Provenance and metamorphic PT conditions of Cryogenian ...

TECTO-126806; No of Pages 19 Tectonophysics xxx (2015) xxx–xxx

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Provenance and metamorphic PT conditions of Cryogenian–Ediacaran metasediments from the Kid metamorphic complex, Sinai, NE Arabian–Nubian Shield: Insights from detrital zircon geochemistry and mineral chemistry Mohammed Z. El-Bialy a,⁎, Kamal A. Ali b, Mahrous M. Abu El-Enen c, Ahmed H. Ahmed b,d a

Geology Department, Faculty of Science, Port Said University, Port Said 42522, Egypt Department of Mineral Resources and Rocks, Faculty of Earth Sciences, King Abdulaziz University, P.O. Box 80206, Jeddah 21589, Saudi Arabia c Geology Department, Faculty of Science, Mansoura University, El-Mansoura, Egypt d Geology Department, Faculty of Science, Helwan University, Cairo, Egypt b

a r t i c l e

i n f o

Article history: Received 13 April 2015 Received in revised form 16 September 2015 Accepted 30 September 2015 Available online xxxx Keywords: Detrital zircon Zircon geochemistry Provenance Metamorphic PT conditions Arabian–Nubian Shield

a b s t r a c t The Malhaq and Um Zariq formations occupy the northern part of the Neoproterozoic Kid metamorphic complex of SE Sinai, NE Arabian–Nubian Shield. This study presents new mineral chemistry data and LA-ICP-MS analyses of the trace element concentrations on zircons separated from metapelites from these formations. The detrital zircons of Um Zariq Formation are more enriched in ΣREE, whereas Malhaq Formation zircons are markedly HREE-enriched with strongly fractionated HREE patterns. The quite differences in the overall slope and size of the Eu and Ce anomalies between REE patterns of the two zircon suites provide a robust indication of different sources. The Ti-in-zircon thermometer has revealed that the zircons separated from Malhaq Formation were crystallized within the 916–1018 °C range, while those from Um Zariq Formation exhibit higher range of crystallization temperatures (1084–1154 °C). The detrital zircons of Malhaq Formation were derived mainly from mafic source rocks (basalt and dolerite), whereas Um Zariq Formation zircons have varied and more evolved parent rocks. Most of the investigated zircons from both formations are concluded to be unaltered magmatic that were lately crystallized from a high LREE/HREE melt. All the studied detrital zircon grains show typical trace elements features of crustal-derived zircons. All of the Um Zariq Formation and most of Malhaq Formation detrital zircons are geochemically discriminated as continental zircons. Both formation metapelites record similar, overlapping peak metamorphic temperatures (537–602 °C and 550–579 °C, respectively), and pressures (3.83– 4.93 kbar and 3.69–4.07 kbar, respectively). The geothermal gradient, at the peak metamorphic conditions, was quite high (37–41 °C/km) corresponding to metamorphism at burial depth of 14–16 km. The peak regional metamorphism of Um Zariq and Malhaq formations is concluded to be generated during extensional regime and thinning of the lithosphere in an island arc setting with heat flow from the underlying arc granitoids. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The exposed Precambrian basement in Sinai (≈ 14,000 km2) constitutes the northwestern segment of the Arabian–Nubian Shield (ANS) that extends over most of NE Africa and the western part of the Arabian Peninsula (Fig. 1). The ANS, formed during the East African Orogeny (EAO; Stern, 1994), is regarded the best-preserved and largest exposed area of Neoproterozoic juvenile continental crust on Earth (Be'eri-Shlevin et al., 2009a, 2012; Hargrove et al., 2006; Stern, 2002; Stern et al., 2004). It is dominated by Neoproterozoic crust that was evolved over about 350 my (between 900 and 550 Ma ) through accretion of juvenile volcanic arc terranes and ophiolite remnants that were ⁎ Corresponding author. Tel.: +20 1223282650; fax: +20 663657601. E-mail address: [email protected] (M.Z. El-Bialy).

amalgamated during the assembly of the eastern part of Gondwana after the closure of the Mozambique Ocean (Avigad and Gvirtzman, 2009; Be'eri-Shlevin et al., 2009a; Cox et al., 2012; Eyal et al., 2014; Fritz et al., 2013; Johnson et al., 2011; Meert, 2003; Stern, 1994, 2002; Stern and Johnson, 2010; Stoeser and Frost, 2006). Prior to the ANS amalgamation (~650 Ma), deposition was dominated by volcano-sedimentary assemblages in volcanic arcs, whereas after 650 Ma deposition changed to volcano-sedimentary assemblages in post-amalgamation basins overlying newly amalgamated arc terranes (Johnson et al., 2011, 2013; Nasiri Bezenjani et al., 2014). Most of the still preserved Neoproterozoic volcano-sedimentary successions in the ANS, including the studied Malhaq Formation metasediments (615– 607 Ma; Moghazi et al., 2012), have deposited in late Cryogenian– Ediacaran (650–542 Ma) post-amalgamation basins (e.g., Breitkreuz et al., 2010; Eyal et al., 2014; Johnson et al., 2011, 2013; Moghazi et al.,

http://dx.doi.org/10.1016/j.tecto.2015.09.036 0040-1951/© 2015 Elsevier B.V. All rights reserved.

Please cite this article as: El-Bialy, M.Z., et al., Provenance and metamorphic PT conditions of Cryogenian–Ediacaran metasediments from the Kid metamorphic complex, Sinai, NE Arabian–Nubia..., Tectonophysics (2015), http://dx.doi.org/10.1016/j.tecto.2015.09.036

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M.Z. El-Bialy et al. / Tectonophysics xxx (2015) xxx–xxx

Fig. 1. Geological map of the central-northern Kid Metamorphic Complex and its environs in south Sinai (compiled and modified after El-Bialy, 2010, 2013). Location of the studied samples and their age, previously determined by Moghazi et al. (2012), are further indicated. The inset figure on the left is a map of the Sinai Peninsula showing the regional extent of the Neoproterozoic basement rocks. Darker reddish-brown areas represent the four metamorphic complexes of the southern Sinai: Kid, Feiran–Solaf, Sa'al-Zaghra and Taba. For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.

2012; Nasiri Bezenjani et al., 2014; Wilde and Youssef, 2000; Willis et al., 1988). The older middle to early Cryogenian (770–870 Ma) volcano-sedimentary assemblages, predating the ANS amalgamation, are more common in the northernmost ANS exposures in Sinai (e.g., Um Zariq Formation metasediments; Moghazi et al., 2012), southern Jordan (e.g., Abu-Baraka Schist: Jarrar et al., 1983) and Israel (e.g., Elat schist: Kröner et al., 1990; Eyal et al., 1991). The oldest known rocks belonging to these pre-amalgamation assemblages are the Mesoproterozoic (~1.1–0.95 Ga detrital zircon age) Ra'ayan Formation metasediments, Sa'al metamorphic complex, Sinai ((Be'eri-Shlevin et al., 2012; Andresen et al., 2014). The large spatial distribution of these Precambrian sedimentary successions across the northwestern terrains of the ANS occurs throughout Egypt (Sinai and Eastern Desert), Saudi Arabia and southern Jordan, and Israel, and their wide range of depositional ages could provide clues for interpretation of the fundamental processes that influenced the Neoproterozoic evolution of the ANS. The Neoproterozoic crustal evolution of Sinai (870–550 Ma; Bentor, 1985; Stern and Manton, 1987; Be'eri-Shlevin et al., 2009a,b, 2012; Eyal et al., 2010, 2014; Moghazi et al., 2012) included the formation of basins filled with volcano-sedimentary successions encompassed by subsequent granitic intrusions (Be'eri-Shlevin et al., 2011, 2012; Brooijmans et al., 2003; El-Bialy, 2013; El-Bialy and Ali, 2013; Eliwa et al., 2008; Moghazi et al., 2012). This geologically critical segment of the Neoprotoerozoic ANS comprises four main metamorphic complexes that are disrupted and isolated from each other by large tracts of granitoid intrusions (i.e., Kid, Taba, Feiran–Solaf, and Sa'al-Zaghra metamorphic complexes; Fig. 1). These metamorphic complexes comprise infracrustal ortho-gneisses and supracrustal meta-sedimentary and volcanic sequences that experienced polyphase deformation and greenschist to upper amphibolite facies metamorphism (Abu El-Enen, 2008, 2011; Abu El-Enen et al., 2004; Abu-Alam and Stüwe, 2009; Be'eri-Shlevin et al., 2012; Brooijmans et al., 2003; El-Bialy, 2013; Eliwa et al., 2008; Fowler et al., 2010a,b; Khalifa et al., 2011; Shimron, 1980, 1984). The southernmost Kid metamorphic complex (KMC) has

received considerable interest and investigations only in aspects of structure, metamorphism and tectonics. The KMC was divided by Shimron (1984, 1987) and Furnes et al. (1985) into four tectonstratigraphic units, namely: the Malhaq, Um Zariq, Heib and Tarr formations. The Malhaq and Um Zariq formations, the target of this study, occupy roughly the northern part of the KMC and are represented by mildly metamorphosed volcano-sedimentary and pelitic sedimentary sequences, respectively. El-Bialy (2013) has revealed that the metasediments of both formations are collectively geochemically immature, predominately derived from felsic to intermediate igneous sources and were originally deposited in a continental arc setting despite the considerable time span between their maximum depositional ages inferred from detrital zircons (Um Zariq Fm. = 813 ± 6 Ma, Malhaq Fm. = 615 ± 6 Ma; Moghazi et al., 2012). Zircon is an extraordinary mineral because of its omnipresent occurrence, existing not only in crustal rocks but also in mantle xenoliths, lunar rocks and meteorites (e.g., Demidova et al., 2014; Liati and Gebauer, 2002; Liu et al., 2010; Nikitina et al., 2012; Orejana et al., 2011; Page et al., 2007; Zheng et al., 2006). The importance of this accessory mineral lies in its unique chemical and physical durability and its remarkable resistance to weathering and recycling, eclogite/granulitefacies metamorphism, mantle storage and crustal anatexis (Belousova et al., 2002; Liu et al., 2010; Scherer et al., 2007; Siebel et al., 2012; Whitehouse and Platt, 2003; Zeck and Williams, 2002). Another reason for its geological significance arises from its tendency to incorporate many geochemically important trace elements (e.g., Sc, Y, Ti, Hf, Th, U, Nb, Ta, V, P, and REE). Zircon trace element geochemistry has proven to be a unique tool for evaluating the genesis of parental magmatic rocks (Barth and Wooden, 2010; Belousova et al., 2006; El-Bialy and Ali, 2013; Hanchar and van Westrenen, 2007; Hinton and Upton, 1991; Hoskin and Schaltegger, 2003; Nardi et al., 2012; Wang et al., 2012), provenance of sedimentary rocks (Barros et al., 2010; Belousova et al., 2002; Heaman et al., 1990; Hoskin and Ireland, 2000; Nardi et al., 2013; Wang et al., 2010; Xie



et al., 2012) and high-grade metamorphism (Baldwin et al., 2007; Hoskin and Black, 2000; Hoskin and Schaltegger, 2003; Jian et al., 2012; Peters et al., 2014; Rubatto, 2002; Slama et al., 2007; Turkina et al., 2012; Whitehouse and Platt, 2003; Wu et al., 2009; Yao et al., 2012). Clastic sediments and low-grade metasedimentary rocks preserve populations of detrital zircons because of the durability of this mineral (Dempster et al., 2004). Therefore, detrital zircon would be a potential powerful tool for the identification of source-rock lithology via its trace element composition (Belousova et al., 2002; Hoskin and Ireland, 2000). Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) has permitted highly sensitive elemental analysis to be performed in-situ on zircon and other accessory minerals. This work presents LA-ICP-MS analyses of major and trace (including REE) elements of detrital zircons from samples from the two spatially associated Ediacaran Malhaq and the Cryogenian Um Zariq formation metasediments, KMC, Sinai (MH-2 and UMZ-1, respectively; Fig. 1), which were previously analyzed for U–Pb geochronology and Sr–Nd isotopes by Moghazi et al. (2012). The results provide new insights into Precambrian crustal evolution of the ANS through determining sediment provenance, and improving our understanding of the regional geological events and the development of sedimentary basins during the Cryogenian–Ediacaran. Following the pioneer work of the first two authors on the use of zircon trace element geochemistry in elucidating the petrogenesis of igneous rocks from the ANS (El-Bialy and Ali, 2013), this contribution is the first application of detrital zircon geochemistry to provenance research in the ANS. Further, we have carried out extensive microprobe analyses for various minerals from the two studied metapelitic samples to estimate the metamorphic P-T conditions at the metamorphic peak using conventional geothermobarometry. 2. Geological setting The Malhaq and Um Zariq formations belong to the Neoproterozoic Kid metamorphic complex (KMC) which encompasses metamorphosed folded thick volcano-sedimentary successions and meta-plutonic rocks (Abu El-Enen, 2008; Abu El-Enen et al., 2003; El-Bialy, 2013; Fowler et al., 2010a,b; Furnes et al., 1985; Khalifa et al., 2011; Reymer et al., 1984; Shimron, 1980, 1984, 1987; Stern et al., 2010). The complex is extruded by the non-metamorphosed felsic to intermediate Dokhan Volcanics and intruded by granitoid rocks ranging in composition from quartz diorite to seynogranite (the Older and Younger Granites) (Fig. 1). Shimron (1984, 1987) and Furnes et al. (1985) have divided the Kid metamorphic complex sequences into Malhaq and Um Zariq formations (northern Kid area), and Heib and Tarr formations (southern Kid area). The total stratigraphic thickness of these four formations is argued between a maximum of 12,000 m (Shimron, 1987) and a minimum of about 1500 m (Blasband et al., 1997; Eliwa et al., 2008). These formations had underwent polyphase deformation and metamorphism ranging from greenschist to upper amphibolite facies (Abu El-Enen et al., 2003, 2004; Brooijmans et al., 2003; Eliwa et al., 2008; Fowler et al., 2010a,b; Reymer et al., 1984). The northern Malhaq and Um Zariq formations, the focus of this research, have undergone more deformation and are higher at metamorphic grade than the southern Heib and Tarr formations, exposed further southward of the mapped area of Fig. 1 (Abu El-Enen, 2008; Brooijmans et al., 2003; Eliwa et al., 2008; Khalifa et al., 2011). The Malhaq Formation is the northernmost volcano-sedimentary unit of the KMC (Fig. 1). It is a mixed succession comprising sequences of dark gray massive to schistose metavolcanics interbedded and intercalated with fine- to medium-grained foliated metasediments. The metasediments prevail in the northern exposures of this formation, while the metavolcanic rocks increase progressively towards the central and southern parts though still variably interbedded and intercalated with metasediments. In the latter case, both form centimeter- to meter-scale rhythmic layering of foliated metasedimentary phyllites

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and mica schists alternating with dark gray to black bands of lavas, metatuffs and other coarser metapyroclastics. The metavolcanic rocks are largely felsic to intermediate in composition, consisting of lavas and fine to coarse pyroclastics (i.e., breccias, lapillistone, lapilli tuff and finely bedded tuffs). The lavas are commonly structureless 0.5–4 m thick flows, infrequently associated with pillowed lava bodies. The Um Zariq Formation, located in the western central part of the KMC (Fig. 1), is an entirely metasedimentary sequence. It is essentially represented by well-bedded metapelitic schists along with occasional graphite-bearing phyllites. Minor intercalation of metapsammites and calcareous metapelites do occur within the abovementioned foliated metasediments and further form substantially thick beds (up to 10 m) in the lower horizons of the formation. The Heib Formation is a hybrid sequence consisting of abundant weakly metamorposed felsic to intermediate volcanics with subordinate sediments. The metasedimentary part of the succession comprises a sequence of turbiditic slate and greywacke (Beda Turbidites) and beds and lenses of cobble and boulder conglomerate with volcanogenic clasts (Kid Conglomerate). Exposed at the southern margin of the KMC, the Tarr Formation covers quite variable lithologies. It comprises lowgrade metamorphosed dacitic to andesitic lavas, ignimbrites, volcanic breccias and tuffs, mudstones and pebbly volcanogenic greywackes, and calcareous pelites (Fowler et al., 2010a; Khalifa et al., 2011). Further, the Tarr Formation involves minor intrusions of albitite (Azer et al., 2010). Geochronologically, the deposition age of the KMC sediments and volcanics was constrained at 615 ± 15 Ma (Bielski, 1982). Recently, Stern et al. (2010) have obtained a younger U–Pb zircon age of 598 ± 8 Ma for a volcanogenic clast from the Heib Formation. However, the detrital zircons used in this study, which were extracted from the two metapelite samples MH-2 (Malhaq Formation) and UMZ-1 (Um Zariq Formation) have yielded 206Pb/238U weighted mean ages of 615 ± 6 Ma and 813 ± 6 Ma, respectively (Moghazi et al., 2012). Along with the aforementioned metamorphosed volcanosedimentary formations, the KMC includes the meta-plutonic units Quneia Formation and Shahira metagabbro-diorite complex. The Quneia Formation rocks are exposed at the western and eastern margins of the KMC as irregular NE- to NW-trending masses (Fig. 1), and include strongly foliated diorites, tonalites and granodiorites. These gneissic rocks have been dated at 560–590 Ma using Rb–Sr techniques (Moghazi et al., 1998) and at 580–595 Ma using U–Pb zircon geochronology (Ali et al., 2009). The Shahira metagabbro-diorite complex crops out as a large intrusive NE-trending body (~ 75 km2) along the northern margin of the KMC (Fig. 1). Recently, the Shahira complex rocks have yielded a U–Pb zircon age of 632 ± 4 Ma (Be'eri-Shlevin et al., 2009a). The aforementioned KMC rock units are bound to the north, west and south and both intruded and extruded by unmetamorphosed late Neoproterozoic magmatic rocks. These younger rocks include the synorogenic Older Granites and the post-collisional Dokhan Volcanics and Younger Granites. Older Granites (650–630 Ma: Stern and Manton, 1987; Moussa et al., 2008) range in composition from quartz diorite through quartz monzodiorite to tonalite, and in texture from granular to porphyritic and gneissose-textured and commonly enclose ovoid to sub-rounded microgranular mafic enclaves of microgabbro or melanocratic microdiorite composition (El-Bialy, 2004). The Dokhan volcanics exposed in the mapped area (609 ± 10 Ma; Bielski, 1982; Moghazi et al., 2012), consist of non-metamorphosed varicolored alternating succession of porphyritic lava flows of dominantly rhyolite– dacite composition, intercalated with compositionally corresponding pyroclastic layers (commonly ignimbrites). Younger Granites (~ 602– 612 Ma; Be'eri-Shlevin et al., 2009a; Moghazi et al., 2012) represent the concluding major magmatic activity in the region, encompassing the KMC rocks except from east and extending edgewise outside the mapped area. They constitute vast expanses of large plutons and are distinguished in the field into biotite monzogranites and leuco-


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syenogranites with blurred gradational mutual contacts. Further, they vary in texture from fine- to coarse-grained granular to porphyritic with pink alkali feldspar megacrysts. 3. Sample description and analytical methods Two metapelites, one from Malhaq Formation (MH-2) and another from Um Zariq Formation (UMZ-1), from the KMC were selected for studying their detrital zircon geochemistry and mineral chemistry. The metapelite sample UMZ-1 was collected from Wadi Um Zariq (28° 18′ 27.6″ N; 34° 15′ 09.1″ E). It is highly foliated medium-grained schist, showing obvious compositional layering of andalusite-rich and andalusite-poor bands (Fig. 2a). It is characterized by the occurrence of equilibrium peak and post-peak assemblage garnet + staurolite + andalusite porphyroblasts in a matrix of biotite, muscovite, quartz and plagioclase (Fig. 2c). The metapelite sample MH-2 was collected from the middle course of Wadi Malhaq (28° 18′ 59.4″ N; 34° 19′ 23″ E), where alternating layers of metaplites and metapyroclastic rocks crop out (Fig. 2b). It is fine-grained, highly foliated garnet phyllite, consisting of muscovite–garnet–biotite–quartz–plagioclase assemblage with minor tourmaline and magnetite (Fig. 2d). Minerals were analyzed in the two investigated samples using a JEOL electron-probe micro-analyzer JXA-8200 at the Faculty of Geosciences, King Abdalaziz University, Saudi Arabia. The chemical compositions of the identified silicate and oxide minerals were performed on polished thin sections. Operating conditions were 15 kV accelerating voltage, 20 nA probe current, and 3 μm probe diameter. The raw data were corrected with an online ZAF program. The following standards were used: quartz for Si, eskolite for Cr, fayalite for Fe, wollastonite for Ca,

corundum for Al, periclase for Mg, manganosite for Mn, jadeite for Na, orthoclase for K, and nickel oxide for Ni. Representative mineral assemblages are given in Table (1). In addition, the full electron microprobe data are presented in Appendices A and B (online supplementary materials). Zircons were separated at the University of Texas at Dallas (UTD) using standard crushing, heavy liquid 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 in half and then they were imaged by cathodoluminescence (CL) using a scanning electron microscope prior to gold coating. In situ trace element measurements were performed by LA-ICP-MS using an Agilent 7500a Q-ICPMS connected to a 193 nm Excimer laser ablation system at the Institute 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 fifty-one zircons (26 from sample MH-2 and 25 from sample UMZ-1). The isotopes free from isobaric interference were selected for measurements of trace elements. Standard silicate glass NIST SRM610 was used as external standard for the concentration of trace elements in conjunction with the internal standardization using 29Si (32.8 SiO2 in zircon). Oxide and molecular interference 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 %). High grade-argon gas, carrying away the ablated materials into the mass spectrometer, was measured twice to establish a blank prior to starting each analysis. Limits of detection (LOD) are typically better than 10 to 30 ppb for REE, Nb, Ta, Ba, Hf, Th, and U; 0.1 to 0.3 ppm for Rb, Sr, and Y; and 5 to 10 ppm for P.

Fig. 2. (a) Close view of the outcrop surface of sample UMZ-1 sampling site showing high foliation and compositional layering of andalusite-rich and andalusite-poor metapelites, Um Zariq Formation.( b) Alternating bands of metapelites–metapyroclastics from the sampling site of sample MH-2, Malhaq Formation. (c) Photomicrograph of sample UMZ-1 showing grt + and + st + bt + ms + qtz mineral assemblage. Staurolite is enclosed in a larger andalusite poikiloblast. (d) Photomicrograph of sample MH-2 showing subidioblastic garnet in matrix of ms + bt + turn + qtz + pl + mag assemblage.



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Table 1 Representative electron microprobe analyses (wt.%) of minerals from the metapelites of Um Zariq (sample UMZ-1) and Malhaq (sample MH-2) formations. Mineral

Garnet

Sample

UMZ-1

Mineral

Biotite

Sample

UMZ-1

MH-2

UMZ-1

MH-2

Sample

Site

Rim

Core

Rim

UMZ-1

Core

Site

Mantle

Mantle

Mantle

Mantle

Site

Rim

Core

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MgO CaO MnO Total Cations Si Ti Al Cr Fe3+ Mg Ca Mn Fe2+ Xfe End members Adr Grs Alm Sps Pyp

37.87 0.06 21.27 0.00 1.26 34.41 1.75 2.11 2.05 100.78

37.57 0.16 20.97 0.00 0.99 26.91 1.02 2.81 9.39 99.82

37.60 0.06 21.16 0.00 1.06 28.77 1.24 1.78 8.87 100.54

37.41 0.26 20.77 0.00 1.03 27.94 1.41 1.66 9.96 100.44

34.99 2.31 19.52 0.00 23.86 6.50 0.04 0.00 0.14 8.64 95.98

34.69 1.83 18.22 0.01 25.38 6.75 0.03 0.00 0.13 8.70 95.73

46.00 0.34 35.75 0.75 0.28 0.46 8.31 0.08 0.02 0.00 91.99

45.81 0.24 34.17 2.80 0.45 0.39 9.03 0.04 0.01 0.03 92.99

27.28 0.56 53.96 0.01 12.73 0.93 0.001 0.03 0.01 95.51

6.05 0.01 4.01 0.00 0.13 0.30 0.31 1.21 3.87 0.93

6.04 0.03 0.95 0.00 0.13 0.34 0.29 1.36 3.77 0.92

6.24 0.02 5.48 0.09 0.01 0.00 0.32 0.10 1.57 0.78

3.89 0.07 8.91 0.00 1.50 0.00 0.21 0.00 0.01 0.00 0.88

3.84 0.03 9.02 0.01 1.52 0.00 0.18 0.00 0.02 0.00 0.89

3.57 1.83 68.10 21.27 5.23

4.07 0.90 65.49 23.65 5.98

5.40 0.21 3.34 1.57 0.01 0.00 3.30 0.04 1.73 0.68 0.57 0.27

6.22 0.03 5.70 0.06 0.01 0.00 0.09 0.12 1.43 0.60

3.70 4.90 64.32 22.73 4.35

5.37 0.27 3.53 1.49 0.01 0.00 3.06 0.04 1.69 0.67 0.54 0.26

SiO2 TiO2 Al2O3 Cr2O3 FeO MgO CaO Na2O K2O Total Cations Si Ti Al Cr Fe2+ Mn Mg Ca Na K Xfe

27.53 0.68 53.59 0.01 12.71 1.00 0.00 0.02 0.00 95.55

6.07 0.02 3.99 0.00 0.12 2.46 0.49 1.29 3.64 0.94

SiO2 TiO2 Al2O3 Cr2O3 FeO MgO CaO MnO Na2O K2O Total Cations Si Ti Al Mg Ca Mn Fe Ba Na K XFe XAnn XPhl XMs XPg

6.05 0.01 4.00 0.00 0.15 0.42 0.36 0.28 4.60 0.92 4.22 2.17 81.33 4.91 7.37

0.84 0.04

0.75 0.05

MH-2

Mineral

Feldspar

Sample

UMZ-1

Site

Rim

Core

Rim

SiO2 Al2O3 FeO CaO Na2O K2O BaO Total Cations Si Al Mg Ca Fe Ba Na K Total End members An Ab Or

60.06 24.62 0.11 6.25 7.91 0.05 0.00 99.00

60.29 24.86 0.04 6.51 7.75 0.06 0.00 99.51

2.70 1.30 0.00 0.30 0.00 0.00 0.69 0.00 5.00 30.31 69.41 0.28

Muscovite

Mineral

Saturolite

Mineral

Chlorite

Mineral

Tourmaline

Mineral

Ilmenite

Magnetite

Sample

MH-2

Sample

MH-2

Sample

UMZ-1

MH-2

Core

Site

Mantle

Site

Mantle

Site

Mantle

Mantle

61.92 23.31 0.14 4.28 9.12 0.18 0.07 99.03

61.61 23.31 0.05 4.48 9.01 0.28 0.03 98.78 2.77 1.23 0.00 0.22 0.00 0.00 0.78 0.02 5.02

36.13 0.61 30.81 3.19 7.58 5.17 0.94 1.85 0.05 10.42 0.15 96.89

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MgO Total Cations Si Ti Al Cr Fe + 3 Fe + 2 Mn Mg

0.03 12.26 0.03 0.01 0.00 11.60 0.00 0.08

0.03 0.03 0.01 0.01 15.87 8.04 0.00 0.01

31.60 68.06 0.35

20.36 78.65 0.99

21.26 77.17 1.57

SiO2 TiO2 Al2O3 Fe2O3 FeO MgO CaO Na2O K2O B2O3* Li2O* Total Cations Si Ti B Al Mg Ca Fe3+ Fe2+ Li Na K Vac XMg

0.09 0.12 0.03 0.04 68.45 31.22 0.03 99.98

2.77 1.23 0.00 0.21 0.01 0.00 0.79 0.01 5.02

27.31 0.61 20.05 0.01 1.77 24.62 9.49 0.06 0.14 2.16 86.21

0.09 53.24 0.07 0.04 0.00 45.34 0.17 98.95

2.69 1.31 0.00 0.31 0.00 0.00 0.67 0.00 4.99

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MgO CaO Na2O K2O Total Cations Si Al Ti Cr Fe3+ Fe2+ Mn Mg Ca Na K

MH-2

5.85 5.14 0.10 0.00 0.28 4.41 0.00 3.03 0.01 0.11 1.18

6.03 0.08 3.00 6.06 1.29 0.17 0.40 1.06 0.10 0.60 0.01 0.23 0.47

Cation normalization: garnet = 24 oxygen; biotite & muscovite = 22 oxygen; staurolite = 23 oxygen; feldspar = 8 oxygen; chlorite = 36 oxygen; tourmaline = 24.5 oxygen; ilmenite = 6 oxygen; magnetite = 32 oxygen.

4. Results 4.1. Mineral chemistry Garnet varies considerably in the two investigated samples, but Almandine is the main end-member (Table 1). Garnets of sample UMZ-1 are characterized by strong growth zoning, defined by sharp decreases of Mn-content and increase of Mg- and Fe contents from cores

to rims (Table 1). It has wide compositional range of Alm65.7–82.2 Prp4.1–7.6 Grs6.2–8.9 with slightly more variable Sps3.5–22.2 and XFe = Fe/ (Fe + Mg) values in the range of 0.91–0.94. On the other hand, garnet of sample MH-2 exhibits weak growth zoning and narrow compositional range of Alm66.4–70.5 Prp5–6 Grs3.9–6.5 Sps19.2–23 with XFe values in the range of 0.92–0.93. Biotite and muscovite are analyzed from the peak assemblages, and are individually uniform. Biotite of sample UMZ-1 is iron-rich (XFe =


6


0.67–0.72) and is characterized by moderate TiO2 content (2.07– 3.29 wt.%, averages 2.50 wt.%), which imply affiliation to the garnet and staurolite zones of amphibolite facies metamorphism (Guidotti, 1984; Henry and Guidotti, 2002). Biotite of sample MH-2 has slightly lower Fe-content (XFe = 0.67–0.68) and TiO2 content (1.73–1.94, average = 1.82 wt.%), indicating lower metamorphic condition relative to sample UMZ-1. Muscovite of sample UMZ-1 has low celadonite (0.03–0.05), higher XMs (0.83–0.84), lower XPg (0.07) and Si in the range 6.22–6.26 (per 22 oxygen). On the other hand, muscovite of sample MH-2 has low XCel (0.09–0.10), high XMs (0.74–0.76), lower XPg (0.04–0.05) and Si in the range 6.20–6.30 apfu. Chlorite is reported in sample MH-2 as a retrograde phase with XFe = 0.61 and Si content being in the range of 5.85–5.93. It plots in the brunsvigite field in the classification diagram of Hey (1954) (not shown). Staurolite analyzed from sample UMZ-1 has no detectable MnO with rather low TiO2 content within the narrow 0.56–0.68 wt.% range from core to rim, respectively. The staurolite is almost unzoned and remarkably iron-rich, with XFe in the range 0.87–0.89. Andalusite is nearly pure Al2SiO5, and has trivial contents of FeO (0.15–0.37 wt.%) and Cr2O3 (0.03–0.09 wt.%). Tourmaline analyzed from sample MH-2 is aluminous schorl-dravite with XMg = 0.47–0.53 and Ca/(Ca + Na) = 0.18–0.29. Tourmaline XMg values imply metamorphism in the upper greenschist facies (Abu El-Enen and Okrusch, 2007). Feldspar in both samples is represented by slightly oscillatory zoned plagioclase with anorthite-rich cores, oscillatory zoned mantles and low-anorthite rims (Table 1). Plagioclase of sample UMZ-1 is andesine (An30–35) with difference between individual grains not exceeding 3 mole%, while it is of oligoclase composition (An20–24) in sample MH-2. Opaque ilmenite and magnetite exist in UMZ-1 and MH-2, respectively, and are close to the ideal member composition with noticeable MgO contents (0.11–0.17 wt.%) in the former. 4.2. Cathodoluminescence imaging Cathodoluminescence (CL) imaging is regularly used to interpret zircons prior to isotope and trace element analyses. Herein, CL imaging was used for spot selection during trace element analyses on the La-

ICP-MS, which ablates a ~30 μm spot on each zircon grain (Fig. 3). Representative CL images of some of the detrital zircon grains from samples UMZ-1 and MH-2 (some have been analyzed, others not) are presented in Fig. 3a and b, respectively. Straight to curved bright Cl fractures, more frequent in zircons of UMZ-1 sample, are possibly consequences of internal stress caused by differential swelling during metamictization (e.g., Geisler et al., 2003). Zircons extracted from the metapelitic schist sample UMZ-1 are transparent and colorless to pale brown in color. However, zircons separated from this sample are variable in terms of their shape (long or short prisms) and size (50–200 μm), form (euhedral vs. rounded) and inclusion content. Such wide variations, perhaps suggest that they were derived from different sources. In spite of their textural variability, the majority of the zircons are euhedral, whereas rounded to sub-rounded grains are less common. With few exceptions (e.g., grains #9 and 25), the grain interiors exhibit very low CL intensity and are either almost homogenous or display slightly chaotic pattern. The darker CL cores are either surrounded by thin discontinuous highluminescent rims, or mantled by thicker blurred oscillatory-zoned rims (e.g., zircons #10 and 21). These rims may represent new zircon overgrowths during a late low-grade metamorphic event (e.g., Dempster et al., 2004; Siebel et al., 2012; Wang et al., 2010, 2012). On the other hand, there does exist a characteristic population of euhedral short prismatic multi-faceted zircons showing strong core-to-rim oscillatory zoning of unquestionable magmatic origin (e.g., zircons #22 and 14) (Barth and Wooden, 2010; Cavosie et al., 2006; Hoskin and Schaltegger, 2003). Zircons separated from the metapelitic schist sample MH-2 are mostly euhedral to subhedral stubby to long prismatic and pale brown in color. The zircon grains of this sample are quite larger and tend to be more euhedral, compared with those of UMZ-1, ranging from 80 to 250 μm in length, with an aspect ratio of 1:1–3:1. This may point to their relative lesser degree of recycling and maturation. CL images occasionally reveal core-rim features with homogeneous cores mantled by oscillatory-zoned rims, but some grains are unzoned. In a few grains, the internal parts contain unidentified dark inclusions, which were avoided during selection of spot analysis.

Fig. 3. Cathodoluminescence (CL) images of some detrital zircons from the two studied pelitic schist samples. (a) Um Zariq Formation sample (UMZ-1). (b) Malhaq Formation sample (MH-2). Locations of the LA-ICP-MS spots (30 μm) are shown by white dotted circles and zircon sample numbers (sp) are also indicated.



4.3. Zircon geochemistry Fifty-one zircon grains were analyzed from samples MH-2 (n = 26) and UMZ-1 (n = 25). The major (Si, Zr and Hf) and trace element (including REE) abundances were in-situ analyzed at only one spot in each zircon grain, and the entire data are presented in Tables 2 and 3. Chondrite-normalized (McDonough and Sun, 1995) spider diagrams and REE patterns of the analyzed zircon samples are shown in Figs. 5 and 6, respectively. Zircon is the main reservoir for Zr and Hf, and consequently the abundances of both elements are apparently correlated in all of the analyzed zircons, accounting for concentrations corresponding to several ten thousand times of their values in chondrites (Fig. 4). Tables 2 and 3 reveal considerable variation in the Zr/Hf ratios of zircons separated from both samples (UMZ-1; 35.2–51.3, average = 42 and MH-2; 38.1–70.6, average = 59.5), with noticeable deviation towards higher values than the chondritic Zr/Hf one (34.2 ± 0.3; Weyer et al., 2002). The detrital zircons from Um Zariq Formation (sampleUMZ-1) and Malhaq Formation (sample MH-2) exhibit considerable overlap between their REE abundances (Fig. 5c). However, the REE contents of zircons from sample UMZ-1 range from 494 ppm to 2769 ppm (average = 1553 ppm), while zircons from sample MH-2 are less enriched with contents ranging between 268 ppm and 1945 ppm (average = 766 ppm) (Tables 2 and 3). Although the chondrite-normalized REE patterns of both zircon groups generally show quite steeply rising slope REE patterns of low LREE/HREE and noticeable positive Ce and negative Eu anomalies of variable magnitudes, which are typical of magmatic zircons (e.g., Belousova et al., 2002; Hinton and Upton, 1991; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003; Whitehouse and Kamber, 2002), there are some profound dissimilarities in between (Fig. 5) which may reflect differences in source or subsequent modification by post-magmatic processes (e.g., hydrothermal alteration and metamorphism). Even though several zircons recovered from the metapelite sample MH-2 display quite gentle REE patterns (very low YbN/LaN ratios down to 2.55), the overall steepness of the REE patterns of zircons from Malhaq metasediments is greater than that of the Um Zariq metasediments as revealed from the twice higher average YbN/LaN ratio of the former (averages; 2377 and 1070 respectively) (Tables 2 and 3; Fig. 5a–c). Zircons from the Malhaq and Um Zariq formation samples have strongly enriched and somewhat uniform HREE concentrations (average YbN = 2336, 1534; average LuN = 2788, 2038 respectively). Nevertheless, Malhaq Formation zircons are markedly HREEenriched (Fig. 5c) and display strongly fractionated HREE patterns (average YbN/GdN = 250) relative to the moderately fractioned HREE patterns of Um Zariq Formation zircons ((average YbN/GdN = 87). On the other hand, the LREE (La-Nd) and MREE (Sm-Tb) exhibit disparate patterns in both rock samples and more specifically in MH-2 in which the zircon LREE profiles vary between steeply positive and nearly flat with extreme divergence in the SmN/LaN ratio between a minimum of 0.71 and a maximum of 302 (Table 3; Fig. 5b). Compared to chondritic values, many of the analyzed zircons from both samples show an obvious overabundance in Light REEs. This feature is clear in zircons from the sample MH-2 whose normalized La values (LaN) surpass the normal range of igneous zircons (≤ 10 ×; Hoskin and Schaltegger, 2003) with values up to 693 (14 zircons out of 26), resulting in an extraordinarily elevated average for all of these sample zircons (Av. LaN = 124) (Table 3). Although Cerium is typically high in igneous zircons leading to its common positive anomaly, several zircons from sample MH-2 possess prodigious normalized values surpassing the normal limit of 100 × chondritic abundance (Hoskin and Schaltegger, 2003) as indicated in Table 3 and Fig. 5b. The serious LREE enrichment and the flatter LREE pattern of lots of Malhaq Formation detrital zircons can be attributed to the alteration by fluids since these attributes are typical of hydrothermally altered zircons (e.g., El-Bialy and Ali, 2013; Fu et al., 2009; Hoskin, 2005; Pettke et al.,

7

2005; Rayner et al., 2005; Whitehouse and Kamber, 2002; Xia et al., 2010). Alternatively, Heavy REEs in almost all the studied zircons vary within the typical range of 103 and 104 chondrite (Barth and Wooden, 2010; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003; Poller et al., 2001) (Fig. 5; Tables 2 and 3). Yttrium, which behaves similar to the HREE, has chondrite-normalized abundances similar to them in all zircon samples (Fig. 4). Although zircons from sample MH-2 are more enriched in Y (623–3734; average = 1840 ppm) relative to those of sample UMZ-1 (417–2824; average = 1005 ppm), the Y concentrations obtained in this study fall within the range of crustal zircons, in opposition to the mantle-derived zircons that have very low Y contents (Belousova et al., 2002, 2006; Hoskin and Schaltegger, 2003; Poller et al., 2001). Without exception, the zircons extracted from Um Zariq metasedimentary sample UMZ-1 display positive Ce anomalies of variable magnitudes (2.29–24.52; average = 9.1) (Table 2). Alternatively, the LREE-enriched detrital zircons (9 out of 26 zircon samples) of Malhaq Formation sample (SmN/LaN = 0.71–1.0) show negligible to no Ce anomaly (Ce/Ce* = 0.97–1.32) (Table 3). Generally, such a characteristic results from the recycling of crustal material, a process that leads to the enrichment of the other light REEs and consequently the reduction or even declination of the positive Ce anomaly (e.g., Peck et al., 2001; Yao et al., 2012). The rest of zircons from this sample exhibit variably sized positive Ce anomalies (Ce/Ce* = 1.8–39.42). In both samples, it is notable that the size of the positive Ce anomaly is inversely correlated with the degree of LREE enrichment (Fig. 5a, b) (cf. El-Bialy and Ali, 2013; Pettke et al., 2005). Positive Ce anomalies may be linked to favorable incorporation of small quantities of Ce4+ from a relatively oxidized melt (Belousova et al., 2002; Hoskin and Schaltegger, 2003). Almost all the zircons analyzed show negative Eu anomalies of variable sizes. Zircons that lack a negative or have positive Eu anomaly have been explained as having crystallized in plagioclase-absent assemblages (Belousova et al., 2002). Detrital zircons from sample MH-2 show restricted range of small Eu/Eu*values (0.19–0.53, average = 0.32), thus displaying pronounced negative anomalies (Fig. 5b), while zircons of UMZ-1exhibit wider spectrum of shallow and deep negative Eu anomalies (0.04–0.60; Fig. 5a). The negative Eu anomalies in zircon REE profiles imply zircon crystallization from Eu-depleted magma due to preceding or simulations plagioclase fractionation (Belousova et al., 2002; Hoskin and Schaltegger, 2003; Hoskin et al., 2000; Wang et al., 2012). Zircons from the Um Zariq metapelites (sample UMZ-1) have significantly high concentrations in Pb, Th and U (averages; 36, 100 and 239 ppm, respectively) compared with zircons from the Malhaq Formation metapelites (averages; 11, 42 and 82 ppm, respectively) (Tables 2 and 3). Unusually high Pb amounts in zircon samples UMZ-1.07 and UMZ-1.17 (167 and 114 ppm, respectively) reflect the radioactive decay of U and Th and imply metamectization of these grains which are already the richest in both radioactive elements among other zircons (Table 2). The perfect proportional linear relation between lead versus U and Th in the investigated zircons (Fig. 6a) may indicate that the Pb occurring in zircon is largely radiogenic nuclide. The zircons extracted from each rock sample show considerable large inter-grain compositional variation in U, Th and Th/U ratio that may occasionally exceed one order of magnitude (e.g., U in sample UMZ-1 zircons; 37– 1144 ppm). Such variation is a characteristic feature of most zircons (Ahrens et al., 1967; Belousova et al., 1998, 2002; Heaman et al., 1990; Hidaka et al., 2002; Hoskin and Ireland, 2000; Pettke et al., 2005). The high-field strength elements (HFSE) Nb, Ta and Ti, substituting for Zr (Hoskin and Schaltegger, 2003), are found in significant detectable amounts in all the analyzed zircons. While zircons from sample MH-2 show constrained range of normal Ti contents (60–119; average = 83 ppm), those of sample UMZ-1 contain exceptionally high Ti abundances (184–282; average = 228 ppm) that seriously exceed the normal abundance of Ti in zircon (≤ 75 ppm; Hoskin and Schaltegger, 2003), which accordingly has yielded very high Ti-in-


8

Sample

Si

P

Ca

Ti

Cr

Rb

Sr

Y

Zr

Nb

Ba

Hf

Ta

Pb

Th

U

Ti-in-Zr C

Th/U

Zr/Hf

Nb/Ta

Y/Ho

U/Yb

La

Ce

Pr

UMZ-1 01 UMZ-1 02 UMZ-1 03 UMZ-1 04 UMZ-1 05 UMZ-1 06 UMZ-1 07 UMZ-1 08 UMZ-1 09 UMZ-1 10 UMZ-1 11 UMZ-1 12 UMZ-1 13 UMZ-1 14 UMZ-1 15 UMZ-1 16 UMZ-1 17 UMZ-1 18 UMZ-1 19 UMZ-1 20 UMZ-1 21 UMZ-1 22 UMZ-1 23 UMZ-1 24 UMZ-1 25 Average

153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225

1029 1216 1231 1383 1478 1558 1561 1831 1310 1924 1849 1793 1789 1508 1809 2236 1684 1957 1907 1860 2612 2007 2672 2226 2727 1806

50,390 60,015 51,798 55,272 70,475 66,696 57,521 69,654 55,413 69,031 63,320 65,120 53,906 54,672 69,730 73,169 56,261 59,283 55,133 62,873 66,564 53,829 55,933 47,165 59,728 60,118

184 213 221 219 282 245 230 246 227 270 246 281 210 218 224 272 187 195 230 205 279 191 210 202 211 228

249 232 237 323 304 399 329 311 263 327 291 387 271 191 324 342 308 261 239 401 310 298 301 333 283 301

3.1 3.3 4.1 4.0 4.6 4.3 3.7 4.8 3.9 5.0 4.3 4.6 4.0 3.9 4.7 4.4 3.8 4.0 3.6 4.6 4.3 4.1 4.5 3.8 4.0 4.1

2.8 1.2 1.1 1.1 1.5 1.8 1.6 2.6 1.3 2.0 5.0 3.0 4.4 1.3 1.4 4.4 2.2 1.3 1.3 1.6 2.0 1.5 1.3 1.1 1.4 2.0

814 2485 2824 731 1851 354 328 1793 1752 638 520 547 434 1088 805 786 1363 957 1074 306 630 417 602 1348 668 1005

438,587 414,390 432,800 401,444 430,356 413,843 375,141 398,824 414,236 424,437 382,017 414,101 397,810 357,123 366,849 389,789 395,814 346,089 372,688 427,428 413,010 361,397 410,001 368,622 375,354 396,886

1.1 0.7 0.8 2.7 5.4 0.8 1.6 2.4 6.0 0.7 0.7 0.6 0.8 2.1 0.8 2.8 3.0 5.6 1.2 0.6 0.7 0.6 0.7 0.7 0.6 1.74

1.60 1.42 1.76 1.78 2.66 30.22 2.17 5.59 1.64 2.20 1.82 2.46 1.60 1.47 2.05 2.10 1.69 1.56 1.71 1.93 2.16 5.86 2.26 1.59 1.45 3.31

10,141 11,100 10,963 9338 9478 10,054 10,658 7541 8687 9827 8984 9785 9574 9312 8471 10,212 9945 6744 8720 11,242 9693 9303 10,497 9732 9686 9588

0.4 0.3 0.3 1.2 1.5 0.2 0.9 0.5 1.4 0.3 0.3 0.3 0.3 0.5 0.3 0.6 1.0 1.3 0.3 0.2 0.2 0.2 0.4 0.2 0.2 0.53

41.1 35.3 36.1 42.0 53.7 20.9 167.1 38.8 7.1 21.2 30.9 17.7 23.2 42.9 19.2 51.3 114.3 4.3 15.8 11.1 14.2 20.1 22.9 21.4 22.4 36

202 31 71 221 145 73 61 198 36 98 140 68 87 182 51 183 166 13 67 41 61 51 71 90 98 100

251 261 263 230 346 131 1144 228 75 129 176 110 147 252 121 399 851 37 94 74 88 151 152 134 133 239

1084 1107 1113 1112 1154 1130 1119 1131 1118 1147 1131 1153 1105 1111 1115 1148 1086 1093 1119 1100 1152 1089 1105 1098 1105 1117

0.80 0.12 0.27 0.96 0.42 0.56 0.05 0.87 0.47 0.76 0.80 0.62 0.59 0.72 0.42 0.46 0.20 0.35 0.71 0.55 0.70 0.34 0.47 0.67 0.74 0.54

43.25 37.33 39.48 42.99 45.40 41.16 35.20 52.89 47.68 43.19 42.52 42.32 41.55 38.35 43.31 38.17 39.80 51.32 42.74 38.02 42.61 38.85 39.06 37.88 38.75 41.75

2.58 2.23 2.53 2.15 3.63 3.38 1.67 5.29 4.13 2.35 2.64 2.36 2.86 4.64 2.50 4.60 2.87 4.34 3.46 4.27 3.32 3.59 1.72 2.83 2.67 3.14

30.49 29.06 28.09 29.80 28.63 31.61 35.66 29.51 31.77 29.34 32.45 29.84 31.94 30.09 30.05 32.37 29.79 27.70 29.61 30.51 30.46 32.68 31.19 29.38 30.14 30.49

0.85 0.34 0.33 0.98 0.60 0.97 6.35 0.39 0.19 0.60 0.86 0.57 0.85 0.79 0.43 1.31 2.00 0.14 0.23 0.60 0.43 0.97 0.75 0.40 0.57 0.90

3.91 0.25 0.25 0.26 0.35 0.19 1.61 2.12 9.44 0.85 2.84 0.70 11.85 0.62 0.21 8.90 3.02 2.39 0.79 0.30 0.27 0.40 0.21 0.20 0.54 2.10

46.91 3.52 2.24 28.13 30.71 14.67 33.10 61.93 67.81 18.78 22.50 17.07 38.43 44.90 7.34 40.92 59.68 15.39 12.78 9.76 12.95 6.90 13.77 12.51 16.47 25.57

2.10 0.18 0.22 0.29 0.60 0.18 1.56 0.93 4.19 0.33 0.62 0.67 2.88 0.80 0.11 4.36 3.81 1.09 0.30 0.12 0.17 0.27 0.15 0.16 0.32 1.06

Sample

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

LaN

CeN

PrN

YbN

LuN

Eu/Eu*

Ce/Ce*

SmN/LaN

YbN/LaN

YbNGdN

LaN/GdN

ΣREE

YbN/SmN

UMZ-1 01 UMZ-1 02 UMZ-1 03 UMZ-1 04 UMZ-1 05 UMZ-1 06 UMZ-1 07 UMZ-1 08 UMZ-1 09 UMZ-1 10 UMZ-1 11 UMZ-1 12 UMZ-1 13 UMZ-1 14 UMZ-1 15 UMZ-1 16 UMZ-1 17 UMZ-1 18 UMZ-1 19 UMZ-1 20 UMZ-1 21 UMZ-1 22 UMZ-1 23 UMZ-1 24 UMZ-1 25 Average

12.75 2.02 4.01 3.18 5.83 1.15 9.07 7.45 28.19 3.25 4.30 3.98 12.73 5.77 0.92 26.79 23.80 5.99 2.58 0.79 1.27 1.44 1.12 1.96 2.07 6.90

6.29 4.21 9.29 5.14 7.16 2.12 4.49 8.65 20.98 3.24 2.48 2.97 3.07 7.40 2.06 12.48 11.94 6.08 3.27 0.93 2.99 1.72 2.02 4.87 3.39 5.57

4.49 0.24 0.29 1.28 1.20 0.59 1.57 3.50 1.66 0.87 0.90 0.73 0.70 3.02 0.55 3.13 4.05 1.87 0.97 0.34 1.16 0.60 0.47 0.74 1.03 1.44

20.05 34.17 58.81 19.56 38.31 8.30 7.84 40.06 53.68 15.36 9.78 12.96 9.13 31.55 13.60 20.25 37.58 27.17 17.98 5.94 15.26 8.09 9.91 29.32 16.09 22.43

6.33 15.69 21.50 5.95 13.79 2.72 2.12 13.16 17.13 4.96 3.36 4.17 2.65 9.65 5.19 5.26 11.45 10.01 6.61 1.87 4.97 3.08 4.06 10.68 5.12 7.66

70.30 212.11 266.29 67.55 166.43 30.18 21.84 158.28 164.69 57.16 39.46 47.40 33.80 102.86 65.88 59.15 124.03 104.82 87.14 25.40 55.60 33.53 49.72 124.01 57.80 89.02

26.70 85.52 100.55 24.55 64.66 11.19 9.21 60.77 55.14 21.75 16.02 18.34 13.58 36.17 26.77 24.28 45.75 34.56 36.28 10.04 20.67 12.75 19.30 45.88 22.15 33.70

123.43 389.10 439.60 108.54 289.79 51.27 49.40 276.54 220.89 98.02 78.53 84.46 66.86 155.29 127.54 118.87 197.26 146.09 176.40 49.04 93.34 61.05 91.01 196.32 103.00 151.67

28.07 82.53 87.67 23.66 60.43 12.18 14.03 59.50 44.87 21.33 18.37 18.55 16.36 32.47 27.93 27.67 42.86 28.95 39.59 11.22 20.42 14.32 20.69 37.76 22.90 32.57

295.37 763.82 802.59 235.92 577.77 135.09 180.20 583.17 398.79 215.94 204.67 191.16 173.45 318.24 280.56 303.97 426.16 267.87 410.11 123.96 202.05 155.10 202.10 336.27 233.18 320.70

64.20 145.25 151.68 47.98 111.11 30.75 49.44 116.99 75.91 44.19 45.18 39.60 39.12 63.73 58.33 71.04 89.30 51.43 86.88 27.94 41.95 35.66 41.58 62.12 49.24 65.62

12.61 0.79 0.79 0.84 1.13 0.60 5.19 6.84 30.45 2.74 9.16 2.26 38.23 2.00 0.69 28.71 9.74 7.71 2.55 0.97 0.87 1.29 0.67 0.64 1.74 6.77

58.06 4.36 2.77 34.81 38.01 18.16 40.97 76.65 83.92 23.24 27.85 21.13 47.56 55.57 9.08 50.64 73.86 19.05 15.82 12.08 16.03 8.54 17.04 15.48 20.38 31.64

17.21 1.43 1.84 2.40 4.92 1.46 12.79 7.62 34.34 2.67 5.08 5.49 23.61 6.56 0.88 35.74 31.23 8.93 2.46 1.02 1.42 2.18 1.24 1.28 2.62 8.66

1413.3 3654.6 3840.1 1128.8 2764.5 646.4 862.2 2790.3 1908.1 1033.2 979.3 914.6 829.9 1522.7 1342.4 1454.4 2039.0 1281.7 1962.2 593.1 966.7 742.1 967.0 1608.9 1115.7 1534.5

1993.8 4510.9 4710.6 1490.1 3450.6 955.0 1535.4 3633.2 2357.5 1372.4 1403.1 1229.8 1214.9 1979.2 1811.5 2206.2 2773.3 1597.2 2698.1 867.7 1302.8 1107.5 1291.3 1929.2 1529.2 2038.0

1.22 0.06 0.04 0.39 0.22 0.43 0.81 0.58 0.15 0.38 0.56 0.36 0.40 0.60 0.32 0.60 0.59 0.45 0.39 0.44 0.53 0.49 0.32 0.19 0.43 0.44

3.94 4.08 2.30 24.52 16.13 19.45 5.03 10.62 2.60 8.59 4.08 6.00 1.58 15.34 11.70 1.58 4.23 2.29 6.32 12.18 14.42 5.09 18.66 17.09 9.55 9.10

2.56 27.19 60.00 31.42 32.52 18.21 4.43 6.49 3.53 6.06 1.39 6.75 0.41 18.97 15.38 2.23 6.29 4.04 6.58 4.93 17.60 6.84 15.37 38.90 9.98 13.92

112.0 4602.8 4836.5 1345.4 2448.6 1082.7 166.0 408.0 62.7 376.8 106.9 405.1 21.7 761.3 1954.0 50.7 209.3 166.2 770.1 612.7 1109.9 575.3 1434.7 2506.1 640.5 1070.6

77.41 131.93 227.07 75.52 147.92 32.05 30.27 154.67 207.26 59.31 37.76 50.04 35.25 121.82 52.51 78.19 145.10 104.90 69.42 22.93 58.92 31.24 38.26 113.21 62.12 86.60

0.16 0.01 0.00 0.01 0.01 0.02 0.17 0.04 0.15 0.05 0.24 0.05 1.08 0.02 0.01 0.37 0.07 0.07 0.04 0.04 0.01 0.04 0.02 0.01 0.03 0.11

710.9 1738.6 1945.0 572.0 1368.1 300.6 385.5 1393.1 1163.4 506.0 449.0 442.8 424.6 812.5 617.0 727.1 1080.7 703.7 881.7 267.7 473.1 334.9 456.1 862.8 533.3 766.0

43.81 169.27 80.61 42.82 75.29 59.45 37.44 62.90 17.73 62.19 77.00 60.05 52.71 40.12 127.07 22.73 33.30 41.11 117.02 124.37 63.05 84.13 93.35 64.42 64.18 68.65

Ce/Ce* = CeN / (LaN × (PrN)0.5. Eu/Eu* = EuN / (SmN × GdN)0.5. T Ti-in-zrc: Ti-in-zircon crystallization temperatures (Watson et al., 2006).



Table 2 LA-ICP-MS major and trace element contents and Ti-in-zircon temperatures of detrital zircon grains from Umm Zariq Formation metapelites (sample UMZ-1).

Sample

Si

P

Ca

Ti

Cr

Rb

Sr

Y

Zr

Nb

Ba

Hf

Ta

Pb

Th

U

i-in-Zr

Th/U

Zr/Hf

Nb/Ta

Y/Ho

U/Yb

La

Ce

Pr

MH2 01 MH2 02 MH2 03 MH2 04 MH2 05 MH2 06 MH2 07 MH2 08 MH2 09 MH2 10 MH2 11 MH2 12 MH2 13 MH2 14 MH2 15 MH2 16 MH2 17 MH2 18 MH2 19 MH2 20 MH2 21 MH2 22 MH2 23 MH2 24 MH2 25 MH2 26 Average

153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225 153,225

565 869 629 4181 971 754 641 760 535 545 529 1278 493 501 656 646 744 685 546 574 669 666 509 547 413 418 782

15,602 13,509 18,518 15,768 17,385 20,387 19,635 18,514 17,607 21,895 25,108 22,970 20,986 18,272 30,926 28,698 32,651 41,871 24,937 31,749 30,326 33,099 29,078 24,797 23,520 26,605 24,016

65 60 67 55 68 72 80 71 60 81 81 74 74 65 98 105 101 103 89 119 110 112 106 86 79 83 83

142 111 138 127 147 178 176 154 127 150 148 130 143 142 171 201 153 173 145 173 153 162 147 154 113 113 149

1.6 1.2 1.6 1.3 1.4 1.5 1.5 1.4 1.2 1.7 1.4 1.7 1.6 1.3 2.1 2.3 2.0 2.3 1.5 2.5 2.3 2.4 2.0 1.8 1.8 1.5 1.72

2.0 2.1 0.6 27.1 3.7 0.7 0.6 1.4 0.5 0.7 0.7 8.4 0.5 1.2 1.7 3.1 0.8 1.7 0.7 2.3 1.3 0.8 0.7 5.0 0.5 2.3 2.73

2594 1632 1610 1720 1385 1816 2355 1760 2935 1354 1787 1402 724 1539 1775 2374 1117 2024 1319 1674 1797 623 2746 3734 1186 2853 1840

411,298 430,152 425,454 392,277 428,413 426,431 404,240 432,006 405,726 433,918 400,374 395,811 413,554 424,618 437,774 445,411 441,226 430,084 429,435 428,000 423,942 439,836 417,685 397,912 403,589 423,033 420,854

22.9 8.2 14.6 17.1 9.9 0.9 3.3 12.5 2.2 6.6 6.6 5.7 4.7 11.9 15.1 15.9 3.4 1.7 0.6 7.7 11.8 0.8 6.3 40.1 7.0 10.6 9.53

6.0 0.7 0.8 3.1 0.8 0.7 0.7 0.8 0.7 0.9 0.8 12.3 0.6 0.7 1.2 17.2 1.0 87.8 0.9 1.2 1.1 1.5 0.9 2.4 0.6 0.8 5.63

6005 6368 6499 5582 6982 8882 5986 7134 5750 7299 6249 5819 7175 6845 6962 8482 7431 9139 10,674 7315 6935 11,546 6647 6878 6742 7474 7261

0.97 1.8 2.6 1.5 2.3 0.4 0.9 2.8 0.7 1.3 1.5 1.2 1.2 2.5 3.2 3.8 1.0 0.6 0.3 1.8 2.4 0.3 1.5 3.1 1.5 2.5 1.68

6.1 6.5 8.0 6.0 6.5 28.8 7.2 9.3 6.7 6.1 6.2 5.7 3.6 7.3 9.6 12.0 4.9 15.3 47.4 6.7 8.2 22.0 8.6 12.9 5.7 12.5 11

30 25 26 19 21 101 36 40 37 24 22 23 10 23 31 42 19 61 183 23 29 77 36 65 24 60 42

51 55 71 45 56 191 60 78 55 52 53 44 31 63 85 117 43 115 314 58 71 146 74 103 50 103 84

938 928 941 916 943 951 964 948 927 966 965 954 953 937 991 1001 996 998 978 1018 1008 1009 1002 973 962 968 967

0.59 0.44 0.36 0.42 0.38 0.53 0.60 0.51 0.67 0.46 0.42 0.52 0.33 0.37 0.36 0.36 0.44 0.52 0.58 0.39 0.41 0.53 0.49 0.63 0.49 0.58 0.48

68.49 67.55 65.47 70.28 61.36 48.01 67.54 60.55 70.57 59.45 64.07 68.02 57.64 62.03 62.88 52.51 59.38 47.06 40.23 58.51 61.14 38.09 62.84 57.85 59.86 56.60 59.54

23.59 4.57 5.65 11.23 4.38 2.34 3.79 4.51 3.18 5.10 4.28 4.68 3.84 4.72 4.77 4.23 3.28 2.75 1.97 4.25 4.82 2.54 4.20 12.94 4.62 4.23 5.40

27.49 26.65 27.61 26.70 27.08 29.88 27.98 27.67 28.05 28.09 27.39 27.40 27.40 25.97 25.14 22.64 26.23 26.11 28.92 26.00 25.88 30.61 25.62 25.69 27.07 26.42 26.99

0.08 0.14 0.17 0.12 0.17 0.38 0.10 0.18 0.07 0.13 0.12 0.12 0.14 0.15 0.17 0.16 0.12 0.19 0.82 0.13 0.15 0.65 0.11 0.12 0.15 0.14 0.19

0.36 177.97 0.09 214.79 143.07 0.09 0.98 77.09 0.09 3.38 2.88 122.11 0.08 31.66 11.53 3.58 0.17 3.51 0.10 23.04 93.15 0.09 0.21 8.00 0.07 79.25 38.36

8.62 489.97 25.93 605.72 389.33 9.48 9.23 229.15 5.17 18.91 23.78 338.71 11.55 103.96 60.40 41.51 6.31 22.40 14.94 55.55 269.42 17.01 18.60 46.61 13.96 229.13 117.90

1.57 69.04 0.28 90.84 52.98 0.16 0.72 28.62 0.46 1.33 1.16 48.51 0.06 11.27 4.57 2.53 0.12 2.55 0.23 8.17 35.89 0.07 0.55 4.64 0.17 30.19 15.26

Sample

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

LaN

CeN

PrN

YbN

LuN

Eu/Eu*

Ce/Ce*

SmN/LaN

YbN/LaN

YbNGdN

LaN/GdN

Σ'REE

YbN/SmN

MH2 01 MH2 02 MH2 03 MH2 04 MH2 05 MH2 06 MH2 07 MH2 08 MH2 09 MH2 10 MH2 11 MH2 12 MH2 13 MH2 14 MH2 15 MH2 16 MH2 17 MH2 18 MH2 19 MH2 20 MH2 21 MH2 22 MH2 23 MH2 24 MH2 25 MH2 26 Average

17.16 354.39 3.49 501.53 267.11 3.20 7.35 143.87 7.29 8.85 9.42 252.26 0.71 59.71 25.39 19.04 2.05 16.29 2.20 39.36 187.51 0.53 8.66 38.44 2.47 158.38 82.18

24.01 83.80 6.47 135.22 63.61 7.03 13.10 38.10 16.31 6.39 10.00 58.74 2.49 18.34 12.58 18.91 4.65 14.62 5.60 14.37 47.73 2.07 14.91 47.24 5.19 45.58 27.58

6.50 9.11 1.90 19.95 6.86 1.09 3.35 4.63 5.00 1.22 2.25 7.13 0.62 2.85 2.23 6.09 1.10 2.51 0.81 2.67 5.73 0.43 3.80 15.63 1.22 5.75 4.63

86.76 103.99 36.87 162.33 77.25 39.05 64.28 63.48 84.35 30.13 45.07 75.92 14.38 45.09 46.50 83.99 27.78 57.43 31.05 47.95 76.09 10.19 78.19 170.11 27.26 97.60 64.73

27.40 21.44 13.83 30.34 16.80 13.32 21.22 17.47 27.23 10.60 15.40 16.53 5.45 14.39 16.13 28.84 9.30 19.26 10.78 15.18 19.53 4.05 26.24 54.09 9.97 27.74 18.94

280.31 190.77 160.30 229.32 153.26 161.15 238.94 181.62 299.04 124.22 174.69 152.92 67.29 159.42 189.38 316.75 112.54 212.57 120.93 175.60 200.49 48.75 293.49 492.80 115.97 303.99 198.33

94.39 61.24 58.30 64.43 51.16 60.79 84.17 63.60 104.66 48.19 65.25 51.17 26.44 59.25 70.58 104.86 42.58 77.53 45.61 64.37 69.41 20.34 107.20 145.34 43.81 107.96 68.95

375.52 243.06 246.55 235.30 201.38 268.73 353.91 263.31 424.15 209.70 268.96 208.10 116.50 249.94 298.23 426.92 189.09 332.93 200.15 270.33 282.97 97.59 431.38 530.27 189.01 444.86 283.03

71.58 44.98 47.20 42.77 38.13 53.60 68.02 50.26 81.80 43.32 51.68 39.92 23.62 48.30 57.47 82.85 38.40 66.40 41.70 51.58 53.37 22.16 81.61 97.40 37.29 84.22 54.60

648.53 393.75 418.97 369.74 332.88 498.03 611.18 444.41 742.48 413.03 452.82 360.83 216.58 420.92 508.83 750.21 363.07 605.87 384.60 454.66 466.48 224.52 703.66 838.51 340.19 729.61 488.24

121.27 70.40 74.68 66.33 58.83 94.94 114.07 77.83 138.81 80.47 82.79 67.02 41.19 74.34 91.65 141.87 68.79 113.94 73.30 83.22 82.51 46.40 128.19 147.07 63.10 131.09 89.77

1.17 574.10 0.29 692.87 461.52 0.28 3.17 248.68 0.28 10.90 9.29 393.90 0.24 102.13 37.19 11.55 0.56 11.32 0.33 74.32 300.48 0.30 0.68 25.81 0.22 255.65 123.74

10.67 606.40 32.09 749.65 481.84 11.73 11.42 283.60 6.40 23.40 29.43 419.20 14.30 128.66 74.75 51.37 7.81 27.72 18.49 68.75 333.44 21.05 23.02 57.69 17.28 283.58 145.91

13 566 2 745 434 1 6 235 4 11 10 398 1 92 37 21 1 21 2 67 294 1 4 38 1 247 125

3103 1884 2005 1769 1593 2383 2924 2126 3553 1976 2167 1726 1036 2014 2435 3590 1737 2899 1840 2175 2232 1074 3367 4012 1628 3491 2336

3766 2186 2319 2060 1827 2948 3543 2417 4311 2499 2571 2081 1279 2309 2846 4406 2136 3539 2276 2584 2562 1441 3981 4567 1960 4071 2788

0.44 0.30 0.38 0.41 0.30 0.20 0.35 0.29 0.41 0.27 0.32 0.33 0.32 0.30 0.28 0.47 0.30 0.27 0.19 0.31 0.29 0.28 0.34 0.53 0.31 0.26 0.32

2.76 1.06 39.42 1.04 1.08 19.45 2.64 1.17 6.28 2.15 3.13 1.06 40.10 1.32 2.00 3.32 10.51 1.80 23.27 0.97 1.12 49.95 13.20 1.84 31.00 1.13 10.11

105.69 0.75 112.85 1.00 0.71 128.30 21.17 0.79 301.95 3.01 5.52 0.76 52.76 0.92 1.73 8.40 42.51 6.62 86.50 0.99 0.81 35.74 112.94 9.39 119.35 0.91 44.70

2663.53 3.28 6818.51 2.55 3.45 8480.14 921.33 8.55 12,825.04 181.25 233.22 4.38 4282.10 19.72 65.46 310.83 3096.57 256.02 5542.74 29.27 7.43 3617.03 4973.11 155.47 7299.12 13.66 2377.45

334.98 401.51 142.36 626.76 298.26 150.77 248.19 245.10 325.68 116.33 174.02 293.13 55.52 174.09 179.54 324.29 107.26 221.74 119.88 185.14 293.78 39.34 301.89 656.80 105.25 376.83 249.94

0.00 1.43 0.00 1.11 1.55 0.00 0.01 1.01 0.00 0.09 0.05 1.34 0.00 0.59 0.21 0.04 0.01 0.05 0.00 0.40 1.02 0.01 0.00 0.04 0.00 0.68 0.37

1763.98 2313.91 1094.86 2768.61 1852.65 1210.66 1590.52 1683.44 1936.83 999.74 1206.15 1799.87 526.96 1299.44 1395.47 2027.95 865.95 1547.81 932.00 1306.05 1890.28 494.20 1896.69 2636.15 849.68 2475.35 1552.51

25.21 4.38 60.42 2.55 4.88 66.10 43.53 10.88 42.47 60.31 42.25 5.73 81.15 21.41 37.74 37.02 72.85 38.67 64.08 29.52 9.12 101.20 44.03 16.56 61.16 14.93 38.39

9

Ce/Ce* = CeN / (LaN × (PrN)0.5. Eu/Eu* = EuN / (SmN × GdN)0.5. T Ti-in-zrc: Ti-in-zircon crystallization temperatures (Watson et al., 2006).



Table 3 LA-ICP-MS major and trace element contents and Ti-in-zircon temperatures of detrital zircon grains from Malhaq Formation metapelites (sample MH-2).

10


Fig. 4. Chondrite-normalized multi-element spider diagrams of zircons from (a) sample UMZ-1 and (b) sample MH-2. C1 chondrite values are after McDonough and Sun (1995).

zircon temperatures (1084–1154 °C; Table 2). However, comparable or even higher Ti-in-zircon temperatures (N1000 °C) have been documented for magmatic, metamorphic and detrital zircons (e.g., Baldwin et al., 2007; Fu et al., 2008; Hiess et al., 2008; Moecher et al., 2014). Nb and Ta contents of zircons from sample MH-2 (average = 9.53 and 1.68 ppm, respectively) are obviously higher (about four to five times) than their abundances in zircons of UMZ-1 (average = 1.74 and 0.53 ppm, respectively). The abundances of Nb and Ta in all sample UMZ-1 zircons and Nb in MH-2 zircons fluctuate within the typical

range of unaltered magmatic zircon (Nb ≤ 62 ppm, Ta b 3 ppm; Hoskin and Schaltegger, 2003). Alternatively, three zircons from sample MH-2 possess slightly high Ta concentrations (3.1–3.8 ppm; Table 3). The enrichment of Ta over Nb in all the analyzed zircons relative to chondritic values is manifest in Fig. 4. Phosphorous abundances in zircons of sample UMZ-1 are roughly more than twice times higher relative to those from sample MH-2 (averages: 1806 and 782 ppm, respectively). The experimental study of Van Lichtervelde et al. (2011) has demonstrated that P can be



11

Fig. 6. Plots of the studied zircons on the: (a) Th + U vs. Pb and (b) Ca vs. P variation diagrams.

between Ca and P in the P-Ca-enriched zircons of Um Zariq metasediments (Fig. 6b) indicate that not all of the measured phosphorous is held in the zircon lattice, but may be ascribed to apatite microinclusions. With regard to the large ion lithophile elements (LILE) Ba, Sr and Rb in the zircons of both metasedimentary samples, they occur in very low concentrations that rarely exceed few ppms (cf. El-Bialy and Ali, 2013; Hoskin and Schaltegger, 2003; Xia et al., 2010). Ba and Sr regularly have subchondritic values, while Rb surpasses with contents up to 10× chondrite (Fig. 4). Fig. 5. Plots of chondrite-normalized (McDonough and Sun, 1995) REE contents in zircons from (a) sample BY-1 and (b) sample KI-1. Plot (c) shows fields of both samples.

5. Discussion 5.1. Ti-in-zircon thermometry

incorporated by coupled substitution in zircon through two different substitution mechanisms (plus Al+3 or Mn+2) to sustain charge balance. Calcium is the subsequent element in abundance after Zr and Si among the analyzed elements, reaching percent level (N10,000 ppm) in all zircon samples (Tables 2 and 3). Similar to phosphorous, calcium content in zircons of sample UMZ-1 is twice to thrice as high as those from sample MH-2 (Tables 2 and 3). These substantial calcium concentrations in zircon along with the reasonably fair positive correlation

The extent of titanium substitution into crystallizing zircon is mainly controlled by temperature and the activity of TiO2 (aTiO2) (Watson and Harrison, 2005). As a result, Ti concentration in zircon is an efficient geochemical tracer to infer zircon crystallization temperature (Ferry and Watson, 2007; Watson and Harrison, 2005; Watson et al., 2006). The capture of Ti4+ in zircon is deceptively facilitated by its substitution without charge compensation most favorably into the Si4 + site


12


(Ferry and Watson, 2007; Harrison et al., 2007). This thermometer (Ti-in-zircon thermometer; Watson et al., 2006) presumes the cocrystallization of rutile (i.e., pure TiO2) with zircon, which is a reasonable assumption for many common igneous rocks. Based on this postulation, equilibrium constant can be calculated from aTiO2 in zircon by setting the rutile activity equal to almost 1 (Watson and Harrison, 2005; Watson et al., 2006). Watson et al. (2006) have used experimentally synthesized zircons for high temperatures (1025–1450 °C) and natural zircons with estimated crystallization temperatures of ~ 580 °C–1170 °C in their experimental calibration of this thermometer. Recently, Ti-in-zircon thermometry has become popular and has been progressively applied in studying natural zircons (e.g., Barth and Wooden, 2010; Cates and Mojzsis, 2009; El-Bialy and Ali, 2013; Ickert et al., 2011; Liu et al., 2015; Moecher et al., 2014; Orejana et al., 2011, 2012; Wang et al., 2013; Wielicki et al., 2012; Xing et al., 2014). Albeit determination of very precise zircon crystallization temperatures (TTi-in-zrc) requires knowledge of aTiO2 and aSiO2 activities (Ferry and Watson, 2007), our calculation is based on the calibration of Watson and Harrison (2005) and Watson et al. (2006), which presumes that zircons crystallized in the presence of quartz and rutile at P ≈ 10 kbar. The largest uncertainties presented to zircon crystallization temperatures (TTi-in-zrc) by unconfined aTiO2 and aSiO2 were quantitatively evaluated to be ≈60–70 °C at 750 °C (Ferry and Watson, 2007). The absolutely very high Ti contents of all the analyzed zircons from both metasedimentary samples have yielded considerably high Ti-inzircon temperatures (Tables 2 and 3; Fig. 7a). Such obviously high

crystallization temperatures rule against formation through metamorphism (cf. Baldwin et al., 2007; Jiao et al., 2013; Sheng et al., 2012; Wu et al., 2009) and emphasize the magmatic origin of the studied zircons. The application of the Ti-in-zircon thermometer (Watson et al., 2006) to the zircons separated from Malhaq Formation metapelites (sample MH-2) returned temperatures within the 916–1018 °C range (average TTi-in-zrc = 967 °C; Table 3). On the other hand, the estimated Ti-in-zircon temperatures of Um Zariq schist (sample UMZ-1) reveal another range of higher zircon crystallization temperatures (1084–1154 °C) that cluster around 1120 °C (Fig. 7a). Since the concentration of titanium in zircon is generally accepted to be mainly controlled by temperature (Ferry and Watson, 2007; Fu et al., 2008; Watson and Harrison, 2005; Watson et al., 2006), the very high Ti-in-zircon temperatures of Um Zariq samples may indicate their derivation from more primitive igneous source (mafic) relative to the low temperature zircons of Malhaq Formation metasediments. Nevertheless, some detailed studies revealed that “even though Ti content in zircons is temperature dependent” zircons derived from most igneous rock types, including mafic rocks, show a similar Ti-temperature range (Coogan and Hinton, 2006; Fu et al., 2008), thus preventing a clear-cut distinction between granitic and more mafic sources on this basis. Also, Nutman (2006) stressed that zircon saturation in hightemperature magmas (i.e., less evolved) is typically only reached upon cooling, and high Ti abundances may then be the result of fractional crystallization rather than the magma composition and/or temperature. Therefore, intrusive rocks underwent slow cooling and extensive fractional crystallizations are expected to contain more Ti-rich zircons relative to their less fractioned compositionally equivalent extrusive counterparts (see Section 5.2). 5.2. Zircon source rocks

Fig. 7. (a) Histograms of crystallization temperatures, based on Ti-in-zircon thermometry, of zircons in samples MH-2 and UMZ-1. (b) Plots of the estimated crystallization temperatures of the studied zircon samples using the thermometer calibration of Watson et al. (2006) against their Hf contents. The boundary value of maximum Ti abundance in unaltered igneous zircon is after Hoskin and Schaltegger (2003).

The trace element characteristics present an effective tool for the recognition the source rock of detrital zircons. The abundances of several diagnostic trace elements, and the shape and slope of chondritenormalized REE patterns are idiosyncratic for zircons of different sources (Belousova et al., 2002; Hermann et al., 2001; Hinton and Upton, 1991; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003; Murali et al, 1983; Whitehouse and Kamber, 2002; Whitehouse and Platt, 2003). Zircons of different origins exhibit a wide range of REE concentrations varying from tens to several tens of thousands ppm (Hinton and Upton, 199; Belousova et al., 2002; Hoskin and Schaltegger, 2003). As mentioned in Section 4.3, the REE contents of zircons from sample UMZ-1 are approximately twice as much as those of sample MH-2 zircons (average = 1553 ppm vs. 766 ppm, respectively). Therefore, this undoubtedly implies different sources for the detrital zircons of Malhaq metasediments (cf. Tables 3–6 in Belousova et al., 2002), with the zircons of Um Zariq Formation being derived from more evolved igneous rock types due to their higher REE abundances. Further, even the MREE and HREE in both sample zircons show comparable and overlapping patterns, the zircons of Malhaq Formation show extremely contrasting LREE patterns from almost flat to steeply positive with Ce spikes (Fig. 5b, c), which also confirm the difference in sources of the two formation zircons. The Zr/Hf ratios of all of the analyzed zircons from both metasediments fall within the 20–70 range typifying common igneous rocks (Bea et al., 2006; Ewing et al., 2014; Hoskin and Schaltegger, 2003). Lower Zr/Hf ratios (b 20) can indicate granitic sources with high silica contents for instance pegmatites and highly fractionated peraluminous granites (e.g., Barros et al., 2010; Bea et al., 2006; Pupin, 2000; Wang et al., 2010). Zircons with extremely low Hf and consequently very high Zr/Hf ratios are usually derived from silicaundersaturated rocks like feldspatoid-bearing syenites (i.e., 73–135) and kimberlites (up to 253) (e.g., Barros et al., 2010; Belousova et al., 2002; Heaman et al., 1990).



Using zircon contents of some trace elements, Belousova et al. (2002) proposed classification and regression trees (CART) for recognizing zircons derived from different igneous rock types. Application of the “short” CART tree of Belousova et al. (2002) using the trace elements Lu, Hf, U, Yb, and Y has revealed that Malhaq Formation zircons are mainly derived from mafic source rocks (basalt and dolerite; 24 zircons out of 26). However, Um Zariq zircons have varied and more evolved parent rocks including high SiO2 (70–75 wt.%) granitoids (Fig. 8). The modest Niobium and Yttrium concentrations in all of the analyzed zircons (b 100 ppm and 10,000 ppm, respectively; Tables 2 and 3) rule against involvement of A-type granites among the granitoid source rocks of both formations (Nardi et al., 2013).

5.3. Metamorphic/hydrothermal vs. magmatic origin Though zircon normally forms by crystallization from magmas, its precipitation from metamorphic and hydrothermal fluids has quite considerable occurrences (e.g., Cavosie et al., 2006; Dempster et al., 2008; Kusiak et al., 2009; Lisowiec et al., 2013; Sheng et al., 2012; Soman et al., 2010). Hydrothermal and metamorphic zircons are quite similar and both comprise lately crystallized grains or overgrowth mantling inherited cores, that precipitate from hydrous melts or aqueous fluids without inherited signature from the protolith (Bouvier et al., 2012; Chen et al., 2010; Fu et al., 2009; Geisler et al., 2003, 2007; Hoskin, 2005; Pelleter et al., 2007; Rubatto, 2002; Rubatto et al., 2008; Sheng

Fig. 8. (a) ‘Short’ CART tree for the recognition of zircons from different rock types (Belousova et al., 2002). (b) Relative abundance of source rock types derived from composition of the detrital zircons of Um Zariq (sample UMZ-1) and Malhaq (sample MH-2) formation metapelites. Numbers between brackets in (b) indicate number of zircons belonging to each rock type.

13

et al., 2012; Whitehouse and Kamber, 2002; Xia et al., 2009; Zheng et al., 2006). Th/U ratios have become a commonly employed criterion for distinguishing zircon formation from magmatic versus metamorphic systems. Specifically, a Th/U value below 0.1 has been widely cited as a discriminant of metamorphic zircon (e.g., Ahrens et al., 1967; Hoskin and Black, 2000; Schaltegger et al., 1999; Rubatto et al., 2001; Rubatto, 2002; Williams et al., 1996). Most zircons of igneous origin have typically Th/U values within the 0.4–1.0 range (Hoskin and Black, 2000; Hoskin and Schaltegger, 2003). Our result (Tables 2 and 3) reveals that none of the studied zircons have the characteristic Th/U signature of metamorphic origin and accordingly they are most plausibly magmatic. However, inexplicable occurrences of low Th/U magmatic and high Th/U metamorphic zircons have been occasionally reported (e.g., Carson et al., 2002; Harley et al., 2007; Hidaka et al., 2002; Hokada and Harley, 2004; Kelly and Harley, 2005; Möller et al., 2003). Hydrothermal zircons exhibit a distinctive REE distribution pattern compared with that of magmatic zircons (e.g., Cavosie et al., 2006; Geisler et al., 2003; Hoskin, 2005; Pelleter et al., 2007; Pettke et al., 2005; Rimsa et al., 2007; Xia et al., 2010). For instance, hydrothermal activities may significantly result in LREE enrichment in igneous zircons. Nevertheless, some works have found that some absolutely hydrothermal zircons are impossible to differentiate from magmatic zircons with respect to their chemical composition and isotopic fingerprint (e.g., Fu et al., 2009; Schaltegger, 2007). In spite of their different formation mechanisms, the hydrothermally altered and hydrothermally precipitated zircons bear common features since they are more LREEenriched and display flatter chondrite-normalized LREE patterns (Low (Sm/La)N) than unaltered magamtic zircon as well as having smaller Ce anomalies (Ce/Ce*) (Hoskin, 2005). As has been comprehensively revealed in Section 4.3, many of the analyzed zircons from both samples, and those from the sample MH-2 in specific, show marked enrichment in Light REEs that surpasses the normal content in igneous zircon. This excessive overabundance in LREE (average (Sm/La)N = 44.7 and 13.92 for MH-2 and UMZ-1, respectively), has led to evidently flat LREE segments for many of the investigated zircons in opposition to the negative steep slope of igneous zircon LREE patterns ((Sm/La)N = 57–547) (Hoskin and Schaltegger, 2003). The probably hydrothermally originated zircons, with (LREE)N ratio N 10, represent the vast majority of grains extracted from both UMZ-1 and MH-2 rock samples (Tables 2 and 3; Fig. 9a). However, the normalized HREE patterns for all of the analyzed zircons from each of the two rock samples exhibit rather similar positively steep slopes (cf. El-Bialy and Ali, 2013; Hoskin, 2005) (Fig. 5). The relationship of LaN vs. PrN provides an efficient discrimination between unaltered magmatic and the LREE-enriched zircons of probable hydrothermal origin (El-Bialy and Ali, 2013). This diagram is based on the combination of LaN N 1 and PrN N 10 as useful discriminants for recognizing possibly hydrothermal zircons with LREE-enriched patterns (cf. Cavosie et al., 2006; Hoskin and Schaltegger, 2003). The relationships of LaN vs. PrN of the studied detrital zircons reveal that appreciable number of zircons (9 from each sample; 18 zircons) fall in the unaltered magmatic zircons field, whereas many of them (22 out of 51) plot in the hydrothermal plus LREE-enriched late-magmatic zircons field, with the remaining zircons (11) falling in an in-between area due to their high LaN ratio (N1) (Fig. 9b). The abovementioned LREE features suggest that hydrous melts and/or aqueous fluids have played a significant role in the formation of these zircons, more especially those hosted in the Malhaq Formation sample (MH-2). An additional evidence to shed light on the difference between igneous and hydrothermal zircons, also according to LREE characteristics, can be made using the (Sm/La)N vs. La discrimination diagram (Hoskin, 2005). In this diagram (Fig. 9c), the Um Zariq and Malhaq formation zircons sub-equally plot in or near the fields defining magmatic and hydrothermal zircon. Most of the data points which plot outside the two fields are still close enough to the limited extension of the fields


14


(dashed line), indicating their possible magmatic and hydrothermal origins (Hoskin, 2005). As revealed earlier, many zircons from Malhaq Formation show more tendencies towards hydrothermal origin

(Fig. 9c). However, the exclusive reliance on LREE abundance and pattern for the distinction of magmatic from hydrothermally altered igneous zircons seems insufficient to some extent. In view of that hydrothermal alteration of zircon is usually associated with coupled addition of Ca and Ba (Rayner et al., 2005), the restrict range of low Ba contents for UMZ-1 and MH-2 zircons (median = 1.82 and 0.9 ppm, respectively) and the constant horizontal variability between Ca and Ba (not shown here), the possibility of subjection to hydrothermal alteration can be definitely excluded. Ti-in-zircon thermometry is a robust indicator for discriminating hydrothermal zircon (TTi-in-zrc b 500 °C) from igneous zircon (TTi-in-zrc N 600 °C) (Fu et al., 2009). The LA-ICP-MS analyses of the studied detrital zircons yielded considerably high Ti content, corresponding to high Ti-in-zircon crystallization temperatures (Tables 2 and 3; Fig. 7). These temperatures are substantially higher than those of hydrothermal zircons, and instead match the crystallization temperatures of magmatic zircons (cf. El-Bialy and Ali, 2013; Fu et al., 2008, 2009; Hoskin, 2005; Ickert et al., 2011; Pettke et al., 2005). The possibility of involvement of hydrothermal fluids in the genesis of some of the analyzed detrital zircons mainly arises from their LREE characteristics including overabundance, flat patterns and trivial positive Ce anomalies. For detrital zircons derived from granitoid source (Fig. 8b), these LREE aspects could be explained by the late crystallization from a high LREE/HREE melt due to the earlier separation of feldspar, hornblend, biotite and early crystallized zircon. The latecrystallized zircons may have been equilibrated with liquids that are assumed to be enriched in LREE, owing to their more incompatibility relative to HREE. The high mineral/melt partition coefficients of HREE in hornblende (Sisson, 1994) and biotite (Nash and Crecraft, 1985) instigate their depletion in the evolved melts, giving rise to flat LREE segments and widening the range of LREE in zircons (cf. Long et al., 2012; Nardi et al., 2012). On the other hand, possible sources of the observed LREE overabundance in the zircons derived mafic source rocks include accidental analysis of LREE-bearing inclusions, Complex REE substitution mechanisms and radiation-induced lattice damage (El-Bialy and Ali, 2013; Whitehouse and Kamber, 2002). 5.4. Mantle versus crustal affinity Zircons crystallized from mantle materials show distinctive REE features that readily distinguish them from crust-derived magmatic zircons (Belousova et al., 2002; Heaman et al., 1990; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003). In rocks of mantle affinity, zircon contains a symptomatically low total REE content that seldom reaches 100 ppm (Belousova et al., 1998; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003). Alternatively, zircon from crustal rocks show wide range of variability in total REE contents that diverge from as low as 250 ppm to as high as 5000 ppm, and typically contain an average between 1500 and 2000 ppm (Heaman et al., 1990; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003). Likewise, yttrium abundances are very low in mantle zircon (b 100 ppm), whereas crustal zircons fluctuate and range from several tens of ppm up to ca. 5000 ppm (Belousova et al., 1998; Hoskin and Ireland, 2000). Without exception, the Y and ∑ REE contents of the detrital zircon grains presented in our study indicate that they are exclusively crust-derived magmatic zircons (Tables 2 and 3). Moreover, mantle-derived zircons have REE patterns characteristically different from crustal affinity zircons as the latter show moderately enriched flat HREE patterns (~100× chondrite) with no significant Eu anomaly (Eu/Eu* ~ 1.0) (Belousova et al., 1998,

Fig. 9. (a) The overabundance of the LREE in the studied zircons illustrated by plotting Hf against LREEN (LREEN = LaN + CeN + PrN). Igneous zircon field data in (a) are from Hoskin and Schaltegger (2003). (b) Distinction between unaltered magmatic zircon and LREEenriched hydrothermal and late-magmatic zircon using a plot of (La)N versus (Pr)N. (c) (Sm/La)N vs. La discrimination diagram (Hoskin, 2005) for zircons from samples UMZ-1 and MH-2.)



2002; Hinton and Upton, 1991; Hoskin and Ireland, 2000; Hoskin and Schaltegger, 2003). All the investigated zircon grains have steeply positive HREE (Tb–Lu) patterns that frequently associated with distinct negative Eu anomalies (Fig. 5), confirming that they are derived from crustal source. The zircons from granitoid source rocks are unquestionably crystallized in continental crust, while those from mafic sources (basalts and dolerites), such as the majority of Malhaq Formation zircons, might have derived from an oceanic crust (e.g., MORB) or more evolved continental crust (e.g., partial melting of more primitive mafic rocks). Grimes et al. (2007) demonstrated that zircons from different source regions have divergent U/Yb ratios, being rather low in ocean gabbros (0.18) and increasing to 1.07 in continental granitoids, and 2.1 in kimberlites. Grimes et al. (op. cit.) introduced a couple of discrimination diagrams employing the relationships of U/Yb ratio versus Hf and Y contents to distinguish between zircons derived from ocean crust, continental crust, and mantle (kimberlite zircon megacrysts). Apart from few zircons from Malhaq Formation that plotted in the field of “ocean crust zircon”, the rest of this formation and all of the Um Zariq Formation detrital zircons are classified as continental zircon (Fig. 10). Further, the older Cryogenian zircons of Um Zariq Formation possess stronger continental crust affinity compared to those from Ediacaran Malhaq Formation that fall within the lower area of the continental crust field. Overall, these findings are consistent with the domination of juvenile continental crust for the magma sources of the late Cryogenian– Ediacaran igneous rocks of the ANS (Fritz et al., 2013; Hargrove et al., 2006; Johnson et al., 2011; Stern, 1994, 2002; Stern and Johnson, 2010). 5.5. Metamorphic PT conditions The metamorphic pressure and temperature conditions of the investigated samples were deduced through conventional geothermometers and geobarometers using chemical composition of minerals in equilibrium at peak metamorphism (Table 4). Sample UMZ-1 peak assemblage which includes garnet(rim) + biotite(S2) + muscovite(S2) + staurolite + analusite, is used for estimation of the peak P-T conditions. Peak temperatures of 555 ± 18 °C and 581 ± 17 °C are estimated using garnet–biotite thermometer of Bhattacharya et al. (1992) and Perchuk and Lavrent’eva (1983) calibrations, respectively. The difference in temperatures between the two previous calibrations is also reported through garnet–muscovite geothermometer of Hoisch (1989) and Wu and Zhao (2006), where

15

Table 4 Conventional geothermobarometery of the metapelites of Um Zariq (sample UMZ-1) and Malhaq (sample MH-2) formations. Sample Temperature estimates (°C) (1) grt-bt geothermometers Bhattacharya et al. (1992) Perchuk and Lavrent’eva (1983) Dasgupta et al. (1991) (2) grt-ms geothermometers Hoisch (1989) Wu and Zhao (2006) (3) grt-turm geothermeter Colopietro and Frieberg (1987) Pressure estimates (kbar) (1) GPBMQ (Hoisch, 1990) (2) GBMAQ (Wu and Zhao, 2006)

UMZ-1

555 ± 18 581 ± 17

559 ± 6 586 ± 16

MH-2

563 ± 6 558 ± 8 543 ± 6

566 ± 13

4.21 ± 0.38 4.50 ± 0.43

3.88 ± 0.19

they yield peak temperatures of 559 ± 6 °C and 586 ± 16 °C, respectively. Peak pressures of sample UMZ-1 were assessed with the GPBMQ (garnet–plagioclase–biotitet–muscovite + qtz) calibration of Hoisch (1990), and GBMAQ (garnet–biotite–muscovite–aluminosilicate– quartz) geobarometer of Wu and Zhao (2006). Pressure inferred using the GPBQ geobarometer of Hoisch (1990) is 4.21 ± 0.38 kbar, while slightly higher pressure of 4.50 ± 0.43 kbar is determined using GBMAQ geobarometer of Wu and Zhao (2006). Sample MH-2 is characterized by peak assemblage of garnet(rim) + biotite(S2) + muscovite(S2). Garnet–biotite thermometers of Perchuk and Lavrent'eva (1983) and Dasgupta et al. (1991) yield close temperatures of 563 ± 6 °C and 558 ± 8 °C, respectively. Consistent peak temperature of 566 ± 13 °C is further confirmed using the bt-turm calibration of Colopietro and Frieberg (1987). However, slightly lower peak temperature of 543 ± 6 °C was calculated using garnet–muscovite calibration of Hoisch (1989). Application of the barometer of Hoisch (1990) gave peak assemblage pressure of 3.88 ± 0.19 kbar. In summary, the estimated metamorphic P-T conditions of the investigated samples are in agreements with previous peak P-T estimates reported by Abu El-Enen et al. (2003) for the Um Zariq Formation (570 °C/4 kbar), Abu El-Enen (2008) for the central and southern parts of Malhaq Formation (533–556 °C/3.8 kbar), and lies in the metamorphic P-T range for the peak metamorphic conditions estimated by Brooijmans et al. (2003) for the KMC.

Fig. 10. Plots of the detrital zircons from Um Zariq and Malhaq formations (samples UMZ-1 and MH-2) on the U/Yb vs. Y (a) and U/Yb vs. Hf (b) diagrams with the fields of Grimes et al. (2007) to discriminate between continental and oceanic crust zircon. Heavy lines indicate the lower limit of zircons from continental crust.


16


5.6. Geothermal gradient during metamorphism The investigated Malhag and Um Zariq metapelitic samples belong to the Barrovian garnet and staurolite metamorphic zones, respectively. Calculation of geothermal gradient for both samples at the peak metamorphic conditions was done assuming 1 kbar of pressure (corresponding to a rock pile 3.5 km thick). The calculated geothermal gradient is quite high, and lies in the range of 37–41 °C/km, which corresponds to metamorphism at burial depth of 14–16 km depth. Our calculated geothermal gradient range is higher than typical Barrovian geothermal gradient of 27–29 °C/km (Wright, 1989). The calculated geothermal gradient for the KMC is in agreement with that determined for the Sa'al-Zaghra metamorphic complex (Hassan et al., 2014; 38–41 °C/km). However, it is higher than the reported geothermal gradient for the metamorphic peak in the Feiran–Solaf metamorphic complex (27 °C/km, Abu El-Enen, 2011). Nevertheless, close geothermal gradient is documented for the isothermal decompression in the Feiran–Solaf metamorphic complex (43 °C/km, Abu El-Enen, 2011), which is attributed to crustal thinning with magmatic input of extensional related magmatism (Abu El-Enen, op cit). Taking into consideration the close age of metamorphism of the metapelites of Feiran–Solaf metamorphic complex (627 ± 7–592 ± 10 Ma; Abu ElEnen and Whitehouse, 2013), the metamorphism of the KMC was contemporaneous with the peak and isothermal decompression in Feiran– Solaf metamorphic complex (Abu El-Enen, 2011), and synchronous with heat flow from concurrent magmatism. On the other hand, the estimated geothermal gradient is slightly lower than that reported in Taba metamorphic complex (Abu El-Enen et al., 2003; Cosca et al., 1999) and Abu-Barqa Metamorphic Suite, SW Jordan (Jarrar et al., 2013), where the geothermal gradient therein (40–55 °C/km) has been concluded to be related to elevated thermal environment in an active arc setting. Our inferred geothermal gradient suggests that the peak regional metamorphism of the Um Zariq and Malhaq formations was generated during extensional regime and thinning of the lithosphere in an island arc setting with heat flow from the underlying arc granitoids. 6. Conclusions New LA-ICP-MS trace element analyses of detrital zircons and electron microprobe data of minerals from two metapelitic samples belonging to the Neoproterozoic Um Zariq and Malhaq formations of the Kid metamorphic complex led to the following conclusions: (1) The Older Cryogenian Um Zariq formation metasediments were subjected to Barrovian metamorphism within the garnet and staurolite zones of amphibolite facies, whereas the subsequent Ediacaran Malhaq Formation metasediments have suffered relatively lower metamorphic conditions down to the upper greenschist facies. (2) Even though the detrital zircons of Um Zariq Formation are more REE-enriched with total REE contents nearly as much as those from Malhaq Formation, zircons separated from the latter are markedly HREE-enriched and display strongly fractionated HREE patterns. Also, the quite differences in the overall slope and size of the Eu and Ce anomalies between REE patterns of the two zircon suites provide robust indication of different sources or subsequent modification by post-magmatic processes and/or metamorphism. (3) The application of the Ti-in-zircon thermometer (Watson et al., 2006) to the zircons separated from Malhaq Formation metapelite sample returned temperatures within the 916– 1018 °C range (average TTi-in-zrc = 967 °C), while those from Um Zariq schist reveal another higher range of zircon crystallization temperatures (1084–1154 °C) that cluster around 1120 °C. (4) The zircons of Um Zariq Formation are likely to be derived from more evolved igneous rock types due to their nearly two-fold

(5)

(6)

(7)

(8)

(9)

higher REE abundances relative to Malhaq zircons. Application of the “short” CART tree of Belousova et al. (2002) to our deitrital zircons has revealed that Malhaq Formation zircons are mainly derived from mafic source rocks (basalt and dolerite; 24 zircon grains out of 26), whereas Um Zaraiq Formation zircons have varied and more evolved parent rocks including high SiO2 (70–75 wt.% ) granitoids. The modest Niobium and Yttrium concentrations in all of the analyzed zircons (b 100 ppm and 10,000 ppm, respectively) rule out involvement of A-type granites among the granitoid source rocks of both formations. The high zircon crystallization temperatures shed doubt on formation through metamorphism. Also, none of the studied zircon samples have the characteristic Th/U signature of metamorphic origin (b0.1) and accordingly all are most plausibly magmatic. Various LREE and trace element features and Ti-in-zircon thermometry indicate the late crystallization from a high LREE/ HREE melt is probably due to the earlier separation of feldspar, hornblend, biotite and early crystallized zircon for most of the investigated zircons from both formations. Nevertheless, some zircons from Malhaq Formation show more likelihood of hydrothermal alteration. Possibility of crystallization from mantle material was eliminated and, without exception, all of the detrital zircon grains in the present study show typical trace elements features of crustderived zircons. Apart from few zircons from Malhaq Formation that are likely derived from oceanic crust, all of the Um Zariq Formation and most of Malhaq Formation detrital zircons are geochemically discriminated as continental zircons. Employing various conventional geothermometers and geobarometers using chemical composition of minerals in equilibrium at peak metamorphic P-T conditions, Um Zariq and Malhaq metapelites record similar overlapping peak temperatures (537–602 °C and 550–579 °C, respectively), and pressures (3.83–4.93 kbar and 3.69–4.07 kbar, respectively). These inferred metamorphic P-T conditions are in agreement with the previous peak P-T estimates determined for the KMC. The geothermal gradient for both samples, at the peak metamorphic conditions, is quite high and lies in the range of 37–41 °C/km, which corresponds to metamorphism at burial depth of 14–16 km. The peak regional metamorphism of the Um Zariq and Malhaq formations was generated during extensional regime and thinning of the lithosphere in an island arc setting with heat flow from the underlying arc granitoids.

Acknowledgment We thank Simon Wilde and Xin Zhou, respectively, for arranging and conducting the geochemical analyses of zircon at the Institute of Geology and Geophysics, Chinese Academy of Sciences in Beijing. We gratefully acknowledge the constructive comments of Yaron Be'eri-Shlevin that seriously improved the manuscript. In addition, we greatly appreciate the editorial handling and comments of Professor Jean-Philippe Avouac (Editor-in-Chief; Tectonophysics). The first author is thankful to Mahmoud M. Hassan for the discussions on the metamorphism of the KMC. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.tecto.2015.09.036. References Abu El-Enen, M.M., 2008. Geochemistry and metamorphism of the Pan-African back-arc Malhaq volcano-sedimentary Neoproterozoic association, Wadi Kid area, SE Sinai, Egypt. J. Afr. Earth Sci. 51, 189–206.


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