Element and isotopic signature of re-fertilized mantle peridotite as determined by nanopowder and olivine LA-ICPMS analyses

Journal Pre-proof Element and isotopic signature of re-fertilized mantle peridotite as determined by nanopowder and olivine LA-ICPMS analyses

Christopher J.M. Lawley, D. Graham Pearson, Pedro Waterton, Alex Zagorevski, Jean H. Bédard, Simon E. Jackson, Duane C. Petts, Bruce A. Kjarsgaard, Shuangquan Zhang, Donald Wright PII:

S0009-2541(20)30003-6

DOI:

https://doi.org/10.1016/j.chemgeo.2020.119464

Reference:

CHEMGE 119464

To appear in:

Chemical Geology

Received date:

3 July 2019

Revised date:

27 December 2019

Accepted date:

3 January 2020

Please cite this article as: C.J.M. Lawley, D.G. Pearson, P. Waterton, et al., Element and isotopic signature of re-fertilized mantle peridotite as determined by nanopowder and olivine LA-ICPMS analyses, Chemical Geology (2020), https://doi.org/10.1016/ j.chemgeo.2020.119464

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© 2020 Published by Elsevier.

Journal Pre-proof Element and isotopic signature of re-fertilized mantle peridotite as determined by nanopowder and olivine LA-ICPMS analyses Christopher J.M. Lawley1†, D. Graham Pearson2, Pedro Waterton2, Alex Zagorevski1, Jean H. Bédard3, Simon E. Jackson1, Duane C. Petts1, Bruce A. Kjarsgaard1 Shuangquan Zhang4 and Donald Wright5 1

Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, K1A 0E8 Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada, T6G 2E3 3 Ressources Naturelles Canada, Commission Géologique du Canada, Québec, Québec, G1K 9A9 4 Department of Earth sciences, Carleton University, Room 2115 Herzberg Laboratories, 1125 Colonel By Drive, Ottawa, Ontario Canada, K1S 5B6 5 Peridot Geoscience Ltd., Ottawa, Ontario, Canada †

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corresponding author: [email protected]

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Keywords: ophiolite, metasomatism, Nahlin, Atlin, LA-ICPMS, nanopowder

Abstract

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The lithospheric mantle should be depleted in base- and precious-metals as these elements are

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transferred to the crust during partial melting. However, some melt-depleted mantle peridotites are enriched in these ore-forming elements. This may reflect re-fertilization of the mantle lithosphere and/or

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sequestering of these elements by residual mantle phase(s). Both processes remain poorly understood because of the low abundances of incompatible elements in peridotite and the nugget-like distribution of

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digestion-resistant mantle phases that pose analytical challenges for conventional geochemical methods.

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Herein we report new major and trace element concentrations for a suite of mantle peridotite and pyroxenite samples from the Late Permian to Middle Triassic Nahlin ophiolite (Cache Creek terrane, British Columbia, Canada) using Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LAICPMS) analysis of nanoparticulate powders and olivine. Compatible to moderately incompatible element concentrations suggest that Nahlin ophiolite peridotites represent residues after ≥ 20% melt extraction. Pyroxenite dykes and replacive dunite bands are folded and closely intercalated with residual harzburgite. These field relationships, coupled with the presence of intergranular base metal sulphide, clinopyroxene and Cr-spinel at the microscale, point to percolating melts that variably re-fertilized melt-depleted mantle peridotite. Radiogenic Pb (206Pb/204Pb = 15.402–19.050;

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Pb/204Pb = 15.127–15.633;

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Pb/204Pb =

Journal Pre-proof 34.980–38.434; n = 45) and Os (187Os/188Os 0.1143–0.5745; n = 58) isotope compositions for a subset of melt-depleted peridotite samples further support metasomatic re-fertilization of these elements. Other oreforming elements are also implicated in these metasomatic reactions because some melt-depleted peridotite samples are enriched relative to the primitive mantle, opposite to their expected behaviour during partial melting. New LA-ICPMS analysis of fresh olivine further demonstrates that a significant proportion of the highly incompatible element budget for the most melt-depleted rocks is either hosted by,

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and/or occurs as trapped inclusions within, the olivine-rich residues. Trapped phases from past melting and/or re-fertilization events are the preferred explanation for unradiogenic Pb isotope compositions and

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Paleozoic to Paleoproterozoic Re-depletion model ages, which predate the Nahlin ophiolite by over one

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billion years.

1 Introduction

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The lithospheric mantle comprises refractory peridotites that underwent variable degrees of melt

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extraction (Bodinier and Godard, 2014; Coleman, 1977; Griffin et al., 2009; Ishiwatari, 1985; Pearce et al., 2000; Pearson et al., 2007; Zhou et al., 2005). Incompatible elements that readily partition into silicate

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liquids are removed during melt extraction, resulting in a mantle residue that is depleted in large ion

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lithophile elements (LILE), high field-strength elements (HFSE) and rare earth elements (REE) relative to the Primitive Mantle (PM)(Hofmann, 1997). Compatible elements (e.g., some first-row transition elements such as Ni + Cr + Co) show the opposite behaviour and become progressively enriched within residual mineral phases as partial melting progresses. The compatible versus incompatible behaviour of base (Cu, Pb, Zn)- and precious (Au, PGE)metals during melt extraction is more complex because these trace (low ppm) to ultratrace elements (low ppb) are strongly dependent on the availability and behaviour of S (Lorand and Luguet, 2016). Due to the extremely high partition coefficients between sulphide and silicate liquids, most base- and preciousmetals are first concentrated in residual base metal sulphide (BMS) during the initial (< 10%) stages of partial melting (Mungall and Brenan, 2014). However, progressive partial melting (> 15–20 %) ultimately

Journal Pre-proof exhausts BMS from the mantle source, at which point some platinum-group elements (PGE) partition into rare, residual alloy phases (Kiseeva et al., 2017; Lorand and Luguet, 2016; O’Driscoll and GonzálezJiménez, 2016). Most other ore-forming elements, including a suite of trace to ultratrace elements included as part of the current study (Cu, Zn, As, Se, Mo, Ag, Cd, In, Sn, Sb, Te, W, Pb, Bi), are generally considered to be highly incompatible during mantle melting after sulphide exhaustion. The base- and precious-metal poor character of mantle residues has led to the general assumption

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that melt-depleted lithospheric mantle represents a poor source region for magmas capable of generating orebodies in the crust (Arndt, 2013; Barnes and Fiorentini, 2012). Instead, ore-bearing magmas are likely

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derived from relatively fertile, sub-lithospheric source regions that have not experienced prior melt

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extraction (Arndt, 2013). However, many melt-depleted peridotite samples yield base- and precious-metal

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concentrations that are unexpectedly fertile (Aulbach et al., 2016; Barnes et al., 2015; Lorand and Luguet, 2016; Luguet and Pearson, 2019), and exceed the composition of the PM in some cases. These unusual

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peridotite trace element compositions are opposite to the expected behaviour of ore-forming elements

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during partial melting. For melt-depleted mantle peridotite samples that yield anomalous concentrations of ore-forming elements, an unusual starting composition and/or the re-fertilization of base- and precious-

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2003).

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metals to a S-depleted source is probably required (Aldanmaz et al., 2012; Edwards, 1990; Lorand et al.,

The fate of most ore-forming elements during partial melting, as well as the impact from the refertilization process, remains poorly understood due primarily to the scarcity of accurate and precise data on these elements in mantle rocks. To address this challenge we utilized direct analysis of nanopowders and olivine grains with laser ablation inductively coupled plasma mass spectrometry (LA-ICPMS). The nanopowder technique is particularly well suited for whole-rock geochemical analyses because it avoids dissolution problems with digestion-resistant mineral phases. The low analytical blanks and excellent sensitivity of ICPMS also make the proposed approach well-suited for the detection of trace to ultratrace elements (Garbe-Schönberg and Müller, 2014; Peters and Pettke, 2017). Herein we present new nanopowder pressed pellet (NPPP) results and compare them with standard, laboratory, fused-disc

Journal Pre-proof ICPMS lithogeochemistry and isotopic analyses (Re-Os and Pb-Pb) for the same sample suite. Wholerock geochemical results are combined with in situ olivine analyses to document the incompatible majorand trace-element contribution of the most abundant residual phase in mantle peridotites (i.e., ≥ 60% olivine). Together these results allow us to critically compare the contributions of residual versus refertilized base- and precious-metals for our mantle peridotite and pyroxenite suite.

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2 Regional Geology The Cache Creek terrane comprises Mississippian to Lower Jurassic oceanic rocks that were

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accreted to the North American margin during the Mesozoic (Monger et al., 1972; Orchard et al., 2001).

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Mantle peridotite and crustal mafic rocks in the northern Cache Creek terrane around the town of Atlin

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have been interpreted as fragments of tectonically dismembered ophiolite (i.e., Nahlin ophiolite; Fig. 1)(Ash and Arksey, 1989; English and Johnston, 2005; English et al., 2010; McGoldrick et al., 2018,

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2017; Mihalynuk et al., 1994; Terry, 1977; Zagorevski et al., 2015). However, the classic Penrose-style

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ophiolite stratigraphy is generally not observed (Zagorevski et al., 2015). Instead, ophiolitic basalts and hypabyssal rocks around the town of Atlin are structurally juxtaposed directly with the underlying mantle

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peridotite massifs, with only rare exposures of middle and lower crust. These relationships suggest that

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the Nahlin ophiolite stratigraphy is controlled by intra-oceanic detachment faults and that much of it formed in a Late Permian to Middle Triassic intra-oceanic core complex in a suprasubduction zone setting (English et al., 2010; McGoldrick et al., 2018, 2017; Zagorevski et al., 2015). The Nahlin ophiolite was obducted onto carbonate platform sequences that are intercalated with Paleozoic alkaline basalt to rhyolite (McGoldrick et al., 2017). The obduction suture was reactivated and imbricated in the early Mesozoic and was stitched by Jurassic granitoids at 183–172 Ma (e.g., Three Sisters Plutonic Suite)(Mihalynuk et al., 1992).

3 Local Geology

Journal Pre-proof Five ultramafic massifs were selected and sampled during regional geological mapping around the town of Atlin, British Columbia (Fig. 1; Monarch Mountain, Peridotite Peak, Hardluck Peak, Nahlin Mountain and Menatatuline Range). Each of the studied massifs are dominated by harzburgite (≤ 5% clinopyroxene) with lesser dunite, pyroxenite and rare lherzolite (Fig. 2). Isotropic to foliated harzburgite samples comprise coarse-grained orthopyroxene porphyroclasts (enstatite; 10–30% estimated modal abundance) set in a finer grained matrix of olivine (60–90%) and lesser Cr-spinel (typically ≤ 1% modal

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abundance). Most of the harzburgite samples analyzed represent the least-serpentinized material available at each location (i.e., fresh olivine and orthopyroxene preserved, uncarbonated). However, a few samples

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that are more strongly serpentinized (antigorite ± bastite)(Hansen et al., 2005) were selected to assess

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element mobility during low-T alteration (discussed below).

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Dunite commonly occurs as bands and pods intercalated and transposed with foliated harzburgite and pyroxenite (Fig. 2c-d). Irregular dunite bands and bodies of variable size (centimetre to 100s m-scale)

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also commonly cut the mantle section. Dunite contains conspicuous spinel-rich laminae and

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disseminations that are notably absent in harzburgite. Contact relationships between the harzburgite and dunite bands vary from sharp to gradational and are defined by the varying modal proportions of coarse

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orthopyroxene. Dunite margins locally contain thick (10s of cm to m scale) orthopyroxene-rich (≥ 20%

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modal abundance) olivine websterite or olivine orthopyroxenite domains that are distinct from the straight-walled, orthopyroxene-rich dykes (see following). The layered appearance of alternating harzburgite, pyroxenite and dunite have been interpreted elsewhere as the products of melt:rock reaction (Bodinier and Godard, 2014; Edwards, 1995; Kelemen et al., 1995). Based on these field relationships, it is unlikely that Nahlin ophiolite peridotites represent simple melt residues. Pyroxenite dyke samples include orthopyroxenite, olivine orthopyroxenite, olivine websterite and websterite. All dykes are coarse grained and the modal abundance of clinopyroxene (0–90%) varies along strike in some individual dykes (Fig. 2h). Orthopyroxenite dykes are locally folded and transposed into banded harzburgite by a ductile, high-T fabric, suggesting that dyke emplacement and transposition occurred in the mantle (Fig. 2a). Other dykes cut the main foliation and are either weakly folded to not

Journal Pre-proof folded, indicating emplacement occurred after development of the high-T mantle fabric. Similar complex relationships are observed in dunites, which are both transposed into, and cut, banded peridotites. Mantle peridotite massifs are mostly devoid of clinopyroxene-bearing lithologies. Rare websterite dykes that show modal banding of clinopyroxene and clinopyroxenite pegmatites occur rarely within and adjacent to orthopyroxenite dykes (Fig. 2g–h). Isotropic lherzolite with patches of coarse Cr-diopside (20% modal abundance) was only observed at a few, isolated outcrops on Peridotite Peak (Fig. 1). These

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clinopyroxene-rich mantle domains could represent relatively undepleted mantle and/or a metasomatism imprinted on melt-depleted harzburgite (Bodinier and Godard, 2014; Griffin et al., 2009; Zhou et al.,

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2005). The interstitial habit of clinopyroxene and its association with Cr-spinel and BMS (pentlandite ±

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pyrrhotite ± chalcopyrite) in these samples at the microscale strongly suggest that at least some of these

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phases were introduced during mantle metasomatism (Fig. 3)(Edwards, 1990; Lorand, 1987; Luguet et al.,

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2004). Petrographic descriptions of the BMS assemblage are discussed in detail in Lawley et al. (2019).

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4 Methods 4.1 Whole-rock lithogeochemistry

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Whole rock samples were submitted to Activation Laboratories (Actlabs; Ancaster, Ontario) for

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conventional lithogeochemical analyses. Samples were crushed in a mild steel jaw crusher prior to pulverization in an agate ball mill. Agate was used to minimize the possibility of contamination of PGE and other trace elements during milling. Powdered samples were analyzed using the multi-element “4Lithores” analytical package, which includes ICPMS analysis of nitric acid-digested lithium metaborate/tetraborate fused glasses. Titration was used to measure FeO versus total Fe. Loss on ignition (LOI) was measured at both 500°C (LOI_500) and 1000°C (measured twice; LOI_1000 and LOI2_1000). Measuring LOI at multiple temperatures provides a more complete assessment of the organic, carbonate and volatile content of the sample suite. In principal, LOI can be used to assess the amount of structurally bound water in serpentinized samples, particularly at the higher ignition temperatures (discussed below).

Journal Pre-proof Total C and S were measured separately by combustion in an oxygen environment using infrared spectroscopy. Precious metals (Au, Pt, and Pd) were determined by fire-assay ICP-MS. Blind powdered reference materials (harzburgite MUH-1; komatiite OKUM; altered peridotite WPR-1a) were incorporated with the submitted sample suite (Supplementary Material Table 1). Based on replicate analysis of the blind reference materials, element concentrations that are routinely above the analytical detection limit are typically within 5% of their accepted values. Duplicate analyses for the three

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blind reference materials are on average within 2% and 4% for major and trace elements routinely above the detection limit, respectively. In addition, multiple reference materials, split duplicates and blanks were

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analyzed by Actlabs as part of the current study (Supplementary Material Table 1). However, we note that

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the vast majority of trace elements in Nahlin ophiolite samples are below the analytical detection limit of

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the 4Lithores Actlabs package (i.e., 57% of elements yield concentrations that were below the analytical detection limit for more than 50% of samples; Supplementary Material Table 1). The low abundance of

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many trace elements within highly melt-depleted mantle peridotite represents a major analytical

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challenge. An alternative approach with improved detection limits for the trace to ultratrace element

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analyses of mantle peridotite samples is described below (Supplementary Material Fig. 1).

4.2 Nanopowder LA-ICP-MS analysis

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Rock powders prepared for lithogeochemistry are typically milled to a median grain size of ≤ 50 µm (D50)(Piercey and Devine, 2014). These conventional powders can be pressed into pellet form and then analyzed directly be LA-ICPMS. However, the relatively coarse grain size of conventional pressed powder pellets can lead to poor analytical results, even using relatively large LA-ICPMS spot sizes (≥ 100 µm)(Garbe-Schönberg and Müller, 2014; Peters and Pettke, 2017). Re-milling these samples as a slurry with a high-speed, planetary ball mills can further reduce conventional rock powders to nanoparticulate grain sizes (D50 ≤ 1 µm) (Supplementary Material Figs. 2–3). Nanoparticulate powders are significantly more homogeneous and, when pressed into pellet form and analyzed directly by LA-ICPMS, can produce analytical results that are comparable with glass reference materials (Garbe-Schönberg and Müller, 2014;

Journal Pre-proof Peters and Pettke, 2017) (Supplementary Material Figs. 4–5). The homogeneity of these powders, coupled with low analytical blanks and the instrument sensitivity of ICPMS, make this an emerging approach for whole-rock analysis of elements at trace (low ppm) to ultratrace (low ppb) concentrations (GarbeSchönberg and Müller, 2014; Peters et al., 2017; Peters and Pettke, 2017). Here we adapt the method described by Garbe-Schönberg and Müller (2014) and Peters and Pettke (2017) using relatively small 15 ml agate milling bowls. Our miniaturized method (i.e., small mill bowls and 5 mm pressed pellets) was

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modified for the analysis of relatively small sample volumes (1–2 g). Recently, Kroner (2019) used the approach documented herein to report results for a suite of small, mafic volcanic-hosted mantle xenoliths

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from British Columbia.

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Peridotite and pyroxenite samples included as part of the current study were trimmed to remove

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weathered surfaces and thinly slabbed (1–5 mm) prior to milling. Saw marks were removed using silica carbide paper before cleaning loose rock powder from sample surfaces and/or within the pore spaces by

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ultra-sonic cleaning each mm-thin rock slab in ultra-pure water for 5–10 min. Samples were then dried

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and crushed in plastic bags with a plastic-covered hammer until almost all of the material passed through a 500 µm mesh disposable, nylon sieve. Coarse rock powders were then milled by hand in an agate

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mortar and pestle for five minutes.

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Two aliquots of the hand-milled rock powders were used for isotope analyses (discussed below). A third aliquot (1–2 grams) of each rock powder was wet-milled in ultrapure water for ~2 hours in an agate planetary ball mill (Fritsch Planetary Micro Mill Pulverisette 7). Wet-milling used 15 ml agate bowls and eight 5 mm agate balls with a fixed water:rock mass ratio (~3). Between samples, agate mills and balls were cleaned in an ultra-sonic bath with ultra-pure water, milled for ~10 minutes using an aliquot of each sample powder and bathed again with ultra-pure water. Laser diffraction particle size analysis (Beckman Coulter LS 13 320, and Fraunhofer and Mie models were used for coarse and fine fractions, respectively) completed over the course of this study suggest that rock powders following our wet-milling approach yield a median grain size of 1–2 µm (Supplementary Material Figs. 2–3). However, field-emission scanning electron microscope imaging

Journal Pre-proof (Tescan Mira-3; FE-SEM) suggest that laser-diffraction results overestimates the true median grain size due to re-aggregation of nanoparticulate powders during analysis (i.e., FE-SEM-based grain size estimates suggest that rock powders are 2–3x finer than laser diffraction results). Based on FE-SEM imaging and the grain size analysis using the “Feature” function of the Oxford Instruments Aztec software, the true median grain size following our wet-milling approach is likely ≤ 1 µm. Each nanoparticulate slurry was then freeze-dried for 48 hours, mixed with microcrystalline cellulose binder

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(Sigma-Aldrich; 20% binder by mass) and pressed into pellets (2 tons for 5 minutes) using a Specac 5 mm evacuable die.

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Pressed pellets were loaded into custom Teflon samples holders and analyzed directly with LA-

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ICPMS (Agilent 7700x ICP-MS coupled to a Photon Machines Analyte G2 193-nm excimer laser

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ablation system) at the Geological Survey of Canada, Ottawa. Data were acquired using a fast, timeresolved acquisition protocol (peak hopping, one point per peak; total cycle time for all masses was 527

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ms) and fixed laser-operating conditions during standard-sample bracketing. Samples were ablated using

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a 110 µm laser spot and a fluence of 5–6 J/cm2 at 10 Hz. The ablation aerosol was transported to the ICPMS using 1 l/min He (MFC-1: 0.6 l/min; MFC-2: 0.4 l/min) and was mixed with approximately 1 l/min

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Ar. Analyses consisted of 40 s of background measurement prior to 80 s of ablation, and ~50 s of washout between samples. The instrument was tuned on NIST-612 to achieve >9,000 cps/ppm 175Lu (50 µm spot,

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~7 J/cm2 at 10 Hz), while minimizing the production of oxides (<0.25% for ThO+/Th+) and maintaining a U/Th signal intensity ratio of ~ 1.0. United States Geological Survey (USGS) doped synthetic basalt glass standard GSD-1G (Guillong et al., 2005) was used as a primary calibration standard. Reference concentrations for primary and secondary standards were from the online geological and environmental reference materials database (GeoReM)(Jochum et al., 2005). Calibration of Te was based on an internal working value (GSE-1G = 279 ppm) calculated from repeated measurement of NIST610. Measured whole-rock MgO concentrations (Supplementary Material Table 1) for each sample were used as the internal standard to correct for different ablation yield between the sample and calibration standard (Longerich et al., 1996). Background

Journal Pre-proof and signal intervals were selected for integration from the time-resolved LA-ICPMS spectra and conversion of these integrated signals to element concentrations was performed in the Glitter software package (Griffin et al., 2008). Reported results represent the median of five replicate analyses for each sample. Three glass reference materials (BCR-2G, USGS; GOR-132, MPI-DING; GSC-1G, USGS) were mounted with unknowns and used to monitor instrument performance over the course of analyses.

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Reference materials (basalt BCR-2, USGS; komatiite OKUM, IAGEO; and harzburgite MUH-1, IAGEO) were also wet-milled to nanopowders, pressed and analyzed with unknowns to compare our NPPP

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approach against glass reference materials during the same analytical session. Replicate analysis of

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harzburgite MUH-1, which represents the closest matrix-match to the unknowns, suggest that

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nanopowder analyses are within 10% of accepted values for the most abundant major and trace elements. Measured SiO2 concentrations for MUH-1 are within 2% of their certified value if re-calculated to an

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anhydrous basis, suggesting that long-mill times and agate abrasion during milling did not impact the

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analytical results. Measurement trueness for trace to ultratrace elements are significantly more variable (Supplementary Material Fig. 4), but are generally within 20% of their target values (e.g., including REE

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concentrations at ~10 ppb in MUH-1). Several elements of interest in MUH-1 do not have certified values

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(i.e., Se, Zr, Mo, Ag, Cd, In, Te, Sn, Te, Ta, W, and Bi), but these elements yield excellent measurement trueness to the certified values of basaltic glass GSC-1G (within 1–20% of target values). Our MUH-1 results are also comparable to concentrations reported by Peters and Pettke (2017) for the same elements and following a similar method; however our method used smaller agate mill bowls and different milling times. Repeatability (%RSD) is excellent for most major (< 5%) and trace elements (mostly < 10%) in MUH-1 (Supplementary Material Fig. 5). Concentrations for REE approach the analytical detection limit of the nanopowder method (i.e., low ppb to ppt) and are subject to large analytical uncertainties, but are promising because they are essentially identical to previously published, high-precision acid-bomb digestion results for other Nahlin ophiolite samples (Babechuk et al., 2010; Canil et al., 2006). The agreement between the nanopowder

Journal Pre-proof REE results and the Babechuk et al. (2010) data highlight the low analytical blanks and excellent sensitivity of the new method, which is important for analysing a suite of trace- to ultratrace-elements within refractory peridotite samples (Fig. 4). For this study, only Tl concentrations were routinely below the analytical detection limit using the nanopowder method (Supplementary Material Fig. 1). Methodology, all analytical results and quality control data are reported in Supplementary Material Table

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2 and 3.

4.3 Whole-rock Re-Os + PGE analyses

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Platinum group elements and Re were analysed using isotope dilution techniques at the Arctic

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Resources Geochemistry Laboratory (ARGL) at the University of Alberta. Approximately 1 g of whole

and

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rock powder and an appropriate amount of a spike, isotopically enriched in 99Ru, 106Pd, 185Re, 190Os, 191Ir, 194

Pt, were weighed precisely and added to 30 ml quartz glass vials. Inverse aqua regia composed of

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~2 ml concentrated HCl (10.6 M) and ~5 ml concentrated HNO3 (15.4 M) was added to the vials. The

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vials were closed with a semi-permeable Teflon seal and heated to 260 °C at a containing pressure of ~130 bar for 16–24 hours in a high pressure asher (HPA), to digest PGE bearing phases and equilibrate

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the spikes and samples. Osmium was separated using CHCl3 solvent extraction, back-extracted into HBr

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(Cohen and Waters, 1996), and purified by micro-distillation (Birck et al., 1997). Following Os extraction, the aqua regia was dried before converting Re and the other PGEs to chloride form by drying repeatedly in HCl. Matrix separation was achieved using either anion exchange chromatography modified from Pearson & Woodland (2000), or cation exchange chromatography modified from Li et al. (2013). Long term repeat analyses (n = 43) of the OKUM standard shows that both methods produce identical ReOs and PGE data, within analytical uncertainty (P. Waterton, unpublished data). Osmium isotopes and abundances were measured using negative thermal ionisation mass spectrometry (N-TIMS), on a Thermo Fisher Triton Plus at the ARGL. Measurements were undertaken using Faraday Cups (FC) equipped with 1012 Ω resistors where sufficient Os was present (Liu and Pearson, 2014). Low Os samples were measured by peak hopping on a secondary electron multiplier

Journal Pre-proof (SEM). Mass fractionation was corrected to

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Os/188Os = 3.082614. Accuracy and precision of the Os

isotope analyses was assessed through repeated measurements of the DrOsS standard, which yielded Os/188Os = 0.16096 ± 0.00013 (2σ, n = 11) for FC analyses of 500 pg – 2.5 ng standards, and

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Os/188Os = 0.16080 ± 0.00036 (2σ, n = 11) for SEM analyses, undertaken during the period of this

study. The mean

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Os/188Os for both FC and SEM DrOsS standards overlap with the accepted value of

0.160924 ± 0.00004 (Luguet et al., 2008) within 2 standard deviations, hence the Os isotopic data are

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therefore considered accurate at this level of precision. PGE and Re abundances were measured on a Nu Attom ICP-MS at the ARGL. Mass

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fractionation was corrected externally using synthetic 1 ppb PGE standards. The mass fractionation

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corrected measured isotopic composition of these standards were repeatable to ± 0.4% for Re, ± 0.6% for

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Ru, and ± 0.8% for Pd, Ir and Pt (2σ relative). These values represent the maximum precision on an individual isotopic measurement. Three total procedural blanks were analysed with the samples, yielding

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average blanks of: Os 8.8 ± 23.9 pg, Ir <8.0 pg, Ru <8.0 pg, Pt 21.2 ± 36.0 pg, Pd <11.3 pg, and Re 2.3 ±

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2.6 pg (2σ uncertainties, maximum values given where one or more blanks were below limit of detection). Accuracy and repeatability of the PGE abundance measurements (including Os) was

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monitored through repeated analyses of the OKUM komatiite reference material. Average values of Os 0.80 ± 0.07 ppb (18%), Ir 0.93 ± 0.07 ppb (14%), Ru 4.43 ± 0.14 ppb (6%), Pt 11.6 ± 0.6 (5%), Pd 11.7 ±

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0.6 (5%), and Re 0.51 ± 0.05 (10%; 2σ uncertainties, percentages in brackets indicate relative 2σ as an indication of repeatability) overlap previously published data for OKUM (Savard et al., 2010), so these data are considered accurate at this level of precision. The repeatabilities are generally poor relative to expected analytical precision. This discrepancy is attributed to nugget effects in the ~1 g samples and/or slight variations in degree of spike-sample equilibration during HPA digestion. All analytical results and quality control data are reported Supplementary Material Table 4.

4.4 Whole-rock Pb-Pb isotope analyses

Journal Pre-proof Whole rock Pb-Pb isotope analyses were completed at the Isotope Geochemistry and Geochronology Research center (IGGRC; Department of Earth sciences, Carleton University) following Cousens (1996). Rock powders were dissolved in a mixture concentrated HF and HNO3. The residue was then re-dissolved in a mixture of HNO3 and HCl sequentially prior to chemical separation. Isotope ratios were measured on a Thermal Ionization Mass Spectrometry (TIMS; ThermoFinnigan Triton) and are corrected for fractionation using the NBS 981 standard values of Todt et al. (1996). For Pb-poor peridotite

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and pyroxenite samples, isotope ratios were acquired using the secondary electron multiplier and are associated with relatively large analytical uncertainties (highlighted in Supplementary Material Table 5).

Pb/204Pb = 15.613 ± 0.020;

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Pb/204Pb = 38.693 ± 0.068;

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Pb/206Pb = 2.0642 ± 0.0020;

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Pb/206Pb =

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Replicate analyses of BCR-2 (n = 7) yield Pb isotope ratios (206Pb/204Pb = 18.744 ± 0.017;

18.754 ± 0.009;

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0.8329 ± 0.0006) that are in good agreement with their preferred values from GeoReM ( 206Pb/204Pb = Pb/204Pb = 15.622 ± 0.005;

208

Pb/204Pb = 38.726 ± 0.022;

208

Pb/206Pb = 2.064 ±

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0.001)(Jochum et al., 2005). The total ranges measured for NBS 981 in a 2 year period bracketing the analyses are (2σ errors): ± 0.017 for 206Pb/204Pb; ± 0.021 for 207Pb/204Pb; ± 0.038 for 208Pb/204Pb; ± 0.0021 208

Pb/206Pb; ± 0.00038 for

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Pb/206Pb. The total procedure blanks were less than 50 picograms. All

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for

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analytical results are reported in Supplementary Material Table 5.

4.5 Olivine EPMA and LA-ICP-MS analyses EPMA analysis was completed on polished olivine epoxy mounts at the University of Alberta using a JEOL JXA-8900R equipped with 5 tunable wavelength dispersive spectrometers. Analyses were completed using a 2 µm beam diameter and a beam energy and current of 20 keV and 20 nA, respectively. Reported signal intensities and concentrations are corrected for dead time, drift, element interference (Donovan et al., 1993) and matrix-effects using a combination of ZAF and PhiRho-Z algorithms (Armstrong, 1988). Most elements were acquired with peak measurement times of 30 s. Each analysis is reported as the average of 3 points. Complete EPMA methodology is presented in Lawley et al. (2018).

Journal Pre-proof Olivine LA-ICP-MS spot analyses (110 µm) were completed at the Geological Survey of Canada, Ottawa using the same instrument set-up as the nanopowder analyses (discussed above) and the methods described in Lawley et al. (2018). Primary calibration standards included the USGS doped synthetic basalt glass standard GSD-1G (Guillong et al., 2005) and the doped synthetic pyrrhotite standard Po726 (prepared by L. Cabri and G. Laflamme at CANMET)(Sylvester et al., 2005). Reference concentrations for these primary and secondary standards are the preferred values taken from GeoReM (Jochum et al.,

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2005) and the Po726 certificate, respectively. For the present study, Te calibration was based on an internal working value (GSE-1G = 279 ppm) calculated from repeated measurement of NIST610. Internal

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standards (i.e., MgO and FeO concentrations from EPMA results for the same spot location) were used to

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correct for different ablation yield between the sample and calibration standard (Longerich et al., 1996).

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Signal integration of the time-resolved LA-ICP-MS spectra and conversion of these integrated signals to element concentrations was completed in the Glitter software package (Griffin et al., 2008).

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Basalt glass (BCR-2G) and a fragment of a well-characterized San Carlos olivine grain

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(Bussweiler et al., 2019, 2017; Lawley et al., 2018) were used to monitor instrument performance. Replicate analysis of BCR-2G suggest that concentrations are generally within 1–10% of their accepted

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values. Measurement repeatability calculated for BCR-2G is typically < 5% RSD for most major and

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trace elements, but increases to ≥ 20% RSD for ultratrace elements at low ppb concentrations. All analytical results and quality control standard data are reported in Supplementary Material Table 6.

5 Results and interpretation 5.1 Whole-rock major and trace element results Lithogeochemistry (Actlabs) results for 66 samples are presented in Supplementary Material Table 1. Nanopowder results for a subset of 45 samples are presented as Supplementary Material Table 3. Elements that yield concentrations routinely above the analytical detection limit and were determined by both methods are in good agreement (e.g., median difference of 9% and 10% for Si and Fe, respectively), suggesting that the nanopowder calibration and methodology is appropriate and that different rock

Journal Pre-proof powders prepared for each method are homogeneous. Three samples yielded contrasting results between Actlabs and nanopowder methods for some elements (i.e., 16LVA-A009-A01 and 16LVA-A018-A01, 17LVA-A011-B01) and are excluded from further discussion. Nahlin ophiolite harzburgite samples yield a narrow range of Al2O3 (0.25–1.62 wt.%; mean = 0.81 ± 0.13 wt.% 2σ; n = 25), which, coupled with REE-poor compositions (typically 0.1 × PM), is consistent with these ultramafic massif samples being residues after significant degrees of melt extraction

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(Fig. 4–7). The most melt-depleted harzburgite samples occur at Hardluck Peak, Nahlin Mountain, Monarch Mountain and Menatatuline Range (mean Al2O3 0.57 wt.%; n = 2; 0.65 wt.%, n = 9; 0.76 wt.%,

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n = 4; 0.79 wt.%, n = 5 respectively), whereas the few samples from Peridotite Peak (1.44 wt.%; Al2O3; n

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= 3) and the town of Atlin (1.05 wt.% Al2O3; n = 2) are relatively fertile.

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Serpentinized harzburgite samples are visually distinct in the field because olivine and orthopyroxene are completely serpentinized, and yield high loss on ignition (LOI) (1.90–10.75 %; mean =

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4.65 ± 2.67 % 2σ; n = 6). However, least-serpentinized harzburgite samples with fresh olivine yield LOI

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that overlap with the highly serpentinized samples (0.57–13.83%; mean = 6.89 ± 1.27 %; n = 25). The variable, but elevated LOI indicates that most Nahlin ophiolite samples contain a significant fraction of

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samples.

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structurally bound water due to the replacement of olivine by serpentine, even in the least-serpentinized

Dunite samples yield lower Al2O3 concentrations (0.10–0.57 wt.%; mean = 0.23 ± 0.12 wt.%, n = 8), and scatter to higher LOI estimates (6.87–13.88 %; mean = 10.11 ± 1.58 %; n = 8) relative to the analyzed harzburgite samples. The major element differences between harzburgite and dunite samples are consistent with the decrease in modal abundance of clinopyroxene and orthopyroxene during increased partial melting. However, most dunite samples reported herein are interpreted as melt metasomatic, rather than representing the most refractory samples in our peridotite suite (Figs. 4–7). The petrogenesis of metasomatic dunite and pyroxenite samples are discussed below. Olivine websterite and orthopyroxenite samples yield compositions that generally overlap with harzburgite, except for one extremely Al2O3-poor (0.07 wt.%) olivine websterite sample from Peridotite

Journal Pre-proof Peak (17LVA-A032-A01; Figs. 4–7). Websterite dykes are Ca-rich (e.g., CaO mean = 9.46 ± 6.27 wt.%; 1.25–16.09 wt.%; n = 5), which contrast with the Ca-poor composition of orthopyroxenite dykes (e.g., CaO mean = 2.09 ± 0.53 wt.%; 1.10–5.46 wt.%; n = 15). Evolved webserite dykes are also notably LREE enriched compared to the other peridotite and pyroxenite samples.

5.2 Whole-rock Re-Os isotope and PGE results

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PGE concentrations and Re-Os isotope results for 45 ultramafic rocks (n.b. not including duplicate analyses) are reported in Supplementary Material Table 4. New isotope dilution (ID)-based PGE

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results yield somewhat scattered concentrations and inter-element relationships (Fig. 8). Replicate

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analysis of the OKUM komatiite reference material (n = 7) suggests that PGE heterogeneity is not an

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analytical artifact. Instead, the PGE heterogeneity is interpreted to reflect a strong nugget effect, which is consistent with ultrafine and heterogeneously distributed alloys for such melt-depleted samples. The

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presence of PGM is further supported by comparisons between the new ID PGE results with previously

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published NiS-FA data for the same sample suite (Lawley et al., 2019) and replicate ID PGE analyses for a subset of the Nahlin ophiolite suite (Supplementary Material Table 4; Supplementary Material Fig. 6). 187

Os/188Os ratios (0.1143–0.1715; n = 44) that do not

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Peridotite analyses yield a wide range of

correlate with melt indices, Re/Os ratios or absolute Os abundances. This range of 187Os/188Os is similar to

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olivine websterite to olivine orthopyroxenite bands intercalated with harzburgite and dunite ( 187Os/188Os ratios = 0.1278–0.1248; n = 5). Orthopyroxenite and websterite dykes scatter to more radiogenic 187

Os/188Os ratios (0.1300–0.5745; n = 9), which, for some samples, exceed the isotopic composition of

present-day chondrite (Shirey and Walker, 1998). Radiogenic dyke samples with suprachondritic 187

Os/188Os ratios yield meaningless future Re-depletion dates (TRD) that are not considered further. Peridotite and pyroxenite analyses with subchondritic 187Os/188Os ratios (n = 35) yield a range of

TRD ages from 63 to 1961 Ma, but with a cluster of model ages at ca. 250 Ma and a second prominent mode at ca. 500 Ma (Fig. 9). The youngest cluster of TRD ages at ca. 250 Ma represents an imprecise, minimum estimate for the timing of melting for these strongly melt-depleted harzburgite samples.

Journal Pre-proof However, the second TRD age mode is older than the Late Permian to Middle Triassic age of the Nahlin ophiolite, which, coupled with anomalous Paleozoic to Paleoproterozoic TRD ages, likely point to ancient Re-depletion events that significantly pre-date ca. 250 Ma ophiolite formation. Available Re-Os model ages (TMA) for a subset of the sample suite are displaced to older ages (i.e., primary mode at ca. 500 Ma; 231–2377 Ma; n = 29). These TMa ages are unlikely to reflect the actual age of past melting events due to recent Re mobility and/or the introduction of variable Os initial isotope signatures during melt:rock

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interaction with pyroxenite.

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5.3 Whole-rock Pb-Pb isotope results

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Isotope results for 45 peridotite samples are reported in Supplementary Material Table 5 and

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presented in Fig. 10. Samples yield a range of Pb isotope ratios (206Pb/204Pb = 15.402–19.050; 207Pb/204Pb = 15.127–15.633; 208Pb/204Pb = 34.980–38.434) and there is considerable overlap between rock types and

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Pb/204Pb = 15.488–15.563), which, along with a subset of harzburgite samples from

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18.648–19.050;

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sample localities (Fig. 10). Websterite dykes yield the most radiogenic Pb isotope ratios (206Pb/204Pb =

Monarch Mountain and Atlin town, plot to the right of the geochron (Fig. 10). Samples with the highest Pb/204Pb and 207Pb/204Pb ratios most likely reflect in-situ growth of radiogenic Pb.

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The vast majority of the remaining samples, including orthopyroxenite dykes, yield low Pb/204Pb and 207Pb/204Pb ratios that are significantly less radiogenic than the composition of the mantle

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206

at ca. 250 Ma (Fig. 10). Because orthopyroxenite dykes cut harzburgite, these relatively unradiogenic Pb isotopic compositions were likely inherited from the previously melt-depleted peridotite and are thus unlikely to reflect a single melting event. The scatter of Pb isotope compositions is also inconsistent with a single secondary isochron and/or simple binary mixing between U and Pb reservoirs. Instead, the Pb isotope results likely point to multiple mantle melting and re-fertilisation events that significantly predate the Nahlin ophiolite. Each event may have fractionated U and Pb to produce a range of initial Pb isotope compositions, which then, in turn, evolved along separate U/Pb trajectories through time (i.e., multiple secondary isochrons; Fig. 10).

Journal Pre-proof The least-radiogenic samples yield

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Pb/204Pb and

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Pb/204Pb ratios that are similar to the

composition of the mantle during the Paleoproterozoic (Stacey and Kramers, 1975). Because U and Pb are relatively low abundance elements, the Pb isotopic system is susceptible to disturbance and/or overprinting during low-T serpentinization. However, serpentinization is unlikely to have produced these Pb isotope signatures, since the two strongly serpentinized samples from the Nahlin Mountain and Menatatuline Range overlap with the Pb isotope composition of the least-serpentinized samples at the

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same localities.

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5.4 Olivine results

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In situ EPMA and LA-ICP-MS results for 460 olivine grains from 11 samples (seven harzburgite;

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one dunite; two olivine websterite; one orthopyroxenite dyke) are reported in Supplementary Material Table 6. Olivine analyses yield a range of Mg# (Mg/Mg+Fe*100; 90.28–93.27; mean = 91.31 ± 0.06 2σ;

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n = 460). Divalent cations with ionic radii similar to olivine’s major cations correlate with Mg# (e.g., Mn;

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Fig. 11) and yield the greatest concentrations, which likely reflects direct substitution involving its sixfold, octahedral M site (De Hoog et al., 2010). Ni is also correlated with Mg#; however, unusually Ni-rich

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inclusions.

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compositions in some olivine websterite and harzburgite (Fig. 11) may reflect ultrafine metasomatic BMS

The concentrations of most other elements do not correlate strongly with melt indices and instead point to some other control on olivine compositions. The strong co-variation between trivalent cations (Al + Cr ± V ± Sc), which partition into olivine by temperature-sensitive substitution reactions (De Hoog et al., 2010; Wan et al., 2008), suggest that equilibration temperature represents a secondary control on olivine chemistry for some elements (Fig. 11)(Lawley et al., 2018). Herein, olivine equilibration temperatures are estimated with the Al-in-olivine thermometer of De Hoog et al. (2010) at constant pressure (10 kbar). Although this thermometer was calibrated for garnet-facies harzburgite and lherzolite using the two-pyroxene and Al (or Ca)-in-orthopyroxene thermometer of Brey and Köhler (1990), results for spinel-facies peridotites are within 30°C of the target values for the calibration sample set (De Hoog et

Journal Pre-proof al., 2010). Results for the Ca-in-olivine thermometer (De Hoog et al., 2010) for our sample set are also in broad agreement with Al- and Cr-based olivine thermometry. Average olivine equilibration temperatures for harzburgite from each massif (Hardluck Peak = 718 ± 15 °C, 618–814 °C, n = 36; Menatatuline Range = 854 ± 12 °C, 785–996 °C, n = 36; Monarch Mountain = 680 ± 19 °C, 591–947 °C, n = 48; Nahlin Mountain = 815 ± 8 °C, 723–1107 °C, n = 174) overlap with previously published olivine-spinel thermometry at Menatatuline Range (mean = 814 ± 26

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°C; 675–970 °C; n = 42) and Hardluck/Peridotite Peak (mean = 756 ± 22 °C; 625–943 °C; n = 63) (McGoldrick et al., 2018). Olivine-spinel thermometry values likely yield minimum equilibration

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temperatures due to the rapid Fe and Mg diffusion between olivine and spinel and the resulting values are

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up to 250 °C lower than temperature estimates based on two-pyroxene thermometry (McGoldrick et al.,

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2018). Olivine thermometry results are also interpreted as metamorphic equilibration temperatures, but tend to suggest that melt-depleted harzburgite samples from each of the Nahlin ophiolite massifs also

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equilibrated at relatively higher temperatures (Fig. 11). McGoldrick et al. (2018) suggested that massifs

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with higher equilibration temperatures cooled faster and/or achieved higher peak temperatures during melting. The different cooling and/or melt-depletion histories of each massif may suggest that the Nahlin

2018).

5.5 Melt models

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ophiolite reflects differential exhumation of the upper mantle along detachment faults (McGoldrick et al.,

Melting curves highlight the compatible versus incompatible element behaviour determined by experimental petrology. Because the starting composition, melting mode and trace element partitioning remain poorly understood, melting curves are subject to large uncertainties (Pearce and Parkinson, 1993). Hence our relatively simple melt models are therefore intended as reference frames rather than an accurate representation of the complete Nahlin ophiolite melt history. Model reference curves are based on fractional melting, using the mineral modes for silicate/oxide and sulphide phases from Niu et al. (1997) and Lee et al. (2012), respectively. Mineral

Journal Pre-proof models were adjusted to simulate the removal of BMS during melting (Lee et al., 2012). To facilitate interpretation, Al2O3 was used as a proxy for the degree of melt extraction and is plotted at 5% increments (Figs. 6–7). The starting composition was assumed to be PM (Palme and O’Neill, 2014), although melt estimates based on some other starting composition, such as Depleted Mid-ocean-ridge-basalt Mantle (DMM), tend to yield lower estimates for the degree of partial melting at any given Al 2O3 (McGoldrick et al., 2017). Lower degrees of partial melting may also be required if the mantle source region was re-

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fertilized prior to second stage melting (Aldanmaz et al., 2012). Partition coefficients for trace to ultratrace elements were assigned a constant value and taken from Pearce and Parkinson (1993), Salter et

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al. (2002) and Kiseeva et al. (2017).

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Compatible (e.g., MgO, total Fe2O3, Ni, Co) to moderately incompatible element (e.g., MnO,

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Al2O3, CaO, Sc, V, Ga and HREE) concentrations for harzburgite, dunite and olivine websterite scatter around melt curves to a first order (i.e., concave-up versus concave-down, respectively)(Figs. 6–7).

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However, important mismatches between our results and predicted melting curves suggest that this

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simplified melt model does not explain our new data (Figs. 6–7). Lab results (both nanopowder and fusion ICPMS) are MgO- and Ni-poor and CaO-rich relative to melt curves (Fig. 6). Measured SiO2 and

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Cr concentrations also yield poor fits to melt models, but are not shown. The mismatch between melt

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models and peridotite compositions is most apparent for the highly incompatible elements (e.g., LREE, LILE and HFSE; Fig. 5). These elements are not known to substitute into any of the primary mantle phases other than clinopyroxene (Pearce and Parkinson, 1993; Salters et al., 2002). Our nanopowder results for melt-depleted harzburgite and dunite samples devoid of clinopyroxene are therefore not consistent with a single-stage melt extraction evolution. Pyroxenite dykes, which, in some cases are particularly clinopyroxene-rich (i.e., websterite), also yield incompatible element-rich compositions that do not plot on melt-deletion trends (Figs. 4-7).

6 Discussion 6.1 Melt depletion

Journal Pre-proof Samples of the ophiolitic mantle were originally interpreted as simple residues of partial melting with compositions that could be explained by variable degrees of melt extraction (Coleman, 1977; Ishiwatari, 1985; Prinzhofer and Allègre, 1985). The complementary history and genetic link between basaltic volcanism and its peridotitic residue is supported to a first-order by the composition and modal mineralogy of mafic to ultramafic assemblages in abyssal settings (Warren, 2016). However, multiple lines of evidence over the last few decades, including variations in the modal abundance of mineral

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phases that deviate from melt indices (Niu et al., 1997) and incompatible element signatures that are inconsistent with partial melting (Aldanmaz et al., 2012; Barth et al., 2008; Elthon, 1992; Hanghøj et al.,

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2010; Takazawa et al., 2000; Uysal et al., 2012), demonstrate that the geological history of most oceanic

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mantle peridotites cannot be attributed to a simple, single-stage partial melt model. A variety of melt-

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and/or fluid-rock interactions contribute to a broad array of metasomatic processes (e.g., modal, cryptic and stealth metasomatism) that can pre-date and/or overprint the partial melting history of mantle

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peridotite in oceanic and continental settings (Bailey, 1982; Griffin et al., 2009; Pearce et al., 2000).

challenge.

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Unravelling the multiple stages of melt-depletion and re-fertilization from the finite rock record remains a

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For Nahlin ophiolite peridotites, compatible and moderately incompatible element concentrations

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increase and/or decrease in accord with predicted melting trends, suggesting that melt extraction represents an important, first-order control on whole-rock compositions (Fig. 4–7). Although the melting mode and process (e.g., fractional versus batch melting), model parameters (e.g., partition coefficients) and starting composition remain poorly constrained, the new nanopowder data suggest that most harzburgite and some dunite samples are strongly melt-depleted (i.e., mostly residues after ≥ 20% partial melting; assuming a PM starting composition)(Palme and O’Neill, 2014). As described in McGoldrick et al. (2017), lower degrees of partial melting are required if the starting composition was more similar to DMM (i.e., 9–20% for Hardluck Peak, Peridotite Peak and Menatatuline Range). Some highly incompatible ultratrace element concentrations (ppb) are also strongly depleted relative to the PM, which is consistent with their expected behaviour during partial melting (Fig. 6–7).

Journal Pre-proof For example, smooth positive-sloped REE profiles highlight the greater incompatibility of LREE relative to MREE and HREE (Fig. 4). The LREE-poor profile of the Nahlin ophiolite samples are similar to meltdepleted harzburgite from the Oman ophiolite (Godard et al., 2000; Hanghøj et al., 2010) and some extremely melt-depleted examples of abyssal peridotites (Godard et al., 2008). Notwithstanding these atypical samples, the Nahlin ophiolite suite is significantly more melt-depleted than most mantle residues in abyssal settings (Bodinier and Godard, 2014), possibly due to extensive, hydrous melting in a supra-

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subduction zone environment (McGoldrick et al., 2017). The Nahlin ophiolite samples may also have experienced multiple stages of partial melting, which would significantly impact these melt estimates

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(Aldanmaz et al., 2012; Uysal et al., 2012). Nevertheless, the agreement between multiple melt indices is

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important because ultratrace elements, including REE, are extremely sensitive to re-fertilization by

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percolating melts.

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6.2 Re-fertilization

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Field relationships document varying degrees of melt:rock interaction between harzburgite, dunite and pyroxenite (Fig. 2), suggesting that Nahlin ophiolite peridotites are not simple residues after

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varying degrees of partial melting. For example, replacive dunite bands at the Hardluck Peak are

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intercalated and folded with orthopyroxene-rich harzburgite and pyroxenite (Fig. 2h). The close structural intercalation of these rocks suggest that at least some dunite bodies represent the product of percolating melts that crystallized olivine during melt:rock reactions (Kelemen et al., 1995). Olivine orthopyroxenite and olivine websterite domains in gradational contact with the margins of dunite bands are therefore interpreted as the product of Si-metasomatism superimposed on the melt-depleted harzburgite by passing silicate melts (Edwards, 1995; Varfalvy et al., 1997). Because these metasomatic domains are locally folded, early-stage(s) metasomatism must have occurred prior to the youngest high-T mantle fabric. In contrast, late orthopyroxenite dykes that cut foliated harzburgite and patches of coarse grained to pegmatitic clinopyroxene overprinting these dykes appear to document multiple stages of metasomatism that both pre- and post-date the high-T mantle fabric.

Journal Pre-proof Orthopyroxenite dykes are depleted in most trace elements and yield compositions that are remarkably similar to the most melt-depleted harzburgite residues (Fig. 4–7). The compositional similarities between orthopyroxenite dykes and residual harzburgite domains may reflect the similar element partitioning behaviour for these olivine- and orthopyroxene-dominated rock types. Similar lowAl pyroxenite dykes have been interpreted elsewhere as segregations from subduction-related boninitc melts (Edwards, 1995; Varfalvy et al., 1997) and are comparable to the low-Al pyroxenite suite from

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other ophiolites globally (Bodinier and Godard, 2014). In contrast, websterite dykes are the most trace element-rich rock type among the Nahlin suite mantle and are interpreted as the crystallization product of

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relatively fractionated melts and/or assimilation of clinopyroxene during melt:rock interaction (Varfalvy

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et al., 1997).

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Both dyke types at the Nahlin ophiolite are distinct from Al-rich, amphibole- and/or garnetbearing pyroxenite dykes in other global ultramafic massifs, which are likely derived from deep mantle

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source regions and are thus unrelated to the host peridotite (Bodinier and Godard, 2014). If correct, the

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orthopyroxenite and websterite dykes of the Nahlin ophiolite are consistent with an arc-like setting for ophiolite formation. The genetic relationship between the different dyke types and their petrogenetic

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significance to the genesis of the Nahlin ophiolite requires further study.

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Because very few Nahlin ophiolite samples are actually as LREE-poor as would be expected from simple batch and/or fractional melt models, the trace element signature of most peridotites is interpreted to contain a small, but variable proportion of trapped melt (Godard et al., 2000). Melt-related re-fertilization may have also contributed to the modest enrichment of LILE, which is superimposed on previously melt-depleted peridotite (Figs. 6–7). The few harzburgite (16LVA-A028-A01), dunite (17LVA-A015-B01) and pyroxenite (17LVA-A032-A01) samples with concave-upwards REE profiles provide the clearest trace element evidence for the re-introduction of highly incompatible elements and are similar to melt-impregnated, plagioclase- and/or amphibole-bearing peridotites that yield pronounced U- or spoon-shaped LREE profiles as described in other ophiolites (Godard et al., 2000; Hanghøj et al., 2010). Websterite dykes are also enriched in these incompatible LREE.

Journal Pre-proof Re-fertilization of incompatible elements also has implications for ore-forming elements because intergranular BMS (pentlandite ± pyrrhotite ± chalcopyrite) is spatially associated with Cr-spinel and clinopyroxene at the microscale (Fig. 3). Based on REE profiles, McGoldrick et al. (2018) demonstrated that some clinopyroxene within the Nahlin ophiolite is metasomatic. If correct, metasomatic intergranular BMS in association with clinopyroxene likely represents the dominant mineral host for most base- and precious-metals and provide a possible explanation for the flat, PM-like to weakly suprachondritic PGE

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profiles of nearly all melt-depleted harzburgite (Fig. 8). However, BMS were not observed or recovered during mineral separation for the majority of samples, suggesting that metasomatic sulphide melt, if

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present, occurs as ultrafine inclusions, concentrated on mineral surfaces and/or are relatively rare (Lawley

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et al., 2019). Nahlin ophiolite harzburgite (± dunite) samples thus appear to offer an opportunity to

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investigate the selective enrichment of ore-forming elements and LILE in the absence of any significant

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trapped, LREE-bearing melt phase.

6.3 Olivine-hosted base and precious metals

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Olivine is the most abundant residual mineral phase in melt-depleted harzburgite (≥ 40%) and is

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the dominant mantle silicate in dunite (≥ 90%). Because of its high modal abundance, residual olivine is

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always a significant, and, in some cases, may be the primary host for trace elements in the most meltdepleted peridotites (De Hoog et al., 2010; Greaney and Rudnick, 2017; Witt-Eickschen et al., 2009a). Herein we quantify the importance of residual olivine as a host phase for incompatible elements within Nahlin ophiolite samples by comparing whole-rock nanopowder results to median olivine concentrations (Figs. 6–7). The mass balance assumes an olivine modal abundance of 70%, which, for most peridotite samples, means that the estimated olivine contribution to the whole rock budget will represent a minimum value. For elements routinely below the analytical detection limit, median olivine concentrations will overestimate its true contribution to the whole rock budget to an unknown extent. The complete mass balance reported for all elements, which is based on median whole-rock nanopowder and olivine LAICPMS results for dunite and harzburgite, is presented in Supplementary Material Fig. 7.

Journal Pre-proof Our results indicate that residual olivine does in fact represent a significant mineral host for a range of incompatible elements (Fig. 6–7). For base- and precious-metals, Cu and Zn are the only two elements known to substitute into olivine to any significant degree (De Hoog et al., 2010). Statistical analyses of the LA-ICPMS signal intensity data from each sweep of the mass spectrometer confirm that the effect of micron-sized inclusions on the Cu concentrations of most olivine is relatively minor, which, coupled with its temperature sensitive partitioning during sub-solidus equilibration (Fig. 11), strongly

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suggest that Cu substitutes into the olivine crystal lattice (De Hoog et al., 2010; Lawley et al., 2018). Olivine-hosted Cu is further supported by apparent, whole-rock melt depletion trends that are comparable

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to other moderately incompatible elements (e.g., Ga), suggesting that some base and precious-metals are

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hosted, at least in part, by trace element substitution into residual mantle phases in addition to re-

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fertilization by metasomatic BMS (Aldanmaz et al., 2012; Kiseeva et al., 2017). If correct, melt models that assume Cu is entirely controlled by BMS will systematically

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underestimate the true composition of melt-depleted mantle peridotite (Fellows and Canil, 2012). Based on the median olivine Cu concentrations from the present study, we suggest that even the most melt-

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depleted peridotite should contain ≥ 0.380 ppm Cu (Fig. 7). Mantle peridotites with lower Cu

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concentrations would require stripping of this base-metal from residual olivine. Whether such Cu-poor

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mantle peridotites are common in nature is unclear because Cu (and other ore-forming elements) concentrations are generally below the analytical detection limits of conventional analytical methods (Supplementary Material Fig. 1).

Compatible to moderately incompatible partial melting trends are also apparent for other oreforming elements including: As + Se + Mo + Ag + Sn + Te + W+ Pb + Bi. Sequestering of these oreforming elements by residual mantle phases could explain the unusual composition of melt-depleted peridotite samples and is partially supported by olivine mass balance estimates (Fig. 7; Supplementary Material Fig. 7)(Babechuk et al., 2010; Liang et al., 2017; Liu et al., 2018; Witt-Eickschen et al., 2009b). However, most ore-forming elements are not known to substitute into the olivine structure, which, coupled with anomalous whole-rock concentrations that exceed the composition of the PM for some

Journal Pre-proof samples, most likely point to metasomatic re-fertilization. A large proportion of olivine analyses also yield ore-forming element concentrations that are below the analytical detection limit, which obscures the true contribution of these residual phases to the whole-rock budget (% of analyses below the analytical detection limit: As = 36%; Se = 74%; Mo = 3%; Ag = 76%; Sn = 3%; Te = 74%; W = 54%; Pb = 7%; Bi = 59%), Since olivine analyses targeted pristine domains that were free of opaque inclusions (e.g.,

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chromite and/or sulphide), we interpret that the olivine-hosted ore-forming elements most likely reflect trapped melt, fluid and/or ultrafine BMS inclusions that were not apparent during petrography (Fig. 3). In

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situ LA-ICPMS analyses are relatively sensitive to these types of fluid and/or mineral inclusions, because

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the concentrations of ore-forming elements within olivine are generally very low. Where present, these

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inclusions may dominate the olivine mass balance estimates for these elements (Supplementary Material Fig. 7). The few available studies that report concentrations for ore-forming elements within other phases,

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such as orthopyroxene, Cr-spinel, and/or mineral phases at grain boundaries, suggest that contributions

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from rare BMS are required to balance the whole-rock budget in most cases (Babechuk et al., 2010; Greaney and Rudnick, 2017; Liang et al., 2017; Liu et al., 2018).

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The origin of this ore-forming element signature in Nahlin ophiolite samples is important because

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estimates for the composition of the bulk silicate Earth are based on the premise that these elements are highly incompatible and yield metal ratios that are uniform and unfractionated by melting processes (Palme and O’Neill, 2014; Sun and McDonough, 1989). We demonstrate that highly incompatible element ratios range over several orders of magnitude and extend to suprachondritic values for the Nahlin ophiolite suite (Fig. 12). Noll et al., (1996) suggested that suprachondritic ratios for highly incompatible elements are a feature of high-T fluid metasomatism in the mantle source regions of arc volcanism. Serpentinized mantle peridotites (i.e., low-T fluid metasomatism) are relatively enriched in a similar suite of incompatible elements and have been proposed as one of the possible sources of this arc signature (Kodolányi et al., 2012). We note that unserpentinized mantle peridotite xenoliths, which were prepared using the same nanopowders method (Kroner, 2019), mostly yield chondritic to subchondritic

Journal Pre-proof incompatible element ratios, suggesting that this signature does not represent an analytical artifact (Fig. 12). Instead, we suggest that the modest re-fertilization of ore-forming elements observed within the Nahlin ophiolite suite may reflect trapped fluids (high- and/or low-T fluid metasomatism), possibly in addition to the precipitation of high-T metasomatic BMS (Andreani et al., 2014; Deschamps et al., 2011; Hattori and Guillot, 2003; Peters et al., 2017). Trapped fluids are also the most likely explanation for the

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abundance of ore-forming elements within otherwise unaltered olivine mineral separates (Fig. 11).

6.4 Timing of melt-depletion and re-fertilization

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Radiogenic isotopes (Os and Pb) are sensitive tracers and chronometers of mantle melting due to

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the incompatible behaviour of Re and U relative to Os and Pb, respectively (Luguet and Pearson, 2019). 187

Os/188Os ratios that likely reflect

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The most radiogenic Nahlin ophiolite samples yield suprachondritic

Re mobility and/or radiogenic Os introduced during melt:rock interaction with pyroxenite dykes. Mixing

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with previously melt-depleted peridotite during melt:rock interaction may also explain the relatively

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unradiogenic Os within some orthopyroxenite dyke samples (discussed below). Nahlin ophiolite samples with relatively unradiogenic Os (i.e., samples with subchondritic Os/188Os ratios) yield imprecise TRD model ages that cluster around the inferred Late Permian to middle

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187

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Triassic age of the Nahlin ophiolite (Zagorevski et al., 2015), depending on the reference reservoir chosen for the model age calculation. Because Nahlin ophiolite peridotites represent melt-depleted residues that are mostly devoid of metasomatic BMS, the youngest cluster of TRD model ages at ca. 250 Ma is interpreted as an imprecise minimum estimate for the ophiolite-related mantle melt depletion. The inferred timing of melt-depletion is also coeval with generation of island arc tholeiite basalts that intrude into and are structurally juxtaposed with the mantle peridotites, and the U-Pb zircon age of a gabbro cutting mantle peridotite north of Atlin (ca. 245 Ma)(Zagorevski et al., 2015). Our new TRD age constraints, coupled with previously published geochemical modelling of the coeval basalts (McGoldrick et al., 2017), support an arc setting for magmatism and its depleted mantle residue, which are now preserved as dismembered ophiolite remnants (Zagorevski et al., 2015).

Journal Pre-proof Other peridotite and pyroxenite samples yield Os model ages that are over one billion years older than the Nahlin ophiolite. The two oldest TRD ages correspond to harzburgite (TRD = 1961 Ma; 16LVAA032-A01; 0.48% wt.% Al2O3) and lherzolite (TRD = 1377 Ma; 17LVA-A031-A01; 1.33 wt.% Al2O3) samples from the Hardluck Peak. These older model ages (i.e., TRD ≥ 250 Ma) are comparable to the Os isotopic signature of low-Al2O3 lherzolite xenoliths from the northern Canadian Cordillera (TRD = 0.06– 0.74 Ga)(Peslier et al., 2000; Polat et al., 2018) and the broader heterogeneity seen in Phanerozoic

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convecting mantle due to global events (Dijkstra et al., 2016; Pearson et al., 2007). For abyssal peridotites, ancient TRD ages have represented something of a paradox because mid-

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ocean ridge basalts (MORB) rarely preserve evidence of interaction with ancient mantle lithosphere (Day

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et al., 2017; Warren and Shirey, 2012). Most authors attribute ancient TRD model ages in these settings to

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rare, refractory Os-bearing alloys that form during partial melting (Dijkstra et al., 2016; Pearson et al., 2007). PGE profiles that yield the characteristic Pd-poor signature of residual BMS and/or alloys are rare

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(Fig. 8), but are consistent with the nuggety distribution of Os and Ir (Supplementary Material Fig. 6).

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The few harzburgite and olivine websterite samples that contain significant metasomatic BMS (16LVAA13-01; 16LVA-A14-A01; 16LVA-A21-A01; 16LVA-A27-A01) yield TRD ages older than ca. 250 Ma,

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suggesting that any residual Os-bearing alloys, if present, may have been consumed and/or mixed during

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re-fertilization (i.e., “re-sulphidation”)(Luguet and Pearson, 2019). A similar alloy origin for the least-radiogenic Pb isotope compositions is unlikely given that silicate minerals and/or rare BMS likely represent the dominant host for this element (Warren and Shirey, 2012). For Nahlin ophiolite samples devoid of BMS, we suggest that unradiogenic Pb isotopic signatures are related to melt and/or fluid inclusions trapped within residual phases such as olivine. Such fluid inclusions are ubiquitous in all of the studied peridotite and pyroxenite samples and likely reflect a mixture of high-T mantle and/or low-T serpentinization fluids. Trapped fluid or melt inclusions may explain the relatively complex Pb isotopic signature of the Nahlin ophiolite suite, which appears to record multiple U-Pb fractionation events that over time evolved to depleted and relatively enriched (µ ≥ 9.7) mantle compositions (Fig. 10). Orthopyroxenite dykes that cut harzburgite yield some of the least

Journal Pre-proof radiogenic Pb isotopic compositions. The ancient Pb isotope composition of these late dykes was presumably inherited from the unradiogenic and previously melt-depleted harzburgite host. However, these same dykes yield suprachondritic

187

Os/188Os ratios, possibly because their source had elevated

time-integrated Re/Os in the form of radiogenic Os from a subduction zone fluids. Combining both isotopic systems therefore provides a complementary picture of the Nahlin ophiolite genesis, which can only be explained by considering the variable contribution of residual phases left-over from past melting

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events and re-fertilization of incompatible elements during melt:rock interaction and/or fluid-related metasomatic processes.

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Whole-rock Os data for Nahlin ophiolite peridotites also provide new constraints on the origin of

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placer PGM, which have remained something of a curiosity because of the importance of placer Au

187

Hattori and Hart (1991) have

Os/188Os using a

187

Os/186Os ratios for placer

186

Os/188Os ratio of 0.11984 (Day et al., 2017)] from

Os/188Os 0.120–0.130 (mean = 0.126 ± 0.001; n = 20) and overlap with

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PGM [recalculated here to

187

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mining to the economic development of the region. Previously published

the Os isotopic signature of the Nahlin ophiolite suite (i.e., harzburgite is the most abundant rock type

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comprising ultramafic massifs that yields 187Os/188Os mean = 0.1269 ± 0.0020; 0.1143–0.1507; n = 31). If

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these PGM nuggets preserve their original 187Os/188Os ratios during transport at surface (Hattori and Hart,

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1991), the Os composition of the Nahlin ophiolite provide an exceptional fit. This supports the notion that the Os signature of PGM in historic placer Au workings may have ultimately been derived from ultramafic massifs exposed around Atlin (Dijkstra et al., 2016; Hattori and Hart, 1991)

7 Conclusions Ultramafic massifs of the Nahlin ophiolite suite represent dismembered remnants of a Late Permian to Early Triassic ophiolite. Major, minor and moderately incompatible trace elements demonstrate that harzburgite and dunite are the residues after varying degrees of melt-extraction (i.e., mostly ≥ 20% partial melting based on single-stage, fractional melting of a PM-like source). However, residual harzburgite is cut by multiple generations of pyroxenite dykes, replaced by dunite, and

Journal Pre-proof overprinted by Si-metasomatized olivine websterite domains. These field relationships suggest that metasomatism is superimposed on the melt residues, possibly in an arc setting. Re-fertilization by passing melt(s) and/or fluid(s) is also consistent with the highly incompatible element (LILE)-rich and radiogenic isotope (Os and Pb) signature of a subset of melt-depleted harzburgite samples. Ore-forming elements are implicated in these metasomatic reactions because rare intergranular BMS (pentlandite ± pyrrhotite ± chalcopyrite) occur with clinopyroxene and Cr-spinel at the microscale. Re-fertilization is further

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supported by melt-depleted peridotite samples that yield ore-forming element concentrations that are greater than the PM, opposite to their expected behaviour during partial melting. We demonstrate that

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residual olivine is a significant, if not primary, host for a suite of ore-forming elements, particularly in

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melt-depleted peridotite samples that are devoid of metasomatic BMS. Olivine-hosted ore-forming

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elements likely occur as fluid and/or melt inclusions that became trapped during recrystallization. Despite their uncertain origin, olivine-hosted inclusions and rare Os-bearing alloys, represent the most likely

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explanation for unradiogenic Pb isotope compositions and TRD ages and that predate the formation of the

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Nahlin ophiolite by over one billion years.

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Acknowledgments

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This work was completed under the auspices of the Targeted Geoscience Initiative (TGI)-5 and the Geomapping for Energy and Minerals (GEM) programs. Andrew Locock (University of Alberta) kindly completed EPMA analyses. Xenoliths submitted for nanopowder analyses were kindly provided by Ryan Kroner and Kelly Russell as part of a MSc project at the University of British Columbia. Alain Grenier (GSC) is thanked for his help freeze-drying samples and for conducting laser diffraction particle size analyses. We thank Norm Graham and Paula Vera of Discovery Helicopters (Atlin) for logistical support and transport during field work.

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Figure Captions

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Figure 1

Regional geologic map of northwestern British Columbia (Cui et al., 2017), with sample localities

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2007).

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indicated. Shuttle radar topography mission (SRTM) elevation data are shown for reference (Farr et al.,

Figure 2 Field photos. (a) Ductily folded orthopyroxenite dykes cutting harzburgite; (b) olivine websterite band intercalated with harzburgite; (c) dunite bands intercalated with harzburgite; (d) dunite pod enclosed within harzburgite; (e) thin dunite bands within orthopyroxenite dykes; (f) dismembered orthopyroxenite dykelet hosted within dunite; (g) very coarse to pegmatitic clinopyroxene at margin of orthopyroxenite dyke; (h) parallel websterite dykes intercalated with dunite-chromite and harzburgite.

Journal Pre-proof Figure 3 Scanning electron microscope (SEM) energy dispersive spectroscopy (EDS) mapping results. (a) Cr spinel spatially associated with rare clinopyroxene within partially serpentinized harzburgite; (b) serpentinized harzburgite with clusters of magnetite and native Fe within the serpentinized rock matrix; (c) serpentinized harzburgite with variably replaced BMS; (d–f) metasomatic BMS, clinopyroxene and Cr spinel within olivine websterite. Low-T fluids replaced BMS with native Fe, native Cu and awaruite

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during serpentinization. Abbreviations: Aw = awaruite; Cpx = clinopyroxene; Cu = native Cu; Fe = native Fe; Mt = magnetite; Ol = olivine; Opx = orthopyroxene; Pn = pentlandite; Spi = Cr spinel; and Srp =

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serpentine.

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Figure 4

Rare earth element (REE) diagrams normalized to chondrite (Palme and O’Neill, 2014). Data are based

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on direct LA-ICPMS analysis of nanopowders. Mantle peridotites are depleted in most rare earth

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elements, consistent with the composition of melt-depleted residues. Trace element re-fertilization is indicated by modest La- and Ce-rich (LREE) profiles for a few metasomatized harzburgite and dunite

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samples. Websterite dykes are enriched in LREE relative to orthopyroxenite dykes.

Multi-element diagrams normalized to primitive mantle (Palme and O’Neill, 2014). Data are based on direct LA-ICPMS analysis of nanopowders. Mantle peridotites are depleted in most trace elements, consistent with the composition of melt-depleted residues. Trace element re-fertilization is indicated by scattered incompatible element concentrations, including large ion lithophile elements (LILE; K, Cs, Pb, Sr). Ore-forming element (As through Bi) concentrations greater than the primitive mantle are inconsistent with simple, single-stage partial melt models and likely reflect the re-introduction of BMS and/or trapped fluid inclusions.

Journal Pre-proof Figure 6 Compatible and incompatible element results. Melt models are based on fractional melting (5% increments) using a PM starting composition (Palme and O’Neill, 2014), with partition coefficients (Pearce and Parkinson, 1993; Salters et al., 2002), and modal mineral abundances from Niu et al. (1997) and Lee et al. (2012). Compatible (e.g., MgO, Ni) to moderately incompatible elements (e.g., CaO, V, Ga, Y) closely follow expected melt trends, suggesting that Nahlin harzburgite represent residues after mostly

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≥ 20% partial melting. Highly incompatible element concentrations are inconsistent with melt depletion models (e.g., orthopyroxenite and websterite dykes). Median LA-ICPMS olivine concentrations re-

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calculated to a 70% modal mineral abundance demonstrate that this residual phase is a significant mineral

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host for Pb within the most melt-depleted samples. Akali basalt-hosted peridotite xenoliths (mostly

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lherzolite and lesser harzburgite) from south-central British Columbia analyzed using the same

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nanopowder method as the Nahlin ophiolite suite are shown for reference (Kroner, 2019).

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Figure 7

Ore-forming element results. Melt models are based on fractional melting (5% increments) using a PM

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starting composition (Palme and O’Neill, 2014), partition coefficients (Kiseeva et al., 2017), and the

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modal mineral abundances from Niu et al. (1997) and Lee et al. (2012). Some ore-forming elements (e.g., Cu ± Se ± Ag ± Mo ± W) concentrations are inconsistent with S-only partitioning (D = 1000)(Kiseeva et al., 2017). Moderately incompatible element behaviour during melting is based on the partitioning coefficients of Cu (Kiseeva et al., 2017). Other ore-forming elements yield weak negative-sloped patterns and apparent compatible element behaviour (e.g., As ± Te ± Bi), which requires re-fertilization and/or sequestering by some unknown residual phase. Median LA-ICPMS olivine concentrations re-calculated to a 70% modal mineral abundance demonstrate that olivine is significant host phase for some ore-forming elements within the most melt-depleted samples. Akali basalt-hosted peridotite xenoliths (mostly lherzolite and less harzburgite) from south-central British Columbia analyzed using the same nanopowders method as the Nahlin ophiolite suite are shown for reference (Kroner, 2019).

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Figure 8 Whole-rock PGE results. Melt models are based on fractional melting (5% increments) using a PM starting composition (Palme and O’Neill, 2014), partition coefficients (Kiseeva et al., 2017), and the varying modal mineral abundances of Niu et al. (1997) and Lee et al. (2012). Previously published NiS-

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FA data for the same sample suite are shown for reference (Lawley et al., 2019).

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Figure 9

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Whole-rock Os model age results (TRD = total Re-depletion ages) for Nahlin ophiolite samples. The

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cluster TRD ages at ca. 250 Ma is in good agreement with the Late Permian to Middle Triassic age of the Nahlin ophiolite. Older TRD model ages (≥ 250 Ma) point to residual alloys and/or BMS leftover from

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Figure 10

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past melting events.

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Pb/204Pb and

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Pb/204Pb

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Whole-rock Pb isotope results for Nahlin samples. Data document a range of

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ratios that likely require a combination of U/Pb fractionation (µ = 8–10) during suspected past melting events and/or radiogenic decay after the re-introduction of U (i.e., secondary isochrons). The leastradiogenic Pb isotope signature of Nahlin peridotites is consistent with Os isotope results, but there is no correlation between isotopic systems. Previously published alkali basalt-hosted lherzolite xenoliths from the subcontinental lithospheric mantle underlying southern British Columbia are mostly unfractionated and are shown for reference (Polat et al., 2018).

Figure 11 In situ electron microprobe determined Mg# versus laser ablation inductively coupled plasma mass spectrometry (LA-ICPMS) data for olivine. Some compatible (e.g., Mn ± Ni ) elements correlate with

Journal Pre-proof melt indices [e.g., Mg# = 100*Mg/(Mg+Fe)]. Other elements (e.g., Ca, Na, V, Sc) correlate with Al and Cr, suggesting temperature-sensitive element partitioning during subsolidus re-equilibration (i.e., olivine equilibration temperatures based on the Al-Cr thermometry)(De Hoog et al., 2010). Trace to ultratrace ore-forming elements (e.g., Cu and Au) tend to be associated with olivine that yield the hottest equilibration temperature and least melt-depleted samples (i.e., low Mg#) for this sample subset (n = 11).

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Figure 12 Whole-rock metal ratio plots used to estimate the composition of the bulk silicate Earth (Palme and

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O’Neill, 2014; Sun and McDonough, 1989). Highly incompatible element ratios are assumed to be

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unaffected by partial melting, consistent with the broadly overlapping ratios of chondrite (CI), primitive

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mantle (PM), lower crust (LC), bulk continental crust (BC) and upper continental crust (CC) (Palme and O’Neill, 2014; Rudnick and Gao, 2003). However, Nahlin samples mostly yield suprachondritic

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incompatible element ratios that vary over several orders of magnitude. Similar suprachondritic ratios for

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arc-related volcanic rocks where attributed to mixing between PM and fluids in the magmatic source region (Noll et al., 1996). Alkali basalt-hosted peridotite xenoliths (mostly lherzolite and subordinate

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harzburgite) from south-central British Columbia analyzed also using the nanopowders method are shown

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for reference (Kroner, 2019).

Journal Pre-proof 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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☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:

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