nd,Y)PO4) inclusions should be easily

Fig. 11. Degree of LREE enrichment (expressed as (Sm/La)N) against geochemicalindicators of possible contaminant minerals in the zircon analysis: (a) P as a generalindicator of phosphate mineral inclusions; (b) Th as an indicator of the presence ofmonazite; (c) Ca/Sr as an indicator of apatite.

distinguished by elevated levels of Th associatedwith LREE enrichment.A plot of (La/Sm)N versus Th (Fig. 11b) shows that there is no positivecorrelation, but instead an inverse correlation in some of BY-1 samplezircons, the cause of which is ambiguous. Nevertheless, this anti-correlation definitely argues against analyzing microinclusions ofmonazite. Apatite (Ca5(PO4)3(OH,F,Cl)) inclusions are common inmany igneous zircons. In contrast to monazite which is commonlyLREE-enriched, the REE patterns of igneous apatites range frommoderately LREE-enriched, through almost flat to MREE-enriched(Whitehouse and Kamber, 2002), and thus their inclusion may notalways result in LREE enrichment. Given that most apatites, particularlythose of igneous origin, are strongly enriched in Sr, a correlationbetween degree of LREE enrichment and Ca/Sr ratio would be expectedfor LREE-enriched apatites, but as shown in Fig. 11c this ratio is variableover almost constant to narrow range of (La/Sm)N values. On account ofthese arguments, we can confidently exclude contributions from eithermonazite or apatite to explain the LREE overabundance in the studiedzircons.

7.2.2. Complex REE substitution mechanismsREE-bearing zircons commonly contain P, and owing to crystal-

chemical resemblance between Y3+ and heavy REE3+, replacement ofZr4+ by REE3+ in zircon is commonly explained by the coupledxenotime-type substitution tomaintain charge balance. In the xenotimesubstitution, it is assumed that P5+ replaces Si4+ at the tetrahedral siteand the trivalent REE replace Zr4+ at the dodecahedral site (Hinton andUpton, 1991; Hanchar et al., 2001; Finch and Hanchar, 2003; Hancharand van Westrenen, 2007). If xenotime substitution is the solemechanism by which charge balance is maintained in REE-substitutedzircon, the atomic ratio REE+Y: P must be unity (Finch and Hanchar,2003; Cavosie et al., 2006). As evidenced by the elevated (REE+ Y)/Pratios of most zircon grains (Fig. 12a), the xenotime-type substitutiondoes not appear to be the major cause of LREE enrichment for thezircons in this study, as the majority of (REE + Y)/P ratios deviatetowards higher values (2–14; Fig. 12a), indicating a more complexcharge-balance mechanism or multiple mechanisms (Whitehouse andKamber, 2002; Finch and Hanchar, 2003; Cavosie et al., 2006). Phos-phorus content appears to have some impact on the substitutionmechanism(s), as the highest (REE + Y)/P values (e.g. N 3) are foundin grains with less than 1000ppm P (Fig. 12a). Because natural zirconsdo not contain sufficient P to charge balance the combined trivalentREE (Hoskin et al., 2000), zircon grains with notably elevated P values(e.g. N 1000 ppm) may possibly have been enriched in P throughsecondary processes rather than by magmatic incorporation.

7.2.3. LREE enrichment due to radiation-induced lattice damageFor zircons that have apparently experienced a complex geological

history, like ours, there is a likelihood of post-magmatic incorporationof LREE into radiation-damaged regions (Whitehouse and Kamber,2002). Many zircons from samples BY-1 and KI-1 have features requireat least one, if not multiple, post-magmatic events of alteration,including discordant U–Pb ages, Pb loss and recrystallization (Moghaziet al., 2012), variable Th–U ratios (Tables 1 and 2, and absent toobscured CL growth zoning (e.g. Fig. 4a). Therefore, several zircon grainsthat are now annealed were areas that may have suffered radiationdamage, and which LREE were preferentially incorporated into theirdamaged lattice sites. Consequently, it is possible that the REE com-position of these zircons is a consequence of the same processes thataffected the U–Th–Pb systematics.

As suggested by Whitehouse and Kamber (2002), the propor-tionality of LREE incorporation to the lattice damage that the zirconexperienced due to alpha particle recoil could be indicated fromthe correlation between degree of LREE enrichment and combinedU and Th content. This probable source of LREE enrichment in thestudied zircons is evaluated by plotting their combined U+Th contentvs. (La/Gd)N (Fig. 12b). The studied zircons show poorly-defined

Fig. 12. Atomic ratios of (REE + Y)/P vs. P content (a) and Chondrite-normalized La/Gdvalues vs. combined Th and U concentration (b) for the analyzed zircons.


vertical trend on the (La/Gd)N) vs. (U+Th) plot (Fig. 12b), in oppositionto the steeppositive trend expected for a longer period of lattice damageaccretion prior to an event(s) leading to LREE incorporation. Thus,radiation-induced lattice damage in the present zircons does not appearto be the source of the observed LREE superenrichment.

7.2.4. Hydrothermal alterationHydrothermal alteration is the last possible process that may lead to

the serious LREE enrichment of the analyzed zircons. Hydrothermalalteration has been proposed to be a widespread source for LREEenrichment in magmatic zircons (e.g. Whitehouse and Kamber, 2002;Hoskin, 2005; Pettke et al., 2005; Rayner et al., 2005; Fu et al., 2009;Xia et al., 2010).The effect of hydrothermal processes in the growthand/or formation of the studied zircons will be explored and discussedin depth in a following separate section.

7.3. Petrogenetic implications from Th/U, Zr/Hf and U/Yb ratios

Th/U ratios in igneous zircons from various rocks generally equals orexceeds 0.5 (Hoskin and Schaltegger, 2003; Xiang et al., 2011), whilezircons grown under metamorphic events show considerably lowerTh/U (~0.01) (Hidaka et al., 2002; Rubatto, 2002; Hoskin andSchaltegger, 2003). In this study, Th/U values of the zircons fromsamples BY-1 and KI-1 (average= 1.0 and 0.55 respectively) confirmtheir igneous origin (Fig. 13a).

Also, Th/U ratio has been shown to be sensitive to temperaturevariations (Bolhar et al., 2008; Gagnevin et al., 2010). Decreasingmagma temperature should promote higher U contents relative to Thcontents, resulting in lower Th/U ratios for zircon crystallized fromlower temperature magma (Xiang et al., 2011). Zircons of sample BY-1

have generally higher Th/U ratios, as much as twice, relative to sampleKI-1 zircons (Fig. 13a; Tables A2 and A3), implying crystallization ofthe zircons of the Dokhan Volcanics at higher temperatures comparedwith those of the Younger Granites. This inference will be latercorroborated using the Ti-in-zircon thermometry (section 7.4).

The incorporation of Th and U into zircon can be highly affected bythese elements availability during zircon crystallization and partitioningwith co-existing minerals. Th–U concentrations in zircon would bechiefly controlled by monazite (Villaseca et al., 2003), which incor-porates Th as a major structural component and gives rise to very highTh/U values up to 14 (e.g. Hermann and Rubatto, 2003). Therefore, thezircons with abnormally high Th/U ratios most probably incorporatedmuch higher Th contents due to destruction of monazite, while lowTh/U zircons would have co-crystallized in equilibrium with monazite(Orejana et al., 2011). Accordingly, the low Th/U content (i.e. b1) of allof the KI-1 and most of B Y-1 zircons (Fig. 13a) may indicate their co-crystallization in equilibrium with monazite.

During magmatic differentiation, the Hf abundance of zirconincreases while the Zr/Hf ratio tend to decrease (Hoskin et al., 2000;Linnen and Keppler, 2002; Hoskin and Schaltegger, 2003). Therefore,Th/U and Zr/Hf ratios can be used effectively together as differentiationindices in zircon. Plots of the studied zircon samples on Th/U vs. Zr/Hfdiagram show good positive correlation as both ratios decline withprogressive magmatic evolution (Fig. 13b). The relative distribution ofthe data points on this diagram illustrates that the crystallization ofsample BY-1 zircons span over longer period and wider range oftemperatures compared with those of sample KI-1.

The U/Yb ratio for zircons from different source regions are distinct,and increase from ocean gabbros (0.18) to continental granitoids (1.07)and kimberlites (2.1) (Grimes et al., 2007). Grimes et al. (2007)proposed two discrimination diagrams based on U/Yb ratio vs. Y andHf content to distinguish between zircons from continental crust,ocean crust, and the mantle (kimberlite megacrysts). In Fig. 13c, d, allthe studied zircons from both samples plot in the field of continentalzircons with slight trend towards the upper limit of ocean crust zirconfield. The identity of the sources that generated the voluminous, mostlyfelsic, post-collisional calc-alkaline magmas (i.e. Dokhan Volcanics +Younger Granites) from which the studied zircons have been crys-tallized are debated (Be'eri-Shlevin et al., 2009a). Hypotheses proposedvary between that these magmas were generated by anatexis of olderisland arc crust (Beyth et al., 1994; Moghazi et al., 1998; El-Bialy,2010) to partial melting of either subducted oceanic crust (Jarraret al., 2003) or mantle-derived mafic lower crust (Be'eri-Shlevin et al.,2009a). The observed explicit continental crust signature of the studiedzircon and their delicate bias towards the continental zircon field lowerlimit (Fig. 13c, d) is consistent with generation of their magmas bypartial melting of older juvenile island arc crust.

7.4. Ti-in-zircon thermometry

Incorporation of titanium into crystallizing zircon is principallycontrolled by temperature and the activity of TiO2 (aTiO2) (Watsonand Harrison, 2005). Therefore, Ti concentration in zircon is a powerfulgeochemical tracer to constrain zircon crystallization temperature(Watson and Harrison, 2005; Watson et al., 2006; Ferry and Watson,2007). This thermometer (Ti-in-zircon thermometer; Watson et al.,2006) assumes coexistence of rutile (i.e. pure TiO2) with zircon attime of crystallization, which is a reasonable postulation for most felsicmagmatic rocks. On account of this assumption, an equilibrium constantcan be calculated from aTiO2 in zircon by setting the rutile activity equalto almost 1 (Watson andHarrison, 2005;Watson et al., 2006). However,Ferry and Watson (2007) based on later thermodynamic models andrevised calibrations for the Ti-in-zircon thermometer have found thatthe solubility of Ti in zircon depends not only on T and aTiO2 but alsoon aSiO2. Experimental calibration of this thermometer has been madeusing experimentally synthesized zircons for high (1025–1450 °C)

Fig. 13.Variation of Th vs. U (a) and Th/U vs. Zr/Hf ratios (b) for zircons from theDokhan Volcanics (BY-1) and Younger Granites (KI-1). Plots of the zircons fromBY-1 and KI-1 on theU/Ybvs.Hf (c) and U/Yb vs. Y (d) diagrams with the fields of Grimes et al. (2007) to discriminate between continental and oceanic crust zircon. Heavy lines indicate the lower limit of zirconsfrom continental crust.


temperatures, and natural zircons with estimated crystallizationtemperatures of ~580 °C–1170 °C (Watson et al., 2006).

The Ti diffusion into zircon is slow (Cherniak and Watson, 2007),and capture of Ti4+ in zircon is deceptively assisted by its substitutionwithout charge compensation most favorably into the Si4+ site (Ferryand Watson, 2007; Harrison et al., 2007). This substitution has beenshown to be pressure sensitive at pressure exceeding 10 Kbar (Ferrisset al., 2008), whereas the pressure correction for the Ti-in-zirconthermometer is approximately 50 °C/10 kbar at 750 °C for most crustalmaterials (Ferry and Watson, 2007).

Application of the Ti-in-zircon thermometer is best in systems thatcontain a pure TiO2 phase (e.g., rutile), although is also suitable formelts with a Ti-saturated phase such as ilmenite (Cates and Mojzsis,2009). Recently, Ti-in-zircon thermometry has become popular andhas been increasingly applied to a growing number of natural zircons(e.g. Claiborne et al., 2006; Harrison et al., 2007; Page et al., 2007; Fuet al., 2008; Kaczmarek et al., 2008; Cates and Mojzsis, 2009; Barthand Wooden, 2010; Ickert et al., 2011; Orejana et al., 2011, 2012;Wielicki et al., 2012; Wang et al., 2013).

Even though calculation of exceedingly precise magmatic zirconcrystallization temperatures (TTi-in-zrc) requires knowledge of aTiO2 andaSiO2 activities (Ferry and Watson, 2007), our estimation is based onthe calibration of Watson and Harrison (2005) and Watson et al.

(2006), which assumes that zircons crystallized in the presence ofquartz and rutile at P≈10kbar. Themaximumuncertainties introducedto zircon crystallization temperatures (TTi-in-zrc) by unconstrained aTiO2and aSiO2 were quantitatively assessed to be ≈ 60–70 °C at 750 °C(Ferry and Watson, 2007). Zircon grains yielding erroneously hightemperatures (BY-1-15 and KI-1-15 = 1373 °C and 1123 °C,respectively) due to their bizarre anomalous Ti content (Tables A1 andA2) are disregarded herein. As mentioned earlier, such unusually highmeasured Ti contents in zircon are not viable for primary igneous zirconand in all likelihood point toward analysis of altered zircon or combinedzircon and Ti-rich inclusions (e.g. rutile). The application of the Ti-in-zircon thermometer (Watson et al., 2006) to the investigated zirconssamples returns temperatures of 790–986 °C (average TTi-in-zrc =901 °C) for sample BY-1 zircons and 781–934 °C (average TTi-in-zrc =847 °C) for sample KI-1 zircons. The higher range of temperatures forsample BY-1 (rhyolite) grains relative to those of sample KI-1(syenogranite) zircons can be considered axiomatic, since it is widelyaccepted that zircons developed from primitive melts and in extrusiverocks crystallize over a considerable range of higher crystallizationtemperatures compared to zircons from evolved magmas and inplutonic rocks (more limited range of lower temperatures).

Because magmatic zircons record steady variation in Hf over asimilar range in Ti abundance, this can be connected with the thermal


evolution of the melt using the Ti thermometer. Plotting Hf contents ofthe studied zircon samples against their calculated crystallizationtemperatures reveals gradual decrease in temperatures with elevationof Hf concentration due to zircon/melt Hf/Zr ratio fractionation duringmagmatic evolution (Fig. 14). With the exception of two samplesyielding extremely high temperatures (i.e. Ti NN 75 ppm), all sampleslie within the range of unaltered magmatic zircon (Fig. 14). Theobtained zircon crystallization temperatures for zircons separatedfrom samples BY-1 and KI-1 are somewhat high or comparable to theclimax temperatures determined for zircons from analogous per-aluminous rhyolites (e.g. Fu et al., 2008; Colombini et al., 2011; Reidet al., 2011; Stelten and Cooper, 2012) and monzo-syenogranites (e.g.Claiborne et al., 2006; Harrison et al., 2007; Barth and Wooden, 2010;Ickert et al., 2011; Orejana et al., 2012), respectively. These relativelyhigh temperatures constrained for zircons from the two igneous rockunits are indicative of a deep level of melting, likely within the lowercrust of the ANS.

7.5. Zircon saturation

Experimental studies have revealed that zircon, when crystalline,has low solubility in crustal melts and fluids (e.g., Watson, 1979;Waston and Harrison, 1983; Ayers and Watson, 1991). These studieshelp to constrain the temperature at which zircon crystallized and theamount of Zr required to saturate and crystallize zircon in the rockfrom which the zircon was extracted (Hanchar and Watson, 2003).Zircon saturation temperature estimates (Tsat) for 25 rock samples ofboth rock units are calculated and listed in Table A1, according to theexperimental work of Waston and Harrison (1983) on Zr saturation inhydrous, low-temperature, non-peralkaline intermediate and felsicmagmas.

Temperature estimates of the Younger Granite samples fall betweena minimum of 762 °C and a maximum of 820 °C, with averagetemperature of 799 °C for the five samples, whereas Dokhan Volcanics(20 samples) display a considerably higher and wide range oftemperatures (769–866 °C; average = 819 °C) (Table A1). The rocksamples BY-1 and KI-1, chosen to study the chemistry of their zircons,have yielded zircon saturation temperatures of 846 °C and 806 °Crespectively. The whole-rock zircon saturation temperatures (Tsat) ofalmost all samples fall within those obtained for the post-collisional

Fig. 14. Plots of the calculated crystallization temperatures of the studied zircon samples usingzircon saturation temperatures Tsat (Waston and Harrison, 1983) for the host rock samples Bigneous zircon is after Hoskin and Schaltegger (2003).

crust-derived Dokhan Volcanics in Sinai (766–882 °C; El-Bialy, 2010).The crystallization temperature spectrum of extracted zircons fromsample KI-1 (average TTi-in-zrc=847 °C) is quite consistent, even beingsomewhat higher, with the average zircon saturation temperature(Tsat) of ca. 806 °C and 799 °C calculated for sample KI-1 and thewhole Younger Granite samples respectively (Table A1). The situationis different when comparing the crystallization temperatures of sampleBY-1 zircons (average TTi-in-zrc = 901 °C) with the obviously lowercalculated zircon saturation temperatures of 20 Dokhan volcanicsamples (average Tsat= 819 °C) and the saturation temperature of thesample itself (Tsat = 846 °C) (Fig. 14). Claiborne et al. (2006) havepresented evidence of a broad range of zircon crystallization tem-peratures (TTi-in-zrc) beginning well above its saturation temperature(Tsat). Further, the results of the model combining both types of zirconthermometry, presented by Watson et al. (2006), have evidentlyshown that the zircon saturation temperature (Tsat) can significantlyunderestimate the onset temperature of zircon crystallization by50–100 °C. Furthermore, the high crystallization temperatures ofsome zircons may suggest that they were entrained during crystalfractionation from a hotter and more primitive melt (Claiborne et al.,2006; Harrison et al., 2007; Barth and Wooden, 2010). Finally, thiswide difference between zircon crystallization and saturation tem-peratures may be justified by the sensitivity of TTi-in-zrc to the activitiesof silica and rutile during zircon crystallization (Fu et al., 2008;Wielicki et al., 2012).

7.6. Hydrothermal versus late-magmatic origin

Although zircon commonly forms by the crystallization of magmas,mostly in felsic rocks, zircon precipitated from hydrothermal andmetamorphic fluids has fairly vast occurrences and distributions (e.g.Tomaschek et al., 2003; Cavosie et al., 2006; Kusiak et al., 2009; Somanet al., 2010; Sheng et al., 2012; Xu et al., 2012; Lisowiec et al., 2013).Hydrothermal zircon is a broad-spectrum term used to depict zirconcrystallized from or modified by an aqueous fluid (Whitehouse andKamber, 2002; Hoskin, 2005; Rayner et al., 2005; Pelleter et al., 2007;Fu et al., 2009; Bouvier et al., 2012). Although the term hydrothermalhas been loosely applied, it is crucial to differentiate between zirconsthat were precipitated from a circulating hydrothermal fluid inopposition to those recrystallized or have altered in the presence of a

the thermometer calibration ofWatson et al. (2006) against their Hf contents. Whole rockY-1 and KI-1 are indicated. The boundary value of maximum Ti abundance in unaltered


high water activity pore fluid (Cavosie et al., 2006; Fu et al., 2009).Hoskin (2005) has suggested three distinct mechanisms that canproduce hydrothermal zircon: (1) direct crystallization from zircon-saturated aqueous fluid; (2) ion-exchange between low-temperatureaqueous fluid and metamict zircon; (3) dissolution-reprecipitation. Itis important at this point to affirm that hydrothermal precipitationand dissolution-reprecipitation mechanisms should be excluded asneither field nor petrographic evidences were found to supportsubjection of the host rocks of the present zircons to apparent hydro-thermal activity. In the following paragraphs, we will discuss somecompositional criteria that may arbitrate the magmatic versus hydro-thermal origin of the analyzed zircons.

Numerous case studies (e.g. Geisler et al., 2003; Hoskin, 2005; Pettkeet al., 2005; Cavosie et al., 2006; Pelleter et al., 2007; Rimsa et al., 2007;Xia et al., 2010) have revealed that hydrothermal zircons display adistinct REE distribution pattern relative to that of igneous zircons. Forexample, hydrothermal processes would be a common cause for LREEenrichment in magmatic zircons. However, some studies found thatsome definitely hydrothermal zircons are indistinguishable frommagmatic zircons with regards to their chemical composition andisotopic signatures (e.g. Schaltegger, 2007; Fu et al., 2009). Despitedifferences in formation mechanisms, the hydrothermal precipitatedand hydrothermally-altered zircons share common characteristics asthey are LREE enriched and have flatter chondrite-normalized LREEpatterns (Low Sm/La)N) than magmatic zircon and they have smallerCe anomalies (Ce/Ce*) than magmatic zircon (Hoskin, 2005).

As previously demonstrated in Section 7.2, the LREE abundance ofstudied zircon samples exceeds the normal content in igneous zircon,indicating their obvious overabundance in LREE (Fig. 10). This excessiveoverenrichment in LREE (average (Sm/La)N=45.61 and 27.19 for BY-1and KI-1, respectively), generates noticeably flat LREE patterns for mostof the studied zircons in contrast to the negative steep slope LREEpatterns of igneous zircon ((Sm/La)N = 57–547) (Hoskin andSchaltegger, 2003). The possibly hydrothermally-related zircon grains,having (La/Sm)N ratio b 10, represent the majority of sample KI-1 (21grain out of 26) and many (9 zircons out of 20) of sample BY-1populations (Tables A2 and A3). These hydrothermally-suspectedzircons have slightly positive to even negative Ce/Ce* anomalies(≈0.5–10) relative to the rest of samples (Ce/Ce* up to 171) (Fig. 9,Tables A2 and A3). However, there is no obvious difference in the slopesof the normalized HREE patterns for all of the studied zircons from thetwo rock samples (cf. Hoskin, 2005) (Fig. 6a, b).

Herein, we adopt a new simple, but very likely efficient, plot of LaNvs. PrN to discriminate between unaltered magmatic and the LREE-enriched zircons of possible hydrothermal affinity (Fig. 15a). Therectilinear discrimination fields are based on the combination ofLaN N 1 and PrN N 10 as useful discriminators for identifying zircondomains with possibly hydrothermal LREE-enriched patterns (cf.Hoskin and Schaltegger, 2003; Cavosie et al., 2006). Plotting our zirconson this diagram shows that limited number of them (10 zircons) fall inthe field of unaltered magmatic zircons while most of them (28 out of46) plot in the hydrothermal zircon field, with the rest in a transitionalarea due to their high LaN ratio (N1). The aforementioned LREEcharacteristics indicates that aqueous fluids and/or hydrous meltsplayed a key role in the formation of these zircons, more particularlythose hosted in the Younger Granite sample (KI-1).

A further distinction between magmatic and hydrothermal zircons,based also on LREE characteristics, may be made using the (Sm/La)Nvs. La and Ce/Ce* vs. (Sm/La)N discrimination diagrams proposed byHoskin (2005). In these plots (Fig. 15b, c), the fields of magmatic andhydrothermal zircons do not wholly enclose all of the studied zirconsfrom the two rock samples, with many zircons plotted outside them.The (Sm/La)N vs. La plot (Fig. 15b) is not particularly robust, as mostof the zircons straddle the gap between the two fields. The Ce/Ce* vs.(Sm/La)N diagram provides more efficient discrimination, in whichthe magmatic zircon field is exclusively occupied with 11 zircons from

BY-1 and bordered by 3 zircons from KI-1, whereas that of hydro-thermal zircon is mainly occupied by KI-1 (17 out of 26) rather thanBY-1 (just 2) zircons. Once more, this outcome reinforce that aqueousfluids and/or hydrous melts have had more influential part in theformation of the Younger Granite zircons (sample KI-1) than of theDokhan Volcanics zircons (sample BY-1). Nonetheless, LREE abundanceand pattern alone appear inadequate for distinguishing magmatic fromhydrothermally altered igneous zircons.

Alteration in zircons has also been investigated using a plot of Caagainst U (Bouvier et al., 2012) (Fig. 15d). Calcium contents areextremely high in all of the analyzed zircon grains, while U valuesfluctuate within the lower half of the normal range of unalteredmagmatic zircon(Tables A2 and A3; Fig. 15d). Consequently, all of thestudied zircons do not plot in any of the three zircon fields (altered,unaltered and porous) because they have uranium abundances similarto unaltered igneous zircon, and in contrary vastly elevated Caconcentrations comparable to hydrothermally altered zircons (cf.Rayner et al., 2005; Grimes et al., 2009; Bouvier et al., 2012). Takinginto consideration that secondary alteration of zircon is commonlyaccompanied by coupled addition of Ca and Ba (Rayner et al., 2005),the narrow range of low Ba concentrations for BY-1 and KI-1 zircons(median = 1.38 and 0.60 ppm, respectively) and the non-correlationbetween Ca and Ba (not shown here), the subsequent hydrothermalalteration of the studied zircons is ruled out. The source of the dreadfullyhigh LA-ICP-MS measured Ca values of the investigated zircons isalmost certainly produced from inadvertent analysis of apatite micro-inclusions as indicated from the good positive correlation between Caand P (Fig. 8b).

Zircon crystallization temperature is a convenient parameter fordistinguishing hydrothermal zircon (generally crystallizing below500 °C) from igneous zircon (N600 °C) (Fu et al., 2009). The LA-ICP-MSanalyses of the present zircons gave relatively high Ti content, corre-sponding to high Ti-in-zircon crystallization temperatures (Tables A2and A3). These crystallization temperatures are obviously higherthan those of hydrothermal zircons, and in return fit well with thecrystallization temperatures of magmatic zircons appeared in literature.For example, crystallization temperatures for magmatic zircons fromigneous rocks of the Lachlan fold belt and adjacent New EnglandOrogen, Australia fall within the range of ≈750° to 950 °C (Hoskin,2005; Pettke et al., 2005; Fu et al., 2008, 2009; Ickert et al., 2011).Further, considering that the Ti-in-zircon temperatures are normallylower than predicted crystallization temperatures for magmatic zircons(Fu et al., 2008, 2009; Ickert et al., 2011), the determined Ti-in-zirconcrystallization temperatures from BY-1 and KI-1 zircons cannot berelated to hydrothermal origin. Therefore, the high Ti-in-zircon tem-peratures for zircons from the Younger Granite and Dokhan Volcanicsamples together with their comparable calculated zircon saturationtemperatures for the whole samples (Table A1) do not support thatthese zircons have formed from or altered by low-temperaturehydrothermal fluids.

The only criteria among all of the previously discussed that may hinttowards involvement of hydrothermal fluids in the genesis of many ofthe studied zircons are their flat-LREE patterns, smaller positive Ceanomalies and the higher LREE contents. These LREE features could beinterpreted as produced in late-crystallized zircons, from a melt with ahigher LREE/HREE value due to the earlier removal of zircon, feldspar,hornblend and biotite. These late-crystallized zircons might have beenequilibrated with liquids that are expected to be enriched in LREE,since these elements will be more incompatible than HREE. Thehigh mineral/melt partition coefficients of HREE in biotite (Nash andCrecraft, 1985) and hornblende (Sisson, 1994) cause their depletion inthemore evolved liquids, leading to flatter LREE patterns and increasingthe range of LREE in zircon grains (cf. Long et al., 2012; Nardi et al.,2012).

As a matter of fact, hydrothermal and magmatic zircons sharemany morphological and compositional features and thus are hardly

Fig. 15. (a) Distinction between unaltered magmatic zircon and LREE-enriched hydrothermal and late-magmatic zircon using a plot of (La)N versus (Pr)N. Discrimination diagrams of (b)(Sm/La)N vs. La (ppm); (c) Ce/Ce* vs. (Sm/La)N for zircons from samples BY-1 and KI-1. Color-shaded fields of “magmatic” and “hydrothermal” are from the Boggy Plain Zoned Pluton(Hoskin, 2005).(d) Plots of the analyzed zircons on the Ca ppm versus U ppm discrimination plot (after Bouvier et al., 2012). The field of porous zircons (dashed line delimited) is fromGrimes et al. (2009). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)


distinguishable fromeach other (Schaltegger, 2007). Thus in conclusion,we suggest that the analyzed zircons from samples BY-1 and KI-1 havebeen formed by equilibrium crystallization in a closedmagmatic systemthat has been progressively augmented with late-magmatic LREE-enriched fluids.

8. Conclusions

In order to contribute to the application of trace element geo-chemistry of zircon to igneous petrogenesis, we reported LA-ICP-MSanalyses of the trace element abundances of zircons separated fromthe Ediacaran (600–614Ma.) post-collisional Dokhan Volcanic sample(20 grains) and the Younger Granite sample (26 grains) from SE Sinaimassive. Younger Granites form huge expanses of batholith-sizedplutons and are mainly represented by subsolvous to transsolvousgranular to slightly porphyritic monzo- and syenogranites, whilethe Dokhan Volcanics consist of non-metamorphosed varicoloredalternating successions of porphyritic lava flows of commonly felsiccomposition (rhyolite–dacite) interlayeredwith compositionally equiv-alent ignimbritic layers. Whole-rock geochemistry of the DokhanVolcanics and Younger Granites hosting the analyzed zircons reflectstheir peraluminous, calc-alkaline and A-type character and points to

the generation of their magmas by partial melting of continental crustor underplated crust in a post-collisional regime. Zircons separatedfrom the Dokhan volcanic sample (BY-1) contain higher abundancesin Hf, ΣREE, Pb, Th and U relative to those from the Younger Granitesample (KI-1), which mirrors crystallization of the former at higherdegree of magmatic evolution. The REE patterns of the two zirconpopulations are characterized by a rather steeply-rising slope due toserious HREE enrichment relative to LREE and MREE with distinctivepositive Ce and negative Eu anomalies, which is typical of unalteredmagmatic zircons. Compared to unaltered magmatic zircons, mostof studied zircons display an evident LREE overabundance, whereasthe vast majority of the analyzed zircons have Th/U ratios≥0.5 commonin igneous zircons. Zircons from sample KI-1 exhibit more Eu depletionand deeper negative anomalies, whilst zircons of sample BY-1 havemarkedly higher Ce contents and more spiky positive anomaliesinferring wetter oxidized melt (Higher fO2) for Dokhan Volcanics. Noneof the common sources of LREE enrichment in zircon, including analysisof LREE-bearing inclusions, complex REE substitution mechanisms,radiation-induced lattice damage and hydrothermal alteration havesucceeded in explaining such extraordinary LREE overabundances.The relation between U/Yb ratio vs. Y and Hf contents suggestcrystallization of the analyzed zircons in continental crust. The application


of the Ti-in-zircon thermometer to the investigated zircons returnstemperatures of 790–986 °C (average TTi-in-zrc = 901 °C) for sampleBY-1 zircons and 781–93 °C (average TTi-in-zrc = 847 °C) for sampleKI-1 zircons. These crystallization temperatures are comparable tothe climax temperatures determined for zircons from analogousperaluminous rhyolites and monzo-syenogranites, which impliesdeep level of melting, likely within the lower crust of the ANS, forthe two host rock magmas. The zircon saturation temperatures forboth rock samples hosting the studied zircons are slightly lowerthan the onset temperature of zircon crystallization (≈ 40–55 °C),which is justified by the sensitivity of TTi-in-zrc to the activities ofsilica and rutile during zircon crystallization. LREE characteristicsindicates that aqueous fluids and/or hydrous melts played a keyrole in the formation of the analyzed two zircon populations, whilein contrary other geochemical trace element indicators (e.g. Ca, U,Ba contents) along with the high Ti-in-zircon temperatures for thesezircons rules out against either direct crystallization from or alterationby hydrothermal fluids. Therefore, the studied zircons from bothsamples are proposed to be formed by equilibrium crystallization in alate-magmatic closed system that has been progressively augmentedwith late orthomagmatic LREE-enriched fluids.

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.chemgeo.2013.10.009.

Acknowledgments

The authors thank Simon Wilde and Xin Zhou, respectively, forarranging and conducting the geochemical analyses of zircon at theInstitute of Geology and Geophysics, Chinese Academy of Sciences inBeijing. We gratefully acknowledge the thorough and constructivecomments of Dr. Yaron Be'eri-Shlevin and of an anonymous reviewerthat seriously improved the manuscript. Also, we greatly appreciatethe editorial comments and suggestions of Professor Klaus Mezger(Editor-in-Chief; Chemical Geology).

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Zircon trace element geochemical constraints on the evolution of the Ediacaran (600–614Ma) post-collisional Dokhan Volcanics and Younger Granites of SE Sinai, NE Arabian–Nubian