fluid evolution of abyssal mantle peridotites from the Nagaland Ophiolite Complex, NE India: Geodynamic significance

Journal Pre-proof P-T-melt/fluid evolution of abyssal mantle peridotites from the Nagaland ophiolite complex, NE India: Geodynamic significance

Aliba AO, Santanu Kumar Bhowmik, Dewashish Upadhyay PII:

S0024-4937(19)30504-3

DOI:

https://doi.org/10.1016/j.lithos.2019.105344

Reference:

LITHOS 105344

To appear in:

LITHOS

Received date:

11 July 2019

Revised date:

12 December 2019

Accepted date:

15 December 2019

Please cite this article as: A. AO, S.K. Bhowmik and D. Upadhyay, P-T-melt/fluid evolution of abyssal mantle peridotites from the Nagaland ophiolite complex, NE India: Geodynamic significance, LITHOS(2019), https://doi.org/10.1016/j.lithos.2019.105344

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© 2019 Published by Elsevier.

Journal Pre-proof P-T-MELT/FLUID EVOLUTION OF ABYSSAL MANTLE PERIDOTITES FROM THE

NAGALAND

OPHIOLITE

COMPLEX,

NE

INDIA:

GEODYNAMIC

SIGNIFICANCE ALIBA AO1,2,* [email protected], SANTANU KUMAR BHOWMIK1 AND DEWASHISH UPADHYAY1 1

DEPARTMENT

OF

GEOLOGY

&

GEOPHYSICS,

INDIAN

INSTITUTE

OF

TECHNOLOGY,

KHARAGPUR, INDIA

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*

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WADIA INSTITUTE OF HIMALAYAN GEOLOGY, DEHRADUN, INDIA

Corresponding author.

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2

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ABSTRACT

Spinel lherzolite from ophiolitic mélange in the Nagaland Ophiolite Complex, Indo-Myanmar

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ranges is used to reconstruct the P-T and chemical evolution of slab mantle in a Neo-Tethyan

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subduction channel. Five sequential high-T stages are identified for lherzolite based on

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petrography, mineral chemistry (major and trace elements), P-T pseudosection modelling and thermobarometry. (1) Large aluminous orthopyroxene + aluminous clinopyroxene + high

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magnesian olivine + chromiferous primary spinel grains formed during dry melting of source abyssal spinel lherzolite at T > 1150-1335 o C and P = 12-16 kbar. (2) Clinopyroxene and spinel exsolution in orthopyroxene, clinopyroxene + spinel coronae around megacrystic orthopyroxene and olivine formed during sub-solidus cooling from ~1100 to ~800 o C. (3) Neoblastic low-Al ortho- and clinopyroxene at ~800 °C formed during high-T mylonitisation, as the subducted lherzolite body accreted to the overlying wedge mantle as a metamorphic sole. Tectonism in this stage switched from being divergent to convergent. (4) Light REEenriched,

high-Mg

clinopyroxene

veins

developed

locally

within

the

orthopyroxene

megacrysts reflect the transit of a hybridised-mantle wedge melt. (5) A transient phase of high-T (T > 800 °C) hydration of clinopyroxene and locally orthopyroxene stabilised

Journal Pre-proof pargasite. Stages (4) and (5) relate to slab-derived melt/fluid migration through the accreted mantle peridotite. These findings provide new insights into the geodynamic evolution of the north-eastern segment of the Neo-Tethyan lithosphere as it evolved from a mid-oceanic ridge to a supra-subduction zone setting.

KEYWORDS: Nagaland Ophiolite Complex, melt/fluid-rock interaction, P-T path, subduction dynamics,

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chemical fingerprinting

1. Introduction

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Studies of ophiolite complexes contribute greatly to our understanding of paleo-

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subduction dynamics and thus play a key role in our understanding of active-subduction (e.g. Parkinson & Pearce 1998). Ophiolites may form in a variety of tectonic environments ranging

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from mid-oceanic-ridge (MOR) to supra-subduction zone (SSZ) settings that can be

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distinguished by their geochemical and mineralogical characteristics (Miyashiro 1973; Pearce

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et al., 1984; Pearce 2014; Dilek & Furnes 2014). Studies reveal that abyssal peridotites collected from MORs are residues of adiabatic decompression melting after extractions of

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small melt fractions (~5-15 %) (e.g. Shervais and Jean 2012; Warren, 2016). Residual pyroxene in abyssal peridotite generally has a pattern of strongly depleted chondritenormalised light rare earth element (LREE) content and relatively flat middle- to heavy rare earth element (MREE/HREE) content (Johnson et al., 1990; Hellebrand et al., 2002). In contrast, mantle wedge peridotite formed above subduction zones are compositionally modified by partial melts/fluids emanating from the descending slab and/or sedimentary rocks (Tatsumi and Eggins, 1995; Plank and Langmuir, 1998; Elliot, 2003; Tatsumi, 2005). Partial melts sourced from such hybridised mantle wedge are enriched in large ion lithophile elements (LILE) (e.g. Cs, Rb, K, Ba, Pb, Sr, La, Ce) and depleted in high-field strength elements (HFSE) (e.g. Ta, Zr, Nb, Ti) (Shervais and Jean 2012).

Journal Pre-proof Tracing the origin of mantle peridotite in an ophiolitic mélange can be complex, as abyssal peridotite (formed below MOR) can be tectonically juxtaposed with SSZ peridotite by plate convergence and thus involve multiple stages of evolution (e.g. Batanova and Sobolev, 2000; Dilek, 2003; Choi et al., 2008; Uysal et al., 2012). Several studies demonstrate such juxtaposed mantle peridotites in rocks exposed along the eastern margin of the Indian plate in the Indo-Myanmar ranges (IMR), namely the Nagaland Ophiolite Complex (NOC) (Venkataramana et al., 1986; Fareeduddin & Dilek, 2015; Dey et al., 2018;

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Abdullah et al., 2018; Ghosh et al., 2018), Manipur Ophiolite Belt (MOB) (Pal et al., 2014; Singh et al., 2017b; Kingson et al., 2017), the Kalaymyo ophiolite (Liu et al., 2016) and the

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Andaman ophiolite (Pal, 2011; Ghosh et al., 2018). Recent studies of mantle peridotites from

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these belts, however, reveal contrasting models for their origin, particularly with relation to

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the NOC and the MOB. For example, Ghosh et al. (2018) proposed a MOR-origin for the NOC, whereas a SSZ origin was proposed by Abdullah et al. (2018). Similarly, the MOB is

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considered to be an abyssal peridotite (MOR setting) subsequently accreted in a forearc

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setting (Singh, 2009, 2013; Ningthoujam et al., 2012; Singh et al., 2017b), or possibly formed

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in a SSZ setting (Pal et al., 2014; Kingson et al., 2017). However, despite signatures of both MOR and SSZ-origin seemingly present in peridotites of the NOC and adjoining MOB, the geodynamic pathways of their transition from oceanic slab mantle to forming part of an overplate mantle wedge are poorly understood. This ambiguity relates to: (1) the structural level of the samples used in previous studies not being identified within the overall architecture of the Nagaland accretionary complex, making it difficult to relate the peridotites to either the descending slab or the mantle wedge material; (2) most previous studies being based on bulk geochemical and isotopic data of samples that thus lacked texturally-controlled trace element analyses of minerals to fingerprint the chemical signatures of protolith and

Journal Pre-proof fluid/melt; (3) a lack of detailed metamorphic investigation of the peridotites to link them to the stages of subduction burial and accretion to the overplate mantle wedge. In view of contradictory interpretations for the origin of the NOC, we undertake a petrological investigation of a suite of samples taken from the base of a spinel peridotite body, which occurs in the structurally lowermost sequence of the Nagaland Accretionary Complex. We have adopted an integrated approach combining textural and mineral chemical

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(major and trace element) studies with geothermobarometric (combined conventional major

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and REE thermometry and isopleth thermobarometry through P-T pseudosection modeling)

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computations. Finally, we combine the current findings and published results to reconstruct the geodynamic and chemical pathways of the eastern segment of the Neo-Tethys from its

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origin in a MOR setting through subduction burial to its accretion at the slab-wedge interface.

2. Geological setting and previous work

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The IMR lie at the interface between easterly subducting Indian oceanic lithosphere in the

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west and Jurassic to Cretaceous rocks of the Burmese plate forming a magmatic arc-fore arc

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in the east (Mitchell et al., 2012) (Fig. 1). The NOC is the largest of a series of dismembered and tectonised suite of ophiolitic rocks that occur all along the IMR. It is part of a ~3000 km long, arcuate ophiolite belt extending from the eastern Himalayan syntaxis in the north through the IMR in the middle to the Andaman-Nicobar Islands in the south (Fig. 1). The NNE-trending NOC is 200 km long and 5-15 km wide, and comprises an ophiolite mélange that includes high-pressure (HP) metamorphic rocks (Chatterjee & Ghose, 2010; Ao & Bhowmik, 2014; Bhowmik & Ao, 2016; Rajkakati et al., 2019). Previous studies of the NOC have demonstrated a tectonised ophiolite stratigraphy with pillow basalts and pelagic ocean sediments at the top, followed progressively downward by sheeted dykes and plagiogranite, cumulate gabbro and a suite of variously serpentinised

Journal Pre-proof ultramafic association of dunite, harzburgite, pyroxenite and lherzolite (Brunnschweiler, 1966; Ghose & Singh, 1980; Acharyya, 1986) (Fig. 2a). Recently Bhowmik & Ao, (2016) subdivided the NOC into two belts: (1) a western belt and (2) an eastern belt. The western belt is a collage of metasedimentary rock, metabasalt, metagabbro and ultramafite. Two varieties of HP metamorphic suite of rocks were identified from the western belt. The dominant variety is a sequence of pumpellyite-diopside, lawsonite- and epidote blueschist facies metabasalt with peak P-T conditions between ~6 kbar, ~335 °C and ~11.5 kbar, ~340

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°C (Ao & Bhowmik, 2014). Based on the evidence for a hairpin clockwise (CW) P-T loop in blueschists and a steep prograde P-T path of the pumpellyite-diopside facies rocks (Fig. 2a), a

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cold and mature stage of an intra-oceanic subduction within the Neo-Tethys has been

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recognised. The second variety are amphibolites, locally garnetiferous and pyroxene-bearing,

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which record two overprinting metamorphic cycles (M1 & M2 ) and looping counter clockwise (CCW) metamorphic P-T paths as part of a single tectonothermal event (Bhowmik & Ao,

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2016) (Fig. 2a). The M1 cycle records peak metamorphism in hornblende-eclogite facies at

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13.8 ± 2.6 kbar, 625 ± 45 °C (error 2σ values) and subsequent cooling and partial exhumation

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to greenschist facies. Whereas the M1 amphibolite metamorphism can be explained by the formation of a metamorphic sole at the base of the overlying mantle wedge during subduction infancy of the Neo-Tethys, the epidote blueschist facies re-metamorphism at 14.4 ± 2 kbar, 540 ± 35 °C (cf. peak M2 metamorphism) and exhumation to greenschist facies condition reflects a more mature stage of subduction in the Neo-Tethys (Bhowmik & Ao, 2016). The mantle peridotite rocks investigated in this manuscript form a lensoidal body at the base of the western belt, south of Satuza (Fig. 2a,b). The eastern belt, which is separated from the western belt by an east-dipping thrust contact,

comprise

volcanic

rocks,

high-level quartz-dolerite to

microgabbro

plutons,

mafic/ultramafic cumulate and serpentinite (Agrawal and Ghose, 1986; Venkataramana et al.,

Journal Pre-proof 1986; Ghose and Agrawal, 1989). Abdullah et al. (2018) recently studied the major and trace element geochemistry of a mafic-ultramafic suite of gabbro, clinopyroxenite, wehrlite and dunite that are considered to represent the Mantle Transition Zone of the NOC. The authors interpret their evolution via melt-rock interaction and fractional crystallisation processes in a SSZ setting. Dey et al. (2018) documented a variety of peridotitic rocks ranging from olivine websterite through websterite to olivine clinopyroxenite, and harzburgite in the NOC. They

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classified them into: (a) refractory; and (b) enriched mantle types. These represent N-MORB and E-MORB source reservoirs respectively. On the other hand, Ghosh et al. (2018) have

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proposed that mantle peridotites from the Nagaland and Manipur ophiolite belts bear

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unequivocal evidence of abyssal peridotite that experienced anhydrous melting of a MORB

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mantle source. From the Manipur Ophiolite (MO), which lies at the southern continuation of the NOC, Singh et al. (2017b) posit that mantle tectonites of spinel lherzolite and

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clinopyroxene-bearing harzburgite equilibrated at temperatures of 970-1155 o C. A two-stage

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evolution for the mantle tectonites of the MO belt was suggested: (1) an early stage marked

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by low degrees of partial melting (< 15%), followed by extraction of basaltic melts in a midoceanic ridge (MOR) tectonic setting, in proximity to the eastern margin of the Indian lithosphere; (2) this residual peridotite was later subducted and subsequently trapped in the forearc region of the subduction zone and then exhumed during the collision of the Indian plate with the Burmese micro-plate (Singh, 2009, 2013; Ningthoujam et al., 2012; Singh et al., 2017b). However, on the basis of the chemical composition of chromitite pods and lenses hosted within the mantle peridotites, Pal et al. (2014) suggested a SSZ setting for the origin of the MO. Constraints from Nd isotopic ratios and trace element abundances in serpentinised peridotites also suggests that the MO represents a buoyant forearc-mantle-wedge system (Kingson et al., 2017). Based on U/Pb zircon ages of plagiogranite and gabbro, the Nagaland-

Journal Pre-proof Manipur ophiolite belt is inferred to have formed during the Early Cretaceous (Singh et al., 2017a; Aitchison et al., 2019). Further south along the IMR, the Kalaymyo ophiolite is characterised by the presence of large peridotite massifs (Liu et al., 2016). These peridotites comprise spinel lherzolite and harzburgites with minor plagioclase peridotite. Geochemical studies indicate both refractory and fertile mantle compositions. Whereas the refractory peridotites have compositions similar to the forearc peridotites, the fertile peridotites are compositionally similar to abyssal

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peridotite (Liu et al., 2016).

In the southernmost end of the IMR, the mantle section in the Andaman Ophiolite (AO)

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consists of serpentinised lherzolite and harzburgite with pods and lenses of dunite and

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chromitite (Pal, 2011). Pal (2011) proposed a combined MOR-SSZ origin for the Andaman

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mantle peridotites, but Ghosh et al. (2018) posited an abyssal peridotite origin for the mantle

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3. Key textures

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peridotites that possibly include a hydrous melting event.

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We have studied several mantle peridotite samples in the NOC ranging in composition from dunite, clinopyroxene-bearing dunite, orthopyroxene-bearing dunite, and lherzolite to harzburgite (see Fig. 2 and supplementary table S1 for their GPS locations and composition). Two mantle peridotite samples of spinel lherzolite composition (117A and 44A) (Fig. 3) selected for detailed petrological study were collected from the basal peridotite body in the western belt of the NOC, south of Satuza (Fig. 2a,b). Both of these samples preserve primary and

later recrystallisation textures, which enable the reconstruction of the complex

mineralogical evolution of the rocks. Key textural features are presented in Figs. 4-5. Both the samples have igneous relics of megacrystic ortho- and clinopyroxene, olivine and spinel set in a serpentine matrix (Fig. 4). Sample 117A has thin, discontinuous but repetitive

Journal Pre-proof layering of clinopyroxene (Fig. 4a). The key textural features observed in the rock are described in detail below.

3.1. Exsolution and corona textures In both samples (117A and 44A), orthopyroxene commonly has thin exsolution-lamellae of clinopyroxene (Fig.

5a-c).

Due to

the effects of subsequent deformation and

recrystallisation, the lamellae are discontinuous and heterogeneous in size and distribution.

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Some orthopyroxene lamellae were recrystallised to form granular discrete grains (Fig. 5c). Rare orthopyroxene lamellae within clinopyroxene are also observed. Pyroxene exsolution in

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spinel lherzolite is generally attributed to cooling at sub-solidus mantle conditions (e.g.

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Hellebrand et al., 2002). Clinopyroxene additionally occurs as part of an intergrowth with

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green spinel, generally as partial coronae to olivine against orthopyroxene (Fig. 5d). This second variety of spinel, which is distinguished from coarse matrix spinel, also occurs as

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small curvilinear blebs within orthopyroxene (Fig. 5c), and is considered to be of

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metamorphic origin. We attribute this corona texture to the combined decomposition of

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primary aluminous and calcic orthopyroxene and olivine as part of the same sub-solidus cooling stage that produced the pyroxene exsolution textures.

3.2. High-temperature deformation fabric Sample 117A records a strong tectonic fabric defined by discontinuous layers of clinopyroxene

alternating

with

layers

rich

in

orthopyroxene

and

olivine

(mostly

serpentinised) (Fig. 4a). Clinopyroxene is either elongated (Fig. 4a) or composed of recystallised mosaics of smaller clinopyroxene grains (Fig. 5e). Orthopyroxene megacrysts have characteristic core-rim structures (Figs. 4a, 5f). Cores of megacrystic orthopyroxene grain (cf. orthopyroxene porphyroclast) have undulatory extinction and bent cleavage traces,

Journal Pre-proof while the rim region shows aggregates of neoblasts of smaller orthopyroxene in association with smaller crystals of olivine and clinopyroxene (Fig. 5g-i). Such textural features of hightemperature deformation (e.g. Linckens et al., 2011 and references therein) are consistent with sample 117A being a mantle tectonite. In contrast, sample 44A is weakly deformed, lacking tectonic layering and porphyroclastic textures (Fig. 4b). The variation in strain is attributed to heterogeneous strain in the spinel lherzolite.

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3.3. Textures for vein formation and high-temperature hydration

Clinopyroxene occurs as megacrysts, exsolution lamellae in orthopyroxene and neoblasts

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around orthopyroxene porphyroclast, and as 50-250 μm thick veins in orthopyroxene in

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sample 117A (Fig. 5j). The high-temperature hydration of the anhydrous mantle peridotite is

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demonstrated by the presence of hornblende. Hornblende occurs as: (a) coronae around vein clinopyroxene (Fig. 5j) and clinopyroxene/orthopyroxene neoblasts in re-crystallised domains

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(Fig. 5i), (b) thin, discontinuous blebs along cleavage traces of clinopyroxene megacrysts

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(Fig. 5k) and (c) pseudomorphs of clinopyroxene lamellae within orthopyroxene.

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Summarising, the Nagaland spinel lherzolite rocks record five stages of equilibration. These are: (a) an early stability of olivine + megacrystic ortho- and clinopyroxenes + primary spinel (cf. stage 1); (b) clinopyroxene and spinel exsolutions in megacrystic orthopyroxene and clinopyroxene + spinel coronae around orthopyroxene and olivine (stage 2);

(c) the

development of neoblasts of ortho- and clinopyroxenes as a result of mylonite deformation (stage 3); (d) the formation of localised clinopyroxene veins within orthopyroxene megacryst (stage 4); and (e) late amphibole partially pseudomorphing clinopyroxene and orthopyroxene (stage 5).

Journal Pre-proof 4. Methods of study 4.1. Analytical techniques 4.1.1. EPMA The major element composition of minerals in the mantle peridotites of the NOC were measured on a CAMECA SX-100 electron microprobe at the EPMA National Facility, hosted at the Department of Geology and Geophysics, Indian Institute of Technology, Kharagpur. The operating conditions were: 1 µm beam diameter, 15 kV accelerating voltage and a 20 nA

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beam current. Natural and synthetic mineral standards used for calibration included orthoclase (Si and K), TiO 2 (Ti), corundum (Al), chromite (Cr), hematite (Fe), rhodonite

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(Mn), diopside (Mg and Ca) and jadeite (Na). ZAF matrix correction were applied through

using

the

AX

program

(T.J.

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recalculated

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the Cameca-supplied PAPSIL program. Mineral analyses (other than that of amphibole) were

http://ftp.esc.cam.ac.uk/research/research-groups/holland/ax).

Holland,

Amphibole

unpublished; analyses

were

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4.1.2. XRF

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normalised following the procedures of Leake et al. (1997).

Bulk-rock major element analyses of sample 44A were carried out at the National Centre for Earth Science Studies (NCESS) Thiruvananthapuram, India, using a Bruker model S4 Pioneer sequential wavelength-dispersive X-ray spectrometer, equipped with a goniometer (which holds seven analysing crystals: OVO-55, PET, LiF 200, LiF 220, Ge, ADP and InSb), 4 kW Rh X-ray tube, 0.23 and 0.46° collimators and SPECTRA plus software for determination of element abundances. Analytical precision was better than 1%. Full details of the precision and accuracy of calibration curves and data reliability are available in the CESS website

(https://www.ncess.gov.in/research-groups/crustal-processes-crp-

group/laboratories/xrf- lab/analytical- method-and-reference-standards.html).

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4.1.3. LA-ICPMS Trace element composition in pyroxene and amphibole were determined using Laser Ablation Inductively Coupled Plasma Mass spectrometry (LA-ICP-MS) at the Department of Geology and Geophysics, IIT Kharagpur. A Cetac 213 nm Nd YAG laser-ablation system was connected to a Varian 820 quadrupole ICP-MS. Samples were ablated at 5 Hz repetition

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rate, ca. 5 J/cm2 fluence and 40-60 μm spot sizes. Analyses were performed in time-resolved

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peak-hopping mode with each analysis consisting of 20 s background measurement with the

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laser turned off, and 40 s peak signal measurement with the laser ablating the sample. The ICP-MS was optimized for maximum sensitivity on Be, Co, In, Ce, and Th by ablating the

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NIST 612 reference glass. The oxide production rate monitored on the

232

Th16 O peak was

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found to be < 2%. External standardisation used standard-sample bracketing technique with ten measurements of the samples bracketed by two measurements of the NIST 612 reference

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glass. The NIST 610 reference glass was measured as an unknown interspersed with sample

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analyses. Data were reduced using the Glitter® software (Griffin et al., 2008) using Si and Ca

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as internal standards. Precision and accuracy were monitored from repeat measurements of the NIST 612 and NIST 610 reference glasses during the analytical session. Accuracies for the NIST 612 glass are: Li, Be, Sc, V, Co, Ni, Cu, Zn, Rb, Sr, Y, Cs, Mo, Ag, Sb, La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, Lu, Hf, Ta, W, Au, Pb, Bi, Th and U: < 1%; As, Zr, Nb and Ba: < 2%; and Ti and Cr: < 3%. Similarly, accuracies for the NIST 610 glass are Be, V, Zn, Y, Zr, Cs, Mo, Ba, Nd, Sm, Gd, Ho, Lu, Hf, W, Pb, Bi, Th and U: < 1%; Sc, Mn, Co, As, Sr, Nb, Ce, Pr, Eu, Tb, Er, Tm and Ta: < 2%; Ni, Rb, Ag, Sb, La, Yb and Au: < 3%; Li, Cr and Ti: < 5%; and Cu: < 7% (2σ). Reproducibilities (2σ) across several analytical sessions were better than 17% for most elements except for Sc (25%), V (25%), Pb (23%), and Bi (27%).

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4.2. Computation of metamorphic pressure (P) and temperature (T) conditions and P-T histories Conventional thermometric and P-T pseudosection modelling approaches were used to calculate P-T conditions for the five stages of equilibration in the spinel lherzolite. We utilize micro-domainal equilibration of mineral assemblages and compositions to obtain the P-T constraints of these stages. For the conventional thermometric approach, we have used a

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variety of two pyroxene and single pyroxene thermometers that are briefly described below.

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4.2.1.1. Trace-element thermometry

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4.2.1. Conventional geothermometry

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The REE-in-two-pyroxene thermometer (TREE) of Liang et al. (2013) is sensitive to both temperature- and pyroxene REE composition. It is built on two internally consistent lattice

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strain models for REE partitioning between pyroxene and basaltic melt (Sun and Liang, 2012;

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Yao et al., 2012). In a plot of ln (D)-A vs. B [where A is the function of major element

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composition of pyroxene and B is the function of ionic radius of a given element (Liang et al., 2013)], a best fit line was drawn for all the measured data and passing through the origin (Supplementary Fig. S1). The slope of the line gives the equilibrium temperature for the REE two-pyroxene thermometer (Liang et al., 2013). A meaningful inversion temperature is obtained by ignoring the scattered data points that fall off this linear trend. Temperature estimates were inverted following the methodology outlined in Appendix A of Liang et al. (2013). Due to the higher closure temperature of REE partitioning between the two pyroxenes (Cherniak and Dimanov, 2010; Liang et al., 2013), TREE is likely to shed critical insights into the

earlier,

higher temperature equilibration stage

metamorphosed spinel lherzolite in this study.

in the polyphase deformed

and

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4.2.1.2. Major element thermometry We have applied two-pyroxene thermometry following the calibrations of Brey & Kohler (1990) (temperature estimate TBKN from now on) and Taylor (1998) (TT from now on). The thermometry is based on the partitioning of enstatite component between coexisting orthopyroxene and clinopyroxene around the temperature-sensitive pyroxene miscibility gap. The TBKN calibration, which was expanded to incorporate other elements within the

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pyroxenes such as Fe, Mg, Ca, and Na is said to reproduce the experimental values to ± 15 °C (Brey and Köhler, 1990). The TT is a recalibration of the TBKN thermometer, for which a

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larger database and an improved activity model that incorporated corrections for minor

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components, specifically Ti, Fe and Na have been used. In general, these thermometers are

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believed to yield reliable mantle temperature estimates owing to a small pressure dependence and the relatively small effect of minor components, specifically ferric iron, on the pyroxene

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mutual solubility in ultramafic systems.

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We have also used Ca-in-orthopyroxene (Brey & Kohler, 1990, *TBKN from now on), and

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Al, Cr-in-orthopyroxene (Witt-Eickschen & Seck, 1991, TWS from now on) thermometers. These thermometers are calibrated for spinel peridotites and reproduce the T-estimates from experiments in natural systems to within ± 20 °C and ± 15 °C respectively (Brey & Kohler, 1990;

Witt-Eickschen

&

Seck,

1991).

Temperatures estimates derived

from both

thermometers may diverge due to superposed thermal events involving high cooling or heating rates (e.g. Witt-Eickschen & Seck, 1991). In general, for fully equilibrated samples, two-pyroxene and single-pyroxene thermometers should yield consistent estimates (e.g. WittEickschen and Seck, 1991). Additionally, the P-T conditions accompanying the late pargasitic amphibole growth have been estimated following an empirical equation developed

Journal Pre-proof by Mandler & Grove (2016). The results of thermometry calculated at a reference pressure of 15 kbar are presented in Table 3.

4.2.2. Phase diagram calculations We have reconstructed metamorphic phase relationships in dry supra- and sub-solidus conditions for a representative Nagaland spinel lherzolite sample (e.g. 44A) using the P-T pseudosection approach. Sample 44A is amenable for this purpose for the following reasons:

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(a) the sample is weakly deformed and lacks late veins; (b) despite partial serpentinisation, the original lherzolitic mineralogy is mostly preserved in the sample. We suggest that

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excepting later hydration, the original rock composition was not significantly modified, and

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hence is suitable for phase diagram computations.

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The calculations were performed with the THERMOCALC 3.4 program (Powell et al., 1998). Activity-composition models for minerals and melt phase (file tc-ds63.txt) follow

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Jennings and Holland (2015). The calculation in the chemical system NCFMASOCr, which is

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analogous to peridotite rock compositions is likely to simulate phase relationships and

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melting in mantle peridotite (e.g. Jennings & Holland, 2015). Ferric iron was set to 0.3 wt% Fe2 O 3 , based on the peridotite xenolith data presented by Canil et al. (1994), which reported the range from 0.1-0.4 wt%. The computed bulk rock compositions (oxide proportions in mole %) for sample 44A is listed below: SiO 2 = 37.86, Al2 O 3 = 1.31, CaO = 1.32, MgO = 51.46, FeO = 7.70, Na2 O = 0.11, O = 0.11, Cr2 O3 = 0.13. Results of this calculation are compared with temperature estimates from major and trace-element thermometry and texturally-supported mineral assemblage evolution. The P-T condition of formation of pargasitic amphibole in the investigated sample has been estimated with a calculated P-T pseudosection in the model NCFMASH system, and using the software Perplex_X 6.7.8 (Connolly, 1990, updated March 30, 2017) and the

Journal Pre-proof internally consistent thermodynamic data set of Holland and Powell (1998, updated 2002). Solution

models

used

are: orthopyroxene,

Opx (HP) (Powell & Holland

1999);

clinopyroxene, Cpx (HP) (Powell & Holland 1999); olivine, O (HP) (Holland & Powell 1998); spinel, Sp (stx) (Stixrude & Lithgow-Bertelloni 2011); amphibole, Amph (DHP) (Dale et al., 2000); chlorite, Chl (HP) (Holland et al., 1998); garnet, Gt (HP) (Holland & Powell 1998); plagioclase, Pl (h) (Newton et al., 1980). The calculated phase diagrams are presented

oo

f

in Figs. 10-11.

5. Results

pr

5.1. Major element chemistry

e-

Representative mineral compositions of orthopyroxene, clinopyroxene, olivine, spinel and

Pr

amphibole in the two samples are presented in Tables 1 & 2. The complete data set is given

rn

5.1.1. Orthopyroxene

al

in Table S2.

Jo u

Regardless of the textural sites, orthopyroxene in the two samples is Mg-rich with XMg (= Mg/Mg+Fe2+) = 0.90-0.92 (Tables 1 & S2a). Although there is some compositional overlap, megacrystic orthopyroxene crystals, particularly in their core regions, are more aluminous (Al2 O3 = 4.3-6.0 wt%) and chromiferous (Cr2 O3 = 0.4-0.7 wt%) than the neoblastic types (Al2 O3 = 2.5-3.5 wt%, Cr2 O3 = 0.2-0.3 wt%) (Fig. 6a). The relatively lower Al and Cr contents in the rims of orthopyroxene megacrysts commonly overlap with compositions of neoblastic orthopyroxene (Fig. 6a) reflecting compositional re-equilibrations during HT deformation.

Journal Pre-proof 5.1.2. Clinopyroxene Clinopyroxene is more magnesian than orthopyroxene with XMg = 0.96-1.00 in sample 117A and 0.94-1.00 in 44A (Tables 2 & S2b), consistent with Mg/Fe partitioning data for these minerals in natural rocks. As in orthopyroxene, megacrystic clinopyroxene is more aluminous (Al2 O3 = 4.0-6.5 wt%) and titaniferous (TiO 2 = 0.2-0.6 wt%) relative to neoblastic (Al2 O3 = 2.8-3.7 wt%, TiO 2 = 0.15-0.28 wt%) and vein-type (Al2 O3 = 2.5-4.0 wt%, TiO 2 = 0.10-0.32 wt%) clinopyroxene (Fig. 6b). In a binary plot of TiO2 vs. Al2 O 3 , megacrystic

oo

f

clinopyroxene compositions fall broadly in the abyssal peridotite field, whereas neoblastic and vein-type clinopyroxene plot between abyssal peridotite and forearc peridotite fields (Fig.

pr

6b). Although clinopyroxene lamellae have the same range of Al-content, they are more

Pr

e-

titaniferous than their megacrystic type (Fig. 6b).

5.1.3. Olivine

al

Irrespective of textural types, XMg (0.90-0.91) and MnO (0.09-0.20 wt%) contents of

rn

olivines in both samples (Table 2 & S2c) are consistent with those of abyssal peridotites (e.g.

5.1.4. Spinel

Jo u

Aldanmaz et al., 2009; Khedr et al., 2014).

Coarse-grained matrix spinel in both of the samples are aluminous (Cr# = 0.10-0.12, where Cr# = Cr3+/Cr3++Al) and magnesian in composition (XMg = 0.76-0.77) (Fig. 6c). Metamorphic spinel, occurring as intergrowths with clinopyroxene around olivine or as small blebs in orthopyroxene, is slightly more chromiferous (Cr# = 0.11-0.15), but has similar magnesian composition (XMg = 0.75-0.77) (Table 1 & S2e). This aluminous spinel composition is generally taken as evidence for a relatively fertile mantle (Dick & Bullen, 1984). Both textural varieties of spinel in the Nagaland mantle peridotites fall within or very

Journal Pre-proof close to the lower limit of the abyssal peridotite field and indicate very low degrees (< 10 %) of partial melting (Fig. 6c). Our data also have close similarities with the spinel compositions of mantle peridotites of the Nagaland Ophiolite Complex and the Andaman ophiolite (e.g. Ghosh et al., 2018; Fig. 6c). Spinel compositions from the Manipur ophiolite belt are, however, more chromiferous (e.g. Singh et al., 2017b; Ghosh et al., 2018; Fig. 6c).

f

5.1.5. Amphibole

oo

Amphibole in both samples is pargasitic (nomenclature after Leake et al., 1997) (Fig. 6d).

pr

Grains are uniformly magnesian (XMg = 0.89-0.86), titaniferous (TiO 2 = 1.05-1.93 wt%) and

e-

chromiferous (Cr2 O3 = 0.61-1.45 wt%) in composition (Table 1 & S2d).

Pr

5.2. Trace element chemistry

The trace element compositions of the pyroxenes, hornblende, olivine and spinel are

Jo u

5.2.1. Orthopyroxene

rn

al

presented in Supplementary tables S3a & S3b.

Orthopyroxene megacrysts show two types of REE patterns: the first group is depleted in LREE with chondrite-normalised abundances gradually increasing from light to heavy REE (Fig. 7a-b), consistent with the LREE depletion trends in the clinopyroxene megacrysts. These megacrysts are also relatively depleted in the LILE and HFSE (Fig. 7c-d). The other group, measured from rim regions of the megacrysts in sample 117A, is relatively enriched in the LREE, LILE, and several HFSE (Fig. 7a,c). Both groups have similar HREE concentrations (Fig. 7a,c).

Journal Pre-proof 5.2.2. Clinopyroxene Clinopyroxene forming megacrysts, granular exsolution lamellae and coronae in both samples has LREE-depleted REE patterns, with relatively flat MREE-HREE trends (Fig. 8ab). In multielement spidergram plots, these clinopyroxene types have marked depletion in HFSE (Th, U, Zr and Hf) and LILE (Rb, Ba and Sr), characteristic of a typical melting residue (Fig. 8c,d). In contrast, vein-type clinopyroxene in 117A is LREE-enriched (up to 30x) (Fig. 8a), which cannot be explained solely by an origin of simple partial melting

oo

f

residue. They also show distinct lower middle and heavy REE concentrations (Fig. 8a). The multielement spidergram plot of clinopyroxene veins also shows strong enrichments in HFSE

pr

(Th, U, Zr and Hf) and LILE (Rb, Ba and Sr) but strong depletion in Ti (Fig. 8c). Apart from

e-

the clinopyroxene veins, one clinopyroxene megacrystic rim composition also has elevated

Pr

trace element concentration similar to the vein-type clinopyroxene (Fig. 8a,c).

al

5.2.3. Olivine

rn

Most of the olivine compositions show a general LREE-depleted trend. Ni concentration

Jo u

in olivine ranges from 3095 to 3642 ppm in sample 44A while the values are slightly lower in sample 117A, being between 1678 and 2253 ppm (Tables S3a & S3b).

5.2.4. Hornblende

Hornblende has both LREE-depleted and LREE-enriched patterns (Fig. 9a), mimicking the REE patterns of the clinopyroxene that they pseudomorph. The LREE depleted patterns of hornblende resemble those of the neoblastic clinopyroxene, while the LREE-enriched patterns are similar to those of the vein-type clinopyroxene. In a multielement plot, the LREE-enriched group also shows enrichments in Ba-Th-U (Fig. 9b).

Journal Pre-proof 5.3. P-T conditions of equilibration of residual mantle 5.3.1. Conventional thermometry A combination of aluminous cores of megacrystic ortho- and clinopyroxenes has been chosen to estimate the temperatures of stage 1 equilibration. In sample 44A, the calculated temperatures for this stage vary from 1335 (TL) through 1115 (TWS)-985 (*TBKN) to 875 (TT )765 (TBKN) (Table 3). We interpret the systematic decrease in the calculated temperatures in

f

terms of differential closure temperatures of REE, Al-Cr, Ca and Mg-Fe in the pyroxenes

oo

(Liang et al., 2013). The highest temperature recorded by the REE thermometry is attributed

pr

to the sluggish diffusion and higher closure temperature of the REE in the two pyroxenes. The lower temperatures obtained by the different major element thermometers can be

e-

attributed to compositional resetting during the stage 2 exsolution process, which seems to be

Pr

extensive for Fe-Mg partitioning between the pyroxenes. We explain the discrepancy in the calculated temperatures in the two single-pyroxene thermometers, *TBKN and TWS (∆T ~130 C) by the different diffusivities of Ca and Al in orthopyroxene, with the diffusion of Ca

al

o

rn

being suggested to be faster than Al by a factor of 4 (e.g. Opper & Seck, 1989, quoted in

Jo u

Witt-Eickschen & Seck, 1991). Whereas Ca significantly re-equilibrated during the stage 2 phase of sub-solidus cooling, the slower diffusivity of Al enabled the retention of primary Al concentration in the cores of megacrystic orthopyroxene, resulting in higher temperature estimates. In the recrystallised mantle tectonite 117A, the calibrations of TL, *TBKN and TWS all yielded broadly consistent temperature estimates between 900 and 1000 o C (Table 3). While this implies an extensive equilibration of REE and Cr-Al and also of Ca in the pyroxenes, the average TL estimate in 117A is lower by ~335-450 o C than that in 44A. We relate this with the resetting of primary REE compositions, even in the cores of the megacrystic pyroxene grains due to the pervasive high-T recrystallization event in the sample. Summarising, T ~1335

o

C is considered as a reliable estimate of the stage 1

Journal Pre-proof equilibration temperature, also supported by the calculated P-T phase diagram of 44A (see the P-T pseudosection modelling section below). For the estimation of temperatures of the stage 2 equilibration, coexisting compositions of clinopyroxene lamellae/corona and orthopyroxene host/contact have been considered. In sample 44A, both the two-pyroxene trace element and single-pyroxene thermometers yielded near identical average temperature estimates of 990-1065 o C, which exceed major element two-pyroxene thermometric results by 280-355 o C (Table 3). As stated above, the divergence

oo

f

in the calculated temperatures is attributed to differences in the closure temperatures of trivalent and divalent cations in the pyroxenes. On the basis of this argument, the REE o

C) is interpreted as the thermal condition of pyroxene

pr

temperature estimate (T ~1050

e-

exsolutions. The major element two-pyroxene thermometric estimate (T ~700

o

C) is

Pr

considered as the lower temperature end of the sub-solidus cooling stage. The coexisting compositions of neoblast ortho- and clinopyroxene in 117A have been

al

chosen to calculate the thermal condition of the stage 3 equilibration. The two calibrations of

C.

Jo u

o

rn

the single-pyroxene thermometer yielded a tight temperature constraint between 780 and 850

The major element two-pyroxene thermometer yielded the lowest temperature estimate of ~600

o

C for both stage 2 and 3 events in 117A. The extensive down-temperature

compositional resetting may be attributed to the effects of widespread low-T serpentinisation in the sample.

5.3.2. P-T pseudosection modelling The computed phase diagram shows that the stability field of spinel lherzolite for sample 44A is demarcated by the pressure-sensitive plagioclase-out curve at low pressure (LP) (between 6 and 9 kbar) and the garnet-in curve at high pressure (HP) (at P > 12-21 kbar) and

Journal Pre-proof by the temperature-sensitive solidus curve at HT (between ~1175 and ~1400 °C) (Fig. 10). This is consistent with the results of experimental studies (e.g. Jenkins & Newton, 1979; O’Neill, 1981; Takahashi et al., 1993; Klemme & O’Neill, 2000; Borghini et al., 2009) and other natural rock data (e.g. Jennings & Holland 2015; Ziberna & Klemme 2016). The calculated pseudosection shows that most of the compositional isopleths of orthopyroxene [e.g. Ca and Al (IV) concentrations in the mineral] and spinel [e.g. Cr# (= Cr3+/(Cr3++Al) and XFe2+ (Fe2+/(Fe2++Mg)] in the spinel-lherzolite field

are strongly T-dependent. The

oo

f

pseudosection also predicts that the concentrations of Ca and Al (IV) in orthopyroxene progressively decrease with cooling, which is also in agreement with the results of

pr

experimental studies (e.g. Lindsley, 1983 for Ca content in orthopyroxene in equilibrium with

e-

clinopyroxene). These compositional parameters can thus be used to constrain the thermal

Pr

and baric conditions at different stages of equilibrations of the rock. Maximum recorded Ca content of 0.05 p.f.u. (per formula unit) in the core of megacrystic

o

C, in agreement with thermometric results (TWS estimate, Table 3).

rn

lherzolite at ~1150

al

orthopyroxene indicates a minimum temperature of sub-solidus equilibration of spinel

Jo u

Considering that orthopyroxene accommodated even more Ca before its exsolution of clinopyroxene, the spinel lherzolite is likely to have been hotter at the onset of cooling across the solidus, as indicated by the results of the trace element thermometry (Table 3). Metamorphic spinel, occurring as blebs within orthopyroxene megacrysts or as composite coronae with clinopyroxene around orthopyroxene and olivine, shows a compositional spread of Cr# (between 0.15 and 0.11) and XFe2+ (between 0.25 and 0.27) with the dominant cluster at Cr# = 0.14-0.15, XFe2+ = 0.25. This, coupled with relatively higher Ca concentration (~0.04 p.f.u.) in coarse co-existing orthopyroxene grains, appears to indicate that the high-T end of the sub-solidus cooling (cf. stage 2 equilibration) that initiated exsolution of spinel in orthopyroxene and formation of coronal spinel + clinopyroxene assemblage can be bracketed

Journal Pre-proof between 1050 and 1150 o C (Fig. 10), also in consonance with the results of conventional thermometry (Table 3). The minimum Ca content (~0.02 p.f.u.) in the rims of orthopyroxene megacrysts (Table 3), when evaluated with the calculated phase diagram, predicts the lower T end of the cooling segment at ~800 °C (Fig. 10).

5.4. P-T conditions of formation of pargasite In Fig. 11, we present the P-T pseudosection showing the calculated amphibole-out line

oo

f

and Na isopleths of pargasitic amphibole (this study) and also the amphibole-out line from several experimental works. Our calculations indicate very high-temperature stability of

pr

pargasite at ~1100 °C and 10 kbar (Fig. 11). The temperature of the computed hornblende-out

e-

curve is also higher by ~200 o C than the experimentally determined curves. This may be

Pr

attributed to bulk-compositional differences between the investigated spinel lherzolites and experimental compositions. Figure 11 is also contoured for isopleths of maximum molar

al

Na2 O + K2 O contents (in %) in pargasite from the investigated sample. The calculations have

rn

been carried out following an empirical equation developed by Mandler & Grove (2016). The

Jo u

calculated isopleths indicate a range of pressures between 14 kbar and 21 kbar (at a TRef of 800 and 1000 o C) for the formation of paragasite in the mantle peridotite.

6. Discussion

6.1. Thermal history and protolith composition of the Nagaland mantle peridotite The Nagaland residual mantle peridotite records five distinct stages of high-temperature mineral assemblage evolution, which in a sequence are: (1) stabilization of an assemblage of coarse-grained aluminous and chromiferous orthopyroxene (Al2 O3 = 4.3-6.0 wt%; Cr2 O3 = 0.7-0.4 wt%) + aluminous and titaniferous clinopyroxenes (Al2 O 3 = 6.5-4.0 wt%; TiO 2 = 0.60.2 wt%) (Fig. 6a, b) + coarse, high magnesian and nickeliferous olivine (XMg = 0.90-0.91;

Journal Pre-proof Ni = 3095-3642 ppm) (mostly serpentinised) (Tables S2c & S3a,b) + chromiferous (Cr# = 0.10-0.13) and magnesian primary spinel (XMg = 0.76-0.77), marking an early equilibrated spinel lherzolite assemblage (Fig. 13); (2) the development of clinopyroxene and spinel exsolution in orthopyroxene (Figs. 5a-c) and clinopyroxene + spinel coronae on megacrystic orthopyroxene and

olivine (Fig.

5d); (3) the formation of neoblastic

ortho-

and

clinopyroxenes at the expense of orthopyroxene megacrysts as a result of pervasive mylonite deformation and high-temperature recrystallization (Fig. 5f); (4) localised LREE- and LILE-

oo

f

enriched clinopyroxene vein formation in orthopyroxene megacrysts (Fig. 5j); and (5) hightemperature hydration of clinopyroxene and locally orthopyroxene, stabilising pargasite (Fig.

pr

5i-k).

e-

Combined textural, mineral compositional (major and trace element), conventional

Pr

thermometric and P-T pseudosection modelling studies indicate that at stage 1, the Nagaland spinel lherzolite rocks equilibrated at mantle depths (P = 12-16 kbar) and at super-solidus T > o

C (Fig. 13). Both REE and isopleth thermometric results indicate high T

al

1150-1335

rn

conditions (1050-1100 o C) for the renewed growth of metamorphic clinopyroxene and spinel

Jo u

(cf. stage 2 evolution), and marking the onset of sub-solidus cooling. The lower temperature end of this cooling segment is constrained at ~700 o C. Compositional features such as lower Al contents of both neoblastic ortho- and clinopyroxene as well as lower Ca content of the same textural type of orthopyroxene, when evaluated with the calculated phase diagram (Fig. 10), indicate equilibration of recystallised pyroxenes (cf. stage 3 evolution) at appreciably lower temperature

conditions.

In banded mylonite sample 117A, the estimation of

temperature by conventional thermometry yields lower temperature estimates of ~800

o

C

(Table 3). Given that the temperature-dependent sub-solidus exchange processes involving Al, Cr and Ca between orthopyroxene and clinopyroxene that operate in mantle peridotite (e.g. Brey and Köhler, 1990; Witt-Eickschen and Seck, 1991), the geological significance of

Journal Pre-proof the estimated temperatures through major element pyroxene thermometry is often debatable with suggestions varying from cooling temperatures (e.g. related to closure temperature of cation diffusion) to temperatures of deformation (e.g. Linckens et al., 2011). In the present context, the chosen fine-grained, neoblast orthopyroxene grains with distinct chemical compositions reflect dynamic recrystallisation of the orthopyroxene megacrysts. We are, therefore, inclined to interpret T ~800 o C from neoblast pyroxenes as the thermal condition for the stage 3 mylonite deformation event.

oo

f

While there is no direct measurement of the temperature condition of LREE- and LILEenriched clinopyroxene vein formation event (cf. stage 4 evolution), the host orthopyroxene

pr

rim and adjacent clinopyroxene vein compositions yield temperature estimates between 850

e-

and 900 o C, and may mark the possible thermal condition of vein formation (Table 3). As will

Pr

be established later, the formation of clinopyroxene vein is likely to be part of a melt infiltration event. This may imply that the infiltrating clinopyroxene-rich melt was even

al

hotter (T ≥ 900 o C). The appearance of extremely magnesian (XMg = 0.89-0.96), aluminous

rn

and titaniferous pargasitic amphibole that replaced both vein and neoblastic clinopyroxene

Jo u

implies an even later hydrous fluid infiltration event (cf. stage 5 evolution). In a recent study, Mandler & Grove (2016) have shown that the molar alkali content of amphibole in fertile and depleted mantle is a linear function of pressure and temperature variations but is insensitive to bulk compositional differences. The empirical calibration of Mandler & Grove (2016) suggests that a wide range of metamorphic pressure between 12.5 and 21 kbar at temperature between 800 and 1000 o C accompanied the stability of pargasite in the Nagaland peridotites. In contrast, the Na content of pargasite in the calculated P-T pseudosection implies an even warmer condition (T ~1050 o C) for the fluid infiltration event. Given the limitations of the solid solution model of hornblende used in P-T pseudosection calculation (as the model does not consider Cr and Ti), it is not possible at this stage to provide tighter constraints of the

Journal Pre-proof temperature of the hydrous fluid infiltration event. Nevertheless, these computations clearly indicate a transient phase of relatively HT melt/hydrous fluid infiltration at mantle depths that affected the partially cooled spinel lherzolite rock. The general LREE- and LILE-depleted compositions of the pyroxene megacrysts (Figs. 7a-d & 8a-d) suggest a residual origin of the Nagaland mantle rocks after melt extraction. Compositions of primary spinel (Cr# = 0.10-0.12; XMg = 0.75-0.77, Fig. 6c) and clinopyroxene (Fig. 12a-b; Fig. 6b), when evaluated with model tectonic discriminant

oo

f

diagrams, reflect a lineage of these rocks with abyssal peridotite and their origin through low degrees (< 6-7 %) of dry melting of a spinel lherzolite mantle source (Dick and Bullen, 1984;

pr

Arai, 1994; Bizimis et al., 2000; Hellebrand et al., 2001). Similar low values (1-5 %) of melt

e-

extraction were also reported by Ghosh et al. (2018) from the Nagaland mantle peridotites

Pr

(Fig. 6c). The REE pattern of clinopyroxene observed in our study are also very similar to clinopyroxene composition of abyssal peridotites from other localities (Fig. 12c), which is

al

indicative of a very small degree of melt extraction from a mantle source similar to a MORB

rn

mantle (e.g. Aldanmaz et al., 2009; Liu et al., 2016). Our finding is also consistent with the

Jo u

results of previous geochemical studies of mantle peridotites in the Nagaland-Manipur ophiolite belt (Singh 2009, 2013; Singh et al., 2017b; Dey et al., 2018; Ghosh et al., 2018). In sample 117A, the overall pattern of the LREE-depleted clinopyroxenes are slightly more enriched compared to the general abyssal clinopyroxene composition trend (Fig. 12c), indicating a possible interaction with the percolating melt/fluid, and as such was not used in calculation of melt trend plot in Fig. 12a-b.

6.2. Origin of clinopyroxene veins and high-T fluid infiltration Although vein clinopyroxene and hornblende are localised in their occurrences in the investigated spinel lherzolite samples, their trace element compositions provide critical

Journal Pre-proof insights into the nature of the source reservoir of the infiltrating melt/fluid. Low-degree partial melts of mantle peridotites can produce such clinopyroxene veins rich in incompatible elements plus a modal and cryptic enrichment in the surrounding peridotites (Pilet et al., 2011 and references therein). Strong negative trends in Nb-Ta-Ti associated with low-HREE concentration of the clinopyroxene vein suggest a significant contribution from the melting of a

subducted

oceanic

basalt

that was transformed

into

a garnet +

rutile-bearing

amphibolite/eclogite (Martin et al., 2005 and references therein). This is because the

oo

f

fractionation of residual mineral phases, such as amphibole and rutile is known to deplete the melt in Nb, Ta and Ti respectively. High Sr and low Y concentrations with corresponding

pr

high Sr/Y ratio are also observed in the clinopyroxene veins, indicating their high-pressure

e-

slab-melt origin (Defant & Drummond, 1990; Drummond & Defant, 1990). Apart from high

Pr

Sr/Y ratios, the trace element composition of clinopyroxene veins also shows high La/Yb values, depletions in Ti and Nb relative to LILE and REE and enrichment in Zr relative to Sm

al

(Fig. 8c). These chemical features are characteristic of adakites (e.g. Leonid et al., 2008).

rn

These adakites are interpreted to have formed by melting of the mantle wedge whose

Jo u

compositions have been modified by reaction with felsic-slab melts (Martin et al., 2005). Conversely, the low M-HREE clinopyroxene pattern is also very similar to that observed in clinopyroxenes of upper mantle wedge peridotites (e.g. Ritter Island peridotites, Tollan et al., 2017). A basalt rather than a sedimentary source is also evident from the very-low Be concentration (~1 ppm, Table. S3a) in the clinopyroxene veins which may indicate very little or no input from the subducting sediments. Variable enrichments in alkali elements such as Rb and Ba in clinopyroxene veins also reveal some degree of hydrothermal alteration of the down-going

basaltic

crust

before

they

were

metamorphosed

to

high-pressure

amphibolite/eclogite facies condition (e.g. Harper, 1984; Rapp et al., 1999). However extremely magnesian composition of vein clinopyroxene (XMg = 0.96) may also form due to

Journal Pre-proof hybridization of adakitic melt as it percolates and interacts with the more magnesian peridotitic mantle wedge material (e.g. Rapp et al., 1999). Consequently, the basalt-melt and mantle wedge peridotite association support the view that these clinopyroxene veins were not just simple slab melt products but are likely to be the products of melting of the mantle wedge, previously modified by reactions with an infiltrating basaltic melt. A hybrid melt source for the origin of clinopyroxene veins is also implied by the binary TiO 2 vs. Al2 O3 diagram (Fig. 6b) that shows a broadly linear array of pyroxene compositions between the

oo

f

abyssal and forearc peridotite fields.

However, a single rim composition of a clinopyroxene megacryst in sample 117A shows

pr

elevated LREE, HFSE (Th, U, Zr, Hf) and LILE (Rb, Ba, Sr) composition similar to that of

e-

clinopyroxene veins but with enriched HREE concentration similar to that of abyssal

Pr

peridotite clinopyroxene (Fig. 8a,c). We suggest that it represents the effects of cryptic metasomatism of the infiltrating hybridised-melt on the host clinopyroxene megacryst.

al

During such cryptic metasomatism, the rim composition of the clinopyroxene megacryst is

rn

enriched in LREE, HFSE and LILE concentration without affecting its original HREE

Jo u

concentration. Similar enrichment patterns are also observed on rims of the corresponding orthopyroxene megacryst that hosts the clinopyroxene veins (Fig. 8c). This suggests that these hybridised melts infiltrated the abyssal peridotites and caused restricted modal metasomatism, forming clinopyroxene veins and cryptic metasomatism affecting the rim composition of the pyroxene megacrysts. Pargasitic amphiboles have been commonly observed in metasomatised mantle rocks and their formation is generally attributed to fluid infiltration and interaction with dry mantle peridotites (Ionov et al., 1997; Bizimis et al., 2000; Liu et al., 2010). Textural analyses of paragasitic amphiboles in both samples (117A and 44A) reveal that they replace predominantly clinopyroxene (of lamellae, vein and neoblast textural modes) and locally

Journal Pre-proof orthopyroxene as well. The two end-member compositional trends of hornblende from this study: LREE-enriched and LREE-depleted (Fig. 9a) patterns are inherently linked with the composition of the replaced clinopyroxene. This implies that the hydrous fluid composition is not uniformly enriched in LREEs. Amphiboles with such low LREE and LILE are rare in mantle rocks but have been reported within fertile subcontinental-type lherzolites (Pirnia et al., 2018 and references therein). Two major processes are invoked to explain the formation of depleted amphiboles in mantle peridotites: (1) LREE- and LILE-enriched amphiboles

oo

f

interacted with the already depleted pyroxenes leading to sub-solidus re-equilibration between the enriched amphiboles and depleted pyroxenes; (2) metasomatism by H2 O-rich

pr

fluids. The first possibility may be discounted for the Nagaland spinel lherzolite amphiboles

e-

because the replaced pyroxenes have geochemical characteristics of a fertile mantle residue

Pr

with very low degree of partial melting (< 6-7 %), and with no evidence for subsequent compositional re-equilibration. The second possibility suggests that the depleted amphiboles

al

are derived from low-density H2 O-rich metasomatic fluids that are originally depleted in

rn

incompatible elements (e.g. Pirnia et al., 2018). Given the depleted nature of the Nagaland

Jo u

spinel lherzolite amphiboles, and the presence of hybridised mantle wedge-melts as documented above, we propose the formation of these amphiboles by H2 O-rich metasomatic fluids that are originally depleted in incompatible elements. Textural evidence seem to suggest that this infiltrating hydrous fluid formed later and was more pervasive than the clinopyroxene-forming melt.

6.3. A geodynamic model We now reconstruct the evolution of the Nagaland mantle peridotite in the context of subduction channel dynamics within the north-eastern segment of the Neo-Tethys (with respect to the current co-ordinates). The occurrence of the studied abyssal peridotites within

Journal Pre-proof the basal part of a progressive metamorphic sequence from greenschist through pumpellyitediopside to lawsonite blueschist facies lithological association of basalt-chert-limestone provides additional evidence that the Nagaland spinel lherzolite unit constitutes the mantle section of the down-going Neo-Tethys (Fig. 2a,b). The successive records of a protracted HT cooling history of > 625 °C (from T > 1335 to T ~700 °C) in the peridotites from the suprasolidus to sub-solidus conditions at mantle depths (P = 10-16 kbar) (Fig. 13), followed by a HT dynamic recrystallisation event, can be explained by a progressive movement of the Neo-

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Tethys away from the spreading centre (Stages 1-2; Fig. 14a) (e.g. Linckens et al., 2011), finally entering the subduction zone (Stage 3; Fig. 14b).

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In a recent review, Agard et al. (2016) proposed a mechanism of formation of

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metamorphic soles of different thermal and baric conditions (e.g. HT and LT soles). They also

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related the origin of other HP metamorphic rocks (e.g. HT and LT eclogites) in the context of metamorphic sole formation and subduction channel dynamics (see Fig. 13) for the P-T

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stability fields and schematic P-T paths of metamorphic sole rocks and eclogites). The

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authors attribute the mechanical coupling of thin, top layer of subducted oceanic crust to the

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base of the mantle wedge at contrasting thermal conditions and invoke the presence of two sequential rheological switches across the subduction interface. The welding occurs when the mechanical strength of the accreted crust closely approaches to that of the progressively cooled, anhydrous to partially hydrated mantle wedge (Agard et al., 2016). The first rheological switch (between mantle wedge and basalts) leads to the accretion of predominantly mafic crust at greater depths (P ~10 ± 2 kbar) and temperatures (T ~800 ± 50 °C) (Agard et al., 2016 and references cited therein), producing a HT metamorphic sole of mafic granulites. The second rheological switch (between mantle wedge and sediments) takes place at appreciably cooler (T ~550-650 °C) and shallower depth (P ~5 kbar) conditions, and gives rise to the welding of metasedimentary-dominated LT sole rocks of amphibolites and

Journal Pre-proof greenschists to the base of partially exhumed HT sole rocks. Taken together, these metamorphic sole rocks which are said to develop within the first 3 Myrs of subduction initiation constitute a composite stack, which from top to bottom are: basal peridotite of the ophiolitic mantle sequence → HT sole → LT sole. In this context, the stage 3 mylonite deformation event at T ~800 °C in the investigated mantle tectonite rock, the latter broadly falling in the same P-T window as that of the global HT metamorphic sole rocks can be linked with the HT accretion stage (Fig. 14b). We suggest

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the accretion of the underthrusted basaltic crust together with the abyssal mantle peridotite beneath the overlying mantle wedge to be linked with a switch over from an initial divergent

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to a later convergent plate tectonic setting (Fig. 14a-b). Although, mafic granulites in the HT

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sole are yet to be recorded from the NOC, such rocks do occur in the north-central segment of

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the Neo-Tethys (with respect to current geographic co-ordinates) (e.g. in the Saga ophiolitic mélange, south Tibet, Guilmette et al., 2012; Yarlung Zangbo ophiolites, Xigaze area, Tibet,

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Guilmette et al., 2008). This implies that mechanical coupling of the subducted basaltic crust

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at the onset of subduction, along warmer thermal gradients may be a general feature in the

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south-eastern segment of the Neo-Tethys (as exposed in the NOC) as well. Since the NOC is an ophiolitic mélange, the original contact between the basal peridotite and the HT metamorphic sole has been lost due to subsequent exhumation tectonics (see below). However, unlike the general model, which predicts a very short-lived residence of the HT metamorphic sole at their depth of formation (e.g. Agard et al., 2016), the accreted Nagaland peridotite appears to reflect a prolonged stay at depth along with the HT sole rocks. This delayed exhumation is indicated by our recognition of post-accretion, localised HT hybridmelt and hydrous fluid percolations in the accreted peridotite tectonite at the same depth level as the inferred HT sole rocks. This is because the HT metasomatic event in the mantle would imply transfer of melt and HT hydrous fluid from significantly deeper sources, including slab

Journal Pre-proof crust or hydrous mantle (Fig. 14c). This implies that the Nagaland spinel lherzolite is likely to be a piece of abyssal mantle peridotite accreted at the base of the overlying mantle during an early metamorphic sole formation stage. These accreted abyssal peridotites are then infiltrated by hybrid melt/fluid of high-pressure origin (> 12 kbar) during the stage 4 evolution (Fig. 14c). Progressive hydration of the mantle wedge further weakens its rheology allowing upwelling of the asthenospheric mantle wedge and its incorporation in an upward circulating

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subduction channel (Cloos, 1982; Gerya et al., 2002; Gorczyk et al., 2007). During this process, the channel expands in width and depth and the HP rocks are detached from the

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overlying mantle and incorporated in the convective upward circulating channel (e.g. Gerya

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et al., 2002) (Fig. 14d). The oceanic crust that enters the subduction channel at this stage (cf.

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cold mature stage of subduction) is metamorphosed under LT blueschist facies conditions and follows a hair-pin clockwise P-T path (with a TMax ~340 °C and ~11.5 kbar) for rocks with a

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history of single cycle burial-exhumation cycle (Ao & Bhowmik, 2014) (Fig. 14d). The final

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exhumation and emplacement of the associated mélange rocks appears to have taken place at

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shallower crustal level during the development of an accretionary complex (Fig. 14e). Although we are unable to record a supportive decompression path in the peridotite rocks, thermo-mechanical predictions do record a steep adiabatic decompression path at this later stage of evolution (Gorczyk et al., 2007). The current assorted occurrence of the modified abyssal mantle peridotite and HP metamorphic rocks of different thermal history in the NOC is attributed to combined exhumation tectonics, involving convective circulations in a hydrated mantle, and later thrust stacking in the accretionary wedge.

7. Conclusions The major results/conclusions of our study can be summarised as follows:

Journal Pre-proof 1. The basal Nagaland mantle spinel lherzolite records five sequential stages of hightemperature mineral assemblage evolution from its origin at a MOR to subductionaccretion with the overplate mantle. 2. The primary megacrystic clinopyroxene, orthopyroxene and spinel in the spinel lherzolite samples record equilibrations at mantle depths (P = 12-16 kbar) and at T > 1150 °C and low degrees (< 6-7 %) of dry melting of source mantle reservoir at a MOR setting.

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3. Pyroxene exsolutions and development of spinel + clinopyroxene coronae reveal a high-T cooling history from initial supra-solidus conditions to the sub-solidus stage.

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This can be related with the thermo-mechanical response following upwelling beneath

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the mid-oceanic ridge, and progressive movement away from the spreading centre.

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4. Formation of pyroxene neoblasts and mylonite banding indicates a high-T (~800°C) deformation and metamorphic recrystallisation event. This is linked with the switch

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over from an initial divergent to a later convergent plate tectonic setting as the slab

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mantle got accreted beneath the wedge mantle as part of a HT metamorphic sole.

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5. Clinopyroxene veins record an infiltration phase of silica-undersaturated melts in accreted slab mantle. The melts are likely to have originated by partial melting of the mantle wedge peridotite, whose composition were previously modified by reaction with an infiltrating, slab-derived basaltic melt. 6. Highly magnesian pargasite that replaced recrystallised pyroxene grains as well as vein clinopyroxene records a transient phase of high-temperature (T > 800 °C) hydration at mantle depths. 7. The exhumation of accreted and partially modified slab mantle and its tectonic juxtaposition with greenschist-pumpellyite-diopside and lawsonite blueschist facies metamorphic ensemble in an ophiolitic mélange in the Nagaland accretionary

Journal Pre-proof complex are attributed to combined exhumation tectonics involving convective circulations in a hydrated mantle and later thrust stacking in the accretionary wedge.

Acknowledgements We thank Saptarshi for his help in EPMA analyses in the EPMA National Facility, Indian Institute of Technology Kharagpur (IIT KGP). We thank Valentin Basch and an anonymous reviewer for the detailed reviews, and Marco Scambelluri for competent editorial handling.

also

from suggestions and

Accommodation provided

editorial corrections by Geoffrey Clarke.

by the Nagaland, State Mining Development Corporation,

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Aulbach and

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The manuscript has also benefited from constructive comments on its earlier version by Sonja

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Government of Nagaland, in their field hostel in Weziho is gratefully acknowledged. A.A.

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also acknowledges the Director, Wadia Institute of Himalayan Geology for providing

acknowledges financial support from the Indian Institute of Technology

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S.K.B.

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Funding

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facilities to carry out the final stages of work.

Kharagpur, in the form of a Cumulative Professional Development Allowance and Indian Space Research Organisation (ISRO) (Grant No. IIT / KCSTC / Chairman / New Approval / 15-16 / 09) to undertake geological fieldwork in remote and inaccessible part of the Nagaland state and relevant analytical work. A.A. acknowledges post-doctoral fellowship from the ISRO (Grant No. SRIC/NAG/2015/SEP/863).

Declaration of interests The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

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Rajkakati, M., Bhowmik, S.K., Ao, A., Ireland, T., Avila, J., Clarke, G., Bhandari, A., Aitchison, J., 2019. Early Jurassic eclogite facies metamorphism in the Nagaland

al

Ophiolite Complex, NE India: New insights into the antiquity and subduction tectonics

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within the Neo-Tethys 346, 105166.

Jo u

Rapp, R.P., Shimizu, N., Norman, M.D., Applegate, G.S., 1999. Reaction between slabderived melts and peridotite in the mantle wedge: experimental constraints at 3.8 GPa. Chemical Geology 160, 335-356. Schmidt, M.W., Poli, S., 1998. Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation. Earth and Planetary Science Letters 163, 361379. Searle, M.P., Noble, S.R., Cottle, J.M., Waters, D.J., Mitchell, A.H.G., Tin Hlaing. & Horstwood, M.S.A., 2007. Tectonic evolution of the Mogok Metamorphic belt, Burma (Myanmar) constrained by U-Th-Pb dating of metamorphic and magmatic rocks. Tectonics, 26, TC3014, doi:10.1029/2006TC002083.

Journal Pre-proof Shervais, J.W., Jean, M.M., 2012. Inside the subduction factory: Modeling fluid mobile element enrichment in the mantle wedge above a subduction zone. Geochimica et Cosmochimica Acta 95, 270-285. Singh, A.K., 2009. High-Al chromian spinel in peridotites of Manipur Ophiolite Complex, Indo-Myanmar Orogenic Belt: implication for petrogenesis and geotectonic setting. Current Science 96, 973-978. Singh, A.K., 2013. Petrology and geochemistry of Abyssal Peridotites from the Manipur

oo

f

Ophiolite Complex, Indo-Myanmar Orogenic Belt, Northeast India: implication for melt generation in mid-oceanic ridge environment. Journal of Asian Earth Sciences

pr

66, 258-276.

e-

Singh, A.K., Chung, S.L., Bikramaditya, R.K., Lee, H.Y., 2017a. New U-Pb zircon ages of

Pr

plagiogranites from the Nagaland-Manipur Ophiolites, Indo-Myanmar Orogenic Belt, NE India. Journal of the Geological Society 174, 170-179. A.K.,

Nayak,

R.,

S.,

Subramanyam,

K.S.V.,

Thakur,

S.S.,

R.K., Satyanarayanan, M., 2017b. Genesis and tectonic

rn

Bikramaditya Singh,

Khogenkumar,

al

Singh,

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implications of cumulate pyroxenites and tectoniteperidotites from the NagalandManipur ophiolites, Northeast India: Constraints from mineralogical and geochemical characteristics. Geological Journal 52, 415-436. Stixrude, L., Lithgow-Bertelloni, C., 2011. Thermodynamics of mantle minerals-II. Phase equilibria. Geophysical Journal International 184, 1180-1213. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. Geological Society, London, Special Publications 42, 313-345.

Journal Pre-proof Sun, C., Liang, Y., 2012. Distribution of REE between clinopyroxene and basaltic melt along a mantle adiabat: Effects of major element composition, water, and temperature. Contributions to Mineralogy and Petrology 163, 807-823. Takahashi, E., Shimazaki, T., Tsuzaki, Y., Yoshida, H., 1993. Melting study of a peridotite KLB-1 to 6.5 GPa, and the origin of basaltic magmas. Philosophical Transactions of the Royal Society of London, Series A 342, 105-120. Tatsumi, Y., 2005. The subduction factory: how it operates in the evolving Earth. GSA

oo

f

Today 15, 4.

Tatsumi, Y., Eggins, S., 1995. Subduction zone magmatism (Vol. 1). Wiley.

pr

Taylor, W.R., 1998. An experimental test of some geothermometer and geobarometer

e-

formulations for upper mantle peridotites with application to the thermobarometry of

Pr

fertile lherzolite and garnet websterite. N Jarbh Mineral Abh 172, 381-408. Tollan, P.M.E., Dale, C.W., Hermann, J., Davidson, J.P., Arculus, R.J., 2017. Generation and

al

Modification of the Mantle Wedge and Lithosphere beneath the West Bismarck Island

rn

Arc: Melting, Metasomatism and Thermal History of Peridotite Xenoliths from Ritter

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Island. Journal of Petrology 58, 1475-1510. Uysal, İ., Ersoy, E.Y., Karslı, O., Dilek, Y., Sadıklar, M.B., Ottley, C.J., Tiepolo, M., Meisel, T., 2012. Coexistence of abyssal and ultra-depleted SSZ type mantle peridotites in a Neo-Tethyan Ophiolite in SW Turkey: Constraints from mineral composition, wholerock

geochemistry

(major-trace-REE-PGE),

and

Re-Os

isotope

systematics.

Lithos132, 50-69. Venkataramana, P., Dutta, A.K., Acharyya, S.K., 1986. Petrography and petrochemistry of Naga Hills Ophiolite. Geological Survey of India, Memoirs 119, 33-63.

Journal Pre-proof Wallace, M., Green, D.H., 1991. The effect of bulk rock composition on the stability of amphibole

in

the

upper mantle: implications for solidus positions and

mantle

metasomatism. Mineralogy and Petrology 44, 1-19. Warren, J.M., 2016. Global variations in abyssal peridotite compositions. Lithos 248, 193219. Witt-Eickschen, G., Seck, H.A., 1991. Solubility of Ca and Al in orthopyroxene from spinel peridotite: an improved version of an empirical geothermometer. Contributions to

oo

f

Mineralogy and Petrology 106, 431-439.

Yao, L., Sun, C., Liang, Y., 2012. A parameterized model for REE partitioning between low-

pr

Ca pyroxene and basaltic melts with applications to adiabatic mantle melting and

e-

pyroxenitederived melt and peridotite interaction. Contributions to Mineralogy and

Pr

Petrology 164, 261-280.

Ziberna L., Klemme S., 2016. Application of thermodynamic modelling to natural mantle

al

xenoliths: examples of density variations and pressure-temperature evolution of the

Jo u

rn

lithospheric mantle. Contributions to Mineralogy & Petrology 171, 16.

Fig. 1. Location of the Indo-Myanmar ranges (IMR) in the tectonic framework of Indian, Asian and Burmese plates (modified after Searle et al., 2007). The marked box in the eastern margin of the IMR marks the location of the study area, given in detail in Fig. 2a.

Fig. 2. (a) Detailed geological map of the NOC (after Anon, 1986) showing the different litho-units of the ophiolite complex. Locations of the studied samples along with the sample numbers are shown. Also shown are the locations of pumpellyite-diopside (PD), blueschist (BS) and amphibolite (AM) units, along with their peak P-T conditions and P-T paths after Ao and Bhowmik (2014) (AB14) and Bhowmik and Ao (2016) (BA16). (b) Geologic cross-

Journal Pre-proof section of the ophiolite belt along the transect A-B, marked in (a). The section shows the studied spinel peridotites occurring at the base of the ophiolite unit. Other abbreviations used: GS, Greenschist; EC, Eclogite; LBS/EBS, Lawsonite blueschist/epidote blueschist.

Fig. 3. Field photographs of the spinel lherzolite samples (a) 117A and (b) 44A.

Fig. 4. (a) False-colour thin-section map of a tectonised spinel lherzolite sample 117A, and

oo

f

generated using a photo editing software, Picasa shows the occurrences of pyroxene megacrysts in a mosaic of serpentine group of minerals. Some of the pyroxene crystals are

pr

deformed and elongated, giving rise to a banded appearance to the rock. Small boxes mark

e-

the locations of BSE images shown in detail in Fig. 5. (b) False-colour thin-section image of

Pr

a spinel lherzolite sample 44A showing the distributions of different minerals. Small boxes

al

mark the locations of the BSE images shown in detail in Fig. 5.

rn

Fig. 5. BSE (a-f, i-k) and Mg- and Ca-X-ray element (g-h) images of spinel lherzolite sample

Orthopyroxene

Jo u

117A and 44A. Rectangular box in Fig. 5f shows the location of Fig. 5g-i. (a-c) megacrysts

showing

exsolution

lamella

of

clinopyroxene

and

an

anastomosing network of fracture-controlled replacement to serpentine (a-b). Some of these lamellae escape orthopyroxene host producing a granular exsolution (c). Bleb-like aluminous spinels locally occur within orthopyroxene. (d) Coronal clinopyroxene-spinel assemblage developing around olivine inclusion (mostly serpentinised) against orthopyroxene host. (e) An

elongated

band

of

fine-grained,

neoblast

aggregates

of clinopyroxene

around

clinopyroxene megacryst. (f-i) Orthopyroxene megacryst is surrounded by fine-grained, neoblast aggregates of orthopyroxene, clinopyroxene and locally olivine. Late hornblende replacing neoblast orthopyroxene and clinopyroxene (i). (j) Part of a vein of clinopyroxene in

Journal Pre-proof deformed orthopyroxene megacryst. Clinopyroxene in the vein is rimmed by hornblende. (k) The occurrence of hornblende as thin, discontinuous blebs parallel to cleavage traces in host clinopyroxene megacryst.

Fig. 6. Binary plots showing the compositional variations of different textural varieties of orthopyroxene (a), clinopyroxene (b), spinel (c), and amphibole (d). (a) Megacrystic orthopyroxene core contains higher Al2 O3 and Cr2 O3 compared to its neoblast variety. Some

oo

f

of the orthopyroxene megacryst rim compositions overlap with that of its neoblast variety. (b) Megacrystic clinopyroxene contains higher Al2 O3 and TiO 2 compared to its neoblast and vein

pr

textural types. Abyssal peridotite and forearc peridotite fields are after Uysal et al. (2012). (c)

e-

Spinel Cr# vs. XMg plot. Abyssal peridotite field is after Dick and Bullen (1984) and Arai

Pr

(1994). Forearc peridotite field is from Ishii et al. (1992) and Parkinson and Pearce (1998). Partial melting trend is from Arai (1992). Published spinel data from Andaman ophiolite

al

(AO), Nagaland ophiolite (NO), Manipur ophiolite (MO) belts are plotted for a comparison.

Neo,

Neoblast;

Jo u

core/rim;

rn

(d) Amphiboles are high magnesian pargasite. Abbreviations used: Mc(C)/Mc(R): Megacryst Rel,

Relics;

Lam/Cor,

Lamellae/corona;

Met./Mag.,

Metamorphic/magmatic.

Fig.

7.

(a-d)

Chondrite-normalised

REE

and

multielement

spidergram

plots

of

orthopyroxenes in samples 117A (a, c) and 44A (b, d). Chondrite composition in this and Figs. 8-9 is after Sun and McDonough (1989).

Fig. 8. (a-d): Chondrite-normalised REE and spidergram plots of clinopyroxenes in samples 117A (a, c) and 44A (b, d).

Journal Pre-proof Fig. 9. (a-b) Chondrite-normalised REE and spidergram plots of hornblende in sample 117A.

Fig. 10. NCFMASOCr P-T pseudosection of spinel lherzolite sample 44A. Mineral abbreviations: g, garnet; cpx, clinopyroxene; opx, orthopyroxene; pl, plagioclase; ol, olivine; spl, Al-rich spinel; cm, Cr-rich spinel; liq, liquid. The bold dashed line shows the position of solidus (liq-in with increasing T). The sub-solidus region is contoured with isopleths of ortjopyroxene (Aliv and Ca p.f.u., where p.f.u. = per formula unit) and spinel (Cr# and XFe2+).

oo

f

The P-T field of initiation of stage 2 sub-solidus cooling is marked by the isopleths of metamorphic spinel (Cr# = 0.14-0.15; XFe2+ = 0.25) and orthopyroxene (Ca p.f.u = 0.04The P-T field of the lower temperature end of the cooling stage is defined by the

pr

0.05).

Pr

e-

isopleth of orthopyroxene rim composition (Ca p.f.u. = 0.02).

Fig. 11. NCFMASH P-T pseudosection calculations in sample 44A. Mineral abbreviations: g,

al

garnet; cpx, clinopyroxene; opx, orthopyroxene; pl, plagioclase; ol, olivine; spl, spinel; hbl,

rn

hornblende; tlc, talc; chl, chlorite. Thin dashed lines (green) are maximum Na (p.f.u.) content

Jo u

in the pargasitic amphibole. Also plotted are calculated isopleths of maximum molar (Na2 O + K2 O) contents (in %) in pargasite (dashed red line) following the methodology of Mandler & Grove (2016). Thick dashed lines (in shades of blue) are amphibole-out curves from experimental studies. Abbreviations used: G 73, Green, 1973; WG 91, Wallace & Green, 1991; M 74: Millhollen et al., 1974; SP 98: Schmidt & Poli, 1998.

Fig. 12. (a-b) Binary Ti vs. Nd (a) and Ti vs. Zr (b) plots of megacrystic clinopyroxene in the spinel lherzolite sample 44A. Abyssal and supra-subduction zone (SSZ) peridotite fields and melting trends are from Bizimis et al. (2000). Clinopyroxene compositional trends in peridotites from the Nagaland (NO), Manipur (MO) and Kalaymyo (KP) ophiolite belts are

Journal Pre-proof also plotted for a comparison. (c) REE plot of the megacrystic clinopyroxene in the investigated lherzolite samples. Data are compared with clinopyroxene compositions in abyssal peridotites (America-Antartica and Southwest Indian Ridge, Johnson et al., 1990), NO, MO, KP and Andaman ophiolite (AO). Data source: NO/MO/AO, Ghose et al. (2018); KP, Liu et al. (2016).

Fig. 13. Reconstructed metamorphic P-T path of evolution of the Nagaland spinel lherzolite

oo

f

rocks (N-Spl-L). The near isobaric cooling path links different stages of metamorphic reconstitutions from supra-solidus to sub-solidus regions (see text for more details). The

pr

stability fields of plagioclase lherzolite (Pl-L), spinel lherzolite (Spl-L) and garnet lherzolite

e-

(Grt-L) are taken from the computed phase diagram (this study). The transition from spinel to

Pr

garnet lherzolite has been additionally demarcated from experimental works (e.g. curves KO 04 & ZK 16) and P-T pseudosection calculation on model peridotite rock composition (curve

al

JH 15). Solidus curves for an average dry peridotite (McKenzie & Bickle 1998) and wet (0.2

rn

% H2 O) pyrolite (Green, 1973) compositions are also plotted for a comparison. The P-T

Jo u

stability fields of HT sole (HTS), LT sole (LTS), HT eclogite (HTE) and LT eclogite (LTE) and the P-T boundary demarcating accretion from non-accretion are from Agard et al. (2016). Arrows in HTE mark counterclockwise (CCW) P-T paths in these rocks. P-T gradient for hot subduction is after Liu et al. (1996). Looping CCW metamorphic P-T paths for the M1 and M2 metamorphic cycles of the Nagaland amphibolites (referred to as NAM 1 and NAM2 respectively) are after Bhowmik & Ao (2016). Area marked by dashed region shows the P-T stability field of metamorphic soles worldwide. Abbreviations used: KO 00, Klemme & O’Neill (2000); JH, Jenkins & Holland (2015); ZK 16, Ziberna & Klemme (2016).

Journal Pre-proof Fig. 14. Cartoon sketches showing the geodynamic evolution of the Nagaland Ophiolite Complex from generation of the oceanic lithosphere at the Mid-Oceanic Ridge setting (a) through subduction initiation (b) to subduction maturation and exhumation (c-e).

hornblende and spinel in

Lhz

Lh z

Lh z 44 A

Lh z 11 7 A

Lh z 11 7 A

44A

44 A

Prg 2_1 5 Cor >Cp x

Sp 3_ 17

Sp 2_ 2

Sp 1_ 36

Sp 3_ 33

Me t

M et

M ag

43.2 1

0.0 1

0. 04

M ag b. d.l .

0. 05 53 .7 8 13 .5 2 12 .2 0

0.0 7

b.d .l.

0.0 3

0.0 6

0. 07

0.0 7

1.06

1.41

1.9 3

1.31

0.0 3

Al2 O3

5.8 3

4.8 2

3.3 5

4.3 5

2.4 0

4. 43

3.8 5

15.1 9

15.0 1

15. 57

15.4 6

56. 54

Cr2 O3

0.6 0

0.5 1

0.2 1

0.3 8

0.1 2

0. 41

0.1 7

1.23

0.98

1.4 5

1.44

10. 82

Fe O

6.5 3

6.5 0

6.5 3

6.3 9

6.5 3

6. 59

6.4 6

3.75

3.53

3.8 1

3.55

12. 42

Mn O

0.1 5

0.2 9

0.1 3

0.1 6

0.1 4

0.1 4

b.d.l.

0.02

b.d .l.

0.03

0.1 9

Mg O

32. 40

33. 31

33. 49

33. 16

34. 23

0. 12 32 .6 1

33. 80

17.1 9

17.6 1

16. 45

17.3 7

19. 81

0. 12 19 .3 3

Ca O

0.5 3

1.4 1

0.4 5

0.4 5

0.4 2

0. 85

1.2 0

12.8 5

12.7 7

12. 34

12.5 2

b.d .l.

0. 04

Na2 O K2 O

b.d .l. 0.0 1

0.0 7 0.0 2

0.0 2 b.d .l.

0.0 2 b.d .l.

0.0 1 b.d .l.

0. 0.0 02 1 b. b.d d.l .l.

2.62 b.d.l .

0.0 1 0.0 2

0. 03 b. d.l

rn

al

2

0.1 3

Jo u

TiO

Pr

e-

pr

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f

Table 1: Representative mineral chemical analysis of orthopyroxene, spinel lherzolite samples from the NOC. Roc Lh Lh Lh Lh Lh Lh Lh Lh k z z z z z z z Lhz Lhz z Sa mpl 11 44 11 44 11 44 44 117 117 44 e 7A A 7A A 7A A A A A A Min Op Op Op Op Op O Op Pr eral x x x x x px x Prg Prg g Ana 3_ 3_ 2_ 1_ 2_ 2_ 3_ 1_ l. # 56 30 24 7 27 4 12 1_2 2_5 2 Lam Cor I^ Mc Mc Mc Mc Ne Ne ^Cp >Cp Cp Site (C) (C) (R) (R) o o Rel x x x 54 SiO 53. 54. 55. 55. 56. .8 55. 43.3 43.6 42. 72 22 74 07 12 7 16 8 0 51 2

2.98

2.91

b.d.l.

0.02

3.2 5 0.0 2

0. 05 57 .6 6 9. 40 11 .9 9 b. d.l . 19 .9 8 b. d.l . b. d.l . b. d.l

0. 04 b. d.l . 56 .7 4 10 .7 1 11 .8 1 0. 07 19 .6 1 0. 01 0. 04 b. d.l

Journal Pre-proof

+

Mn Mg Ca Na K Su m XMg

97.8 6

97. 33

97.5 1

10 0.0 5

6 1.8 6

6 1.9 2

6 1.8 9

6 1.9 3

6 1. 90

6 1.8 9

23

23

23

4

4

4

4

6.15

6.15

6.10

-

-

-

-

0.1 6 0.0 1 0.0 4 0.1 5

0.11

0.15

0.14

2.54

2.50

0.14

0.11

23 6.0 6 0.2 1 2.6 2 0.1 6 0.0 1 0.4 5

0.2 4 0.0 2 0.0 2 0.1 7

0.1 4 0.0 1 0.0 1 0.1 8

0.0 3 0.1 6

0. 18 0. 01 0. 01 0. 18

1.7 2 0.0 2

0.1 8 0.0 1 0.0 2 0.1 8 0.0 1 1.7 0 0.0 2

0.1 0

1.6 7 0.0 2

0.2 0 0.0 1 0.0 4 0.1 5 0.0 1 1.7 0 0.0 5

1.7 3 0.2 2 0.0 4 0.2 3

1. 68 0. 28 0. 04 0. 23

1. 77 0. 19 0. 04 0. 23

1. 75 0. 22 0. 03 0. 23

1.7 6 0.0 2

1. 68 0. 03

1.7 3 0.0 4

-

-

3.63

3.70

3.66

0.7 7

0. 76

0. 78

0. 77

1.89

-

-

-

-

4.0 0 0.9 1

4.0 2 0.9 2

4.0 0 0.9 1

4.0 0 0.9 0

4.0 0 0.9 2

4. 00 0. 90

4.0 1 0.9 2

0.72 15.7 3

3.0 0 0.7 7 0.1 1

3. 00 0. 77 0. 14

3. 00 0. 77 0. 10

3. 00 0. 77 0. 11

0.05 0.40

2.57 0.16

f

6 1.8 6

10 0.9 8

oo

Fe2

10 0.0 3

0.09 0.33

pr

+

10 0.0 1

e-

Cr Fe3

99. 92

Pr

Al

10 1.3 6

al

Ti

97.6 3

. 99 .1 1

rn

Si

99. 90

Jo u

Tot al Ox yge n

. 99 .9 7

1.95

1.93

0.82 15.8 0

0.80 15.7 8

0.90

0.92

3.5 0 1.8 9 0.9 0 15. 79 0.8 9

0.18 0.24 -

0.94

. 99 .0 8

. 99 .0 3

Cr# Mc: Megacryst; C: Core; R: Rim; Lam: Lamellae; Neo: Neoblast; Cor: Corona; I: Inclusion; Rel: Relics; Met: metamorphic; Mag: magmatic; b.d.l.: below detection limit; ^: within; >: around; X Mg = Mg/(Mg+Fe2+); Cr# = Cr/Cr+Al. Table 2: Representative mineral chemical lherzolite samples from the NOC. Roc k Lhz Lhz Lhz Lhz Lhz Sam 117 117 ple A 44A A 44A 117A Min eral Cpx Cpx Cpx Cpx Cpx Anal 1_3 2_1 3_5 .# 1_1 8 6 4 3_22 Mc( Mc( Mc( Mc( Lam^ Site C) C) R) R) Opx SiO2 52.1 52.0 51. 51.6 50.60

analysis of clinopyroxene and olivine in spinel

Lhz 44A

Lhz 117 A

Lhz 117 A

Cpx

Cpx

1_11 Lam^ Opx 50.72

2_1 Neo (V) 52.6

Cpx 2_3 3 Neo 52.6

Lh z 44 A Cp x 1_ 17 Co r 51.

Lhz 117 A Ol 1_4 5 Rel 40.7

Lh z 11 7A

44A

Ol 2_ 38 Ne o 41.

Ol 3_1 9 I^O px 40.4

Lhz

Lh z 44 A Ol 1_ 22 Rel 40.

Journal Pre-proof

K2O Tota l Oxy gen

0.37 0.74 b.d. l. 0.04 100. 100. 53 38 6

6

1.88 1.88

Ti

0.01 0.01

Al

0.19 0.22

Cr

0.02 0.02

Fe3+

0.03 0.05

Fe2+ Mn

0.03 0.03 0.01

Mg

0.86 0.87

Ca

0.95 0.88

Na K

0.03 0.05 -

Sum

4.00 4.00

XMg

0.40

0.25

0.18

6.13

5.30

6.43

3.60

2.87

0.95

0.73

0.95

0.46

0.26

2.62

2.31

2.26

2.35

2.37

0.08 15.9 6 23.2 9

0.04

0.08

15.49

15.18

24.58

23.69

0.09 16.6 3 23.8 0

0.06 17.0 2 24.4 8

0.84 b.d. l. 102. 12

0.35

0.73

b.d.l. 100.0 0

b.d.l. 100.4 4

6

6

6

1.84

1.84

0.01

0.96 0.96

0.39 0.38 b.d.l . 0.01 100. 100. 19 30 6

6

1.83

1.91

1.90

0.02

0.01

0.01

0.01

0.23

0.27

0.15

0.12

0.02

0.03

0.01

0.01

0.07

0.06

0.07

0.04

0.07

-

0.01 -

-

0.03 -

-

0.85

0.84

0.82

0.90

0.92

0.89

0.96

0.92

0.92

0.95

0.06 -

0.03 -

0.05 -

0.03 -

0.03 -

4.00

4.00

4.00

4.00

4.00

1.00

0.99

1.00

0.96

1.00

0.26 0.03

Jo u

Si

6 1.8 5 0.0 1 0.2 3 0.0 2 0.0 5 0.0 3 0.8 5 0.9 3 0.0 3 4.0 0 0.9 7

0.60

f

CaO Na2 O

0.03 0.17 15.9 16.0 9 9 24.4 22.6 1 2

0.35

06 0.3 9 6.6 2 1.0 1 2.2 4 b.d .l. 14. 71 22. 95 0.8 8 0.0 1 99. 86

oo

2.03 2.75

0.70 0.64

7

pr

FeO Mn O Mg O

4.53 5.05

2

e-

0.31 0.28

0

Pr

TiO2 Al2 O3 Cr2 O3

02 0.4 1 5.4 4 0.6 5 2.4 7 0.0 3 15. 72 23. 80 0.4 4 b.d. l. 99. 98

al

2

rn

6

6 1.8 6 0.0 1 0.2 8 0.0 3 0.0 2 0.0 5 0.8 0 0.8 9 0.0 6 4.0 0 0.9 4

4 b.d. l. b.d. l. b.d. l. 9.62 0.14 49.6 7 b.d. l. b.d. l. b.d. l. 100. 17 4

00 b.d .l. b.d .l. b.d .l. 9.7 3 0.1 6 49. 06 b.d .l. 0.0 3 b.d .l. 99. 98

2 b.d. l. 0.02 b.d. l. 9.21 0.11 50.3 8 0.03 b.d. l. b.d. l. 100. 22

0.99

4 1.0 0

0.99

4 1.0 0

-

-

-

-

-

-

-

-

-

-

-

-

0.01

0.01

1.81

0.2 0 1.7 9

1.83

0.2 0 1.7 8

-

-

-

-

-

3.0 0 0.9 0

-

3.0 0 0.9 0

0.19 -

3.00 0.91

4

92 b.d .l. 0.0 1 b.d .l. 9.8 7 0.2 0 48. 78 0.0 2 b.d .l. 0.0 2 99. 82

0.18 -

3.01 0.91

Mc: Megacryst; C: Core; R: Rim; Lam: Lamellae; Neo: Neoblast; V: Vein; Cor: Corona; Rel: Relics; b.d.l.: below detection limit; ^: within; XMg = Mg/(Mg+Fe 2+).

Journal Pre-proof Table 3: Results of thermometry at different stages of evolution of the abyssal mantle peridotite, Nagaland Ophiolite Complex.

Ca

Cp x

g

Mc(C) 3_ 29

Mc( C) 0.2 0

0.0 3

0.9 2

3_ 31

Mc(C) 44 A

3_ 27

Mc( C) 0.2 4

0.0 3

0.9 2

Mean (1 σ) Cor

Mc(C) 1_ 24

3_ 31

0.2 2

0.0 4

0.9 0

1_ 17

Mc(C) 1_ 34

0.2 1

0.0 2

0.9 0

1_ 32

Mc(R) 0.2 3

0.0 2

0.9 1

0.1 8 Mc(R)

2_ 13

0.2 4

Mc(C) 1_ 42

0.0 2

0.9 0

Jo u

1_ 7

rn

Mc(R)

0.2 1

0.0 2

0.0 3

1_ 3

0.9 0

0.9 0

1_ 9

2_ 12

1_ 40

Mc(C) 3_ 23

0.1 5

0.0 2

0.9 2

3_ 22

Mc(R) 3_ 47 44 A

0.1 4

0.0 2

0.9 2

3_ 46

Mc(R) 3_ 26

0.1 3

0.0 3

0.9 2

3_ 25

0.1 5

0.0 2

0.9 1

2_ 7

Mc(R) 2_ 5

A l 0 . 3 0 0 . 3 0

0 . 2 8 0 . 2 5 0 . 2 9 0 . 2 4 0 . 3 0 0 . 2 8 0 . 2 2 0 . 2 4 0 . 1 5 0 . 2

Ex. La m

al

1_ 4

Ex. La m

Ex. La m Ex. La m Ex. La m Ex. La m Ex. La m Ex. La m Ex. La m

Al-Cr in Opx Therm

XMg

TL (°C)

*TBKN (°C)

TWS (°C)

Two Pyroxene Therm TB TT KN (° (° C) C)

C a 0 . 8 9 0 . 8 9

1.00

132 0

985

1070

87 0

77 0

76 5 76 5± 05

0 . 8 9 0 . 9 1

Type

Sta ge 1

134 5 133 5±1 5

980

1155

985± 05

1115± 45

88 0 87 5± 05

0.94

110 5

1075

1075

74 0

74 5

0.98

102 0

900

1030

70 5

67 5

0.97

-

900

1080

73 0

73 0

1.00

-

880

955

65 5

65 5

0.99

-

900

1115

66 0

66 0

0.95

-

925

1045

77 0

77 5

1.00

-

950

925

69 0

70 0

1.00

-

855

900

67 0

69 0

1.00

-

980

870

68 0

68 0

1.00

-

930

930

77 5

79 0

1.00

oo

XM Al

Ca-in Opx Ther m

Compositi on

Composition

O px

RE E The rm

f

Tex ture

pr

Texture

e-

M in er al

Pr

Sa m pl e

0 . 9 0 . 9 2 0 . 9 1 0 . 8 9 0 . 9 1 0 . 9 1 0 . 9 3 0 . 8

Sta ge 2

Journal Pre-proof 5

Mc(C) 4_ 4

0.2 0

0.0 2

0.2 0

0.0 2

0.2 0

0.0 2

Mc(C) 4_ 4 Mc(C) 4_ 11

0 . 2 6 0 . 2 6 0 . 2 3

0 . 9 4 0 . 9 2 0 . 9 3

106 5±0 2

990± 85

1000± 85

71 0± 40

71 0± 05

1.00

980

890

975

66 0

60 5

1.00

985

890

975

74 5

71 0

73 0 68 0± 55

Sta ge 1

103 0 100 0±2 2

870

1050

885± 10

1000± 35

74 0 71 5± 40

-

900

1000

52 5

51 0

1.00

-

870

985

56 5

54 5

0.98

-

890

980

63 0

62 5

1.00

-

855

1010

67 0

67 0

-

900

1065

-

885± 20

1010± 30

56 5 59 0± 50

56 0 58 0± 60

1.00

-

800

760

56 0

55 5

1.00

-

890

820

56 5

55 5

1.00

-

900

830

61 0

59 5

-

800 850± 50

700 780±5 0

63 5 59 5±

64 5 59 0±

1.00

3_ 25

0.0 2

0.1 8

0.0 2

0.2 1

0.0 2

Mc(C) 3_ 27 Mc(C) 3_ 29 Mc(C) 3_ 74

0.2 2

Neo 2_ 29

0.0 2

Jo u

11 7 A

Mean (1 σ) Neo

0.0 8

0.0 1

0.9 1

2_ 22

Neo 2_ 31

Neo 0.1 1

0.0 2

0.9 0

2_ 22

Neo 2_ 44 11 7 A

Neo 0.1 2

0.0 2

0.9 1

2_ 33

Neo 2_ 37

Neo 0.0 4

0.0 1

pr

0.1 8

Mc(C)

0.9 1

2_ 41 Mean (1 σ)

0.99

e-

0.0 2

0 . 9 6 0 . 9 6 0 . 9 5 0 . 9 3 0 . 9 5

Pr

0.2 0

0 . 2 3 0 . 1 7 0 . 2 1 0 . 2 7 0 . 2 7

al

3_ 23

rn

Mc(C)

Mean (1 σ) Ex. La 0.9 3_ m 1 22 Ex. La 0.9 3_ m 1 24 Ex. La 0.9 3_ m 1 26 Ex. La 0.9 3_ m 1 28 Ex. La 0.9 3_ m 1 73

oo

f

11 7 A

Mean (1 σ) Mc( C) 0.9 2_ 2 14 Mc( C) 0.9 4_ 2 5 Mc( C) 0.9 4_ 2 12

9

0 . 1 4 0 . 1 4 0 . 1 2 0 . 1 3

0 . 9 5 0 . 9 5 0 . 9 5 0 . 9 5

Sta ge 2 1.00

Sta ge 3 0.98

-

Journal Pre-proof 30 11 7 A

2_ 7 4_ 18

Mc( 0.1 R) 3 M c( R ) 0.12

0.0 2

0.0 2

0.9 1

0.9 2

2_ 1

4_ 17

Vei n

Vei n

0.15

0.92 0.96

-

890

830

35 77 0

77 0

Sta ge 4

84 5

84 5 81 900± 845±1 0± Mean (1 σ) 10 3 40 Abbreviations used: TL , Liang et al., (2013); TBKN , Brey and Kohler (1990); TT, Taylor (1998); *TBKN , Brey Kohler (1990); TWS, Witt-Eickschen and Seck (1991); Mc(C), megacryst core; Mc(R), megacryst rim; Cor, Corona; Neo, neoblast; Ex. Lam, exsolution lamellae. 0.14

0.91

1.00

-

910

855

81 0± 40 and

Jo u

rn

al

Pr

e-

pr

oo

f

Highlights:  The basal mantle peridotite of spinel lherzolite composition in the NOC originated in a MOR setting.  The lherzolite records initial equilibration at mantle depths (P = 12-16 kbar) and at T > 1150 °C.  The peridotite was later transferred into the subduction zone and accreted as a HT (T ~800 o C at mantle depths) metamorphic sole beneath the mantle wedge.  The sole rock was later infiltrated by hybridised mantle-wedge melts in an SSZ setting.  A HT (T > 800 o C) hydrous fluid fluxing stabilised pargasite in the metasomatized peridotite.