U–Pb zircon geochronology and Nd–Hf–O isotopic systematics of the Neoproterozoic Hadb adh Dayheen ring complex, Central Arabian Shield, Saudi Arabia

    U-Pb zircon geochronology and Nd-Hf-O isotopic systematics of the Neoproterozoic Hadb adh Dayheen ring complex, Central Arabian Shield, Saudi Arabia Kamal A. Ali, Heejin Jeon, Arild Andresen, Shuang-Qing Li, Hesham M. Harbi, Ernst Hegner PII: DOI: Reference:

S0024-4937(14)00267-9 doi: 10.1016/j.lithos.2014.07.030 LITHOS 3358

To appear in:

LITHOS

Received date: Revised date: Accepted date:

19 October 2013 23 July 2014 25 July 2014

Please cite this article as: Ali, Kamal A., Jeon, Heejin, Andresen, Arild, Li, Shuang-Qing, Harbi, Hesham M., Hegner, Ernst, U-Pb zircon geochronology and Nd-Hf-O isotopic systematics of the Neoproterozoic Hadb adh Dayheen ring complex, Central Arabian Shield, Saudi Arabia, LITHOS (2014), doi: 10.1016/j.lithos.2014.07.030

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ACCEPTED MANUSCRIPT U-Pb zircon geochronology and Nd-Hf-O isotopic systematics of the Neoproterozoic Hadb adh Dayheen ring complex, Central Arabian Shield, Saudi

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Arabia

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Kamal A. Ali a,*, Heejin Jeon b, Arild Andresen c, Shuang-Qing Li d,e, Hesham M.

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Harbi a , Ernst Hegner d

a

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Department of Mineral Resources and Rocks, Faculty of Earth Sciences, King Abdulaziz University, P.O. Box 80206, Jeddah 21589, Kingdom of Saudi Arabia b

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Department of Geosciences, Swedish Museum of Natural History, P.O. Box 50007, SE-104 05 Stockholm, Sweden

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c

d

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Department of Geosciences, P.O. Box 1047, University of Oslo, Blindern, 0316 Oslo, Norway Department of Geo- and Environmental Sciences and Geo-BioCenter, University of Munich, D-80333 München, Germany e

University of Science and Technology of China, School of Earth and Space Sciences,

230026 Hefei, China

Manuscript: text; 7 Figures; 3 tables and supplementary data (3 Tables) *

Corresponding author: E-mail [email protected]

Telephone: +966 2 640 0579; Fax: +966 2 695 2095

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ACCEPTED MANUSCRIPT Abstract A combined study of single zircon U-Pb dating, Hf-O zircon isotopic analyses

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and whole-rock Nd isotopic compositions was carried out to infer the magma sources of

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Neoproterozoic post-collisional A-type granitoids in Saudi Arabia. U-Pb zircon dating of magmatic zircons of two samples from the Hadb adh Dayheen ring complex yielded

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ages of 625 ± 11 Ma for a hornblende-biotite granite sample, and 613 ± 4 Ma for a

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monzogranite sample. The granitic rocks show initial Nd values of +4.1 to +5.3 and Hf of +4.5 to +8.4 that are lower than those of a model depleted mantle (Hf ~ +14 and Nd

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~ +6.5) and consistent with melting of subduction-related crustal protoliths that were formed during the Neoproterozoic assembly of the Arabian-Nubian Shield (ANS).

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Crustal-model ages (Hf-tNC) of 0.81 to 1.1 Ga are inconsistent with depleted-mantle Nd model ages of 0.71 to 0.81 Ga and indicate that the post-collisional Hadb adh Dayheen

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granites were derived mostly from juvenile crust formed in Neoproterozoic time. Single zircons data show a wide range in δ18O values from +3.2‰ to +6.4‰, possibly indicating crystallization of zircon from magma derived from magmatic rocks altered by meteoric water in a magma chamber-caldera system.

Keywords: Arabian-Nubian Shield; Sm-Nd isotopes; single zircon Hf-O isotopes; Atype granites; U-Pb zircon dating

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ACCEPTED MANUSCRIPT 1. Introduction There is no general agreement about the origin of the A-type alkaline granites

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(Bonin, 2004). Petrogenetic models include: (1) mixing of magma derived from

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juvenile crust and older crust (Eby, 1990, 1992; Kemp et al., 2006; Zhang et al., 2012); (2) partial melting of lower crustal sources (Collins et al., 1992; Frost and Frost, 1997;

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Huang et al., 2011; Patiňo Douce, 1997); or (3) extreme differentiation of mantle-

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derived tholeiitic or alkaline basaltic magma (Bonin and Giret, 1990; Mushkin et al., 2003; Turner et al., 1992). Alkaline A-type granites are commonly inferred to be

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emplaced in anorogenic settings (Bonin, 2007), and often form ring complexes such as the Cretaceous Ossipee Mountain in the White Mountain Magma Series of New

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Hampshire (Billings, 1945), the Jurassic Kudaru Hills in the younger Granites of Northern Nigeria (Jacobson et al., 1958), the Pikes Peak batholith of Colorado (Barker

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et al., 1975), and as well as many Precambrian ring complexes in Saudi Arabia (Moghazi et al., 2011). However, not all ring complexes contain A-type alkaline rock suites (Bonin, 2007).

The Arabian-Nubian Shield (ANS, Fig. 1a) is dominated by Neoproterozoic crust that was formed from 900 to 550 Ma by the accretion of intra-oceanic arcs during closure of the Mozambique Ocean and amalgamation of Gondwana (Ali et al., 2010a; Collins and Pisarevsky, 2005; Johnson and Woldehaimanot, 2003; Stern, 1994; Stern and Johnson, 2010). The ANS records ~300 Ma of orogeny from intra-oceanic subduction, arc and back-arc magmatism (~870 to 700 Ma), through terrane amalgamation (~800 to 650 Ma) to terminal collision between major fragments of East and West Gondwana. This was followed by tectonic escape, strike-slip faulting,

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ACCEPTED MANUSCRIPT delamination, and extension of the newly formed continental crust at 630 to 550 Ma (Avigad and Gvirtzman, 2009; Genna et al., 2002; Hargrove et al., 2006a, 2006b;

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Johnson et al., 2011; Kröner, 1985; Stoeser and Camp, 1985). These events shaped the

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northern part of the East African Orogen (Johnson and Woldehaimanot, 2003; Stern, 1994).

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The ANS is characterized by four main rock associations: predominantly

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juvenile arc-related supracrustal sequences, ophiolites, and gneiss terranes with core complexes and granitoid intrusions (Abd El-Naby et al., 2000; Shalaby et al., 2005).

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These lithotectonic units were tectonically intercalated by thrusting during accretion and left-lateral transcurrent movement along the Najd and other NW-striking shear

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zones at ~600 Ma (Johnson and Woldehaimanot, 2003).

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I-type plutons (Cryogenian and Early Ediacaran) are major components of the

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ANS; “arc assemblage intrusives” make up ca. 30% of the exposed ANS (Johnson et al., 2003). These are metaluminous to slightly peraluminous calc-alkaline, subductionrelated arc batholiths such as the Nabitah batholiths and peripheral belts and gneiss domes, which formed synchronously with large molasse basins (Genna et al., 2002; Johnson, 1998; Nehlig et al., 2002; Stern and Johnson, 2010). The orogeny was followed by ongoing convergence and tectonic escape from 610 Ma to 545 Ma (Genna et al., 2002; Greiling et al., 1994; Stern, 1994). The intrusion of post-collision alkaline to peralkaline plutonic rocks and formation of pull-apart basins (Johnson, 1998, 2003) marked crustal extension after the Neoproterozoic orogenic event, and the final suturing of the Arabian Shield (Fig. 1b; Nehlig et al., 2002). The alkaline to peralkaline granites constitute about 2 % of the

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ACCEPTED MANUSCRIPT Arabian Shield (Stoeser, 1986; Fig. 1b) and represent one of the largest fields of alkaline granites in the world (Stoeser, 1986). Rb-Sr radiometric ages suggest

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emplacement of these granites from 680 and 510 Ma (Stoeser, 1986). More recently U-

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Pb zircon ages of 620 Ma (Jabal Al-Hassir; Moufti et al., 2013), 613 Ma (Jabal Sitarah granite; Stoeser and Frost, 2006), 596 ± 10 Ma (Hanak granite; Hargrove et al., 2006a),

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and 574 Ma (Sawdah granitoids; Stoeser and Frost, 2006) to 566 Ma (Hail granitoids;

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Stoeser and Frost, 2006) have been reported for these granitoids. The alkaline granites generally form small intrusions (< 10 km2), although large plutons (e.g., Jabal Radwa,

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450 km2; Moufti et al., 2002) and batholiths (e.g., Aja batholith, 1600 km2; Qadhi, 2007) are also common. The alkaline to peralkaline granite intrusions form oval to sub-

(Küster, 2009).

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circular ring complexes (i.e. Harris, 1985) and many of them are rare-metal-bearing

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In several recent studies it has been shown that single zircon isotopic studies (UPb, Hf and O) give important information about the petrogenetic history of granitic rocks and reflect the contribution of mantle and crustal material in their protoliths (Goodge and Vervoort, 2006; Kovalenko et al., 2007; Stern et al., 2010). Zircons are not only the principal mineral for determining precise ages using the U-Pb isotopic system but are also the principal host mineral of Hf. These features, combined with low Lu/Hf ratios and their resistance to isotopic disturbance, make zircons ideal repositories of both age (U-Pb geochronology) and tracer (Lu-Hf isotope system) information (Belousova et al., 2006, 2010; Goodge and Vervoort, 2006; Kinny and Maas, 2003; Lenting et al., 2010). O isotopes in zircon are eminently suitable to recognize whether a rock underwent supracrustal processes or not (Be‟eri-Shlevin et al., 2009a, 2009c,

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ACCEPTED MANUSCRIPT 2010; Valley, 2003; Zheng et al., 2004, 2006). The δ18O record of non-metamict zircon is generally preserved from the time of crystallization despite high-grade

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metamorphism or hydrothermal alteration due to robustness of the mineral (Valley,

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2003; Valley et al., 2005). Therefore, the δ18O of zircon can be used to relate the magmatic environment of the zircon to the U-Pb age (Valley, 2003).

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Several studies have argued, based on mass balance calculations, that at least

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80% of the continental crust was created along destructive plate margins (Hawkesworth et al., 2010; Rudnick, 1995). Previous isotopic studies, focusing on crust-forming

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processes, have been carried out on rocks from the ANS (Ali et al., 2009, 2010a, 2010b, 2013; Hargrove et al., 2006a, 2006b; Lundmark et al., 2012; Stern et al., 2010). Some

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of these studies have concluded that parts of the inferred ANS crust are not as juvenile as previously thought (Ali et al., 2009, 2013; Hargrove et al., 2006b; Morag et al.,

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2011, 2012). Other workers have argued that some of the gneiss domes in the Eastern Desert of Egypt represent a pre-Neoproterozoic (pre-Pan-African) basement (e.g., ElGaby et al., 1984, 1988, 1990; Hamimi et al., 1994; Khudeir et al., 1995, 2008). For this purpose, we have studied U-Pb ages and Hf-O isotopic compositions in single zircons in addition to whole-rock Sm-Nd isotopic analyses on late Neoproterozoic Atype granitic rocks from Hadb adh Dayheen ring complex in the central Arabian Shield of western Saudi Arabia.

2. Geology and petrography of Hadb adh Dayheen ring complex Western

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contains

late

Precambrian

rock

assemblages

characterized by widespread post-collisonal granitoid ring complexes with A-type

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ACCEPTED MANUSCRIPT granites as an important rock-type.These are arcuate bodies interpreted to represent successive igneous intrusions in the form of ring-dykes, cone sheets, plugs, and diapirs

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(Roobol and White, 1985; Shang et al., 2010; Vail, 1970, 1976). The geology of Hadb

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adh Dayheen ring complex (Fig. 2) has been described by Radain et al. (1981) and Moghazi et al. (2011). It occupies an area of ~ 117 square kilometers and is located

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between latitudes 23º30‟ and 23º37‟ N, and longitudes 41º 70‟ and 41º 15‟ E (Fig. 2). It

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is located ~30 km northeast of the Mahd adh Dahab gold mine. The Hadb adh Dayheen complex was intruded into an older diorite-granodiorite

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suite. It has a central wadi similar to many cauldron subsidences in the Arabian Nubian Shield (Hassanen, 1997; Roobol and White, 1985; Shang et al., 2010; Vail, 1976,

alkali feldspar granite.

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1983). The main rock types that constitute the ring structure are monzogranite and

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The monzogranite comprises the core of the complex and occurs as relicts in low-lying land due to its highly weathered nature. It is coarse- to medium-grained and occasionally porphyritic with pale pink to grayish pink color. It is composed of plagioclase, quartz, perthite and biotite. Plagioclase feldspar crystals (33 to 36 %) occur as subhedral twinned grains. Quartz (34 to 40 %) occurs as interstitial grains or enclosed in feldspar crystals. Alkali feldspar (20 to 23 %) occurs as anhedral grains of microperthite interstitial between plagioclase grains. Biotite crystals are subhedral and strongly pleochroic with tiny inclusions of zircon, forming pleochroic haloes. Few crystals of biotite are partially altered to green chlorite. Titanite is the most abundant accessory mineral, while zircon and apatite are less abundant. The contacts between the

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ACCEPTED MANUSCRIPT monzogranite and the other granitoids of the complex are not exposed due to the highly weathered and eroded intrusive rocks.

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The alkali feldspar granite body consists predominantly of a core of porphyritic

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hornblende biotite granite partly mantled by porphyritic aegerine riebeckite granite. Locally, the porphyritic hornblende-biotite granite is intruded by a microgranite (Fig.

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2), which has the same mineralogical composition, but a finer-grained texture.

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The porphyritic hornblende-biotite granite is fine-grained, massive, has a subsolvus groundmass and pink to reddish color. It is composed of consists of quartz,

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perthites, microcline, plagioclase, hornblende and biotite. Quartz occurs as fine-grained crystals in the groundmass and phenocrysts containing abundant inclusions of

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plagioclase laths. K-feldspar is represented by perthites (patch- and mesoperthites) and minor microcline. Both types of K-feldspar occur as phenocrysts (up to 1cm long) and

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as fine-to medium-grained anhedral crystals. Plagioclase feldspar is subhedral with prismatic to lath-like shape, which occasionally contains irregular fluorite crystals. Hornblende is subhedral prismatic and pleochroic and some crystals exhibit spongy form and sieve-like texture due to abundant inclusions (i.e. skeletal crystals). The biotite crystals are pale reddish brown and moderately pleochroic. Titanite, flourite, zircon, apatite and allanite are common accessory minerals. The porphyritic hornblende biotite granite also occurs as roof-pendant to the aegerine riebeckite granite. Moreover, numerous angular and flattened xenoliths up to 50 cm long of the porphyritic hornblende biotite granite are common constituents of the aegerine riebeckite granite. The latter is white to grayish white with an average thickness of ca. 2.5 km. It is porphyritic with large feldspar phenocrysts up to 2cm long and has a medium-grained

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ACCEPTED MANUSCRIPT groundmass. It consists of quartz, alkali feldspar, albite, aegerine and riebeckite. Accessory minerals are titanite, rutile, zircon, allanite, fluorite, and apatite. The rocks

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are inequigranular and exhibit porphyritic texture with perthite, quartz, and aegerine

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being the most common phenocrysts. The contacts with the porphyritic hornblende biotite granite and the other older country rocks are sharp and distinct. Pegmatitic

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dykes, quartz veinlets and few variolitic cavities are common in the aegerine riebeckite

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granite. The pegmatite dykes are of irregular form and size and range from few centimeters up 25m in length. Most of the pegmatite dykes appear in the southern,

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eastern, and northeastern parts of the complex.

The age of Hadb adh Dayheen ring complex is not well constrained. A Pb-

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model age of 500 Ma was inferred from galena collected in a small pegmatite vein

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(Goldsmith, 1971), a Rb-Sr age of 573  22 Ma has been reported for the alkaline-

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peralkaline granites (Kemp et al., 1982), and K-Ar ages of 574 and 557 Ma were reported for granite porphyry dikes and riebeckite granite, respectively (Radain et al., 1981).

3. Results

3.1. U-Pb zircon geochronology

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ACCEPTED MANUSCRIPT To constrain the age of the Hadb adh Dayheen ring complex, we have dated zircons from two granitic samples. Sample locations are shown in Fig. 2 and analytical

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results are listed in Table 1. Detailed analytical procedures are presented in the

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Appendix.

Sample HD-3 (23°34'1.50"N, 41°14'8.25"E) is a medium to coarse-grained

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hornblende-biotite granite. Zircons separated from it are euhedral (100 – 200 μm) and

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yellow to pale brown. Cathodoluminescence (CL) images show well-developed zoning as expected for magmatic zircons (Fig. 3; Corfu et al., 2003). U contents vary from

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134-1200 ppm, Th contents from 26-2200 ppm and Th/U from 0.10-4.2. We analyzed 20 zircon grains of this sample (Table 1) and excluded sixteen spot analyses from age

Pb/238U ratios indicative of Pb loss and/or are excessively discordant. The remaining

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calculation because they have a high common Pb content, show significantly lower

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Pb/238U ages of 625 ± 11 Ma

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four analyses are concordant (Fig. 4a) and yield mean (2ζ, n = 4, MSWD = 1.8).

Sample HD-30 (23°32'49.09"N, 41°12'3.94"E) is a medium to coarse-grained monzogranite. Zircon recovered from this sample is subhedral to euhedral and yellow to pale brown. Zircon grains are slender, and needle shaped, and exhibit well-preserved oscillatory growth zoning (Fig. 3). They have U contents from 147-2810 ppm, Th contents from 56-860 ppm and Th/U from 0.29-0.60 (Table 1). Out of twenty-five zircon grains analyzed (Fig. 4b), six grains yielded concordant mean 206Pb/238U ages of 613 ± 4 Ma (2ζ, n = 6, MSWD = 0.86) (Table 1). Eighteen analyses were excluded from age calculation because they yield much younger ages probably due to Pb loss, show a high common Pb content, and/or are discordant. One zircon analysis (spot 25)

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ACCEPTED MANUSCRIPT produced a concordant data point with 206Pb/238U age of 775 ± 6 Ma, interpreted as that

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of a zircon xenocryst derived from older source material.

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3.2. Sm-Nd isotope compositions

The Sm-Nd isotopic data for five granitoid samples of the ring complex are

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listed in Table 2 and initial εNd values are plotted in Fig. 5a. The isotopic data suggest a

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zoned composition of the ring complex: the two monzogranite samples (HD-22 and HD-30) from the core show large differences in Sm and Nd concentrations and Sm/Nd

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ratios probably due to variable magma fractionation of a common parental magma as indicated by their within analytical error identical initial εNd values of ~+5.3 (error

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estimated at ~0.25 ε-units, initial values calculated for an age of 610 Ma). These values are lower than those for the depleted mantle at ~600 Ma (~ +6.5, DePaolo model) and

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consistent with melting of crustal protoliths derived from subduction-overprinted mantle sources. Depleted mantle Nd model ages of 0.71 to 0.81 Ga suggest that the protoliths were formed during the late Proterozoic which is in agreement with the data of previously published studies (Ali et al., 2009, 2010c; Hargrove et al., 2006b; Liégeois and Stern, 2010; Moussa et al., 2008; Stoeser and Frost, 2006). A single sample of hornblende-biotite granite (sample HD-3) forming the mantle around the core of monzogranite shows a lower initial εNd value of +4.1 indicating different crustal sources with a minor proportion of older material and this sample gives a little older 2stage Nd model age of 0.81 Ga. Two samples of aegerine-riebeckite A-type granite (samples HD-9B and HD-17) forming the margin of the ring complex are isotopically identical to the monzogranitic core of the complex indicating a similar parental magma

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ACCEPTED MANUSCRIPT as for the core that was differentiated to a much higher degree. The initial εNd values are plotted in Fig. 5a together with data for the overall juvenile magmatic arc assemblages

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of the ANS (Ali et al., 2009, 2010c, 2012a, 2012b; Hargrove et al., 2006b; Liégeois and

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Stern, 2010; Moussa et al., 2008; Stoeser and Frost, 2006 ). The data of the postcollisonal A-type to I-type granitoid ring complex of this study plot on the ANS data

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envelope and suggest melting of arc root material. However, the isotopic data of the

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samples are not capable to resolve any mantle input in the ring complex (i.e., mantle source or recycling of juvenile arc crust). We suggest that involvement of upper mantle

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magmas during formation of the ring complex as heat and protolith source is very

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likely.

3.3. Hf isotopic compositions in zircons

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The Lu-Hf isotopic data for the two dated granitic samples from Hadb adh Dayheen ring complex are listed in Table 3. Twenty-two Lu-Hf isotopic analyses of single zircons are presented in Figs. 5b and 6. Ten spot-analyses were carried out on 10 zircon grains from sample HD-3 which were dated by SIMS, and the analytical data are shown in Table 3. As shown on the Hf isotopic evolution diagram of Fig. 6a, the zircons have initial Hf(t) values of +4.5 to +8.0 and crustal model ages (Hf-tNC) vary between 0.87 to 1.1 Ga. Twelve Lu-Hf spot analyses were done on zircon grains from sample HD-30 (Table 3) and yielded initial Hf(t) values of +6.6 to +8.3 (Fig. 6b). Hf crustal model ages (tNC) vary between 0.81 to 0.97 Ga (Table 3).

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ACCEPTED MANUSCRIPT In the Nd(t) versus Hf(t) in Fig. 5b, single zircon Hf(t) and whole-rock Nd(t) are plotted together with data fields for the whole-rock Nd(t) and single zircon Hf(t)

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analyses for eastern desert post-collisional granites and arc-metavolcanic rocks (Ali et

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al., 2012a, 2013; Be‟eri-Shlevin et al., 2010). The samples of this study extend the data

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field for post-collisional granitoids to lower Hf(t) and Nd(t) values and indicate heterogeneous sources for post-collisional granitoids with higher proportion of older

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3.4. Single zircon δ18O results

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recycled material in the ring complex of this study.

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The results of ion-microprobe δ18O analyses of the concordant single zircon

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grains for two granitoid samples from Hadb adh Dayheen ring complex are presented in Table 3 and Figure 7 along with Lu-Hf isotopic analyses for the same grains used for

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U-Pb age determination.

For hornblende-biotite granite sample HD-3 from the core-mantling lithology of the ring complex, twelve analyses yielded δ18O values of +3.2 to +5.7 ‰. The four grains of concordant U-Pb ages gave a relatively small δ18O range from 3.2 to 4.1 ‰, whereas those of younger ages showed notably higher and dispersed compositions (4.15.7 ‰). It might indicate source heterogeneity of the granite, but it is more reasonable not to consider the latter values probably disturbed by a syn- or post-magmatic event that also caused significant Pb loss from igneous zircons. The four concordant single zircon grains yield a mean δ18O value of +3.2 ± 0.65 ‰ that is more or less homogenous or slightly heterogeneous. One concordant analysis (spot 05) which is rejected for age calculation (206Pb/238U = 646 Ma) shows similar low δ18O value (+4.0 ±

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ACCEPTED MANUSCRIPT 0.3 ‰). The very low δ18O values in this sample (Fig. 7) relative to mantle values of ~ +5.3 ± 0.6 ‰ (Valley et al., 1998), can be explained with post-crystallization oxygen

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isotope exchange with late meteoric fluids during hydrothermal alteration of the ring

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complex or perhaps an isotopically heterogeneous magma formed from protoliths that were pervasively altered by meteoric fluids before melting.

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The δ18O values for single zircon grains of monzogranite sample HD-30 range

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from +5.2 to +6.4 ‰ (mean = +5.5 ± 0.2 ‰; MSWD = 1.4; n = 15; Fig. 7) and these values are similar to δ18O values in mantle-derived rocks (e.g., Valley et al., 1998). The

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O isotopic compositions do not vary and no difference was observed between the six concordant zircon grains which have been used in age calculation (+5.3 to +6.4 ‰) and

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4. Discussion

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the others (Table 3) unlike in the sample HD-3.

4.1. Geochronology of Hadb adh Dayheen granites The set of zircon U-Pb ages from this study enables us to infer the crystallization age of Hadb adh Dayheen ring complex at ~ 625 ± 11 to 613 ± 4 Ma. Our results found many discordant and Pb-loss affected zircons in the studied granites. This may be due to radiation damage in high U zircons that causes these to act as open systems (Ewing et al., 2003) subject to gain or loss of U, Pb. This damage may have been exacerbated by alteration due to circulation of late F-rich fluids, which can dissolve zircons (cf., Watson and Harrison, 1983). Combined effects of radiation damage and F-rich magmatic fluids could be responsible for elevated proportions of common Pb in dated zircons (Table 1).

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ACCEPTED MANUSCRIPT In summary, the proposed crystallization ages are in agreement with previous U-Pb zircon ages for the post-tectonic alkaline/peralkaline granites in the Arabian

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Shield: e.g., Jabal Al-Hassir ring complex (620 ± 3 Ma; Moufti et al., 2013); Jabal

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Sitarah granites (613 Ma; Stoeser and Frost, 2006); Hanak granite (596 ± 10 Ma; Hargrove et al., 2006a). Similar ages were reported in the northern ANS for the within-

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plate alkaline granites (608-580 Ma; Be‟eri-Shlevin et al., 2009b; Morag et al., 2011).

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The zircon Hf isotope data and Sm-Nd whole rocks data indicate a largely juvenile magma source of Neoproterozoic age and negate the involvement of large amounts of

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an older crustal component.

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4.2. Isotopic features of the A-type Hadb adh Dayheen granites There are few published Sm-Nd, Lu-Hf and O isotopic data for the post-

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collisional granites of the ANS. Published Sm-Nd whole-rock isotopic data, indicate that most of the ANS crust was derived from depleted (e.g. island arc-like) mantle sources (Ali et al., 2009, 2010c, 2012a, 2012b; Augland et al., 2012; Liégeois and Lundmark et al., 2012; Moussa et al., 2008; Stern, 2010; Stern et al., 2010). Depletedmantle Nd model ages and initial Nd(t) for ANS arc-generated and post-collisional rocks (Fig. 5a, b) are similar and have been interpreted as evidence for little or no preNeoproterozoic material in the lower and middle crust (Eyal et al., 2010; Stern, 2002; Stern et al., 2010; Stoeser and Frost, 2006). This inference is supported by U–Pb dating of single zircons from various plutonic rocks (Be'eri-Shlevin et al., 2009c) and by RbSr and Sm-Nd radiogenic isotopic studies of whole-rock Neoproterozoic granite-gneiss samples from the Eastern Desert, Egypt (Liégeois and Stern, 2010).

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ACCEPTED MANUSCRIPT The analysis of Hf isotopes in zircons (Belousova et al., 2006; Patchett et al., 1981; Scherer et al., 2000, 2001; Vervoort et al., 1996) has advantages over the Sm-Nd

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whole-rock system as a tracer of magma source and petrogenetic processes (Belousova et al., 2010; Dickin, 1995) because (1) zircons have high Hf and low Lu concentrations, 176

Lu/177Hf, so that their present-day Hf isotopic compositions approximate

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hence low

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those of magmas from which the zircons crystallized (Kinny and Maas, 2003). These

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zircons have 176Hf/177Hf ratios consistent with evolution in a reservoir with low or high Lu/177Hf, i.e. continental crust or mantle; (2) Hf is an essential structure element in

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the crystal structure of zircon (Patchett et al., 1981); in addition, zircon is very robust with remarkable resistance to re-equilibration of the Hf isotopic composition (Watson,

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1996; Watson and Cherniak, 1997). Therefore, its Hf isotopic composition is generally not disturbed by magmatic processes or high-grade metamorphism (e.g., Huang et al.,

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2006; Lenting et al., 2010).

The initial Nd values of the studied granitoid samples, on first inspection, appear to be a coherent group as plotted in Fig. 5a. The granitic rocks are characterized by initial Nd, and Hf isotopic ratios (Nd of +5.2 to 5.4 and Hf of +3.6 to +8.0) that are lower than those predicted for depleted mantle values at ~ 600 Ma [Hf of about +14, Nd of about +6.5 using depleted mantle evolutionary line of DePaolo, (1981)] (Figs. 5, 6). The positive Hf and Nd values of the post-collisional granitic samples indicate derivation mainly from a juvenile Neoproterozoic mantle source, but also suggesting the involvement of at least some older crustal material in the formation of ANS crust. This is in agreement with the extensive zircon U-Pb and Hf data set reported from northern Sinai (Morag et al., 2011, 2012). They show that the average εHf values in

16

ACCEPTED MANUSCRIPT the post-collisional rocks are all positive and range between +6 to +9, consistent with mantle source and juvenile crust reworking.

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Avigad and Gvirtzman (2009) suggested delamination of the lithospheric root as

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the mechanism of the post-collisional magmatism. Replacement of parts of the lithospheric mantle by hot asthenosphere strongly caused partial melting at the base of

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the remaining lithospheric mantle and the generated mafic melts pass through the base

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of the crust to induce crustal melting and post-collisional magmatism (Morag et al., 2011). Previous geochemical data set is further evidence that the mantle-crust magma

MA

mixing was important during the formation of the post-collisional Hadb adh Dayheen magmatism (Moghazi et al., 2011).

D

The δ18O values of the single zircons of the two studied granitic samples (Table

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3) are not similar. Sample HD-30 (monzogranite) displays zircon δ18O values of ca. 5.5

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‰, which is mantle-like (+5.3 ± 0.6 ‰; Valley et al., 1998). In contrast, sample HD-3 (hornblende-biotite granite) yields δ18O values of ca. 3.2 ‰, lower than δ18O mantle values. Zircon which crystallized from magmas that incorporated crustal material previously affected by low-temperature alteration has values above +6.5 ‰ (Valley et al., 2005). Zircon with low δ18O below-mantle values (< 4.6 ‰) is more difficult to explain (van Schijndel, 2013). It could originate from magmas formed by melting of rocks that were hydrothermally altered by meteoric water at high temperature (>300 °C), such as lower oceanic crust (van Schijndel, 2013; Wei et al., 2002) or crystallized from low δ18O magma (Liu and Zhang, 2013). The retention of initial magmatic δ18O values of zircon relies on several factors, including temperature and cooling rate in the process of zircon-fluid interaction and zircon degree of metamictization (Liu and

17

ACCEPTED MANUSCRIPT Zhang, 2013). High metamictization can lead to zircon recrystallization during zirconfluid interaction by dissolution-reprecipitation to produce porous zircons with a mottled

PT

CL texture (Booth et al., 2005; Liu and Zhang, 2013; Tomaschek et al., 2003). Such zircons display highly discordant ages, decreasing Th/U and sometimes systematic

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change of δ18O values (Booth et al., 2005). The considered zircon grains for O isotopic

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compositions of HD-3 were concordant and not affected by serious Pb loss. Also zircon

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δ18O values are not correlated with their U and Th contents (Table 3), suggesting that their low δ18O values are not related to zircon metamict degree. Therefore, we prefer

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the suggestion that low δ18O values must be related to low δ18O magmatic origin. Low δ18O values of single zircon grains of Hadb adh Dayheen ring complex may be formed

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through the re-melting of previous magmatic rocks altered by meteoric water in a

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system of magma chamber-caldera, as demonstrated for the low δ18O of the rhyolites at

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Yellowstone (Bindeman and Valley, 2001; Bindeman et al., 2008) and the plutonic complex to the north of the Yaolinghe Group (Liu and Zhang, 2013). Caldera collapse brought altered previous magmatic rocks into magma chamber, which resulted to their re-melting to produce low δ18O magma. Previous study shows that whole-rock δ18O values of Hadb adh Dayheen monzogranite have a uniform isotopic composition (+8.1 ± 0.2 ‰; Radain et al., 1981), whereas riebeckite-bearing peralkaline granites exhibit scattered δ18O values from +7.7 to +9.6 ‰ (Radain et al., 1981). The elevated δ18O values require participation of supracrustal material in the petrogenesis of the Hadb adh Dayheen rocks, during melting or assimilation (Be'eri-Shlevin et al., 2009a, 2009c, 2010; Taylor, 1978; Taylor et al., 1981) or may indicate variations in the extent of crustal material involved or due

18

ACCEPTED MANUSCRIPT to isotopic exchange with high δ18O rocks during generation and emplacement of the

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magmas (Radain et al., 1981).

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4.3. Nd and Hf model ages

The interpretation of the zircon Hf-isotopic data in terms of magma source

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composition is less straightforward than the model age calculation using Sm-Nd whole-

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rock data because the Lu-Hf ratio of the zircon is not that of the host magma. Depleted mantle model (TDM) ages assume that the magma or its precursor separated from a

MA

depleted mantle source (Be'eri-Shlevin et al., 2010).

Calculating Nd-tDM using the model of DePaolo (1981) is preferred over those

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D

of Goldstein et al. (1984) for rocks formed in arc-settings as have been inferred for the ANS (Ali et al., 2009; Be'eri-Shlevin et al., 2010; Hargrove et al., 2006b; Stern, 2002).

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The Nd-tDM ages calculated after DePaolo (1981) for the Hadb adh Dayheen samples (Table 2) range from 0.71 to 0.81 Ga, consistent with melting of juvenile crust as has been documented elsewhere in the ANS (Liégeois and Stern, 2010; Stern et al., 2010), and in agreement with the absence of inherited zircon cores. A Hf depleted-mantle model age (Hf-tDM) for the magmatic host rock of a given zircon is calculated using the measured

176

Hf/177Hf and

176

Lu/177Hf of the zircon. This

gives a minimum age for the source rock of the host magma (e.g., Kemp et al., 2006, 2007). Mogahzi et al. (2011) suggested that magmas of Hadb adh Dayheen were derived from juvenile lower crust; thus a more realistic "crustal" model age (Hf-tDMC) may be calculated by assuming that the source rocks of the magma had the

176

Lu/177Hf

ratio of average crust (0.015; Griffin et al., 2002). Zircons in samples HD-3 and HD-30

19

ACCEPTED MANUSCRIPT yielded tDMC ages from 1.1 to 1.3 and 1.0 to 1.2 Ga, respectively, suggesting that according to this mantle model their crustal precursor contained old crustal material

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(Fig. 6). Discrepancies between whole-rock Nd model ages (Nd-tDM) and Hf model age

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(Hf-tDMC) is another problem that needs to be addressed. Traditionally, Hf model ages

SC

are calculated considering the isotopic composition of the depleted mantle similar to

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that of new continental crustal material (Dhuime et al., 2011). However, the isotopic composition of island arc rocks (e.g., ANS rocks) and hence the new continental crust

MA

is different from that of the depleted mantle (Dhuime et al., 2011). The best estimate for the present-day composition of average new crust formed at destructive plate margins

D

and the corresponding mantle source maybe inferred from the weighted average of 13

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modern island arcs worldwide with εHf of 13.2 ± 1.1 (Dhuime et al., 2011). Thus, the

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model ages (Hf-tNC) calculated from the composition of the new crust are younger than model ages inferred from the depleted mantle model (Fig. 6; Table 3). Using the latter model, zircons of samples HD-3 and HD-30 yielded tNC ages of 0.89 to 1.1 Ga and 0.83 to 0.97 Ga, respectively. The better agreement of these values with the Nd model ages derived from the DePaolo model (1981) support our assumption that the Hf isotope evolution model of Dhuime et al. (2011) may be more applicable to infer model ages for Hf-in-zircon than the one for the depleted mantle of Vervoort and Blichart-Toft (1999), an inference that intuitively is plausible considering the subduction-modified mantle composition under magmatic arcs. We conclude that the “new-crust” model ages (Hf-tNC) are consistent with the Nd whole-rock model ages (Nd-tDM) and support a mostly juvenile Neoproterozoic crust (Fig. 6; Table 3).

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ACCEPTED MANUSCRIPT In summary isotopic data of the post-collisional granitic from Hadb adh Dayheen ring complex indicate derivation from a juvenile Neoproterozoic mantle

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source and some older crustal material supporting the petrogenetic mixing magma

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model of Eby (1990, 1992), Kemp et al. (2006), and Zhang et al. (2012).

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5. Conclusions

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The following conclusions can be drawn from this study along with published data: 1. Magmatic zircons from Hadb adh Dayheen ring complex yielded

206

Pb/238U

MA

ages of 625 ± 11 and 613 ± 4 Ma interpreted to represent the crystallization age of the granites.

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2. Initial Nd of +4.1 to 5.3 and Hf of +4.5 to +8.4 of the post-collisional granitoid samples are consistent with melting of arc-related crustal protoliths

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comprising the ANS and contributions from subduction-overprinted mantle sources from which the ANS was derived during the Neoproterozoic. 3. The Hf-crustal ages (Hf-tNC) using the present-day composition of average of modern island arcs (Dhuime et al., 2011) are consistent with the depleted mantle Nd model ages of DePaolo (1981) and indicate melting of predominantly late Neoproterozoic crustal protoliths. 4. Average δ18O values of +5.7 for domains in single zircons of monzogranite sample HD-30 are like those in mantle-derived rocks whereas average δ18O value of +3.56 ± 0.65 in hornblende-biotite granite sample HD-3 is interpreted as due to crystallization of zircon from magma derived from magmatic rocks altered by meteoric water in a magma chamber-caldera system.

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ACCEPTED MANUSCRIPT 5. Nd-Hf-O isotopic compositions of Hadb adh Dayheen ring complex indicate heterogeneous magma sources and support the petrogenetic model of Eby

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(1990, 1992), Kemp et al. (2006), and Zhang et al. (2012).

Acknowledgments

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The authors would like to thank the Geosciences Department at University of Oslo

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(Norway) for performing zircon Hf- isotope analyses. We sincerely thank Swedish Museum of Natural History for analyzing U-Pb and O isotope. The NordSIM ion

MA

microprobe facility is financed and operated under an agreement between the research councils of Denmark, Sweden, and the Geological Survey of Finland and the Swedish

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Museum of Natural History. We are grateful to Dr. Cosmas Shang, an anonymous reviewer and Editor-in-Chief Prof. Nelson Eby for critical reviews that improved this

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manuscript. This is Nordsim contribution # .

Appendix : Analytical techniques Based on previous petrographic and chemical investigations (Mogazi et al., 2011), representative samples, covering the different granite varieties (Table A1 – supplementary data), were selected for U-Pb zircon age determinations and Nd, Hf and O isotopic compositions. Sample locations are shown in Fig. 2. These data are listed in Tables 1, 2 and 3. U-Pb zircon age determinations, zircons were separated using standard procedure: crushing, heavy liquids and magnetic separation. Zircon from the least

22

ACCEPTED MANUSCRIPT magnetic fractions were handpicked under a binocular microscope in distilled isopropyl alcohol to obtain the most transparent and inclusion-free grains. U-Th-Pb analyses were

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performed by secondary ion mass spectrometry (SIMS) using a CAMECA IMS-1280

RI

instrument at the Swedish Museum of Natural History, Stockholm, Sweden (Nordsim facility) closely following methods outlined by Whitehouse and Kamber (2005). Zircon

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grains from the non-magnetic fractions were mounted on glass slides by using double –

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sided tape, cast in epoxy and polished to expose their internal structure-zoning. Prior to the analysis the zircon grains were photographed at high magnification and imaged by

MA

cathodoluminescence (CL) using a scanning electron microscope. Grain mounts were then washed, rinsed in distilled water, dried in a vacuum oven, and coated with gold.

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An O2- primary ion beam of 23 kV with 10 nA O2- ion was focused on the surface to a

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≤ 25 μm diameter spot, and the zircon sputtered under an oxygen flooding to enhance

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the secondary ion yield of Pb. Measurements were made at mass resolutions of 5000. Pb/U ratios were calibrated using an empirical power law relationship between measured Pb/U and UO2/U ratios. The 1065 Ma Geostandard zircon 91500 (Wiedenbeck et al., 1995), which was used both for Pb/U calibration and U concentration estimates, was analyzed frequently during the analytical session. A total of 45 spots were analyzed and analytical uncertainties are listed in the Table 1 and plotted as two-sigma error ellipses on concordia diagrams (Terra and Wasserburg, 1972). Ages mentioned in the text are given with two sigma error and MSWD values on concordia ages are those of combined concordance and equivalence. Calculations use the routines of Isoplot (Ludwig, 2001).

23

ACCEPTED MANUSCRIPT The Sm–Nd isotopic analyses were carried out at Munich University (LMU), Germany. Sample powders were digested in HF-HClO4 in PFA vessels for three days,

PT

and Sm and Nd were separated using chromatographic procedures as outlined in

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Hegner et al. (1995). Sm and Nd isotopic compositions were measured on a Spectromat-upgraded MAT 261 using a dynamic 2-line triple mass collection mode and

146

Nd/144Nd = 0.512108 ± 8 (2ζ

Nd/144Nd = 0.7219 for the JNdi-1 standard. The La Jolla Nd

NU

mean) normalized to

143

SC

monitoring 147Sm. During this study the MAT yielded

standard solution yielded over a longer period

143

Nd/144Nd = 0.511847 ± 8 (2SD, N =

MA

10). Two analyses of the USGS rock standard BCR-1 (Basalt Columbia River) are listed in Table 2. The external reproducibility of the

143

Nd/144Nd ratio measured on the 147

Sm/144Nd = 0.1960 and

Nd/144Nd = 0.512630 for the present chondritic uniform reservoir (Bouvier et al,

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143

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MAT is ca. 1.1x10-5. The εNd values were calculated with

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2008).

Concordant zircon spots previously analyzed by SIMS using a CAMECA IMS1280 instrument were selected for Hf isotopic analyses on a Neptune MC-ICP-MS (Nu Instruments, UK), coupled to a 193 nm excimer laser ablation system (Resolution M50, Resonetics LLC, USA), installed in the Department of Geosciences, University of Oslo. Data were acquired by ablating 55 μm (diameter) laser spots, and a 10 Hz repetition rate was used. The analyses consisted of a 30 s blank measurement prior to the start of ablation and 40 s of ablation. The 179

176

Hf/177Hf ratios were normalized to

Hf/177Hf = 0.7325, using an exponential law for mass fractionation. Isobaric Lu on

176

respectively. The in-situ measured

173

interference of

176

Yb and

176

Hf was corrected by monitoring

172

Yb and

175

Lu,

Yb/172Yb ratio was used for mass bias correction

24

ACCEPTED MANUSCRIPT for both Yb and Lu because of their similar physicochemical properties. Ratios used for the corrections were 0.5887 for

176

Yb/172Yb and 0.02655 for

176

Lu/

175

Lu (Vervoort et

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al., 2004). External corrections were applied to all unknowns, and standard zircons

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Mud Tank, Temora-2 and Malawi-1 were used as external standards (Table A2 – Supplementary data) and were analyzed frequently during the analytical session. The 177

Hf ratios and a

176

Lu decay constant of 1.867 x 10-11 as reported by

SC

measured 176Lu/

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Scherer et al. (2001, 2007) were used to calculate initial 176Hf/ 177Hf ratios. Calculation of εHf values is based on the chondritic values of

176

Hf/177Hf and

176

Lu/177Hf as

MA

reported by Blichert-Toft and Albarède (1997). The mantle extraction model age (TDM) was calculated using the measured

176

Lu/177Hf of the zircon, but this only provides a

D

minimum age for the source material of the magma from which the zircon crystallized.

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Therefore, we also calculated „crustal‟ model ages tc, (Table 3), which assume that the

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parental magma of the zircons was produced from an average continental crustal source with (176Lu/177Hf = 0.015 (Griffin et al., 2000, 2002) that was ultimately derived from the depleted mantle. Crustal model ages tNC, (Table 3) are calculated for the new crust assuming a value of 0.0113 for average continental crust, and a juvenile crust 176

Lu/177HfNC and

176

Hf/177HfNC of 0.0384 and 0.283165, respectively (Chauvel et al.,

2008). The depleted mantle line in Fig. 6 is defined by present-day 0.28325 and

176

176

Hf/177Hf =

Lu/177Hf = 0.0384 (εHf = 17; Griffin et al., 2002, 2004; Vervoort and

Blichert-Toft, 1999) and NC evolution curve is the line evolution from a uniform reservoir (CHUR) to εHf =13.2 for the present-day island-arc crust (Dhuime et al., 2013).

25

ACCEPTED MANUSCRIPT Oxygen isotope ratios of concordant zircon grains previously analyzed for U-Pb ages were measured using a CAMECA IMS-1280 instrument at the Swedish Museum

PT

of Natural History, Stockholm, Sweden. Analytical procedures follow Be'eri-Shlevin et

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al. (2010). Ion-microprobe δ18O was performed on twenty eight zircon grains for two samples from Hadb adh Dayheen ring complex (HD-3 and HD-30) in the same U-Pb

SC

dated spots or very close proximity. Prior to ion-microprobe δ18O analysis the U-Pb

NU

analysis spots were removed from the zircons by minor polishing follow by recoating with ~ 30 nm gold. Results from the unknown are presented in Table 3. Geostandards

MA

zircon 91500 (mean = 9.86 ± 0.12; n = 18; MSWD = 0.57) and Temora-2 (mean = 8.17 ± 0.18; n = 7; MSWD = 0.82), were analyzed frequently during the analytical session

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TE

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(Table A3 – supplementary data).

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Figure Captions: Fig. 1: (a, inset) Geological sketch map outlining the Arabian-Nubian Shield. (b) Simplified map of the Arabian Shield, showing major tectonostratigraphic terranes, ophiolite belts, sutures and fault zones, and post-accretionary basins, modified after Johnson and Woldehaimanot (2003); Nehlig et al. (2002); Stern and Johnson (2010).

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ACCEPTED MANUSCRIPT The distribution of the alkaline/peralkaline granites in the Arabian Shield is from Stoeser (1986); locations of the ring complexes (numbers) are from Roobol and White

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(1985).

Fig. 2: Geological Map of Hadb adh Dayheen ring complex, central Arabian Shield,

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showing major rock units, faults, and sample locations (modified after Moghazi et al.,

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2011; Radain and Kerrich, 1979).

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Fig. 3. Cathodoluminescence images of zircons from Hadb adh Dayheen ring complex samples (HdD-3 and HD-30) analyzed during this study. Location of ion-microprobe

Fig.

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circles, scale is 50 μm.

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U-Pb ages and O isotope spots, and Lu-Hf (MC-ICP-MS) analyses are shown by white

Concordia

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ionprobe

U-Pb

zircon

analyses

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alkaline/peralkaline granites of Hadb adh Dayheen ring complex. (a) Sample HD-3 (hornblende-biotite granite). (b) Sample HD-30 (monzogranite). Dashed ellipses indicate rejected data from the mean age calculation. For analytical data see Table 1.

Fig. 5. (a) Nd isotopic evolution for Hadb adh Dayheen ring complex analyzed in this study. The reference line for the chondritic uniform reservoir (CHUR) and depleted mantle are from Goldstein et al. (1984) and DePaolo (1981). The data field for the juvenile Neoproterozoic ANS is from Ali et al. (2009), Hargrove et al. (2006b), Moussa et al. (2008), and Stoeser and Frost (2006). (b) Hf-in-zircon versus whole-rock Nd diagram showing data of this study and from the literature. The field for Nd(t) - Hf(t) 45

ACCEPTED MANUSCRIPT in juvenile crust is from Ali et al. (2012a), Be‟eri-Shlevin et al. (2010), Katz et al. (2004), and Vervoort and Blichert-Toft (1999). The field for Nd(t) - Hf(t) in arc-

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metavolcanics is from Ali et al. (2013).

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Fig. 6. Epsilon Hf(t) versus zircon age diagrams showing zircon data for Hadb adh

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Dayheen granitic rocks. (a) Sample HD-3 (hornblende-biotite granite). (c) Sample HD-

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30 (monzogranite). The depleted mantle (DM) and new crust (NC) growth curves (solid lines) from a chondritic uniform reservoirs (CHUR) value at the Earth's formation (i.e.,

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4.56 billion years ago) to εHf = 17 at the present time for the depleted mantle (Vervoort and Blichert-Toft, 1999), and εHf = 13.2 for the average mean of 13 modern island arcs

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(Dhuime et al., 2011).

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Fig. 7. Single zircons δ18O data with 206Pb/238U ages from the Neoproterozoic Hadb adh Dayheen granitic rocks. Field of mantle zircon is from Valley et al. (1998), and field of Sinai (Egypt) and S Israel from Be‟eri-Shlevin et al. (2010).

46

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Nd(610Ma)

tDM (Ga)

0.512695 ± 25

5.3

0.77

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Table 2. Sm-Nd isotopic data for granitoid samples from the alkaline Hadb adh Dayheen ring complex of the Arabian Shield in Saudi Arabia.

0.1061

0.512544 ± 10

5.2

0.72

65.52

0.1722

0.512746 ± 10

4.1

0.81*

61.53

0.1418

0.512691 ± 8

5.3

0.76

31.82

120.9

0.1591

0.512758 ± 9

5.3

0.71*

6.530

28.58

0.1382

0.512625 ± 9

--

--

Sm/144Nd

143

5.715

0.1439

21.26

Age (Ma)

Rock type

Sm [µg/g]

Nd [µg/g]

HD-22

~610

Monzogranite

1.360

HD-30

613

3.733

HD-3

625

HD-9B

~610

HD-17

~610

BCR-1

Recent

Monzogranite Hornblende-biotite granite Aegerine-riebeckite granite Aegerine-riebeckite granite USGS Columbia River Basalt rock standard

PT ED

14.43

MA

18.66

SC

Sample

147

Nd/144Nd (m.)

AC

CE

Sm-Nd isotope analyses were carried out at LMU according to the procedures of Hegner et al. 1995. 143Nd/144Nd normalized to 146Nd/144Nd = 0.7219. External precision for 143Nd/144Nd is ~1.1 x 10-5 (2s.d.). Error for 147Sm/144Nd ~0.15% (2s.d.). The La Jolla Nd standard solution yielded 143Nd/144Nd = 0.511847 ± 8 (2s.d., N = 10). JNdi-1 measured with this batch of samples gave 0.512108 ± 9. Sm and Nd concentrations determined by isotope dilution, m. = measured ratio, t = initial ratio. Model-age calculation according to DePaolo (1981); * Sm/Nd ratio too high for calculation of 1-stage Nd model ages; value refers to a 2-stage evolution according to Liew and Hofmann (1988) and using the DePaolo (1981) depleted mantle model.

55

AC

CE

PT ED

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

56

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

AC CE P

TE

D

MA

Graphical abstract

57

ACCEPTED MANUSCRIPT Highlights Hadb adh Dayheen granitic rocks are Neoproterozoic in age (625 – 613 Ma)

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The granitic rocks yielded εHf(t) of +4.5 to +8.4 and εNd(t) of +4.1 to +5.3

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Single zircon grains show a wide range in δ18O values of +3.2‰ to +6.4‰

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Hadb adh Dayheen rocks are consistent with melting of juvenile crustal

RI

PT

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AC CE P

TE

D

MA

NU

SC

protoliths

58