Contribution of slab-derived fluid and sedimentary melt in the incipient arc magmas with development of the paleo-arc in the Oman Ophiolite

    Contribution of slab-derived fluid and sedimentary melt in the incipient arc magmas with development of the paleo-arc in the Oman Ophiolite Yuki Kusano, Susumu Umino, Ryuichi Shinjo, Anzu Ikei, Yoshiko Adachi, Sumio Miyashita, Shoji Arai PII: DOI: Reference:

S0009-2541(16)30659-3 doi:10.1016/j.chemgeo.2016.12.012 CHEMGE 18182

To appear in:

Chemical Geology

Received date: Revised date: Accepted date:

23 February 2016 4 November 2016 6 December 2016

Please cite this article as: Kusano, Yuki, Umino, Susumu, Shinjo, Ryuichi, Ikei, Anzu, Adachi, Yoshiko, Miyashita, Sumio, Arai, Shoji, Contribution of slab-derived fluid and sedimentary melt in the incipient arc magmas with development of the paleo-arc in the Oman Ophiolite, Chemical Geology (2016), doi:10.1016/j.chemgeo.2016.12.012

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Contribution of slab-derived fluid and sedimentary melt in the incipient arc magmas with development of the paleo-arc in the Oman Ophiolite

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Yuki KUSANO1, Susumu UMINO2, Ryuichi SHINJO3, Anzu IKEI4, Yoshiko ADACHI5, Sumio MIYASHITA5, Shoji ARAI2 Yuki KUSANO ([email protected]), Corresponding author Research Institute of Earthquake and Volcano Geology, Geological Survey of Japan,

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AIST, Tsukuba Central 7, 1-1-1 Higashi, Tsukuba, Ibaraki 305-8567, Japan TEL: +81-29-861-0470 2 Susumu UMINO ([email protected]) Division of Earth and Environmental Sciences, Graduate School of Natural Science and Technology, Kanazawa University, Kakuma-machi, Kanazawa 920-1192, Japan 3 Ryuichi SHINJO ([email protected]) Department of Physics and Earth Sciences, University of the Ryukyus, 1 Senbaru, Nishihara, Okinawa 903-0213, Japan 4 Anzu IKEI ([email protected]) School of Earth Sciences, University of Melbourne, Victoria 3010, Australia Yoshiko ADACHI ([email protected]) Center for Transdisciplinary Research, Niigata University, 8050 Ikarashi 2-no-cho, Nishi-ku, Niigata 950-2181, Japan 6 Sumio MIYASHITA ([email protected]) Institute of Science and Technology, Niigata University, 8050 Ikarashi 2-no-cho, Nishi-ku, Niigata 950-2181, Japan 7 Shoji ARAI ([email protected]) Division of Earth and Environmental Sciences, Graduate school of Natural Science

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and Technology, Kanazawa University, Kakuma-machi, Kanazawa 920-1192, Japan

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Abstract (260 words)

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Major-element and trace-element geochemistry of fresh volcanic glass, and the whole-rock Hf and Nd isotopic compositions of volcanic rocks from the Oman Ophiolite were used to understand the extent to which slab-derived fluids and melts were involved in magma generation during incipient arc development. The spreading stage (V1) samples have trace-element characteristics similar to those of

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mid-ocean-ridge basalts, with Hf and Nd isotopic compositions of the Indian mantle domain. The incipient arc stage (V2) volcanic glasses have lower abundances of high field strength elements (HFSEs) and middle and heavy rare earth elements (MREEs and HREEs), and higher abundances of large ion lithophile elements (LILEs) than the V1 glasses. Progressive increases in B, Pb, and LILEs in the V2 stage glasses with age indicate an increasing contribution of slab-derived fluids from earlier arc tholeiite (LV2) to later boninite (UV2). The LV2 was generated by the contribution of amphibolite-derived fluid. The UV2 is subdivided into low-Si UV2 and high-Si UV2 magmas, with the former having lower SiO2, and LILE, and higher Na2O, HFSE and

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REE concentrations than the latter. The low-Si UV2 magma was generated with the involvement of high-temperature slab-derived fluid. In contrast, the high-Si UV2 shows a spoon-shaped trace-element pattern and Hf and Nd isotope chemistry, indicating the involvement of sediment melt in its magma genesis. Decreasing HFSEs and HREEs in the magmas with age indicate progressive depletion of the source mantle during the V2 arc magmatism, suggesting an absence of convection in the mantle wedge that resulted from subduction of a hot, buoyant slab beneath young, hot lithosphere. Keywords: 6

Oman Ophiolite, boninite, Hf–Nd isotope chemistry, slab-derived fluid, incipient arc, volcanic glass.

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1. Introduction Knowledge of the processes operating in the deep mantle and the subducted slab are essential for a full understanding of the material and thermal evolution of the sub-arc mantle and the development of arc systems. Arc magmas are considered to receive recycled materials from the slab but the most part of subducted slab is recovered into the mantle. Therefore, material balance between the arc magmas and subducted slab is still uncertain. Some ophiolites, such as those in Oman and the United Arab Emirates,

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are favorable targets for research in that the entire sequence of volcanic and magmatic products before and after the initiation of subduction. They are well preserved and exposed on land, including huge masses of sub-arc mantle peridotites with underlying metamorphic rocks. The metamorphic rocks beneath ophiolite nappes are called metamorphic sole (e.g. Boudier et al., 1988).

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The metamorphic rock-derived fluids are considered to be related to the arc magmatism (Ishikawa et al., 2005). The V2 extrusive rocks which unconformably overlie the V1 extrusive rocks have strong arc signatures and include boninite (Umino et al., 1990;

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Ishikawa et al., 2002; Gilgen et al., 2014; Kusano et al., 2014). Ishikawa et al. (2005) demonstrated that the trace-element characteristics of the V2 arc tholeiite and boninite could be explained by partial melting of the depleted source mantle with the introduction of hydrous fluids in equilibrium with amphibolite in the metamorphic sole. On the other hand, granitic dikes are found in the mantle section which have sedimentary feature for their origin (Haase et al., 2015; Rollinson et al., 2015). Since these dikes may be formed at high-temperature condition (Rollinson et al., 2015) similar to the metamorphic rocks, the metamorphic rocks are possibly the source of the granitic dikes, in other words, the remnants of the subducted slab. In the northern Oman Ophiolite, the pathways of these slab-derived fluids and melts are preserved as 5–10 km wide zones of highly refractory harzburgite which run subparallel to each other at a space of ca. 5 km and extend from below the Moho all the way down to the base of the mantle peridotite massifs just above the metamorphic sole (Kanke and Takazawa, 2014). These zones are interpreted to be the channels of V2 arc magmas that passed through and interacted with the peridotites in the mantle wedge. Therefore, the Oman Ophiolite can provide invaluable information on the magmatic processes in the mantle wedge involving the recycle system of subducted slab in the infant Oman arc, which is inaccessible in the present arc system. However, contribution of sediment melt to the V2 magmas is not well constrained by the previous studies.

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In addition to the composition of the source mantle, the composition of arc magmas is largely varied by the composition of slab-derived fluids, which are best evaluated by analyzing mobile elements such as large ion lithophile element (LILE) liberated through the dehydration and partial melting of the subducted slab. To better understand the temporal and spatial evolution of the incipient arc and the mantle wedge, understanding of the genetic relationships between the arc magmas and the subducted slab is necessary. However, these LILE are susceptible to low-temperature alteration and weathering,

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which prevail in the extrusive rocks of the Oman Ophiolite.

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The V2 arc magmatism ended after a short period (age ranges of 96–94 Ma; Hacker et al., 1996) with the extrusion of high-Ca and low-Si boninite. Based on the stratigraphy and geochemistry, Stern et al. (2012) applied the model of Izu–Bonin–Mariana (IBM) proto-arc development to ophiolites and proposed that ophiolites originate in the region of forearc extension that followed subduction initiation and resulted in the formation of a new arc–trench system. Although the Oman Ophiolite bears some resemblance in terms of magmatic history to the IBM proto-arc (Ishizuka et al., 2011), the incipient arc

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magmatism in the Oman Ophiolite significantly differs in that: 1) it started with the production of arc tholeiite and ended with boninite, and was followed by intraoceanic volcanism that produced alkali basalt (Alabaster et al., 1982; Umino et al., 1990; Umino, 2012; Kusano et al., 2014); 2) it was short lived and lasted for only < 3 m.y. (96–94 Ma; Hacker et al., 1996); and 3) the subduction initiated between young and buoyant oceanic plates including the ridge axis (Nicolas, 1989; Umino et al., 1990; Ishikawa et al., 2002; Kusano et al., 2014). It is currently uncertain whether the lowermost V1 stage rocks were formed at a mid-ocean ridge or represent oceanic lithosphere generated at a backarc or a forearc spreading axis, as the subtle arc-like signature of the V1 rocks is obscured by pervasive alteration and weathering of the volcanic products (e.g. Pearce et al., 1984; Nicolas, 1989; Moores et al., 2000; Stern et al., 2012; MacLeod et al., 2013). The incipient arc V2 magmatism was emplaced in this such an oceanic basement as plutonic bodies and on as the extrusive sequences (Umino et al., 1990; Ishikawa et al., 2002; Nagaishi, 2008; Yamazaki, 2013; Kusano et al., 2014). We have recovered fresh volcanic glasses from the V1 and V2 volcanic sequences extending over ~80 km in the northern Oman Ophiolite. By analyzing these fresh volcanic glasses, we were able to obtain data on the pristine major- and trace-element compositions, including LILEs, of the V1 and V2 volcanic rocks. On the basis of these

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geochemical data and the Hf and Nd isotopic compositions of the bulk V1 and V2 rocks, we present a geochemical model of the V2 magma genesis and discuss the evolutionary processes of the intraoceanic subduction zone recorded in the Oman Ophiolite.

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2. Geological background of the volcanic glass samples The lowermost V1 unit was extruded and emplaced on and near the spreading axis at 98–96 Ma (Einaudi et al., 2003; Kusano et al., 2012; Rioux et al., 2012) and is comagmatic with the underlying sheeted dike complex and gabbros (e.g. Alabaster et al.,

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1982; Lippard et al., 1986). The V1 sequence consists mainly of lava flows with minor pyroclastic and hyaloclastic breccias, including pillow flows, pahoehoe and lobate flows, and sheet flows, with a few massive lavas (Kusano et al., 2012). The lower V1 (LV1) sequence is composed of evolved and homogeneous on-axis lava flows, interbedded with and subsequently overlain by the upper V1 (UV1) extrusive rocks emplaced off-ridge. The UV1 sequence was extruded in situ from off-ridge vents and is more primitive than the LV1. The V1 rocks are usually either aphyric or carry < 10% phenocrysts, containing plagioclase (Pl), clinopyroxene (Cpx) and lesser olivine (Ol) (Kusano et al., 2012). Most of the Pl and Ol crystals are replaced by secondary minerals

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such as albite, clay minerals and some zeolites filling vesicles. Fresh V1 volcanic glass was discovered within the LV1 at two localities along wadis Jizi and Shafan. Samples 12Jiz1, OM17-67A, and OM17-67B are from glassy chilled margins on pahoehoe lobes and OM11-22A is hyaloclastite breccia intercalated with pillows. These volcanic glasses are generally fresh except thin hydration halos surrounding glass fragments and along cracks (Fig. 2a). The V2 sequence is divided into the lower V2 (LV2) and upper V2 (UV2) subsequences (Kusano et al., 2014) in the Wadi Bidi. The LV2 comprises arc tholeiite interbedded with thin boninite pyroclastic fall deposits at stratigraphically higher levels. The LV2 consists mainly of lower pahoehoe and upper sheet flows, but is locally intercalated with pyroclastic deposits at wadis Suq and Asher (Fig. 1). Fresh LV2 volcanic glass fragments were obtained at four localities over a distance of 80 km along the inferred paleoarc. Sample OM16-113 consists of glassy clasts in hyaloclastite intercalated with pahoehoe lobes, and samples OM16-73B, OM17-59, 12SH12 and 12SH13 are glassy clasts or scoriae of subaqueous pyroclastic fall deposits. These volcanic glasses contain Pl and Cpx microlites (Fig. 2b). More crystalline parts of the samples contain 10% phenocrysts of Pl, Ol, orthopyroxene (Opx), and Cpx (Kusano et al., 2014).

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The UV2 sequence consists of boninite sheet flows and pyroclastic fall deposits. The UV2 consists of Ol–Opx–Cpx phyric boninite with a glassy groundmass (Ishikawa et al., 2002; Kusano et al., 2014). Most of the Ol phenocrysts have been replaced by clay minerals, but some Ol relicts with Fo87–91 have remained unaltered (Fig. 2c). The UV2 volcanic glass samples were collected from five localities spread over 50 km along the paleoarc. Samples OM14-2C, OM16-46C, OM17-62, OM17-63, OM17-80A and OM17-28C are chilled margin glasses of boninite pillow lobes, OM16-84 is scoria in a pyroclastic fall deposit, and OM17-97B is a glassy fragment from an epiclastic tuff

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breccia.

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3. Geochemistry The major elements of 21 volcanic glass samples were analyzed using a JEOL JXA-8800 electron probe microanalyzer (EPMA) at Kanazawa University, Japan, with an accelerating voltage of 15 kV and a beam current of 12 nA. Counting time was 20 seconds for the peak and 10 seconds for the background using a broad beam of 20 μm in diameter, following the procedures of Noguchi et al. (2004) to correct the time-depending Na-loss. Corrections were made using the ZAF method. Analysis are

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monitored by laboratory-standard clinopyroxene and BCR-2G compositions. Relative analytical errors (1σ) of BCR-2G standard were generally better than ~3%. Hydration haloes and microcrystals (Fig. 2) were avoided for the geochemical analysis.

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The trace-element compositions (Li, B, Sc, V, Cr, Co, Ni, Rb, Sr, Y, Zr, Nb, Cs, Ba, REEs, Hf, Ta, Pb, Th, and U) of the volcanic glass were determined by laser ablation (193 nm ArF excimer; MicroLas GeoLas Q-plus)–inductively coupled plasma–mass spectrometry (LA–ICP–MS) (Agilent 7500S) at Kanazawa University (Ishida et al., 2004; Morishita et al., 2005a, b). Each analysis was performed by single-spot ablation (60 µm diameter) at a 6-Hz repetition rate with an energy density of 8 J/cm2 per pulse. Signal integration times were 50 seconds for the gas blank interval and 60 seconds for the ablation interval. BCR-2G was used as the primary calibration standard. The reference values of BCR-2G were selected from the GeoReM database (see Jochum and Nohl, 2008). NIST 612 was also measured for quality control. Data reduction was performed using 42Ca as an internal standard element, based on CaO contents obtained by EPMA analysis, following a protocol essentially identical to that outlined by Longerich et al. (1996). The major- and trace-element compositions of the volcanic glass samples are listed in Table 1. Repeated analyses of NIST 612 demonstrated that the reproducibility was better than 3% (Table 1).

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Whole-rock Nd and Hf isotopic compositions were analyzed using a multi-collector ICP–MS (Neptune Plus; Thermo Fischer Scientific) at the University of Ryukyus, Japan. The major- and trace-element contents of 71 samples were reported by Kusano et al. (2012, 2014). Sample preparation and chemical separation followed the procedures of Shinjo et al. (2010, 2015). The mass fractionation factor was determined from measured 179 Hf/177Hf, and other isotope ratios were normalized to 179Hf/177Hf = 0.7325 using an exponential law. Repeated measurements of the standards JMC-475 for Hf, JNdi-1 for

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Nd, and La Jolla provided 176Hf/177Hf = 0.282153±0.000020 (n = 20), 143Nd/144Nd = 0.512109±0.000023 (n = 17), and 0.511825±0.000030 (n = 50). The measured 176 Hf/177Hf values for the samples were normalized to a JMC-475 value of 176Hf/177Hf = 0.28216 and 143Nd/144Nd values were normalized to a La Jolla value of 143Nd/144Nd = 0.511858. εHf(t) and εNd(t) values were calculated assuming an age of 95 Ma and relative to chondritic uniform reservoir values of 0.282785 for 176Hf/177Hf and 0.512630 for 143Nd/144Nd (Bouvier et al., 2008). The whole-rock Nd and Hf isotopic compositions are listed in Table 2.

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3.1. Major and trace elements The presence of some hydration halos surrounding the volcanic glass fragments along cracks suggests that some fluid-mobile element compositions could be modified by seawater alteration or palagonitization (Staudigel and Hart, 1983). Yb-K2O diagram shows a positive correlation for both the V1 and V2 samples (Fig. 3). Although K is susceptible to low-temperature alteration, Yb is immobile in such alterations. Because both elements are expected to be incompatible during the fractional crystallization of the magmas, the tight positive correlation of Yb and K2O for the V2 samples indicates that the K variation is magmatic composition. However the V1 samples show both high and low K2O contents at a given Yb. The low-K2O V1 samples show high Rb contents, particularly the 12Jiz1 glass has significantly high Rb, Sr, and Ba contents. Thus, LILE contents of these two V1 samples are considered to have been affected by alteration and do not preserve the primary igneous features. On the other hand, the V2 samples show clusters or positive correlations in Rb-, Sr- and Ba-Yb diagram (Fig. 3), and are hence considered to have retained their primary compositions. Thus, we used these geochemical compositions of volcanic glasses for the following discussion of the V2 arc magmatism. The only exception among the LV2 samples is 12SH13, which has higher K2O at a given TiO2.

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The V1 glass samples contain 55.1–62.0 wt% SiO2, 2.0–4.0 wt% MgO, 1.3–3.7 wt% Na2O, and 0.5–1.2 wt% K2O. The V1 glass samples plot in the evolved compositional range of the bulk V1 and sheeted dikes and show fractionation trends consistent with the bulk rocks on SiO2–MgO and TiO2–MgO diagrams (Fig. 4). The bulk Na2O and CaO contents are higher and lower than the volcanic glasses, which is consistent with the hydrothermal alteration under greenschist facies that modified the bulk V1 compositions. The V1 glasses show flat trace-element patterns normalized against primitive mantle (Sun and McDonough, 1989), which overlap the evolved V1 bulk compositions

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(Kusano et al., 2012) (Fig. 5a). The overall multi-element patterns of the V1 glasses plot within the Indian MORB field but do not show positive Ba or negative Th anomalies. The negative anomalies in Ti can be explained by fractionation of Fe–Ti oxide minerals in andesitic compositions, as indicated by a positive correlation on the TiO2–MgO diagram (Fig. 4). Other features of the V1 glasses that differ from the Indian MORBs are positive anomalies in K and relatively low LILE and high middle rare earth element (MREE) and heavy rare earth element (HREE) content.

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The LV2 glass samples contain 55.2–62.7 wt% SiO2, 2.4–7.7 wt% MgO, 1.3–2.3 wt %

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Na2O, and 0.2–0.5 wt% K2O. The major-element abundances of these samples are similar to those of the bulk rocks. The LV2 samples show higher SiO2 and CaO, and lower TiO2, MnO, Na2O, and K2O contents at a given MgO value than the V1 samples (Fig. 4). They show subparallel multi-element patterns normalized against normal MORB (N-MORB; Sun and McDonough, 1989) and have positive K and Pb, and negative Ti, Th, and Nb–Ta anomalies (Fig. 5b). Three LV2 glasses show negative Li anomalies, and two have positive Li anomalies. Compared with the V1 glasses, the LV2 glasses have much lower concentrations of Sr, La, and Th, and higher concentrations of B, Rb, Cs, and Pb (Fig. 3 and 5). The UV2 samples contain 54.2–56.1 wt% SiO2, 6.7–8.7 wt% MgO, 0.8–1.1 wt% Na2O, and 0.1–0.2 wt% K2O. Of all the extrusive rocks, the UV2 samples show the lowest TiO2, Na2O, Sr, Ba, and La contents, and the highest MgO, CaO, B, Cs, and Pb contents at a given SiO2 value (Figs 3, 4 and A1). The UV2 glasses can be subdivided into low-Si and high-Si groups on the basis of MgO and N-MORB-normalized REE patterns. The high-Si UV2 group has higher SiO2, K2O, Rb, Cs, and Pb, and lower TiO2, high field strength element (HFSE), MREE, and HREE contents than the low-Si UV2 group (Fig. 4). In addition to the anomalies in U, K, Pb, Sr, and Li, the high-Si UV2 samples show overall spoon-shaped multi-element patterns. In contrast, the low-Si UV2 samples

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display relatively flat patterns for HFSEs and REEs, with slightly higher LREE (light rare earth element)/HREE ratios and lower LILE than the high-Si UV2 (Fig. 5c).

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3.2. Nd and Hf isotopes Although the bulk-rock samples widely vary generally compatible with fractional crystallization (Fig. 4), Hf isotopic compositions vary little with variable SiO2 or Zr contents (Fig. A2). The V1 samples have the largest range in Hf isotopic ratio (initial εHf(t) = 13.04–16.82; t = 95 Ma). The Hf isotopic variations of the LV2 (13.36–16.03)

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and UV2 (13.21–14.20) samples are within the range of the V1 except one UV2 sample which have lower Hf isotopic ratios. Also, Nd isotopic compositional range of the V1 is wider (initial εNd(t) = 7.24–9.26) than that of the LV2 (8.05–8.97). The UV2 samples show lower εNd(t) than the V1 and LV2, clustering around 0.88–2.15 and 4.57–5.69. The samples with lower Hf isotopic ratios have both low and high Nd isotopic composition. One LV2 sample with significantly low εNd(t) of 3.38 is considered to be altered based on the LREE compositions (10SH21 in Kusano et al., 2014). The compositional difference with the UV2 is shown in Fig. A3.

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The Nd and Hf isotopic compositions of the V1 bulk-rock samples plot in the Indian mantle field of Pearce et al. (1999) and lie on the mantle array (Vervoort et al., 1999; Fig. 6a). The initial εNd(t) values are in the lower end of previously reported ranges of plutonic and extrusive rocks from the Oman Ophiolite (Benoit et al., 1996; Godard et al., 2006; Tsuchiya et al., 2013). The LV2 bulk-rock samples have a similar and narrower range of initial εNd(t) and εHf(t) values to the V1 samples. In contrast, the UV2 samples show a much lower and broader range of εNd values than the other rocks, whereas the initial εHf values are identical to those of LV2. The low εNd values of UV2 lie on the mixing line between the V1 samples and the Indian pelagic sediments and the granitic intrusions into the mantle peridotite rather than the continental materials from the Indian Ocean (Fig. 6b). The εNd(t) and εHf(t) variations of the UV2 samples can be reproduced by 0.005–0.03% sediment-derived felsic melt mixed with the V1 residue (Fig. 6b). 4. Discussion 4.1. Geochemical evolution of the source mantle and subduction components 4.1.1. Implications for the Indian MORB mantle as the source of the V1 magma The fresh V1 glasses have relatively higher CaO and lower Na2O concentrations than the bulk V1 (Fig. 4), indicating that the latter have been significantly modified by

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low-temperature hydrothermal alteration, consistent with the widespread occurrence of the greenschist-facies mineral assemblage of albite, chlorite, actinolite, and epidote (e.g. Lippard et al., 1986). Also, comparison of the altered rinds and fresh cores of glasses clearly demonstrates LILE-enrichment in the former relative to the latter. Therefore, besides positive anomalies in LILEs, most likely due to alteration, the overall geochemical characteristics of the V1 glasses are identical to those of Indian MORB (Fig 5a). This finding is also supported by the relationships between εHf and εNd in the V1 lavas of the present and previous study (Fig. 6a; A‟Shaikh et al., 2005; Godard et al.,

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2006). The wider range of εNd reported in previous studies (Benoit et al., 1996; Godard et al., 2006; Tsuchiya et al., 2013) and εHf along mantle array suggests heterogeneity of the sub-Tethyan mantle along the paleo-spreading axis.

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4.1.2. Progressive depletion of the source mantle from V1 to V2 magmatism In the IBM proto-arc, the onset of subduction of the Pacific Plate at 52–50 Ma caused upwelling of depleted MORB-mantle-like asthenosphere and rifting of the northeastern margin of the Philippine Sea Plate, leading to forearc spreading and the formation of MORB-like basalt (Reagan et al., 2010; Ishizuka et al., 2011, 2014; Umino et al., 2015).

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The MORB activity was followed by the extrusion of severely depleted high-silica boninite at 48–46 Ma, followed in turn by less-depleted low-silica boninite at 45 Ma, and finally by fertile calc-alkaline and arc tholeiitic rocks at 45–40 Ma (Kanayama et al., 2012, 2014; Umino et al., 2015). These temporal geochemical variations in proto-arc magma indicate the beginning and establishment of convection within the mantle wedge that replaced the ultra-depleted source of the boninitic magmas with the fertile source of the 45–40 Ma arc magmas. In contrast, Kusano et al. (2014) described progressive decreases in incompatible elements (e.g., TiO2, Na2O, HFSE, MREE, and HREE) from the V1 through LV2 to UV2 (Figs 4 and 5) and ascribed this pattern to stepwise depletion of the source mantle as the MORB-like V1 residue was remelted by the introduction of slab-derived fluids after the onset of subduction. The peridotite melting model of Stolz et al. (1996) describes how the Nb/Ta ratio of the residue decreases with increasing partial melting. The range of Nb/Ta ratios for LV2 (13–16; Fig. 7) largely overlaps with but extends lower than that of the V1, which could be explained by remelting of the V1 residue with retained V1 melt left in the subaxial mantle. The even lower Nb/Ta ratios of the UV2 would have been formed by either a higher degree of partial melting of the V1 residue or remelting of the LV2 residue. Furthermore, the similar εHf values of these extrusive rocks confirm gradual depletion of the same source mantle rather than the introduction of depleted mantle elsewhere. Thus, the stepwise

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melt extraction and depletion of the same source mantle throughout V2 magmatism strongly suggests the absence of convection within the mantle wedge in the Oman subduction zone (Kusano et al., 2014).

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4.1.3. Secular variations in subduction components in the V2 magmas Positive spike of Pb in multi-element diagram (Fig. 5) is a common characteristic of arc magmas ascribed to the addition of Pb-enriched hydrous slab-derived fluid (Miller et al., 1994; Brenan et al., 1995; Kogiso et al., 1997) or partial melting of basalt and/or

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sediments of the slab (Tatsumi, 2000; Kelemen et al., 2003). However, those fluid-mobile incompatible elements are either increased by fractional crystallization of magmas (Fig. A1) or reduced with the increase in depletion of the source mantle. To minimize the effects of fractional crystallization and source depletion, we evaluated the contributions of slab-derived fluid components based on [highly fluid mobile]/[immobile element] ratios (Ishizuka et al., 2003). Ba/Zr and Th/Zr ratios are indicators of fluid components from altered oceanic crust and sediment melt, respectively (Elliot et al., 1997; Pearce et al., 2005), and both increase from V1 through LV2, to low-Si and high-Si UV2 glasses (Fig. 7). This finding is consistent with the

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increasing values of the B/La, Th/La and Pb/Ce ratios, indicating progressively greater contributions of slab-derived fluids and/or melts over time. The LV2 samples have a similar Th/La ratio to oceanic basalts and the characteristic Pb/Ce ratios of arc magma (Miller et al., 1994; Plank, 2005), suggesting the contribution of hydrous fluids during the generation of the LV2 magma. The similar εNd and εHf values for LV2 and V1 support the lack of involvement of sediment- and igneous crust-derived melts in the generation of the LV2 magma (Fig. 6). The higher Th/La ratios of UV2 resulted from the inclusion of sediment (Plank, 2005). The variation of the bulk UV2 initial εNd versus εHf cannot be explained by the addition of hydrothermal fluid but requires the involvement of high proportions of pelagic sediment melt or sediment-derived granitic melt (Fig. 6b; Fig. A3). Off-ridge deep hydrothermal circulation has been proposed to cause hydration of the off-ridge mantle peridotite, resulted in the formation of pyroxenite veins (Nicolle et al., 2016), which could be a potential source for the V2 arc magmas. However, these pyroxenites were crystallized from melts with LREE-depleted chondrite-normalized REE patterns, which are discriminated from the flat to high LREE/MREE-patterns of the LV2 and UV2. The εNd values of 6.2–7.9 for the pyroxenites are significantly higher than the UV2 samples. The low-temperature hydrothermal fluid circulating through the uppermost sediment on the oceanic basement cannot carry sufficient Nd to cause the observed large reduction in εNd for the UV2

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compared to other V1 and LV2. Thus, we conclude that the sedimentary component must have come from the subducted slab which reached 770–900 °C at 1.1–1.3 GPa (Cox, 2000; Cowan et al., 2014).

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The slightly elevated Th/La and Pb/Ce ratios in the low-Si UV2 may have resulted from a minor contribution of sediment-derived felsic melt or high-temperature hydrous fluid, as Th is more mobile in high-temperature fluids than is Ba (Elliot et al., 1997; Pearce et al., 2005). The similar Pb/K ratios of the low-Si and high-Si UV2 glasses indicate the

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involvement of a high-temperature hydrous fluid in the generation of the high-Si UV2 magma. However, the higher values and larger variations of B/La and Nb/Ta ratios of the high-Si UV2 glasses suggest greater and variable contributions of fluid and sediment-derived felsic melt, and mantle heterogeneity.

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The spoon-shaped N-MORB-normalized patterns for the high-Si UV2 glasses are common characteristics of bulk boninite from the Oman Ophiolite (Ishikawa et al., 2002; Kusano et al., 2014). Despite the depleted HFSEs and HREEs, these rocks have distinct positive anomalies in Li, Pb, and K. In contrast, the low-Si UV2 glasses are

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distinguished by their less-depleted HFSEs and HREEs, and lower LILEs, K, Pb, and Li compared with the high-Si UV2 glasses, and are considered to be transitional between the LV2 and high-Si UV2.

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4.2. Geochemical modeling of the composition of the primary magma To understand the geochemical evolution of the mantle wedge during V2 arc magmatism, we conducted geochemical modeling to reproduce the V2 magma compositions by estimating the degree of depletion of the source mantle and the varying contributions of hydrous fluids and sediment melts. As discussed above, the LV2 magmas can be modeled by batch modal melting of the LV2 sub-arc mantle, which is the residual mantle after extraction of the V1 magma that had been metasomatized by hydrous fluids (Table 3). The subsequent UV2 magmas are modeled by batch modal melting of the UV2 sub-arc mantle, which is assumed to be the residual mantle after extraction of either the V1 or LV2 magma, plus hydrous fluid and/or sediment melt as shown by the εNd(t)-εHf(t) relationships. We used the hydrous fluid compositions reported by Ishikawa et al. (2005), which were estimated from amphibolite (type I fluid) and Cpx-bearing amphibolite (type II fluid) in the metamorphic sole based on partitioning of selected trace elements between fluid and

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amphibolites. The type I fluid contains high abundances of B, K, Sr, and Li, while the type II fluid is enriched in Nb, LREE, and Li. As the agent of sediment melt for the UV2, we used the metachert associated with the amphibolite at wadi Tayin (Ishikawa et al. 2005). The trace element abundance of pelagic sediments shown in Fig. 6b (Plank and Langmuir, 1998) is roughly intermediate between metachert and amphibolite-derived fluid and suggests that the trace element contents can be reproduced by a mixture of metachert and amphibolite-derived fluids (Fig. 8a). The model LV2 and UV2 magma compositions were used to fit the primary magma

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compositions estimated from the most primitive glass and bulk compositions, to determine the degree of melting and the proportion of amphibolite-derived hydrous fluid and sediment melt.

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The most plausible candidates for the V1 residues are type I lherzolite with Fo 90.4 olivine, and harzburgite with Fo91.5 olivine in the mantle section described by Takazawa et al. (2003). Type I lherzolite is moderately depleted in REE and is interpreted as the residue of V1 MORB beneath the Oman spreading ridge (Takazawa et al., 2003). In contrast, the harzburgite is more depleted and experienced a higher degree of melting,

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and thus is the most likely candidate for the residue of the depleted V1 lava. The type I lherzolite in the base of the mantle section in the Oman Ophiolite was refertilized by LREE-depleted melt and changed to type II lherzolite (Takazawa et al., 2003). As suggested by the Hf–Nd isotope systematics, the slab component involved in the generation of LV2 magma did not include sediment melt, which is in agreement with the LREE-depleted melt responsible for the refertilization of type II lherzolite. Therefore, we assume the type II lherzolite with Fo90.7 to be the residue of the LV2 magma. We also assume the highly refractory harzburgite (Kanke and Takazawa, 2014) to be the residue of the UV2 magma because the Fo91.5 is in equilibrium with that of the UV2 olivine phenocryst with Fo91. The modal compositions of the harzburgite, type II lherzolite, and depleted harzburgite are thus regarded as those of the residual mantle peridotites of the V1, depleted V1, LV2 and UV2 magmas, respectively. However, the fluid-mobile element abundances of these peridotites have been raised by secondary alteration (Takazawa et al., 2003), and hence cannot be used as the source composition in the geochemical modeling. Therefore, we estimated the trace-element compositions of the residual mantle on the basis of the V1 bulk, and LV2 and UV2 glass compositions by using peridotite/melt partition coefficients (Kelemen et al., 2003). Only Fo contents of the prescribed residues assumed above were used to put constraints on the Mg#s of the estimated primary magmas of the V1, depleted V1, LV2 and UV2.

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Because all V1 glasses have evolved compositions (MgO < 4.0 wt%), primitive V1 flows and sheeted dikes with MgO > 8.0 wt%, Mg# > 60, and without positive anomalies of LILEs in N-MORB-normalized patterns were used to estimate the trace element abundances of the primary magmas of less-depleted V1 and depleted V1. The MgO-CaO and MgO-Al2O3 variations of MORB from the East Pacific Rise are controlled by Ol-Pl-Cpx fractionation below 8.0 wt% MgO. Between 8 and 9 wt% MgO, fractionation of Ol and Pl buffers CaO and Al2O3 contents at the maxima, beyond which

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CaO and Al2O3 are suggested to decrease with increasing MgO because of fractionation of solely Ol as shown by the fractional crystallization model by using pMELTS (Asimow and Longhi, 2004). Although the V1 data is highly disturbed by the hydrothermal alteration on the CaO-MgO plot (Fig. A4), they also tend to show the maximum CaO at around 8.0 wt% MgO. The fractionation model of the V1 magma based on MELTS suggests that Pl joins the fractionating phase at about 9 wt% or less, above which Ol is the only fractionating phase (MacLeod et al., 2013). To estimate the primary LV2 magma, the most magnesian LV2 glass OM16-73 (7.7 wt% MgO and 11.7 wt% CaO) was used. Although there are some uncertainties of the amount of phases

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fractionated, the difference is subtle between the trace element compositions of the original LV2 sample and the estimated primary LV2 magma because the amount of Ol added (6.5 wt%) was minimal as described below. Also, MgO content of the UV2 volcanic glass is relatively low within 6–12 wt% MgO of the UV2 bulk rocks and volcanic glass (Fig. 4), but multi-element patterns are similar to each other (Fig.5). Therefore, we directly compared with the estimated UV2 magma and the UV2 volcanic glass compositions. Assuming Ol is the solely fractionated phase, the primary V1 and LV2 magma compositions were restored by adding equilibrium Ol to the most primitive compositions selected above. The steps in the process are as follows: 1) the Fe3+/(Fe2+ + Fe3+) ratios of the V1 and LV2 magmas were assumed to be 0.16 (Cottrell and Kelley, 2011) and 0.2 (Kelley and Cottrell, 2009), respectively; 2) the equilibrium Ol composition was calculated based on the Fe–Mg exchange partition coefficient of 0.3 between Ol and melt, and assuming Ol stoichiometry; and 3) 1 wt% Ol was added to the melt at each step until Fo content of the equilibrium Ol became equal to that of the assumed residue: Fo90.4 of type I lherzolite for the less depleted V1 (07Fz22), Fo91.5 of the residual harzburgite for the depleted V1 (96om112), and Fo90.7 of type II lherzolite for the LV2 (OM16-73). The trace-element compositions of the primary magmas were

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calculated based on the amount of Ol added and the Ol–melt partition coefficients of Kelemen et al. (2003) (Table 4). The estimated trace-element compositions of the residual mantle are consistent with that of Takazawa et al. (2003) without LILE abundance.

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The LV2 can be successfully modeled by 9% partial melting of the less-depleted V1 residue as estimated above combined with 1.7% mixed fluids of 88% type I and 12% type II (Fig. 8b). Alternatively, 2% partial melting of the depleted V1 lherzolite residue

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with 0.39% mixed fluid can also reproduce the LV2 magma. Both models fail to reproduce the highly positive Pb anomaly displayed by the LV2 magma. The positive Pb anomaly in the LV2 magma requires the introduction of a high-temperature fluid with a high concentration of Pb, leached from the slab (Ishikawa et al., 2005).

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The trace-element compositions of the low-Si UV2 glasses are characterized by moderate depletion in REE and enrichment in LILE, Li, and Pb (Fig. 7). This transitional magma between LV2 and high-Si UV2 was probably generated by remelting of either the V1 residue or the LV2 residue (Fig. 8c). The low-Si UV2 primary magma

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composition can be explained by 5% partial melting of the depleted V1 residue and 0.47% mixed fluid of type I and II. Alternatively, 6% partial melting of the LV2 residue and 0.48% mixed fluid could have generated the low-Si UV2 primary magma.

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Although the spoon-shaped N-MORB-normalized patterns of high-Si UV2 (Fig. 5c) can be reproduced by partial melting of a depleted source mantle with the introduction of fluids in equilibrium with amphibolite in the metamorphic sole (Ishikawa et al., 2005), the εHf–εNd relationships require an additional contribution of 0.005–0.03% sedimentary melt for the high-Si UV2 magma. Referring the contribution of sedimentary melt to isotopic estimation, 7% partial melt of the depleted V1 residue and 0.03% metachert and 0.85% amphibolite fluids can reproduce the trace element abundance of the high-Si UV2 (Fig. 8d). Also, 7% partial melt of the LV2 residue combined with 0.005–0.03 wt% metachert and 0.87–0.88 wt% amphibolite-derived fluid can well reproduce the trace-element composition of the high-Si UV2 primary magma. A higher abundance of Pb could be explained by the addition of Pb-rich fluid liberated from the deeper and hotter part of the slab (Ishikawa et al., 2005), because the slab surface temperature in the UV2 stage was sufficiently higher than in the LV2 stage to melt slab sediments. Thus, a contribution of felsic melt derived from pelagic sediment and amphibolite fluids to the LV2 residue could have generated the UV2

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magma.

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4.3. Evolution of the hot and shallow subduction zone The Nd and Hf isotopic relationships of the present study exclude the involvement of continental materials in the generation of UV2 magma, suggesting that the subduction zone formed in an intraoceanic environment isolated from continents. The thermal and chemical conditions of this subduction zone are constrained by the following observations: 1) the preservation of diapiric structures in the upper mantle (Nicolas,

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1989); 2) the small timelag between the V1 and V2 magmatism (< 2 m.y. difference; Hacker et al., 1996; Rioux et al., 2012, 2013); 3) conditions of 1320 °C and 0.5 GPa as magma genetic conditions for a primary UV2 low-silica boninite (Umino et al., 2014); and 4) peak temperature–pressure conditions of 770–900 °C and 1.1–1.3 GPa for the amphibolite in the metamorphic sole (Cox, 2000; Ishikawa et al., 2005; Cowan et al., 2014), which was in equilibrium with the fluids involved in formation of the V2 magmas, as shown above. This indicates that the subducted slab was unusually hot and was hence buoyant (Gutcher et al., 2000). These lines of evidence are most readily explained by intraoceanic thrusting initiated near the ridge axis that developed into a

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shallow and hot subduction zone (Nicolas, 1989; Ishikawa et al., 2005; Kusano et al., 2014). In addition, a stress field of E–W compression prevailed during the V2 magmatism in the northern Oman Ophiolite, as demonstrated by conjugate sets of NW–SE-striking sinistral shear zones and NE–SW-striking dextral shear zones, which show mutual cross-cutting relationships with E–W-striking V2 gabbronorite dikes (Umino et al., 1990). During the process of intraoceanic thrusting, thin and hot subaxial asthenosphere was sandwiched as a hot mantle wedge between the overlying thin lithosphere and the underlying buoyant slab. The hot slab was metamorphosed to Cpx amphibolite under conditions of > 770 °C and ~1 GPa, and released fluids (Cox, 2000; Ishikawa et al., 2005; Cowan et al., 2014) that ascended into the overlying mantle wedge and caused remelting of lherzolitic residue of the less-depleted V1 magma and generated the LV2 magma (Fig. 9a). As subduction progressed, the slab temperature increased and higher-temperature hydrous fluid was released that caused remelting of the lherzolitic residue of LV2 magma and the harzburgite residue of depleted V1 magma, resulted in the formation of the low-Si UV2 magma. The increasing supply of fluid and felsic melt derived from Cpx amphibolite and metachert (Fig. 9b) promoted further remelting of the harzburgitic depleted V1 residue, producing a broad distribution (> 100 km along

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the arc) of the UV2 volcanic rocks. There are subtle spatial changes in the chemical composition and degree of contribution of metachert, which probably reflect the uneven distribution of sedimentary covers on the subducted slab. Granitic intrusive rocks reported from the mantle harzburgite in the Haylayn block (Haase et al., 2015), ~80 km to the southeast of the present study area, could represent such sediment-derived granitic melts frozen during ascent in the mantle wedge. Numerical modeling of a subhorizontal subduction zone suggested that melting of the

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slab and mantle wedge occurred only in the early stage and ceased as the mantle wedge cooled because of the absence of convection (Gutcher et al., 2000). The shallow and buoyant subduction zone hindered the development of convection within the mantle wedge, and resulted in the progressive remelting and depletion of the source mantle wedge through the V2 arc magmatism (Figs 5, 7 and 8) (Nicolas, 1989; Kusano et al., 2014). As melt was progressively extracted, the residual mantle became more refractory and ultimately ceased melting within 3 m.y. Intraplate V3 volcanism at ca. 90 Ma (Alabaster et al., 1982; Umino et al., 1990) lacks any trace of an arc signature (cf. Nb–Ta depletion in the multi-element diagram; Fig 11c in Umino, 2012), indicating the

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termination of the V2 arc magmatism with the extrusion of the UV2.

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5. Conclusions Major- and trace-element compositions of fresh glasses, and the bulk Nd and Hf isotopic compositions of the V1, LV2 (mainly arc tholeiite), and UV2 (boninite) flows and pyroclastic rocks are reported. The V1 samples have Indian MORB-like major-element and trace-element values, and Nd and Hf isotope compositions. The V2 volcanic glasses show arc signatures with lower HFSE, MREE, and HREE, and higher LILE abundances than the V1 glasses. Higher concentrations of B, Pb, and LILE can be ascribed to an increasing contribution of slab-derived fluid in the LV2 to UV2 magmas. LV2 magma was formed by remelting of the V1 residue with the addition of 1.7% amphibolite-derived fluid. UV2 can be divided into low-Si UV2 and high-Si UV2 based on major- and trace-element characteristics. The low-Si UV2 has a transitional composition between LV2 and high-Si UV2, and is distinguished from the high-Si UV2 by lower SiO2, and LILE, and higher Na2O, HFSE, and REE concentrations. The low-Si UV2 magmas were generated by the introduction of high-temperature hydrous fluids and lesser amounts of sedimentary melt. In contrast, generation of the high-Si UV2 magma required both hydrous fluid and sedimentary melt, as indicated by the Hf and Nd isotope geochemistry. The high-temperature mantle wedge indicated by the genetic

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conditions of the primary UV2 magma and the high-temperature slab recorded by the Cpx amphibolite in the metamorphic sole are consistent with a shallow and buoyant subduction zone, in which the absence of convection within the mantle wedge caused progressive remelting and depletion of the source mantle wedge, resulting in the short-lived (< 3 m.y.) V2 arc magmatism. Acknowledgements We thank Salim Al-Ibrahim, Mahommed Al-Battashi, Mohamed Al-Araimi (Ministry of

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Commerce and Industry of Oman), and the Japanese Embassy in Oman for their kind support during field surveys. We also thank K. Kanayama and R. Otsuka for their support during field surveys. We are indebted to N. Tsuchiya, E. Takazawa, T. Kurihara, and K. Hara for fruitful discussions. We appreciate the assistance of A. Tamura and Y. Soda with EPMA and LA–ICP–MS analyses at Kanazawa University. This study was supported by Fukada Grant-in-Aid 2013 from The Fukada Geological Institute and JSPS KAKENHI Grant No. JP25400446 to Y. Kusano, and by Monkasho Special Budget Project No. 812000009 to S. Arai.

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fractionation on quantitative analysis. Geochem. Jour. 39, 327–40. Morishita, T., Ishida, Y., Arai, S., Shirasaka, M., 2005b. Determination of multiple trace element compositions in thin (30 mm) layers of NIST SRM 614 and 616 using laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS). Geostandards and Geoanalytical Research. 29, 107–22. Moores, E. M., Kellogg, L. H., Dilek, Y., 2000. Tethyan ophiolites, mantle convection, and tectonic „„historical contingency‟‟: A resolution of the „„ophiolite conundrum,‟‟ in Dilek Y., et al. (Eds.), Ophiolites and Oceanic Crust: New Insights From Field Studies and the Ocean Drilling Program. Spec. Pap. Geol. Soc. Am., 349, pp. 3–12, Murton, B. J., Tindle, A. G., Andrew Milton, J., and Sauter, D. 2005. Heterogeneity in southern Central Indian Ridge MORB: Implications for ridge-hot spot interaction. Geochem. Geophys. Geosys. Doi: 10.1029/2004GC000798. Nagaishi, K., 2008. Fluid transfer and magmatism in the initial stage of subduction: inference from the Oman Ophiolite (Ph.D. thesis). Shizuoka University, Japan.

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Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Academic Publication, Norwell, Massachusetts. Nicolle, M., Jousselin, D., Reisberg, L., Bosch, D., Stephant, A., 2016. Major and trace element and Sr nd Nd isotopic results from mantle diapirs in the Oman ophiolite:

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Implications for off-axis magmatic processes. Earth. Planet. Sci. Lett., 473, 138-149. http://dx.doi.org/10.1016/j.epsl.2015.12.005. Noguchi, S., Morishita, T., Toramaru, A., 2004. Corrections for Na-loss on micro-analysis of glasses by electron probe X-ray micro analyzer. Japan. Mag.

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Mineral. Petrol. Sci., 33, 85-95 (in Japanese with English abstract). Pearce, J.A., Lippard, S.J., and Roberts, S., 1984. Characteristics and tectonicsignifi cance of supra-subduction zone ophiolites, in Kokelaar, B.P., and Howells, M.F., eds., Marginal basin geology: Geological Society of London Special Publication 16, p. 77–94, doi:10.1144/GSL.SP.1984.016.01.06. Pearce, J.A., Stern, R.J., Bloomer, S.H., Fryer, P., 2005. Geochemical mapping of the Mariana arc-basin system: implications for the nature and distribution of Geophys.

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Pearce, J. A., Kempton, P. D., Nowell, G. M., and Noble, S. R., 1999. Hf–Nd element and isotope perspective on the nature and provenance of mantle and subduction components in Western Pacific arc-basin systems. Jour. Petrol. 40, 1579–611. Plank, T., 2005. Constraints from thorium/lanthanum on sediment recycling at

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subduction zones and the evolution of the continents. Jour. Petrol. 46, 921-944. Plank, T., Langmuir, C. H., 1998. The chemical composition of subducting sediment and its consequences for the crust and mantle. Chem. Geol. 145, 325-394. Reagan, M.K., Ishizuka, O., Stern, R.J., Kelley, K.A., Ohara, Y., lichert-Toft, J., Bloomer, S.H., Cash, J., Fryer, P., Hanan, B.B., Hickey-Vargas, R., Ishii, T., Kimura, J., Peate, D.W., Rowe, M.C. and Woods, M., 2010. Fore-arc basalts and subduction initiation in the Izu-Bonin-Mariana system. Geochem. Geophys. Geosyst., Q03X12, doi:10.1029/2009GC002871. Rioux, M., Bowring, S., Kelemen, P., Gordon, S., Dudás, F., Miller, R., 2012. Rapid crustal accretion and magma assimilation in the Oman-U.A.E. ophiolite: High precision U-Pb zircon geochronology of the gabbroic crust. J. Geophys. Res. 117, doi:10.1029/2012JB009273. Rioux, M., Bowring, S., Kelemen, P., Gordon, S., Miller, R., Duda´ s, F., 2013. Tectonic development of the Samail ophiolite: High-precision U-Pb zircon geochronology and Sm-Nd isotopic constraints on crustal growth and emplacement. J. Geophys.

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Res. Solid Earth. 118, 2085–2101, doi:10.1002/jgrb.50139. Rollinson, H. 2015. Slab and sediment melting during subduction initiation: granitoid dykes from the mantle section of the Oman ophiolite. Contrib. Mineral. Petrol. Doi: 10.1007/x00410-015-1177-9. Shinjo, R., Ginoza, Y., Meshesha, D., 2010. Improved method of Hf separation from silicate rocks for isotopic analysis using the Ln-spec resin column. J. Mineral. Petrol. Sci. 105, 297-302. Shinjo, R., Goeku, N., Ikei, A. 2015. Development of single-step column separation

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method for Hf and Nd isotopic analyses of geological samples. Goldschmidt Abstracts 2015, 2877. Staudigel, H., Hart, S. R. 1983. Alteration of basaltic glass: Mechanisms and significance for the oceanic crust-seawater budget. Geochimica et Cosmochimica Acta, 47, 337–350. Stern, R. J., Reagan, M., Ishizuka, O., Ohara, Y., Whattam, S., 2012. To understand subduction initiation, study forearc crust: To understand forearc crust, study

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ophiolites. Lithosphere. 4, 469–483. Stolz, A. J., Jochum, K. P., Spettel, B., Hofmann, A. W., 1996. Fluid- and melt-related enrichment in the subarc mantle: Evidence from Nb/Ta variations in island-arc

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basalts. Geology. 24, 587–590. Sun, S.S., McDonough, W. F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A. D., Norry, M. J. (Eds), Magmatism in the Ocean Basins. Geological Society, London,

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Special Publications, 42, 313–345, doi:10.1144/GSL.SP.1989.042.01.19. Takazawa, E., Okayasu, T., Satoh, K., 2003. Geochemistry and origin of the basal lherzolites from the northern Oman ophiolite (northern Fizh block). Geochem. Geophys. Geosys. 4, 1021, doi:10.1029/2001GC000232. Tatsumi, Y. 2000. Continental crust formation by crustal delamination in subduction zones and complementary accumulation of the enriched mantle I component in the mantle. Geochem. Geophys. Geosys. 1, 2000GC000094. Tsuchiya, N., Shibata, T., Yoshikawa, M., Adachi, Y., Miyashita, S., Adachi, T., Nakano, N., Osanai, Y., 2013. Petrology of Lasail plutonic complex, northern Oman ophiolite, Oman: An example of arc-like magmatism associated with ophiolite detachment. Lithos. 156–159, 120–138. Umino, S., Yanai, S., Jaman, A.R., Nakamura, U., Iiyama, J. T., 1990. The transition from spreading to subduction: Evidence from the Semail ophiolite, northern Oman mountains. In: Malpas, J., Moores, E.M., Panayiotou, A., Xenophontos, C.

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(Eds), Ophiolites: Oceanic Crustal Analogues, Proceedings of the Symposium, Troodos, 1987. Geological Survey Department, Nicosia, pp. 375-384. Umino, S., Miyashita, S., Hotta, F., Adachi, Y., 2003. Along-strike variation of the sheeted dike complex in the Oman Ophiolite: Insights into subaxial ridge segment

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structures and the magma plumbing system. Geochem. Geophys. Geosys. 4, 8618, doi:10 .1029/2001GC000233. Umino, S., 2012. Emplacement mechanism of off-axis large submarine lava field from

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the Oman Ophiolite. J. Geophys. Res. 117, doi:10.1029/2012JB009198. Umino, S., Kitamura, K., Kanayama, K., Tamura, A., Sakamoto, N., Ishizuka, O. and Arai, S., 2015. Thermal and chemical evolution of the subarc mantle revealed by spinel-hosted melt inclusions in boninite from the Ogasawara (Bonin)

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Archipelago, Japan. Geology, 43, 151-154, doi: 10.1130/G36191.1 Umino, S., Kitamura, K., Kanayama, K., Kusano, Y., Nagaishi, K., Ishikawa, T., Ishizuka, O. 2014. Prolonged vs. Failed Subduction Zone. Goldschmidt2014

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abstracts, 2532. Vervoort, J. D., Patchett, P. J., Blichert-Toft, J., Albarede, F., 1999. Relationships between Lu–Hf and Sm–Nd isotopic systems in the global sedimentary system.

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Earth. Planet. Sci. Lett. 168, 79–99. Yamazaki, S., 2013. Incipient island arc crust formation within oceanic crustal sequence: Geology, geochemistry and geochronology of late intrusive rocks in the Oman ophiolite (Ph.D. thesis), Niigata University, Japan.

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Caption of figures

Fig. 1. Geological map of the study area, modified from the 1:250,000 geological maps of Buraymi, Ibri, Nizwa, and Seeb (Ministry of Petroleum and Minerals, Sultanate of Oman, 1992). Sample localities of volcanic glasses and rocks (details are described in Kusano et al., 2012, 2014) are shown. Fig. 2. Microphotographs of volcanic glass. (a) V1 volcanic glass with thin hydration halos along cracks (BSE-SEM image, sample 12Jiz1). (b) LV2 volcanic glass containing Pl and Cpx microlites (BSE-SEM image, sample OM16-113). (c) UV2 containing partially preserved Ol phenocrysts 0.5–1.0 mm in diameter with chrome spinel inclusions (open polars, sample OM14-2C). Most of the Ol has been replaced by carbonate and clay minerals. Fig. 3 K2O, Rb, Sr, Ba plotted against Yb concentrations. Fig. 4. SiO2, TiO2, Al2O3, FeO, MnO, CaO, Na2O, K2O, and K2O/TiO2 ratio plotted against MgO on an anhydrous basis. White circles and diamonds show altered

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samples of the V1 and LV2 glasses. Sheeted dikes and V1 tholeiite (gray circles) (Einaudi et al., 2003; Godard et al., 2003, 2006; Miyashita et al., 2003; Umino et al., 2003) and V2 tholeiite and boninite (black diamonds and gray crosses) (Ishikawa et al., 2002; Godard et al., 2003, 2006) are plotted for comparison. Large gray circles and black diamonds represent the V1 and LV2 samples, respectively (Kusano et al., 2012, 2014), utilized for Hf and Nd isotopic analyses. Fig. 5. (a) Primitive-mantle-normalized multi-element diagrams for the V1 samples. The hatched and shaded areas represent the V1 compositional ranges after Kusano

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et al. (2012) and the Indian MORB glass taken from PetDB (available at www.earthchem.org/petdb: Mahoney et al., 2002; Meyzen et al., 2003; Murton et al., 2005), respectively. (b) N-MORB-normalized multi-element diagrams for LV2. The shaded area in (b) indicates the compositional range of the V1 glass. (c) N-MORB-normalized multi-element diagrams for UV2. The primitive mantle and N-MORB values are after Sun and McDonough (1989).

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Fig. 6. εNd(t) versus εHf(t) variations for the bulk V1 and V2 samples. (a) All samples from extrusive sequences plot in the Indian mantle domain (above the black line) of Pearce et al. (1999). The V1 and LV2 samples plot along the mantle array

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(dashed line) of Vervoort et al. (1999). The shaded and cross hatched areas are the εNd(t) ranges of V1 and V2, respectively, from Benoit et al. (1996), Godard et al. (2006), and Tsuchiya et al. (2013). (b) The UV2 samples plot on the mixing line of V1 residue and pelagic sediment. Correlations between the volcanic product and residue are shown in Table 3. The V1 residue is assumed to have the average LV1 isotopic compositions (143Nd/144Nd initial = 0.512942, 176Hf/177Hf initial = 0.283126), and Nd and Hf compositions of the mantle peridotite (Nd = 0.037 μg/g, Hf = 0.022 μg/g; Takazawa et al., 2003). Compositions of pelagic sediments and continental materials in the Indian Ocean are from Handley et al. (2011). The area enclosed by the dashed line indicates the compositional range of felsic dikes intruded into the mantle section (Haase et al., 2015). Data from the Oman Ophiolite are corrected at 95 Ma. Fig. 7. Nb/Ta–Nb, Th/Zr–Ba/Zr, and (Pb/Nd)n–(K/Nd)n diagrams, and B/La, Th/La, and Pb/Ce ratios plotted against Nb/Ta ratios of the Oman Ophiolite extrusive sequences. Large gray circles, black diamonds, and gray crosses denote the bulk V1, LV2, and UV2 compositions, respectively (Kusano et al., 2012, 2014); small symbols indicate data of Godard et al. (2003, 2006). The values Th/La = 0.2 and Pb/Ce = 0.5 are from Plank (2005) and Miller et al. (1994). Pb/Nd and K/Nd ratios are normalized by N-MORB values are after Sun and McDonough (1989).

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Fig. 8. (a) Trace element spectra of subduction component and model trace-element spectra of the V2 magmas compared with the compositions of (b) LV2, (c) low-Si UV2, and (d) high-Si UV2 glasses. The parameters used for the modeling are listed in Table 4.

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Fig. 9. Schematic cartoon showing the development and cessation of V2 arc magmatism in the Oman Ophiolite (not to scale). Intraoceanic thrusting was initiated near a ridge axis and developed into a shallow, hot subduction zone (Nicolas, 1989; Ishikawa et al., 2005; Kusano et al., 2014). (a) LV2 magmatism started at ca. 94

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Ma, within 2 m.y. after the spreading stage (V1 magmatism). The island-arc tholeiitic magma was produced by wet partial melting of the V1 residue with hydrous fluid. (b) UV2 magmatism occurred later in the subduction stage. At the peak temperature and pressure conditions for the highest-grade amphibolite recorded in the metamorphic sole (770–900 °C at 1.1–1.3 GPa; Cowan et al., 2014), high-temperature hydrous fluid and felsic melt liberated from the subducted slab caused remelting of the LV2 residue to produce boninite magma at 1320 °C and 0.5 GPa (Umino et al., 2014). As the hot and buoyant slab suppressed mantle convection, the overlying mantle wedge was cooled, leading to

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the end of arc magmatism.

Fig. A1. Concentrations of B, Rb, Sr, Cs, Ba, La, Pb, and Th plotted against MgO. Fig. A2. εHf(t) and εNd(t) versus SiO2 and Zr diagrams for the bulk V1 and V2 samples.

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Fig. A3. εNd(t) versus Th/Yb diagram for the bulk V1 and V2 samples. The UV2 samples show a negative correlation in the diagram, but low εNd(t) sample of the LV2 show similar Th/Yb ratio to most of the LV2. The low εNd(t) seems to be mixed with cretaceous seawater of Kawahata et al. (2001), suggesting that the low εNd(t) produced by secondary alteration rather than sediment melt mixing. On the other hand, the UV2 samples are plotted along the mixing line between the V1 residue and mixed fluid. Composition of V1 residue: Th = 0.0010 μg/g, Yb =0.210 μg/g (Takazawa et al., 2003), εNd(t) = same as Fig. 6; pelagic sediment: Handley et al. (2011) and Plank and Langmuir (1998); mixed fluid: 143Nd/144Nd initial = 0.512461, Th/Yb = 4.804 (pelagic sediment and amphibolite-derived fluid are mixed by 0.27:0.63 based on the mixing ratio of Fig. 8). Fig. A4. MgO-CaO variation diagram of East Pacific rise and V1 lavas and sheeted dikes. Bulk and glass compositions are from Gale et al. (2013), Einaudi et al. (2003), Godard et al. (2003; 2006), Miyashita et al. (2003), Umino et al. (2003),

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Kusano et al. (2012) and this study.

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Title of tables

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Table 1 Major- and trace-element compositions of volcanic glass samples from the Oman Ophiolite.

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Table 2 Whole-rock Hf and Nd isotopic compositions. Table 3 Lava-source-residue pairs of geochemical modeling. Table 4 Estimated primary magmas and parameters used for the geochemical modeling.

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LV2

IP

V1

V1

V1

V1

LV2

LV2

LV2

Sampl

12Jiz

OM11-2

OM17-

OM17

OM16-1

OM17-

OM17-

e No.

1

2A

67Aβ

-67B

13

59C

OM16

12SH

12SH

59D

-73B

12

13

Pyrocl

Pyrocl

Pyrocl

Pyrocl

Pyrocl

astic

astic

astic

astic

astic

bomb

clasts

clasts

clasts

clasts

W.

W.

W.

W. Jizi

Majan

Majan Hatta

Hilti

Hilti

Crust Pillow

Crust of

Occurr

lobe

Hyalo-cl

of

Hyalo-cl

ence

margi

astite

pahoeh

pahoe

astite hoe

n

oe lobe

W.

W.

MA

lobe Localit

LV2

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LV2

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Table 1. Volcanic glass major and trace element compositions.

W.

Suq

Shafan

Latitud

24.20

23.5236

24.342

24.34

24.2983

24.376

24.376

24.76

24.11

24.11

274

9

67

267

1

78

78

269

839

839

56.30

56.4805

56.506

56.50

56.4253

56.514

56.514

56.37

56.56

56.56

160

7

48

648

0

48

48

422

087

087

Longit ude

SiO2 55.09

CE P

e

Suq

TE

y

D

W. Suq

59.74

61.90

60.37

62.73

58.69

60.12

56.57

55.24

58.17

1.51

1.41

1.62

0.98

0.66

0.71

0.47

0.49

0.40

14.35

15.42

14.85

14.60

15.36

15.58

14.95

16.04

17.11

0.03

n.a.

0.02

0.01

0.02

0.02

0.03

0.05

0.01

0.03

11.52

11.18

8.84

9.72

9.85

8.12

8.07

6.22

7.72

6.68

MnO

0.18

0.16

0.20

0.18

0.15

0.14

0.14

0.19

0.14

0.15

MgO

4.01

2.09

2.00

2.32

2.44

4.95

4.07

7.70

7.06

4.94

CaO

10.53

7.83

5.61

6.02

7.29

9.42

8.66

11.69

11.78

10.41

Na2O

1.29

2.40

3.49

3.72

1.51

2.29

2.23

1.93

1.34

1.65

K2O

0.45

0.74

1.12

1.19

0.44

0.35

0.40

0.23

0.18

0.48

0.24

0.49

0.79

0.73

0.45

0.53

0.56

0.49

0.37

1.21

2.411

2.220

1.393

1.912

2.462

7.336

1.293

6.879

1.671

n.a.

3.480

4.878

5.878

1.288

8.877

7.756

8.599

11.02

6.337

n.a.

(wt%) 1.89

Al2O3

15.00

Cr2O3 FeO*

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TiO2

K2O/Ti O2 Li (μg/g) B

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0 20.48

17.38

21.12

34.16

32.49

32.34

39.01

37.13

n.a.

331

81

56

105

317

302

305

272

268

n.a.

Cr

12.83

4.25

2.79

5.25

2.97

65

62

315

58

n.a.

Co

31.9

16.6

12.9

17.3

29.3

31.3

29.7

47.6

37.3

n.a.

Ni

12.2

3.4

3.2

4.6

7.2

36.9

40.6

122.7

44.5

n.a.

Rb

15.88

21.07

3.61

5.63

7.74

8.58

10.13

5.69

3.41

n.a.

Sr

433.3

151.7

137.3

142.1

132.9

68.9

84.9

78.4

101.8

n.a.

Y

45.4

67.0

61.9

63.6

30.2

19.5

23.3

15.9

13.7

n.a.

Zr

146.9

194.2

212.4

219.6

68.4

45.2

55.6

34.7

26.2

n.a.

Nb

3.514

4.128

4.818

5.041

1.264

1.059

1.200

0.609

0.493

n.a.

Cs

0.393

0.088

0.065

0.059

0.765

0.546

0.612

0.314

0.104

n.a.

Ba

77.6

54.0

56.0

54.8

49.7

42.6

52.8

33.3

33.2

n.a.

La

6.003

7.116

8.315

8.416

2.727

1.685

2.018

1.174

0.977

n.a.

22.296

24.450

7.803

4.931

5.674

3.606

2.932

n.a.

1.299

0.799

0.934

0.591

0.484

n.a.

7.202

4.346

5.121

3.309

2.743

n.a.

18.53 Ce

3.725

15.76 Nd

4.028

20.769

Sm

5.079

Eu

1.723

Gd

6.568

Tb

1.170

Dy

AC

CE P

6

Ho

3.930

IP

21.41

20.380

7.984

SC R

9

TE

2.963

NU

25.18

9 Pr

MA

V

T

27.98

D

Sc

4

6.839

6.670

6.903

2.635

1.567

1.869

1.309

1.063

n.a.

2.269

2.102

2.196

0.940

0.576

0.671

0.479

0.446

n.a.

9.249

8.584

8.733

3.773

2.364

2.787

1.897

1.637

n.a.

1.637

1.540

1.569

0.695

0.433

0.526

0.363

0.309

n.a.

5.019

3.174

3.802

2.606

2.265

n.a.

11.446

10.81 10.385 0

1.675

2.411

2.211

2.266

1.088

0.693

0.833

0.562

0.497

n.a.

Er

5.048

7.382

6.675

6.968

3.419

2.185

2.617

1.792

1.544

n.a.

Tm

0.710

1.054

0.957

0.988

0.504

0.317

0.390

0.257

0.225

n.a.

Yb

4.985

7.113

6.475

6.645

3.423

2.264

2.709

1.839

1.575

n.a.

Lu

0.740

1.076

0.964

1.000

0.501

0.355

0.417

0.282

0.237

n.a.

Hf

4.030

5.136

5.639

5.709

2.062

1.360

1.723

1.083

0.809

n.a.

Ta

0.243

0.279

0.311

0.338

0.077

0.080

0.087

0.044

0.031

n.a.

Pb

0.960

1.285

1.475

1.344

1.546

2.539

2.792

1.451

0.781

n.a.

Th

0.367

0.404

0.535

0.541

0.236

0.178

0.216

0.125

0.075

n.a.

U

0.148

0.148

0.210

0.215

0.144

0.164

0.179

0.109

0.049

n.a.

* Total Fe as

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FeO. Major and trace element compositions are average values of four analytical points on each sample.

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IP

2008) and recalculated to unhydrous basis. RSD: relative standard deviation.

T

Reference values of BCR-2G and NIST 612 are selected from the GeoReM database (Jochum and Nohl,

UV2

UV2

Low

Low

Low

UV2

UV2

Low Si Si

Si

Sampl

OM17

OM17

OM17

OM17-

e No.

-62A

-62D

-62E

63C-1

Pillow

Pillow

Pillow

margi n

margi n

Localit Majan

UV2

UV2

UV2

High

High

High

High

High

Si

Si

Si

Si

Si

OM17-

OM16

OM1

OM16

OM17

OM17

OM17-

63C-2

-84A

4-2C

-46C

-28C

-80A

97B-2

Pillow

Pillow

Pillow

lobe

lobe

lobe

Epiclas

margi

margi

margi

t

n

n

n

W.

W.

W.

W.

Humay

Suq

Suq

Salahi

Suq

dah

Low Si

Majan

Pillo

Pillow

Pillow

Pyrocl

w

lobe

lobe

astic

lobe

margin

margin

clasts

margi

lobe

margi

Majan

UV2 High Si

n

AC

y Latitud

lobe

TE

ence

lobe

CE P

Occurr

UV2

D

Si

UV2

MA

UV2

NU

Table 1. Volcanic glass major and trace element compositions.

n Majan

Majan

Bidah

24.36

24.36

24.36

24.360

24.360

24.64

24.33

24.33

24.15

24.33

24.575

184

184

184

35

35

326

134

143

135

884

32

56.51

56.51

56.51

56.508

56.508

56.39

56.48

56.48

56.55

56.51

56.388

050

050

050

59

59

441

871

859

197

056

58

54.77

54.52

54.76

54.21

54.38

56.00

55.38

55.73

55.49

55.04

56.15

TiO2

0.41

0.35

0.35

0.36

0.35

0.42

0.26

0.29

0.32

0.29

0.28

Al2O3

15.09

14.75

14.91

14.82

14.81

15.26

14.59

14.88

14.74

15.21

15.68

Cr2O3

0.03

0.06

0.04

0.03

0.06

0.00

0.04

0.04

0.03

0.03

0.02

FeO*

7.67

7.54

7.55

7.70

7.61

7.60

8.06

7.02

7.69

7.80

7.94

MnO

0.14

0.16

0.11

0.15

0.13

0.15

0.14

0.21

0.15

0.15

0.13

MgO

7.99

8.66

8.38

8.71

8.68

7.28

8.03

7.71

8.24

8.02

6.68

e Longit ude

SiO2 (wt%)

ACCEPTED MANUSCRIPT

12.61

12.89

12.78

12.90

12.88

12.09

12.28

12.95

12.30

12.43

12.12

Na2O

1.13

0.93

0.97

0.98

0.95

0.99

0.98

0.96

0.82

0.84

0.75

K2O

0.17

0.14

0.15

0.14

0.15

0.21

0.24

0.21

0.22

0.19

0.24

0.41

0.41

0.42

0.40

0.43

0.48

0.91

0.71

0.67

0.64

0.84

5.566

4.768

4.814

4.890

4.774

n.a.

2.871

5.823

6.558

9.294

9.868

7.885

8.932

n.a.

T

CaO

Li (μg/g)

11.22 B

7.352

14.44

16.31

10.32 8.175

0

29.421 0

41.89

42.83

42.97

n.a.

42.30

45.27

33.25

43.25

42.02

V

260

264

261

268

261

n.a.

278

298

216

270

273

Cr

200

354

339

365

352

n.a.

241

186

236

192

89

Co

37.0

38.3

37.8

38.5

37.3

n.a.

42.0

45.9

28.4

38.4

34.8

Ni

50.1

64.4

62.7

61.3

61.2

n.a.

67.7

70.2

50.0

59.2

41.1

Rb

5.29

4.24

4.27

4.56

4.41

n.a.

7.03

7.75

6.10

6.32

8.31

Sr

42.0

36.9

37.6

37.5

37.6

n.a.

29.3

32.2

20.9

27.3

25.7

Y

12.4

10.3

10.6

10.5

10.7

n.a.

9.4

10.5

7.4

9.4

8.8

Zr

19.6

13.3

13.6

13.6

13.9

n.a.

9.6

10.7

8.1

8.7

11.1

Nb

0.597

0.429

0.436

0.443

0.458

n.a.

0.461

0.527

0.489

0.448

0.632

Cs

0.444

0.383

0.377

0.408

0.396

n.a.

0.636

0.703

0.642

0.609

0.820

Ba

19.6

16.0

16.6

16.5

16.5

n.a.

26.3

29.6

17.4

21.0

21.4

La

0.891

0.630

0.643

0.657

0.659

n.a.

0.382

0.428

0.556

0.481

0.817

Ce

2.509

1.723

1.749

1.787

1.773

n.a.

1.008

1.102

1.337

1.157

2.001

Pr

0.395

0.268

0.276

0.276

0.277

n.a.

0.154

0.172

0.184

0.168

0.256

Nd

2.180

1.492

1.536

1.525

1.545

n.a.

0.945

1.048

0.945

0.955

1.324

Sm

0.873

0.623

0.644

0.657

0.669

n.a.

0.455

0.534

0.401

0.466

0.509

Eu

0.347

0.254

0.270

0.266

0.269

n.a.

0.199

0.234

0.163

0.196

0.203

Gd

1.336

1.037

1.074

1.071

1.074

n.a.

0.870

1.002

0.696

0.849

0.809

Tb

0.266

0.215

0.219

0.225

0.227

n.a.

0.183

0.212

0.147

0.184

0.173

Dy

1.996

1.640

1.662

1.697

1.699

n.a.

1.452

1.669

1.156

1.452

1.347

Ho

0.452

0.369

0.385

0.390

0.393

n.a.

0.344

0.377

0.269

0.334

0.311

Er

1.437

1.202

1.229

1.247

1.263

n.a.

1.117

1.255

0.879

1.120

1.032

Tm

0.209

0.182

0.184

0.188

0.192

n.a.

0.162

0.189

0.135

0.171

0.160

Yb

1.492

1.314

1.346

1.348

1.343

n.a.

1.244

1.359

0.975

1.236

1.177

Lu

0.229

0.202

0.205

0.206

0.210

n.a.

0.190

0.216

0.156

0.194

0.179

D

MA

41.78

CE P

39.18

TE

1

AC

Sc

6.735

NU

2

SC R

O2

IP

K2O/Ti

ACCEPTED MANUSCRIPT

0.612

0.436

0.463

0.469

0.468

n.a.

0.360

0.409

0.281

0.343

0.385

Ta

0.048

0.033

0.037

0.036

0.038

n.a.

0.042

0.048

0.042

0.042

0.053

Pb

1.503

1.397

1.406

1.499

1.415

n.a.

2.297

2.407

1.890

1.812

2.355

Th

0.124

0.097

0.099

0.107

0.103

n.a.

0.096

0.107

0.126

0.208

U

0.102

0.099

0.100

0.101

0.099

n.a.

0.135

0.144

0.191

T

Hf

IP

0.139 0.155

TE

D

MA

NU

SC R

0.141

Table 1. Volcanic glass major and trace element compositions.

Sample No.

CE P

Reference standard

BCR-2G Average

Occurrence Locality Latitude

AC

(n=6)

NIST612

Standard

Reference

Standard

deviation

Reference RSD

value

(n=4)

Longitude

SiO2 (wt%)

Average

RSD

56.03

0.54

0.01

55.37

TiO2

2.23

0.07

0.03

2.31

Al2O3

13.71

0.07

0.01

13.64

Cr2O3

0.002

0.004

2.46

0.003

FeO*

12.20

0.27

0.02

12.62

MnO

0.16

0.06

0.37

0.19

MgO

3.66

0.06

0.02

3.62

CaO

7.18

0.06

0.01

7.19

Na2O

3.02

0.25

0.08

3.29

deviation

value

ACCEPTED MANUSCRIPT

K2O

1.80

0.05

0.03

1.77

B

34.60

Sc

39.94

V

0.95

0.02

40.20

IP

38.31

0.78

0.02

34.30

0.23

0.01

SC R

Li (μg/g)

T

K2O/TiO2

37.80

0.39

0.01

38.80

38.00

0.99

0.03

36.40

35.42

0.41

0.01

35.50

30.84

1.34

0.04

38.80

30.31

0.36

0.01

31.40

80.15

0.47

0.01

78.40

42.98

0.42

0.01

38.30

40.95

0.47

0.01

37.90

38.36

0.18

0.00

38.90

41.82

0.37

0.01

42.70

38.74

0.23

0.01

39.30

36.35

0.36

0.01

36.00

39.13

0.24

0.01

38.40

39.17

0.33

0.01

37.90

37.01

0.16

0.00

35.50

39.14

0.35

0.01

37.70

36.42

0.50

0.01

35.60

40.23

1.09

0.03

37.30

40.71

1.49

0.04

37.60

38.94

0.29

0.01

35.50

Ho

40.63

0.58

0.01

38.30

Er

41.87

0.47

0.01

38.00

Tm

40.50

0.64

0.02

36.80

Yb

41.37

0.66

0.02

39.20

Lu

40.84

0.58

0.01

37.00

Hf

39.44

0.71

0.02

36.70

Ta

39.86

1.20

0.03

37.60

Pb

38.76

0.90

0.02

38.57

Th

39.08

0.36

0.01

37.79

U

37.17

0.24

0.01

37.38

Cr Co

NU

Ni Rb Sr

MA

Y Zr Nb

D

Cs

TE

Ba La

Nd Sm Eu Gd Tb Dy

AC

Pr

CE P

Ce

AC

CE P

TE

D

MA

NU

SC R

IP

T

ACCEPTED MANUSCRIPT

ACCEPTED MANUSCRIPT

Sub

Sa Lo

qu

seq

mpl

(μ

cal en

uen

e

ce

No.

Sm

143

(μg/

d/

g)

Nd

N

144

g/

ity ce

147

Nd

ε Lu

N 2S E*

Sm /

144

d(

Nd

Fiz h

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4.25

0.51

5

3071

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NU

LV1

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0.

00

0.2

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V1

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(μ

(μg/

/

g)

f

8.

176

ε

for

2S

Lu/

Hf

geolo

E*

177

(t)

gy

Hf

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and

Hf

177

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petrol ogy Kusan 0.

0.

0.28 3.11

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Refer ence

g)

*

W.

176

g/

t)*

g)

Hf

SC R

Se

IP

T

Table 2. Bulk-rock Hf and Nd isotope compositions.

27 00

1

11

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V1

LV1

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h3

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(2014 1 )

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Kusan

ACCEPTED MANUSCRIPT

Fiz

Fiz

04

h

h4

7

6

2764

00

345

00

1

16

5

2

5

325

00

75

3.

o et

0

00

8

6

al.

9

(2012

10

IP

T

12

error; **: t = 95 Ma

143

εHf (t) = [(

Nd/

144

Nd)CHUR - 1]

176

Hf) sample/(

Hf/

177

Hf)CHUR - 1]

Hf)CHUR = 0.282785

TE

D

MA

Hf/

177

143

Nd)CHUR = 0.512630

177

Hf/

Nd) sample/(

CE P

*10000, (

144

Nd/

176

176

144

Nd/

AC

*10000, (

143

NU

εNd(t) = [(

SC R

*: SE = standard

)

ACCEPTED MANUSCRIPT

Table 3. Lava-source-residue pairs of geochemical modeling. V1

V1

(Lherzolite)

LV2

Slab components

Residue

T

Source

IP

Lava

depleted V1

(Lherzolite)

depleted V1

(Lherzolite)

LV2

type I Lherzolite

low-Si UV2

type II Lherzolite

SC R

Unit

hydrous fluid (type I and II fluids)

UV 2

hydrous fluid (type I and II fluids)

type II Lherzolite

hydrous fluid (type I and II) and

Harzburgite

sediment melt

MA

high-Si UV2

NU

Harzburgite

Lherzolites and harzburgite of Takazawa et al. (2003).

b

Estimated amphibolite-derived fluid of Ishikawa et al. (2005).

c

Refractory harzburgite of Kanke and Takazawa (2014).

AC

CE P

TE

D

a

b

type I Lherzolite Harzburgite

a

a

type II Lherzolite

a

Refractory harzburgite c

Refractory harzburgite

ACCEPTED MANUSCRIPT

T

Table 4. Estimated primary magmas and parameters used for the geochemical modeling. Estimated primary

IP

Residues equilibrium with primary magma magmas Depleted

Depleted

LV

Less

depleted V1

V1

V1

2

depleted V1

96om11

96om11

07Fz22

d

no.

2

e

2 4

90.4

90.4

V1

V1

type I

type I

harzburg

lherzolite

lherzolite

ite

LV2

OM16-73

e

Added 18

Depleted

6

Ol%

6.5

MA

Fo

NU

Sample

Depleted

SC R

Less

content

90.

of

91.5

7

primary

D

magmas

TE

Modal composit

K

Li B

2668

AC

residue

CE P

ion of

17.60

802

4.11

type II lherzolit e 17 786

3.04

0.76

0.23

79 6.4 4.02

0.09 4

30 Ti

7336

3766

3693

774

399

286

247

0

0

0

0.08

2.969

1.999

0.689

1.769

3.458

1.688

1.080

1.051

3.025

1.039

0.709

0.575

0.006

0.001

0.001

0.001

06 5.3 Rb

3.80

1.18

1.15 3 73.

Sr

167.72

112.93

110.68 40 14.

Y

33.86

16.53

16.20 90 32.

Zr

97.03

33.32

32.66 53

Nb

2.11

0.43

0.42

0.5

ACCEPTED MANUSCRIPT

7 97.91

9.34

9.15

0.009

0.001

0.000

0.380

0.026

T

31. Ba

0.009

0.003

0.007

0.131

0.050

0.018

0.033

0.034

0.013

0.005

0.009

0.264

0.106

0.041

0.070

0.146

0.059

0.025

0.089

0.058

0.029

0.013

0.020

0.264

0.114

0.054

0.095

0.055

0.026

0.013

0.020

0.387

0.187

0.096

0.156

0.090

0.045

0.025

0.035

0.290

0.145

0.087

0.106

0

0

0.315

0.155

0.102

0.144

0.051

0.026

0.018

0.025

0.119

0.041

0.022

0.039

0

0

0.000

19 1.1 3.60

1.29

1.26

IP

La

3.3 Ce

10.83

4.15

4.07 8 0.5

Pr

1.80

0.71

0.70 5

9.71

3.90

3.83

NU

3.1 Nd

SC R

0

0

1.2 3.29

1.33

1.30

MA

Sm

3

0.4

Eu

1.11

0.55

0.54

Dy

Ho

Er

Tm

TE

0.79

1.89

0.37

CE P

Tb

4.38

5.02

1.14

AC

Gd

3.35

0.51

D

5

2.43

0.57

1.7

1.86 8 0.3 0.37 4 2.4 2.38 4 0.5 0.56 3 1.6

1.68

1.64 8 0.2

0.26

0.26 4 1.7

Yb

3.23

1.60

1.57 3 0.2

Lu

0.48

0.24

0.24 6 1.0

Hf

2.44

0.84

0.82 1 0.0

Ta

0.14

0.05

0.05

0.0000

4 Pb

0.00

0.19

0.19

1.3

6 0

0.0020

0.001

0

ACCEPTED MANUSCRIPT

6 0.1 Th

0.08

0.05

0.0003

0.04

0.00018

0.00010

0.00003 9

0.1 0.08

0.04

0.04

0.00003 0

Estimated amphibolite-derived fluid composition of Ishikawa et al.

(2005).

Indian pelagic sediment composition of Plank and Langmuir (1998).

d

Primary magma composition from Kusano et al.

MA

(2012). Primary magma composition from Miyashita et al.

(2003). f

NU

c

e

0

Assumed value based on the reported mineral/melt distribution coefficients (Kelemen et al., 2003) and

mineral assemblages (Takazawa et al., 2003). b

0.0006 0.00000

SC R

a

0.00001

IP

U

T

2

Whole-rock metachart composition in Wadi Tayin of Ishikawa et al.

AC

CE P

TE

D

(2005).

Table 4. Estimated primary magmas and parameters used for the geochemical modeling. Indian Bulk partition coefficient peridotite/melt

Amphibolite-derived

a

Fluid

Metachar pelagic

b

t sediment YY04mc

V1 residue Sample no. Added Ol%

V1 residue

LV2 residue

type I

type II

f

c

ACCEPTED MANUSCRIPT

Fo content of primary

T

magmas type I

harzburgit

type II

lherzolite

e

lherzolite

IP

Modal

n of residue K

0.0009

0.0003

0.0007

Li

0.073

0.023

0.051

54400

4183

21916

45.6

161

8.86

100

100

0

19900

773

4459

0.078

0.089

Rb

0

0

0

41.3

64

15.94

82.6

Sr

0.02

0.01

0.01

6.34

2370

27.97

251

Y

0.10

0.07

0.08

0

14.7

19.24

22.1

Zr

0.03

0.02

0.03

0

244

25.39

165

Nb

0.003

0.002

0.002

0

38.1

3.17

11.26

Ba

0.00009

0.00003

0.00006

129

510

63.51

543

La

0.007

0.003

0.005

0

38.6

13.08

31.15

Ce

0.012

0.005

0.009

0

93.5

22.16

67

Pr

0.019

0.007

0.014

0

12.7

3.56

Sm Eu Gd Tb Dy Ho Er

D

TE

CE P

Nd

NU

0.106

0.027

0.011

0.020

0

53.4

14.63

28.36

0.044

0.019

0.033

0

11.2

3.34

5.24

0.052

0.024

0.039

0

3.74

0.76

1.06

0.060

0.029

0.046

0

8.68

3.45

4.01

0.069

0.035

0.053

0

1.09

0.49

0.077

0.041

0.059

0

4.41

3.21

0.078

0.044

0.062

0

0.62

0.63

0.087

0.053

0.070

0

0.97

1.85

0

0

0.25

AC

Ti

30900

MA

B

SC R

compositio

Tm

3.98

2.3

Yb

0.097

0.065

0.081

0

0

1.56

2.14

Lu

0.107

0.076

0.091

0

0

0.23

0.322

Hf

0.049

0.027

0.038

0

0

0.68

6.18

Ta

0

0.003

0

0

0

0.21

0.803

Pb

0.010

0.004

0.007

1.29

3.11

3.81

24.5

Th

0.002

0.001

0.001

0

0

2.15

10.28

0.00040

0.00013

0.00028

0

0

0.26

2.26

U a

Assumed value based on the reported mineral/melt distribution coefficients (Kelemen et al., 2003) and

ACCEPTED MANUSCRIPT

mineral assemblages (Takazawa et al., 2003). Estimated amphibolite-derived fluid composition of Ishikawa et al. (2005).

c

Indian pelagic sediment composition of Plank and Langmuir (1998).

d

Primary magma composition from Kusano et al. (2012).

e

Primary magma composition from Miyashita et al. (2003).

CE P

TE

D

MA

IP

NU

SC R

Whole-rock metachart composition in Wadi Tayin of Ishikawa et al. (2005).

AC

f

T

b

AC

CE P

TE

D

MA

NU

SC R

IP

T

ACCEPTED MANUSCRIPT

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CE P

TE

D

MA

NU

SC R

IP

T

ACCEPTED MANUSCRIPT

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D

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SC R

IP

T

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D

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SC R

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T

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D

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SC R

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T

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D

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SC R

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T

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D

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SC R

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T

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TE

D

MA

NU

SC R

IP

T

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CE P

TE

D

MA

NU

SC R

IP

T

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CE P

TE

D

MA

NU

SC R

IP

T

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CE P

TE

D

MA

NU

SC R

IP

T

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CE P

TE

D

MA

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SC R

IP

T

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CE P

TE

D

MA

NU

SC R

IP

T

ACCEPTED MANUSCRIPT