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Recycled oceanic crust as a source for tonalite intrusions in the mantle section of the Khor Fakkan block, Semail ophiolite (UAE)
Geoscience Frontiers 10 (2019) 1187e1210
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Research Paper
Recycled oceanic crust as a source for tonalite intrusions in the mantle section of the Khor Fakkan block, Semail ophiolite (UAE) Hosung Joun a, Sotirios Kokkalas a, *, Stylianos Tombros b a b
Department of Earth Sciences, Khalifa University, 2533, Abu Dhabi, United Arab Emirates Department of Geology, University of Patras, 26442, Greece
a r t i c l e i n f o
a b s t r a c t
Article history: Received 1 November 2017 Received in revised form 10 June 2018 Accepted 11 September 2018 Available online 12 October 2018 Handling Editor: Sanghoon Kwon
Several types of felsic granitoid rocks have been recognized, intrusive in both the mantle and the crustal sequence of the Semail ophiolite. Several models have been proposed for the source of this suite of tonalites, granodiorites, trondhjemites intrusions, however their genesis is still not clearly understood. The sampled Dadnah tonalites that intruded in the mantle section of the Semail ophiolite display arctype geochemical characteristics, are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, show positive peaks for Ba, Pb, Eu, negative troughs for U, Ti and occur with low d18OH2O, moderate εSr and negative εNd values. They have crystallized at temperatures that range from w550 C to w720 C and pressure ranging from 4.4 kbar to 6.5 kbar. The isotopic ages from our tonalite samples range between 98.6 Ma and 94.9 Ma, slightly older and overlapping with the age of the metamorphic sole. Our field observations, mineralogical, petrological, geochemical, isotopic and melt inclusion data suggest that the Dadnah tonalites formed by partial melting (w10%e15% continuous or w12% batch partial melting), accumulation of plagioclase, fractional crystallization (w55%e57%), and interaction with their host harzburgites. These tonalites were the end result of partial melting and subsequent contamination and mixing of w4% oceanic sediments with w96% oceanic lithosphere from the subducted slab. This MORB-type slab melt composed from w97% recycled oceanic crust and w3% of the overlying mantle. We suggest that a possible protolith for these tonalites was the basaltic lavas from the subducted oceanic slab that melted during the initial stages of the supra-subduction zone (SSZ), which was forming synchronously to the spreading ridge axis. The tonalite melts mildly modified due to low degree of mixing and interaction with the overlying lithospheric mantle. Subsequently, the Dadnah tonalites emplaced at the upper part of the mantle sequence of the Semail ophiolite and are geochemically distinct from the other mantle intrusive felsic granitoids to the south. 2018, China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/ licenses/by-nc-nd/4.0/).
Keywords: Tonalites Felsic granitoids Recycled oceanic crust Source contribution Partial melting Melt inclusions
1. Introduction Different types of felsic intrusives, although volumetrically minor, are common in the oceanic crust and lithospheric mantle, varying from diorite and trondhjemite-to-tonalite and granite (Koepke et al., 2007; rock types generally called plagiogranites; Coleman and Peterman, 1975) but are mostly concentrated at the gabbro-sheeted dykes boundary (Lippard et al., 1986; Amri et al., 1996), reflecting either extreme fractionation in shallow melt
* Corresponding author. E-mail address: [email protected] (S. Kokkalas). Peer-review under responsibility of China University of Geosciences (Beijing).
lenses or re-melting of hydrothermally altered mafic crustal rocks (Amri et al., 1996; Brophy, 2008). Their variable composition, abundance across the oceanic lithosphere and geochemical characteristics make their formation and source still poorly understood (Lippard et al., 1986; Briqueu et al., 1991; Amri et al., 1996; Rollinson, 2009). Such felsic intrusions yield important insights into complex processes, such as the formation of earth’s continental crust (Rollinson, 2009), evolution of oceanic crust by magma intrusion, recycling of crustal components (Liu et al., 2015; Spencer et al., 2017; Xu et al., 2018) or oceanic crust, with or without minor amounts of continental input (Hofmann and White, 1982) and origin of MORB enrichment (Hemond et al., 2006). Crustal component may be incorporated into basalts by either shallow contamination or deep mixing (>120 km; Liu et al., 2015) carried by
https://doi.org/10.1016/j.gsf.2018.09.006 1674-9871/ 2018, China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the CC BY-NCND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
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subducting crust into the asthenospheric mantle, producing both soluble and insoluble materials at the slab-mantle interface (Spencer et al., 2017; Xu et al., 2018). However, the percentage of crustal component in these felsic melts is highly variable (Cox et al., 1999; Haase et al., 2015; Rollinson, 2015, and this study). The plagiogranites that intrude the Semail ophiolite can be categorized in three main types (Rollinson, 2009) based on their location in the ophiolite section: (i) early crustal plagiogranites that associated with partial melting in the upper gabbros (Lippard et al., 1986), (ii) late crustal plagiogranites and minor gabbro emplaced in the sheeted dyke and/or pillow lavas of the ophiolites (Lippard et al., 1986), and (iii) plagiogranites intrusive in the mantle section (Briqueu et al., 1991; Peters and Kamber, 1994; Cox et al., 1999). Alternatively, Spencer et al. (2017) used a geochemical distinction scheme into three suites of granitoids: a metaluminous crustal plagiogranite suite that ranges from granodiorite to tonalite (SiO2 of 54e72 wt.% and low K2O) and two other mantle intrusive suites, one metaluminous that ranges from granodiorite to tonalite (SiO2 of 67e75 wt.%), while the other is peraluminous with higher SiO2 values (72e78 wt.%). More specifically, in our broader study area of the Khor Fakkan block (Fig. 1), the plagiogranites commonly occur in the cumulate gabbros of the crustal sequence (Peters and Kamber, 1994) and in the harzburgites of the mantle sequence (Alleman and Peters, 1972). Previous studies on these plagiogranites have focused on the fact that potassic granitoids are relatively more frequent than in other parts of the Semail ophiolite (Lippard et al., 1986; Python and Ceuleneer, 2003). These granitoids are compositionally distinct from the typical oceanic plagiogranites, which have an extremely low K content (e.g., Coleman and Peterman, 1975). In contrast, the felsic granitoids in the Khor Fakkan block are quite potassic and are enriched in REE’s (Lippard et al., 1986; Cox et al., 1999; Amri et al., 2007). The genesis of plagiogranites is still poorly understood due to their numerous petrological and geochemical constraints. To explain their genesis, several models have been proposed: (1) magmatic differentiation, including fractional crystallization of primitive basaltic magma (e.g. Coleman and Peterman, 1975; Pallister and Hopson, 1981; Pallister and Knight, 1981; Lippard et al., 1986; Peters and Kamber, 1994) and silicate liquid immiscibility of a tholeiitic liquid (e.g. Dixon and Rutherford, 1979), (2) partial melting of sediments, gabbros and basalts (e.g. Malpas, 1979; Amri et al., 1996; Cox et al., 1999; Searle and Cox, 1999; Koepke et al., 2004, 2007; Searle et al., 2015) or pure sediment melting for the peraluminous type intrusive in mantle (Spencer et al., 2017) and (3) mixture of melted crust and sediments (Cox et al., 1999; Rollinson, 2009, 2015; Haase et al., 2015). Cox et al. (1999) and Amri et al. (2007) mainly discussed the formation of the felsic granitoids in the Khor Fakkan block. Based on whole-rock geochemistry and Rb-Sr and Sm-Nd isotopic compositions, Cox et al. (1999) proposed that the monzogranite-toleucogranite dikes were formed due to partial melting of a sedimentary protolith (i.e. Haybi and Hawasina complex). However, the dikes emplaced at higher structural levels (i.e., Dadnah area) are compositionally distinctive from the felsic granitoids of other localities in the Khor Fakkan block, as they thought to be the products of partial melting of both hybrid metasedimentary and oceanic crustal components (Cox et al., 1999). Experimental data on potassic granitoids (e.g., tonalites and trondhjemites) of Amri et al. (2007) suggest that there is a genetic relationship between them and a sedimentary protolith. The granitoids are metaluminous and are likely to be associated with both crustal and mantle components. Amri et al. (2007) concluded that these granitoids were derived by the partial melting of depleted mantle and then were metasomatized due to low-T hydrothermal alteration. The variation
in their composition (i.e., Dadnah felsic granitoids) are thought to be related to local differences in the physicochemical conditions during their formation. Later, more detailed work was done by Rollinson (2009, 2015) in which he set the broader context related to the felsic granitoids of the Semail ophiolite, sampling from six different locations, e.g., the Wadi Hemli, Wadi Hajr, Wadi Zikt, Wadi Fizh, Al-Bithnah and AlDadnah, and using data from Amri et al. (1996) about Maqsad area. Rollinson (2015) compared the major element trends between the crustal and mantle felsic granitoids in Oman. In contrast to Cox et al. (1999), these felsic granitoids are scarcely potassic, appear relatively enriched in MgO, Cr and Ni and Th, U, Ta, La and Ce than their crustal homologs and have REE patterns similar to the host gabbros of the crustal sequence (Rollinson, 2009, 2015). Rollinson (2009) suggested that these felsic granitoids were the products of mixing, contamination of crustal and mafic components. Our study focused on the felsic granitoids that intrude the mantle section in the Dadnah area of the Khor Fakkan block, where high degree of fractionation have been reported in previous studies (Peters and Kamber, 1994; Cox et al., 1999). We studied a smaller area with more detail comprising new petrographic, geochemical, isotope data, and geo-thermobarometry in order to identify, based on our first reported melt inclusion analysis, the least fractionated samples that can shed some more light on the formation and source contribution for this distinct granitoid melts. We discuss the analytical results and complement them with all recent work done in the Semail ophiolite (for further details see Rollinson, 2015 and references therein), in order to propose a tectono-petrogenetic model that could explain better the varied petrological and geochemical characteristics of the north structural block of the Semail ophiolite. 2. Geological setting 2.1. The Khor Fakkan block The eastern part of the Arabian platform is made up of the UAEOman ophiolite complex also known as Semail ophiolite, which is the largest and most widely studied ophiolite complex worldwide (Fig. 1). The Semail ophiolite, which is part of the Cretaceous “Alpine ophiolite chain”, forms a well exposed allochthonous arcuate thrust sheet, w500 km in length, 100 km wide and 10e15 km thick, parallel to the Gulf of Oman (Glennie et al., 1974; Lippard et al., 1986). The first reported ages for the plutonic rocks in the ophiolite came from Tilton et al. (1981), from zircons in the plagiogranites and the dates span in the interval between 95.9 Ma and 93.5 Ma, with most of them clustering around 95 Ma. This age range has been widely accepted as the mean age of the ophiolite. Recently, Rioux et al. (2016) reported higher precision UePb dates, suggesting that the main portion of the ophiolite crust formed between 96.12 Ma and 95.50 Ma. The Semail ophiolite thrusted towards SW over the Neo-Tethyan Hawasina basin for more than 400e500 km and finally was emplaced onto the PermianeMesozoic passive continental margin of the Arabian plate between 84 Ma and 72 Ma (Glennie et al., 1974; Tilton et al., 1981; Tippit et al., 1981; Cox et al., 1999). Petrologic and structural evidence suggests that the initial phase of the Semail ophiolite formed in an extensional setting, however, there is still debate concerning whether it was mainly formed at a mid-oceanic ridge (e.g. Coleman, 1981; Boudier et al., 1988; Ernewein et al., 1988; Nicolas, 1989; Pflumio, 1991) or at a supra-subduction zone (e.g. Searle and Malpas, 1980; Pearce et al., 1981; Alabaster et al., 1982; Searle and Cox, 1999, 2002). The UAE part of the Semail ophiolite comprises two intact ophiolite blocks, the Aswad and the Khor Fakkan block (Fig. 1). The
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1990; Searle and Ali, 2009). These were overthrust onto the Arabian passive margin, which comprises a w4 km thick carbonate sequence (Hajar Supergroup) that was deposited from midPermian to Cenomanian (Glennie et al., 1974). Recent studies have documented a more complex magmatic history for the Semail ophiolite than the standard layered Penrose ophiolite succession (Anonymous, 1972), involving younger, ultramafic-to-felsic island-arc type intrusive rocks (known as phase 2) that cross-cut the earlier magmatic rocks of MORB-composition, known as phase 1 (section 2.3, Grimes et al., 2013; Goodenough et al., 2014, for more details see also section 2.3; Thomas et al., 2014). In the Khor Fakkan block, the phase 2 magmatic rocks mainly appear at the uppermost part of mantle sequence and are restricted in the eastern part of the block. The phase 2 rocks include gabbros, wehrlites and dunites, which cut both the harzburgites of the mantle sequence and the layered gabbros of the crustal sequence and are cut by later sills, dikes and pods of quartz-diorite and tonalities (Goodenough et al., 2014). 2.2. Metamorphic sole and formation of felsic granitoids
Figure 1. Simplified geological map of the UAE ophiolite showing the main tectonic structures and the distribution of mantle and crustal rocks (modified from Gnos and Nicolas, 1996; Thomas et al., 2014). Our study area is shown with a red rectangle (see Fig. 2 for details), and Rollinson’s (2015) sampling area with a cyan colored blue rectangle (R).
Khor Fakkan block occupies the northernmost part of the Semail ophiolite, covering an area of approximately 750 km2, and mainly exposes the mantle sequence of the ophiolite. The large-scale NEtrending Wadi Sidr and NW-trending Wadi Ham fault zones bound the north and south margin of this block, respectively (Fig. 1). The southeast flank is juxtaposed against the Masafi-Ismah and Bani Hamid metamorphic thrust sheets that represent the metamorphic sole of the Semail ophiolite (Fig. 1; Searle, 1980; Goodenough et al., 2006; Searle et al., 2015). The Khor Fakkan ophiolite block forms the structurally uppermost thrust sheet of the nappe stack. Below the mantle sequence and the metamorphic sole, several thrust nappes were stacked during the obduction process. These allochthonous units include Neo-Tethyan oceanic sediments (Hawasina Complex), deep-ocean and subduction-related sediments, volcanics and mélange (Haybi Complex) and proximal carbonate slope-facies rocks (Sumeini complex) (Glennie et al., 1974; Lippard et al., 1986; Bechennec et al.,
The Semail metamorphic sole is a thin (10 m to w500 m thick) amphibolite to granulite sliver of oceanic crust with variable proportions of basaltic and sedimentary material that was strained and metamorphosed, by heat transfer from the hot upper plate mantle toward the slab, during the initial stage of intra-oceanic subduction (Soret et al., 2017). The protoliths of metamorphic sole were the TriassiceJurassic and Early Cretaceous Haybi volcanics, limestones, Mn-rich chert, melange and other deep marine sediments of Hawasina complex (Ghent and Stout, 1981; Searle and Cox, 2002). The uppermost part of the sole, directly below a thick mylonitic peridotite comprises high-temperature (HT) amphibolite-togranulite facies metabasalt, while below that there is the low temperature (LT) sole which is mainly composed of greenschist facies metachert with imbrications of metabasalts and metasedimentary rocks (Searle and Malpas, 1980; Ghent and Stout, 1981; Boudier et al., 1988; Soret et al., 2017). The estimated P-T for the HT sole are on the range of 700e900 C and 0.4e1.3 GPa and decreases in the LT sole to w500 C and 0.45e0.55 GPa with an inverted thermal gradient, grading steeply from granulite to greenschist facies (>1000 C/km, Searle and Malpas, 1982; Hacker and Mosenfelder, 1996; Searle and Cox, 2002). The age of the metamorphic sole is considered almost coeval or slightly postdates the time of the Semail ophiolite obduction (Soret et al., 2017). Cox et al. (1999) assumed that HP metamorphism was roughly contemporaneous to the formation of the metamorphic sole throughout the Semail ophiolite, based on contiguous metamorphism ages (i.e., 96 2 Ma, Ar/Ar of As Sifah eclogite age, 93.5 0.1 Ma, Ar/Ar sole metamorphism mean age of Wadi Tayin area of Cox et al. (1999)). However, the published ages of the sole were estimated by the cooling ages of minerals (e.g., hornblende, biotite, and garnet) and display various age ranges. Indeed, the K/Ar dating on a metamorphic hornblende separated from the amphibolite yielded ages in the broad range 101e89 Ma and cluster near 98 Ma (Gnos and Peters, 1993). Recent data from sole peak metamorphism (Rioux et al., 2016) reported dates between 96.16 Ma and 94.82 Ma, suggesting that metamorphic rocks were formed either prior to or during the formation of the oceanic crust. 2.3. Magmatism in Semail ophiolite In the Semail ophiolite two distinct phases of magmatism have been described. An earlier sea-floor spreading magmatism termed as phase 1, composed of gabbros, sheeted dykes and pillow lavas and a later phase 2 more hydrous magma suite, composed of
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wehrlites, clino-pyroxenites and a complex network of gabbroic, diorite and tonalite intrusions, recording the initial stages of arc magmatism (Styles et al., 2006; Goodenough et al., 2014). The different magmatic phases are marked by the lavas and the relevant magmatic rocks found within the crustal and mantle sequence. Three major lavas were distinguished based on their geochemical affinities: (1) V1 lavas (i.e., Geotimes) have MORB-like affinity and are formed at spreading axis, (2) V2 lavas (i.e., Lasail or Alley) are formed off-axis and have arc-like petrochemical signatures (Alabaster et al., 1982; Ernewein et al., 1988) and (3) V3 lavas (i.e., Salahi) are enriched in incompatible elements and range compositionally from alkaline to transitional tholeiites. The V3 lavas formed by off-axis magmatism and postdate the other two lavas (Alabaster et al., 1982; Lippard et al., 1986). Furthermore, due to the fact that V2 and V3 lavas are more abundant in the Khor Fakkan block i.e., phase 2 magmatism, the V1 lavas are relatively rare (Goodenough et al., 2014; Thomas and Ellison, 2014). The volcanic rocks related to the phase 2 magmas are characterized by higher amount of plagioclase and clinopyroxene, with variable amounts of olivine, orthopyroxene, hornblende and quartz and are considered to form by melting of a hydrous magma in a SSZ setting (Goodenough et al., 2014). Haase et al. (2016) have grouped the plutonic rocks from the crustal sequence of Oman ophiolite into two groups, with each one mimic the geochemical signatures of the V1 and V2 lavas, respectively. V2 magmas are more voluminous in the northern part of the ophiolite showing the stronger input of slab component, regardless if that was sedimentary or not. 3. Sampling and analytical methods 3.1. Field observations and sampling The study area, which is located between the Dadnah town and east of Wadi Zikt, is structurally comprising the uppermost part of the mantle sequence and the Moho transition zone (MTZ) of the Khor Fakan block of the Semail ophiolite (Fig. 2).
The felsic granitoids in this area appear as veins, sills, and dikes, extending up to several hundreds of meters laterally and few meters in thickness (Fig. 3aec). Almost all of the examined granitoids are hosted by serpentinized harzburgites. They appear fine-tomedium-grained with typical granitic textures, although a few of them display pegmatitic textures. They usually contain up to 20 cm in length mafic enclaves and xenoliths in their rock body (i.e. location D6, Figs. 2 and 3a). At the same location two different granitoid types can be distinguished. There is a grayish-colored coarser grained, intense foliated and more sheared type with hornblende clinopyroxene that comprises the biggest part of the sill, while in the eastern part there is a composite body together with a less foliated, fine-to medium grained, less deformed leucocratic rock type containing abundant biotite crystals. The leucocratic granitoid intrudes in places the foliated grayish type, suggesting that probably was a later intrusive phase. The majority of the felsic sills and dikes observed in this area are related to pre-existing ductile or reactivated brittle-ductile faults, as their margins in many cases controlled by ductile shear zones or hydrothermally altered striated surfaces. Most of the dikes are oriented subparallel-parallel to the fault/fracture planes (Fig. 2, stereonets). Dikes exhibit various degree of deformation showing development of quartzose mylonites, gneissic fabrics or biotite grain-shape preferred orientations often aligned almost parallel to the dike margins due to intense ductile shearing. Many of the meter-to-hectometre scale in length felsic granitoids are feeder dikes, developing around them cm-scale secondary veins. No clear crosscutting relationships among the meter-to-hectometre scale felsic dikes was observed, while the cm-scale veins frequently crosscut each other, suggesting coeval development for these secondary veins. The felsic dikes in the study area trend between a frequent WNWeESE orientation (95 e105 ) and two less prominent orientations NEeSW (30 e50 ) and NWeSE (300 e310 ) (Fig. 2, stereonet). This latter trend follows the map-scale NW-trending fault/ fracture set (Fig. 2, map) and is frequent among the meter-scale
Figure 2. Aerial view (Google Earth image) of the study area (red rectangle of Fig. 1) at the west side of the Dadnah city in the Khor Fakkan block. The yellow arrows indicate the sampling locations and the red lines show the main map-scale fault trends in the study area. Stereonets of the tonalite dikes and fault planes are shown as great circles on the right side of the figure.
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Figure 3. (a) Field photograph of a tonalite sill that outcrops at locality D6. The sill contains xenoliths of the host harzburgite and is offset by later stage brittle faults. (b) Field photograph of the same tonalite sill, showing in detail the xenoliths contained in the sill. The xenoliths appear partially resorbed and recrystallized with intense mylonitic textures, comprising, olivine, ortho- and clino-pyroxene and plagioclase. (c) Sample from the tonalite sill composed of plagioclase, quartz and hornblende that is cross-cut by late quartz veins. (d) Representative thin section image of tonalite composed of plagioclase (Pl), quartz (Qz), hornblende (Hbl) and clinopyroxene (Cpx). At the left bottom corner the plagioclase crystal contain relic plagioclase (sample PGD2, XPL, 4x). The XPL colors of minerals are not typical due to the fact that thin sections are thicker and were constructed for fluid and melt inclusions analyses. (e) Thin section image of a tonalite medium-to-intense mylonitized (sample PGD7_2, XPL, 2x), with clinopyroxene and plagioclase porphyroblasts rotating in a foliated matrix of quartz, plagioclase, alkali feldspar and mica. (f) Cathodoluminescence (CL) image of the tonalite that contains xenoliths of the host harzburgite. The plagioclase of the host harzburgite (purple Pl) appears with green CL colors and intergrowths with the non-luminescent clinopyroxene (Cpx). The plagioclase of the tonalite (cyan Pl) appears with blue CL colors and replaces the plagioclase from the harzburgite, and very commonly resorbs these plagioclases (sieve texture). (g) Backscattered electron (BSE) image of the tonalite plagioclase. The plagioclases appear with regular zonation with cores composing of andesine (An44e56) and rims of oligoclase (An25e38). At the down-left corner partially resorbed plagioclases of higher An contents (An53e56), are also shown. At the up-right corner the plagioclase composition is shown on the ternary Ab-Or-An plot (Ab: Albite, Ol: Oligoclase, And: Andesine, La: Labradorite, By: Bytownite, An: Anorthite, Ano: Anorthoclase). Solid pink squares symbolize the cores, deep yellow squares the rims, and light ochre squares the relict plagioclase cores (Mineral abbreviations are after Whitney and Evans, 2010).
felsic dikes. The mesoscale fault orientations cluster around a NEeSW trend (w30 e45 ), while the WNWeESE trend (110 e120 ) display sinistral shear motion along the fault surfaces and is the dominant map-scale trend in the area between Wadi Zikt and the village (Fig. 2, map).
3.2. Analytical methods In total, thirteen granitoid samples as well as few representative host harzburgite samples (e.g., OPD2) were collected from seven sites in the study area (Fig. 2, Table 1). From these samples 25 thin
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Table 1 Major and trace element chemistry of the tonalites. Melt inclusions are shown also on table. Samples
PG
OP
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
D1
D2
D2_2
D3_1
D3_2
D4_Mi
D4_Ma
D5
D6_1
D7
D7_2
M1
M9
M12
65.99 13.76 6.12 0.12 2.77 7.05 1.11 0.27 0.91 0.22 2.13 100.45
72.90 13.90 0.73 0.01 0.28 0.21 3.42 6.54 0.02 0.06 0.61 98.68
73.59 12.48 1.64 0.02 2.18 1.52 3.88 0.99 0.11 0.03 2.40 98.84
74.57 13.57 0.82 0.02 0.42 0.80 4.36 2.73 0.06 0.06 1.50 98.91
73.41 13.11 1.87 0.02 1.51 3.15 2.91 0.80 0.17 0.01 2.21 0.02 <0.01 99.37
75.00 12.97 1.15 0.02 0.91 2.81 3.29 1.31 0.11 0.03 1.08 98.68
76.99 12.94 0.98 0.01 0.50 2.94 4.01 0.26 0.09 0.02 0.96 0.01 0.01 99.90
75.92 12.29 1.58 0.04 1.77 2.17 2.10 0.45 0.26 <0.01 3.03 0.01 <0.01 99.72
75.18 13.49 1.38 0.11 0.39 0.88 5.01 1.05 0.01 0.03 1.46 98.99
76.40 13.52 1.09 0.01 0.22 1.34 6.17 0.25 0.03 0.02 0.22 <0.01 0.01 99.32
73.45 15.78 0.01 0.00 0.00 0.00 3.59 4.01 0.02 0.01 2.71 0.02 0.01 99.58
74.98 14.78 1.32 0.00 0.02 0.03 5.47 1.27 0.01 0.01 2.01 0.02 <0.01 100.00
76.45 13.29 0.65 0.00 0.00 0.01 7.84 0.02 0.00 0.01 1.45 0.01 <0.01 99.72
-
28.8 63.7 0.4 2.9
56.0 27.6 0.2 3.8
47.9 44.7 0.4 1.1
57.3 21.2 0.1 6.5
55.7 27.4 0.2 4.8
62.4 14.2 0.1 5.2
64.8 14.2 0.1 7.7
56.9 33.7 0.2 2.6
59.7 32.7 0.1 1.9
<1 18 85 20.17 7.26 371.00 37.00 245.00 12.26 0.10 192.00 1.34 0.85 1.58 4.52 1.03 0.18 10.28 24.61 3.75 18.36 5.34 2.90 3.86 1.41 7.19 1.25 4.66 0.47 3.55 0.46 0.51 3.6 0.28 0.21 0.66 0.89 2.17 5.47 28.33 29.71 64.2 0.02
<1 <1 <5 2.66 4.05 5.00 6.00 9.00 27.87 0.43 9.00 0.78 1.20 2.26 8.66 1.23 3.45 16.66 39.78 4.34 14.41 3.79 1.16 3.34 0.96 4.51 0.70 2.75 0.31 2.48 0.33 1.04 62.7 0.64 0.50 0.81 1.08 1.25 0.00 48.19 8.62 60.3 0.81
<1 1 7 4.31 28.24 152.00 18.00 32.00 11.13 0.61 583.00 0.75 1.17 0.82 5.24 0.49 8.09 65.32 117.16 11.77 38.30 5.45 3.40 7.66 1.38 2.93 0.46 2.68 0.19 1.58 0.23 0.97 6.9 1.74 3.05 1.93 0.97 2.55 0.01 10.42 38.83 84.0 0.19
<1 4 <5 4.33 130.15 12.00 24.00 10.00 21.03 0.34 11.00 0.46 0.97 1.65 6.45 1.61 2.10 6.45 15.98 1.89 6.88 2.43 1.49 2.39 0.69 3.03 0.49 2.06 0.24 2.19 0.27 1.05 26.4 0.39 0.22 0.66 1.05 2.36 0,02 81.96 46.97 67.0 10.85
4 <1 2 12 6.58 20.09 261.00 8.00 59.00 6.67 0.46 142.00 0.72 0.29 0.46 5.78 1.16 6.50 46.08 85.57 9.00 32.56 5.89 3.43 10.69 1.95 2.92 0.47 3.68 0.18 1.79 0.22 0.77 4.2 1.14 3.34 1.88 0.97 2.07 0.02 14.59 22.39 50 76.2 0.08
<1 1 8 7.67 20.10 266.00 8.00 46.00 7.93 0.46 196.00 1.29 1.05 0.92 3.66 1.36 10.70 67.72 121.73 12.52 44.88 7.10 3.31 14.57 2.48 2.39 0.36 4.19 0.13 1.50 0.18 0.76 24.2 1.39 1.90 2.62 0.96 1.61 0.01 13.20 17.04 75.8 0.08
6 <1 <1 <5 3.23 80.40 275.00 9.00 79.00 3.05 0.11 154.00 0.64 0.19 0.85 3.09 1.97 4.47 30.58 58.25 5.93 20.90 3.62 3.18 6.88 1.22 1.02 0.15 1.92 0.06 0.69 0.09 0.74 17.3 1.23 3.26 2.02 0.99 3.10 0.01 11.80 60.53 16.8 66.9 0.29
6 <1 1 8 5.60 15.26 236.00 11.00 124.00 10.84 0.88 173.00 0.69 0.70 0.31 2.26 1.62 0.43 3.75 6.43 0.64 2.31 0.52 2.66 1.00 0.19 0.53 0.10 0.58 0.06 0.57 0.08 1.04 25.2 1.06 0.49 1.07 0.95 7.13 0.27 78.13 109.08 16.7 81.6 0.06
9 5 <5 3.52 81.98 540.00 40.00 28.00 19.30 5.79 23.00 1.27 2.62 2.06 3.01 1.83 5.04 6.58 16.14 1.73 6.24 2.83 1.86 2.85 0.92 4.84 0.78 3.06 0.39 3.56 0.43 1.08 8.2 0.34 0.24 0.68 1.10 2.36 0.00 39.96 51.45 52.8 0.15
4 2 <1 5 4.73 23.84 152.00 28.00 26.00 17.30 1.06 304.00 2.02 1.58 0.81 3.83 0.69 9.15 70.06 120.83 12.01 41.99 6.00 3.94 16.87 2.74 1.70 0.24 4.31 0.09 1.15 0.14 0.89 26.2 1.70 4.52 1.92 0.95 1.94 0.00 7.16 118.93 22.5 44.4 0.16
42.4 51.7 0.1 2.6 0.77 -
55.8 37.9 0.1 2.7 0.71 -
60.1 36.3 0.1 1.1 0.72 -
-
55.7 -
-
wt.% 66.92 SiO2 Al2O3 19.43 Fe2O3t 0.62 MnO 0.01 MgO 0.56 CaO 1.16 Na2O 8.93 K2O 1.60 TiO2 0.01 P2O5 0.05 1.06 LOI/H2O Cl F Total 100.35 C.I.P.W. normative Qz 36.4 An 57.8 Or 0.3 Ab 0.5 ppm B Be <1 Sc 3 V <5 Ge 1.32 Rb 39.91 Sr 187.00 Y 24.00 Zr 36.00 Nb 29.85 Cs 0.22 Ba 282.00 Hf 1.36 Ta 1.26 W 1.19 Pb 2.79 U 1.45 Th 2.17 La 17.96 Ce 36.05 Pr 3.66 Nd 11.67 Sm 2.81 Eu 2.15 Gd 2.67 Tb 0.73 Dy 3.59 Ho 0.59 Er 2.35 Tm 0.26 Yb 2.05 Lu 0.28 ASI 0.95 LREE/HREE 7.8 (La/Sm)N 0.93 (La/Yb)N 0.65 (Gd/Yb)N 0.79 Ce/Ce* 1.02 Eu/Eu* 3.08 Ti/Ti* 0.01 Pb/Pb* 29.14 Y/Ho 40.41 Cl/B Mg# 78.2 Rb/Sr 0.21
PG ¼ Tonalite; OP ¼ Host harzburgite; M(1,9,12) ¼ Melt inclusion glass.
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sections, double polished and blue epoxy impregnated were constructed (50 mm thick, 24 mm 46 mm, Wagner Petrographic Lab, USA). An Olympus BX51 petrographic microscope was used for examining the thin sections. We should note that these thin sections were constructed primarily for fluid and melt inclusion analyses and due to that they do not display the typical XL mineral colors, as shown in Fig. 3d and e. The same samples were prepared by removing exterior surfaces with a rock splitter. Firstly, we separated the xenoliths of the host harzburgite from the felsic granitoids. Mineral separation from the granitoids involved handpicking after examination under a binocular microscope to ensure purity of better than 98%, based on their paragenetic relations and textural equilibrium. In this way, only contiguous, non-tectonized and unaltered crystals of pyroxene, plagioclase, hornblende, biotite, quartz and calcite were selected. In order to ensure primary isotopic signature, before the isotopic analysis the samples were heated between 100 C and 200 C to remove the involvement of secondary fluid inclusions. Mineral compositions were determined using a JEOL 8900 Superprobe equipped with wavelength, energy dispersive and back-scattered detectors and a xClent system for ppm-level resolution at the Microprobe Center of the Department of Earth and Planetary Sciences Department, at McGill University. Operating conditions included an acceleration voltage of 15 kV, a beam current of 10 nA, and a counting time of 20 s. Scans in WDS mode were also used to detect trace element contents. Standards used were natural olivine, clino- and ortho-pyroxene, amphibole, chlorite, epidote, K-feldspar, plagioclase, barite, muscovite, magnetite, ilmenite, pyrite, chalcopyrite and galena, plus the native metals Ag, Sb, Au, Se, Ta and Cd. ZAF corrections were made with proprietary JEOL software. A minimum of ten analyses were obtained from each sample and three from each grain. Scanning Electron Microscopy (SEM) was also used in order to determine the chemical composition of certain minerals of the felsic granitoids. These were carried out at the Laboratory of Electron Microscopy and Microanalysis, University of Patras (Greece), using a Jeol JSM-6300 Scanning Electron Microscope, equipped with EDS and WDS and a THETA software. The operating conditions were 15 kV accelerating voltage and 3.3 nA beam current with 4 mm diameter beam. Also, at the same laboratory X-ray power diffraction patterns of oriented granitoid samples were obtained using a Bruker D-8 Focus diffractometer, with Ni-filtered Cu Ka radiation (with a 2q angle of 3 e70 ), using the XRD-data method of Moore and Reynolds (1997). The samples already used for probe analyses, prior to the microthermometric study of the melt inclusions were examined by cathodoluminescence (CL) using the Luniscope CL Reliotron spectrometer (Laboratory of Electron Microscopy and Microanalysis, University of Patras, Greece), equipped on a Quanta 200F ESEM with 45-s scanning time. The applied acceleration voltage and current were 15 kV and 120 nA and the CL images were collected with a resolution of 1280 960 pixels and 256 gray level. Major elements from the least altered felsic granitoid samples, as their Ishiwaka Alteration Index is 8.5%, were detected by inductively coupled plasma (ICP), using a lithium tetraborate fusion, at the Actlabs laboratory in Ontario, Canada. The concentrations of trace elements from the felsic granitoid samples were obtained by Agilent 7700s inductively coupled plasma mass spectrometry (ICP-MS), at Chonnam National University, South Korea. JG-2 gabbro and JGb-2 granite, obtained from the Geological Survey of Japan were used as standards. Repeated measurements of the data are between 0.05% and 1.4% relative for major elements, 7.1% relative for TiO2, 0.06% and 0.1% relative for REE. To quantify the Ce, Eu, Ti, Pb anomalies, we have used the equations of Taylor and McLennan (1985), Bau et al. (1996), Peters and Day (2014) and
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the normalizing values of McDonough and Sun (1995). The results of whole-rock analyses are presented in Table 1. Furthermore, we have microthermometrically analyzed only primary melt inclusions on one doubly polished 50e100 mm thick section, since primary inclusions were infrequent and showed signs of metastability in the other sections. The melt inclusions in the granitoid sample were completely crystallized in the quartz and plagioclase grains, and so they were homogenized using the conventional horizontal cold-seal pressure vessel technique. Hightemperature measurements were obtained with a MDSG600 heating-freezing stage, at the USGS Denver Inclusion Analysis Laboratory. Heating-freezing rates range from 0.1 to 130 C/min. The stage was calibrated at the critical temperature of H2O (374.1 C), the melting points of ultrapure NaCl (801 C), and silver (961 C). The accuracy and precision of homogenization temperatures of melt inclusions are better than 5 C. From room temperature to 400 C, a heating rate of 30 C/min was used. The rate was slowed to 10 C/min from 400 C to 800 C and then held at that temperature for thirty minutes before proceeding with homogenization runs. Melt inclusion glasses were analyzed using the EPMA difference technique (Table 1), with a Cameca SX50 and an SX100 electron microprobe at the same laboratory equipped with both WDS and EDS. Quantitative analyses were conducted using WDS with silicate, oxide, and phosphate standards. The data were corrected according to PAP methodology (Pichou and Pouchoir, 1985) using a vendor supplied software. Isotopic compositions of oxygen, hydrogen, carbon and silicon analyzed using a MAT-253 stable isotope ratio mass spectrometer. Analyses were performed at the Chinese Academy of Geological Sciences (CAGS), Beijing, China. Samples used for isotopic analyses were cross checked under the CL microscope to avoid any possible secondary isotopic signature due to alteration. Oxygen and hydrogen were released from pyroxene, plagioclase, hornblende, biotite, quartz and calcite using the BrF5 extraction technique (Clayton and Mayeda, 1963; Friedman and O’Neil, 1977). Silicon was measured using the SiF4 technique of Ding (2004), whereas carbon and oxygen in calcite were liberated as CO2 after Clayton et al. (1972). The isotopic ratios are reported in standard d notation per mil relative to SMOW for oxygen and hydrogen, NBS-28 for silicon and PDB belemnite for carbon. Analytical precision was better than 0.2& for d18O and d13C, 2& for dD and 0.1& for d30Si. We have used the AlphaDelta software of Beaudoin and Therrien (2009) to compute the isotopic fractionation factors and equilibrium temperatures. The same minerals were analyzed for their Rb/Sr and Sm/Nd isotopic compositions. The method used for Sr and Nd isotopic analysis is described by Papanastassiou et al. (1977) and De Paolo (1980). Samples were analyzed using a VG-354 ionization mass spectrometer at the Modern Analysis Center, Nanjing University, Nanjing, China. The total procedure blanks for Rb and Sr were 20 pg and 50 pg and for Nd and Sm 1 ng and 0.2 ng, respectively. Analytical precision for the 87Sr/86Sr and 143Nd/144Nd ratios was better than 0.00001 (all errors are reported at 2s absolute). The standards used for 87Sr/86Sr, 146Nd/144Nd and 143Nd/144Nd ratios were NBS987 (0.710223 8) and JNdi-1 (0.7129 5 and 0.51112 5, Tanaka et al., 2000). The normalizing factors used to correct for isotopic fractionation of Sr and Nd were 86 Sr/88Sr ¼ 0.1194 and 146Nd/144Nd ¼ 0.7219, respectively. Isotopic data for Nd and Sr are presented in terms of the epsilon notation, i.e., εSr and εNd (De Paolo, 1980). All ages were calculated using the decay constants for 87Rb and 47Sm which are 1.42 1011 and 6.54 1012/yr, respectively. Ages were calculated using the leastsquare regression technique. The Rb-Sr and Sm-Nd isochron ages were calculated using the York’s model 3 fit in Isoplot (Ludwig, 2001). We have used 2s errors of 0.5% for the 147Sm/144Nd
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ratios, 0.0004% for the 143Nd/144Nd ratios, 0.055% for the 87 Rb/86Sr and 0.0055% for the 87Sr/86Sr ratios. 4. Petrography and mineral chemistry 4.1. Petrography The felsic granitoids based on crystal size are fine and mediumgrained with hypidiomorphic-to-allotriomorphic primary granular textures and preserve deformational textures, including granophyric or mylonitic (Fig. 3d and e). In addition, sutured grain boundaries in most minerals, undulose grain extinction and wider grains of quartz are common because of ductile deformation. The felsic minerals of the granitoids are composed of plagioclase (70e60 vol.%), quartz (20e15 vol.%) and K-feldspar (5 vol.%). The mafic minerals (5e15 vol.%) are mostly hornblende, clinopyroxene and biotite. Epidote, titanite, titanomagnetite and ilmenite (5e7 vol.%) occur as accessory phases. In few samples, hydrothermal alteration resulted in the formation of chlorite, muscovite, serpentine and syntaxial veinlets comprising quartz, calcite, siderite, muscovite, and chlorite (Fig. 3d). A distinctive feature of the PGD1 to PGD4 samples is that they contain xenoliths of their host harzburgite, which in places comprise up to 50 vol.% of the felsic granitoids. These partially resorbed and recrystallized xenoliths with intense mylonitic textures, comprise ortho- and clino-pyroxene porphyroblasts (60e70 vol.%) and plagioclase, with accessory minerals magnetite and chromite (Fig. 3a, b and f). Perthite clino- from ortho-pyroxene wispy exsolution textures and plagioclase mottled by clinopyroxene occur in the xenoliths contained in the sample PGD3. The exsolution lamellae are planar, and their orientation is crystallographically controlled along the (110) cleavage of orthopyroxene. Corona textures with cores of olivine mantled by rims of clinopyroxene are common. Sometimes clinopyroxene, in turn, is rimmed by hornblende. In the PGD5 to PGD7 samples, the recrystallized pyroxene-bearing xenoliths are surrounded by finegrained matrix comprising of quartz, hornblende biotite and muscovite. The wall-rocks are harzburgites, comprising relict olivine, othroand clino-pyroxene, and a hydrothermal assemblage that includes talc, serpentine, chlorite, muscovite and minor quartz. Serpentinization is common among these rocks. Magnetite and chromite occur as accessory phases. Olivine and ortho- and clino-pyroxene initially alter towards biotite, which in turn alters to Fe-rich chlorite and muscovite. 4.2. Mineral chemistry The plagioclases from the felsic granitoids are interstitial, subhedral to anhedral composing of either fine-grained matrix or coarse-grained squat crystals. They appear with regular zonation with cores composing of andesine (An44e56) and rims of oligoclase (An25e38) (Fig. 3g), where the cores are enriched in the anorthite component by w15%e20%. The K-feldspars are uncommon and always appear with allotriomorphic, medium-grained crystals intergrown with plagioclases, forming intergranular textures. Their chemistry is persistent and characterizes them as orthoclase (Or100). In rare cases, relict plagioclases of higher An contents (Añ53e56, Fig. 3f and g) were observed under CL microscopy, characterized by resorption texture. The mafic minerals from the felsic granitoids analyzed are pyroxene, hornblende and biotite (e.g., sample PGD6_1). Clinopyroxenes appear with the same crystallographic habit as plagioclases, either as a fine-grained matrix or with coarse-grained crystals (Fig. 3d). The clinopyroxenes analyzed are classified as diopside-to-
augite (En59.8e61.8 and Wo27.7e28.2). Their Mg# (Mg# ¼ 100 Mg/ (Mg þ Fe2þþFe3þþMn)) values range from 0.84 to 0.855, and are relatively enriched in Al2O3 (up to 0.11 apfu; based on 6 oxygens) and depleted in TiO2, Cr2O3 and NiO (0.01 apfu). Hornblende form hypidiomorphic crystals showing intensive pleochroism in shades of green. Some crystals appear with light colors or even colorless and intergrown with plagioclases irregularly distributed due to intensive tectonism (Fig. 3d). Hornblende is classified as tschermatitic-hornblende towards tschermakite, with Mg# values ranging from 0.73 to 0.81. Biotite appears with the same textural characteristics to hornblende and based on its Sitotal and Mg# values (w0.45) is categorized as biotite. Magnetite occurs as medium-grained accessory phase, is rimmed by plagioclase and/or clinopyroxene and on the basis of its Fe, Cr, and Ti composition is almost pure magnetite. Mixtures of fine-grained serpentine, muscovite, chlorite and minor albite and calcite are the main alteration products of the felsic granitoid minerals. Based on their Si and Al content, the analyzed muscovites are rich-muscovites-to-low-phengites. Chlorite appears fine-grained and fibrous. Chlorite is clinochlore due to its anomalous polarization colors and its Al, Si and Fe content. Serpentine is classified as solid solution of w25% sepiolite and w75% serpentine (e.g., sample PGD2_2). Olivine from the xenoliths occurs as a fine-grained matrix as well as with medium-grained crystals that are highly altered towards serpentine. The analyzed olivines are classified as forsterite (Fo86.7e88.3Fa11.7e13.3) and their Mg# ratios range from 0.867 to 0.883. Orthopyroxenes are classified as enstatite (En93.5e96.8 and Fs3.2e6.5), with Mg# ratios ranging from 0.935 to 0.968, and the clinopyroxenes as augite (En72.3e75.6 and Fs10.5e11.4). The ortho- and clino-pyroxenes are enriched in Al2O3, and depleted in NiO. Spinel contained in the host fragments occur as medium-grained accessory phase with anhedral crystals and many of them appear in clusters of ortho- or clino-pyroxene. The spinel is categorized as Al Mg-chromite. 4.3. Geothermobarometry and physicochemical conditions of alteration The compositions of the analyzed minerals were used for geothermobarometry. For the temperature and pressure estimations, seven different geothermobarometers were used (Table 2). The two-feldspar geothermometer is based on the distribution of the albite component between plagioclase and alkali feldspar crystals during magma solidification (Stormer, 1975; Putirka, 2008). The temperature obtained from the application on the two-feldspar geothermometer of Stormer (1975) in our granitoids is 681 C. The Ti-in-biotite geothermometer of Henry et al. (2005) is based on the Ti-saturation of near-isobaric natural biotite equilibrated at 4e6 kbar. Based on the Ti-in-biotite geothermometer of Henry et al. (2005), the calculated temperature is from 553 C to 583 C. Pressure estimates for the granitoid samples were made using the clinopyroxene barometer. According to Ashchepkov et al. (2001) pressure is the main variable that controls the chemical behavior of clinopyroxene in the magmatic environment. The exchange of jadeite-diopside components on clinopyroxene allows the calibration of the clinopyroxene thermobarometer (Ashchepkov, 2002, 2003). Based on the clinopyroxene barometer the estimated pressure ranged from 5.7 kbar to 6.5 kbar. In addition, the Al-in-hornblende geobarometer of Hammarstrom and Zen (1986) was employed. In the magmatic hornblendes the Altotal content correlates linearly with crystallization pressure of the intrusion. The Al-in-hornblende geobarometer was re-evaluated, amongst others, by Johnson and Rutherford (1989) which emphasized that the Al-in-hornblende geobarometer should be applied
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Table 2 Summary of the geothermo-barometrical results and melt inclusion analyses. Geothermometer-geobarometer
T ( C)
P (kbar)
logfO2
Reference
Two-feldspar Magnetite-ilmenite Clinopyroxene Al-in-hornblende Hornblende-plagioclase Ti-in-biotite Chlorite Melt inclusions Host Homogenization temperature (Th) Water content Trapping pressure (Pt)
682.1 677e682 620e626 553e583 222e251
5.7e6.5 5.0e6.0 5.0e6.0 -
21 to e20.9 -
Stormer (1975) Buddington and Lindsley (1964) Ashchepkov (2002, 2003) Johnson and Rutherford (1989) Blundy and Holland (1990) Henry et al. (2005) Cathelineau (1988) and Kranidiotis and MacLean (1987)
Quartz 635e650 C 1.5e2.1 wt.% H2O 4.38e4.41 kbar
for the assemblage hornblende þ biotite þ plagioclase þ Kfeldspar þ quartz þ titanite þ Fe-Ti oxides. Based on this geobarometer, the estimated pressure ranged from 5.0 kbar to 6.0 kbar. The Al-in-hornblende geobarometer was also combined with the hornblende-plagioclase geothermometer after Blundy and Holland (1990). This geothermometer also uses the Altotal values in hornblende as well as the An and Ab components of the plagioclase that co-exists with it. Based on geothermometer of Blundy and Holland (1990), the obtained temperatures are 620e626 C. Buddington and Lindsley (1964) proposed the use of the compositions of coexisting ilmenite-hematite and magnetiteulvospinel solid solutions in the system FeO-Fe2O3-TiO2 as a geothermometer and oxygen barometer (i.e., magnetite-ilmenite geothermo-oxygen-barometer) which has been re-evaluated by Powell and Powell (1977) and Stormer (1983). The temperature obtained based on the application on the magnetiteilmenite geothermo-oxygen-barometer of Buddington and Lindsley (1964) (calculated with ILMAT excel, Lepage, 2003) for this co-existence is estimated to range from 677 C to 682 C. For these temperatures the logfO2 values are w e21 or DlogfO2 (HM) ¼ e8.5 at w 680 C, suggesting a highly oxidized melt. The used geothermobarometers and the crystallization temperatures and pressures of the felsic granitoid minerals are summarized in Table 2. The non-stoichiometric behavior of chlorite makes it a geothermometer, as chlorite composition (i.e., AlIV content of chlorite) unequivocally relates to temperature of formation (Cathelineau and Nieva, 1985; Kranidiotis and MacLean, 1987; Cathelineau, 1988; Jowett, 1991). Based on the chlorite geothermometer the calculated temperatures that represent the temperatures of hydrothermal alteration range from 222 to 251 C. The chemical conditions of the highly serpentinized sample PGD2_2 were estimated from phase stability relationships using SUPCRT92 (Johnson et al., 1992) with thermodynamic properties from the 2007 database (slop07.dat; Shock and Helgeson, 1988). These estimations are based on the coexistence among silicates for NaCl-saturated aqueous liquids. Temperatures of 230 C (averaged T obtained from chlorite geothermometer) and pressure of 1 kbar are assumed. All solid phases considered to have ideal behavior. Individual ion activity coefficients of dissolved species were calculated using the B-gamma extension of Helgeson et al. (1981).
Plagioclase 664e733 C 2.1e2.7 wt.% H2O 4.44e4.59 kbar
The ionic strength calculated values of the serpentinized solution are I ¼ 0.3 0.01. From the equilibria at 230 C, the obtained ionic activity values occurred during the alteration of the plagiogranites are listed in Table 3. Fluid-wall rock interaction played a major role in the alteration of the granitoids. Alteration involved hydration of the mafic minerals (Table 3, reaction 1) and ion exchange between the fluid phase and these rocks (Table 3, reactions 2e4). The most significant parameter during alteration was pH neutralization. The pH increase modified the stability fields of the alteration minerals, and subsequently drove reactions (reactions 1e4, Table 3) to the right leading to their precipitation. 5. Melt inclusions 5.1. Petrography of melt inclusions Melt inclusions in the Wadi Zikt felsic granitoids occur as isolated inclusions or in assemblages defining crystal growth zones within plagioclase and quartz crystals. Melt inclusions that occur in the same growth zone suggest synchronous crystallization at the time the melt was being trapped (Fig. 4b) and provides the necessary evidence that the melt was in equilibrium with the host plagioclases and quartz. Melt inclusions vary in sizes from 10 mm to 30 mm. The analyzed melt inclusions are composed of homogeneous glass with or without bubble or daughter minerals (Fig. 4a,b, respectively). The shapes of melt inclusions are rounded, negative or elongate (Fig. 4c). The melt inclusions are assumed to be primary as they are trapped along crystal growth faces or between growth zones with no visible fractures, following the nomenclature of Roedder (1984). 5.2. Microthermometry of melt inclusions Microthermometric measurements included the complete melting temperature (Tm) of all silicate phases (except host) within the melt inclusion and the homogenization temperatures (Th). For undersaturated melts, the Th constrains the minimum trapping temperature (Tt). The pressure inside an inclusion (internal trapping pressure Pt) can be estimated using the measured
Table 3 Summary of the alteration chemistry.
1. 2. 3. 4.
Reaction
Minerals
Physicochemical conditions (T ¼ 230 C)
2Mg2Si2O6(s) þ 2Hþ(aq) þ H2O(l) ¼ Mg3Si2O5(OH)4(s) þ Mg2þ(aq) þ 2SiO2(s) 3KMg3AlSi3O10(OH)2(s) þ 20Hþ(aq) ¼ KAl3Si3O10(OH)2(s) þ 2Kþ(aq) þ 9 Mg2þ(aq) þ 6SiO2(aq) þ 12H2O(g) CaAl2Si2O8(s) þ 5Mg2þ(aq) þ 8H2O(l) þ SiO2(aq) ¼ Mg5Al2Si13O10(OH)8(s) þ Ca2þ(aq) þ 8Hþ(aq) CaAl2Si2O8(s) þ 2Naþ(aq) þ 4SiO2(aq) ¼ 2NaAlSi3O8(s) þ Ca2þ(aq)
Enstatite-serpentine Biotite-muscovite Andesine-chlorite Andesine-albite
log[aMg2þ/(aHþ)2] ¼ 5.2 log(aK þ/aHþ) ¼ 3.8 log[aCa2þ/(aHþ)2] ¼ 6.5 log(aNaþ/aHþ) ¼ 4.9
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Homogenization temperatures of the melt inclusions hosted in plagioclase ranged from 664 C to 733 C whereas the ones in quartz from 635 C to 650 C (Table 2). These temperatures are w þ50 C higher than the ones obtained from the geothermometry, suggesting that some of these temperatures represent cooling or closure temperatures (Table 2). The water contents (Cw) and entrapment pressures of melt inclusions were also calculated with the MELTS software (Asimow and Ghiorso, 1998). The water contents of the melt inclusions based on their measured melting temperatures of plagioclase range of 1.5e2.1 equivalent wt.% H2O, and that of quartz range of 2.1e2.7 equivalent wt.% H2O. The calculated pressures range from the melt inclusions hosted in quartz and plagioclase, i.e., from 4.4 kbar to 4.6 kbar (Table 2) are determined by using the equation Holtz et al. (1992). These pressures are lower w2 kbar than the pressures obtained from the geobarometry, and attributed to ductile deformation of the pyroxene and hornblende (Table 2, see also Searle et al., 2015). Based on the model the Ebadi and Johannes (1991) for the haplogranite system (Ab-An-Or-Qz-H2O) and their equations, the glass has a composition of An36.3e51.7, Ab1.1e2.7, Or0.06e0.1 and Qz42.4e60.1. EPMA analyses of the glasses showed that they compose of (all in wt.%) 73.45e76.45 SiO2, 13.29e15.78 Al2O3, 0.01e0.65 FeO, 3.59e7.84 Na2O, and 0.02e4.01 K2O, whereas TiO2, MnO, MgO and CaO are 0.03 wt.% (Table 1). The glass composition in plagioclase has (K/Ti)N values ranging from 10.1 to 17, which is considered as typical of MORB composition (Sobolev, 1996; Michael et al., 2002). The analyzed halogen and H2O contents were 0.012e0.021 Cl, 0.0002e0.005 F and 1.45e2.71 wt.%, respectively (Table 1). 6. Geochemistry 6.1. Major element geochemistry
Figure 4. Petrography of melt inclusions trapped in quartz (a and b) and plagioclase (c) crystals. The melt inclusions contain glass, and sometimes bubble or daughter minerals (M.I.: Melt inclusions, P.F.I.: Primary fluid inclusions, S.F.I.: Secondary fluid inclusions in arrays, Qz: Quartz, Pl: Plagioclase). Photographed at þ 22.5 C in plane-polarized, transmitted light.
temperature, isochores for associated melts, and the appropriate melt-solid phase equilibrium, using the MELTS software (Asimow and Ghiorso, 1998). All of the studied melt inclusions were homogenized to liquid upon heating.
The Dadnah felsic granitoids have high SiO2 (66e77 wt.%), low to moderate total alkali contents (Na2O þ K2O ¼ 1.4e10.5 wt.%) and Mg# values ranging from 44.8 to 84.0 (clustering at 65 and 75, Table 1). Utilizing the total alkali-silica (TAS) discrimination diagram of Middlemost (1994) and the normative Qz-An-Or and AnAb-Or ternary plot (after Barker, 1979), the majority of the analyzed granitoids and melt inclusion glass are classified as subalkaline trondhjemites and tonalites (for simplicity from now on are referred as tonalites, Fig. 5a and b). The tonalites and melt inclusion glass of Dadnah area display a wide range of An contents (plagioclase normative compositions from w14 to w 64) with low values of Or, suggesting plagioclase fractionation (Table 1, Fig. 5b). Similar distributions of Ab, An, and Or contents have been reported from previous studies, especially, for the crustal tonalites-togranites (Rollinson, 2009, 2014, 2015; Grimes et al., 2013). Based on the ASI values and silica contents, our tonalites (ASI ¼ 0.74e1.08), trondhjemite (ASI ¼ 0.95) and melt inclusion glass (ASI ¼ 0.71e0.77) are from strong-to-mildly metaluminous to weakly peraluminous and are mainly plotted as I-type granitoids (Fig. 5c). On the Harker diagrams, in which we have added the EPMA analyses of the glass obtained from the melt inclusions and the available data of Rollinson (2015) from the Dadnah area, the Fe2O3, CaO, K2O and MgO contents of our tonalites show strong-to-weak depletion with increasing SiO2. Moreover, Na2O reveals positive correlation with increasing SiO2 (Fig. 6aef). These trends suggest that the tonalitic magmas were evolved due to fractional crystallization of clinopyroxene and plagioclase (alkali feldspar), which was accompanied with an increase of the Ab content of the plagioclase (cores to rims). Scatter of K2O and Na2O indicates that
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K-feldspars and albite were not significant crystallizing phases (Fig. 6c and d, Table 1). However, the fact that the compositions of glass measured from the melt inclusions are similar to the analyzed compositions of the tonalites (e.g., PGD2_2, PGD3_2, PGD6_1, PGD7, Table 1) indicates that some of them represent least-fractionated melts. 6.2. Trace element geochemistry
Figure 5. Tonalite classification plots using major elements: (a) Alkali versus silica classification plot (TAS in wt.%) after Middlemost (1994). The alkaline and subalkaline boundary of Irvine and Baragar (1971) is shown (after Rickwood, 1989). (b) Normative quartz-anorthite-orthoclase (Qz-An-Or) ternary diagram, where the tonalitetrondhjemite fields are shown (after Barker, 1979) (blue triangles refer to composition of the analyzed melt inclusions, D: diorite, T and Tj: tonalite-trondhjemite). The experimental basaltic melt compositions are from Koepke et al. (2004). (c) SiO2 (wt.%) versus ASI plot (ASI ¼ the aluminum saturation index Al/[2(Ca 1.67P) þ Na þ K]. The metaluminous and peraluminous fields are adapted from Maniar and Piccoli (1989).
Chondrite-normalized multi-element plot of the tonalites shows that the high field strength elements (HFSE) are rather depleted relative to the large-ion lithophile elements (LILE). Also U and Ti (Ti/Ti* ¼ 0.00e0.27) display strong negative and the Nb, Ta, Zr, Y, Hf, and Th moderate negative anomalies, relative to the adjacent elements of similar incompatibility (Zr and REE’s). Zr, Y, Hf, Nb and Ta depletion is related to fractionation of hornblende and zircon. From the more incompatible elements, Ba is strongly enriched and Pb (Pb/Pb* ¼ 7.16e78.13), La and Ce show a small positive anomaly. This trace element distribution pattern is probably attributed to subduction of recycled oceanic crust (Hofmann, 2004; Pilet et al., 2011, based on Monte-Carlo simulations) with minor contamination of oceanic sediments (Fig. 7a). The total REE abundance of the tonalites is up to w283 ppm. Chondrite-normalized REE patterns of the studied samples display two different forms. The first type is essentially sub-parallel with relatively flat and gently sloping patterns for the middle- (MREE) and heavy-rare earth elements (HREE) with (Gd/Yb)N values ranging from 0.8 to 1.1. The undepleted patterns for the MREE and HREE suggest low-pressure during the formation of the tonalitic melts (e.g., garnet-free), which is also consistent with the pressure determined by melt inclusions (Table 2). They appear mildly-tostrong enriched in the light rare earth elements (LREE) relative to the HREE, with (La/Sm)N and LREE/HREE values ranging from 1.1 to 1.7, and 4.2 to 62.7, respectively (Table 1). They also show weak positive Eu and negative Ce, Sm and Ho anomalies. These trends suggest fractional crystallization of plagioclase and hornblende. These tonalites which contain plagioclase, quartz, hornblende and minor clinopyroxene are metaluminous with (La/Yb)N values of 1.9e4.5 indicating that are moderate fractionated (Fig. 7b, Table 1). Fractionation and enrichment of LREE is considered typical for felsic granitoids related to SSZ, and is also found in the tonalitegranodiorite suite from the mantle section of Oman ophiolite (Wadi Hemli or Hamiliya; Rollinson, 2014, 2015). The second type occurs with the same overall form (e.g., (Gd/ Yb)N ¼ 0.8e2.1 and (La/Sm)N ¼ 0.9e1.2). In addition, they characterized by intense positive Eu (Eu/Eu* ¼ 2.6e7.1), no or slightly positive Ce and negative Sm anomalies, coupled with high Ba and low Rb/Sr values which are attributed to accumulation of plagioclase (Table 1). These tonalites that contain plagioclase, quartz, biotite and minor hornblende are peraluminous with (La/Yb)N values of 0.5 suggesting that they are least fractionated (Fig. 7b, Table 1). For comparison, we have also plotted the host harzburgite (e.g., sample OP_2) that displays similar U and Th depletion and Ba enrichment to the tonalities (Fig. 7a). The host harzburgite shows the lowest values of (La/Sm)N ¼ 0.28, (La/Yb)N ¼ 0.21, (Gd/ Yb)N ¼ 0.66 and Mg# ¼ 64, negative Ce (w0.9) and strong positive Eu (w2.2) anomaly. Few of Dadnah tonalites appear with (La/ Sm)N, (La/Yb)N, (Gd/Yb)N and Mg# values similar to their host harzburgite. This suggests that there was interaction of the tonalite melt with its host harzburgite. In one of the thickest dikes, we have sampled both the margins and the median plane of the dike (e.g., PG4Ma and PG4Mi). This dike appears internally geochemically variable from the margins towards the median plane. For the major oxides there is a decrease of the Fe2O3t and
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Figure 6. (aee) Harker diagrams for selected major elements (wt.%). The black circles represent the tonalites, the red squares the host harzburgite, the blue triangles the analyzed melt inclusions and the purple rhombs the data of Rollinson (2015) from the Dadnah area. (f) Ti versus Al2O3/TiO2 plot depicting fractional crystallization for the tonalites.
Figure 7. Chondrite-normalized (a) multi-element plot (concentrations are normalized to chondrite average composition of McDonough and Sun, 1995) and (b) rare earth element (REE) concentrations (normalized after chondrite of Sun and McDonough, 1989). The solid lines represent the metaluminous tonalites, and the dashed lines the peraluminous tonalites. For comparison, we have also plotted the host harzburgite (e.g., sample OP_2, red circles).
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MgO (w60%) and CaO and increase of their K2O (w40%) and Na2O contents (Table 1). Also regarding the trace elements, there is a decrease of the Ba and SREE (w30%), Hf and Th (w40%) and Ta (w70%) and an increase of the Zr (w30%) and Pb (w60%) contents. These relationships support that fractionation increased towards the center of the dike and there was also interaction with the host harzburgite at the margins. Pearce et al. (1984) proposed the Y versus Nb trace elements for the classification of granitoids, on the basis that Y and Nb are relatively more independent for alteration. In Fig. 8a, it is supported that our samples plot mainly in the volcanic arc granites field (VAG) which is also evident from the trace element patterns (Fig. 7a). Additionally, it appears that the majority of the tonalites and glass intruded under compression, suggesting that the melts are associated with the subduction process (Fig. 8b, see discussion also). 6.3. Characterization of source and magmatic differentiation Based on the Nd versus Nd/Ce plot, we have been able to distinguish the two plausible processes responsible for the magmatic differentiation of the tonalite melts (Fig. 8c). The positive trend between Nd and Nd/Ce (Fig. 8c) implies that the tonalites were produced from partial melting, whereas the constant values of Nd/Ce over a wide range of the Nd content as well as the increasing of SiO2 with decreasing (La/Sm)N at almost constant LaN values, indicate that they were subjected to fractional crystallization (Fig. 8c and d). Following the evidence for partial melting involvement in the formation of the tonalites, the CaO/(MgO þ FeOt) versus Al2O3/ (MgO þ FeOt) plot indicates that the tonalites and related melt inclusion glass derived from partial melting of basaltic and greywacke sources (Fig. 8e). The majority of them, however, relate to a basaltic protolith (Mg# values of 44e53, Kelemen et al., 1997, and TiO2 values for the melt inclusions glass of 0.03 wt.%, Table 1). These results are consistent with the ones obtained from halogens (Table 1). It is apparent that there are two types of tonalites. The first group has elevated Cl/B ratios (Cl/B ¼ 50) depicting the presence of tourmaline, while the second group shows lower Cl/B ratios (Cl/B 22.5) depicting the presence of apatite. The presence of tourmaline and apatite could be related to both partial melting of sedimentary and igneous protoliths. The majority of the tonalites, according to the log(Nb/Y) versus log(Ti/Y) diagram, are plotted in the transitional field from tholeiitic to calc-alkaline affinity (Fig. 8f). These trends are consistent with various degrees of partial melting of subducting oceanic crust (e.g. Martin et al., 2005; Thorkelson and Breitsprecher, 2005). 6.4. Isotope geochemistry Previous studies of oxygen isotopes of the Semail plagiogranites (obtained from tonalites-trondhjemites and minerals: plagioclase, pyroxene, hornblende, quartz and zircon) indicated that d18O values range from 2.4& to 14.0& (e.g. McCulloch et al., 1980, 1981; Gregory and Taylor, 1981; Stakes and Taylor, 1992; Grimes et al., 2013). McCulloch et al. (1980, 1981), and Gregory and Taylor (1981) suggested that the d18OH2O values have been modified due to interaction of seawater with the oceanic crust. Gregory and Taylor (1981) attributed the average d18OH2O values (e.g. 5.7& 0.2&) of the oceanic crust to seawater buffering due to hydrothermal circulation. Unfortunately, there are no calculated d18Omelt or d18OH2O values for the isotopic equilibrium of the plagiogranites or their minerals. Additionally, the d18OH2O values given by Gregory and Taylor (1981) were calculated based on the assumption that both the granitoids and all of their minerals were in equilibrium with a fluid phase and not with the melt.
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Stakes and Taylor (1992) have argued that the plagiogranites represent the final magmatic event that occurred at the spreading axis. Grimes et al. (2013) have interpreted the low d18O values of zircons in tonalites-trondhjemites due to re-melting of the crust, which was altered by hydrothermal fluids (i.e., seawater at higher temperatures). Recently, Spencer et al. (2017) reported unusually high d18O values from peraluminous, S-type granite samples in UAE and Oman that range from 14& to 28& (whole-rock: 14&e23&, quartz: 20&e22&, and zircon: 14&e28&). They suggest that these extremely high d18O values indicate that the protolith of these granitoids was marine pelitic/siliceous muds that entered the mantle through the subduction zone. In our study, the oxygen and hydrogen isotope compositions were obtained from the tonalite minerals, including pyroxene, plagioclase, quartz, hornblende, biotite, and calcite (Table 4). For the calculations, we have used the average trapping temperature of melt inclusions as well as the temperatures obtained from the application of the different geothermometers (Table 2). We have utilized the mineral-melt and -H2O equations of Zhao and Zheng (2003) and Zheng (1993a,b). The calculated d18OMelt, d18OH2O and d18DH2O values are listed in Table 4. The d18Omelt or d18OH2O values range from 5.3& to 6.7&, clustering at w5.3& and 6.2&. The d18OH2O values of hydrous mafic minerals of the tonalites are higher (e.g., 7.3&e7.8&, Table 4), whereas the d18OH2O values of quartz and calcite from the later veins are lower (e.g., 1.8&e3.0&, Table 4). Valley (2003) and Carmody et al. (2013) suggested that the values of w5.5& represent a lithospheric mantle source, whereas the values that are w 6.8& are enriched relative to the mantle and retain isotopic characteristics of typical oceanic crust related to MORB (Zi et al., 2012). Additionally, the d13CCO2 values of the vein calcite are plotted in the ‘Igneous calcite box’ indicating a MOR-type source for CO2 (Fig. 9c, Chaussidon et al., 1991). We report, for the first time, nine helium isotope compositions obtained from pyroxene, plagioclase, hornblende and biotite in the Dadnah tonalites (Table 4). The R/RA compositions range from 6.05 to 7.99 (Table 4) and are close to or lower than the values attributed to MORB (R/RA ¼ 8, Savelieva et al., 2008; Lustrino and Anderson, 2015). It is implied that the possible helium source for our analyzed tonalites is related to MOR-type oceanic crust. Most of these minerals were also analyzed for their hydrogen isotopes. Previous studies of hydrogen isotopes from Gregory and Taylor (1981) obtained data from whole rock and hornblende. Their measured dD values range from e56& to 0.0&, and are attributed to magmatic H2O, which interacted with seawater, however, there are no data for the d18DH2O values. We have obtained a dDmelt value of e85& from pyroxene, which is also related to MOR-type oceanic crust (Chaussidon et al., 1991). Spooner (1977), Gregory and Taylor (1981) and McCulloch et al. (1981) reported that 87Sr/86Sr initial ratios of the mantle and crust lithologies of the Semail ophiolite display extremely large variation (e.g., 0.7030e0.7065). They suggested that they were modified due to hydrothermal interaction of oceanic crust with seawater. McCulloch et al. (1981) also reported that the initial 143Nd/144Nd ratios of these rocks have a limited range of εNd values of 7.5e8.6, indicating that all have distinctive oceanic affinity. Cox et al. (1999) reported negative εNd values and intermediate-to-high initial 87 Sr/86Sr ratios of the analyzed monzogranites-to-leucogranites from the Khor Fakkan block, and suggested mixing of a LILE enriched mantle component, i.e., metabasalts and a continental sedimentary component. Haase et al. (2016) proposed that the almost constant age-corrected εNd values of w8 from the tonalites over a range of 87Sr/86Sr initial values (e.g., 0.7035e0.7060) resemble to MORB oceanic crust affected by seawater. The rubidium, strontium, samarium and neodymium isotope results were obtained from fifteen samples of the tonalite minerals,
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Figure 8. Diagrams indicating the possible tectonic settings of Dadnah tonalites (symbol color same as Fig. 6). (a) Y (ppm) versus Nb (ppm) plot. The WPG (Within Plate Granite), VAG þ Syn-COLG (Volcanic Arc Granite, Syn-Collisional Granite) and ORG (Ocean Ridge Granite) fields are adapted from Pearce et al. (1984). (b) SiO2 (wt.%) versus log (CaO/ (Na2O þ K2O)) plot. The extensional and compressional fields are adapted from Wu and Kerrich (1986). (c) Nd (ppm) versus Nd/Ce plot. The partial melting and fractional crystallization trends are adapted from Pearce (1980) and Ahmad and Tarney (1991). (d) LaN versus (La/Sm)N plot. The partial melting and fractionation trends are adapted from Xu et al. (2015). (e) Molar CaO/(MgO þ FeOt) versus molar Al2O3/(MgO þ FeOt) plot (A: partial melts with metapelitic source, B: partial melts with metagreywackes source and C: partial melts with meta-basaltic and -tonalitic source). The meta-pelitic, -greywackes and meta-basaltic and -tonalitic fields are adapted from Altherr et al. (2000). (f) Nb/Y versus Ti/Y plot. The tholeiitic, transitional and calc-alkaline fields are adapted from Winchester and Floyd (1977) as modified by Pearce (1996).
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Table 4 Oxygen, hydrogen, carbon, silicon, and helium isotope data of the studied tonalites. Utilized the mineral-melt-H2O and -CO2 equations and isotope fractionation values from Zhao and Zheng (2003) and Zheng (1993a,b). Samples
Minerals
d18O (&) T ( C) d18Omelt d18OH2O dD (&) dDmelt dDH2O d13C d13Cfluid d30Si (&)
PGD6_1a PGD6_1b PGD6_1c PGD6_1a PGD6_1c PGD6_1c PGD6_1c PGD6_1b PGD6_1c PGD6_1b PGD6_1a PGD6_1c PGD6_1b PGD6_1a PGD6_1c
Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, coarse-grained Pyroxene, coarse-grained Plagioclase, fine-grained Plagioclase, coarse-grained Hornblende Biotite Calcitea Calcitea Calcitea Quartza Quartza Quartza
4.4 4.7 4.3 4.3 4.2 4.2 5.1
730 730 730 700 700 680 620
5.0 5.3 5.9 5.7 5.4 5.8 5.5 5.6
620 550 430 430 430 430 430 430
(&)
(&) 37
6.4 6.7 6.3 6.2 6.1 5.3 6.1 7.3 7.8 3.0 2.8 2.5 2.1 1.8 1.9
55 61
(&)
(&)
(&)
(&)
85
77.1 95.2
He (107 cm3 3He (1013 cm3 3He/4He R/RA (106) STP/g) STP/g) 4
3.56 3.77 3.64 2.43 4.19 3.37 2.28
38.49 39.88 37.92 22.95 38.56 37.45 23.17
10.81 10.58 10.42 9.44 9.20 11.11 10.16
7.78 7.61 7.49 6.79 6.62 7.99 7.31
4.23 4.84
37.93 40.67
8.97 8.40
6.45 6.05
6.2 4.7 5.6 4.1 5.8 4.3 62
e 0.43 e 0.42 e 0.41
a Temperatures were determined independently based on the calcite-quartz equilibrium pairs (based on the equations of Zhao and Zheng, 2003 and Zheng, 1993a, b). The isotopic temperatures obtained are w430 C.
including pyroxene, plagioclase, hornblende, biotite, quartz and calcite. Unfortunately, we were able to obtain only two 143Nd/144Nd isotopic values from hornblende and biotite. Accordingly, we have calculated their initial isotopic ratios, epsilon values (εSr and εNd) and isotopic ages (Table 5). Their initial 143Nd/144Nd and εNd values (e.g., 0.511968 and 0.512032, and e13.07 and e11.82, Table 5) are lower than those of the MORB lithospheric mantle (εNd z þ19) and the chondritic reservoir (εNd ¼ 0.0). These values are similar to those of MORB protolith and suggest contamination from oceanic sediments (Chaussidon et al., 1991; Savelieva et al., 2008; Zi et al., 2012; Xu et al., 2015 and references therein). From the hydrogen versus oxygen isotope plot, it is evident that all of the calculated d18OMelt, d18OH2O and d18DH2O values relate to the Subduction-Related Vapor, Arc and Crystal Felsic Magma (Fig. 9a, Kus¸cu et al., 2011). The possibility of mixing of different end-members to the formation of the tonalite melts can be depicted from d18Omelt or d18OH2O versus age-corrected εSr values (Fig. 9b). MOR-type lithospheric mantle is characterized by lower d18OMelt and εSr value (e.g., d18OMelt w 5.6& and εSr w e30) relative to oceanic crust (e.g., d18OMelt w 6.8& and εSr w 0e10), and oceanic sediments (e.g., d18OMelt > 10& and εSr > 35) (Ito and Stern, 1986; Zi et al., 2012; Carmody et al., 2013; Roberts and Spencer, 2014; White and Klein, 2014; Lustrino and Anderson, 2015, and references therein). The tonalite samples on Fig. 9b are plotted on the mixing line between the oceanic lithospheric mantle and oceanic crust. This suggests that from these two sources, the oceanic crust is dominant (Fig. 9b). However, this plot cannot distinguish between the genesis of the melt due to direct mantle-derived differentiation or to re-melting of the oceanic crust. Moreover, the εSr coupled with the εNd values of the tonalites suggest that their parental magmas were formed from partial melting of mainly recycled oceanic crust (Table 4, see Stracke et al., 2003). The composition of silicon isotopes from the tonalites suits with the silicon isotope composition of recycled oceanic crust related to partial melting (i.e. d30SiNBS e0.4&; Pringle et al., 2016). Our helium isotopic data concur with the hypothesis of recycled oceanic crust of MORB affinity that subducted in a lithospheric mantle (Savelieva et al., 2008; Graham et al., 2014; Lustrino and Anderson, 2015). Additionally, in the R/RA versus (87Sr/86Sr)i plot (Fig. 9d) there is a negative trend between the R/ RA and initial 87Sr/86Sr isotopic values, which is consistent with contamination from oceanic sediments of the subducted slab (Savelieva et al., 2008; Lustrino and Anderson, 2015).
7. Discussion 7.1. Formation mechanisms and source for the Dadnah tonalities Previous studies in the Khor Fakkan block have frequently depicted the monzogranites-to-leucogranites as low to high potassic and mostly peraluminous, with high contents of LILE and elevated 87Sr/86Sr ratios relative to mantle (Briqueu et al., 1991; Peters and Kamber, 1994; Cox et al., 1999 and references therein). These geochemical characteristics have been interpreted as contribution from a crustal source, attributed to the partial melting of metasedimentary rocks (e.g., Cox et al., 1999; Spencer et al., 2017). Our analytical data suggest that the Dadnah tonalites are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, (Rb, Ba, Sr, Pb, Eu), and Ta, La, Ce, depleted in U, and occur with low d18OH2O, moderate εSr and negative εNd values. The Dadnah tonalites most probably formed from a complex process involving partial melting of the subducted slab of MOR-type composition, accumulation of plagioclase and fractional crystallization, and interaction of the tonalite melt with the host harzburgite. Accumulation of plagioclase, as a mechanism for the genesis of the Dadnah tonalites, is evidenced by the intense positive Eu (Eu/ Eu* ¼ 2.6e7.1, Fig. 7b), negative Sm anomaly coupled with high Ba and low Rb/Sr values (Table 1). The Eu anomaly may also indicate the temperature of solidification of the melts during the plagioclase accumulation. Temperatures decreased from 730 C down to 665 C (Table 2) as the melts ponded at the Moho Transition zone, due to its behavior as a permeability barrier (Kelemen et al., 1997). The strong positive Eu, Ba and negative U and Th anomalies (Fig. 7a and b) coupled with Mg# of w64, similar to their host harzburgite, suggest that some of the tonalites have interacted with their host rock (Fig. 3f). Fractionation is suggested by the linear trends on the normative An-Ab-Or ternary plot (Fig. 5b), the depletion of the Fe2O3, CaO, K2O and MgO coupled with the positive correlation of Na2O with increasing SiO2 on the Harker diagrams (Fig. 6aee) and the linear trend in the Ti versus Al2O3/TiO2 plot (Fig. 6f). Fractionation is also supported by the enrichment of the tonalities with Mg# values 75 in the LREE relative to HREE (LREE/HREE z 4.2e62.7), the constant values of Nd/Ce over a wide range of Nd content (Fig. 8c) and negative trend of increasing SiO2 wt.% with decreasing (La/Sm)N at almost constant LaN values (Fig. 8d).
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Figure 9. Diagrams indicating origins of melts and fluids involved in the formation of the tonalites: (a) Hydrogen versus oxygen isotope plot, showing the calculated water compositions from Dadnah tonalites. The ‘Meteoric Water Line’ and ‘Kaolinite Line’, ‘Subduction-Related Vapor, Arc and Crystal Felsic Magma’, ‘Metamorphic water box’, ‘Magmatic melts and Traditional Magmatic Water box’ are obtained from Kuþcu et al. (2011) and references therein. (b) εSr versus d18Omelt or d18OH2O compositions of Dadnah tonalites, relative to possible magma sources. ‘Lithospheric mantle’, ‘Oceanic crust’ and ‘Sediments’ fields are after De Paolo (1980), Ito and Stern (1986), Cox et al. (1999), Rudnick and Gao (2003), White and Klein (2014), and Lustrino and Anderson (2015). The inset histogram is a detail of the d18Omelt or d18OH2O axis. The colored bars represent the analyzed minerals, i.e., pyroxene, plagioclase, quartz, calcite, hornblende and biotite. (c) Carbon versus oxygen isotope diagram for vein calcites showing the calculated isotope systematics of the tonalites. The ‘Marine Limestone’ and ‘Igneous Calcite’ boxes are from Bowman (1998). (d) R/RA versus (87Sr/86Sr)i plot, relative to possible magma sources. The ‘MORB line’ and ‘Oceanic crust’ field are after Savelieva et al. (2008) and Lustrino and Anderson (2015). (e) 87Rb/86Sr versus (87Sr/86Sr)i plot, relative to magma mixing and possible sources. ‘Lithospheric mantle’ and ‘Oceanic crust’ are after Ito and Stern (1986), Rudnick and Gao (2003), and Lustrino and Anderson (2015).
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Table 5 RbeSr and SmeNd isotope results of the studied tonalites. Samples
Minerals
87
Rb/86Sr measured
87 Sr/86Sr measured
87 Sr/86Sr initial
Age (Ma)
εSr
PGD6_1a PGD6_1b PGD6_1c PGD6_1b PGD6_1a PGD6_1b PGD6_1c PGD6_1c PGD6_1a PGD6_1b PGD6_1b PGD6_1c PGD6_1a PGD6_1b PGD6_1c
Pyroxene, coarse-grained Pyroxene, coarse-grained Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, fine-grained Plagioclase, fine-grained Plagioclase, coarse-grained Quartz Quartz Quartz Calcite Calcite Calcite Hornblende Biotite
0.24 0.36 0.39 0.37 0.26 0.15 0.19 0.20 0.19 0.18 0.15 0.16 0.17
0.705325 0.705841 0.704958 0.704977 0.704886 0.704865 0.705337 0.707357 0.707760 0.707342 0.708002 0.707460 0.707190
0.705650 0.706166 0.705283 0.705295 0.705211 0.705190 0.705662 0.707682 0.708085 0.707667 0.708327 0.707785 0.707515
98.22a,c 98.5 2a,c 97.82b,c 97.32b,c 97.52b,c 97.72d 98.61d 95.61e 95.71e 95.71e 95.4 1 94.9 1 95.2 1
16.3 23.6 11.1 11.3 10.1 9.8 16.5 45.2 50.9 45.0 54.3 46.6 42.8
147 Sm/144Nd measured
143 Nd/144Nd measured
143 Nd/144Nd initial
Age (Ma)
εNd
0.246 0.257
0.512468 0.512475
0.511968 0.512032
96.51f 96.31g
13.07 11.82
The MSWD (Mean Square of Weighted Deviates) shown in the table is a measure of the ratio of the observed scatter of the points (from the best-fit line) to the expected scatter (from the assigned errors and error correlations). a MSWD ¼ 0.83 for coarse-grained pyroxene. b MSWD ¼ 0.88 for fine-grained pyroxene. c MSWD ¼ 0.85 for all analyzed pyroxene. d MSWD ¼ 0.95 for fine- and coarse-grained plagioclase. e MSWD ¼ 0.97 for quartz. f MSWD ¼ 0.80 for hornblende. g MSWD ¼ 1.12 for biotite.
The degree of fractionation that the tonalite melts were subjected was performed with the AFC-Modeler software (Ersoy and Helvaci, 2010). For this model we used two incompatible trace elements Rb and Zr (D ¼ 0.01 and 0.002, respectively) and the least and most evolved samples (e.g., PG1, SiO2 ¼ 66.92 wt.%, Zr ¼ 36 ppm and Rb ¼ 79 ppm, and PG5, SiO2 ¼ 77.99 wt.%, Zr ¼ 79 ppm and Rb ¼ 80.4 ppm, Table 1). Assuming Rayleigh crystal fractionation, we estimate that the Dadnah tonalites experienced w55%e57% fractional crystallization (F ¼ 73). Partial melting is implied by the fact that the Dadnah tonalities are enriched in LILE, the positive trends of Nd versus Nd/Ce and LaN with (La/Sm)N (Fig. 8c and d). The CaO/(MgO þ FeOt) versus Al2O3/ (MgO þ FeOt) plot indicates that the tonalites and related melt inclusion glass were derived from the partial melting of basaltic and greywacke sources (Fig. 8e). The majority of them, however, relate to a basaltic protolith (Mg# values of 44e53, Kelemen et al., 1997, and TiO2 values for the melt inclusions glass of 0.03 wt.%, Table 1). Their elevated εSr values relative to mantle and negative εNd values, coupled with the composition of silicon isotopes, suggest that their parental magma originated from partial melting of mainly recycled oceanic crust with minor contribution of melting from oceanic sediments (Table 5). The degree of partial melting in the oceanic slab of MOR-type composition was performed with the AFC-Modeler software (Ersoy and Helvaci, 2010). For this model, we used the typical LaN and (La/Sm)N composition of the oceanic lithosphere (Sun and McDonough, 1989) and the least fractionated tonalite sample PGD7 with (La/Yb)N values of w0.3 (Table 1, Fig. 8d). Assuming the continuous and batch melting models, we obtain that the Dadnah tonalites formed due to w10%e15% continuous melting or 11.6% (w12%) batch partial melting of the oceanic slab (Fig. 8d). In an attempt to estimate the possible contribution of the oceanic lithosphere representing the slab (as defined by Jacobs et al., 2015, his Fig. 6 for the Semail ophiolite) and the overlying sediments that possibly involved in the formation of Dadnah tonalites, we have employed the mixing model method of De Paolo (1980) (Fig. 9b and d). For our calculations, we have used the AFCModeler software (Ersoy and Helvaci, 2010). The tonalite melt consists of two end-members, i.e., oceanic lithosphere (with an
age-corrected isotopic composition of (87Sr/86Sr)i,96Ma ¼ 0.703 and Rb/86Sr ¼ 0.001, Allégre and Sutcliffe, 2008; White and Klein, 2014; and oceanic sediments (87Sr/86Sr)i,96Ma ¼ 0.716 and 87 Rb/86Sr ¼ 0.32, Cox et al., 1999; Rudnick and Gao, 2003). The oceanic lithosphere component as a source of the least-fractionated and fractionated Dadnah tonalites is contributing 93% to 97.5%, with an average of w96%, while the melted sediments contribution is 2.5%e7%, with an average of w4% (Fig. 9b). The low contribution of oceanic sediments (w4%) in the melt that formed the Dadnah tonalites is also supported by the observation of Ernewein et al. (1988) which argue that no major sedimentary layer was involved during the transition of V1 and V2 volcanism. We have also estimated the contribution of the oceanic crust component of MOR-type composition on the tonalite partial melt. For our calculations, we have used the incompatible element Rb. A typical RbMean value for the oceanic crust is considered as 1.4 ppm (White and Klein, 2014). The least-fractionated tonalites have Rb values ranging from 4.1 ppm to 82 ppm (average of w34 ppm, Table 1), while the fractionated ones have values ranging from 7.3 ppm to 130.3 ppm (average of w43 ppm, Table 1). Since the average contribution of the oceanic lithosphere, as estimated above, is w96%, the concentration of Rb in the least-fractionated and fractionated tonalites ranges of 3.9e78.7 ppm and 7e124.9 ppm, respectively. Applying these values on the MELTS software (Smith and Asimow, 2005) the obtained F values are 1.8%e7.1% for the least-fractionated and 2.2%e4.1% for the fractionated tonalites. Therefore, the least-fractionated tonalite melts originated from the partial melting of w1.8%e7.1% lithospheric mantle and w92.9%e 98.2% recycled oceanic crust, while for the fractionated tonalites the lithospheric mantle contribution was 2.2%e4.1% and 95.9%e97.8% of recycled oceanic crust (average values for all tonalites, 97% recycled oceanic crust and 3% lithospheric mantle; Fig. 9e). A possible origin of the recycled oceanic crust is evident from the occurrence of mafic enclaves, the LREE/HREE ratios of 26e63, the R/RA and lower εSr values (e.g., w10, Tables 4 and 5), the association on CaO/ (MgO þ FeOt) versus Al2O3/(MgO þ FeOt) plot in which the tonalites and the glass contained in the melt inclusions display Mg# values of 44e56, strong Ti depletion (Ti/Ti* 0.27) and the Pb spike in the spidergram (Pd/Pb* up to w 82, Fig. 7a, Table 1; see also Hofmann 87
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and White, 1982; Hofmann, 2004) suggest mainly a meta-basaltic protolith field (Fig. 8e, field C). It is quite difficult to define with confidence the protolith (i.e., volcanic rocks, gabbros, amphibolites etc.) that comprised the recycled oceanic crust which partially melted to generate the Dadnah tonalite melt, since the magmatism in the Khor Fakkan block was almost continuous during the late Cretaceous and the transition from V1 to V2 phase occurred almost at the same time, i.e., 96.4e95 Ma (Goodenough et al., 2014; Thomas and Ellison, 2014). Furthermore, the subducted slab should be also comprised by older oceanic crust formed during the Arabian passive margin development (Triassic to Cretaceous Haybi volcanics, cherts and shallow water exotic limestones; Searle et al., 2015). However we cannot rule out the possibility of layered or massive gabbros below those units as a source too. The Dadnah tonalities contains clinopyroxene as a minor mafic mineral (Fig. 3d, with high Mg# values of 84e85.5 and TiO2, Cr2O3 and NiO depletion), while their plagioclases are less calcic (An44e56, Figs. 3g and 5b) than the experimental products of Koepke et al. (2004). Koepke et al. (2004) have melted gabbro samples, amongst others, at T 900 C which contain clinopyroxene (with Mg# values of 75e78) and plagioclase (An53e55) to produce plagiogranites enriched in Al and depleted in Ti and Fe. These An contents can be explained due to the smaller influence of water during the plagioclase crystallization (Koepke et al., 2004). Alternatively, these plagioclases with cores An53e56, which occur in the least-fractionated tonalites with (La/Yb)N values 0.5, may represent residual plagioclase (Fig. 3d and g) after the partial melting event of the recycled oceanic crust. The least-fractionated, weakly peraluminous tonalites can be also derived from the fractional crystallization and plagioclase accumulation of a metaluminous melt (Lee and Morton, 2015; Gao et al., 2016). Plagioclase accumulation is an effective way for the tonalite melt, which formed by partial melting of the subducted slab, to evolve from strong-to-mildly metaluminous to weakly peraluminous (ASI 1.08). In this scenario, which can also give an explanation for the two different tonalite types (metaluminous and peraluminous types in the sill of D6 location; Fig. 2), the Dadnah metaluminous tonalites were subjected to feldspar fractionation (ASI z 1, Bilal and Giret, 1999). Feldspar accumulation is not expected to significantly change the Al content of the residual tonalite melt (Bilal and Giret, 1999; Frost and Frost, 2008). Textural evidence (Fig. 3d) suggests that feldspar accumulation which occurred at temperatures ranging from 665 C to 730 C (Table 2) was followed by fractionation of hornblende (ASI ¼ 0.5, Bilal and Giret, 1999) at T z 620 C (Table 2) that resulted in the formation of the moderate fractionated, metaluminous tonalites (ASI 0.71) coupled with an increase of the Al content in the residual melt. Then, the residual was mixed and contaminated with the oceanic sediments and so was relatively enriched in water (higher PH2O), which facilitated the crystallization of biotite (ASI ¼ 1e1.5, Bilal and Giret, 1999). Biotite that commonly replaces hornblende in Dadnah tonalites incorporates the remaining Al from the residual melt (Frost and Frost, 2008) and its crystallization, at T z 570 C (Table 2) differentiate the tonalite melt to evolve to a more leucocratic and less fractionated peraluminous type (ASI 1.08). 7.2. Effect of water content and alteration on tonalites A small volume of water can significantly increase the degree of partial melting for the generation of the plagiogranites (Koepke et al., 2004). The melt inclusion glass in our plagioclase obtained from the uppermost part of mantle sequence contains 2.1e2.7 equivalent wt.% H2O (Table 2), showing that considerable amount of water was involved during the fractional crystallization. From our mineralogy results, the plagioclase, hornblende and biotite
mainly contribute to the aluminous character of the Dadnah tonalites. Hornblende and biotite can derived from the partial melting of metasedimentary rocks, but our hydrogen and oxygen isotopes (e.g., d18OH2O ¼ 7.3&e7.8& and d18DH2O ¼ e95.2& to e77.1&, Table 4), suggest that these minerals are associated with typical magmatic fluids (Fig. 9a). Therefore, we suggest that the mafic minerals of the tonalites are most probably related to slabderived fluids (Fig. 9a and c). The low d18OH2O values obtained from quartz and calcite (Table 4) coupled with the Y/Ho ratios that are w 100 (Table 1) suggest that some of the tonalites may have interacted with seawater (Bau et al., 1996). The influx of seawater and its interaction with the tonalites resulted in the precipitation of the late quartz-calcite veins at T w430 C (based on the temperatures obtained from calcite-quartz isotopic pair, Table 4). Possibly, few sub-localities in our study area (i.e. sub-locality D6 and D7; Fig. 2) associate to water excess through fault zones. The intense talc and serpentine assemblages observed along these faults indicate that the fluids used these faults as pathways. Our geothermobarometry results suggest that alteration occurred at temperatures of w230 C. Therefore, the Dadnah tonalites appear to have experienced minor excess from seawater and are affected only locally where fault zones developed. 7.3. Genetic model for Dadnah tonalites The plagiogranites can formed at any level in the ophiolite section, through various magmatic processes with diverse source origin (see Rollinson, 2009, 2015). The variation of local conditions influences significantly the plagiogranite formation and thus, is hard to define a simple genetic model. In previous plagiogranite related genetic models, the source was suggested by: (1) the location where the plagiogranite melts generated (i.e., spreading axis or off-axis), (2) the structural level of emplacement (mantle or crustal section of ophiolites) and (3) the variety of magmatic processes, such as partial melting, fractional crystallization, magma mixing and contamination. In order to propose a genetic model for the Dadnah tonalites, we briefly summarize the already published genetic models for the northern and other parts of the Semail ophiolite, e.g., Searle and Cox (1999), Rollinson (2009), Grimes et al. (2013), France et al. (2014), Haase et al. (2015, 2016). The genesis of felsic crustal intrusive rocks in the Semail ophiolite was documented by Haase et al. (2016), where they suggest that these granitoids formed by fractional crystallization from mafic melts that were genetically related. These melts formed by the fluid-induced melting of the mantle wedge and mixing with subducted sediments and emplaced on the crustal sequence under crustal extension (Haase et al., 2016). For the mantle leucogranites Haase et al. (2015) argue that they represent partial melts of pelagic sediments from the subducted slab. The Dadnah tonalites are emplaced on the uppermost part of the mantle sequence of the northernmost block of the ophiolite during compression, while the primary melts of Dadnah tonalite generated by w12% partial melting of the oceanic slab (Fig. 8d) with minor contribution from mantle and sediments. Such a difference in the source contribution between our study and Haase et al. (2015, 2016) suggests spatial heterogeneity and variety of processes in the genesis of the felsic granitoids along the different localities along the Semail ophiolite. The genetic model of Rollinson (2009) included tonalite-togranite samples from both the crustal and mantle sequence of the ophiolite. He argues that the mantle-hosted granitoids formed due to either mixing or contamination at the upper part of the depleted mantle during the early stage of ophiolite emplacement (i.e., from Wadi Hamaliya, Oman). Mixing between 10% and 30% of a metasedimentary melt into the melt of a mafic source is considered as the main genetic process for the tonalites-to-granites in the mantle
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sequence of Oman ophiolite (Rollinson, 2015). This model bears many similarities to our results in the aspect of the source involvement that contributed to the formation of the tonalite melts. However, the Dadnah tonalites have quite different geochemical character and source component contribution, related to the melting of the subducted slab. Searle and Malpas (1982) and Searle and Cox (1999) reported centimeter-to-meter scale tonalitic melt pods and veins, formed by 5%e10% partial melting of the subducted continental crustal basement at the late stages of obduction. Furthermore, Briqueu et al. (1991) based on Sr, Nd and Pb isotopes obtained from felsic granitoids of the mantle sequence in the northeastern Oman ophiolite, suggested anatexis of the metamorphic sole as source for the formation of these granitoids. An alternative model for the genesis of the shallow level intrusive rocks (tonalites-trondhjemites) at/or near the dike-
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gabbro transition involves seawater circulation from an active hydrothermal system in close proximity to the active-melt lenses at the top of the magma chamber (e.g. Grimes et al., 2013; France et al., 2014). However, this model aims to explain the generation of the intrusives at the crustal sequence of the ophiolite. The seawater can infiltrate the deeper level gabbros in a ductile regime through oceanic detachment faults (e.g. Koepke et al., 2007), but is unlikely to reach and react with the melt-lenses in Moho level due to the higher depth. Moreover, Nicolas et al. (2000) reported absence of a magma chamber or mantle diapir in the Dadnah area according to their structural mapping in the Semail ophiolite. Based on our geochemical, isotope and geochronology data (Fig. 10), complemented with new melt inclusion data, we make an attempt to propose a tectono-petrogenetic model for the northern Khor Fakkan ophiolite block with reference to the source
Figure 10. Summary chart for the ages of the Semail Ophiolite, plagiogranite intrusions and metamorphic sole. Our calculated Rb/Sr isochron ages from sample PG 6_1 on (a) pyroxene (coarse and fine grained), (b) plagioclase (coarse and fine grained) and (c) quartz range from 98.1e95.4 Ma. In general our dating results span between 98.5 Ma and 94.9 Ma (Red square -7-, see also Table 5). The green area indicates the possible time interval for the formation of plagiogranites, based on the published ages for the plagiogranites from Tilton et al. (1981) and Warren et al. (2005). The green arrow indicates the maximum age range from subduction-related metamorphism (110 Ma). Blue arrow indicates the metamorphism of the sole. The obtained ages from previous studies are: (1) and (2) U/Pb plagiogranite ages (93.5e95.9 Ma and 97e97.9 Ma) from Tilton et al. (1981). (3) SmeNd isochron age (98.8 9.5 Ma) from Cox et al. (1999). (4) 40Ar/39Ar on garnet-bearing granite in UAE (89.1 Ma and 89.9 0.4 Ma) and (5) from hornblende-bearing granite (93.8 0.2 Ma) from Hacker et al. (1996). (6) U/Pb resampling of the plagiogranite of Tilton et al. (1981) (93e83 Ma, 95.5 0.24 Ma, and 95.16 0.2 Ma) from Warren et al. (2005). (8) The oldest age of ophiolite from McCulloch et al. (1981). (9) 40Ar/39Ar on hornblende from amphibolite (95.7e92.4 Ma) from Hacker (1994). (10) K/Ar on metamorphic hornblende from the amphibolite with ages in the wide range 101 Ma to 89 Ma that cluster age near 98 Ma of Gnos and Peters (1993). (11) 40Ar/39Ar on mafic dike intruding the metamorphic sole and lowest part of the mantle sequence (93.7 0.8 Ma) from Hacker and Gnos (1997). (12) 40Ar/39Ar on muscovite from metamorphic sole (92.4 0.2 Ma) from Hacker and Gnos (1997). (13) Hornblende from the metamorphic sole in northern part of the Semail ophiolite (94.9 0.2 Ma) and (14) in the southern part (93.5 0.1 Ma) from Hacker and Gnos (1997). (15) K/Ar biotite from the metamorphic sole (89.2 0.4 Ma) of Gnos and Peters (1993). (16) U/Pb age (99.8 3.3 Ma) of peraluminous granite dikes from zircon cores and rims from Spencer et al. (2017). (17) New high precision UePb dates on zircons (96.169 0.022 Ma to 96.146 0.035 Ma) from Sumeini sole (Rioux et al., 2016) and (17*) UePb dates on zircons (94.815 0.030 Ma) from Wadi Tayin sole (Rioux et al., 2016). Stars in the plagiogranite field are the ages of 98.5 Ma of the composite sill of Ra’s Dadnah and 95.3 Ma trondhjemite. The ages of tonalites-trondhjemites (stars), metamorphism of the sole (blue arrow), exhumation (80e70 Ma) and end of obduction (72 Ma) are obtained from McCulloch et al. (1980, 1981), Gregory and Taylor (1981), Stakes and Taylor (1992), Hacker and Gnos (1997), Cox et al. (1999), Warren et al. (2005), Grimes et al. (2013), Searle et al. (2015), Rioux et al. (2016).
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contribution, the water effect and related magmatic evolution processes (Fig. 11). The subducted older and denser oceanic crust (Triassic to late Cretaceous Haybi complex) was overridden by young hot asthenospheric mantle providing the thermal gradient that facilitated melting at shallower than the expected depth. That intra-oceanic subduction was more evolved in the northern ophiolite blocks (Khor Fakkan and Aswad blocks) compared to the central and south blocks, based on the gradual increase in volume phase 2 magmatism, the influence of fluids from the subducting slab (Goodenough et al., 2014) and the timing of ophiolite emplacement onto the continental margin at 93e92 Ma (Styles et al., 2006) in the north compared to the younger emplacement in the south (82e79 Ma; Warren et al., 2005; Fig. 10). Subduction appears to be almost contemporaneous with the operation of the fast-spreading ridge, based on the almost continuous magmatism from MORB-like composition (phase 1) to the more hydrous island arc-type composition (phase 2) related to the SSZ setting (Goodenough et al., 2014; Haase et al., 2016). During the evolution
of this complex tectono-magmatic process, the input of small amount of sediments on the subducted slab led to the mixing of the recycled oceanic crust (i.e., upper basaltic lavas and lower gabbros) which interacted with the mantle wedge (e.g. Haase et al., 2016), but with the influence of lithospheric mantle being relatively minor. Cox (2000), using mixing calculations on subducted metabasalt (i.e. metamorphic equivalents of Haybi volcanics), suggested that Ra-Dibba monzogranites contain w65% of the basaltic endmember, the Ra-Dadnah composite dikes 30%e50%, and the Wadi Hulw dikes less than 5%, emphasizing the high degree of heterogeneity of the source. Additionally, Rollinson (2015) proposed mixing of 10%e30% metasedimentary melt into a mafic source. Our estimations suggest that the tonalites of Dadnah were the end result of partial melting and subsequent contamination and mixing of w4% oceanic sediments with w96% oceanic lithosphere from the subducted slab. This MORB-type slab melt composed from w97% recycled oceanic crust and w3% of the overlying mantle. Partial melting of the slab assisted also by the fluids released from the
Figure 11. (a) A proposed schematic petrogenetic model for the formation of the Dadnah tonalites intrusive in the mantle sequence at the northern part of the Semail ophiolite. The anatexis of recycled oceanic crust (upper part of slab with basaltic primary composition mixed with minor sedimentary component) produced the off-axis melts and their subsequent minor interaction and mixing with the overlying mantle produced the tonalite melts. The melts recorded temperatures from 720 C to 680 C, ascent and pond at the upper part of mantle, just below the Moho transition zone (plate setting modified from Cox et al., 1999, Goodenough et al., 2014; Searle et al., 2015, Roberts et al., 2016; Soret et al., 2017). (b) Schematic illustration for the process of the tonalite generation at the upper part of the subducting slab. Oceanic lithosphere and sediments on the slab mantle interface can melt at low pressures (3e7 kbar, i.e. 10e20 km; Kostopoulos and Murton, 1992) due to unusually hot conditions within the mantle wedge during the first 5e10 Myr after subduction initiation (Stern et al., 2012). The higher degree of depletion in the lithospheric mantle of the SSZ ophiolite is balanced by the introduction in the subduction zone of older oceanic enriched crust (Haybi complex and below) with the associated sediments carried above the slab. The discontinuous sediment cover along and across the slab interface can be responsible for the variety of granitoid types and ages observed. On the slab interface, the metamorphic sole accretion is controlled by the percentage of partial melting. In our case for the northern part, the sole accretion is inhibited and mechanical decoupling is favored triggering less sedimentary input and more oceanic crust to be involved in the partial melt (see also Soret et al., 2017).
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hydrous subducted sediments. We imply that the composition and thickness of the subducted slab and its sedimentary cover was highly heterogeneous laterally and vertically, consisting by older oceanic crust (mainly of Late Triassic exotic reefal limestones (seamount) above an basaltic substrate, cherts, and mélanges; Haybi complex) and was discontinuously overlain by marine sediments (Searle et al., 2015). Another potential source could be the metabasalts and metagreywackes comprising the basement unit (Hatat fm.) below the Haybi and Hawasina allocthons (Gray and Gregory, 2003). Oceanic crust and sediments on the slab mantle interface can melt at low pressures (3e7 kbar, i.e. 10e20 km; Kostopoulos and Murton, 1992) due to unusually hot conditions within the mantle wedge during the first 5e10 Myr after subduction initiation (Stern et al., 2012). Mantle melting is enhanced due to SSZ fluid input (Pearce et al., 1984). The higher degree of depletion in the lithospheric mantle of the SSZ ophiolite is balanced by the introduction in the subduction zone of the older oceanic enriched crust (Haybi complex) with the associated sediments carried above the slab. The discontinuous sedimentary cover along and across the slab interface can be responsible for the variety of granitoid types and ages observed along the ophiolite structural grain (from UAE to Oman) and across the oceanic mantle and crust (crustal and mantle plagiogranites). Soret et al. (2017) argued that on the slab interface, the metamorphic sole accretion is controlled by the percentage of melting since low melt fractions < 7% do not affect the inter-plate viscous coupling and only contribute to strain localization and weakening on the detachment of the sole thrust slices. For melt fractions >7% (w12% partial melting in our case) sole accretion is inhibited and mechanical decoupling is favored (Soret et al., 2017). This can explain the low percentage of sedimentary component in our tonalite melt, since sole accretion stacks up thicker thrust slices of metasediments, while in the decoupling the oceanic crust carries only thinner sedimentary layers. Partial melting and mixing followed by plagioclase accumulation and fractional crystallization and interaction with the host harzburgite. The melts ascended to higher structural levels (pressure difference from 6.5 kbar to 4.4 kbar, corresponding to w7 km upward flow, Table 2) and ponded at the uppermost part of mantle sequence and in the MTZ of the Khor Fakkan block, which represents a permeability barrier, forming the unique in geochemical and isotopic character Dadnah tonalites. Our calculated Rb/Sr isochron ages for the tonalites yielded a range between 98.6 Ma and 94.9 Ma (Fig. 10, Table 5) which indicates that the crystallization of these melts predates slightly the suggested 96.1e94.8 Ma age range of the metamorphic sole (Rioux et al., 2016). In addition, our ages obtained from sample PG6_1 (Fig. 10a, b and c; Table 5) in which pyroxene has a weighted mean Rb/Sr isochron of 97.86 1.5 Ma (Mean Square of Weighted Deviates, MSWD ¼ 0.85; Fig. 10a), plagioclase of 98.1 1 Ma (MSWD ¼ 0.95; Fig. 10b) and quartz of 95.4 1 Ma (MSWD ¼ 0.97; Fig. 10c). Spencer et al. (2017) reported ages (UePb on zircon) of 99.8 3.3 Ma obtained from Stype, granite dikes that intruded the Semail peridotite. The dates from quartz are identical with the ages (e.g., 96.1e94.0 Ma and within a narrow age range of 95.3 0.5 Ma) reported from Warren et al. (2005), Goodenough et al. (2014), and Rioux et al. (2016). These authors have dated zircon grains from several tonalite and trondhjemite intrusions along the length of the ophiolite. We attribute the older ages obtained to possibly record the partial melting event, i.e., relate to relict plagioclases of higher An contents (An53e56; Fig. 3g) and pyroxenes or to depict tonalites genetically related to phase 1 magmatism (Rioux et al., 2013; Goodenough et al., 2014). Our Sm/Nd isochron of 96.5 1 Ma on Hbl (MSWD ¼ 0.80; Table 5) and 96.3 1 Ma on Bt (MSWD ¼ 1.12; Table 5) are very close to the older ages (e.g. 96.1 Ma) for the tonalites given by Rioux et al. (2016). Our geochronological data,
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despite some analytical uncertainties, leave also a hint that the inta-oceanic subduction might initiated a bit earlier than the ophiolite generation at the spreading ridge. Since the ages of the metamorphic sole, the ophiolite formation and the plagiogranite melt are very close and overlapping, a mechanism including partial melting of the subducted slab with synchronous formation and exhumation for the metamorphic sole as a direct consequence of spreading above a nascent subduction zone is required, explaining in a way the varied estimations for the sedimentary input in the tonalite melts. 8. Conclusions Our concluding remarks based on the synthesis and analysis of geological, mineralogical, petrological, geochemical and isotopic data in the Khor Fakkan ophiolite block in the Eastern Emirates can be summarized as follows: (1) The Dadnah tonalites, intrusive in the mantle section of the Khor Fakkan block, are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, Ta, La, Ce, depleted in U, and occur with low d18OH2O, moderate εSr and negative εNd values. Geochemical and isotopic compositions link our tonalites with magmatic processes occurring at the initial stages of a supra-subduction zone setting. (2) Melt inclusions hosted in plagioclase indicated the leastfractionated tonalites and the PeT conditions (T ¼ 664e733 C, P ¼ 4.4e4.6 kbar) for their formation. (3) The Rb/Sr isochron suggest an age range of 98.6e94.9 Ma for the Dadnah tonalites. (4) The Dadnah tonalites formed by partial melting (w10%e15% continuous or w 12% batch partial melting), accumulation of plagioclase, fractional crystallization (w55%e57%), and mild interaction with the host harzburgite. (5) The stable and noble gas (He) isotopes indicate that the tonalites were formed by melting of mainly recycled oceanic crust that was hydrothermally altered. (6) From the estimation of source contributions, the tonalite melt evolved from w96% of oceanic lithosphere and w4% of oceanic sediments. The oceanic lithosphere was mainly recycled oceanic crust (w97%) with minor component from the overlying mantle (w3%). Acknowledgements We express our appreciation to the Petroleum Institute (part of Khalifa University, UAE) for providing us access to the ADRIC lab facilities and field support during this study. Especially, we would like to thank Dr. Kil and his staff of Chonnam National University, S. Korea for the permission to use their ICP-MS facility. We also thank the staff of the USGS Denver Inclusion Analysis Laboratory for the melt inclusions microthermometry and EPMA analyses, and the Chinese Academy of Geological Sciences (CAGS), Beijing, China for the isotopic analyses. This work is part of H.J.’s MSc thesis project (PIPGSTS-16-16) conducted at the Petroleum Institute, UAE. H. Rollinson and three anonymous reviewers are warmly thanked for their constructive reviews and suggestions and S. Kwon for his editorial handling that lead to a significant improvement of the final version of the manuscript. References Ahmad, T., Tarney, J., 1991. Geochemistry and petrogenesis of Garhwal volcanics: implications for evolution of the North India lithosphere. Precambrian Research 50, 69e88.
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H. Joun et al. / Geoscience Frontiers 10 (2019) 1187e1210
Alabaster, T., Pearce, J.A., Malpas, J., 1982. The volcanic stratigraphy and petrogenesis of the Oman ophiolite complex. Contributions to Mineralogy and Petrology 81, 168e183. Allégre, C.J., Sutcliffe, C., 2008. Isotope Geology. Cambridge University Press, 512 p. Alleman, F., Peters, T., 1972. The ophiolite-radiolarite belt of the North Oman mountains. Eclogae Geologicae Helvetiae 65, 657e698. Altherr, R., Holl, A., Hegner, E., Langer, C., Kreuzer, H., 2000. High potassium, calcalkaline I-type plutonism in the European Variscides: Northern Vosges (France) and Northern Schwarzwald (Germany). Lithos 50, 51e73. Amri, I., Benoit, M., Ceuleneer, G., 1996. Tectonic setting for the genesis of oceanic plagiogranites: evidence from a paleo-spreading structure in the Oman ophiolite. Earth and Planetary Science Letters 139, 177e194. Amri, I., Ceuleneer, G., Benoit, M., Python, M., Puga, E., Targuisti, K., 2007. Genesis of granitoids by interaction between mantle peridotites and hydrothermal fluids in oceanic spreading setting in the Oman Ophiolite. Geogaceta 42, 23e26. Anonymous, 1972. Penrose field conference on ophiolites. Geotimes 17, 24e25. Asimow, P.D., Ghiorso, M.S., 1998. Algorithmic modifications extending MELTS to calculate subsolidus phase relations. American Mineralogist 83, 1127e1132. Ashchepkov, I.V., Vladykin, N.V., Gerasimov, P.A., Saprykin, A.I., Khmel’nikova, O.S., Anoshin, G.N., 2001. Temperature gradient and structure of the lithospheric block beneath the southeastern margin of the Siberian craton: disintegrated xenolith evidence from kimberlitic pipes of the Aldan shield. Doklady Earth Sciences 378, 495e499. Ashchepkov, I.V., 2002. Empirical clinopyroxene thermobarometers for mantle rocks based on the diopside-jadeite exchange. Doklady Earth Sciences 382, 366e370. Ashchepkov, I.V., 2003. More Precise Equation of the Jd-en Barometer. Herald of the Department of Earth Sciences RAS, p. 21. Brophy, J.G., 2008. A study of rare earth element (REE)eSiO2 variations in felsic liquids generated by basalt fractionation and amphibolite melting: a potential test for discriminating between the two different processes. Contributions to Mineralogy and Petrology 156, 337e357. Barker, F., 1979. Trondhjemite: definition, environment and hypotheses of origin. In: Barker, F. (Ed.), Trondhjemites, dacites, and related rocks. Amsterdam, Elsevier. Developments in Petrology 6, 1e12. Bau, M., Koschinsky, A., Dulski, P., Heinz, J., 1996. Comparison of the partitioning behaviours of yttrium, rare-earth elements and titanium between hydrogenic marine ferromanganese crusts and seawater. Geochimica et Cosmochimica Acta 60, 1709e1725. Beaudoin, G., Therrien, P., 2009. The updated web stable isotope fractionation calculator. In: De Groot, P.A. (Ed.), Handbook of Stable Isotope Analytical Techniques 2. Elsevier, pp. 1120e1122. Bechennec, F., Le Metour, J., Rabu, D., Bourdillon-de-Grissac, C., de Wever, P., Beurrier, M., Villey, M., 1990. The Hawasina Nappes; stratigraphy, paleogeography and structural evolution of a fragment of the South-Tethyan passive continental margin. In: Rorbertson, A.H.F., Searle, M.P., Ries, A.C. (Eds.), The Geology and Tectonics of the Oman Region, vol. 49. Geological Society, London, Special Publications, pp. 213e223. Bilal, E., Giret, A., 1999. The aluminum saturation index and the MgO/TiO2 ratio: two parameters influenced by PH2O and their use to discriminate magma series. Revista Brasileira de Geociencias 29, 55e58. Blundy, J.D., Holland, T.J.B., 1990. Calcic amphibole equilibria and a new amphiboleplagioclase geothermometer. Contributions to Mineralogy and Petrology 104, 208e224. Boudier, F., Ceuleneer, G., Nicolas, A., 1988. Shear zones, thrusts and related magmatism in the Oman ophiolite: initiation of thrusting at an oceanic ridge. Tectonophysics 151, 275e296. Bowman, J.R., 1998. Stable-isotope systematics of skarns. In: Lentz, D.R. (Ed.), Mineralized Intrusion-related Skarn Systems, vol. 26. Mineralogical Association of Canada Short Course Series, pp. 99e145. Briqueu, L., Mevel, C., Boudier, F., 1991. Sr, Nd and Pb isotopic constraints in the genesis of a calc-alkaline plutonic suite in the Oman ophiolite related to the obduction process. In: Peters, T.J., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer, Netherlands, pp. 517e542. Buddington, A.F., Lindsley, D.H., 1964. Iron-titanium oxide minerals and synthetic equivalents. Journal of Petrology 5, 310e335. Carmody, L., Barry, P.H., Shervais, J.W., Kluesner, J.W., Taylor, L.A., 2013. Oxygen isotopes in subducted oceanic crust: a new perspective from Siberian diamondiferous eclogites. Geology, Geophysics and Geosystems 14, 3479e3493. https://doi.org/10.1002/ggge.20220. Cathelineau, M., 1988. Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals 23, 471e485. Cathelineau, M., Nieva, D., 1985. A chlorite solid solution geothermometer: the Los Azufres (Mexico) geothermal system. Contributions to Mineralogy and Petrology 91, 235e244. Chaussidon, M., Sheppard, S.M., Michard, A., 1991. Hydrogen, sulfur and neodymium isotope variations in the mantle beneath the EPR at 12 50΄N. In: Taylor, H.P., O’Neil, J.R., Kaplan, I.R. (Eds.), Stable Isotope Geochemistry: A Tribute to Samuel Epstein. Geochemical Society Special Publications, pp. 325e337. Clayton, R.N., Mayeda, T.K., 1963. The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochimica et Cosmochimica Acta 27, 43e52. Clayton, R.N., Muffler, L.J.P., White, D.E., 1972. Oxygen isotope fractionation study of calcite and silicates of the River Ranch, California. American Journal of Science 266, 968e979.
Coleman, R.G., 1981. Tectonic setting for ophiolite obduction in Oman. Journal of Geophysical Research 86, 2497e2508. Coleman, R.G., Peterman, Z.E., 1975. Oceanic plagiogranite. Journal of Geophysical Research 80, 1099e1108. Cox, J.S., 2000. Subduction-obduction Related Petrogenetic and Metamorphic Evolution of the Semail Ophiolite Sole in Oman and the United Arab Emirates (Ph.D. thesis). University of Oxford, Oxford, UK, 169 p. Cox, J., Searle, M., Pedersen, R., 1999. The petrogenesis of leucogranitic dykes intruding the northern Semail ophiolite, United Arab Emirates: field relationships, geochemistry and Sr/Nd isotope systematics. Contributions to Mineralogy and Petrology 137, 267e287. De Paolo, D.J., 1980. Crustal growth and mantle evolution; Inferences from models of element transport and Nd and Sr isotopes. Geochimica et Cosmochimica Acta 44, 85e96. Ding, T.P., 2004. Analytical methods for silicon isotope determinations. In: De Groot, P.A. (Ed.), Handbook of Stable Isotope Analytical Techniques. Elsevier, pp. 523e537. Dixon, S., Rutherford, M.J., 1979. Plagiogranites as late-stage immiscible liquids in ophiolite and mid-ocean ridge suites: an experimental study. Earth and Planetary Science Letters 45, 45e60. Ebadi, A., Johannes, W., 1991. Beginning of melting and composition of first melts in the system Qz-Ab-Or-H2O-CO2. Contributions to Mineralogy and Petrology 106, 286e295. Ernewein, M., Pflumio, C., Whitechurch, H., 1988. The death of an accretion zone as evidenced by the magmatic history of the Sumail ophiolite (Oman). Tectonophysics 151, 247e274. Ersoy, Y., Helvaci, C., 2010. FC-AFC-FCA and mixing modeler: a Microsoft Excel spreadsheet program for modeling geochemical differentiation of magma by crystal fractionation, crustal assimilation and mixing. Computers and Geosciences 36, 383e390. France, L., Koepke, J., MacLeod, C., Ildefonse, B., Godard, M., Deloule, E., 2014. Contamination of MORB by anatexis of magma chamber roof rocks: constraints from a geochemical study of experimental melts and associated residues. Lithos 202, 120e137. Friedman, I., O’Neil, J.R., 1977. Compilation of Stable Isotope Fractionation Factors of Geochemical Interest, 440KK. U.S. Geological Survey Professional Paper, p. 12. Frost, B.R., Frost, C.D., 2008. A geochemical classification for feldspathic igneous rocks. Journal of Petrology 49, 1955e1969. Ghent, E.D., Stout, M.Z., 1981. Metamorphism at the base of the Semail ophiolite. Journal of Geophysical Research 86, 2557e2571. Glennie, K.W., Boeuf, M.G., Hughes-Clarke, M.H.W., Moody-Stuart, M., Pilaar, W.F., Reinhardt, B., 1974. Geology of the Oman Mountains. Verhandelingen Koninklijk, The Nederlands, Geologisch Mijnbouwkundig Genootschap, p. 31. Gnos, E., Nicolas, A., 1996. Structural evolution of the northern end of the Oman Ophiolite and enclosed granulites. Tectonophysics 254, 111e137. Gnos, E., Peters, T., 1993. K-Ar ages of the metamorphic sole of the Semail ophiolite; implications for ophiolite cooling history. Contributions to Mineralogy and Petrology 113, 325e332. Graham, D.W., Barry, B., Hanan, B.W., Hémond, Ch, Blichert-Toft, J., Albarède, F., 2014. Helium isotopic textures in Earth’s upper mantle. Geology, Geophysics and Geosystems 15, 2048e2074. https://doi.org/10.1002/2014GC005264. Gray, D.R., Gregory, R.T., 2003. Ophiolite obduction and the Samail Ophiolite: the behaviour of the underlying margin. Ophiolites in Earth History 218 (1), 449e465. https://doi.org/10.1144/GSL.SP.2003.218.01.23. Goodenough, K.M., Phillips, E.R., Styles, M.T., Thomas, R.J., Farrant, A.R., Arkley, S.L.B., 2006. Geology of the Khor Fakkan 1:50000 Map Sheet. British Geological Survey, United Arab Emirates, pp. 1e84. Goodenough, K.M., Thomas, R.J., Styles, M.T., Schofield, D.I., MacLeod, C.J., 2014. Records of ocean growth and destruction in the Oman-UAE ophiolite. Elements 10, 109e114. Gregory, R.T., Taylor Jr., H.P., 1981. An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail Ophiolite, Oman: evidence for d18O buffering of the oceans by deep (> 5 km) seawater-hydrothermal circulation at mid-ocean ridges. Journal of Geophysical Research 86, 2737e2755. Grimes, C.B., Ushikubo, T., Kozdon, R., Valley, J.W., 2013. Perspectives on the origin of plagiogranite in ophiolites from oxygen isotopes in zircon. Lithos 179, 48e66. Gao, P., Zheng, Y.-F., Zhao, Z.-F., 2016. Distinction between S-type and peraluminous I-type granites: Zircon versus whole-rock geochemistry. Lithos 258-259, 77e91. Haase, K.M., Freund, S., Koepke, J., Hauff, F., Erdmann, M., 2015. Melts of sediments in the mantle wedge of the Oman ophiolite. Geology 43 (4), 275e278. https:// doi.org/10.1130/G36451.1. Haase, K.M., Freund, S., Beier, C., Koepke, J., Erdmann, M., Hauff, F., 2016. Constraints on the magmatic evolution of the oceanic crust from plagiogranite intrusions in the Oman ophiolite. Contributions to Mineralogy and Petrology 171, 1e16. Hacker, B.R., 1994. Rapid emplacement of young oceanic lithosphere. Science 265, 1563e1565. Hacker, B.R., Mosenfelder, J.L., 1996. Metamorphism and deformation along the emplacement thrust of the Samail ophiolite, Oman. Earth and Planetary Science Letters 144, 435e451. Hacker, B.R., Gnos, E., 1997. The conundrum of samail: explaining the metamorphic history. Tectonophysics 279, 215e226. Hacker, B.R., Mosenfelder, J.L., Gnos, E., 1996. Rapid emplacement of the Oman ophiolite: thermal and geochronological constraints. Tectonics 15, 1230e1247.
H. Joun et al. / Geoscience Frontiers 10 (2019) 1187e1210 Hammarstrom, J.M., Zen, E., 1986. Aluminum in hornblende: an empirical igneous geobarometer. American Mineralogist 71, 1297e1313. Helgeson, H.C., Kirkham, D.H., Flowers, G.C., 1981. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures: calculation of activity coefficients, osmotic coefficients, and apparent molal, standard, and relative partial molal properties to 600 C and 5 kb. American Journal of Science 281, 1249e1516. Hemond, C., Hofmann, A.W., Vlastelic, I., Nauret, F., 2006. Origin of MORB enrichment and relative trace element compatibilities along the Mid-Atlantic Ridge between 10 and 24 N. Geochemistry, Geophysics, Geosystems 7, Q12010. https://doi.org/10.1029/2006GC001317. Henry, D.J., Guidotti, C.V., Thomson, J.A., 2005. The Ti-saturation surface for low-to medium pressure metapelitic biotites: implications for geothermometry and Tisubstitution mechanisms. American Mineralogist 90, 316e328. Hofmann, A.W., White, W.M., 1982. Mantle plumes from ancient oceanic crust. Earth and Planetary Science Letters 57, 421e436. Hofmann, A.W., 2004. Sampling mantle heterogeneity through oceanic basalts: isotopes and trace elements. Treatise in Geochemistry 2, 61e101. Holtz, F., Pichavant, M., Barbey, P., Johannes, W., 1992. Effects of H2O on liquidus phase relations in the haplogranitic system at 2 and 5 kbar. American Mineralogist 77, 1223e1241. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences 8, 523e548. Ito, E., Stern, R.J., 1986. Oxygen- and strontium-isotopic investigations of subduction zone volcanism: the case of the Volcano Arc and the Mariana Island Arc. Earth and Planetary Science Letters 76, 312e320. Jacobs, J., Thomas, R.J., Ksienzyk, A.K., Dunkl, I., 2015. Tracking the Oman ophiolite to the surface e new fission track and (U-Th)/He data from the Aswad and Khor Fakkan Massifs, United Arab Emirates. Tectonophysics 644, 68e80. Johnson, J.W., Oelkers, E.H., Helgeson, H.C., 1992. SUPCRT92: a software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species and reactions from 1 to 5000 bars and 0 to 1000 C. Computers and Geoscience 18, 899e947. Johnson, M.C., Rutherford, M.J., 1989. Experimental calibration of the aluminum-inhornblende geobarometer with application to Long Valley caldera (California) volcanic rocks. Geology 17, 837e841. Jowett, E.C., 1991. Fitting iron and magnesium into the hydrothermal chlorite geothermometer. In: GAC/MAC/SEG Joint Annual Meeting, Toronto, A62, p. 16. Kelemen, P.B., Koga, K., Shimizu, N., 1997. Geochemistry of gabbro sills in the crust/ mantle transition zone of the Oman ophiolite: implications for the origin of the oceanic lower crust. Earth and Planetary Science Letters 146, 475e488. Kostopoulos, D.K., Murton, B.J., 1992. Origin and distribution of components in boninite genesis: significance of the OIB component, 60. Geological Society, London, Special Publications, pp. 133e154. Koepke, J., Feig, S.T., Snow, J., Freise, M., 2004. Petrogenesis of oceanic plagiogranites by partial melting of gabbros: an experimental study. Mineralogy and Petrology 146, 414e432. Koepke, J., Berndt, J., Feig, S.T., Holtz, F., 2007. The formation of SiO2-rich melts within the deep oceanic crust by hydrous partial melting of gabbros. Contributions to Mineralogy and Petrology 153, 67e84. Kranidiotis, P., MacLean, W.H., 1987. Systematics of chlorite alteration at the Phelps Dodge massive sulfide deposit, Matagami, Quebec. Economic Geology 82, 1898e1911. lu-Kus¸cu, G., Kus¸cu, I., Yılmazer, E., Güleç, N., Bayır, S., Demirela, G., ençaliog Gezer, G., Kuru, S., Kaymakcı, N., 2011. U-Pb and 40Ar-39Ar Geochronology and isotopic constraints on the genesis of copper-gold-bearing iron oxide deposits in the Hasançelebi District, Eastern Turkey. Economic Geology 106, 261e288. Lee, C.-T.A., Morton, D.M., 2015. High silica granites: Terminal porosity and crystal settling in shallow magma chambers. Earth and Planetary Science Letters 409, 23e31. Lepage, L.D., 2003. ILMAT: an excel worksheet for ilmenite-magnetite geothermometry and geobarometry. Computer Geoscience 29, 673e678. Lippard, S.J., Shelton, A.W., Gass, I.G., 1986. The Ophiolite of Northern Oman. The Geological Society Memoir no. 11. Blackwell Scientific Publications, Oxford, p. 178. Liu, Y., He, D., Gao, C., Foley, S., Gao, S., Hu, Z., Zong, K., Chen, H., 2015. First direct evidence of sedimentary carbonate recycling in subduction-related xenoliths. Scientific Reports 5, 11547. https://doi.org/10.1038/srep11547. Ludwig, K.R., 2001. Users Manual for Isoplot/Ex, rev. 2.49. A Geochronological Toolkit for Microsoft Excel, vol. 1a. Berkeley Geochronology Center, Special Publication, pp. 1e55. Lustrino, M., Anderson, D.L., 2015. The mantle isotopic printer: basic mantle plume geochemistry for seismologists and geodynamicists. In: Foulger, G.R., Lustrino, M., King, S.D. (Eds.), The Interdisciplinary Earth: A Volume in Honor of Don L. Anderson, vol. 514. Geological Society of America Special Paper, pp. 257e279. Malpas, J., 1979. Two contrasting plagiogranite associations from transported ophiolites in western Newfoundland: initial report. In: Barker, F. (Ed.), Trondhjemites, Dacites and Related Rocks. Elsevier, Amsterdam, pp. 465e487. Maniar, P.D., Piccoli, P.M., 1989. Tectonic discrimination of granitoids. Geological Society of America Bulletin 101, 635e643. Martin, H., Smithies, R.H., Rapp, R., Moyen, J.F., Champion, D., 2005. An overview of adakite, tonalite-trondhjemite-granodiorite (TTG), and sanukitoid: relationships and some implications for crustal evolution. Lithos 79, 1e24. McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., Taylor Jr., H.P., 1980. A neodymium, strontium, and oxygen isotopic study of the Cretaceous Samail Ophiolite and
1209
implications for the petrogenesis and seawater-hydrothermal alteration of oceanic crust. Earth and Planetary Science Letters 46, 201e211. McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., Taylor Jr., H.P., 1981. Sm-Nd, Rb-Sr and 18O-16O isotropic systematics in an oceanic crustal section; evidence from the Semail ophiolite. Journal of Geophysical Research 86, 2721e2735. McDonough, W.F., Sun, S., 1995. The composition of the Earth. Chemical Geology 120, 223e253. Michael, P.J., McDonough, W.F., Nielsen, R.L., Cornell, W.C., 2002. Depleted melt inclusions in MORB plagioclase: messages from the mantle or mirages from the magma chamber? Chemical Geology 183 (1e4), 43e61. https://doi.org/10.1016/ S0009-2541(01)00371-0. Middlemost, E., 1994. Naming materials in the magma/igneous rock system. Earth Science Reviews 37, 215e224. Moore, D.M., Reynolds, R.C., 1997. X-Ray diffraction and the identification and analysis of clay minerals. Oxford University Press, Oxford, p. 378. Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Academic, Amsterdam, p. 369. ISBN 978-94-010-7569-5. Nicolas, A., Ildefonse, B., Boudier, F., Lenoir, X., Ben Ismail, W., 2000. Dike Distribution in the Oman-United Arab Emirates Ophiolite, vol. 21. Kluwer Academic Publishers, pp. 269e287. Pallister, J.S., Hopson, C.A., 1981. Samail ophiolite plutonic suite: field relations, phase variation, cryptic variation and layering, and a model of a spreading ridge magma chamber. Journal of Geophysical Research 86, 2593e2644. Pallister, J.S., Knight, R.J., 1981. Rare-earth element geochemistry of the samail ophiolite near Ibra, Oman. Journal of Geophysical Research 86, 2673e2697. Papanastassiou, D.A., DePaolo, D.J., Wasserburg, G.J., 1977. Rb-Sr and Sm-Nd chronology and genealogy of mare basalts from the Sea of Tranquility. In: Lunar Science Conference, 8th Houston, Proceedings, vol. 2. Pergamon Press, New York, pp. 1639e1672. Pearce, J.A., 1980. Geochemical evidence for the genesis and eruptive setting of lavas from Tethyan ophiolites. In: Panayiotou, A. (Ed.), Ophiolites: Proceedings of the International Ophiolite Symposium, 1979, Cyprus, Nicosia, Cyprus Ministry Agriculture Natural Resources. Geological Survey Department, pp. 261e272. Pearce, J.A., 1996. A user’s guide to basalt discrimination diagrams. In: Wyman, D.A. (Ed.), Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration, vol. 12. Geological Association of Canada Short Course Notes, pp. 79e113. Pearce, J.A., Alabaster, T., Shelton, A.W., Searle, M.P., 1981. The Oman ophiolite as a Cretaceous arcbasin complex: evidence and implications. Philosophical Transactions of the Royal Society of London. Series A, Mathematical and Physical Sciences 300 (1454), 299e317. Pearce, J.A., Harris, N.W., Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology 25, 956e983. Peters, T.J., Kamber, B.S., 1994. Peraluminous, potassium rich granitoids in the Semail ophiolite. Contributions to Mineralogy and Petrology 118, 229e238. Peters, B.J., Day, J.M.D., 2014. Assessment of relative Ti, Ta and Nb (TITAN) enrichments in ocean island basalts. Geochemistry, Geophysics, Geosystems 15. https://doi.org/10.1002/2014GC005506. Pflumio, C., 1991. Evidences for polyphased oceanic alteration of the extrusive sequence of the Semail ophiolite from the Salahi block (northern Oman). In: Peters, T., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Petrology and Structural Geology 5. Kluwer Academic Publishers, Dordrecht, pp. 313e352. Pichou, J.L., Pouchoir, F., 1985. “PAP” procedure for improved quantitative microanalysis. Microbeam Analysis 20, 104e105. Pilet, S., Baker, B.M., Müntener, O., Stolper, M.E., 2011. Monte-Carlo simulations of metasomatic enrichment in the lithosphere and implications for the source of alkaline basalts. Journal of Petrology 52, 1415e1442. Pringle, E.A., Moynier, F., Savage, P.S., Jackson, M.G., Moreira, M., Day, J.M.D., 2016. Silicon isotopes reveal recycled altered oceanic crust in the mantle sources of ocean island basalts. Geochimica et Cosmochimica Acta 189, 282e295. Putirka, K., 2008. Thermometers and barometers for volcanic systems. In: Putirka, K., Tepley, F. (Eds.), Minerals, Inclusions and Volcanic Processes. Reviews in Mineralogy and Geochemistry 69, pp. 61e120. Powell, R., Powell, M., 1977. Geothermometry and oxygen barometry using coexisting iron-titanium oxides: a reappraisal. Mineralogical Magazine 41, 257e263. Python, M., Ceuleneer, G., 2003. Nature and distribution of dykes and related melt migration structures in the mantle section of the Oman ophiolite. Geology, Geophysics and Geosystems (G-Cubed) 4, 8612e8646. Rickwood, P., 1989. Boundary lines within petrologic diagrams, which use oxides of major and minor elements. Lithos 22, 247e263. Rioux, M., Bowring, S., Kelemen, P., Gordon, S., Miller, R., Dudás, F., 2013. Tectonic development of the Samail ophiolite: high-precision U-Pb zircon geochronology and Sm-Nd isotopic constraints on crustal growth and emplacement. Journal of Geophysical Research: Solid Earth 118, 2085e2101. Rioux, M., Garber, J., Bauer, A., Bowring, S., Searle, M., Kelemen, P., Hacker, B., 2016. Synchronous formation of the metamorphic sole and igneous crust of the Semail ophiolite: new constraints on the tectonic evolution during ophiolite formation from high-precision U-Pb zircon geochronology. Earth and Planetary Science Letters 451, 185e195. Roberts, N.M.W., Spencer, C.J., 2014. The zircon archive of continent formation through time. In: Roberts, N.M.W., Van Kranendonk, M., Parman, S., Shirey, S.,
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H. Joun et al. / Geoscience Frontiers 10 (2019) 1187e1210
Clift, P.D. (Eds.), Continent Formation through Time, vol. 389. Geological Society, London, Special Publications. https://doi.org/10.1144/SP389.14. Roberts, N.M.W., Thomas, R.J., Jacobs, J., 2016. Geochronological constraints on the metamorphic sole of the Semail ophiolite in the United Arab Emirates. Geoscience Frontiers 7, 609e619. Rollinson, H., 2009. New models for the genesis of plagiogranites in the Oman ophiolite. Lithos 112, 352. Rollinson, H., 2014. Plagiogranites from the mantle section of the Oman Ophiolite: models for early crustal evolution. In: Rollinson, H., Searle, M.P., Abbasi, I., AlLazki, A., Al Kindi, M.H. (Eds.), Tectonic Evolution of the Oman Mountains, vol. 392. Geological Society, London, Special Publications, pp. 247e261. Roedder, E., 1984. Fluid inclusions. Reviews in Mineralogy, 12. Mineralogical Society of America, p. 644. Rollinson, H., 2015. Slab and sediment melting during subduction initiation: granitoid dykes from the mantle section of the Oman ophiolite. Contributions to Mineralogy and Petrology 170, 1e20. Rudnick, R.L., Gao, S., 2003. Composition of the continental crust. In: Rudnick, R.L. (Ed.), The Crust. Elsevier, p. 170. Savelieva, G.N., Bortnikov, N.S., Bayanova, T.B., Ikorskii, S.V., Kamenskii, I.L., 2008. Sm-Nd and Rb-Sr isotopic systems and captured He and hydrocarbon gases as markers of melt sources and fluid regime under which the oceanic crust of the Mid-Atlantic Ridge was formed at 5 -6 N. Geochemistry International 46, 745e758. https://doi.org/10.1134/S0016702908080016. Searle, M.P., 1980. The Metamorphic Sheet and Underlying Volcanic Rocks beneath the Oman Ophiolite in the Northern Oman Mountains of Arabia (Unpublished) (PhD thesis). Open University, Milton Keynes. Searle, M.P., Malpas, J., 1980. Structure and metamorphism of rocks beneath the Semail ophiolite of Oman and their significance in ophiolite obduction. Transactions of the Royal Society of Edinburgh, Earth Sciences 71, 247e262. Searle, M.P., Malpas, J., 1982. Petrochemistry and origin of sub-ophiolite metamorphic and related rocks in the Oman Mountains. Journal of the Geological Society of London 139, 235e248. Searle, M.P., Cox, J., 1999. Tectonic setting, origin, and obduction of the Oman ophiolite. Geological Society of America Bulletin 111, 104e122. Searle, M.P., Cox, J., 2002. Subduction zone metamorphism during formation and emplacement of the Semail ophiolite in the Oman Mountains. Geological Magazine 139, 241e255. Searle, M.P., Ali, M., 2009. Structural and tectonic evolution of the Jabal Sumeini-Al Ain-Buraimi region, northern Oman and eastern United Arab Emirates. GeoArabia 14, 115e142. Searle, M.P., Waters, D.J., Garber, J.M., Rioux, M., Cherry, A.G., Ambrose, T.K., 2015. Structure and metamorphism beneath the obducting Oman ophiolite: evidence from the Bani Hamid granulites, northern Oman Mountains. Geosphere 11, 1812e1836. Shock, E.T., Helgeson, H.C., 1988. Calculation of the thermodynamic and transport properties of aqueous species at high pressures and temperatures: correlation algorithms for ionic species and equation of state prediction to 5 kb and 1000 C. Geochimica et Cosmochimica Acta 52, 2009e2036. Smith, P.M., Asimow, P.D., 2005. Adiabat_1ph: a new public front-end to the MELTS, pMELTS, and pHMELTS models. Geology, Geophysics and Geosystems (GCubed) 6, 1e8. https://doi.org/10.1029/2004GC000816. Q02004. Soret, M., Agard, P., Dubacq, B., Plunder, A., Yamato, P., 2017. Petrological evidence for stepwise accretion of metamorphic soles during subduction infancy (Semail ophiolite, Oman and UAE). Journal of Metamorphic Geology 35, 1051e1080. https://doi.org/10.1111/jmg.12267. Sobolev, A.V., 1996. Melt inclusions in minerals as a source of principle petrological information. Petrology 4, 209e220. Spencer, C.J., Cavosie, A.J., Raub, T.D., Rollinson, H., Jeon, H., Searle, M.P., Miller, J.A., McDonald, B.J., Evans, N.J., The Edinburgh Ion Microprobe Facility (EIMF), 2017. Evidence for melting mud in Earth’s mantle from extreme oxygen isotope signatures in zircon. Geology 45, 975e978. https://doi.org/ 10.1130/G39402.1. Spooner, E.T.C., 1977. Hydrodynamic model for the origin of the ophiolitic cupriferous pyrite ore deposits of Cyprus. Geological Society, London, Special Publications 7, 58e72. Stakes, D.S., Taylor, H.P., 1992. The northern Samail ophiolite: an oxygen isotope, microprobe and field study. Journal of Geophysical Research 97, 7043e7080. Stern, R.J., Reagan, M., Ishizuka, O., Ohara, Y., Whattam, S., 2012. To understand subduction initiation, study forearc crust: To understand forearc crust, study ophiolites. Lithosphere 4, 469e483.
Stormer, J.C., 1975. A practical two-feldspar geothermometer. American Mineralogist 60, 667e674. Stormer, J.C., 1983. The effects of recalculation on estimates of temperature and oxygen fugacity from analyses of multi-component iron-titanium oxides. American Mineralogist 68, 586e594. Stracke, A., Bizimis, M., Salters, V.J.M., 2003. Recycling oceanic crust: quantitative constraints. Geology, Geophysics and Geosystems (G-Cubed) 4, 1e33. https:// doi.org/10.1029/2001GC000223. Styles, M.T., Ellison, R.A., Phillips, E.R., Arkley, S., Schofield, D.I., Thomas, R.J., Goodenough, K.M., Farrant, A.R., McKervey, J.A., Crowley, Q.G., Pharoah, T.C., 2006. The geology and geophysics of the United Arab Emirates, vol. 2. Geology, Ministry of Energy, United Arab Emirates, Abu Dhabi, p. 351. Sun, 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.), 42. Geological Society, London, Special Publications, pp. 313e345. Tanaka, T., Togashi, S., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, T., Yuhara, M., Orihashi, Y., Yoneda, S., Shimizu, H., Kunimaru, T., Takahashi, K., Yanagi, T., Nakano, T., Fujimaki, H., Shinjo, R., Asahara, Y., Tanimizu, M., Dragusanu, C., 2000. JNdi-1: a neodymium isotopic reference in consistency with La Jolla neodymium. Chemical Geology 168, 279e281. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Oxford, 312. Thomas, R.J., Ellison, R.A., 2014. Geological Evolution of the United Arab Emirates. Over Six Hundred Million Years of Earth History. British Geological Survey, Keyworth, Nottingham, p. 287. Thomas, R.J., Styles, M.T., Goodenough, K.M., 2014. The UAE-Oman Ophiolite - a cross section through the Earth’s oceanic lithosphere on land. In: Thomas, R.J., Ellison, R.A. (Eds.), Geological Evolution of the United Arab Emirates: Over Six Hundred Million Years of Earth History. British Geological Survey, Ministry of Energy United Arab Emirates, pp. 153e176. Thorkelson, D.J., Breitsprecher, K., 2005. Partial melting of slab window margins: genesis of adakitic and non-adakitic magmas. Lithos 79, 25e41. Tilton, G.R., Hopson, C.A., Wright, J.E., 1981. Uranium-lead isotopic ages of the Semail ophiolite Oman, with applications to Tethyan ocean ridge tectonics. Journal of Geophysical Research 86, 2763e2776. Tippit, P.R., Pessagno, E.A., Smewing, J.D., 1981. The biostratigraphy of sediments in the volcanic unit of the Samail Ophiolite. Journal of Geophysical Research 862756e862762. Valley, J.W., 2003. Oxygen isotopes in zircon. Reviews in Mineralogy and Geochemistry 53, 343e385. https://doi.org/10.2113/0530343. Warren, C.J., Parrish, R.R., Waters, D.J., Searle, M.P., 2005. Dating the geologic history of Oman’s Semail ophiolite: insights from U-Pb geochronology. Contributions to Mineralogy and Petrology 150, 403e422. White, W.M., Klein, E.M., 2014. Composition of the oceanic crust. Treatise on Geochemistry 2, 457e496. Whitney, D., Evans, B., 2010. Abbreviations for names of rock-forming minerals. American Mineralogist 95, 185e187. Winchester, J.A., Floyd, P.A., 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology 20, 325e343. Wu, T.W., Kerrich, R., 1986. Combined oxygen isotope compositional studies of some granitoids from the Grenville, Ontario, Canada. Implications for source regions. Canadian Journal of Earth Sciences 23, 1420e1432. Xu, W., Zhang, H., Luo, B., Guo, L., Yang, H., 2015. Adakite-like geochemical signature produced by amphibole-dominated fractionation of arc magmas: an example from the late cretaceous magmatism in Gangdese belt, south Tibet. Lithos 232, 197e210. Xu, Z., Zheng, Y.F., Zhao, Z.F., 2018. Zircon evidence for incorporation of terrigenous sediments into the magma source of continental basalts. Scientific Reports 8, 178. https://doi.org/10.1038/s41598-017-18549-7. Zhao, Z., Zheng, Y., 2003. Calculation of oxygen isotope fractionation in magmatic rocks. Chemical Geology 193, 59e80. Zheng, Y.F., 1993a. Calculation of oxygen isotope fractionation in anhydrous silicate minerals. Geochimica et Cosmochimica Acta 57, 1079e1091. Zheng, Y.F., 1993b. Calculation of oxygen-isotope fractionation in hydroxyl-bearing silicates. Earth and Planetary Science Letters 120, 247e263. Zi, J.W., Cawood, P.A., Fan, W.M., Wang, Y.J., Tohver, E., 2012. Contrasting rift and subduction-related plagiogranites in the Jinshajiang ophiolitic mélange, southwest China, and implications for the Paleo-Tethys. Tectonics 31, TC2012. https://doi.org/10.1029/2011TC002937.
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Research Paper
Recycled oceanic crust as a source for tonalite intrusions in the mantle section of the Khor Fakkan block, Semail ophiolite (UAE) Hosung Joun a, Sotirios Kokkalas a, *, Stylianos Tombros b a b
Department of Earth Sciences, Khalifa University, 2533, Abu Dhabi, United Arab Emirates Department of Geology, University of Patras, 26442, Greece
a r t i c l e i n f o
a b s t r a c t
Article history: Received 1 November 2017 Received in revised form 10 June 2018 Accepted 11 September 2018 Available online 12 October 2018 Handling Editor: Sanghoon Kwon
Several types of felsic granitoid rocks have been recognized, intrusive in both the mantle and the crustal sequence of the Semail ophiolite. Several models have been proposed for the source of this suite of tonalites, granodiorites, trondhjemites intrusions, however their genesis is still not clearly understood. The sampled Dadnah tonalites that intruded in the mantle section of the Semail ophiolite display arctype geochemical characteristics, are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, show positive peaks for Ba, Pb, Eu, negative troughs for U, Ti and occur with low d18OH2O, moderate εSr and negative εNd values. They have crystallized at temperatures that range from w550 C to w720 C and pressure ranging from 4.4 kbar to 6.5 kbar. The isotopic ages from our tonalite samples range between 98.6 Ma and 94.9 Ma, slightly older and overlapping with the age of the metamorphic sole. Our field observations, mineralogical, petrological, geochemical, isotopic and melt inclusion data suggest that the Dadnah tonalites formed by partial melting (w10%e15% continuous or w12% batch partial melting), accumulation of plagioclase, fractional crystallization (w55%e57%), and interaction with their host harzburgites. These tonalites were the end result of partial melting and subsequent contamination and mixing of w4% oceanic sediments with w96% oceanic lithosphere from the subducted slab. This MORB-type slab melt composed from w97% recycled oceanic crust and w3% of the overlying mantle. We suggest that a possible protolith for these tonalites was the basaltic lavas from the subducted oceanic slab that melted during the initial stages of the supra-subduction zone (SSZ), which was forming synchronously to the spreading ridge axis. The tonalite melts mildly modified due to low degree of mixing and interaction with the overlying lithospheric mantle. Subsequently, the Dadnah tonalites emplaced at the upper part of the mantle sequence of the Semail ophiolite and are geochemically distinct from the other mantle intrusive felsic granitoids to the south. 2018, China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/ licenses/by-nc-nd/4.0/).
Keywords: Tonalites Felsic granitoids Recycled oceanic crust Source contribution Partial melting Melt inclusions
1. Introduction Different types of felsic intrusives, although volumetrically minor, are common in the oceanic crust and lithospheric mantle, varying from diorite and trondhjemite-to-tonalite and granite (Koepke et al., 2007; rock types generally called plagiogranites; Coleman and Peterman, 1975) but are mostly concentrated at the gabbro-sheeted dykes boundary (Lippard et al., 1986; Amri et al., 1996), reflecting either extreme fractionation in shallow melt
* Corresponding author. E-mail address: [email protected] (S. Kokkalas). Peer-review under responsibility of China University of Geosciences (Beijing).
lenses or re-melting of hydrothermally altered mafic crustal rocks (Amri et al., 1996; Brophy, 2008). Their variable composition, abundance across the oceanic lithosphere and geochemical characteristics make their formation and source still poorly understood (Lippard et al., 1986; Briqueu et al., 1991; Amri et al., 1996; Rollinson, 2009). Such felsic intrusions yield important insights into complex processes, such as the formation of earth’s continental crust (Rollinson, 2009), evolution of oceanic crust by magma intrusion, recycling of crustal components (Liu et al., 2015; Spencer et al., 2017; Xu et al., 2018) or oceanic crust, with or without minor amounts of continental input (Hofmann and White, 1982) and origin of MORB enrichment (Hemond et al., 2006). Crustal component may be incorporated into basalts by either shallow contamination or deep mixing (>120 km; Liu et al., 2015) carried by
https://doi.org/10.1016/j.gsf.2018.09.006 1674-9871/ 2018, China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the CC BY-NCND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
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subducting crust into the asthenospheric mantle, producing both soluble and insoluble materials at the slab-mantle interface (Spencer et al., 2017; Xu et al., 2018). However, the percentage of crustal component in these felsic melts is highly variable (Cox et al., 1999; Haase et al., 2015; Rollinson, 2015, and this study). The plagiogranites that intrude the Semail ophiolite can be categorized in three main types (Rollinson, 2009) based on their location in the ophiolite section: (i) early crustal plagiogranites that associated with partial melting in the upper gabbros (Lippard et al., 1986), (ii) late crustal plagiogranites and minor gabbro emplaced in the sheeted dyke and/or pillow lavas of the ophiolites (Lippard et al., 1986), and (iii) plagiogranites intrusive in the mantle section (Briqueu et al., 1991; Peters and Kamber, 1994; Cox et al., 1999). Alternatively, Spencer et al. (2017) used a geochemical distinction scheme into three suites of granitoids: a metaluminous crustal plagiogranite suite that ranges from granodiorite to tonalite (SiO2 of 54e72 wt.% and low K2O) and two other mantle intrusive suites, one metaluminous that ranges from granodiorite to tonalite (SiO2 of 67e75 wt.%), while the other is peraluminous with higher SiO2 values (72e78 wt.%). More specifically, in our broader study area of the Khor Fakkan block (Fig. 1), the plagiogranites commonly occur in the cumulate gabbros of the crustal sequence (Peters and Kamber, 1994) and in the harzburgites of the mantle sequence (Alleman and Peters, 1972). Previous studies on these plagiogranites have focused on the fact that potassic granitoids are relatively more frequent than in other parts of the Semail ophiolite (Lippard et al., 1986; Python and Ceuleneer, 2003). These granitoids are compositionally distinct from the typical oceanic plagiogranites, which have an extremely low K content (e.g., Coleman and Peterman, 1975). In contrast, the felsic granitoids in the Khor Fakkan block are quite potassic and are enriched in REE’s (Lippard et al., 1986; Cox et al., 1999; Amri et al., 2007). The genesis of plagiogranites is still poorly understood due to their numerous petrological and geochemical constraints. To explain their genesis, several models have been proposed: (1) magmatic differentiation, including fractional crystallization of primitive basaltic magma (e.g. Coleman and Peterman, 1975; Pallister and Hopson, 1981; Pallister and Knight, 1981; Lippard et al., 1986; Peters and Kamber, 1994) and silicate liquid immiscibility of a tholeiitic liquid (e.g. Dixon and Rutherford, 1979), (2) partial melting of sediments, gabbros and basalts (e.g. Malpas, 1979; Amri et al., 1996; Cox et al., 1999; Searle and Cox, 1999; Koepke et al., 2004, 2007; Searle et al., 2015) or pure sediment melting for the peraluminous type intrusive in mantle (Spencer et al., 2017) and (3) mixture of melted crust and sediments (Cox et al., 1999; Rollinson, 2009, 2015; Haase et al., 2015). Cox et al. (1999) and Amri et al. (2007) mainly discussed the formation of the felsic granitoids in the Khor Fakkan block. Based on whole-rock geochemistry and Rb-Sr and Sm-Nd isotopic compositions, Cox et al. (1999) proposed that the monzogranite-toleucogranite dikes were formed due to partial melting of a sedimentary protolith (i.e. Haybi and Hawasina complex). However, the dikes emplaced at higher structural levels (i.e., Dadnah area) are compositionally distinctive from the felsic granitoids of other localities in the Khor Fakkan block, as they thought to be the products of partial melting of both hybrid metasedimentary and oceanic crustal components (Cox et al., 1999). Experimental data on potassic granitoids (e.g., tonalites and trondhjemites) of Amri et al. (2007) suggest that there is a genetic relationship between them and a sedimentary protolith. The granitoids are metaluminous and are likely to be associated with both crustal and mantle components. Amri et al. (2007) concluded that these granitoids were derived by the partial melting of depleted mantle and then were metasomatized due to low-T hydrothermal alteration. The variation
in their composition (i.e., Dadnah felsic granitoids) are thought to be related to local differences in the physicochemical conditions during their formation. Later, more detailed work was done by Rollinson (2009, 2015) in which he set the broader context related to the felsic granitoids of the Semail ophiolite, sampling from six different locations, e.g., the Wadi Hemli, Wadi Hajr, Wadi Zikt, Wadi Fizh, Al-Bithnah and AlDadnah, and using data from Amri et al. (1996) about Maqsad area. Rollinson (2015) compared the major element trends between the crustal and mantle felsic granitoids in Oman. In contrast to Cox et al. (1999), these felsic granitoids are scarcely potassic, appear relatively enriched in MgO, Cr and Ni and Th, U, Ta, La and Ce than their crustal homologs and have REE patterns similar to the host gabbros of the crustal sequence (Rollinson, 2009, 2015). Rollinson (2009) suggested that these felsic granitoids were the products of mixing, contamination of crustal and mafic components. Our study focused on the felsic granitoids that intrude the mantle section in the Dadnah area of the Khor Fakkan block, where high degree of fractionation have been reported in previous studies (Peters and Kamber, 1994; Cox et al., 1999). We studied a smaller area with more detail comprising new petrographic, geochemical, isotope data, and geo-thermobarometry in order to identify, based on our first reported melt inclusion analysis, the least fractionated samples that can shed some more light on the formation and source contribution for this distinct granitoid melts. We discuss the analytical results and complement them with all recent work done in the Semail ophiolite (for further details see Rollinson, 2015 and references therein), in order to propose a tectono-petrogenetic model that could explain better the varied petrological and geochemical characteristics of the north structural block of the Semail ophiolite. 2. Geological setting 2.1. The Khor Fakkan block The eastern part of the Arabian platform is made up of the UAEOman ophiolite complex also known as Semail ophiolite, which is the largest and most widely studied ophiolite complex worldwide (Fig. 1). The Semail ophiolite, which is part of the Cretaceous “Alpine ophiolite chain”, forms a well exposed allochthonous arcuate thrust sheet, w500 km in length, 100 km wide and 10e15 km thick, parallel to the Gulf of Oman (Glennie et al., 1974; Lippard et al., 1986). The first reported ages for the plutonic rocks in the ophiolite came from Tilton et al. (1981), from zircons in the plagiogranites and the dates span in the interval between 95.9 Ma and 93.5 Ma, with most of them clustering around 95 Ma. This age range has been widely accepted as the mean age of the ophiolite. Recently, Rioux et al. (2016) reported higher precision UePb dates, suggesting that the main portion of the ophiolite crust formed between 96.12 Ma and 95.50 Ma. The Semail ophiolite thrusted towards SW over the Neo-Tethyan Hawasina basin for more than 400e500 km and finally was emplaced onto the PermianeMesozoic passive continental margin of the Arabian plate between 84 Ma and 72 Ma (Glennie et al., 1974; Tilton et al., 1981; Tippit et al., 1981; Cox et al., 1999). Petrologic and structural evidence suggests that the initial phase of the Semail ophiolite formed in an extensional setting, however, there is still debate concerning whether it was mainly formed at a mid-oceanic ridge (e.g. Coleman, 1981; Boudier et al., 1988; Ernewein et al., 1988; Nicolas, 1989; Pflumio, 1991) or at a supra-subduction zone (e.g. Searle and Malpas, 1980; Pearce et al., 1981; Alabaster et al., 1982; Searle and Cox, 1999, 2002). The UAE part of the Semail ophiolite comprises two intact ophiolite blocks, the Aswad and the Khor Fakkan block (Fig. 1). The
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1990; Searle and Ali, 2009). These were overthrust onto the Arabian passive margin, which comprises a w4 km thick carbonate sequence (Hajar Supergroup) that was deposited from midPermian to Cenomanian (Glennie et al., 1974). Recent studies have documented a more complex magmatic history for the Semail ophiolite than the standard layered Penrose ophiolite succession (Anonymous, 1972), involving younger, ultramafic-to-felsic island-arc type intrusive rocks (known as phase 2) that cross-cut the earlier magmatic rocks of MORB-composition, known as phase 1 (section 2.3, Grimes et al., 2013; Goodenough et al., 2014, for more details see also section 2.3; Thomas et al., 2014). In the Khor Fakkan block, the phase 2 magmatic rocks mainly appear at the uppermost part of mantle sequence and are restricted in the eastern part of the block. The phase 2 rocks include gabbros, wehrlites and dunites, which cut both the harzburgites of the mantle sequence and the layered gabbros of the crustal sequence and are cut by later sills, dikes and pods of quartz-diorite and tonalities (Goodenough et al., 2014). 2.2. Metamorphic sole and formation of felsic granitoids
Figure 1. Simplified geological map of the UAE ophiolite showing the main tectonic structures and the distribution of mantle and crustal rocks (modified from Gnos and Nicolas, 1996; Thomas et al., 2014). Our study area is shown with a red rectangle (see Fig. 2 for details), and Rollinson’s (2015) sampling area with a cyan colored blue rectangle (R).
Khor Fakkan block occupies the northernmost part of the Semail ophiolite, covering an area of approximately 750 km2, and mainly exposes the mantle sequence of the ophiolite. The large-scale NEtrending Wadi Sidr and NW-trending Wadi Ham fault zones bound the north and south margin of this block, respectively (Fig. 1). The southeast flank is juxtaposed against the Masafi-Ismah and Bani Hamid metamorphic thrust sheets that represent the metamorphic sole of the Semail ophiolite (Fig. 1; Searle, 1980; Goodenough et al., 2006; Searle et al., 2015). The Khor Fakkan ophiolite block forms the structurally uppermost thrust sheet of the nappe stack. Below the mantle sequence and the metamorphic sole, several thrust nappes were stacked during the obduction process. These allochthonous units include Neo-Tethyan oceanic sediments (Hawasina Complex), deep-ocean and subduction-related sediments, volcanics and mélange (Haybi Complex) and proximal carbonate slope-facies rocks (Sumeini complex) (Glennie et al., 1974; Lippard et al., 1986; Bechennec et al.,
The Semail metamorphic sole is a thin (10 m to w500 m thick) amphibolite to granulite sliver of oceanic crust with variable proportions of basaltic and sedimentary material that was strained and metamorphosed, by heat transfer from the hot upper plate mantle toward the slab, during the initial stage of intra-oceanic subduction (Soret et al., 2017). The protoliths of metamorphic sole were the TriassiceJurassic and Early Cretaceous Haybi volcanics, limestones, Mn-rich chert, melange and other deep marine sediments of Hawasina complex (Ghent and Stout, 1981; Searle and Cox, 2002). The uppermost part of the sole, directly below a thick mylonitic peridotite comprises high-temperature (HT) amphibolite-togranulite facies metabasalt, while below that there is the low temperature (LT) sole which is mainly composed of greenschist facies metachert with imbrications of metabasalts and metasedimentary rocks (Searle and Malpas, 1980; Ghent and Stout, 1981; Boudier et al., 1988; Soret et al., 2017). The estimated P-T for the HT sole are on the range of 700e900 C and 0.4e1.3 GPa and decreases in the LT sole to w500 C and 0.45e0.55 GPa with an inverted thermal gradient, grading steeply from granulite to greenschist facies (>1000 C/km, Searle and Malpas, 1982; Hacker and Mosenfelder, 1996; Searle and Cox, 2002). The age of the metamorphic sole is considered almost coeval or slightly postdates the time of the Semail ophiolite obduction (Soret et al., 2017). Cox et al. (1999) assumed that HP metamorphism was roughly contemporaneous to the formation of the metamorphic sole throughout the Semail ophiolite, based on contiguous metamorphism ages (i.e., 96 2 Ma, Ar/Ar of As Sifah eclogite age, 93.5 0.1 Ma, Ar/Ar sole metamorphism mean age of Wadi Tayin area of Cox et al. (1999)). However, the published ages of the sole were estimated by the cooling ages of minerals (e.g., hornblende, biotite, and garnet) and display various age ranges. Indeed, the K/Ar dating on a metamorphic hornblende separated from the amphibolite yielded ages in the broad range 101e89 Ma and cluster near 98 Ma (Gnos and Peters, 1993). Recent data from sole peak metamorphism (Rioux et al., 2016) reported dates between 96.16 Ma and 94.82 Ma, suggesting that metamorphic rocks were formed either prior to or during the formation of the oceanic crust. 2.3. Magmatism in Semail ophiolite In the Semail ophiolite two distinct phases of magmatism have been described. An earlier sea-floor spreading magmatism termed as phase 1, composed of gabbros, sheeted dykes and pillow lavas and a later phase 2 more hydrous magma suite, composed of
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wehrlites, clino-pyroxenites and a complex network of gabbroic, diorite and tonalite intrusions, recording the initial stages of arc magmatism (Styles et al., 2006; Goodenough et al., 2014). The different magmatic phases are marked by the lavas and the relevant magmatic rocks found within the crustal and mantle sequence. Three major lavas were distinguished based on their geochemical affinities: (1) V1 lavas (i.e., Geotimes) have MORB-like affinity and are formed at spreading axis, (2) V2 lavas (i.e., Lasail or Alley) are formed off-axis and have arc-like petrochemical signatures (Alabaster et al., 1982; Ernewein et al., 1988) and (3) V3 lavas (i.e., Salahi) are enriched in incompatible elements and range compositionally from alkaline to transitional tholeiites. The V3 lavas formed by off-axis magmatism and postdate the other two lavas (Alabaster et al., 1982; Lippard et al., 1986). Furthermore, due to the fact that V2 and V3 lavas are more abundant in the Khor Fakkan block i.e., phase 2 magmatism, the V1 lavas are relatively rare (Goodenough et al., 2014; Thomas and Ellison, 2014). The volcanic rocks related to the phase 2 magmas are characterized by higher amount of plagioclase and clinopyroxene, with variable amounts of olivine, orthopyroxene, hornblende and quartz and are considered to form by melting of a hydrous magma in a SSZ setting (Goodenough et al., 2014). Haase et al. (2016) have grouped the plutonic rocks from the crustal sequence of Oman ophiolite into two groups, with each one mimic the geochemical signatures of the V1 and V2 lavas, respectively. V2 magmas are more voluminous in the northern part of the ophiolite showing the stronger input of slab component, regardless if that was sedimentary or not. 3. Sampling and analytical methods 3.1. Field observations and sampling The study area, which is located between the Dadnah town and east of Wadi Zikt, is structurally comprising the uppermost part of the mantle sequence and the Moho transition zone (MTZ) of the Khor Fakan block of the Semail ophiolite (Fig. 2).
The felsic granitoids in this area appear as veins, sills, and dikes, extending up to several hundreds of meters laterally and few meters in thickness (Fig. 3aec). Almost all of the examined granitoids are hosted by serpentinized harzburgites. They appear fine-tomedium-grained with typical granitic textures, although a few of them display pegmatitic textures. They usually contain up to 20 cm in length mafic enclaves and xenoliths in their rock body (i.e. location D6, Figs. 2 and 3a). At the same location two different granitoid types can be distinguished. There is a grayish-colored coarser grained, intense foliated and more sheared type with hornblende clinopyroxene that comprises the biggest part of the sill, while in the eastern part there is a composite body together with a less foliated, fine-to medium grained, less deformed leucocratic rock type containing abundant biotite crystals. The leucocratic granitoid intrudes in places the foliated grayish type, suggesting that probably was a later intrusive phase. The majority of the felsic sills and dikes observed in this area are related to pre-existing ductile or reactivated brittle-ductile faults, as their margins in many cases controlled by ductile shear zones or hydrothermally altered striated surfaces. Most of the dikes are oriented subparallel-parallel to the fault/fracture planes (Fig. 2, stereonets). Dikes exhibit various degree of deformation showing development of quartzose mylonites, gneissic fabrics or biotite grain-shape preferred orientations often aligned almost parallel to the dike margins due to intense ductile shearing. Many of the meter-to-hectometre scale in length felsic granitoids are feeder dikes, developing around them cm-scale secondary veins. No clear crosscutting relationships among the meter-to-hectometre scale felsic dikes was observed, while the cm-scale veins frequently crosscut each other, suggesting coeval development for these secondary veins. The felsic dikes in the study area trend between a frequent WNWeESE orientation (95 e105 ) and two less prominent orientations NEeSW (30 e50 ) and NWeSE (300 e310 ) (Fig. 2, stereonet). This latter trend follows the map-scale NW-trending fault/ fracture set (Fig. 2, map) and is frequent among the meter-scale
Figure 2. Aerial view (Google Earth image) of the study area (red rectangle of Fig. 1) at the west side of the Dadnah city in the Khor Fakkan block. The yellow arrows indicate the sampling locations and the red lines show the main map-scale fault trends in the study area. Stereonets of the tonalite dikes and fault planes are shown as great circles on the right side of the figure.
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Figure 3. (a) Field photograph of a tonalite sill that outcrops at locality D6. The sill contains xenoliths of the host harzburgite and is offset by later stage brittle faults. (b) Field photograph of the same tonalite sill, showing in detail the xenoliths contained in the sill. The xenoliths appear partially resorbed and recrystallized with intense mylonitic textures, comprising, olivine, ortho- and clino-pyroxene and plagioclase. (c) Sample from the tonalite sill composed of plagioclase, quartz and hornblende that is cross-cut by late quartz veins. (d) Representative thin section image of tonalite composed of plagioclase (Pl), quartz (Qz), hornblende (Hbl) and clinopyroxene (Cpx). At the left bottom corner the plagioclase crystal contain relic plagioclase (sample PGD2, XPL, 4x). The XPL colors of minerals are not typical due to the fact that thin sections are thicker and were constructed for fluid and melt inclusions analyses. (e) Thin section image of a tonalite medium-to-intense mylonitized (sample PGD7_2, XPL, 2x), with clinopyroxene and plagioclase porphyroblasts rotating in a foliated matrix of quartz, plagioclase, alkali feldspar and mica. (f) Cathodoluminescence (CL) image of the tonalite that contains xenoliths of the host harzburgite. The plagioclase of the host harzburgite (purple Pl) appears with green CL colors and intergrowths with the non-luminescent clinopyroxene (Cpx). The plagioclase of the tonalite (cyan Pl) appears with blue CL colors and replaces the plagioclase from the harzburgite, and very commonly resorbs these plagioclases (sieve texture). (g) Backscattered electron (BSE) image of the tonalite plagioclase. The plagioclases appear with regular zonation with cores composing of andesine (An44e56) and rims of oligoclase (An25e38). At the down-left corner partially resorbed plagioclases of higher An contents (An53e56), are also shown. At the up-right corner the plagioclase composition is shown on the ternary Ab-Or-An plot (Ab: Albite, Ol: Oligoclase, And: Andesine, La: Labradorite, By: Bytownite, An: Anorthite, Ano: Anorthoclase). Solid pink squares symbolize the cores, deep yellow squares the rims, and light ochre squares the relict plagioclase cores (Mineral abbreviations are after Whitney and Evans, 2010).
felsic dikes. The mesoscale fault orientations cluster around a NEeSW trend (w30 e45 ), while the WNWeESE trend (110 e120 ) display sinistral shear motion along the fault surfaces and is the dominant map-scale trend in the area between Wadi Zikt and the village (Fig. 2, map).
3.2. Analytical methods In total, thirteen granitoid samples as well as few representative host harzburgite samples (e.g., OPD2) were collected from seven sites in the study area (Fig. 2, Table 1). From these samples 25 thin
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Table 1 Major and trace element chemistry of the tonalites. Melt inclusions are shown also on table. Samples
PG
OP
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
PG
D1
D2
D2_2
D3_1
D3_2
D4_Mi
D4_Ma
D5
D6_1
D7
D7_2
M1
M9
M12
65.99 13.76 6.12 0.12 2.77 7.05 1.11 0.27 0.91 0.22 2.13 100.45
72.90 13.90 0.73 0.01 0.28 0.21 3.42 6.54 0.02 0.06 0.61 98.68
73.59 12.48 1.64 0.02 2.18 1.52 3.88 0.99 0.11 0.03 2.40 98.84
74.57 13.57 0.82 0.02 0.42 0.80 4.36 2.73 0.06 0.06 1.50 98.91
73.41 13.11 1.87 0.02 1.51 3.15 2.91 0.80 0.17 0.01 2.21 0.02 <0.01 99.37
75.00 12.97 1.15 0.02 0.91 2.81 3.29 1.31 0.11 0.03 1.08 98.68
76.99 12.94 0.98 0.01 0.50 2.94 4.01 0.26 0.09 0.02 0.96 0.01 0.01 99.90
75.92 12.29 1.58 0.04 1.77 2.17 2.10 0.45 0.26 <0.01 3.03 0.01 <0.01 99.72
75.18 13.49 1.38 0.11 0.39 0.88 5.01 1.05 0.01 0.03 1.46 98.99
76.40 13.52 1.09 0.01 0.22 1.34 6.17 0.25 0.03 0.02 0.22 <0.01 0.01 99.32
73.45 15.78 0.01 0.00 0.00 0.00 3.59 4.01 0.02 0.01 2.71 0.02 0.01 99.58
74.98 14.78 1.32 0.00 0.02 0.03 5.47 1.27 0.01 0.01 2.01 0.02 <0.01 100.00
76.45 13.29 0.65 0.00 0.00 0.01 7.84 0.02 0.00 0.01 1.45 0.01 <0.01 99.72
-
28.8 63.7 0.4 2.9
56.0 27.6 0.2 3.8
47.9 44.7 0.4 1.1
57.3 21.2 0.1 6.5
55.7 27.4 0.2 4.8
62.4 14.2 0.1 5.2
64.8 14.2 0.1 7.7
56.9 33.7 0.2 2.6
59.7 32.7 0.1 1.9
<1 18 85 20.17 7.26 371.00 37.00 245.00 12.26 0.10 192.00 1.34 0.85 1.58 4.52 1.03 0.18 10.28 24.61 3.75 18.36 5.34 2.90 3.86 1.41 7.19 1.25 4.66 0.47 3.55 0.46 0.51 3.6 0.28 0.21 0.66 0.89 2.17 5.47 28.33 29.71 64.2 0.02
<1 <1 <5 2.66 4.05 5.00 6.00 9.00 27.87 0.43 9.00 0.78 1.20 2.26 8.66 1.23 3.45 16.66 39.78 4.34 14.41 3.79 1.16 3.34 0.96 4.51 0.70 2.75 0.31 2.48 0.33 1.04 62.7 0.64 0.50 0.81 1.08 1.25 0.00 48.19 8.62 60.3 0.81
<1 1 7 4.31 28.24 152.00 18.00 32.00 11.13 0.61 583.00 0.75 1.17 0.82 5.24 0.49 8.09 65.32 117.16 11.77 38.30 5.45 3.40 7.66 1.38 2.93 0.46 2.68 0.19 1.58 0.23 0.97 6.9 1.74 3.05 1.93 0.97 2.55 0.01 10.42 38.83 84.0 0.19
<1 4 <5 4.33 130.15 12.00 24.00 10.00 21.03 0.34 11.00 0.46 0.97 1.65 6.45 1.61 2.10 6.45 15.98 1.89 6.88 2.43 1.49 2.39 0.69 3.03 0.49 2.06 0.24 2.19 0.27 1.05 26.4 0.39 0.22 0.66 1.05 2.36 0,02 81.96 46.97 67.0 10.85
4 <1 2 12 6.58 20.09 261.00 8.00 59.00 6.67 0.46 142.00 0.72 0.29 0.46 5.78 1.16 6.50 46.08 85.57 9.00 32.56 5.89 3.43 10.69 1.95 2.92 0.47 3.68 0.18 1.79 0.22 0.77 4.2 1.14 3.34 1.88 0.97 2.07 0.02 14.59 22.39 50 76.2 0.08
<1 1 8 7.67 20.10 266.00 8.00 46.00 7.93 0.46 196.00 1.29 1.05 0.92 3.66 1.36 10.70 67.72 121.73 12.52 44.88 7.10 3.31 14.57 2.48 2.39 0.36 4.19 0.13 1.50 0.18 0.76 24.2 1.39 1.90 2.62 0.96 1.61 0.01 13.20 17.04 75.8 0.08
6 <1 <1 <5 3.23 80.40 275.00 9.00 79.00 3.05 0.11 154.00 0.64 0.19 0.85 3.09 1.97 4.47 30.58 58.25 5.93 20.90 3.62 3.18 6.88 1.22 1.02 0.15 1.92 0.06 0.69 0.09 0.74 17.3 1.23 3.26 2.02 0.99 3.10 0.01 11.80 60.53 16.8 66.9 0.29
6 <1 1 8 5.60 15.26 236.00 11.00 124.00 10.84 0.88 173.00 0.69 0.70 0.31 2.26 1.62 0.43 3.75 6.43 0.64 2.31 0.52 2.66 1.00 0.19 0.53 0.10 0.58 0.06 0.57 0.08 1.04 25.2 1.06 0.49 1.07 0.95 7.13 0.27 78.13 109.08 16.7 81.6 0.06
9 5 <5 3.52 81.98 540.00 40.00 28.00 19.30 5.79 23.00 1.27 2.62 2.06 3.01 1.83 5.04 6.58 16.14 1.73 6.24 2.83 1.86 2.85 0.92 4.84 0.78 3.06 0.39 3.56 0.43 1.08 8.2 0.34 0.24 0.68 1.10 2.36 0.00 39.96 51.45 52.8 0.15
4 2 <1 5 4.73 23.84 152.00 28.00 26.00 17.30 1.06 304.00 2.02 1.58 0.81 3.83 0.69 9.15 70.06 120.83 12.01 41.99 6.00 3.94 16.87 2.74 1.70 0.24 4.31 0.09 1.15 0.14 0.89 26.2 1.70 4.52 1.92 0.95 1.94 0.00 7.16 118.93 22.5 44.4 0.16
42.4 51.7 0.1 2.6 0.77 -
55.8 37.9 0.1 2.7 0.71 -
60.1 36.3 0.1 1.1 0.72 -
-
55.7 -
-
wt.% 66.92 SiO2 Al2O3 19.43 Fe2O3t 0.62 MnO 0.01 MgO 0.56 CaO 1.16 Na2O 8.93 K2O 1.60 TiO2 0.01 P2O5 0.05 1.06 LOI/H2O Cl F Total 100.35 C.I.P.W. normative Qz 36.4 An 57.8 Or 0.3 Ab 0.5 ppm B Be <1 Sc 3 V <5 Ge 1.32 Rb 39.91 Sr 187.00 Y 24.00 Zr 36.00 Nb 29.85 Cs 0.22 Ba 282.00 Hf 1.36 Ta 1.26 W 1.19 Pb 2.79 U 1.45 Th 2.17 La 17.96 Ce 36.05 Pr 3.66 Nd 11.67 Sm 2.81 Eu 2.15 Gd 2.67 Tb 0.73 Dy 3.59 Ho 0.59 Er 2.35 Tm 0.26 Yb 2.05 Lu 0.28 ASI 0.95 LREE/HREE 7.8 (La/Sm)N 0.93 (La/Yb)N 0.65 (Gd/Yb)N 0.79 Ce/Ce* 1.02 Eu/Eu* 3.08 Ti/Ti* 0.01 Pb/Pb* 29.14 Y/Ho 40.41 Cl/B Mg# 78.2 Rb/Sr 0.21
PG ¼ Tonalite; OP ¼ Host harzburgite; M(1,9,12) ¼ Melt inclusion glass.
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sections, double polished and blue epoxy impregnated were constructed (50 mm thick, 24 mm 46 mm, Wagner Petrographic Lab, USA). An Olympus BX51 petrographic microscope was used for examining the thin sections. We should note that these thin sections were constructed primarily for fluid and melt inclusion analyses and due to that they do not display the typical XL mineral colors, as shown in Fig. 3d and e. The same samples were prepared by removing exterior surfaces with a rock splitter. Firstly, we separated the xenoliths of the host harzburgite from the felsic granitoids. Mineral separation from the granitoids involved handpicking after examination under a binocular microscope to ensure purity of better than 98%, based on their paragenetic relations and textural equilibrium. In this way, only contiguous, non-tectonized and unaltered crystals of pyroxene, plagioclase, hornblende, biotite, quartz and calcite were selected. In order to ensure primary isotopic signature, before the isotopic analysis the samples were heated between 100 C and 200 C to remove the involvement of secondary fluid inclusions. Mineral compositions were determined using a JEOL 8900 Superprobe equipped with wavelength, energy dispersive and back-scattered detectors and a xClent system for ppm-level resolution at the Microprobe Center of the Department of Earth and Planetary Sciences Department, at McGill University. Operating conditions included an acceleration voltage of 15 kV, a beam current of 10 nA, and a counting time of 20 s. Scans in WDS mode were also used to detect trace element contents. Standards used were natural olivine, clino- and ortho-pyroxene, amphibole, chlorite, epidote, K-feldspar, plagioclase, barite, muscovite, magnetite, ilmenite, pyrite, chalcopyrite and galena, plus the native metals Ag, Sb, Au, Se, Ta and Cd. ZAF corrections were made with proprietary JEOL software. A minimum of ten analyses were obtained from each sample and three from each grain. Scanning Electron Microscopy (SEM) was also used in order to determine the chemical composition of certain minerals of the felsic granitoids. These were carried out at the Laboratory of Electron Microscopy and Microanalysis, University of Patras (Greece), using a Jeol JSM-6300 Scanning Electron Microscope, equipped with EDS and WDS and a THETA software. The operating conditions were 15 kV accelerating voltage and 3.3 nA beam current with 4 mm diameter beam. Also, at the same laboratory X-ray power diffraction patterns of oriented granitoid samples were obtained using a Bruker D-8 Focus diffractometer, with Ni-filtered Cu Ka radiation (with a 2q angle of 3 e70 ), using the XRD-data method of Moore and Reynolds (1997). The samples already used for probe analyses, prior to the microthermometric study of the melt inclusions were examined by cathodoluminescence (CL) using the Luniscope CL Reliotron spectrometer (Laboratory of Electron Microscopy and Microanalysis, University of Patras, Greece), equipped on a Quanta 200F ESEM with 45-s scanning time. The applied acceleration voltage and current were 15 kV and 120 nA and the CL images were collected with a resolution of 1280 960 pixels and 256 gray level. Major elements from the least altered felsic granitoid samples, as their Ishiwaka Alteration Index is 8.5%, were detected by inductively coupled plasma (ICP), using a lithium tetraborate fusion, at the Actlabs laboratory in Ontario, Canada. The concentrations of trace elements from the felsic granitoid samples were obtained by Agilent 7700s inductively coupled plasma mass spectrometry (ICP-MS), at Chonnam National University, South Korea. JG-2 gabbro and JGb-2 granite, obtained from the Geological Survey of Japan were used as standards. Repeated measurements of the data are between 0.05% and 1.4% relative for major elements, 7.1% relative for TiO2, 0.06% and 0.1% relative for REE. To quantify the Ce, Eu, Ti, Pb anomalies, we have used the equations of Taylor and McLennan (1985), Bau et al. (1996), Peters and Day (2014) and
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the normalizing values of McDonough and Sun (1995). The results of whole-rock analyses are presented in Table 1. Furthermore, we have microthermometrically analyzed only primary melt inclusions on one doubly polished 50e100 mm thick section, since primary inclusions were infrequent and showed signs of metastability in the other sections. The melt inclusions in the granitoid sample were completely crystallized in the quartz and plagioclase grains, and so they were homogenized using the conventional horizontal cold-seal pressure vessel technique. Hightemperature measurements were obtained with a MDSG600 heating-freezing stage, at the USGS Denver Inclusion Analysis Laboratory. Heating-freezing rates range from 0.1 to 130 C/min. The stage was calibrated at the critical temperature of H2O (374.1 C), the melting points of ultrapure NaCl (801 C), and silver (961 C). The accuracy and precision of homogenization temperatures of melt inclusions are better than 5 C. From room temperature to 400 C, a heating rate of 30 C/min was used. The rate was slowed to 10 C/min from 400 C to 800 C and then held at that temperature for thirty minutes before proceeding with homogenization runs. Melt inclusion glasses were analyzed using the EPMA difference technique (Table 1), with a Cameca SX50 and an SX100 electron microprobe at the same laboratory equipped with both WDS and EDS. Quantitative analyses were conducted using WDS with silicate, oxide, and phosphate standards. The data were corrected according to PAP methodology (Pichou and Pouchoir, 1985) using a vendor supplied software. Isotopic compositions of oxygen, hydrogen, carbon and silicon analyzed using a MAT-253 stable isotope ratio mass spectrometer. Analyses were performed at the Chinese Academy of Geological Sciences (CAGS), Beijing, China. Samples used for isotopic analyses were cross checked under the CL microscope to avoid any possible secondary isotopic signature due to alteration. Oxygen and hydrogen were released from pyroxene, plagioclase, hornblende, biotite, quartz and calcite using the BrF5 extraction technique (Clayton and Mayeda, 1963; Friedman and O’Neil, 1977). Silicon was measured using the SiF4 technique of Ding (2004), whereas carbon and oxygen in calcite were liberated as CO2 after Clayton et al. (1972). The isotopic ratios are reported in standard d notation per mil relative to SMOW for oxygen and hydrogen, NBS-28 for silicon and PDB belemnite for carbon. Analytical precision was better than 0.2& for d18O and d13C, 2& for dD and 0.1& for d30Si. We have used the AlphaDelta software of Beaudoin and Therrien (2009) to compute the isotopic fractionation factors and equilibrium temperatures. The same minerals were analyzed for their Rb/Sr and Sm/Nd isotopic compositions. The method used for Sr and Nd isotopic analysis is described by Papanastassiou et al. (1977) and De Paolo (1980). Samples were analyzed using a VG-354 ionization mass spectrometer at the Modern Analysis Center, Nanjing University, Nanjing, China. The total procedure blanks for Rb and Sr were 20 pg and 50 pg and for Nd and Sm 1 ng and 0.2 ng, respectively. Analytical precision for the 87Sr/86Sr and 143Nd/144Nd ratios was better than 0.00001 (all errors are reported at 2s absolute). The standards used for 87Sr/86Sr, 146Nd/144Nd and 143Nd/144Nd ratios were NBS987 (0.710223 8) and JNdi-1 (0.7129 5 and 0.51112 5, Tanaka et al., 2000). The normalizing factors used to correct for isotopic fractionation of Sr and Nd were 86 Sr/88Sr ¼ 0.1194 and 146Nd/144Nd ¼ 0.7219, respectively. Isotopic data for Nd and Sr are presented in terms of the epsilon notation, i.e., εSr and εNd (De Paolo, 1980). All ages were calculated using the decay constants for 87Rb and 47Sm which are 1.42 1011 and 6.54 1012/yr, respectively. Ages were calculated using the leastsquare regression technique. The Rb-Sr and Sm-Nd isochron ages were calculated using the York’s model 3 fit in Isoplot (Ludwig, 2001). We have used 2s errors of 0.5% for the 147Sm/144Nd
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ratios, 0.0004% for the 143Nd/144Nd ratios, 0.055% for the 87 Rb/86Sr and 0.0055% for the 87Sr/86Sr ratios. 4. Petrography and mineral chemistry 4.1. Petrography The felsic granitoids based on crystal size are fine and mediumgrained with hypidiomorphic-to-allotriomorphic primary granular textures and preserve deformational textures, including granophyric or mylonitic (Fig. 3d and e). In addition, sutured grain boundaries in most minerals, undulose grain extinction and wider grains of quartz are common because of ductile deformation. The felsic minerals of the granitoids are composed of plagioclase (70e60 vol.%), quartz (20e15 vol.%) and K-feldspar (5 vol.%). The mafic minerals (5e15 vol.%) are mostly hornblende, clinopyroxene and biotite. Epidote, titanite, titanomagnetite and ilmenite (5e7 vol.%) occur as accessory phases. In few samples, hydrothermal alteration resulted in the formation of chlorite, muscovite, serpentine and syntaxial veinlets comprising quartz, calcite, siderite, muscovite, and chlorite (Fig. 3d). A distinctive feature of the PGD1 to PGD4 samples is that they contain xenoliths of their host harzburgite, which in places comprise up to 50 vol.% of the felsic granitoids. These partially resorbed and recrystallized xenoliths with intense mylonitic textures, comprise ortho- and clino-pyroxene porphyroblasts (60e70 vol.%) and plagioclase, with accessory minerals magnetite and chromite (Fig. 3a, b and f). Perthite clino- from ortho-pyroxene wispy exsolution textures and plagioclase mottled by clinopyroxene occur in the xenoliths contained in the sample PGD3. The exsolution lamellae are planar, and their orientation is crystallographically controlled along the (110) cleavage of orthopyroxene. Corona textures with cores of olivine mantled by rims of clinopyroxene are common. Sometimes clinopyroxene, in turn, is rimmed by hornblende. In the PGD5 to PGD7 samples, the recrystallized pyroxene-bearing xenoliths are surrounded by finegrained matrix comprising of quartz, hornblende biotite and muscovite. The wall-rocks are harzburgites, comprising relict olivine, othroand clino-pyroxene, and a hydrothermal assemblage that includes talc, serpentine, chlorite, muscovite and minor quartz. Serpentinization is common among these rocks. Magnetite and chromite occur as accessory phases. Olivine and ortho- and clino-pyroxene initially alter towards biotite, which in turn alters to Fe-rich chlorite and muscovite. 4.2. Mineral chemistry The plagioclases from the felsic granitoids are interstitial, subhedral to anhedral composing of either fine-grained matrix or coarse-grained squat crystals. They appear with regular zonation with cores composing of andesine (An44e56) and rims of oligoclase (An25e38) (Fig. 3g), where the cores are enriched in the anorthite component by w15%e20%. The K-feldspars are uncommon and always appear with allotriomorphic, medium-grained crystals intergrown with plagioclases, forming intergranular textures. Their chemistry is persistent and characterizes them as orthoclase (Or100). In rare cases, relict plagioclases of higher An contents (Añ53e56, Fig. 3f and g) were observed under CL microscopy, characterized by resorption texture. The mafic minerals from the felsic granitoids analyzed are pyroxene, hornblende and biotite (e.g., sample PGD6_1). Clinopyroxenes appear with the same crystallographic habit as plagioclases, either as a fine-grained matrix or with coarse-grained crystals (Fig. 3d). The clinopyroxenes analyzed are classified as diopside-to-
augite (En59.8e61.8 and Wo27.7e28.2). Their Mg# (Mg# ¼ 100 Mg/ (Mg þ Fe2þþFe3þþMn)) values range from 0.84 to 0.855, and are relatively enriched in Al2O3 (up to 0.11 apfu; based on 6 oxygens) and depleted in TiO2, Cr2O3 and NiO (0.01 apfu). Hornblende form hypidiomorphic crystals showing intensive pleochroism in shades of green. Some crystals appear with light colors or even colorless and intergrown with plagioclases irregularly distributed due to intensive tectonism (Fig. 3d). Hornblende is classified as tschermatitic-hornblende towards tschermakite, with Mg# values ranging from 0.73 to 0.81. Biotite appears with the same textural characteristics to hornblende and based on its Sitotal and Mg# values (w0.45) is categorized as biotite. Magnetite occurs as medium-grained accessory phase, is rimmed by plagioclase and/or clinopyroxene and on the basis of its Fe, Cr, and Ti composition is almost pure magnetite. Mixtures of fine-grained serpentine, muscovite, chlorite and minor albite and calcite are the main alteration products of the felsic granitoid minerals. Based on their Si and Al content, the analyzed muscovites are rich-muscovites-to-low-phengites. Chlorite appears fine-grained and fibrous. Chlorite is clinochlore due to its anomalous polarization colors and its Al, Si and Fe content. Serpentine is classified as solid solution of w25% sepiolite and w75% serpentine (e.g., sample PGD2_2). Olivine from the xenoliths occurs as a fine-grained matrix as well as with medium-grained crystals that are highly altered towards serpentine. The analyzed olivines are classified as forsterite (Fo86.7e88.3Fa11.7e13.3) and their Mg# ratios range from 0.867 to 0.883. Orthopyroxenes are classified as enstatite (En93.5e96.8 and Fs3.2e6.5), with Mg# ratios ranging from 0.935 to 0.968, and the clinopyroxenes as augite (En72.3e75.6 and Fs10.5e11.4). The ortho- and clino-pyroxenes are enriched in Al2O3, and depleted in NiO. Spinel contained in the host fragments occur as medium-grained accessory phase with anhedral crystals and many of them appear in clusters of ortho- or clino-pyroxene. The spinel is categorized as Al Mg-chromite. 4.3. Geothermobarometry and physicochemical conditions of alteration The compositions of the analyzed minerals were used for geothermobarometry. For the temperature and pressure estimations, seven different geothermobarometers were used (Table 2). The two-feldspar geothermometer is based on the distribution of the albite component between plagioclase and alkali feldspar crystals during magma solidification (Stormer, 1975; Putirka, 2008). The temperature obtained from the application on the two-feldspar geothermometer of Stormer (1975) in our granitoids is 681 C. The Ti-in-biotite geothermometer of Henry et al. (2005) is based on the Ti-saturation of near-isobaric natural biotite equilibrated at 4e6 kbar. Based on the Ti-in-biotite geothermometer of Henry et al. (2005), the calculated temperature is from 553 C to 583 C. Pressure estimates for the granitoid samples were made using the clinopyroxene barometer. According to Ashchepkov et al. (2001) pressure is the main variable that controls the chemical behavior of clinopyroxene in the magmatic environment. The exchange of jadeite-diopside components on clinopyroxene allows the calibration of the clinopyroxene thermobarometer (Ashchepkov, 2002, 2003). Based on the clinopyroxene barometer the estimated pressure ranged from 5.7 kbar to 6.5 kbar. In addition, the Al-in-hornblende geobarometer of Hammarstrom and Zen (1986) was employed. In the magmatic hornblendes the Altotal content correlates linearly with crystallization pressure of the intrusion. The Al-in-hornblende geobarometer was re-evaluated, amongst others, by Johnson and Rutherford (1989) which emphasized that the Al-in-hornblende geobarometer should be applied
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Table 2 Summary of the geothermo-barometrical results and melt inclusion analyses. Geothermometer-geobarometer
T ( C)
P (kbar)
logfO2
Reference
Two-feldspar Magnetite-ilmenite Clinopyroxene Al-in-hornblende Hornblende-plagioclase Ti-in-biotite Chlorite Melt inclusions Host Homogenization temperature (Th) Water content Trapping pressure (Pt)
682.1 677e682 620e626 553e583 222e251
5.7e6.5 5.0e6.0 5.0e6.0 -
21 to e20.9 -
Stormer (1975) Buddington and Lindsley (1964) Ashchepkov (2002, 2003) Johnson and Rutherford (1989) Blundy and Holland (1990) Henry et al. (2005) Cathelineau (1988) and Kranidiotis and MacLean (1987)
Quartz 635e650 C 1.5e2.1 wt.% H2O 4.38e4.41 kbar
for the assemblage hornblende þ biotite þ plagioclase þ Kfeldspar þ quartz þ titanite þ Fe-Ti oxides. Based on this geobarometer, the estimated pressure ranged from 5.0 kbar to 6.0 kbar. The Al-in-hornblende geobarometer was also combined with the hornblende-plagioclase geothermometer after Blundy and Holland (1990). This geothermometer also uses the Altotal values in hornblende as well as the An and Ab components of the plagioclase that co-exists with it. Based on geothermometer of Blundy and Holland (1990), the obtained temperatures are 620e626 C. Buddington and Lindsley (1964) proposed the use of the compositions of coexisting ilmenite-hematite and magnetiteulvospinel solid solutions in the system FeO-Fe2O3-TiO2 as a geothermometer and oxygen barometer (i.e., magnetite-ilmenite geothermo-oxygen-barometer) which has been re-evaluated by Powell and Powell (1977) and Stormer (1983). The temperature obtained based on the application on the magnetiteilmenite geothermo-oxygen-barometer of Buddington and Lindsley (1964) (calculated with ILMAT excel, Lepage, 2003) for this co-existence is estimated to range from 677 C to 682 C. For these temperatures the logfO2 values are w e21 or DlogfO2 (HM) ¼ e8.5 at w 680 C, suggesting a highly oxidized melt. The used geothermobarometers and the crystallization temperatures and pressures of the felsic granitoid minerals are summarized in Table 2. The non-stoichiometric behavior of chlorite makes it a geothermometer, as chlorite composition (i.e., AlIV content of chlorite) unequivocally relates to temperature of formation (Cathelineau and Nieva, 1985; Kranidiotis and MacLean, 1987; Cathelineau, 1988; Jowett, 1991). Based on the chlorite geothermometer the calculated temperatures that represent the temperatures of hydrothermal alteration range from 222 to 251 C. The chemical conditions of the highly serpentinized sample PGD2_2 were estimated from phase stability relationships using SUPCRT92 (Johnson et al., 1992) with thermodynamic properties from the 2007 database (slop07.dat; Shock and Helgeson, 1988). These estimations are based on the coexistence among silicates for NaCl-saturated aqueous liquids. Temperatures of 230 C (averaged T obtained from chlorite geothermometer) and pressure of 1 kbar are assumed. All solid phases considered to have ideal behavior. Individual ion activity coefficients of dissolved species were calculated using the B-gamma extension of Helgeson et al. (1981).
Plagioclase 664e733 C 2.1e2.7 wt.% H2O 4.44e4.59 kbar
The ionic strength calculated values of the serpentinized solution are I ¼ 0.3 0.01. From the equilibria at 230 C, the obtained ionic activity values occurred during the alteration of the plagiogranites are listed in Table 3. Fluid-wall rock interaction played a major role in the alteration of the granitoids. Alteration involved hydration of the mafic minerals (Table 3, reaction 1) and ion exchange between the fluid phase and these rocks (Table 3, reactions 2e4). The most significant parameter during alteration was pH neutralization. The pH increase modified the stability fields of the alteration minerals, and subsequently drove reactions (reactions 1e4, Table 3) to the right leading to their precipitation. 5. Melt inclusions 5.1. Petrography of melt inclusions Melt inclusions in the Wadi Zikt felsic granitoids occur as isolated inclusions or in assemblages defining crystal growth zones within plagioclase and quartz crystals. Melt inclusions that occur in the same growth zone suggest synchronous crystallization at the time the melt was being trapped (Fig. 4b) and provides the necessary evidence that the melt was in equilibrium with the host plagioclases and quartz. Melt inclusions vary in sizes from 10 mm to 30 mm. The analyzed melt inclusions are composed of homogeneous glass with or without bubble or daughter minerals (Fig. 4a,b, respectively). The shapes of melt inclusions are rounded, negative or elongate (Fig. 4c). The melt inclusions are assumed to be primary as they are trapped along crystal growth faces or between growth zones with no visible fractures, following the nomenclature of Roedder (1984). 5.2. Microthermometry of melt inclusions Microthermometric measurements included the complete melting temperature (Tm) of all silicate phases (except host) within the melt inclusion and the homogenization temperatures (Th). For undersaturated melts, the Th constrains the minimum trapping temperature (Tt). The pressure inside an inclusion (internal trapping pressure Pt) can be estimated using the measured
Table 3 Summary of the alteration chemistry.
1. 2. 3. 4.
Reaction
Minerals
Physicochemical conditions (T ¼ 230 C)
2Mg2Si2O6(s) þ 2Hþ(aq) þ H2O(l) ¼ Mg3Si2O5(OH)4(s) þ Mg2þ(aq) þ 2SiO2(s) 3KMg3AlSi3O10(OH)2(s) þ 20Hþ(aq) ¼ KAl3Si3O10(OH)2(s) þ 2Kþ(aq) þ 9 Mg2þ(aq) þ 6SiO2(aq) þ 12H2O(g) CaAl2Si2O8(s) þ 5Mg2þ(aq) þ 8H2O(l) þ SiO2(aq) ¼ Mg5Al2Si13O10(OH)8(s) þ Ca2þ(aq) þ 8Hþ(aq) CaAl2Si2O8(s) þ 2Naþ(aq) þ 4SiO2(aq) ¼ 2NaAlSi3O8(s) þ Ca2þ(aq)
Enstatite-serpentine Biotite-muscovite Andesine-chlorite Andesine-albite
log[aMg2þ/(aHþ)2] ¼ 5.2 log(aK þ/aHþ) ¼ 3.8 log[aCa2þ/(aHþ)2] ¼ 6.5 log(aNaþ/aHþ) ¼ 4.9
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Homogenization temperatures of the melt inclusions hosted in plagioclase ranged from 664 C to 733 C whereas the ones in quartz from 635 C to 650 C (Table 2). These temperatures are w þ50 C higher than the ones obtained from the geothermometry, suggesting that some of these temperatures represent cooling or closure temperatures (Table 2). The water contents (Cw) and entrapment pressures of melt inclusions were also calculated with the MELTS software (Asimow and Ghiorso, 1998). The water contents of the melt inclusions based on their measured melting temperatures of plagioclase range of 1.5e2.1 equivalent wt.% H2O, and that of quartz range of 2.1e2.7 equivalent wt.% H2O. The calculated pressures range from the melt inclusions hosted in quartz and plagioclase, i.e., from 4.4 kbar to 4.6 kbar (Table 2) are determined by using the equation Holtz et al. (1992). These pressures are lower w2 kbar than the pressures obtained from the geobarometry, and attributed to ductile deformation of the pyroxene and hornblende (Table 2, see also Searle et al., 2015). Based on the model the Ebadi and Johannes (1991) for the haplogranite system (Ab-An-Or-Qz-H2O) and their equations, the glass has a composition of An36.3e51.7, Ab1.1e2.7, Or0.06e0.1 and Qz42.4e60.1. EPMA analyses of the glasses showed that they compose of (all in wt.%) 73.45e76.45 SiO2, 13.29e15.78 Al2O3, 0.01e0.65 FeO, 3.59e7.84 Na2O, and 0.02e4.01 K2O, whereas TiO2, MnO, MgO and CaO are 0.03 wt.% (Table 1). The glass composition in plagioclase has (K/Ti)N values ranging from 10.1 to 17, which is considered as typical of MORB composition (Sobolev, 1996; Michael et al., 2002). The analyzed halogen and H2O contents were 0.012e0.021 Cl, 0.0002e0.005 F and 1.45e2.71 wt.%, respectively (Table 1). 6. Geochemistry 6.1. Major element geochemistry
Figure 4. Petrography of melt inclusions trapped in quartz (a and b) and plagioclase (c) crystals. The melt inclusions contain glass, and sometimes bubble or daughter minerals (M.I.: Melt inclusions, P.F.I.: Primary fluid inclusions, S.F.I.: Secondary fluid inclusions in arrays, Qz: Quartz, Pl: Plagioclase). Photographed at þ 22.5 C in plane-polarized, transmitted light.
temperature, isochores for associated melts, and the appropriate melt-solid phase equilibrium, using the MELTS software (Asimow and Ghiorso, 1998). All of the studied melt inclusions were homogenized to liquid upon heating.
The Dadnah felsic granitoids have high SiO2 (66e77 wt.%), low to moderate total alkali contents (Na2O þ K2O ¼ 1.4e10.5 wt.%) and Mg# values ranging from 44.8 to 84.0 (clustering at 65 and 75, Table 1). Utilizing the total alkali-silica (TAS) discrimination diagram of Middlemost (1994) and the normative Qz-An-Or and AnAb-Or ternary plot (after Barker, 1979), the majority of the analyzed granitoids and melt inclusion glass are classified as subalkaline trondhjemites and tonalites (for simplicity from now on are referred as tonalites, Fig. 5a and b). The tonalites and melt inclusion glass of Dadnah area display a wide range of An contents (plagioclase normative compositions from w14 to w 64) with low values of Or, suggesting plagioclase fractionation (Table 1, Fig. 5b). Similar distributions of Ab, An, and Or contents have been reported from previous studies, especially, for the crustal tonalites-togranites (Rollinson, 2009, 2014, 2015; Grimes et al., 2013). Based on the ASI values and silica contents, our tonalites (ASI ¼ 0.74e1.08), trondhjemite (ASI ¼ 0.95) and melt inclusion glass (ASI ¼ 0.71e0.77) are from strong-to-mildly metaluminous to weakly peraluminous and are mainly plotted as I-type granitoids (Fig. 5c). On the Harker diagrams, in which we have added the EPMA analyses of the glass obtained from the melt inclusions and the available data of Rollinson (2015) from the Dadnah area, the Fe2O3, CaO, K2O and MgO contents of our tonalites show strong-to-weak depletion with increasing SiO2. Moreover, Na2O reveals positive correlation with increasing SiO2 (Fig. 6aef). These trends suggest that the tonalitic magmas were evolved due to fractional crystallization of clinopyroxene and plagioclase (alkali feldspar), which was accompanied with an increase of the Ab content of the plagioclase (cores to rims). Scatter of K2O and Na2O indicates that
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K-feldspars and albite were not significant crystallizing phases (Fig. 6c and d, Table 1). However, the fact that the compositions of glass measured from the melt inclusions are similar to the analyzed compositions of the tonalites (e.g., PGD2_2, PGD3_2, PGD6_1, PGD7, Table 1) indicates that some of them represent least-fractionated melts. 6.2. Trace element geochemistry
Figure 5. Tonalite classification plots using major elements: (a) Alkali versus silica classification plot (TAS in wt.%) after Middlemost (1994). The alkaline and subalkaline boundary of Irvine and Baragar (1971) is shown (after Rickwood, 1989). (b) Normative quartz-anorthite-orthoclase (Qz-An-Or) ternary diagram, where the tonalitetrondhjemite fields are shown (after Barker, 1979) (blue triangles refer to composition of the analyzed melt inclusions, D: diorite, T and Tj: tonalite-trondhjemite). The experimental basaltic melt compositions are from Koepke et al. (2004). (c) SiO2 (wt.%) versus ASI plot (ASI ¼ the aluminum saturation index Al/[2(Ca 1.67P) þ Na þ K]. The metaluminous and peraluminous fields are adapted from Maniar and Piccoli (1989).
Chondrite-normalized multi-element plot of the tonalites shows that the high field strength elements (HFSE) are rather depleted relative to the large-ion lithophile elements (LILE). Also U and Ti (Ti/Ti* ¼ 0.00e0.27) display strong negative and the Nb, Ta, Zr, Y, Hf, and Th moderate negative anomalies, relative to the adjacent elements of similar incompatibility (Zr and REE’s). Zr, Y, Hf, Nb and Ta depletion is related to fractionation of hornblende and zircon. From the more incompatible elements, Ba is strongly enriched and Pb (Pb/Pb* ¼ 7.16e78.13), La and Ce show a small positive anomaly. This trace element distribution pattern is probably attributed to subduction of recycled oceanic crust (Hofmann, 2004; Pilet et al., 2011, based on Monte-Carlo simulations) with minor contamination of oceanic sediments (Fig. 7a). The total REE abundance of the tonalites is up to w283 ppm. Chondrite-normalized REE patterns of the studied samples display two different forms. The first type is essentially sub-parallel with relatively flat and gently sloping patterns for the middle- (MREE) and heavy-rare earth elements (HREE) with (Gd/Yb)N values ranging from 0.8 to 1.1. The undepleted patterns for the MREE and HREE suggest low-pressure during the formation of the tonalitic melts (e.g., garnet-free), which is also consistent with the pressure determined by melt inclusions (Table 2). They appear mildly-tostrong enriched in the light rare earth elements (LREE) relative to the HREE, with (La/Sm)N and LREE/HREE values ranging from 1.1 to 1.7, and 4.2 to 62.7, respectively (Table 1). They also show weak positive Eu and negative Ce, Sm and Ho anomalies. These trends suggest fractional crystallization of plagioclase and hornblende. These tonalites which contain plagioclase, quartz, hornblende and minor clinopyroxene are metaluminous with (La/Yb)N values of 1.9e4.5 indicating that are moderate fractionated (Fig. 7b, Table 1). Fractionation and enrichment of LREE is considered typical for felsic granitoids related to SSZ, and is also found in the tonalitegranodiorite suite from the mantle section of Oman ophiolite (Wadi Hemli or Hamiliya; Rollinson, 2014, 2015). The second type occurs with the same overall form (e.g., (Gd/ Yb)N ¼ 0.8e2.1 and (La/Sm)N ¼ 0.9e1.2). In addition, they characterized by intense positive Eu (Eu/Eu* ¼ 2.6e7.1), no or slightly positive Ce and negative Sm anomalies, coupled with high Ba and low Rb/Sr values which are attributed to accumulation of plagioclase (Table 1). These tonalites that contain plagioclase, quartz, biotite and minor hornblende are peraluminous with (La/Yb)N values of 0.5 suggesting that they are least fractionated (Fig. 7b, Table 1). For comparison, we have also plotted the host harzburgite (e.g., sample OP_2) that displays similar U and Th depletion and Ba enrichment to the tonalities (Fig. 7a). The host harzburgite shows the lowest values of (La/Sm)N ¼ 0.28, (La/Yb)N ¼ 0.21, (Gd/ Yb)N ¼ 0.66 and Mg# ¼ 64, negative Ce (w0.9) and strong positive Eu (w2.2) anomaly. Few of Dadnah tonalites appear with (La/ Sm)N, (La/Yb)N, (Gd/Yb)N and Mg# values similar to their host harzburgite. This suggests that there was interaction of the tonalite melt with its host harzburgite. In one of the thickest dikes, we have sampled both the margins and the median plane of the dike (e.g., PG4Ma and PG4Mi). This dike appears internally geochemically variable from the margins towards the median plane. For the major oxides there is a decrease of the Fe2O3t and
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Figure 6. (aee) Harker diagrams for selected major elements (wt.%). The black circles represent the tonalites, the red squares the host harzburgite, the blue triangles the analyzed melt inclusions and the purple rhombs the data of Rollinson (2015) from the Dadnah area. (f) Ti versus Al2O3/TiO2 plot depicting fractional crystallization for the tonalites.
Figure 7. Chondrite-normalized (a) multi-element plot (concentrations are normalized to chondrite average composition of McDonough and Sun, 1995) and (b) rare earth element (REE) concentrations (normalized after chondrite of Sun and McDonough, 1989). The solid lines represent the metaluminous tonalites, and the dashed lines the peraluminous tonalites. For comparison, we have also plotted the host harzburgite (e.g., sample OP_2, red circles).
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MgO (w60%) and CaO and increase of their K2O (w40%) and Na2O contents (Table 1). Also regarding the trace elements, there is a decrease of the Ba and SREE (w30%), Hf and Th (w40%) and Ta (w70%) and an increase of the Zr (w30%) and Pb (w60%) contents. These relationships support that fractionation increased towards the center of the dike and there was also interaction with the host harzburgite at the margins. Pearce et al. (1984) proposed the Y versus Nb trace elements for the classification of granitoids, on the basis that Y and Nb are relatively more independent for alteration. In Fig. 8a, it is supported that our samples plot mainly in the volcanic arc granites field (VAG) which is also evident from the trace element patterns (Fig. 7a). Additionally, it appears that the majority of the tonalites and glass intruded under compression, suggesting that the melts are associated with the subduction process (Fig. 8b, see discussion also). 6.3. Characterization of source and magmatic differentiation Based on the Nd versus Nd/Ce plot, we have been able to distinguish the two plausible processes responsible for the magmatic differentiation of the tonalite melts (Fig. 8c). The positive trend between Nd and Nd/Ce (Fig. 8c) implies that the tonalites were produced from partial melting, whereas the constant values of Nd/Ce over a wide range of the Nd content as well as the increasing of SiO2 with decreasing (La/Sm)N at almost constant LaN values, indicate that they were subjected to fractional crystallization (Fig. 8c and d). Following the evidence for partial melting involvement in the formation of the tonalites, the CaO/(MgO þ FeOt) versus Al2O3/ (MgO þ FeOt) plot indicates that the tonalites and related melt inclusion glass derived from partial melting of basaltic and greywacke sources (Fig. 8e). The majority of them, however, relate to a basaltic protolith (Mg# values of 44e53, Kelemen et al., 1997, and TiO2 values for the melt inclusions glass of 0.03 wt.%, Table 1). These results are consistent with the ones obtained from halogens (Table 1). It is apparent that there are two types of tonalites. The first group has elevated Cl/B ratios (Cl/B ¼ 50) depicting the presence of tourmaline, while the second group shows lower Cl/B ratios (Cl/B 22.5) depicting the presence of apatite. The presence of tourmaline and apatite could be related to both partial melting of sedimentary and igneous protoliths. The majority of the tonalites, according to the log(Nb/Y) versus log(Ti/Y) diagram, are plotted in the transitional field from tholeiitic to calc-alkaline affinity (Fig. 8f). These trends are consistent with various degrees of partial melting of subducting oceanic crust (e.g. Martin et al., 2005; Thorkelson and Breitsprecher, 2005). 6.4. Isotope geochemistry Previous studies of oxygen isotopes of the Semail plagiogranites (obtained from tonalites-trondhjemites and minerals: plagioclase, pyroxene, hornblende, quartz and zircon) indicated that d18O values range from 2.4& to 14.0& (e.g. McCulloch et al., 1980, 1981; Gregory and Taylor, 1981; Stakes and Taylor, 1992; Grimes et al., 2013). McCulloch et al. (1980, 1981), and Gregory and Taylor (1981) suggested that the d18OH2O values have been modified due to interaction of seawater with the oceanic crust. Gregory and Taylor (1981) attributed the average d18OH2O values (e.g. 5.7& 0.2&) of the oceanic crust to seawater buffering due to hydrothermal circulation. Unfortunately, there are no calculated d18Omelt or d18OH2O values for the isotopic equilibrium of the plagiogranites or their minerals. Additionally, the d18OH2O values given by Gregory and Taylor (1981) were calculated based on the assumption that both the granitoids and all of their minerals were in equilibrium with a fluid phase and not with the melt.
1199
Stakes and Taylor (1992) have argued that the plagiogranites represent the final magmatic event that occurred at the spreading axis. Grimes et al. (2013) have interpreted the low d18O values of zircons in tonalites-trondhjemites due to re-melting of the crust, which was altered by hydrothermal fluids (i.e., seawater at higher temperatures). Recently, Spencer et al. (2017) reported unusually high d18O values from peraluminous, S-type granite samples in UAE and Oman that range from 14& to 28& (whole-rock: 14&e23&, quartz: 20&e22&, and zircon: 14&e28&). They suggest that these extremely high d18O values indicate that the protolith of these granitoids was marine pelitic/siliceous muds that entered the mantle through the subduction zone. In our study, the oxygen and hydrogen isotope compositions were obtained from the tonalite minerals, including pyroxene, plagioclase, quartz, hornblende, biotite, and calcite (Table 4). For the calculations, we have used the average trapping temperature of melt inclusions as well as the temperatures obtained from the application of the different geothermometers (Table 2). We have utilized the mineral-melt and -H2O equations of Zhao and Zheng (2003) and Zheng (1993a,b). The calculated d18OMelt, d18OH2O and d18DH2O values are listed in Table 4. The d18Omelt or d18OH2O values range from 5.3& to 6.7&, clustering at w5.3& and 6.2&. The d18OH2O values of hydrous mafic minerals of the tonalites are higher (e.g., 7.3&e7.8&, Table 4), whereas the d18OH2O values of quartz and calcite from the later veins are lower (e.g., 1.8&e3.0&, Table 4). Valley (2003) and Carmody et al. (2013) suggested that the values of w5.5& represent a lithospheric mantle source, whereas the values that are w 6.8& are enriched relative to the mantle and retain isotopic characteristics of typical oceanic crust related to MORB (Zi et al., 2012). Additionally, the d13CCO2 values of the vein calcite are plotted in the ‘Igneous calcite box’ indicating a MOR-type source for CO2 (Fig. 9c, Chaussidon et al., 1991). We report, for the first time, nine helium isotope compositions obtained from pyroxene, plagioclase, hornblende and biotite in the Dadnah tonalites (Table 4). The R/RA compositions range from 6.05 to 7.99 (Table 4) and are close to or lower than the values attributed to MORB (R/RA ¼ 8, Savelieva et al., 2008; Lustrino and Anderson, 2015). It is implied that the possible helium source for our analyzed tonalites is related to MOR-type oceanic crust. Most of these minerals were also analyzed for their hydrogen isotopes. Previous studies of hydrogen isotopes from Gregory and Taylor (1981) obtained data from whole rock and hornblende. Their measured dD values range from e56& to 0.0&, and are attributed to magmatic H2O, which interacted with seawater, however, there are no data for the d18DH2O values. We have obtained a dDmelt value of e85& from pyroxene, which is also related to MOR-type oceanic crust (Chaussidon et al., 1991). Spooner (1977), Gregory and Taylor (1981) and McCulloch et al. (1981) reported that 87Sr/86Sr initial ratios of the mantle and crust lithologies of the Semail ophiolite display extremely large variation (e.g., 0.7030e0.7065). They suggested that they were modified due to hydrothermal interaction of oceanic crust with seawater. McCulloch et al. (1981) also reported that the initial 143Nd/144Nd ratios of these rocks have a limited range of εNd values of 7.5e8.6, indicating that all have distinctive oceanic affinity. Cox et al. (1999) reported negative εNd values and intermediate-to-high initial 87 Sr/86Sr ratios of the analyzed monzogranites-to-leucogranites from the Khor Fakkan block, and suggested mixing of a LILE enriched mantle component, i.e., metabasalts and a continental sedimentary component. Haase et al. (2016) proposed that the almost constant age-corrected εNd values of w8 from the tonalites over a range of 87Sr/86Sr initial values (e.g., 0.7035e0.7060) resemble to MORB oceanic crust affected by seawater. The rubidium, strontium, samarium and neodymium isotope results were obtained from fifteen samples of the tonalite minerals,
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Figure 8. Diagrams indicating the possible tectonic settings of Dadnah tonalites (symbol color same as Fig. 6). (a) Y (ppm) versus Nb (ppm) plot. The WPG (Within Plate Granite), VAG þ Syn-COLG (Volcanic Arc Granite, Syn-Collisional Granite) and ORG (Ocean Ridge Granite) fields are adapted from Pearce et al. (1984). (b) SiO2 (wt.%) versus log (CaO/ (Na2O þ K2O)) plot. The extensional and compressional fields are adapted from Wu and Kerrich (1986). (c) Nd (ppm) versus Nd/Ce plot. The partial melting and fractional crystallization trends are adapted from Pearce (1980) and Ahmad and Tarney (1991). (d) LaN versus (La/Sm)N plot. The partial melting and fractionation trends are adapted from Xu et al. (2015). (e) Molar CaO/(MgO þ FeOt) versus molar Al2O3/(MgO þ FeOt) plot (A: partial melts with metapelitic source, B: partial melts with metagreywackes source and C: partial melts with meta-basaltic and -tonalitic source). The meta-pelitic, -greywackes and meta-basaltic and -tonalitic fields are adapted from Altherr et al. (2000). (f) Nb/Y versus Ti/Y plot. The tholeiitic, transitional and calc-alkaline fields are adapted from Winchester and Floyd (1977) as modified by Pearce (1996).
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Table 4 Oxygen, hydrogen, carbon, silicon, and helium isotope data of the studied tonalites. Utilized the mineral-melt-H2O and -CO2 equations and isotope fractionation values from Zhao and Zheng (2003) and Zheng (1993a,b). Samples
Minerals
d18O (&) T ( C) d18Omelt d18OH2O dD (&) dDmelt dDH2O d13C d13Cfluid d30Si (&)
PGD6_1a PGD6_1b PGD6_1c PGD6_1a PGD6_1c PGD6_1c PGD6_1c PGD6_1b PGD6_1c PGD6_1b PGD6_1a PGD6_1c PGD6_1b PGD6_1a PGD6_1c
Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, coarse-grained Pyroxene, coarse-grained Plagioclase, fine-grained Plagioclase, coarse-grained Hornblende Biotite Calcitea Calcitea Calcitea Quartza Quartza Quartza
4.4 4.7 4.3 4.3 4.2 4.2 5.1
730 730 730 700 700 680 620
5.0 5.3 5.9 5.7 5.4 5.8 5.5 5.6
620 550 430 430 430 430 430 430
(&)
(&) 37
6.4 6.7 6.3 6.2 6.1 5.3 6.1 7.3 7.8 3.0 2.8 2.5 2.1 1.8 1.9
55 61
(&)
(&)
(&)
(&)
85
77.1 95.2
He (107 cm3 3He (1013 cm3 3He/4He R/RA (106) STP/g) STP/g) 4
3.56 3.77 3.64 2.43 4.19 3.37 2.28
38.49 39.88 37.92 22.95 38.56 37.45 23.17
10.81 10.58 10.42 9.44 9.20 11.11 10.16
7.78 7.61 7.49 6.79 6.62 7.99 7.31
4.23 4.84
37.93 40.67
8.97 8.40
6.45 6.05
6.2 4.7 5.6 4.1 5.8 4.3 62
e 0.43 e 0.42 e 0.41
a Temperatures were determined independently based on the calcite-quartz equilibrium pairs (based on the equations of Zhao and Zheng, 2003 and Zheng, 1993a, b). The isotopic temperatures obtained are w430 C.
including pyroxene, plagioclase, hornblende, biotite, quartz and calcite. Unfortunately, we were able to obtain only two 143Nd/144Nd isotopic values from hornblende and biotite. Accordingly, we have calculated their initial isotopic ratios, epsilon values (εSr and εNd) and isotopic ages (Table 5). Their initial 143Nd/144Nd and εNd values (e.g., 0.511968 and 0.512032, and e13.07 and e11.82, Table 5) are lower than those of the MORB lithospheric mantle (εNd z þ19) and the chondritic reservoir (εNd ¼ 0.0). These values are similar to those of MORB protolith and suggest contamination from oceanic sediments (Chaussidon et al., 1991; Savelieva et al., 2008; Zi et al., 2012; Xu et al., 2015 and references therein). From the hydrogen versus oxygen isotope plot, it is evident that all of the calculated d18OMelt, d18OH2O and d18DH2O values relate to the Subduction-Related Vapor, Arc and Crystal Felsic Magma (Fig. 9a, Kus¸cu et al., 2011). The possibility of mixing of different end-members to the formation of the tonalite melts can be depicted from d18Omelt or d18OH2O versus age-corrected εSr values (Fig. 9b). MOR-type lithospheric mantle is characterized by lower d18OMelt and εSr value (e.g., d18OMelt w 5.6& and εSr w e30) relative to oceanic crust (e.g., d18OMelt w 6.8& and εSr w 0e10), and oceanic sediments (e.g., d18OMelt > 10& and εSr > 35) (Ito and Stern, 1986; Zi et al., 2012; Carmody et al., 2013; Roberts and Spencer, 2014; White and Klein, 2014; Lustrino and Anderson, 2015, and references therein). The tonalite samples on Fig. 9b are plotted on the mixing line between the oceanic lithospheric mantle and oceanic crust. This suggests that from these two sources, the oceanic crust is dominant (Fig. 9b). However, this plot cannot distinguish between the genesis of the melt due to direct mantle-derived differentiation or to re-melting of the oceanic crust. Moreover, the εSr coupled with the εNd values of the tonalites suggest that their parental magmas were formed from partial melting of mainly recycled oceanic crust (Table 4, see Stracke et al., 2003). The composition of silicon isotopes from the tonalites suits with the silicon isotope composition of recycled oceanic crust related to partial melting (i.e. d30SiNBS e0.4&; Pringle et al., 2016). Our helium isotopic data concur with the hypothesis of recycled oceanic crust of MORB affinity that subducted in a lithospheric mantle (Savelieva et al., 2008; Graham et al., 2014; Lustrino and Anderson, 2015). Additionally, in the R/RA versus (87Sr/86Sr)i plot (Fig. 9d) there is a negative trend between the R/ RA and initial 87Sr/86Sr isotopic values, which is consistent with contamination from oceanic sediments of the subducted slab (Savelieva et al., 2008; Lustrino and Anderson, 2015).
7. Discussion 7.1. Formation mechanisms and source for the Dadnah tonalities Previous studies in the Khor Fakkan block have frequently depicted the monzogranites-to-leucogranites as low to high potassic and mostly peraluminous, with high contents of LILE and elevated 87Sr/86Sr ratios relative to mantle (Briqueu et al., 1991; Peters and Kamber, 1994; Cox et al., 1999 and references therein). These geochemical characteristics have been interpreted as contribution from a crustal source, attributed to the partial melting of metasedimentary rocks (e.g., Cox et al., 1999; Spencer et al., 2017). Our analytical data suggest that the Dadnah tonalites are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, (Rb, Ba, Sr, Pb, Eu), and Ta, La, Ce, depleted in U, and occur with low d18OH2O, moderate εSr and negative εNd values. The Dadnah tonalites most probably formed from a complex process involving partial melting of the subducted slab of MOR-type composition, accumulation of plagioclase and fractional crystallization, and interaction of the tonalite melt with the host harzburgite. Accumulation of plagioclase, as a mechanism for the genesis of the Dadnah tonalites, is evidenced by the intense positive Eu (Eu/ Eu* ¼ 2.6e7.1, Fig. 7b), negative Sm anomaly coupled with high Ba and low Rb/Sr values (Table 1). The Eu anomaly may also indicate the temperature of solidification of the melts during the plagioclase accumulation. Temperatures decreased from 730 C down to 665 C (Table 2) as the melts ponded at the Moho Transition zone, due to its behavior as a permeability barrier (Kelemen et al., 1997). The strong positive Eu, Ba and negative U and Th anomalies (Fig. 7a and b) coupled with Mg# of w64, similar to their host harzburgite, suggest that some of the tonalites have interacted with their host rock (Fig. 3f). Fractionation is suggested by the linear trends on the normative An-Ab-Or ternary plot (Fig. 5b), the depletion of the Fe2O3, CaO, K2O and MgO coupled with the positive correlation of Na2O with increasing SiO2 on the Harker diagrams (Fig. 6aee) and the linear trend in the Ti versus Al2O3/TiO2 plot (Fig. 6f). Fractionation is also supported by the enrichment of the tonalities with Mg# values 75 in the LREE relative to HREE (LREE/HREE z 4.2e62.7), the constant values of Nd/Ce over a wide range of Nd content (Fig. 8c) and negative trend of increasing SiO2 wt.% with decreasing (La/Sm)N at almost constant LaN values (Fig. 8d).
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Figure 9. Diagrams indicating origins of melts and fluids involved in the formation of the tonalites: (a) Hydrogen versus oxygen isotope plot, showing the calculated water compositions from Dadnah tonalites. The ‘Meteoric Water Line’ and ‘Kaolinite Line’, ‘Subduction-Related Vapor, Arc and Crystal Felsic Magma’, ‘Metamorphic water box’, ‘Magmatic melts and Traditional Magmatic Water box’ are obtained from Kuþcu et al. (2011) and references therein. (b) εSr versus d18Omelt or d18OH2O compositions of Dadnah tonalites, relative to possible magma sources. ‘Lithospheric mantle’, ‘Oceanic crust’ and ‘Sediments’ fields are after De Paolo (1980), Ito and Stern (1986), Cox et al. (1999), Rudnick and Gao (2003), White and Klein (2014), and Lustrino and Anderson (2015). The inset histogram is a detail of the d18Omelt or d18OH2O axis. The colored bars represent the analyzed minerals, i.e., pyroxene, plagioclase, quartz, calcite, hornblende and biotite. (c) Carbon versus oxygen isotope diagram for vein calcites showing the calculated isotope systematics of the tonalites. The ‘Marine Limestone’ and ‘Igneous Calcite’ boxes are from Bowman (1998). (d) R/RA versus (87Sr/86Sr)i plot, relative to possible magma sources. The ‘MORB line’ and ‘Oceanic crust’ field are after Savelieva et al. (2008) and Lustrino and Anderson (2015). (e) 87Rb/86Sr versus (87Sr/86Sr)i plot, relative to magma mixing and possible sources. ‘Lithospheric mantle’ and ‘Oceanic crust’ are after Ito and Stern (1986), Rudnick and Gao (2003), and Lustrino and Anderson (2015).
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Table 5 RbeSr and SmeNd isotope results of the studied tonalites. Samples
Minerals
87
Rb/86Sr measured
87 Sr/86Sr measured
87 Sr/86Sr initial
Age (Ma)
εSr
PGD6_1a PGD6_1b PGD6_1c PGD6_1b PGD6_1a PGD6_1b PGD6_1c PGD6_1c PGD6_1a PGD6_1b PGD6_1b PGD6_1c PGD6_1a PGD6_1b PGD6_1c
Pyroxene, coarse-grained Pyroxene, coarse-grained Pyroxene, fine-grained Pyroxene, fine-grained Pyroxene, fine-grained Plagioclase, fine-grained Plagioclase, coarse-grained Quartz Quartz Quartz Calcite Calcite Calcite Hornblende Biotite
0.24 0.36 0.39 0.37 0.26 0.15 0.19 0.20 0.19 0.18 0.15 0.16 0.17
0.705325 0.705841 0.704958 0.704977 0.704886 0.704865 0.705337 0.707357 0.707760 0.707342 0.708002 0.707460 0.707190
0.705650 0.706166 0.705283 0.705295 0.705211 0.705190 0.705662 0.707682 0.708085 0.707667 0.708327 0.707785 0.707515
98.22a,c 98.5 2a,c 97.82b,c 97.32b,c 97.52b,c 97.72d 98.61d 95.61e 95.71e 95.71e 95.4 1 94.9 1 95.2 1
16.3 23.6 11.1 11.3 10.1 9.8 16.5 45.2 50.9 45.0 54.3 46.6 42.8
147 Sm/144Nd measured
143 Nd/144Nd measured
143 Nd/144Nd initial
Age (Ma)
εNd
0.246 0.257
0.512468 0.512475
0.511968 0.512032
96.51f 96.31g
13.07 11.82
The MSWD (Mean Square of Weighted Deviates) shown in the table is a measure of the ratio of the observed scatter of the points (from the best-fit line) to the expected scatter (from the assigned errors and error correlations). a MSWD ¼ 0.83 for coarse-grained pyroxene. b MSWD ¼ 0.88 for fine-grained pyroxene. c MSWD ¼ 0.85 for all analyzed pyroxene. d MSWD ¼ 0.95 for fine- and coarse-grained plagioclase. e MSWD ¼ 0.97 for quartz. f MSWD ¼ 0.80 for hornblende. g MSWD ¼ 1.12 for biotite.
The degree of fractionation that the tonalite melts were subjected was performed with the AFC-Modeler software (Ersoy and Helvaci, 2010). For this model we used two incompatible trace elements Rb and Zr (D ¼ 0.01 and 0.002, respectively) and the least and most evolved samples (e.g., PG1, SiO2 ¼ 66.92 wt.%, Zr ¼ 36 ppm and Rb ¼ 79 ppm, and PG5, SiO2 ¼ 77.99 wt.%, Zr ¼ 79 ppm and Rb ¼ 80.4 ppm, Table 1). Assuming Rayleigh crystal fractionation, we estimate that the Dadnah tonalites experienced w55%e57% fractional crystallization (F ¼ 73). Partial melting is implied by the fact that the Dadnah tonalities are enriched in LILE, the positive trends of Nd versus Nd/Ce and LaN with (La/Sm)N (Fig. 8c and d). The CaO/(MgO þ FeOt) versus Al2O3/ (MgO þ FeOt) plot indicates that the tonalites and related melt inclusion glass were derived from the partial melting of basaltic and greywacke sources (Fig. 8e). The majority of them, however, relate to a basaltic protolith (Mg# values of 44e53, Kelemen et al., 1997, and TiO2 values for the melt inclusions glass of 0.03 wt.%, Table 1). Their elevated εSr values relative to mantle and negative εNd values, coupled with the composition of silicon isotopes, suggest that their parental magma originated from partial melting of mainly recycled oceanic crust with minor contribution of melting from oceanic sediments (Table 5). The degree of partial melting in the oceanic slab of MOR-type composition was performed with the AFC-Modeler software (Ersoy and Helvaci, 2010). For this model, we used the typical LaN and (La/Sm)N composition of the oceanic lithosphere (Sun and McDonough, 1989) and the least fractionated tonalite sample PGD7 with (La/Yb)N values of w0.3 (Table 1, Fig. 8d). Assuming the continuous and batch melting models, we obtain that the Dadnah tonalites formed due to w10%e15% continuous melting or 11.6% (w12%) batch partial melting of the oceanic slab (Fig. 8d). In an attempt to estimate the possible contribution of the oceanic lithosphere representing the slab (as defined by Jacobs et al., 2015, his Fig. 6 for the Semail ophiolite) and the overlying sediments that possibly involved in the formation of Dadnah tonalites, we have employed the mixing model method of De Paolo (1980) (Fig. 9b and d). For our calculations, we have used the AFCModeler software (Ersoy and Helvaci, 2010). The tonalite melt consists of two end-members, i.e., oceanic lithosphere (with an
age-corrected isotopic composition of (87Sr/86Sr)i,96Ma ¼ 0.703 and Rb/86Sr ¼ 0.001, Allégre and Sutcliffe, 2008; White and Klein, 2014; and oceanic sediments (87Sr/86Sr)i,96Ma ¼ 0.716 and 87 Rb/86Sr ¼ 0.32, Cox et al., 1999; Rudnick and Gao, 2003). The oceanic lithosphere component as a source of the least-fractionated and fractionated Dadnah tonalites is contributing 93% to 97.5%, with an average of w96%, while the melted sediments contribution is 2.5%e7%, with an average of w4% (Fig. 9b). The low contribution of oceanic sediments (w4%) in the melt that formed the Dadnah tonalites is also supported by the observation of Ernewein et al. (1988) which argue that no major sedimentary layer was involved during the transition of V1 and V2 volcanism. We have also estimated the contribution of the oceanic crust component of MOR-type composition on the tonalite partial melt. For our calculations, we have used the incompatible element Rb. A typical RbMean value for the oceanic crust is considered as 1.4 ppm (White and Klein, 2014). The least-fractionated tonalites have Rb values ranging from 4.1 ppm to 82 ppm (average of w34 ppm, Table 1), while the fractionated ones have values ranging from 7.3 ppm to 130.3 ppm (average of w43 ppm, Table 1). Since the average contribution of the oceanic lithosphere, as estimated above, is w96%, the concentration of Rb in the least-fractionated and fractionated tonalites ranges of 3.9e78.7 ppm and 7e124.9 ppm, respectively. Applying these values on the MELTS software (Smith and Asimow, 2005) the obtained F values are 1.8%e7.1% for the least-fractionated and 2.2%e4.1% for the fractionated tonalites. Therefore, the least-fractionated tonalite melts originated from the partial melting of w1.8%e7.1% lithospheric mantle and w92.9%e 98.2% recycled oceanic crust, while for the fractionated tonalites the lithospheric mantle contribution was 2.2%e4.1% and 95.9%e97.8% of recycled oceanic crust (average values for all tonalites, 97% recycled oceanic crust and 3% lithospheric mantle; Fig. 9e). A possible origin of the recycled oceanic crust is evident from the occurrence of mafic enclaves, the LREE/HREE ratios of 26e63, the R/RA and lower εSr values (e.g., w10, Tables 4 and 5), the association on CaO/ (MgO þ FeOt) versus Al2O3/(MgO þ FeOt) plot in which the tonalites and the glass contained in the melt inclusions display Mg# values of 44e56, strong Ti depletion (Ti/Ti* 0.27) and the Pb spike in the spidergram (Pd/Pb* up to w 82, Fig. 7a, Table 1; see also Hofmann 87
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and White, 1982; Hofmann, 2004) suggest mainly a meta-basaltic protolith field (Fig. 8e, field C). It is quite difficult to define with confidence the protolith (i.e., volcanic rocks, gabbros, amphibolites etc.) that comprised the recycled oceanic crust which partially melted to generate the Dadnah tonalite melt, since the magmatism in the Khor Fakkan block was almost continuous during the late Cretaceous and the transition from V1 to V2 phase occurred almost at the same time, i.e., 96.4e95 Ma (Goodenough et al., 2014; Thomas and Ellison, 2014). Furthermore, the subducted slab should be also comprised by older oceanic crust formed during the Arabian passive margin development (Triassic to Cretaceous Haybi volcanics, cherts and shallow water exotic limestones; Searle et al., 2015). However we cannot rule out the possibility of layered or massive gabbros below those units as a source too. The Dadnah tonalities contains clinopyroxene as a minor mafic mineral (Fig. 3d, with high Mg# values of 84e85.5 and TiO2, Cr2O3 and NiO depletion), while their plagioclases are less calcic (An44e56, Figs. 3g and 5b) than the experimental products of Koepke et al. (2004). Koepke et al. (2004) have melted gabbro samples, amongst others, at T 900 C which contain clinopyroxene (with Mg# values of 75e78) and plagioclase (An53e55) to produce plagiogranites enriched in Al and depleted in Ti and Fe. These An contents can be explained due to the smaller influence of water during the plagioclase crystallization (Koepke et al., 2004). Alternatively, these plagioclases with cores An53e56, which occur in the least-fractionated tonalites with (La/Yb)N values 0.5, may represent residual plagioclase (Fig. 3d and g) after the partial melting event of the recycled oceanic crust. The least-fractionated, weakly peraluminous tonalites can be also derived from the fractional crystallization and plagioclase accumulation of a metaluminous melt (Lee and Morton, 2015; Gao et al., 2016). Plagioclase accumulation is an effective way for the tonalite melt, which formed by partial melting of the subducted slab, to evolve from strong-to-mildly metaluminous to weakly peraluminous (ASI 1.08). In this scenario, which can also give an explanation for the two different tonalite types (metaluminous and peraluminous types in the sill of D6 location; Fig. 2), the Dadnah metaluminous tonalites were subjected to feldspar fractionation (ASI z 1, Bilal and Giret, 1999). Feldspar accumulation is not expected to significantly change the Al content of the residual tonalite melt (Bilal and Giret, 1999; Frost and Frost, 2008). Textural evidence (Fig. 3d) suggests that feldspar accumulation which occurred at temperatures ranging from 665 C to 730 C (Table 2) was followed by fractionation of hornblende (ASI ¼ 0.5, Bilal and Giret, 1999) at T z 620 C (Table 2) that resulted in the formation of the moderate fractionated, metaluminous tonalites (ASI 0.71) coupled with an increase of the Al content in the residual melt. Then, the residual was mixed and contaminated with the oceanic sediments and so was relatively enriched in water (higher PH2O), which facilitated the crystallization of biotite (ASI ¼ 1e1.5, Bilal and Giret, 1999). Biotite that commonly replaces hornblende in Dadnah tonalites incorporates the remaining Al from the residual melt (Frost and Frost, 2008) and its crystallization, at T z 570 C (Table 2) differentiate the tonalite melt to evolve to a more leucocratic and less fractionated peraluminous type (ASI 1.08). 7.2. Effect of water content and alteration on tonalites A small volume of water can significantly increase the degree of partial melting for the generation of the plagiogranites (Koepke et al., 2004). The melt inclusion glass in our plagioclase obtained from the uppermost part of mantle sequence contains 2.1e2.7 equivalent wt.% H2O (Table 2), showing that considerable amount of water was involved during the fractional crystallization. From our mineralogy results, the plagioclase, hornblende and biotite
mainly contribute to the aluminous character of the Dadnah tonalites. Hornblende and biotite can derived from the partial melting of metasedimentary rocks, but our hydrogen and oxygen isotopes (e.g., d18OH2O ¼ 7.3&e7.8& and d18DH2O ¼ e95.2& to e77.1&, Table 4), suggest that these minerals are associated with typical magmatic fluids (Fig. 9a). Therefore, we suggest that the mafic minerals of the tonalites are most probably related to slabderived fluids (Fig. 9a and c). The low d18OH2O values obtained from quartz and calcite (Table 4) coupled with the Y/Ho ratios that are w 100 (Table 1) suggest that some of the tonalites may have interacted with seawater (Bau et al., 1996). The influx of seawater and its interaction with the tonalites resulted in the precipitation of the late quartz-calcite veins at T w430 C (based on the temperatures obtained from calcite-quartz isotopic pair, Table 4). Possibly, few sub-localities in our study area (i.e. sub-locality D6 and D7; Fig. 2) associate to water excess through fault zones. The intense talc and serpentine assemblages observed along these faults indicate that the fluids used these faults as pathways. Our geothermobarometry results suggest that alteration occurred at temperatures of w230 C. Therefore, the Dadnah tonalites appear to have experienced minor excess from seawater and are affected only locally where fault zones developed. 7.3. Genetic model for Dadnah tonalites The plagiogranites can formed at any level in the ophiolite section, through various magmatic processes with diverse source origin (see Rollinson, 2009, 2015). The variation of local conditions influences significantly the plagiogranite formation and thus, is hard to define a simple genetic model. In previous plagiogranite related genetic models, the source was suggested by: (1) the location where the plagiogranite melts generated (i.e., spreading axis or off-axis), (2) the structural level of emplacement (mantle or crustal section of ophiolites) and (3) the variety of magmatic processes, such as partial melting, fractional crystallization, magma mixing and contamination. In order to propose a genetic model for the Dadnah tonalites, we briefly summarize the already published genetic models for the northern and other parts of the Semail ophiolite, e.g., Searle and Cox (1999), Rollinson (2009), Grimes et al. (2013), France et al. (2014), Haase et al. (2015, 2016). The genesis of felsic crustal intrusive rocks in the Semail ophiolite was documented by Haase et al. (2016), where they suggest that these granitoids formed by fractional crystallization from mafic melts that were genetically related. These melts formed by the fluid-induced melting of the mantle wedge and mixing with subducted sediments and emplaced on the crustal sequence under crustal extension (Haase et al., 2016). For the mantle leucogranites Haase et al. (2015) argue that they represent partial melts of pelagic sediments from the subducted slab. The Dadnah tonalites are emplaced on the uppermost part of the mantle sequence of the northernmost block of the ophiolite during compression, while the primary melts of Dadnah tonalite generated by w12% partial melting of the oceanic slab (Fig. 8d) with minor contribution from mantle and sediments. Such a difference in the source contribution between our study and Haase et al. (2015, 2016) suggests spatial heterogeneity and variety of processes in the genesis of the felsic granitoids along the different localities along the Semail ophiolite. The genetic model of Rollinson (2009) included tonalite-togranite samples from both the crustal and mantle sequence of the ophiolite. He argues that the mantle-hosted granitoids formed due to either mixing or contamination at the upper part of the depleted mantle during the early stage of ophiolite emplacement (i.e., from Wadi Hamaliya, Oman). Mixing between 10% and 30% of a metasedimentary melt into the melt of a mafic source is considered as the main genetic process for the tonalites-to-granites in the mantle
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sequence of Oman ophiolite (Rollinson, 2015). This model bears many similarities to our results in the aspect of the source involvement that contributed to the formation of the tonalite melts. However, the Dadnah tonalites have quite different geochemical character and source component contribution, related to the melting of the subducted slab. Searle and Malpas (1982) and Searle and Cox (1999) reported centimeter-to-meter scale tonalitic melt pods and veins, formed by 5%e10% partial melting of the subducted continental crustal basement at the late stages of obduction. Furthermore, Briqueu et al. (1991) based on Sr, Nd and Pb isotopes obtained from felsic granitoids of the mantle sequence in the northeastern Oman ophiolite, suggested anatexis of the metamorphic sole as source for the formation of these granitoids. An alternative model for the genesis of the shallow level intrusive rocks (tonalites-trondhjemites) at/or near the dike-
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gabbro transition involves seawater circulation from an active hydrothermal system in close proximity to the active-melt lenses at the top of the magma chamber (e.g. Grimes et al., 2013; France et al., 2014). However, this model aims to explain the generation of the intrusives at the crustal sequence of the ophiolite. The seawater can infiltrate the deeper level gabbros in a ductile regime through oceanic detachment faults (e.g. Koepke et al., 2007), but is unlikely to reach and react with the melt-lenses in Moho level due to the higher depth. Moreover, Nicolas et al. (2000) reported absence of a magma chamber or mantle diapir in the Dadnah area according to their structural mapping in the Semail ophiolite. Based on our geochemical, isotope and geochronology data (Fig. 10), complemented with new melt inclusion data, we make an attempt to propose a tectono-petrogenetic model for the northern Khor Fakkan ophiolite block with reference to the source
Figure 10. Summary chart for the ages of the Semail Ophiolite, plagiogranite intrusions and metamorphic sole. Our calculated Rb/Sr isochron ages from sample PG 6_1 on (a) pyroxene (coarse and fine grained), (b) plagioclase (coarse and fine grained) and (c) quartz range from 98.1e95.4 Ma. In general our dating results span between 98.5 Ma and 94.9 Ma (Red square -7-, see also Table 5). The green area indicates the possible time interval for the formation of plagiogranites, based on the published ages for the plagiogranites from Tilton et al. (1981) and Warren et al. (2005). The green arrow indicates the maximum age range from subduction-related metamorphism (110 Ma). Blue arrow indicates the metamorphism of the sole. The obtained ages from previous studies are: (1) and (2) U/Pb plagiogranite ages (93.5e95.9 Ma and 97e97.9 Ma) from Tilton et al. (1981). (3) SmeNd isochron age (98.8 9.5 Ma) from Cox et al. (1999). (4) 40Ar/39Ar on garnet-bearing granite in UAE (89.1 Ma and 89.9 0.4 Ma) and (5) from hornblende-bearing granite (93.8 0.2 Ma) from Hacker et al. (1996). (6) U/Pb resampling of the plagiogranite of Tilton et al. (1981) (93e83 Ma, 95.5 0.24 Ma, and 95.16 0.2 Ma) from Warren et al. (2005). (8) The oldest age of ophiolite from McCulloch et al. (1981). (9) 40Ar/39Ar on hornblende from amphibolite (95.7e92.4 Ma) from Hacker (1994). (10) K/Ar on metamorphic hornblende from the amphibolite with ages in the wide range 101 Ma to 89 Ma that cluster age near 98 Ma of Gnos and Peters (1993). (11) 40Ar/39Ar on mafic dike intruding the metamorphic sole and lowest part of the mantle sequence (93.7 0.8 Ma) from Hacker and Gnos (1997). (12) 40Ar/39Ar on muscovite from metamorphic sole (92.4 0.2 Ma) from Hacker and Gnos (1997). (13) Hornblende from the metamorphic sole in northern part of the Semail ophiolite (94.9 0.2 Ma) and (14) in the southern part (93.5 0.1 Ma) from Hacker and Gnos (1997). (15) K/Ar biotite from the metamorphic sole (89.2 0.4 Ma) of Gnos and Peters (1993). (16) U/Pb age (99.8 3.3 Ma) of peraluminous granite dikes from zircon cores and rims from Spencer et al. (2017). (17) New high precision UePb dates on zircons (96.169 0.022 Ma to 96.146 0.035 Ma) from Sumeini sole (Rioux et al., 2016) and (17*) UePb dates on zircons (94.815 0.030 Ma) from Wadi Tayin sole (Rioux et al., 2016). Stars in the plagiogranite field are the ages of 98.5 Ma of the composite sill of Ra’s Dadnah and 95.3 Ma trondhjemite. The ages of tonalites-trondhjemites (stars), metamorphism of the sole (blue arrow), exhumation (80e70 Ma) and end of obduction (72 Ma) are obtained from McCulloch et al. (1980, 1981), Gregory and Taylor (1981), Stakes and Taylor (1992), Hacker and Gnos (1997), Cox et al. (1999), Warren et al. (2005), Grimes et al. (2013), Searle et al. (2015), Rioux et al. (2016).
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contribution, the water effect and related magmatic evolution processes (Fig. 11). The subducted older and denser oceanic crust (Triassic to late Cretaceous Haybi complex) was overridden by young hot asthenospheric mantle providing the thermal gradient that facilitated melting at shallower than the expected depth. That intra-oceanic subduction was more evolved in the northern ophiolite blocks (Khor Fakkan and Aswad blocks) compared to the central and south blocks, based on the gradual increase in volume phase 2 magmatism, the influence of fluids from the subducting slab (Goodenough et al., 2014) and the timing of ophiolite emplacement onto the continental margin at 93e92 Ma (Styles et al., 2006) in the north compared to the younger emplacement in the south (82e79 Ma; Warren et al., 2005; Fig. 10). Subduction appears to be almost contemporaneous with the operation of the fast-spreading ridge, based on the almost continuous magmatism from MORB-like composition (phase 1) to the more hydrous island arc-type composition (phase 2) related to the SSZ setting (Goodenough et al., 2014; Haase et al., 2016). During the evolution
of this complex tectono-magmatic process, the input of small amount of sediments on the subducted slab led to the mixing of the recycled oceanic crust (i.e., upper basaltic lavas and lower gabbros) which interacted with the mantle wedge (e.g. Haase et al., 2016), but with the influence of lithospheric mantle being relatively minor. Cox (2000), using mixing calculations on subducted metabasalt (i.e. metamorphic equivalents of Haybi volcanics), suggested that Ra-Dibba monzogranites contain w65% of the basaltic endmember, the Ra-Dadnah composite dikes 30%e50%, and the Wadi Hulw dikes less than 5%, emphasizing the high degree of heterogeneity of the source. Additionally, Rollinson (2015) proposed mixing of 10%e30% metasedimentary melt into a mafic source. Our estimations suggest that the tonalites of Dadnah were the end result of partial melting and subsequent contamination and mixing of w4% oceanic sediments with w96% oceanic lithosphere from the subducted slab. This MORB-type slab melt composed from w97% recycled oceanic crust and w3% of the overlying mantle. Partial melting of the slab assisted also by the fluids released from the
Figure 11. (a) A proposed schematic petrogenetic model for the formation of the Dadnah tonalites intrusive in the mantle sequence at the northern part of the Semail ophiolite. The anatexis of recycled oceanic crust (upper part of slab with basaltic primary composition mixed with minor sedimentary component) produced the off-axis melts and their subsequent minor interaction and mixing with the overlying mantle produced the tonalite melts. The melts recorded temperatures from 720 C to 680 C, ascent and pond at the upper part of mantle, just below the Moho transition zone (plate setting modified from Cox et al., 1999, Goodenough et al., 2014; Searle et al., 2015, Roberts et al., 2016; Soret et al., 2017). (b) Schematic illustration for the process of the tonalite generation at the upper part of the subducting slab. Oceanic lithosphere and sediments on the slab mantle interface can melt at low pressures (3e7 kbar, i.e. 10e20 km; Kostopoulos and Murton, 1992) due to unusually hot conditions within the mantle wedge during the first 5e10 Myr after subduction initiation (Stern et al., 2012). The higher degree of depletion in the lithospheric mantle of the SSZ ophiolite is balanced by the introduction in the subduction zone of older oceanic enriched crust (Haybi complex and below) with the associated sediments carried above the slab. The discontinuous sediment cover along and across the slab interface can be responsible for the variety of granitoid types and ages observed. On the slab interface, the metamorphic sole accretion is controlled by the percentage of partial melting. In our case for the northern part, the sole accretion is inhibited and mechanical decoupling is favored triggering less sedimentary input and more oceanic crust to be involved in the partial melt (see also Soret et al., 2017).
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hydrous subducted sediments. We imply that the composition and thickness of the subducted slab and its sedimentary cover was highly heterogeneous laterally and vertically, consisting by older oceanic crust (mainly of Late Triassic exotic reefal limestones (seamount) above an basaltic substrate, cherts, and mélanges; Haybi complex) and was discontinuously overlain by marine sediments (Searle et al., 2015). Another potential source could be the metabasalts and metagreywackes comprising the basement unit (Hatat fm.) below the Haybi and Hawasina allocthons (Gray and Gregory, 2003). Oceanic crust and sediments on the slab mantle interface can melt at low pressures (3e7 kbar, i.e. 10e20 km; Kostopoulos and Murton, 1992) due to unusually hot conditions within the mantle wedge during the first 5e10 Myr after subduction initiation (Stern et al., 2012). Mantle melting is enhanced due to SSZ fluid input (Pearce et al., 1984). The higher degree of depletion in the lithospheric mantle of the SSZ ophiolite is balanced by the introduction in the subduction zone of the older oceanic enriched crust (Haybi complex) with the associated sediments carried above the slab. The discontinuous sedimentary cover along and across the slab interface can be responsible for the variety of granitoid types and ages observed along the ophiolite structural grain (from UAE to Oman) and across the oceanic mantle and crust (crustal and mantle plagiogranites). Soret et al. (2017) argued that on the slab interface, the metamorphic sole accretion is controlled by the percentage of melting since low melt fractions < 7% do not affect the inter-plate viscous coupling and only contribute to strain localization and weakening on the detachment of the sole thrust slices. For melt fractions >7% (w12% partial melting in our case) sole accretion is inhibited and mechanical decoupling is favored (Soret et al., 2017). This can explain the low percentage of sedimentary component in our tonalite melt, since sole accretion stacks up thicker thrust slices of metasediments, while in the decoupling the oceanic crust carries only thinner sedimentary layers. Partial melting and mixing followed by plagioclase accumulation and fractional crystallization and interaction with the host harzburgite. The melts ascended to higher structural levels (pressure difference from 6.5 kbar to 4.4 kbar, corresponding to w7 km upward flow, Table 2) and ponded at the uppermost part of mantle sequence and in the MTZ of the Khor Fakkan block, which represents a permeability barrier, forming the unique in geochemical and isotopic character Dadnah tonalites. Our calculated Rb/Sr isochron ages for the tonalites yielded a range between 98.6 Ma and 94.9 Ma (Fig. 10, Table 5) which indicates that the crystallization of these melts predates slightly the suggested 96.1e94.8 Ma age range of the metamorphic sole (Rioux et al., 2016). In addition, our ages obtained from sample PG6_1 (Fig. 10a, b and c; Table 5) in which pyroxene has a weighted mean Rb/Sr isochron of 97.86 1.5 Ma (Mean Square of Weighted Deviates, MSWD ¼ 0.85; Fig. 10a), plagioclase of 98.1 1 Ma (MSWD ¼ 0.95; Fig. 10b) and quartz of 95.4 1 Ma (MSWD ¼ 0.97; Fig. 10c). Spencer et al. (2017) reported ages (UePb on zircon) of 99.8 3.3 Ma obtained from Stype, granite dikes that intruded the Semail peridotite. The dates from quartz are identical with the ages (e.g., 96.1e94.0 Ma and within a narrow age range of 95.3 0.5 Ma) reported from Warren et al. (2005), Goodenough et al. (2014), and Rioux et al. (2016). These authors have dated zircon grains from several tonalite and trondhjemite intrusions along the length of the ophiolite. We attribute the older ages obtained to possibly record the partial melting event, i.e., relate to relict plagioclases of higher An contents (An53e56; Fig. 3g) and pyroxenes or to depict tonalites genetically related to phase 1 magmatism (Rioux et al., 2013; Goodenough et al., 2014). Our Sm/Nd isochron of 96.5 1 Ma on Hbl (MSWD ¼ 0.80; Table 5) and 96.3 1 Ma on Bt (MSWD ¼ 1.12; Table 5) are very close to the older ages (e.g. 96.1 Ma) for the tonalites given by Rioux et al. (2016). Our geochronological data,
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despite some analytical uncertainties, leave also a hint that the inta-oceanic subduction might initiated a bit earlier than the ophiolite generation at the spreading ridge. Since the ages of the metamorphic sole, the ophiolite formation and the plagiogranite melt are very close and overlapping, a mechanism including partial melting of the subducted slab with synchronous formation and exhumation for the metamorphic sole as a direct consequence of spreading above a nascent subduction zone is required, explaining in a way the varied estimations for the sedimentary input in the tonalite melts. 8. Conclusions Our concluding remarks based on the synthesis and analysis of geological, mineralogical, petrological, geochemical and isotopic data in the Khor Fakkan ophiolite block in the Eastern Emirates can be summarized as follows: (1) The Dadnah tonalites, intrusive in the mantle section of the Khor Fakkan block, are high siliceous, low-potassic, metaluminous to weakly peraluminous, enriched in LILE, Ta, La, Ce, depleted in U, and occur with low d18OH2O, moderate εSr and negative εNd values. Geochemical and isotopic compositions link our tonalites with magmatic processes occurring at the initial stages of a supra-subduction zone setting. (2) Melt inclusions hosted in plagioclase indicated the leastfractionated tonalites and the PeT conditions (T ¼ 664e733 C, P ¼ 4.4e4.6 kbar) for their formation. (3) The Rb/Sr isochron suggest an age range of 98.6e94.9 Ma for the Dadnah tonalites. (4) The Dadnah tonalites formed by partial melting (w10%e15% continuous or w 12% batch partial melting), accumulation of plagioclase, fractional crystallization (w55%e57%), and mild interaction with the host harzburgite. (5) The stable and noble gas (He) isotopes indicate that the tonalites were formed by melting of mainly recycled oceanic crust that was hydrothermally altered. (6) From the estimation of source contributions, the tonalite melt evolved from w96% of oceanic lithosphere and w4% of oceanic sediments. The oceanic lithosphere was mainly recycled oceanic crust (w97%) with minor component from the overlying mantle (w3%). Acknowledgements We express our appreciation to the Petroleum Institute (part of Khalifa University, UAE) for providing us access to the ADRIC lab facilities and field support during this study. Especially, we would like to thank Dr. Kil and his staff of Chonnam National University, S. Korea for the permission to use their ICP-MS facility. We also thank the staff of the USGS Denver Inclusion Analysis Laboratory for the melt inclusions microthermometry and EPMA analyses, and the Chinese Academy of Geological Sciences (CAGS), Beijing, China for the isotopic analyses. This work is part of H.J.’s MSc thesis project (PIPGSTS-16-16) conducted at the Petroleum Institute, UAE. H. Rollinson and three anonymous reviewers are warmly thanked for their constructive reviews and suggestions and S. Kwon for his editorial handling that lead to a significant improvement of the final version of the manuscript. References Ahmad, T., Tarney, J., 1991. Geochemistry and petrogenesis of Garhwal volcanics: implications for evolution of the North India lithosphere. Precambrian Research 50, 69e88.
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Alabaster, T., Pearce, J.A., Malpas, J., 1982. The volcanic stratigraphy and petrogenesis of the Oman ophiolite complex. Contributions to Mineralogy and Petrology 81, 168e183. Allégre, C.J., Sutcliffe, C., 2008. Isotope Geology. Cambridge University Press, 512 p. Alleman, F., Peters, T., 1972. The ophiolite-radiolarite belt of the North Oman mountains. Eclogae Geologicae Helvetiae 65, 657e698. Altherr, R., Holl, A., Hegner, E., Langer, C., Kreuzer, H., 2000. High potassium, calcalkaline I-type plutonism in the European Variscides: Northern Vosges (France) and Northern Schwarzwald (Germany). Lithos 50, 51e73. Amri, I., Benoit, M., Ceuleneer, G., 1996. Tectonic setting for the genesis of oceanic plagiogranites: evidence from a paleo-spreading structure in the Oman ophiolite. Earth and Planetary Science Letters 139, 177e194. Amri, I., Ceuleneer, G., Benoit, M., Python, M., Puga, E., Targuisti, K., 2007. Genesis of granitoids by interaction between mantle peridotites and hydrothermal fluids in oceanic spreading setting in the Oman Ophiolite. Geogaceta 42, 23e26. Anonymous, 1972. Penrose field conference on ophiolites. Geotimes 17, 24e25. Asimow, P.D., Ghiorso, M.S., 1998. Algorithmic modifications extending MELTS to calculate subsolidus phase relations. American Mineralogist 83, 1127e1132. Ashchepkov, I.V., Vladykin, N.V., Gerasimov, P.A., Saprykin, A.I., Khmel’nikova, O.S., Anoshin, G.N., 2001. Temperature gradient and structure of the lithospheric block beneath the southeastern margin of the Siberian craton: disintegrated xenolith evidence from kimberlitic pipes of the Aldan shield. Doklady Earth Sciences 378, 495e499. Ashchepkov, I.V., 2002. Empirical clinopyroxene thermobarometers for mantle rocks based on the diopside-jadeite exchange. Doklady Earth Sciences 382, 366e370. Ashchepkov, I.V., 2003. More Precise Equation of the Jd-en Barometer. Herald of the Department of Earth Sciences RAS, p. 21. Brophy, J.G., 2008. A study of rare earth element (REE)eSiO2 variations in felsic liquids generated by basalt fractionation and amphibolite melting: a potential test for discriminating between the two different processes. Contributions to Mineralogy and Petrology 156, 337e357. Barker, F., 1979. Trondhjemite: definition, environment and hypotheses of origin. In: Barker, F. (Ed.), Trondhjemites, dacites, and related rocks. Amsterdam, Elsevier. Developments in Petrology 6, 1e12. Bau, M., Koschinsky, A., Dulski, P., Heinz, J., 1996. Comparison of the partitioning behaviours of yttrium, rare-earth elements and titanium between hydrogenic marine ferromanganese crusts and seawater. Geochimica et Cosmochimica Acta 60, 1709e1725. Beaudoin, G., Therrien, P., 2009. The updated web stable isotope fractionation calculator. In: De Groot, P.A. (Ed.), Handbook of Stable Isotope Analytical Techniques 2. Elsevier, pp. 1120e1122. Bechennec, F., Le Metour, J., Rabu, D., Bourdillon-de-Grissac, C., de Wever, P., Beurrier, M., Villey, M., 1990. The Hawasina Nappes; stratigraphy, paleogeography and structural evolution of a fragment of the South-Tethyan passive continental margin. In: Rorbertson, A.H.F., Searle, M.P., Ries, A.C. (Eds.), The Geology and Tectonics of the Oman Region, vol. 49. Geological Society, London, Special Publications, pp. 213e223. Bilal, E., Giret, A., 1999. The aluminum saturation index and the MgO/TiO2 ratio: two parameters influenced by PH2O and their use to discriminate magma series. Revista Brasileira de Geociencias 29, 55e58. Blundy, J.D., Holland, T.J.B., 1990. Calcic amphibole equilibria and a new amphiboleplagioclase geothermometer. Contributions to Mineralogy and Petrology 104, 208e224. Boudier, F., Ceuleneer, G., Nicolas, A., 1988. Shear zones, thrusts and related magmatism in the Oman ophiolite: initiation of thrusting at an oceanic ridge. Tectonophysics 151, 275e296. Bowman, J.R., 1998. Stable-isotope systematics of skarns. In: Lentz, D.R. (Ed.), Mineralized Intrusion-related Skarn Systems, vol. 26. Mineralogical Association of Canada Short Course Series, pp. 99e145. Briqueu, L., Mevel, C., Boudier, F., 1991. Sr, Nd and Pb isotopic constraints in the genesis of a calc-alkaline plutonic suite in the Oman ophiolite related to the obduction process. In: Peters, T.J., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer, Netherlands, pp. 517e542. Buddington, A.F., Lindsley, D.H., 1964. Iron-titanium oxide minerals and synthetic equivalents. Journal of Petrology 5, 310e335. Carmody, L., Barry, P.H., Shervais, J.W., Kluesner, J.W., Taylor, L.A., 2013. Oxygen isotopes in subducted oceanic crust: a new perspective from Siberian diamondiferous eclogites. Geology, Geophysics and Geosystems 14, 3479e3493. https://doi.org/10.1002/ggge.20220. Cathelineau, M., 1988. Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals 23, 471e485. Cathelineau, M., Nieva, D., 1985. A chlorite solid solution geothermometer: the Los Azufres (Mexico) geothermal system. Contributions to Mineralogy and Petrology 91, 235e244. Chaussidon, M., Sheppard, S.M., Michard, A., 1991. Hydrogen, sulfur and neodymium isotope variations in the mantle beneath the EPR at 12 50΄N. In: Taylor, H.P., O’Neil, J.R., Kaplan, I.R. (Eds.), Stable Isotope Geochemistry: A Tribute to Samuel Epstein. Geochemical Society Special Publications, pp. 325e337. Clayton, R.N., Mayeda, T.K., 1963. The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochimica et Cosmochimica Acta 27, 43e52. Clayton, R.N., Muffler, L.J.P., White, D.E., 1972. Oxygen isotope fractionation study of calcite and silicates of the River Ranch, California. American Journal of Science 266, 968e979.
Coleman, R.G., 1981. Tectonic setting for ophiolite obduction in Oman. Journal of Geophysical Research 86, 2497e2508. Coleman, R.G., Peterman, Z.E., 1975. Oceanic plagiogranite. Journal of Geophysical Research 80, 1099e1108. Cox, J.S., 2000. Subduction-obduction Related Petrogenetic and Metamorphic Evolution of the Semail Ophiolite Sole in Oman and the United Arab Emirates (Ph.D. thesis). University of Oxford, Oxford, UK, 169 p. Cox, J., Searle, M., Pedersen, R., 1999. The petrogenesis of leucogranitic dykes intruding the northern Semail ophiolite, United Arab Emirates: field relationships, geochemistry and Sr/Nd isotope systematics. Contributions to Mineralogy and Petrology 137, 267e287. De Paolo, D.J., 1980. Crustal growth and mantle evolution; Inferences from models of element transport and Nd and Sr isotopes. Geochimica et Cosmochimica Acta 44, 85e96. Ding, T.P., 2004. Analytical methods for silicon isotope determinations. In: De Groot, P.A. (Ed.), Handbook of Stable Isotope Analytical Techniques. Elsevier, pp. 523e537. Dixon, S., Rutherford, M.J., 1979. Plagiogranites as late-stage immiscible liquids in ophiolite and mid-ocean ridge suites: an experimental study. Earth and Planetary Science Letters 45, 45e60. Ebadi, A., Johannes, W., 1991. Beginning of melting and composition of first melts in the system Qz-Ab-Or-H2O-CO2. Contributions to Mineralogy and Petrology 106, 286e295. Ernewein, M., Pflumio, C., Whitechurch, H., 1988. The death of an accretion zone as evidenced by the magmatic history of the Sumail ophiolite (Oman). Tectonophysics 151, 247e274. Ersoy, Y., Helvaci, C., 2010. FC-AFC-FCA and mixing modeler: a Microsoft Excel spreadsheet program for modeling geochemical differentiation of magma by crystal fractionation, crustal assimilation and mixing. Computers and Geosciences 36, 383e390. France, L., Koepke, J., MacLeod, C., Ildefonse, B., Godard, M., Deloule, E., 2014. Contamination of MORB by anatexis of magma chamber roof rocks: constraints from a geochemical study of experimental melts and associated residues. Lithos 202, 120e137. Friedman, I., O’Neil, J.R., 1977. Compilation of Stable Isotope Fractionation Factors of Geochemical Interest, 440KK. U.S. Geological Survey Professional Paper, p. 12. Frost, B.R., Frost, C.D., 2008. A geochemical classification for feldspathic igneous rocks. Journal of Petrology 49, 1955e1969. Ghent, E.D., Stout, M.Z., 1981. Metamorphism at the base of the Semail ophiolite. Journal of Geophysical Research 86, 2557e2571. Glennie, K.W., Boeuf, M.G., Hughes-Clarke, M.H.W., Moody-Stuart, M., Pilaar, W.F., Reinhardt, B., 1974. Geology of the Oman Mountains. Verhandelingen Koninklijk, The Nederlands, Geologisch Mijnbouwkundig Genootschap, p. 31. Gnos, E., Nicolas, A., 1996. Structural evolution of the northern end of the Oman Ophiolite and enclosed granulites. Tectonophysics 254, 111e137. Gnos, E., Peters, T., 1993. K-Ar ages of the metamorphic sole of the Semail ophiolite; implications for ophiolite cooling history. Contributions to Mineralogy and Petrology 113, 325e332. Graham, D.W., Barry, B., Hanan, B.W., Hémond, Ch, Blichert-Toft, J., Albarède, F., 2014. Helium isotopic textures in Earth’s upper mantle. Geology, Geophysics and Geosystems 15, 2048e2074. https://doi.org/10.1002/2014GC005264. Gray, D.R., Gregory, R.T., 2003. Ophiolite obduction and the Samail Ophiolite: the behaviour of the underlying margin. Ophiolites in Earth History 218 (1), 449e465. https://doi.org/10.1144/GSL.SP.2003.218.01.23. Goodenough, K.M., Phillips, E.R., Styles, M.T., Thomas, R.J., Farrant, A.R., Arkley, S.L.B., 2006. Geology of the Khor Fakkan 1:50000 Map Sheet. British Geological Survey, United Arab Emirates, pp. 1e84. Goodenough, K.M., Thomas, R.J., Styles, M.T., Schofield, D.I., MacLeod, C.J., 2014. Records of ocean growth and destruction in the Oman-UAE ophiolite. Elements 10, 109e114. Gregory, R.T., Taylor Jr., H.P., 1981. An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail Ophiolite, Oman: evidence for d18O buffering of the oceans by deep (> 5 km) seawater-hydrothermal circulation at mid-ocean ridges. Journal of Geophysical Research 86, 2737e2755. Grimes, C.B., Ushikubo, T., Kozdon, R., Valley, J.W., 2013. Perspectives on the origin of plagiogranite in ophiolites from oxygen isotopes in zircon. Lithos 179, 48e66. Gao, P., Zheng, Y.-F., Zhao, Z.-F., 2016. Distinction between S-type and peraluminous I-type granites: Zircon versus whole-rock geochemistry. Lithos 258-259, 77e91. Haase, K.M., Freund, S., Koepke, J., Hauff, F., Erdmann, M., 2015. Melts of sediments in the mantle wedge of the Oman ophiolite. Geology 43 (4), 275e278. https:// doi.org/10.1130/G36451.1. Haase, K.M., Freund, S., Beier, C., Koepke, J., Erdmann, M., Hauff, F., 2016. Constraints on the magmatic evolution of the oceanic crust from plagiogranite intrusions in the Oman ophiolite. Contributions to Mineralogy and Petrology 171, 1e16. Hacker, B.R., 1994. Rapid emplacement of young oceanic lithosphere. Science 265, 1563e1565. Hacker, B.R., Mosenfelder, J.L., 1996. Metamorphism and deformation along the emplacement thrust of the Samail ophiolite, Oman. Earth and Planetary Science Letters 144, 435e451. Hacker, B.R., Gnos, E., 1997. The conundrum of samail: explaining the metamorphic history. Tectonophysics 279, 215e226. Hacker, B.R., Mosenfelder, J.L., Gnos, E., 1996. Rapid emplacement of the Oman ophiolite: thermal and geochronological constraints. Tectonics 15, 1230e1247.
H. Joun et al. / Geoscience Frontiers 10 (2019) 1187e1210 Hammarstrom, J.M., Zen, E., 1986. Aluminum in hornblende: an empirical igneous geobarometer. American Mineralogist 71, 1297e1313. Helgeson, H.C., Kirkham, D.H., Flowers, G.C., 1981. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures: calculation of activity coefficients, osmotic coefficients, and apparent molal, standard, and relative partial molal properties to 600 C and 5 kb. American Journal of Science 281, 1249e1516. Hemond, C., Hofmann, A.W., Vlastelic, I., Nauret, F., 2006. Origin of MORB enrichment and relative trace element compatibilities along the Mid-Atlantic Ridge between 10 and 24 N. Geochemistry, Geophysics, Geosystems 7, Q12010. https://doi.org/10.1029/2006GC001317. Henry, D.J., Guidotti, C.V., Thomson, J.A., 2005. The Ti-saturation surface for low-to medium pressure metapelitic biotites: implications for geothermometry and Tisubstitution mechanisms. American Mineralogist 90, 316e328. Hofmann, A.W., White, W.M., 1982. Mantle plumes from ancient oceanic crust. Earth and Planetary Science Letters 57, 421e436. Hofmann, A.W., 2004. Sampling mantle heterogeneity through oceanic basalts: isotopes and trace elements. Treatise in Geochemistry 2, 61e101. Holtz, F., Pichavant, M., Barbey, P., Johannes, W., 1992. Effects of H2O on liquidus phase relations in the haplogranitic system at 2 and 5 kbar. American Mineralogist 77, 1223e1241. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Canadian Journal of Earth Sciences 8, 523e548. Ito, E., Stern, R.J., 1986. Oxygen- and strontium-isotopic investigations of subduction zone volcanism: the case of the Volcano Arc and the Mariana Island Arc. Earth and Planetary Science Letters 76, 312e320. Jacobs, J., Thomas, R.J., Ksienzyk, A.K., Dunkl, I., 2015. Tracking the Oman ophiolite to the surface e new fission track and (U-Th)/He data from the Aswad and Khor Fakkan Massifs, United Arab Emirates. Tectonophysics 644, 68e80. Johnson, J.W., Oelkers, E.H., Helgeson, H.C., 1992. SUPCRT92: a software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species and reactions from 1 to 5000 bars and 0 to 1000 C. Computers and Geoscience 18, 899e947. Johnson, M.C., Rutherford, M.J., 1989. Experimental calibration of the aluminum-inhornblende geobarometer with application to Long Valley caldera (California) volcanic rocks. Geology 17, 837e841. Jowett, E.C., 1991. Fitting iron and magnesium into the hydrothermal chlorite geothermometer. In: GAC/MAC/SEG Joint Annual Meeting, Toronto, A62, p. 16. Kelemen, P.B., Koga, K., Shimizu, N., 1997. Geochemistry of gabbro sills in the crust/ mantle transition zone of the Oman ophiolite: implications for the origin of the oceanic lower crust. Earth and Planetary Science Letters 146, 475e488. Kostopoulos, D.K., Murton, B.J., 1992. Origin and distribution of components in boninite genesis: significance of the OIB component, 60. Geological Society, London, Special Publications, pp. 133e154. Koepke, J., Feig, S.T., Snow, J., Freise, M., 2004. Petrogenesis of oceanic plagiogranites by partial melting of gabbros: an experimental study. Mineralogy and Petrology 146, 414e432. Koepke, J., Berndt, J., Feig, S.T., Holtz, F., 2007. The formation of SiO2-rich melts within the deep oceanic crust by hydrous partial melting of gabbros. Contributions to Mineralogy and Petrology 153, 67e84. Kranidiotis, P., MacLean, W.H., 1987. Systematics of chlorite alteration at the Phelps Dodge massive sulfide deposit, Matagami, Quebec. Economic Geology 82, 1898e1911. lu-Kus¸cu, G., Kus¸cu, I., Yılmazer, E., Güleç, N., Bayır, S., Demirela, G., ençaliog Gezer, G., Kuru, S., Kaymakcı, N., 2011. U-Pb and 40Ar-39Ar Geochronology and isotopic constraints on the genesis of copper-gold-bearing iron oxide deposits in the Hasançelebi District, Eastern Turkey. Economic Geology 106, 261e288. Lee, C.-T.A., Morton, D.M., 2015. High silica granites: Terminal porosity and crystal settling in shallow magma chambers. Earth and Planetary Science Letters 409, 23e31. Lepage, L.D., 2003. ILMAT: an excel worksheet for ilmenite-magnetite geothermometry and geobarometry. Computer Geoscience 29, 673e678. Lippard, S.J., Shelton, A.W., Gass, I.G., 1986. The Ophiolite of Northern Oman. The Geological Society Memoir no. 11. Blackwell Scientific Publications, Oxford, p. 178. Liu, Y., He, D., Gao, C., Foley, S., Gao, S., Hu, Z., Zong, K., Chen, H., 2015. First direct evidence of sedimentary carbonate recycling in subduction-related xenoliths. Scientific Reports 5, 11547. https://doi.org/10.1038/srep11547. Ludwig, K.R., 2001. Users Manual for Isoplot/Ex, rev. 2.49. A Geochronological Toolkit for Microsoft Excel, vol. 1a. Berkeley Geochronology Center, Special Publication, pp. 1e55. Lustrino, M., Anderson, D.L., 2015. The mantle isotopic printer: basic mantle plume geochemistry for seismologists and geodynamicists. In: Foulger, G.R., Lustrino, M., King, S.D. (Eds.), The Interdisciplinary Earth: A Volume in Honor of Don L. Anderson, vol. 514. Geological Society of America Special Paper, pp. 257e279. Malpas, J., 1979. Two contrasting plagiogranite associations from transported ophiolites in western Newfoundland: initial report. In: Barker, F. (Ed.), Trondhjemites, Dacites and Related Rocks. Elsevier, Amsterdam, pp. 465e487. Maniar, P.D., Piccoli, P.M., 1989. Tectonic discrimination of granitoids. Geological Society of America Bulletin 101, 635e643. Martin, H., Smithies, R.H., Rapp, R., Moyen, J.F., Champion, D., 2005. An overview of adakite, tonalite-trondhjemite-granodiorite (TTG), and sanukitoid: relationships and some implications for crustal evolution. Lithos 79, 1e24. McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., Taylor Jr., H.P., 1980. A neodymium, strontium, and oxygen isotopic study of the Cretaceous Samail Ophiolite and
1209
implications for the petrogenesis and seawater-hydrothermal alteration of oceanic crust. Earth and Planetary Science Letters 46, 201e211. McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., Taylor Jr., H.P., 1981. Sm-Nd, Rb-Sr and 18O-16O isotropic systematics in an oceanic crustal section; evidence from the Semail ophiolite. Journal of Geophysical Research 86, 2721e2735. McDonough, W.F., Sun, S., 1995. The composition of the Earth. Chemical Geology 120, 223e253. Michael, P.J., McDonough, W.F., Nielsen, R.L., Cornell, W.C., 2002. Depleted melt inclusions in MORB plagioclase: messages from the mantle or mirages from the magma chamber? Chemical Geology 183 (1e4), 43e61. https://doi.org/10.1016/ S0009-2541(01)00371-0. Middlemost, E., 1994. Naming materials in the magma/igneous rock system. Earth Science Reviews 37, 215e224. Moore, D.M., Reynolds, R.C., 1997. X-Ray diffraction and the identification and analysis of clay minerals. Oxford University Press, Oxford, p. 378. Nicolas, A., 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Kluwer Academic, Amsterdam, p. 369. ISBN 978-94-010-7569-5. Nicolas, A., Ildefonse, B., Boudier, F., Lenoir, X., Ben Ismail, W., 2000. Dike Distribution in the Oman-United Arab Emirates Ophiolite, vol. 21. Kluwer Academic Publishers, pp. 269e287. Pallister, J.S., Hopson, C.A., 1981. Samail ophiolite plutonic suite: field relations, phase variation, cryptic variation and layering, and a model of a spreading ridge magma chamber. Journal of Geophysical Research 86, 2593e2644. Pallister, J.S., Knight, R.J., 1981. Rare-earth element geochemistry of the samail ophiolite near Ibra, Oman. Journal of Geophysical Research 86, 2673e2697. Papanastassiou, D.A., DePaolo, D.J., Wasserburg, G.J., 1977. Rb-Sr and Sm-Nd chronology and genealogy of mare basalts from the Sea of Tranquility. In: Lunar Science Conference, 8th Houston, Proceedings, vol. 2. Pergamon Press, New York, pp. 1639e1672. Pearce, J.A., 1980. Geochemical evidence for the genesis and eruptive setting of lavas from Tethyan ophiolites. In: Panayiotou, A. (Ed.), Ophiolites: Proceedings of the International Ophiolite Symposium, 1979, Cyprus, Nicosia, Cyprus Ministry Agriculture Natural Resources. Geological Survey Department, pp. 261e272. Pearce, J.A., 1996. A user’s guide to basalt discrimination diagrams. In: Wyman, D.A. (Ed.), Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration, vol. 12. Geological Association of Canada Short Course Notes, pp. 79e113. Pearce, J.A., Alabaster, T., Shelton, A.W., Searle, M.P., 1981. The Oman ophiolite as a Cretaceous arcbasin complex: evidence and implications. Philosophical Transactions of the Royal Society of London. Series A, Mathematical and Physical Sciences 300 (1454), 299e317. Pearce, J.A., Harris, N.W., Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology 25, 956e983. Peters, T.J., Kamber, B.S., 1994. Peraluminous, potassium rich granitoids in the Semail ophiolite. Contributions to Mineralogy and Petrology 118, 229e238. Peters, B.J., Day, J.M.D., 2014. Assessment of relative Ti, Ta and Nb (TITAN) enrichments in ocean island basalts. Geochemistry, Geophysics, Geosystems 15. https://doi.org/10.1002/2014GC005506. Pflumio, C., 1991. Evidences for polyphased oceanic alteration of the extrusive sequence of the Semail ophiolite from the Salahi block (northern Oman). In: Peters, T., Nicolas, A., Coleman, R.G. (Eds.), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Petrology and Structural Geology 5. Kluwer Academic Publishers, Dordrecht, pp. 313e352. Pichou, J.L., Pouchoir, F., 1985. “PAP” procedure for improved quantitative microanalysis. Microbeam Analysis 20, 104e105. Pilet, S., Baker, B.M., Müntener, O., Stolper, M.E., 2011. Monte-Carlo simulations of metasomatic enrichment in the lithosphere and implications for the source of alkaline basalts. Journal of Petrology 52, 1415e1442. Pringle, E.A., Moynier, F., Savage, P.S., Jackson, M.G., Moreira, M., Day, J.M.D., 2016. Silicon isotopes reveal recycled altered oceanic crust in the mantle sources of ocean island basalts. Geochimica et Cosmochimica Acta 189, 282e295. Putirka, K., 2008. Thermometers and barometers for volcanic systems. In: Putirka, K., Tepley, F. (Eds.), Minerals, Inclusions and Volcanic Processes. Reviews in Mineralogy and Geochemistry 69, pp. 61e120. Powell, R., Powell, M., 1977. Geothermometry and oxygen barometry using coexisting iron-titanium oxides: a reappraisal. Mineralogical Magazine 41, 257e263. Python, M., Ceuleneer, G., 2003. Nature and distribution of dykes and related melt migration structures in the mantle section of the Oman ophiolite. Geology, Geophysics and Geosystems (G-Cubed) 4, 8612e8646. Rickwood, P., 1989. Boundary lines within petrologic diagrams, which use oxides of major and minor elements. Lithos 22, 247e263. Rioux, M., Bowring, S., Kelemen, P., Gordon, S., Miller, R., Dudás, F., 2013. Tectonic development of the Samail ophiolite: high-precision U-Pb zircon geochronology and Sm-Nd isotopic constraints on crustal growth and emplacement. Journal of Geophysical Research: Solid Earth 118, 2085e2101. Rioux, M., Garber, J., Bauer, A., Bowring, S., Searle, M., Kelemen, P., Hacker, B., 2016. Synchronous formation of the metamorphic sole and igneous crust of the Semail ophiolite: new constraints on the tectonic evolution during ophiolite formation from high-precision U-Pb zircon geochronology. Earth and Planetary Science Letters 451, 185e195. Roberts, N.M.W., Spencer, C.J., 2014. The zircon archive of continent formation through time. In: Roberts, N.M.W., Van Kranendonk, M., Parman, S., Shirey, S.,
1210
H. Joun et al. / Geoscience Frontiers 10 (2019) 1187e1210
Clift, P.D. (Eds.), Continent Formation through Time, vol. 389. Geological Society, London, Special Publications. https://doi.org/10.1144/SP389.14. Roberts, N.M.W., Thomas, R.J., Jacobs, J., 2016. Geochronological constraints on the metamorphic sole of the Semail ophiolite in the United Arab Emirates. Geoscience Frontiers 7, 609e619. Rollinson, H., 2009. New models for the genesis of plagiogranites in the Oman ophiolite. Lithos 112, 352. Rollinson, H., 2014. Plagiogranites from the mantle section of the Oman Ophiolite: models for early crustal evolution. In: Rollinson, H., Searle, M.P., Abbasi, I., AlLazki, A., Al Kindi, M.H. (Eds.), Tectonic Evolution of the Oman Mountains, vol. 392. Geological Society, London, Special Publications, pp. 247e261. Roedder, E., 1984. Fluid inclusions. Reviews in Mineralogy, 12. Mineralogical Society of America, p. 644. Rollinson, H., 2015. Slab and sediment melting during subduction initiation: granitoid dykes from the mantle section of the Oman ophiolite. Contributions to Mineralogy and Petrology 170, 1e20. Rudnick, R.L., Gao, S., 2003. Composition of the continental crust. In: Rudnick, R.L. (Ed.), The Crust. Elsevier, p. 170. Savelieva, G.N., Bortnikov, N.S., Bayanova, T.B., Ikorskii, S.V., Kamenskii, I.L., 2008. Sm-Nd and Rb-Sr isotopic systems and captured He and hydrocarbon gases as markers of melt sources and fluid regime under which the oceanic crust of the Mid-Atlantic Ridge was formed at 5 -6 N. Geochemistry International 46, 745e758. https://doi.org/10.1134/S0016702908080016. Searle, M.P., 1980. The Metamorphic Sheet and Underlying Volcanic Rocks beneath the Oman Ophiolite in the Northern Oman Mountains of Arabia (Unpublished) (PhD thesis). Open University, Milton Keynes. Searle, M.P., Malpas, J., 1980. Structure and metamorphism of rocks beneath the Semail ophiolite of Oman and their significance in ophiolite obduction. Transactions of the Royal Society of Edinburgh, Earth Sciences 71, 247e262. Searle, M.P., Malpas, J., 1982. Petrochemistry and origin of sub-ophiolite metamorphic and related rocks in the Oman Mountains. Journal of the Geological Society of London 139, 235e248. Searle, M.P., Cox, J., 1999. Tectonic setting, origin, and obduction of the Oman ophiolite. Geological Society of America Bulletin 111, 104e122. Searle, M.P., Cox, J., 2002. Subduction zone metamorphism during formation and emplacement of the Semail ophiolite in the Oman Mountains. Geological Magazine 139, 241e255. Searle, M.P., Ali, M., 2009. Structural and tectonic evolution of the Jabal Sumeini-Al Ain-Buraimi region, northern Oman and eastern United Arab Emirates. GeoArabia 14, 115e142. Searle, M.P., Waters, D.J., Garber, J.M., Rioux, M., Cherry, A.G., Ambrose, T.K., 2015. Structure and metamorphism beneath the obducting Oman ophiolite: evidence from the Bani Hamid granulites, northern Oman Mountains. Geosphere 11, 1812e1836. Shock, E.T., Helgeson, H.C., 1988. Calculation of the thermodynamic and transport properties of aqueous species at high pressures and temperatures: correlation algorithms for ionic species and equation of state prediction to 5 kb and 1000 C. Geochimica et Cosmochimica Acta 52, 2009e2036. Smith, P.M., Asimow, P.D., 2005. Adiabat_1ph: a new public front-end to the MELTS, pMELTS, and pHMELTS models. Geology, Geophysics and Geosystems (GCubed) 6, 1e8. https://doi.org/10.1029/2004GC000816. Q02004. Soret, M., Agard, P., Dubacq, B., Plunder, A., Yamato, P., 2017. Petrological evidence for stepwise accretion of metamorphic soles during subduction infancy (Semail ophiolite, Oman and UAE). Journal of Metamorphic Geology 35, 1051e1080. https://doi.org/10.1111/jmg.12267. Sobolev, A.V., 1996. Melt inclusions in minerals as a source of principle petrological information. Petrology 4, 209e220. Spencer, C.J., Cavosie, A.J., Raub, T.D., Rollinson, H., Jeon, H., Searle, M.P., Miller, J.A., McDonald, B.J., Evans, N.J., The Edinburgh Ion Microprobe Facility (EIMF), 2017. Evidence for melting mud in Earth’s mantle from extreme oxygen isotope signatures in zircon. Geology 45, 975e978. https://doi.org/ 10.1130/G39402.1. Spooner, E.T.C., 1977. Hydrodynamic model for the origin of the ophiolitic cupriferous pyrite ore deposits of Cyprus. Geological Society, London, Special Publications 7, 58e72. Stakes, D.S., Taylor, H.P., 1992. The northern Samail ophiolite: an oxygen isotope, microprobe and field study. Journal of Geophysical Research 97, 7043e7080. Stern, R.J., Reagan, M., Ishizuka, O., Ohara, Y., Whattam, S., 2012. To understand subduction initiation, study forearc crust: To understand forearc crust, study ophiolites. Lithosphere 4, 469e483.
Stormer, J.C., 1975. A practical two-feldspar geothermometer. American Mineralogist 60, 667e674. Stormer, J.C., 1983. The effects of recalculation on estimates of temperature and oxygen fugacity from analyses of multi-component iron-titanium oxides. American Mineralogist 68, 586e594. Stracke, A., Bizimis, M., Salters, V.J.M., 2003. Recycling oceanic crust: quantitative constraints. Geology, Geophysics and Geosystems (G-Cubed) 4, 1e33. https:// doi.org/10.1029/2001GC000223. Styles, M.T., Ellison, R.A., Phillips, E.R., Arkley, S., Schofield, D.I., Thomas, R.J., Goodenough, K.M., Farrant, A.R., McKervey, J.A., Crowley, Q.G., Pharoah, T.C., 2006. The geology and geophysics of the United Arab Emirates, vol. 2. Geology, Ministry of Energy, United Arab Emirates, Abu Dhabi, p. 351. Sun, 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.), 42. Geological Society, London, Special Publications, pp. 313e345. Tanaka, T., Togashi, S., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, T., Yuhara, M., Orihashi, Y., Yoneda, S., Shimizu, H., Kunimaru, T., Takahashi, K., Yanagi, T., Nakano, T., Fujimaki, H., Shinjo, R., Asahara, Y., Tanimizu, M., Dragusanu, C., 2000. JNdi-1: a neodymium isotopic reference in consistency with La Jolla neodymium. Chemical Geology 168, 279e281. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Oxford, 312. Thomas, R.J., Ellison, R.A., 2014. Geological Evolution of the United Arab Emirates. Over Six Hundred Million Years of Earth History. British Geological Survey, Keyworth, Nottingham, p. 287. Thomas, R.J., Styles, M.T., Goodenough, K.M., 2014. The UAE-Oman Ophiolite - a cross section through the Earth’s oceanic lithosphere on land. In: Thomas, R.J., Ellison, R.A. (Eds.), Geological Evolution of the United Arab Emirates: Over Six Hundred Million Years of Earth History. British Geological Survey, Ministry of Energy United Arab Emirates, pp. 153e176. Thorkelson, D.J., Breitsprecher, K., 2005. Partial melting of slab window margins: genesis of adakitic and non-adakitic magmas. Lithos 79, 25e41. Tilton, G.R., Hopson, C.A., Wright, J.E., 1981. Uranium-lead isotopic ages of the Semail ophiolite Oman, with applications to Tethyan ocean ridge tectonics. Journal of Geophysical Research 86, 2763e2776. Tippit, P.R., Pessagno, E.A., Smewing, J.D., 1981. The biostratigraphy of sediments in the volcanic unit of the Samail Ophiolite. Journal of Geophysical Research 862756e862762. Valley, J.W., 2003. Oxygen isotopes in zircon. Reviews in Mineralogy and Geochemistry 53, 343e385. https://doi.org/10.2113/0530343. Warren, C.J., Parrish, R.R., Waters, D.J., Searle, M.P., 2005. Dating the geologic history of Oman’s Semail ophiolite: insights from U-Pb geochronology. Contributions to Mineralogy and Petrology 150, 403e422. White, W.M., Klein, E.M., 2014. Composition of the oceanic crust. Treatise on Geochemistry 2, 457e496. Whitney, D., Evans, B., 2010. Abbreviations for names of rock-forming minerals. American Mineralogist 95, 185e187. Winchester, J.A., Floyd, P.A., 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology 20, 325e343. Wu, T.W., Kerrich, R., 1986. Combined oxygen isotope compositional studies of some granitoids from the Grenville, Ontario, Canada. Implications for source regions. Canadian Journal of Earth Sciences 23, 1420e1432. Xu, W., Zhang, H., Luo, B., Guo, L., Yang, H., 2015. Adakite-like geochemical signature produced by amphibole-dominated fractionation of arc magmas: an example from the late cretaceous magmatism in Gangdese belt, south Tibet. Lithos 232, 197e210. Xu, Z., Zheng, Y.F., Zhao, Z.F., 2018. Zircon evidence for incorporation of terrigenous sediments into the magma source of continental basalts. Scientific Reports 8, 178. https://doi.org/10.1038/s41598-017-18549-7. Zhao, Z., Zheng, Y., 2003. Calculation of oxygen isotope fractionation in magmatic rocks. Chemical Geology 193, 59e80. Zheng, Y.F., 1993a. Calculation of oxygen isotope fractionation in anhydrous silicate minerals. Geochimica et Cosmochimica Acta 57, 1079e1091. Zheng, Y.F., 1993b. Calculation of oxygen-isotope fractionation in hydroxyl-bearing silicates. Earth and Planetary Science Letters 120, 247e263. Zi, J.W., Cawood, P.A., Fan, W.M., Wang, Y.J., Tohver, E., 2012. Contrasting rift and subduction-related plagiogranites in the Jinshajiang ophiolitic mélange, southwest China, and implications for the Paleo-Tethys. Tectonics 31, TC2012. https://doi.org/10.1029/2011TC002937.