# Sulfur isotope signatures in the lower crust: A SIMS study on S-rich scapolite of granulites

Johannes Hammerli, Anthony I.S. Kemp, Natasha Barrett, Boswell A. Wing, Malcom Roberts, Richard J. Arculus, Pierre Boivin, Prosper M. Nude, Kai Rankenburg

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To cite this version:

Johannes Hammerli, Anthony I.S. Kemp, Natasha Barrett, Boswell A. Wing, Malcom Roberts, et al. Sulfur isotope signatures in the lower crust: A SIMS study on S-rich scapolite of granulites. Chemical Geology, 2017, 454, pp.54-66. 10.1016/j.chemgeo.2017.02.016. hal-01503019

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# Sulfur isotope signatures in the lower crust: A SIMS study on S-rich scapolite of granulites

1 Johannes Hammerli¹, Anthony I.S. Kemp¹, Natasha Barrett¹,², Boswell A. Wing³*, Malcolm

2 Roberts⁴, Richard J. Arculus⁵, Pierre Boivin⁶, Prosper M. Nude⁷, Kai Rankenburg⁸

3 ¹Centre for Exploration Targeting, School of Earth Sciences, The University of Western Australia, Perth, WA

4 6009, Australia

5 ²Department of Earth and Atmospheric Sciences, University of Alberta, 1-26 Earth Sciences Building,

6 Edmonton, T6G 2E3, Canada

7 ³Department of Earth and Planetary Science, McGill University, Montreal, QC H3A 0E8, Canada

8 ⁴Centre for Microscopy, Characterisation and Analysis, The University of Western Australia, Perth, WA 6009,

9 Australia

10 ⁵Research School of Earth Sciences, Australian National University, ACT 0200, Australia

11 ⁶Université Clermont Auvergne, CNRS, IRD, OPGC, Laboratoire Magmas et Volcans, F-63000 Clermont-
12 Ferrand, France

13 ⁷Department of Earth Science, College of Basic and Applied Sciences, University of Ghana, P.O. Box LG58,

14 Legon-Ghana

15 ⁸School of Earth Sciences, The University of Western Australia, Perth, WA 6009, Australia

16 *present address: Department of Geological Sciences, University of Colorado Boulder, UCB 399, Boulder,
17 Colorado 80309-0399, USA

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Johannes.hammerli@uwa.edu.au

25 **Keywords:** S-rich scapolite; lower crust; stable isotopes; SIMS analyses; scapolite in granulites; sulfur cycle; S flux

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## Abstract

31 Scapolite is an important reservoir for volatiles in the deep crust and provides unique insights
32 into the S isotope signatures at the mantle/crust interface. Here, we document the first
33 scapolite reference material (herein referred to as CB1) for in situ S isotope analysis. The
34 chemical and isotope composition of this euhedral, S-rich scapolite megacryst was
35 characterized via LA-ICP-MS, EPMA, SIMS, and bulk fluorination gas source isotope ratio
36 mass spectrometry. The CB1 scapolite is isotopically homogeneous and our results show that
37 crystal orientation does not affect in situ S isotope SIMS analysis. This makes CB1 an ideal
38 primary calibration standard for in situ analysis of S isotope ratios (³⁶S/³²S, ³⁴S/³²S and
39 ³³S/³²S) in scapolite. With this reference material in hand, we then applied in situ SIMS
analysis of S isotopes for the first time on scapolite in granulite samples from the lower crust/upper mantle. The analysed sample suite comprises rocks from classic granulite xenolith locations in southeastern Australia, as well as a sample from the high-grade suture zone of the Dahomeyides in south-eastern Ghana. The results show that scapolites in the lower crust have δ³⁴S values between ~-0.5 and +4 (‰ VCDT). These values fall within the range of S isotope signatures present in mantle rocks and provide no evidence for the recycling of seawater-derived S into the lower crust. We propose that scapolite formed during granulite facies metamorphism of igneous cumulates, where S was sourced from precursor igneous sulfides. Sulfur isotope heterogeneities between individual scapolite grains in some of the studied samples may reflect non-uniform S-isotope compositions of igneous S-phases, which precipitated from mantle-derived melt.

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## 1. Introduction

The global S cycle is of multidisciplinary interest, as it addresses questions concerning the evolution of the atmosphere including its oxygen balance, the evolution of early life on Earth (e.g., Canfield, 2004 and references therein) as well as the formation of ore-deposits throughout the Earth's history (e.g., Benning and Seward, 1996). While surface or shallow subsurface processes involving S distribution and fluxes can be studied and monitored in great detail (e.g., Canfield, 2004), S behaviour in the crust and mantle is more challenging to comprehend. Previous studies have shown that the recycling of S via subduction might be a key process in the global S cycle (e.g., Alt et al., 1993; 2012; Wallace, 2005) and important for the redox budget of the mantle (e.g., Alt et al., 2012; Evans and Powell, 2015). Our current understanding is that significant amounts of S, as sequestered in seawater, sediments, and altered (oceanic) crust, is recycled via subduction and arc magmatism/volcanism (Fig. 1).

Variations in the stable isotope composition of S (i.e., ³⁶S/³²S, ³⁴S/³²S and ³³S/³²S) compared to Vienna Cañon Diablo Troilite (VCDT) have been used as an essential tool for
tracking S fluxes and provenance during geodynamic processes (e.g., Sasaki and Ishihara, 1979; Alt et al., 1993; Cabral et al. 2013; Labidi et al., 2014). For instance, several studies have reported elevated values of δ³⁴S in subduction-related volcanic rocks (e.g., Sasaki and Ishihara, 1979; Ishihara and Sasaki, 1989; Alt et al., 1993) (Fig. 2). These phenomena might be attributed to fluids released from the subducting slab, derived from seawater and sulfates with strongly positive δ³⁴S values (+20 ‰, Fig. 2), which are recycled through the upper (sub-arc) mantle and overlying crust, eventually reaching the atmosphere via volcanic degassing (Fig. 1). However, using S isotopes to constrain subduction-related and global S mass transfer relies on how well we can constrain the different S reservoirs in terms of their S content and S isotope characteristics. While we have a general idea about the S isotope composition of subducted material (input) and arc magmas (output) (e.g., Alt, 1995; Marini et al., 2011 and references therein), one of the missing links for understanding the S cycle concerns the behaviour and isotope composition of S in the lower crust, at the mantle-lithosphere interface.

Rocks from the lower crust/upper mantle may be the key to better constrain the S isotope characteristics of slab-derived fluids. This is because the S isotope signatures of deep crustal rocks are less likely to be affected by interaction with supracrustal material and degassing reactions compared to volcanic samples, or rocks in the middle/upper crust (e.g., Sasaki and Ishihara, 1979; Ishihara and Sasaki, 1989).

Granulites affiliated with subduction zones are promising candidates for gaining insight into the S isotope signature of the lower crust, and for tracing potential S fluxes from the mantle and subducted slab (Fig. 1). Granulites are generally sulfide-deficient (Porter and Austrheim, 2016), however, metamorphic S-rich scapolite is widespread in lower crustal rocks (e.g., Knorring and Kennedy, 1958; Lovering and White, 1964; Goldsmith, 1976; Moechner et al., 1994; Yoshino and Satish-Kumar, 2001; Aulbach et al., 2010; Mansur et al., 2014; Porter and Austrheim, 2016) and is possibly a major rock-forming mineral in the lower
crust (e.g., Lovering and White, 1964; Goldsmith, 1976). Scapolite in granulite/upper mantle xenoliths rapidly delivered to the surface by volcanic pipes with little to no retrogression might therefore provide a novel window into S sources and S isotope signatures in the lower crust (Fig. 1).

In this study, we measure S isotopes in scapolite formed at lower crustal conditions at high spatial resolution by secondary ionisation mass spectrometry (SIMS). Grains were analysed from petrographic thin sections, preserving textural context and allowing resolution of S isotope signatures on a micrometer-scale. We propose that in the case of scapolite, post-formation alteration and crustal contamination can be readily resolved by means of EPMA, LA-ICP-MS and SIMS analysis by targeting fresh cores of grains. The analysis of unaltered scapolite is particularly important for S isotope quantification, as sulfides might form as an alteration product (Porter and Austrheim, 2016), potentially leading to a significant S isotope fractionation that cannot be accounted for in bulk grain analysis. Our scapolite results fall within the same $\delta^{34}$S range as previously found in mantle xenoliths studies, suggesting that seawater-derived S is unlikely to be a major component in the studied lower crustal samples. Furthermore, some of the studied samples show S isotope variations between individual scapolite grains, which we attribute to isotope heterogeneity of primary magmatic sulfides that acted as the major S source for scapolite formation during granulite facies metamorphism.

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## 2. Background

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Previous studies have improved the general understanding of S behaviour prior to and during subduction and associated magmatism. For instance, Alt (1995) showed that the $\delta^{34}$S values of bulk altered oceanic crust (~0.9 ‰) are only minimally different from their original basaltic source (~0.5 ± 0.6 ‰; e.g., Sakai et al., 1984). Other studies showed when and how S
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is expected to be released due to mineral reactions during subduction, and how it might be
recycled within the slab itself (Fig. 1) (Evans et al., 2014).

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Alt et al. (1993) studied the S isotope signatures from the Mariana island arc and back arc basin and found that δ³⁴S is enriched (mean ~ +3.8 ‰ δ³⁴S) in arc volcanic rocks, which they interpret to reflect a seawater S component in the sub-arc mantle source. The authors estimated that approximately one third of the S in arc magmas might be contributed by a metasomatic slab-derived fluid that mixed with mantle MORB-like S in the magma source region. Back-arc basalts on the other hand showed δ³⁴S values of 1.1 ± 0.5 ‰, which are similar to MORB values. These significantly lower values in back-arc basalts are attributed to the dilution of the slab S isotope signature in the predominantly unmetasomatized mantle source further from the arc. de Hoog et al. (2001) studied S isotope compositions in basaltic and basaltic andesitic lavas from the Indonesian arc. The study concluded that S is derived from a mantle source that is significantly enriched in δ³⁴S compared to typical MORB and OIB sources. The authors estimated that the magma sources had average δ³⁴S values between +5 and +7 ‰. These elevated δ³⁴S values are inferred to be a result of significant S transport by slab-derived fluids into the mantle wedge. Intriguingly, de Hoog et al. (2001) also showed that the range of δ³⁴S values in lavas is rather limited, despite the potentially large variations in S isotope composition of the subducted material.

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Wilson et al. (1996) studied mantle xenoliths and found evidence of metasomatized mantle that led to elevated δ³⁴S isotopes values of up to +7 ‰, attributed to the subduction of crustal S. Ionov et al. (1992) examined a range of upper mantle xenoliths and their results show that most of their analysed peridotite xenoliths fall between -1 and +4 ‰ δ³⁴S, with an overall average of +2.1 ‰ δ³⁴S. These studies on mantle xenoliths imply non-uniform S isotope signatures in the mantle. Several other studies have also challenged the uniformity of the mantle S isotope composition. Chaussidon and Lorand (1990) have found evidence that

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the upper mantle S isotope signatures might range from -3 to +3 ‰ δ³⁴S, which can possibly
explain S isotope variations in MORBs, alkali basalts and continental tholeiites. More recent
work (Labidi et al., 2013, 2015; Cabral et al. 2013) also proposes variable and significantly
depleted δ³⁴S values for MORBs, with a proposed depleted mantle end-member of ~1.3 ‰
(Labidi et al., 2013). OIBs on the other hand were found to range between strongly negative
and positive δ³⁴S values (Cabral et al., 2013; Labidi et al., 2015) (Fig. 2). These findings of
significant S isotope variations in the mantle reservoir further complicate the interpretation of
S isotope compositions of arc magmas and middle crustal rocks.

The occurrence of scapolite at elevated P-T conditions has relevance for the
sequestration of volatiles in lower crustal environments (e.g., Lovering and White, 1964;
Goldsmith, 1976; Moecher et al., 1992). Scapolite stable at these conditions is typically
enriched in C and/or S, approaching the meionite Ca₄(Al₆Si₆O₂₄)CO₃ or silvialite
compositions Ca₄(Al₆Si₆O₂₄)SO₄ (Sokolova and Hawthorne, 2008 and references therein),
depending on the availability of S. According to experimental data by Newton and Goldsmith
(1977), end-member silvialite requires at least 775°C at 17kbar and 1200 °C at 10 kbar to
form. Such conditions were confirmed by the study of Stolz (1987) whose pressure-
temperature calculations indicated that silvialite from the McBride Province, North
Queensland, Australia, crystallized from alkali basaltic magma at ~900-1000° and 8-12 kbar.
This restricted stability field of meionite and silvialite scapolite means that those phases can
be used to track C and S isotope signatures in the lower crust/upper mantle (e.g., Moecher et
al., 1994; Yoshino et al., 2002; Hoefs et al., 1987; Iyver et al., 1992). For example, Moecher
et al. (1994) studied C isotope signatures in scapolites from a range of deep crustal granulites
and xenoliths, and concluded that these minerals derived their C primarily from a mantle
source.

Studying the S isotope composition in scapolite from granulites is key to furthering
our understanding of S recycling during high-grade metamorphism at convergent margins
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(e.g., Alt et al., 2012; Evans et al., 2014, Tomkins and Evans, 2015; Evans and Powell, 2015).

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Moreover, Porter and Austrheim (2016) suggest that the breakdown of S-rich scapolite during retrogression of originally high temperature granulites might be a significant source for S in the middle/upper crust, as an important ligand for the transport of ore metals.

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**3. Samples and standard reference material (CB1)**

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**3.1 Scapolite megacryst CB1 (Massif Central, France)**

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The Enval-Volvic volcanic line, east of the Chaîne des Puys in the Massif Central, France, is one of the few locations where S-rich scapolite megacrysts are found (Boivin and Camus, 1981). This 13 km long fissure is 90 kyr old, and represents one of the first volcanic episodes of the Chaîne des Puys, Massif Central, France. From south to north, the Enval-Volvic fissure is marked by the alignment of the Enval maar and the associated spatter cone of Chuquet-Genestoux; the Puy de Couleyras; the maar-spatter cone couples of Bois de Chanat and Bois de Clerzat; and the Puy de la Bannière above Volvic village. However, despite the current intraplate position of the Enval-Volvic volcanic line, several studies (e.g., Femenias et al., 2004 and references therein) found evidence that some xenoliths delivered by volcanics in the Massif Central record upper mantle metasomatism potentially related to earlier Variscan subduction.

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Cobber 1 (CB1), sampled from the same location as given in Boivin and Camus (1981) is a large (~1 cm), euhedral scapolite crystal of a dark grey/green colour, hosted in tephra. In addition to scapolite, other megacrysts from the tephra include clinopyroxene + amphibole (kaersutite) + Fe-Ti oxides + feldspar (andesine to K-oligoclase) carried up by basalts erupted along the Enval-Volvic fissure (Boivin and Camus, 1981). Liotard et al. (1988 and references therein) suggest that the megacryst suite represents a paragenetic assemblage of phenocrysts that were in equilibrium with a differentiated, deep-seated magma.

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## 3.2 Scapolite in mafic xenoliths from the Monaro Volcanic Province and Delegate Pipes, New South Wales (Australia)

We studied five scapolite-bearing granulites for which detailed petrological and petrographic descriptions, including EPM analyses can be found in Barrett (2014) (Fig. 3, and see appendix for sample location coordinates). The samples MVP99-2-05, MVP99-2-12, and MVP99-2-13 are xenoliths from the Eocene-Oligocene intraplate basaltic Monaro Volcanic Province (MVP) in New South Wales (NSW), southeastern Australia (Wellman and McDougall, 1974; Taylor et al., 1990; Roach, 2004). The MVP volcanic plugs and dykes, and less common maars, cover a range of primary to moderately evolved rocks, including: olivine nephelinite, melanephilinite, nepheline basanite, alkali olivine basalt, tephrite and K-trachybasalt (Roach, 2004). The more than 65 volcanic centers of the MVP delivered a variety of crustal and mantle xenoliths. The samples MVP99-2-12 and MVP99-2-13 represent garnet granulites with a mineral assemblage consisting of garnet (Al₃₅-₃₇Gr₂₁-₂₂Py₄₁-₄₃) + clinopyroxene (En₃₈Fs₁₄Wo₄₈) + plagioclase (An₅₆-₅₇Ab₄₀-₄₁Or₂.₈-2.₉) + scapolite, as well as accessory rutile and magnetite. These samples have a coarse, equant texture with most of the grains being between 0.5 and 2 mm in diameter, forming a polygonal granular assemblage (Fig. 3). In sample MVP99-2-12, clinopyroxene is the most common mineral with an estimated modal abundance of ~40%, followed by plagioclase (~30%), garnet (~25%), and scapolite (~5%). Sample MVP99-2-13 contains clinopyroxene and garnet, both estimated at a modal abundance of ~35%, while plagioclase is estimated to make up ~25-30% of the mineral assemblage, and scapolite occurs as a minor phase (<5%). Garnet has distinct dark kelyphitic rims in both samples (Fig. 3), which have been interpreted to be a product of a breakdown reaction of garnet due to a sudden change of pressure and temperature when the xenoliths were transported to the surface (Keankeo et al., 2000). Scapolite shows conspicuous alteration rims that can vary from a few microns to ~20 microns thick. These fine-grained zones consist of a matrix of plagioclase and possibly sericite. Sample MVP99-2-05 is a two-pyroxene
granulite that contains an estimated modal abundance of ~35% plagioclase and clinopyroxene, ~25% orthopyroxene, and ~5–10% scapolite, which always shows alteration rims of similar composition to those observed in garnet-pyroxene granulite samples.

Sample 10438B also represents a two-pyroxene granulite collected from the Delegate breccia pipes, NSW, Australia, the same location where Lovering and White (1964) described for the first time S-rich scapolite in granulite xenoliths. The Delegate pipes are located approximately ~50km SW from the MVP samples described above. The major mineral phases (clinopyroxene (En₃₈Fs₁₂–₁₄Wo₄₉)+orthopyroxene (En₆₅–₆₇Fs₃₃–₃₄Wo₀–₁.₅)+plagioclase (An₈₅–₈₇Ab₁₃–₁₅Or₀.₂–₀.₃)+scapolite form a granular assemblage with sharp, polygonal grain boundaries (Fig. 3). Plagioclase is estimated to make up ~50% of the mineral assemblage, followed in abundance by ~30% orthopyroxene, and ~15% clinopyroxene. Scapolite occurs as a minor mineral phase (making up <5% of the mineral assemblage), and similar to the MVP samples, a reaction texture on the scapolite rims is always present (Fig. 3F). This sample shows compositional banding, defined by alternating pyroxene (orthopyroxene and clinopyroxene) and plagioclase-rich layers on both a hand specimen- and thin section-scale, interpreted as relict igneous cumulate layering.

Preliminary data from two-pyroxene granulites of the MVP by Barrett (2014) confirm earlier findings by (Chen et al., 1998) who dated two-pyroxene granulites from the nearby Delegate pipes. The work of Chen et al. (1998) revealed ages (U-Pb, zircon) of ~391 and ~398 Ma for the Delegate pipe granulites, confirming that these are significantly older than the volcanic activity at both localities (MVP: ~55–34 Ma; Wellman and McDougall, 1974; Delegate pipes: ~170–160 Ma; Lovering and Richards, 1964). Based on experimental work on samples from the Delegate pipes, Irving (1974) concluded that the two-pyroxene xenoliths represent lower crustal material, whereas layered garnet-pyroxene xenoliths originally formed as cumulates in melt pockets in the mantle. Barrett (2014) and White and Chappell (1989) proposed that the two-pyroxene and garnet-pyroxene granulite xenoliths formed as igneous
cumulates with subsequent (re-)crystallization at granulite facies conditions in the lower crust.

Granulite facies metamorphism in this part of Australia can be linked to the vast subduction-
related accretionary orogenic system along the eastern Palaeo-pacific margin of Gondwana
during the Palaeozoic and Mesozoic (Cawood and Buchan, 2007 and references therein).

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252 **3.3 Scapolite in granulite gneiss from the Shai Hill, south-east Ghana**

We have also studied a scapolite-bearing gneissic granulite from the suture zone of the Dahomeyides orogen in south-east Ghana (sample MW1). This orogen is interpreted to have resulted from the easterly subduction of the rifted margin of the West African Craton during the Pan-African orogenic events (e.g., Affaton et al., 1991; Attoh and Nude, 2008). The studied sample was collected from the high-grade metamorphic suture zone that comprises mafic-ultramafic rocks, eclogites, and granulite gneisses, which crop out in the Mampong Inselberg (Attoh and Nude, 2008). MW1 comes from the same Shai Hill locality where Knorring and Kennedy (1958) originally described garnet-hornblende-pyroxene-scapolite gneiss. Peak metamorphism was likely attained at ~610 Ma when some of the suture zone rocks are thought to have reached mantle depths during a collisional orogeny at the West African Craton margin (Attoh and Nude, 2008 and references therein). MW1 shows a distinct gneissic banding defined by garnet-pyroxene-hornblende and plagioclase-quartz rich zones. Scapolite is more abundant in the plagioclase bands, although it is still a minor mineral phase (<2%) when compared to garnet (~35%), plagioclase (~35%), clinopyroxene (~15%), and hornblende (~10%). Rutile is the most common accessory mineral. Some scapolite grains are heavily altered (generally to K-feldspar), however, fresh scapolite occurs in the plagioclase-rich layers. These fresh scapolites do not contain a reaction rim as typically found in the other samples (see above).

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272 **4. Analytical methods**
273 Matrix matched samples and reference materials are essential for SIMS analysis. Due
274 to the potential for variable sputtering rates, crystallographic orientation of the standard and
275 the unknown might also need to be matched or corrected for in some cases (e.g., Eiler et al.,
276 1997). In order to test the homogeneity of natural scapolite, and to assess the effect of
277 crystallographic orientation for S isotope analysis by SIMS, we chose a single scapolite
278 megacryst from the Chaîne des Puys volcanics in the Massif Central, France (CB1, see above
279 for sample description). After the inclusion-free scapolite crystal was freed from its
280 surrounding tephra, sections were cut along the C- and A-axis, mounted in a 2.5 cm epoxy
281 resin puck, and polished for microanalysis by EPMA, LA-ICP-MS and SIMS. Following
282 EPMA mapping and LA-ICP-MS trace element analysis, we conducted SIMS S isotope
283 traverses across the two fragments to test for homogeneity and crystallographic effects.
284 EPMA and LA-ICP-MS analyses were conducted at the University of Western Australia,
285 using a JEOL JXA8530F electron probe equipped with 5 tunable wavelength dispersive
286 spectrometers and an Analyte G2 laser coupled to an X-series II mass spectrometer,
287 respectively. Detailed information and the experimental set-ups for LA-ICP-MS and EPMA
288 can be found in the electronic appendix.

289 Scapolite grains in the granulite samples were analysed in polished petrographic thin
290 sections (~50-60 microns thick). Following imaging and EPMA analysis, small (2.5 mm)
291 discs containing the targeted scapolite grains were extracted by microdrilling, and pressed
292 into indium for SIMS and LA-ICP-MS analysis (see appendix for details).

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294 **4.1 Fluorination method**

295 The S isotope composition of two aliquots of the same crystal were analysed by
296 chemical extraction of total S, coupled with fluorination and dual-inlet gas-source isotope-
297 ratio mass spectrometry in the Stable Isotope Laboratory of the Department of Earth and
298 Planetary Sciences at McGill University. Sulfur from powdered fragments of CB1 was
extracted using a Kiba Reagent (Tin (II) Strong Phosphoric Acid; Kiba et al., 1955) technique as modified for samples with low S abundances (Hong et al., 2000). The Kiba reagent was prepared by heating a solution of 80.0g NaCl, 80.0g SnCl₂-2H₂O, and 1000g of orthophosphoric acid at 300°C until ≈190ml of liquid has evaporated from the mixture (Hong et al., 2000). A mixture of powdered CB1 and Kiba reagent was heated to 300°C, and held at this temperature for 60 minutes under a steady stream of N₂ gas. This liberated all S in the powder as H₂S, which was then bubbled through a 0.2 M AgNO₃ solution where it was precipitated directly as Ag₂S. The Ag₂S precipitate was collected, washed twice with doubly-deionized water, once with 1M ammonium hydroxide solution, twice again with doubly-deionized water, and then dried overnight. Dried Ag₂S samples were reacted with F₂(g) in nickel bombs at 250 °C to generate pure SF₆(g). The isotopic composition of SF₆(g) was purified cryogenically and chromatographically and analyzed on a Thermo MAT-253 in dual inlet mode. Results were normalized to repeated measurements of international reference material IAEA-S-1, with a defined δ³⁴S value of -0.3‰ on the Vienna Canyon Diablo Troilite (V-CDT) scale. We took the δ³³S value of IAEA-S-1 to be -0.061‰ V-CDT and the δ³⁶S value of IAEA-S-1 to be -1.27‰ V-CDT (Wing and Farquhar, 2015). Sulfur isotope compositions are expressed as:

$$ \delta^{32}\text{S} = \left( \frac{^{32}\text{R}_{\text{sample}}}{^{32}\text{R}_{\text{V-CDT}}} - 1 \right) \times 1000 \qquad (1) $$

where $^{32}\text{R} = ^{32}\text{S}/^{32}\text{S}$ and i is 3, 4, or 6, and

$$ \Delta^{32}\text{S} = \delta^{32}\text{S} - 1000 \times \left( \left( 1 + \frac{\delta^{32}\text{S}}{1000} \right)^{\frac{32}{\lambda}} - 1 \right) \qquad (2) $$

where j is 3 or 6.

We calculated Δ³³S and Δ³⁶S values through reference mass dependent exponents of $^{33}\lambda=0.515$ and $^{36}\lambda = 1.9$ (Wing and Farquhar, 2015). Uncertainty (2SD) on the entire
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analytical procedure is estimated to be better than 0.3‰ for δ³⁴S, 0.02‰ for Δ³³S and 0.3‰ for Δ³⁶S.

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## 4.2 SIMS

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### 4.2.1 Analytical conditions

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The SIMS analyses reported in this study were carried out using the Cameca IMS-1270 (#309) instrument at the School of GeoSciences at the University of Edinburgh, UK.

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Sulfur isotopes ($^{32}$S and $^{34}$S) were analysed as S⁻ ions produced by bombardment of the target by a ~5nA, $^{133}$Cs+ primary beam accelerated at +10kV, resulting in a net impact energy of 20keV at the sample surface. To eliminate charging during the analysis, the samples were coated with ~30nm layer of gold and the exact position of the primary Caesium beam was flooded with low energy electrons produced by the normal incidence electron gun.

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Secondary ions were accelerated at -10kV and analysed at a mass resolution of ~3600, which is sufficient to resolve molecular interference by $^{16}$O-$^{18}$O, $^{33}$S$^{1}$H and $^{32}$S$^{1}$H$_2$ on the $^{34}$S-peak. An energy window of 40 eV was used. After pre-sputtering for 60s, automated secondary ion beam alignment was performed using the DT$_{xy}$ deflection plates to centre the beam in the mass spectrometer field aperture (3,000µm) and entrance slit (80µm) position. The L'2 and H'2 Faraday cups were calibrated at the start of each analytical session using in-built Cameca hardware and software.

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Measurements of S isotopes were made simultaneously in multi-collection mode using two off-axis Faraday cups (L'2 for $^{32}$S and H'2 for $^{34}$S). Count rates were typically ~9×10⁷ cps of $^{32}$S and ~4×10⁶ cps of $^{34}$S. The total acquisition time was 160s comprising of 20 cycles, split into two blocks of 10, with each cycle comprising of a 8s counting period. A single ~5-min analysis (including 60 s pre-sputtering time and beam centring routines) resulted in an internal error of <0.13‰. Standards were analysed after every 10-15 unknowns so that the change in instrumental mass fractionation could be monitored and corrected for, if required.

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## 4.2.2 Measurements

A total of 381 SIMS measurements were carried out over 4 different sessions, of which 224 measurements were of the reference scapolite CB1, fragments of which were mounted in each sample block. The scapolite reference CB1 had been previously checked for major element and isotopic homogeneity, together with confirmation that there was no preferential instrumental mass fractionation associated with crystal orientation (e.g., Kita et al., 2010).

Most sessions showed a slight linear drift in the instrumental fractionation with time, presumably caused by changes in the primary beam current and density. This linear drift was corrected for in each session by linear regression through the standard data acquired during that session. The correction for drift resulted in <0.1‰ correction to the data.

The average ³³S/³⁴S were calculated as cycles accumulated. Cycles outside the 3σ standard deviation were rejected. The remaining cycles were used to calculate the final ³³S/³⁴S and its standard deviation. This standard deviation divided by the square root of the number of retained cycles gives a standard mean error, referred to as internal precision (Fitzsimons et al., 2000). All unknown analyses were standardized with the S isotope values of CB1 obtained by the fluorination method.

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# 5. Results

## 5.1 Major element concentrations of scapolites

Quantitative EPMA S maps of our standard scapolite CB1 show a subtle oscillatory growth zonation (Fig. 4). This difference can be detected by EPMA spot analyses where the difference in the S concentration between the bright zones (e.g., 3.9 wt.% SO₃, Table 1) and the dark zones (e.g., 3.7 wt.% SO₃, Table 1) is beyond the analytical uncertainty (~1.2% at a 99% confidence level, see appendix Table). The average SO₃ content is 3.86 wt.% ± 0.15
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(2SD) and 2.24 wt.% ± 0.14 CO₂, while the average Cl concentration is 0.23 wt.% ± 0.03. The
average Na₂O content is 3.80 wt.% ± 0.22, translating to a meionite component of ~70%
(Table 1 and appendix).

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Analyses of MW1 scapolite show uniform values for all major elements with SO₃
contents being close to 5 wt.% which equals to XS~0.57, where XS is the atomic occupancy
of S at the anion site (A) of the crystal defined as A=1=(S+Cl+C) (e.g., Teertstra et al., 1999)
(Table 1 and appendix). CO₂ fills the rest of the anion site, whereas Cl contents are negligible
(<0.1 wt.%) (Table 1). This also agrees with quantitative EPMA element maps, which show a
homogenous major element distribution and only subtle variations in S contents (Fig. 5).
MW1 scapolite contains the highest Na₂O (~4.75 wt.%) contents, which results in the lowest
meionite component (Me% ~65).

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Scapolite from the Delegate pipe locality (10438B) also contains ~5 wt.% SO₃ and
similar levels of calculated CO₂ (~1.9 wt.%) as determined in MW1 (~2 wt.%). However, the
Na₂O content in 10438B scapolite is lower (~2.8 wt.%) which results in a higher meionite
component (Me% ~79.5). Some scapolite grains show S depletion and Cl and CO₂
enrichment towards the rim (SO₃: ~4.9 wt.% to ~3 wt.%, Cl: ~0.01 wt.% to ~0.02 wt.%, and
CO₂(calc): 2 wt.% to 3 wt.%) (Fig. 5 and appendix table).

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The scapolite compositions of MVP99-2-13 and MVP99-2-12 are homogeneous (Fig. 5, Table 1) and chemically similar, however, scapolite in MVP99-2-12 contains slightly higher SO₃ contents (~4.5 wt.%) compared to MVP99-2-13 (~4.1 wt.%). Scapolites from both samples contain similar amounts of CO₂ (~2–2.4 wt.%), which means that similar proportions of S and C occupy the anion site (Table 1). MVP99-2-12 (~3.8 wt%) shows a minor enrichment in Na₂O relative to MVP99-2-13 (~3.4 wt.%) whereby MVP99-2-12 has a lower meionite component (~70%) compared to MVP99-2-12 (~74%). Sample MVP99-2-05 contains the lowest average SO₃ concentrations (~3.3 wt.%), which means that CO₂ is the major volatile component in the crystal structure (XC~0.6). The Na₂O content of MVP99-2-

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05 is ~2.9 wt.% (Me ~77%) (Table 1). Some of the scapolites show zoned S concentrations, which are mirrored by Cl and CO₂(calc) contents, where cores have higher S contents (~4.4 wt.% SO₃) than rims (~3.3 wt.% SO₃). Chlorine can be slightly enriched in rims (~0.05 wt.%) compared to cores (~0.03 wt.%), which is also true for CO₂(calc) (~2.2 wt.% in core vs. ~2.8 wt.% in rim) (Fig. 5, appendix table and Fig. A-1). K₂O concentrations for all the analysed samples are low (<0.3 wt.%) and FeO and MgO are typically <0.5 wt.%.

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## 5.2 Trace element concentrations LA-ICP-MS

The complete data set for trace elements measured via LA-ICP-MS can be found in the appendix. Scapolites may contain several tens of ppm LREE, whereas HREE concentrations are in the sub-ppm range. Beside sample MW1, which only shows a subtle positive Eu anomaly, all studied scapolites show a distinct positive Eu anomaly and HREE depleted chondrite-normalized trace element patterns (Fig. 6). Scapolites in garnet-bearing granulites are more depleted in HREE than the two-pyroxene granulites and CB1, with Er, Yb, and Lu concentrations often being below the limit of detection. While no trace element variations within individual grains and samples are observed, scapolite from different samples contain variable concentrations of Sr, Ba and Pb (see appendix). For instance, scapolite in samples CB1 and MVP99-2-12 contain ~2000–2400 ppm Sr on average, while sample 10438B contains ~500 ppm Sr. Similarly, Ba contents in scapolite from sample 10438B (~30 ppm) are lower compared to Ba concentrations in MVP99-2-12 and CB1 (~70 ppm and ~125 ppm, respectively). Lead values are highest in sample MW1 (~5.3 ppm on average) and lowest in CB1 (~1.1 ppm).

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## 5.3 Fluorination method

The two repeats of the CB1 aliquots are consistent within analytical uncertainty and give δ³⁴S values of 5.22 ‰ ± 0.30 (2SD) and 4.94 ‰ ± 0.30 (2SD), with mass-dependent

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426 $\Delta^{33}\text{S}$ (-0.008 ‰ ± 0.016; -0.018 ‰ ± 0.016) and $\Delta^{36}\text{S}$ values (-0.21 ‰ ± 0.30; -0.28 ‰ ± 0.30). Uncertainties in the $\delta^{33}\text{S}$ and $\delta^{36}\text{S}$ values for CB1 covary strongly with the uncertainties for $\delta^{34}\text{S}$ values due to mass-dependent fractionation processes in the measurement procedure (cf. Wing and Farquhar, 2015), giving rise to significant statistical covariances among the uncertainties in these measurements (see appendix).

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### 5.4 S isotope analysis by SIMS

#### 5.4.1 CB1 as primary calibration standard for in situ S isotope analysis

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The existence of S-rich scapolite has been known for several decades. Until now, however, in situ methods for S isotope analysis in scapolite have not been applied, due mainly to the scarcity of potential standard reference material and unknown crystallographic effects for SIMS analysis. We found that the $^{34}\text{S}/^{32}\text{S}$ ratio of CB1 measured by SIMS is independent of the crystallographic orientation of the systematically analysed scapolite fragments and of the oscillatory zonation in this crystal (Fig. 7). The concentration of S, which can semi-quantitatively be monitored by the $^{34}\text{S}$ counts, is independent of the $^{34}\text{S}/^{32}\text{S}$ ratio (See appendix Fig. A-2). The reproducibility of $\delta^{34}\text{S}$ values in CB1 over four analytical sessions is 0.40 ‰ (2SD), comparable to the reproducibility of the two bulk-grain aliquots analysed by the fluorination method (see above). These data, coupled with elemental homogeneity and the lack of inclusions make CB1 suitable as a primary calibration standard for S isotope analyses by SIMS. Furthermore, the chemical composition of CB1 (Table 1) falls within the range of all unknowns, which minimises the potential for analytical matrix effects. Scapolite CB1 was subsequently used to standardize all the SIMS analyses of the unknown scapolite from the granulite samples (see below).

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#### 5.4.2 Sulfur isotope analysis in scapolite from granulites
Scapolite grains from sample MVP99-2-05 (two-pyroxene granulite) are internally homogeneous for ³⁴S/³²S, however, δ³⁴S values vary from -0.9 to +0.3 between individual grains within the same thin section (Fig. 8). The second two-pyroxene granulite sample, 10438B, shows a homogeneous S isotope pattern of individual grains with a mean δ³⁴S value of 0.07 ± 0.16 (2SD) ‰. As with sample MVP99-2-05, scapolite of MVP99-2-12 (garnet granulite) shows S isotope differences between individual grains. The highest S isotope ratios measured in this sample are approximately +3 ‰ δ³⁴S, whereas the lowest S isotope ratios are ~+1.7 ‰ δ³⁴S (Fig. 8). Importantly, scapolites from this sample are chemically homogeneous without S depletion towards the rims (Fig. 5). The other garnet-granulite sample, MVP99-2-13, contains scapolite with a homogeneous S isotope composition of +1.25 ± 0.17 ‰ δ³⁴S.

Scapolites from sample MW1, representing the garnet-pyroxene-hornblende granulite from Mampong, Shai Hills, southeastern Ghana, also show a homogeneous (relative to analytical reproducibility) S isotope composition of +4.22 ± 0.37 ‰ δ³⁴S with no resolvable variations between the analysed grains.

In general, scapolite from the garnet bearing samples has elevated δ³⁴S values compared to scapolite of the two-pyroxene granulites MVP99-2-05 and 10438B (Fig. 8). However, CB1 scapolite, part of a garnet-absent assemblage, contains the heaviest S isotope signature (~+5.08 ‰ δ³⁴S, see section 5.3).

# 6. Discussion

## 6.1 Formation conditions of the studied scapolites

The high-resolution quantitative EPMA maps show a subtle but distinct oscillatory growth pattern in CB1 (Fig. 4), which under the assumed P-T formation conditions supports an igneous origin. HREE abundances of CB1 follow the same pattern as scapolite from garnet-absent granulites (Fig. 6 and see below), which suggests the absence of garnet in the
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source and agrees with the lack of garnet megacrysts in the host basaltic tephra. The
formation of scapolite megacrysts in the Chaîne des Puys has previously been explained by a
high fSO₂ content and exceptionally high fO₂ in the relatively undifferentiated alkalic magma
source (Boivin and Camus, 1981). Boivin and Camus (1981) concluded that the megacrysts
formed at high pressure and temperature (~10 kbar and ~1100° C) from a basic magma under
high aH₂O conditions.

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Irving (1974) performed experiments on two-pyroxene and garnet-plagioclase-pyroxene granulite xenoliths from the Delegate pipe locality. The experiments showed that the mineral assemblage, although scapolite-deficient, can be reproduced at 6-10kb and 1100°C for two-pyroxene granulites. Garnet pyroxenite xenolith assemblages were reproduced at 13-17 kbar and ~1050 to 1100 °C. The samples studied here are very similar to those described in Irving (1974). Barrett (2014) studied MVP and Delegate pipe xenoliths in detail and estimated pressure-temperature conditions of ~11kbar and ~1050°C for a scapolite-bearing garnet granulite (MVP99-2-12). Calculations for the P-T conditions of two-pyroxene granulites also support the previously constrained P-T formation environment by Irving (1974). Barrett (2014) concluded that (scapolite-bearing) granulites from NSW (Australia) likely represent cumulates that recrystallized under granulite facies conditions, which agrees with earlier findings by Irving (1974) and White and Chappell (1989). A cumulate origin for these samples is supported by the distinct banding of garnet, pyroxene, and plagioclase-layers and the conspicuous positive Eu anomalies evident in whole-rock analyses.

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Thermobarometric calculations on garnet-bearing granulites from the Dahomeyide suture zone (sample MW1) suggest a peak metamorphic temperature of at least ~800°C and ~13–14kbar for samples from the Shai Hill locality (Attoh and Nude 2008 and references therein). The presence of primary hornblende supports lower formation temperatures than reported for the other studied scapolite-bearing samples.

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Chondrite normalized trace element patterns of scapolite from garnet-bearing and garnet-absent samples are distinctly different (Fig. 6). Scapolites in garnet-bearing samples are more depleted in HREE than their equivalents in garnet-absent granulites. This strongly suggests preferential HREE partitioning into garnet at the time of scapolite formation, resulting in a more prominent HREE depletion in the coexisting scapolite (cf. Hammerli et al., 2014). Based on this, together with textural relationships and polygonal grain boundaries, we infer that scapolite in the studied granulites formed at peak metamorphic conditions, when the major mineral phases crystallized together.

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## 6.2 Major element composition of scapolites

Whereas scapolite grains from samples MVP99-2-12, MVP99-2-13 and MW1 are homogeneous in terms of their major element composition, some scapolite in granulite xenoliths 10438B and MVP99-2-05 show S, Cl, and CO₂ compositional zonation. The replacement of S-rich scapolite by more C- and Cl-rich scapolite, possibly via dissolution-recrystallization (see Putnis and Austrheim, 2010), has been observed in retrogressed scapolite (Porter and Austrheim, 2016). In the samples of the present study, the substitution of Cl for S along cracks further supports retrogressive processes rather than changing physiochemical conditions during the mineral's formation.

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## 6.3 Scapolite formation and the source of S

Meaningful interpretation of S isotope data from high-grade rocks is challenged by the possibility of multiple generations of sulfide formation, and subsequent S mobility during retrograde metamorphism and/or alteration at low temperatures (e.g., Morrison and Valley, 1991). In the studied samples, scapolite is the only S-bearing phase, and due to its restricted stability field (see above) it is likely that scapolite records the local S isotope signature present at the time of its formation in the lower crust.

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Interestingly, some samples (MVP99-2-05 and MVP99-2-12) show isotope variations on an intra-sample-scale beyond the analytical uncertainty (Fig. 8). We interpret this as evidence for isotope heterogeneity of the local S source within the sample (discussed below).

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All samples, except CB1, are inferred to be associated with a convergent margin setting, where granulite facies conditions were reached in deep arc lithosphere. There are four possible scenarios for the origin of the S sequestered by scapolite in granulites: A) sedimentary S-bearing phases (sulfides, sulfates) react to scapolite during high-T metamorphism; B) scapolite crystallizes as a primary igneous phase, C) scapolite receives its S component and isotope signature from an external fluid during high-grade metamorphism and D) primary magmatic S-bearing phases supply S for scapolite formation during the (re-)crystallization of the magmatic mineral assemblage under granulite facies conditions.

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Scenario A can be ruled out for CB1 and also for the NSW samples given the evidence that granulites from the Delegate breccia pipes and MVP represent recrystallized igneous cumulates (see section 5.1); we cannot, however, exclude the possibility that the protolith to sample MW1 had a prior upper crustal history. Stolz (1987) studied scapolite in mafic xenoliths from the McBride Province, Australia. He found idioblastic scapolite inclusions in plagioclase and clinopyroxene as well as garnet reaction coronas around scapolite grains.

Together with the compositional banding and P-T calculations, Stolz (1987) suggested that scapolite crystallized as a primary magmatic phase to subsequently form igneous cumulates.

Despite some similarities to the McBride xenolith samples (e.g., compositional banding), the absence of scapolite inclusions in other co-existing minerals argues against the scapolite grains in the present study forming by direct igneous crystallisation (scenario B). We found no obvious evidence for the infiltration of external S-rich fluids (scenario C) such as scapolite-rich veins, as for example seen in some deep-seated gabbros of the Kohistan Arc (Yoshino and Satish-Kumar, 2001) or the Bergen Arcs region of Western Norway (Porter and Austrheim, 2016). In the studied samples, scapolite is not restricted to certain zones, but
rather homogeneously distributed in the sample and in textural equilibrium with garnet and pyroxene. Furthermore, the mineralogy of the scapolite-bearing granulites suggests a “dry” environment for their formation, hence, the putative fluid would have been H₂O-deficient.

Scenario D, scapolite formation with S derived from reactions with primary magmatic sulfides, seems the most feasible process. It is generally agreed that the formation of S-rich scapolite requires oxidised conditions (Boivin and Camus, 1981; Stolz, 1987), which can be achieved by the infiltration of oxidizing fluids. However, Goldsmith (1976) proposed a simplified reaction between sulfides, oxides and silicates, where oxidation takes place via a ferrous-ferric equilibrium with primary high-T deep crustal minerals:

$$S^{2-}(\text{in sulfide}) + 4Fe_2O_3(\text{in magnetite}) \rightarrow SO^{2-}_4(\text{in scapolite}) + 8FeO(\text{in silicate})$$

If this reaction takes place, Goldsmith (1976) proposes that no oxygen transfer or other oxidizing agents (e.g., fluid) are required to form scapolite. However, fluxing of the lower crust by CO₂ exsolved from underlying crystallizing magmas (see Moecher et al., 1994), might play an important role in facilitating the above reaction. This is because such a process further stabilizes meionitic scapolite over other phases at high P and T (Goldsmith and Newton, 1977), thereby driving the reaction to the right and augmenting (re-)crystallization of the primary (igneous) mineral assemblage under granulite facies conditions (Harlov, 2012 and references therein). The presence of Fe-oxides with scapolite as observed in thin sections supports scapolite formation via the above reaction. Such processes would also explain the rather homogeneous scapolite distribution in the studied samples. Further support for the importance of the coupled redox reaction between ferrous/ferric Fe and sulfate/sulphide comes from the reversed process inferred by Porter and Austrheim (2016). These authors suggest that scapolite sulfate is reduced to sulphide during scapolite retrogression by the oxidation of ferrous Fe in the bulk rock and simultaneous formation of ferric minerals. In our samples, the absence of other S-bearing minerals besides scapolite, and provided that S was not introduced from an external source, could suggest that primary sulfides were the limiting
reaction agent, whereas as the other components (magnetite and Fe-bearing silicates, such as garnet and pyroxene), and potentially externally-derived CO₂ were present in excess. This would imply that the vast majority of S accumulated in the lower crust became subsequently recycled into scapolite. Isotopic fractionation of S would not be expected in this scenario, as supported by the narrow range of S isotope ratios within most samples and the absence of isotopic zonation within individual grains (but see below).

## 6.4. Isotope variability on a sample-scale

There are multiple possible causes for the observed S isotope variation within samples MVP99-2-12 and MVP99-2-05 (Fig. 8). It is possible that in- or out-diffusion on a mineral scale was variable within the sample. However, S isotope profiles through scapolite grains do not reveal any evidence for diffusion, such as e.g. systematic intragrain variations. We therefore regard diffusion as having a minor effect on the S isotope variation. In the case of MVP99-2-05 it could be argued that retrogression (i.e. the localised replacement of S by Cl) led to different isotope signatures in scapolite from grain to grain. This explanation cannot, however, apply to scapolite of sample MVP99-2-12, which lacks compositional zoning or localised Cl enrichment (Fig. 8).

Previous in situ S isotope studies on mantle rocks have also found within-sample δ³⁴S variations of several permil, as well as intra-grain variations outside of analytical uncertainty (e.g. Chaussidon et al., 1989, Giuliani et al., 2016). One explanation for the isotope variability observed in the latter study might be the incomplete isotope homogenization of monosulfide solid solution minerals during their equilibration at relatively low temperatures (~≤600°C) (Giuliani et al., 2016 and references therein). Chaussidon et al. (1989) observed S isotope variation in sulfide globules hosted in pyroxene from pyroxenite cumulates. The authors interpret the isotope variations between sulfide globules to be a direct result of mantle S isotope heterogeneity caused by migrating fluids in the mantle, which may have carried
different S-bearing species with variable S isotope signatures. Conceivably, such initial
isotope variations of primary magmatic sulfides, as reported by Chaussidon et al. (1989), are
mirrored in scapolite that subsequently formed by replacement of sulfide globules during
high-grade granulite facies metamorphism, in the presence or absence of an oxidizing fluid.

Despite the homogeneous spatial distribution of scapolite and the sample-scale isotope
heterogeneities, we cannot entirely exclude S addition to the system by percolating mantle-
derived fluids. This alternative scenario of an external S-source (as described by Austrheim,
2013), would, however, require an isotopically heterogeneous fluid, at least on a sample scale,
to form the observed S isotope heterogeneities between individual scapolite grains.

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6.5. Sulfur isotope signatures in scapolite and their implications for S in the lower crust

There is evidence that parts of the mantle underwent metasomatism, which is reflected
by distinct S isotope signatures of metasomatized peridotite samples (e.g., Wilson et al., 1996;
Giuliani et al., 2016, Fig. 9). However, the length scale of mantle metasomatism is difficult to
constrain, and it is unclear if metasomatized material, including fluids, might be introduced
into the lower crust. The measured S isotope signatures of scapolites from the studied
granulites fall within the range of S isotope ratios typically found in unaltered mantle rocks
(Fig. 9). Based on the initial sample set targeted here, there is no evidence that S-bearing
fluids sourced directly from subducted sediments (e.g., seawater sulfate) control scapolite
formation in the lower crust. If this were the case, δ³⁴S signatures of scapolite would be
expected to be significantly heavier (>>5 ‰ δ³⁴S). It is possible, however, that the S isotope
signature of slab-derived fluid was diluted and camouflaged by interaction with the sub-arc
mantle prior to percolating into the lower crust.

While we cannot exclude the possibility that slab-derived fluid interacted with the sub-
arc mantle, our results, together with those of previous studies show that the S isotope
signatures of granulites and pristine mantle rocks are indistinguishable (Figs. 2, 9). This
observation is in accordance with Moechner et al.'s (1994) findings that lower crustal scapolite acquires C from a mantle source, and challenges the concept that slab-derived fluids play a key role in universally enriching arc magmas in $ ^{34}S $. These findings might be further tested by studying the S isotope signatures of scapolites from deep crustal samples of other arc terranes.

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636 **6.6 Scapolite as a S source in the middle crust**

637 Porter and Austrheim (2016) recently found that S-rich scapolite formed in lower
638 crustal granulites becomes unstable during hydration and deformation at amphibolite facies
639 conditions. Release of S from the breakdown of scapolite can lead to the formation of sulfides
640 in the reaction zone, which might subsequently be mobilized by further hydration or
641 deformation. Levering and White (1964) determined that the scapolite reaction rims in
642 samples from the Delegate pipes consist of very fine-grained plagioclase. We found that these
643 scapolite reaction zones are typically associated with elevated K concentrations. The lack of
644 secondary sulfides in the reaction zones of scapolite suggests that S was fully mobilized after
645 peak-metamorphic scapolite breakdown (see also Porter and Austrheim, 2016). Scapolite
646 retrogression at mid-crust levels therefore likely contributes significant S to the overall S
647 budget, where the composition of the percolating fluids would be crucial for the transport of
648 this liberated S (Porter and Austrheim, 2016). This process is potentially of key importance
649 for the mobilization and concentration of chalcophile ore elements in the surrounding crust.
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650

651 **7. Conclusions**

652 We characterized an isotopically homogeneous S-rich scapolite (CB1), suitable as a
653 primary calibration standard for in situ SIMS analysis. SIMS S isotope analysis in scapolite is
654 independent of crystallographic orientation.

655 EPMA maps of the CB1 megacryst show oscillatory growth zonation, typical for the
656 formation of igneous minerals. This supports earlier findings of an igneous origin of scapolite
657 megacrysts from the Chaine de Puys. The elevated δ³⁴S (~+5 ‰) values of CB1 indicate that
658 the mantle source for intraplate volcanic activity in the Chaine de Puys contains a heavy S
659 isotope reservoir. Given the geodynamic setting, this reservoir is not directly associated with
660 coeval subduction processes, which is also supported by the mantle-like δ¹³C signature of
661 scapolite from the same location (Moecher et al., 1994). This suggests that the upper mantle
662 can in some cases be a reservoir of relatively heavy S, which in this case potentially
663 represents a memory effect of heavy S from previous Variscan subduction in the region
664 (Femenias et al., 2004).

665 Sulfur isotope signatures from lower crustal rocks represented by garnet-pyroxene and
666 two-pyroxene granulites from eastern Australia and south-eastern Ghana, both associated with
667 convergent tectonics, plot between ~-0.5 and ~+4 ‰ δ³⁴S. Sulfur isotope ratios of these
668 samples fall within the range of S isotope signatures reported from mantle rocks. These
669 results therefore provide no evidence for a major seawater component (i.e. slab-derived
670 fluids) in the lower crust beneath these areas, which we anticipate would lead to δ³⁴S values
671 of >>+4 ‰.

672 Our preferred model for S isotope variations in some of the studied samples involves
673 the presence of isotopically heterogeneous magmatic S-phases prior to metamorphism. This
674 would require that the initial magmas from which the primary S-phases separated were not
675 isotopically homogeneous. Isotope heterogeneity of mantle melt, as preserved in scapolite,
676 could be due to the melting of different mantle source domains, or to the local interaction of
677 mantle peridotite with externally-derived fluids. However, the putative external fluid flux in
678 this scenario was evidently insufficient to shift the S isotope signature of the source peridotite
679 away from the established mantle range. Subsequent granulite facies metamorphism of mantle
680 cumulates likely took place under relatively dry conditions (i.e. low fluid/rock ratios) as local
681 S isotope variations prevail.

682
683 Acknowledgements:

684 We greatly appreciate the substantial contribution of John Craven at the Edinburgh Ion
685 Microprobe Facility (EIMF) in sample preparation, technique development and in
686 conducting the SIMS S isotope analyses. Nicola Cayzer (EIMF) is thanked for the
687 measurements of the crystallographic orientations of the CB1 fragments. The authors
688 acknowledge the facilities, and the scientific and technical assistance of the Australian
689 Microscopy & Microanalysis Research Facility at the Centre for Microscopy,
690 Characterisation & Analysis, The University of Western Australia, a facility funded by the
691 University, State and Commonwealth Governments. The laser ablation system used in
692 this study was funded by the Australian Research Council (ARC LE150100013). We
693 thank the two anonymous reviewers for their constructive comments and the editor for
694 handling this paper. This work was supported by Swiss National Science Foundation
695 grant P2SKP2_155067 to J. Hammerli and an ARC fellowship (FT100100059) to T.
696 Kemp.
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698

699

700 **Figure/Table captions**

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702 **Figure 1:** Simplified sketch of hypothesised S recycling in subduction systems with
703 references to key studies of the respective reservoirs. Examples for S isotopes studies in lower
704 crustal rocks (granulites and cumulates) include Hoefs et al. (1987) and Iyver et al. (1992).
705 The dotted box represents the mantle-crust interface, for which comprehensive datasets are
706 lacking.

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708 **Figure 2:** Compilation of S isotope signatures (compared to VCDT) in different reservoirs.

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710 **Figure 3:** Scans of polished thin sections of the studied scapolite-bearing granulite samples.
711 MVP99-2-13 (A), MVP99-2-12 (B) and MW1 (C) are garnet-pyroxene-(hornblende)
712 granulites. MVP99-2-05 (D) and 10438B (E, F) are two-pyroxene granulites. F)
713 Microphotograph in plane-polarized light of sample 10438B, where the white arrows point to
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the alteration rims of scapolite. Opx=orthopyroxene, Scp scapolite, Pl=plagioclase;

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Cpx=clinopyroxene; Grt=garnet.

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**Figure 4:** Quantitative S map of CB1. The larger upper piece represents the A-plane

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(perpendicular to the A-axis) and the smaller lower piece represents the C-plane (a slice along

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the C-axis) of the crystal. Stereonets show the poles for each crystal plane, plotted relative to

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the 001 plane.

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**Figure 5.** Quantitative EPMA S maps of representative scapolite grains. Scapolite grains

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from sample MVP99-2-12 and MVP99-2-13 show homogeneous S concentrations in

724

individual grains. Scapolite in sample MVP99-2-05 and 10438B show S depletion towards

725

their rims. The mapped scapolite grain from sample MW1 shows an irregular S distribution

726

without distinct core-rim correlations.

727

728

**Figure 6:** Chondrite normalized (McDonough and Sun, 1995) REE analyses of scapolite by

729

LA-ICP-MS.

730

731

**Figure 7:** Reproducibility of S isotope measurements from scapolite CB1, as assessed by 224

732

individual SIMS analyses, where the measured ³⁴S/³²S ratios are normalised to the whole

733

grain δ³⁴S value of CB1 (+5.08 ‰). The weighted mean value is shown as the black

734

horizontal line. Black bars represent analyses from a traverse perpendicular to the C-axis of

735

CB1 (average δ³⁴S = +4.91 ‰ ± 0.44 (2SD)), dark grey bars represent measurements from a

736

traverse perpendicular to the A-axis (average δ³⁴S = +5.08 ‰ ± 0.41 (2SD)), and light grey

737

bars represent the combined analyses over 4 analytical sessions on various CB1 crystal

738

fragments (average δ³⁴S = +5.05 ‰ ± 0.41 (2SD)). All data points include 2SD uncertainties,
739 which combine in-run (measured) uncertainties and the uncertainty of the bulk grain reference
740 values, summed in quadrature.

741
742 **Figure 8:** Average S isotope composition of individual scapolite grains of the studied
743 samples. Error bars represent 2SD.

744
745 **Figure 9:** Comparison of S isotope signatures of lower crustal rocks (Hoefs et al., 1981; Iyver
746 et al., 1992), mantle-rocks (Ionov et al., 1992; Wilson et al., 1996), and S isotope analysis in
747 scapolite from granulites (error bars are 2SD). The S isotope variation in Ionov et al. (1992)
748 might be a result of isotope fractionation due to melt extraction (e.g., Chaussidon et al., 1989).
749 *presence of secondary scapolite, **metasomatized.

750
751 **Table 1:** Representative EPMA data of scapolite for each studied sample. The number of ions
752 is calculated on the basis of 12 (Si, Al). CO₂ is calculated on the assumption that the T space
753 is fully occupied by C, S and Cl, hence C=1–S–Cl. Meionite equivalent% is expressed as
754 Me%=100[∑(divalent cations)/4)] and equivalent anorthite content as EqAn = 100(Al–3)/3.
755 CB1(bright) and CB1(dark) refer to the bright and dark zonation shown in Fig. 4.

756
757 **References**
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762
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766
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770

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774

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917

918

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922

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925

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929

930 Liotard, J.M., Briot, D., Boivin, P. 1988 Petrological and geochemical relationships between pyroxene megacrysts and associated alkali-basalts from Massif Central (France). Contrib. Mineral. Petrol. 98, 81–90.

933

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936

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940

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1031
<figure><img src="image_3.png" /></figure>

a) e.g., Alt et al. (1995); Alt et al. (2012)

b) e.g., Farquhar et al. (2010) and references therein

c) e.g., Alt et al. (2012); Evans et al. (2014)

d) e.g., Chaussidon et al. (1989); Cabral et al. (2013); Labidi et al. (2013)

e) e.g., Wilson et al. (1996); Ionov et al. (1992)

f) e.g., Sasaki & Ishihara (1979); Alt et al. (1993); de Hoog et al. (2001)

g) e.g., Sakai et al. (1982)
<figure><img src="image_4.png" /></figure>

*Marine sulfates/sedimentary pyrite:* Farquhar et al. (2010) and therein. *OIBs:* Sakai et al. (1984); Torssander (1989); Chaussidon et al. (1989); Cabral et al. (2013). *MORBs:* Kanehira et al. (1973); Sakai et al. (1984); Labidi et al. (2012, 2013). *Mantle xenoliths:* Chaussidon et al. (1989); Chaussidon & Lorand (1990); Kyser (1990); Ionov et al. (1992); Wilson et al. (1996). *Arc Volcanics:* Ueda & Sakai (1984); Alt et al. (1993); Woodhead et al. (1987); de Hoog et al. (2001), Marini et al. (2011). *Continental tholeiites:* Schneider et al. (1970). *Continental alkali-rich rocks:* Schneider et al. (1970), Harmon et al. (1987)
<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>

<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /></figure>

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>

<figure><img src="image_13.png" /></figure>

S wt.%

<figure><img src="image_14.png" /></figure>

1 mm
S wt.%

<figure><img src="image_15.png" /></figure>

MVP99-2-12

S wt.%

<figure><img src="image_16.png" /></figure>

MVP99-2-05

3

2

1

S wt.%

<figure><img src="image_17.png" /></figure>

MVP99-2-13

S wt.%

<figure><img src="image_18.png" /></figure>

10438B

3

2

1

S wt.%

<figure><img src="image_19.png" /></figure>

MW1

3

2

1

0
<figure><img src="image_20.png" /></figure>
<figure><img src="image_21.png" /></figure>

Mean = 5.069±0.026 [0.51%] 95% conf.

Wtd by data-pt errs only, 0 of 224 rej.

MSWD = 0.96 (error bars are 2s)
<figure><img src="image_22.png" /></figure>
<figure><img src="image_23.png" /></figure>
<table><thead><tr><th>Sample</th><th>10438B</th><th>MVP99-2-13</th><th>MVP99-2-05</th><th>MVP99-2-12</th><th>MW1</th><th>CB1(bright)</th><th>CB1(dark)</th><th>CB1 (n=16)<br>average 2SD</th></tr></thead><tbody><tr><td>(wt.%)</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>SiO<sub>2</sub></td><td>43.85</td><td>46.15</td><td>44.59</td><td>46.52</td><td>47.66</td><td>44.97</td><td>45.59</td><td>45.48 0.60</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>27.63</td><td>26.37</td><td>27.93</td><td>26.00</td><td>25.43</td><td>27.24</td><td>26.78</td><td>26.16 1.03</td></tr><tr><td>FeO</td><td>0.26</td><td>0.33</td><td>0.24</td><td>0.39</td><td>0.19</td><td>0.35</td><td>0.42</td><td>0.43 0.13</td></tr><tr><td>CaO</td><td>18.55</td><td>16.89</td><td>18.40</td><td>16.02</td><td>15.46</td><td>16.89</td><td>16.64</td><td>16.57 0.26</td></tr><tr><td>MgO</td><td>0.17</td><td>0.49</td><td>0.22</td><td>0.44</td><td>0.07</td><td>0.09</td><td>0.09</td><td>0.09 0.02</td></tr><tr><td>Na<sub>2</sub>O</td><td>2.70</td><td>3.41</td><td>2.99</td><td>3.84</td><td>4.76</td><td>3.85</td><td>4.07</td><td>3.80 0.23</td></tr><tr><td>K<sub>2</sub>O</td><td>0.06</td><td>0.13</td><td>0.12</td><td>0.20</td><td>0.07</td><td>0.24</td><td>0.22</td><td>0.23 0.05</td></tr><tr><td>Cl</td><td>0.01</td><td>0.03</td><td>0.04</td><td>0.07</td><td>0.05</td><td>0.20</td><td>0.24</td><td>0.23 0.03</td></tr><tr><td>SO<sub>3</sub></td><td>5.09</td><td>4.13</td><td>3.28</td><td>4.58</td><td>4.92</td><td>3.96</td><td>3.73</td><td>3.86 0.15</td></tr><tr><td>CO<sub>2</sub> (calc)</td><td>1.85</td><td>2.40</td><td>2.88</td><td>2.11</td><td>1.98</td><td>2.27</td><td>2.36</td><td>2.24 0.14</td></tr><tr><td>Total</td><td>100.18</td><td>100.33</td><td>100.68</td><td>100.16</td><td>100.59</td><td>100.08</td><td>100.14</td><td>99.11 1.65</td></tr><tr><td>-O=Cl</td><td>0.00</td><td>0.01</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.05</td><td>0.05</td><td>0.05 0.01</td></tr><tr><td>Total Cl corrected</td><td>100.18</td><td>100.32</td><td>100.67</td><td>100.14</td><td>100.58</td><td>100.03</td><td>100.08</td><td>99.06 1.66</td></tr><tr><td>Total (12/Si+Al)</td><td>9.43</td><td>9.34</td><td>9.30</td><td>9.34</td><td>9.29</td><td>9.35</td><td>9.34</td><td></td></tr><tr><td>Si</td><td>6.89</td><td>7.17</td><td>6.90</td><td>7.23</td><td>7.37</td><td>7.00</td><td>7.09</td><td></td></tr><tr><td>Al</td><td>5.11</td><td>4.83</td><td>5.10</td><td>4.77</td><td>4.63</td><td>5.00</td><td>4.91</td><td></td></tr><tr><td>Fe</td><td>0.03</td><td>0.04</td><td>0.03</td><td>0.05</td><td>0.02</td><td>0.05</td><td>0.05</td><td></td></tr><tr><td>Ca</td><td>3.12</td><td>2.81</td><td>3.05</td><td>2.67</td><td>2.56</td><td>2.82</td><td>2.77</td><td></td></tr><tr><td>Mg</td><td>0.04</td><td>0.11</td><td>0.05</td><td>0.10</td><td>0.02</td><td>0.02</td><td>0.02</td><td></td></tr><tr><td>Na</td><td>0.82</td><td>1.03</td><td>0.90</td><td>1.16</td><td>1.43</td><td>1.16</td><td>1.23</td><td></td></tr><tr><td>K</td><td>0.01</td><td>0.03</td><td>0.02</td><td>0.04</td><td>0.01</td><td>0.05</td><td>0.04</td><td></td></tr><tr><td>Cl</td><td>0.00</td><td>0.01</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.05</td><td>0.06</td><td></td></tr><tr><td>S</td><td>0.60</td><td>0.48</td><td>0.38</td><td>0.53</td><td>0.57</td><td>0.46</td><td>0.43</td><td></td></tr><tr><td>C</td><td>0.40</td><td>0.51</td><td>0.61</td><td>0.45</td><td>0.42</td><td>0.48</td><td>0.50</td><td></td></tr><tr><td>Me%</td><td>79.30</td><td>73.83</td><td>77.27</td><td>70.20</td><td>64.35</td><td>70.45</td><td>69.16</td><td></td></tr><tr><td>EqAn#</td><td>70.48</td><td>60.99</td><td>69.90</td><td>58.90</td><td>54.47</td><td>66.66</td><td>63.67</td><td></td></tr><tr><td>XS</td><td>0.6</td><td>0.48</td><td>0.38</td><td>0.53</td><td>0.57</td><td>0.46</td><td>0.43</td><td></td></tr><tr><td>XC</td><td>0.4</td><td>0.51</td><td>0.61</td><td>0.45</td><td>0.42</td><td>0.48</td><td>0.50</td><td></td></tr></tbody></table>