Sulfides, Native Silver, and Associated Trace Minerals of the Skaergaard

Home , Skaergaard intrusion

8 Sulfides, native silver, and associated trace minerals of the Skaergaard 9 Intrusion, Greenland: Evidence of late hydrothermal fluids 10 11 Ben Wernette1†, Peishu Li1, and Alan Boudreau1 12 1Division of Earth and Ocean Sciences, Duke University, Durham, N.C., 27705, USA 13 †Corresponding Author: [email protected] 14 15

24 Abstract

25 Sulfide assemblages and accessory phases throughout the Skaergaard were characterized

26 to better understand the role of magmatic volatiles in modifying the Skaergaard metal budget and

27 distribution. Sulfides in and below the Platinova Reef are readily replaced by low-Ti magnetite;

28 in thin section, the ratio of the low-Ti magnetite mode)/(sulfide mode) reaches maximum

29 values at the Platinova Reef. Sulfide assemblages below the Reef are accompanied by trace

30 quantities of clinopyroxene, orthopyroxene, biotite, apatite, and minor calcite. Native Ag,

31 commonly accompanied by trace amounts of Cl, occurs both in and below the Platinova Reef.

32 Evidence of coexisting precious metal + brine assemblages exists where native metals are

33 accompanied by sylvite ± halite and Ag is accompanied by Ag-halides. Ag occurrences in the

34 Platinova Reef are of irregular morphology with trace Cl ± S ± calcite. Further evidence

35 supportive of a metal + brine assemblage is observed where Ag + quartz are found in an apparent

36 fluid inclusion consisting of Na, Si, Cl, Ca, K, and S.

37 In agreement with earlier studies, the observed assemblage is consistent with the

38 Skaergaard being a S-poor intrusion with S continually lost during cooling and crystallization.

39 Partitioning of Ag into an exsolving fluid phase is a function of Cl concentration. Numerical

40 modeling suggests that, in a sulfide bearing system, residual Ag concentrations and

41 concentrations in the exsolved fluid are most affected at the point where sulfide is lost to a

42 separating volatile fluid phase. It is suggested that, owing to the low S nature of the Skaergaard

43 system, fractional crystallization produced enrichment of Ag in the interstitial silicate much

44 higher than normal due to delayed sulfide saturation. As this interstitial liquid evolved, Ag was

45 lost to an exsolved volatile phase of high salinity and migrated upward along grain boundaries

46 and fluid pathways. A similar process likely occurred for Au and other elements with high

47 affinities for Cl such as platinum-group elements.

49 1. Introduction

50 In layered mafic intrusions (LMIs), evidence for the existence of volatile phases is

51 preserved in fluid inclusions, apatite composition, and halogen geochemistry (e.g., Hanley et al.

52 2008; Boudreau and McCallum 1989; Boudreau et al. 1997). Of debate however, is whether

53 these volatile fluids, play a significant role in the processes responsible for generating, or

54 modifying, the stratiform platinum-group element (PGE)-sulfide ore bodies for which LMIs are

55 known. Conventionally, these deposits are thought to form as the result of gravitationally driven

56 downward movement of immiscible sulfide liquid droplets. Large sulfide liquid/silicate liquid

57 partition coefficients (e.g., Brenan and Mungall 2014) and large silicate liquid/sulfide liquid

58 mass ratios (R-factor, Campbell and Naldrett 1979) are thought to allow for the extreme

59 enrichment of transition and noble metals as well as PGE in immiscible sulfide droplets. Indeed,

60 experimental results confirm that these are reasonable assumptions (e.g., Fleet et al. 1993).

61 Conversely, others have suggested that stratiform mineralization is the result of upward

62 moving exsolved fluid, rich in halides, allowing for the efficient transport of transition and

63 precious metals vertically through an igneous body (e.g., Boudreau and McCallum 1992).

64 Theoretical work (Shinohara 1994) supports this model and experimental work performed on

65 rhyolitic and granitic systems (Candela and Holland 1984; Simon et al. 2008; Frank et al. 2011)

66 and theoretical work done on LMIs (Boudreau and McCallum 1992; Muerer et al. 1999) suggest

67 the importance of exsolved halide-bearing volatiles. Together these end-member models are

68 referred to as the “orthomagmatic” and “hydromagmatic” models, respectively.

69 Complicating our understanding of LMIs is that the larger intrusions (i.e., Bushveld and

70 Stillwater Complex) show evidence of multiple magma injections and prolonged cooling

71 histories for which the preservation of original igneous textures and chemistry is uncertain. The

72 small, single-pulse intrusion of the Paleogene Skaergaard Complex of southeast Greenland

73 contains a zone enriched in Au and the PGEs known as the Platinova Reef. The Skaergaard

74 Intrusions provides a unique opportunity to examine for evidence supportive of one or both of

75 the competing genetic models without the complication of potential magma mixing events. In

76 this study, we examine sulfide assemblages and their associated phases and characterize the

77 occurrence of native silver to better understand the role of exsolved volatiles in the distribution

78 and modification of metals in LMIs.

80 2. Geologic Background and Summary of Previous Studies

81 2.1 Skaergaard Complex

82 For extensive reviews, readers are directed to the early studies of Wager and Deer (e.g.,

83 Wager and Deer 1939) and the more recent work of Nielsen (2004) and references therein.

84 Briefly, the Skaergaard Complex is a ~ 55 Ma (Brooks and Gleadow 1977; Hirschman et al.

85 1997) intrusion located in southeast Greenland (Fig. 1). The intrusion is related to the Paleogene

86 rifting of Greenland from Eurasia (Nielsen 1975). The Complex is hosted uncomfortably in

87 Archean gneiss and is principally comprised of the Layered, Upper Border, and Marginal Border

88 Series. The Layered Series is thought to have crystallized from the floor of the intrusion while

89 the Upper Border and Marginal Border Series crystallized from the roof and the walls of the

90 intrusion, respectively. The Layered Series is subdivided into the Lower, Middle, and Upper

91 Zone according to the presence of index minerals. The Platinova Reef is a diffuse (~ 60 m in

92 thickness, Anderson et al. 1998) metal-rich zone with distinct PGE and Au rich subzones hosted

93 in the upper part of the Middle Zone.

94 Wager et al. (1957) describe a Middle Zone Cu-rich sulfide assemblage consisting of

95 chalcopyrite and bornite. Anomalously high Cu/S ratios observed throughout the stratigraphy

96 have been attributed to either early shallow-level degassing and S loss (Li and Boudreau 2017),

97 late S loss to hydrothermal fluids (Andersen 2006), or anomalous Cu/S ratios in the Skaergaard

98 parental magma source region (Keays and Tenger 2016). Andersen (2006) concluded that sulfide

99 texture and mineralogy in the Lower Zone is well explained by an exsolved hydrothermal fluid

100 replacing preexisting sulfides with low-Ti magnetite and redistributing Cu, S, and precious

101 metals to different stratigraphic levels. Li and Boudreau (2017) arrived at a similar conclusion

102 after conducting a modal analysis of sulfides found in the Lower Zone and Marginal Border

103 Series. Conversely, Godel et al. (2014) note that LZ sulfides are generally not accompanied by

104 low-Ti magnetite, ultimately suggesting that late hydrothermal fluids did not oxidize magmatic

105 sulfide assemblages.

106 Detailed studies of precious metal distribution through the Skaergaard stratigraphy abound

107 (Holwell et al. 2016; Keays and Tenger 2016; Nielsen et al. 2015; Godel et al. 2014). These

108 studies make little to no mention of Ag, understandably focusing largely on the volumetrically

109 significant Cu, Au, Pd, and Pt. In the Platinova Reef, researchers have observed the occurrence

110 of Ag in Au-Cu-Ag alloys (Bird et al. 1991 and Cabri et al. 2005) as well as Au-PGE alloys

111 (Andersen et al. 1998). Additionally, Holwell et al. (2014) note silver concentrations of ~ 250

112 ppm in sulfides from the Triple Group. Moreover, Ag has been reported in Au-Cu-Ag alloys

113 found in marginal basement schists (Andersen et al. 2017).

114

115 3. Sampling and Methods

116 Samples were collected in 1990 by A. Boudreau. Two samples, SK90-5 and SK90-13,

117 were published in a previous study (Li and Boudreau 2017) and were re-examined and included

118 in this study for completeness. Individual polished thin-sections were systematically scanned

119 using the Cameca CAMEBAX microprobe at Duke University. Sulfide assemblages, and

120 accessory phases were characterized using energy dispersive spectrometry (EDS) and X-ray

121 composite maps. The same methods were used to identify and characterize the occurrences of

122 native Ag. Sulfides, Ag, and accessory phases were imaged and imported into the photo

123 processing software IMAGE-J (Schneider et al. 2012) for area and perimeter measurements.

124 Because trace assemblages are those that do not occur in large quantities, extreme caution was

125 used when characterizing occurrences of native Ag. Ag without associated trace minerals or

126 elements (e.g., calcite, Cl, K, Na, S) were only included in the dataset when found with other late

127 crystallizing phases (apatite) or associated with silver halides. This filtering removed

128 approximately two thirds of the original observations of Ag.

129

130

131

132

133 4. Results

134 4.1 Summary of Stratigraphic Trends

135 As summarized in Fig. 2 and discussed in detail below, stratigraphic trends show a

136 general decrease in chalcopyrite (and a corresponding increase in bornite) from the Lower Zone

137 through the Middle Zone before chalcopyrite and digenite join the sulfide assemblage in the

138 Upper Zone. Similarly, the relative modal abundance of low-Ti magnetite as a fraction of the

139 sulfide assemblage is highest in the Middle Zone. As noted above, the Ti-free magnetite has been

140 previously interpreted to be replacements of the original sulfide assemblage as the result of S-

141 loss (e.g., Andersen 2006; Li and Boudreau 2017). The sulfide-magnetite assemblages are seen

142 at both grain boundaries and within primocryst silicate minerals. Further, the relative proportion

143 of accessory minerals (orthopyroxene, clinopyroxene, calcite, biotite, and apatite; trace

144 clinopyroxene is identified as having no exsolution lamellae and little iron content, see Fig. 4a) is

145 greatest in those samples from the Lower Zone. Occurrences of native Cu, Ag, and other

146 precious metals are generally restricted to Middle and Lower Zone samples. Similarly, halite and

147 sylvite are observed only in those samples in and below the Middle Zone. Native silver is

148 commonly associated with trace quantities of Cl. Broadly, the size range of individual silver

149 occurrences increases up stratigraphy, accompanied by a general increase in the total number of

150 Ag grains observed.

151 Sulfide morphology varies systematically through the stratigraphy with the area of

152 individual sulfide assemblages being greatest for those occurrences in the MZ (Fig. 3a).

153 Similarly, the morphologic parameter elongation, or the deviation of a given sulfide from

154 circular, is greatest for MZ samples (Fig. 3b). Silver does not display any systematic behavior

155 perhaps owing to their rarity, but overall occurrence is greatest for those samples from the MZ.

156 A comprehensive description of the individual sulfide (n = 165) and silver (n = 12) occurrences

157 are presented in supplementary tables 1 and 2, respectively.

158

159 4.2 General Description of Sulfide Assemblages and Associated Trace Minerals

160 LZa: Sulfides, mainly chalcopyrite and bornite, are typically found interstitial to olivine

161 and plagioclase. Pentlandite is observed in limited quantities and generally comprises a small

162 proportion of individual assemblages (~ average of 4%). Sulfides are commonly non-uniform in

163 morphology and are frequently replaced in part by low-titanium magnetite (Fig. 4a, b). Native

164 copper is observed where sulfides are extensively replaced by low-titanium magnetite (Fig. 4a).

165 Hydrous minerals, such as biotite and amphibole, are observed truncating sulfides or mantling

166 sulfide assemblages (Fig. 4b, c). Calcite is found in limited abundance mantling sulfides with no

167 evidence of other low or high temperature silicate alteration. When observed, calcite is rounded

168 and commonly no greater than 20 µm in size. Anhedral clino- and orthopyroxene are found

169 mantling sulfides. When adjacent, plagioclase is increasingly calcic near sulfide assemblages

170 (Fig 4a, b, c).

171 LZc: Observed sulfide proportions by area average 9% chalcopyrite and 91% bornite.

172 Sulfides are found both within silicates and at silicate ground boundaries. Low-titanium

173 magnetite is less abundant relative to LZa. Biotite truncates silicate grain boundaries or replaces

174 sulfides. Calcite was observed in a single sulfide assemblage. As for LZa, plagioclase becomes

175 increasing calcic towards sulfide assemblages.

176 MZ: The sulfide assemblage consists entirely of bornite with variable amounts of low-

177 titanium magnetite. Sulfides are interstitial to, or found entrained within, silicates. Sulfide

178 morphology is typically irregular and is positively correlated with the abundance of secondary

179 silicates or low-Ti magnetite. Biotite and amphibole are observed in limited quantities mantling

180 sulfides and calcite is only associated with native silver or silver-halides. At pyroxene grain

181 boundaries, olivine is observed enclosing sulfide assemblages and in some cases low-titanium

182 magnetite as well (Fig. 4d). When present, this is broadly accompanied by an increase in

183 plagioclase anorthite content in plagioclase towards the sulfide assemblage and associated

184 olivine.

185 UZa: In order of relative abundance bornite, chalcopyrite, and digenite are 82, 13, and

186 5% of the observed relative modal abundance, respectively. Low-titanium magnetite is observed

187 replacing sulfides in quantities similar to LZc. Apatite is regularly associated with sulfides,

188 typically surrounding the periphery of the mineral. Sulfides are notably non-uniform in

189 composition with copper-rich, sulfur-poor areas (Fig. 4e). Anhydrite is observed in limited

190 abundance where sulfides are in direct contact with plagioclase. Additionally, plagioclase

191 becomes increasingly calcic towards sulfide assemblages.

192 UZb: Bornite and minor digenite comprise the entire sulfide assemblage. Sulfides are

193 commonly rounded and found at silicate-silicate or silicate-oxide grain boundaries and less

194 commonly within high-titanium magnetite. Additionally, sulfides are reliably found interstitial to

195 modally abundant apatite. Low-titanium magnetite is seen in minor abundance and biotite is

196 observed in only one assemblage.

197

198

199

200 4.3 Occurrence of Precious Metals and Fluid Inclusions

201 Below the Middle Zone, native Ag occurs as discrete rounded grains within major

202 minerals such as plagioclase, interstitial to late phases such as clinopyroxene and apatite (Fig.

203 5a), or at primocryst grain boundaries (Fig. 5b, c). Within the Middle Zone, silver occurs both as

204 well-rounded grains and amoeboidal, or irregular in shape, morphologies at grain boundaries

205 (Fig. 5d). Those occurrences of irregular morphology are typically accompanied by trace

206 amounts of chlorine ± sulfur ± calcite (Fig. 5d, e) and silver halides (Fig. 5d). One occurrence of

207 note, which is currently interpreted as a ruptured fluid inclusion, is native silver found within

208 clinopyroxene with quartz and trace quantities of Na, Si, Cl, Ca, K, and S (Fig. 6). Less common,

209 silver occurs with gold intergrowths and trace copper with notable trace quantities of Cl (Fig. 5e,

210 f). Silver was not observed above UZa.

211 Native copper is seen where significant replacement, or loss, of sulfur from sulfide

212 minerals has occurred (LZa). Copper “colloids” are observed at sulfide grain boundaries without

213 clear evidence of sulfur loss or silicate replacement in LZa. Trace quantities of native Cu are

214 observed in pyroxenes and amphiboles (LZc and MZ, respectively). Additionally, Ag-Zn alloys

215 are observed with native copper in clinopyroxene in the MZ with significant trace Cl (Fig. 5e).

216 Native platinum is observed within high-Ti magnetite (LZa). Further, platinum sulfides

217 (likely skaergaardite) and Pt-Au alloys are observed in the MZ. Gold is observed as intergrowths

218 with silver in the MZ and singular occurrences with halite and sylvite in plagioclase in LZa.

219 Below the Middle Zone, fluid inclusions are rarely observed, however the former

220 presence of a high temperature fluid is inferred by the presence of sylvite inclusions in

221 primocrysts (LZa and LZc) and halite crystals at silicate grain boundaries (LZa). Within the

222 Middle Zone, fluid inclusions are common and observed as singular occurrence in pyroxene or

223 trails in olivine (Fig. 7). Inclusions observed in pyroxene and, inadvertently opened late during

224 thin section preparation, are commonly no greater than 10 µm in diameter and contain Ag, Cu-

225 Zn alloys, and trace quantities of Cl, Na, Ca, and K. Fluid inclusions observed in olivine can be

226 categorized into three principle groups that commonly form inclusion trails, consisting of

227 elongate or oblate shaped fluid inclusions less than 10 µm in diameter. One population of

228 inclusions parallels olivine fracture surfaces, while the other two are not associated with any

229 present fracture surface and differ only in size and concentration (or density) of inclusions.

230 Broadly, inclusions in olivine are composed of a single phase; that is to say that no salts or

231 exsolved phases are observed within these inclusions.

232

233 5. Discussion

234 5.1 Sulfide Mineralogy and Associated Minerals

235 Our observations are consistent with work reported by other researchers (Andersen 2006;

236 Li and Boudreau 2017). Specifically, sulfides in and below UZa are commonly replaced by low-

237 Ti magnetite as the result of S loss via ƒO2-neutral reactions:

238 1 CuFeS2 + 1.15 H2O → 0.27 Fe3O4 + 0.19 Cu5FeS4 + 1.15 H2S + 0.04 SO2 (1) 239 cpy vapor mt bn --- vapor --- 240 and

241 1 CuFeS2 + 1.11 H2O + 2 HCl + SO2 → 0.33 Fe3O4 + 1 CuCl2 + 2.11 H2S (2) 242 cpy ------vapor ------mt --- vapor --- 243 244 (Li and Boudreau 2017). Evidence of sulfur loss is ubiquitous from LZa through UZa.

245 Additionally, the abrupt decrease in secondary, or hydrous, minerals above the MZ is consistent

246 with loss on ignition (LOI) data reported by Keays and Tenger (2016) where values approach 1

247 wt. % below UZa but are closer to 0.2 wt. % in, and above, UZa. Similarly, the enrichment in the

248 An content in plagioclase suggests silica and Na loss to a silica-under saturated fluid via:

+ 2+ + 249 2 CaNaAlSi2.5O8 + 4 H2O + 4 H → CaAl2Si2O8 + 3 H4SiO4 + Ca + 2 Na (3) 250 plagioclase -- vapor -- anorthite ------vapor ----- 251

252 is inferred as calcium content, thus An, increase toward sulfide assemblages (LZa through UZa,

253 see Fig. 4a, b, c, and f). Further, silica loss can produce olivine at pyroxene grain boundaries via

254 the reaction:

255 2 (Mg0.5Fe0.5)SiO3 + 2 H2O → (MgFe)SiO4 + H4SiO4 (4) 256 orthopyroxene vapor olivine vapor 257 258 259 Similar observations have been made at the 2.7 Ga Stillwater Complex (Meurer et al. 1997)

260 where discordant troctolites were explained by silica loss to reactive fluids.

261

262 5.2 Silver and an Exsolved Volatile Phase

263 The reactants involved in equations 1 through 4 require a sulfide-undersaturated volatile

264 phase to produce the observed assemblages. Direct evidence for a fluid include the observed

265 fluid inclusions with associated silver and quartz that retain trace quantities of Ca, Na, Cl, and S

266 (products seen in reactions in Section 5.1) and MZ clinopyroxene host Ag-Zn alloys with Cu and

267 trace Cl. Moreover, the association of Ag and Cl in MZ and LZ samples suggests their

268 concentration mechanism is the same. Further, native Au with proximal halite and sylvite (LZa,

269 Fig. 8) is supportive of coexisting metal + brine assemblage.

270 Magmatic volatile phases have been argued as an important control on the distribution of

271 precious metals in epithermal systems (e.g., Heinrich et al. 2004), Cu-Au porphyry systems (e.g.,

272 Candela 1989; Heinrich et al. 2004; Simon et al. 2008), and layered mafic intrusions (e.g.,

273 Boudreau and McCallum 1992; Meurer et al. 1999, Hanley et al. 2008). Previous studies of the

274 Skaergaard Complex have invoked magmatic volatile phases to explain dendritic anorthosites

275 and pegmatites found in UZa (Sonnenthal 1998), oxidized sulfides found throughout the

276 stratigraphy (Andersen 2006), and unusually high Cu/S ratios (Anderson 2006). Further, Godel

277 et al. (2014) assert that Au textures observed in the uppermost zone (Pd5) of the Platinova Reef

278 are not consistent with a magmatic origin.

279 Several researchers (e.g., Simon et al. 2008; Frank et al. 2011; Zajacz et al. 2013; Yin and

280 Zajacz 2018) have studied the behavior of silver in sulfur poor granitic and rhyolitic systems.

281 Theoretical and experimental studies highlight the importance of secondary boiling in the

282 generation of a magmatic volatile phases and the pathways through which they migrate (Huber et

283 al. 2013 and Candela and Holland 1984, respectively). Zajacz et al. (2008) show that Ag fluid-

푓푙푢푖푑/푙푖푞푢푖푑 284 silicate liquid partition coefficient (퐷퐴푔 ) increase linearly with the molarity of Cl, mCl,

푓푙푢푖푑/푙푖푞푢푖푑 285 in the exsolved fluid phase such that 퐷퐴푔  4 x mCl. Indeed, Cl has an affinity for the

푓푙푢푖푑/푠푖푙푖푐푎푡푒 286 aqueous phase with 퐷퐶푙 values between 20 and 40 (Holland, 1972) or between .9 and

287 6 in molten basalts (Webster et al. 1999). A pertinent question is whether an exsolved Cl-bearing

288 fluid phase can transport significant quantities of Ag via:

289 AgOmelt + HCl → AgCl + 0.5H2O (5) 290 vapor ----- vapor ----- 291 292

293 (Simon et al. 2008) and by inference, other Cl soluble metals such as Au, Cu, and Zn. Yin and

294 Zajacz (2018) suggest that at higher temperatures (~ 900°C), the neutral NaAgCl2 is the favored

295 Ag species in Cl-bearing exsolved fluids via the modified H2-neutral reaction:

296 AgOmelt + NaCl + HCl + 0.5SO2 → NaAgCl2 + 0.5H2SO4 (6) 297 ------vapor------vapor---- 298 299 300 Na cations can come from reaction (3) suggesting that the conversion of plagioclase to anorthite

301 by a silica-undersaturated fluid is intimately related to the transport of Ag as illustrated by

302 reaction (6).

303 Less well understood is the behavior of Ag in basaltic systems. Early studies investigated

304 the concentration of Ag and other precious metals in MORBs to understand their potential as

305 protoliths to greenstone Au deposits (Keays and Scott 1978). Recent experimental results have

306 allowed for the determination of Ag sulfide-silicate liquid partition coefficients on the order of

307 1000 (Li and Audetet 2012). Indeed, the investigation of sulfide globules in MORB glasses lead

308 to the determination of a similar Ag sulfide-silicate liquid partition coefficient of 1138 ± 245

309 (Patten et al. 2013). Workers have observed that MORB Cu/Ag values remain constant at ~ 3500

310 with decreasing Mg#, suggesting the two elements behave similarly throughout MORB

311 fractionation (Jenner 2017). Additionally, similar Cu/Ag values are observed in arc and back-arc

312 environments prior to sulfide saturation (Jenner et al. 2010). Further support of the associated

313 nature of Cu and Ag is found in mantle peridotites where Cu/Ag values are nearly identical to

314 those observed in MORBs (Wang and Becker 2015). Zajacz et al. (2013) determine Ag

315 solubilities for basaltic liquids to be ~ 11 and ~ 5.5 ppm for sulfide-bearing and sulfide-free

316 compositions, respectively. Because of the relatively late S-saturation in the S-poor Skaergaard

317 system, it is appropriate to investigate the partitioning of Ag into coevolving vapor in basalts to

318 better understand Ag behavior in exsolved hydrothermal fluids in igneous intrusions.

319

320 5.3 Temperature, Timing, and other Considerations

321 Temperature constraints might allow for better understanding of Ag mobilization during

322 the Skaergaard cooling history. Physical relations such as the interstitial nature of Ag in LZ

323 samples and close proximity to grain boundaries suggests that Ag precipitation occurred late in the

324 Intrusion’s cooling history. The transgressive nature of Au mineralization described by Goldel et

325 al. (2014) further supports this. Manning and Bird (1986) determined minimum temperature of

326 formation for late clinopyroxene rich veins to be between ~ 500 and 750°C. Using carbonate-

327 dolomite geothermometry, Aird and Boudreau (2013) were able to separate Stillwater carbonate

328 assemblages based on temperature and textures. A similar method would provide valuable

329 temperature information however much of the carbonate associated with Ag is too small to analyze

330 with a high degree of certainty (see Fig. 5d, e). Carbonate occurrences are also observed in late,

331 cross-cutting, titanite + albite + calcite veins. In this instance, carbonates are more typically end-

332 member calcite, consistent with temperature estimates (~ 350-500°C) for these veins by other

333 workers (Bird et al. 1986).

334 Sonnenthal (1992) described Cl bearing fluid inclusions in late anorthositic pegmatites of

335 the Upper Zone while Larsen et al. (1992) describe saline fluid inclusions containing up to ~ 22

336 wt. % NaCl in gabbroic pegmatites in and below the Middle Zone. This is consistent with Ag

337 transport by an aqueous NaAgCl2 complex, as predicted by Yin and Zajacz (2018). Further,

338 Larsen et al. (1992) suggest that fluid inclusions present within interstitial phases suggests a

339 crystal + (evolved) silicate liquid + exsolved fluid assemblage, consistent with the presence of

340 apatite, biotite, and salts. This is in agreement with our observation of fluid inclusions in

341 interstitial olivine (MZ) and plagioclase with apparent silicate and iron-rich melt inclusions

342 (MZ). Na enrichment will occur during reaction (3) where plagioclase is converted to anorthite

343 by a silica under-saturated fluid. It is noted that, in some instances, MZ sulfide assemblages

344 preserve An# ~ 62 near oxidized sulfides while An# of ~ 52 is preserved away from sulfides

345 (Fig. 4a, b, c). UZa plagioclase preserve larger variations, with one instance of near sulfide An#

346 ~ 63 while An# away from the sulfide is ~ 33 (Fig. 4d), suggesting this process becomes more

347 efficient, and total Na+ increases, up stratigraphy.

348 What remains unclear is why Ag is observed exclusively in native form in lithologies

349 below the MZ while Ag occurs as alloys (Bird et al. 1991; Andersen et al. 1998; Cabri et al.

350 2005), in addition to native Ag, and in sulfides (Holwell et al. 2014) in the MZ. High Ag sulfide

351 liquid/silicate liquid partition coefficients of ~ 1000 (Li and Audetet, 2012; Patten et al., 2013)

352 suggest that Ag should be retained in sulfide, if any is present. One possible explanation is that

353 the Lower Zone magma was simply under-saturated in sulfide. Another scenario may exist where

354 LZ S was lost to an exsolving volatile phase (e.g., Li and Boudreau 2017) leaving behind small

355 amounts of Ag with little or no S.

356

357 5.3 Modelling Silver Solubility

358 To illustrate Ag enrichment during crystallization and degassing, a simple partitioning

359 model is calculated using Rayleigh fractionation conditions (Shaw 1970):

퐷̅−1 360 퐶푙푖푞 = 퐶표 ∗ 퐹 (1)

361 where 퐶푙푖푞 and 퐶표 represent the concentration of Ag in the silicate liquid and the initial

362 concentration, respectively. 퐹 corresponds to the weight fraction of silicate liquid remaining; 퐷̅ is

363 the bulk partition coefficient of Ag in the system given by:

̅ 푣푎푝/푙푖푞 푥푡푎푙/푙푖푞 364 퐷 = 퐷 ∗ 푋푣푎푝 + ∑(퐷푖 ∗ 푋푖) (2)

푣푎푝/푙푖푞 푥푡푎푙/푙푖푞 365 퐷 is the Ag vapor/liquid partition coefficient, 퐷푖 corresponds to the Ag

366 crystal/liquid partition coefficient for each mineral phase 푖. 푋푖 and 푋푣푎푝 are the weight fraction

367 of crystal 푖 and vapor separating from a silicate liquid, respectively. Because Ag does not readily

푥푡푎푙/푙푖푞 368 partition into common silicate minerals (Adam and Green 2006; Greaney et al. 2017), 퐷푖

369 is equal to 0. Ultimately, the concentration of Ag in the vapor as crystallization proceeds is given

370 by:

퐷̅−1 푣푎푝/푙푖푞 371 퐶푣푎푝표푟 = 퐶표 ∗ 퐹 ∗ 퐷 (3)

372

푣푎푝/푙푖푞 373 Simon et al. (2008) determine that 퐷퐴푔 in rhyolitic systems is 32. For the purposes of this

푣푎푝/푙푖푞 374 calculation, a conservative 퐷퐴푔 of 10 is used. Because the initial Ag concentration of the

375 Skaergaard silicate liquid is unknown, end-member calculations are carried out for a mid-ocean

376 ridge basalt (MORB) (퐶퐴푔 = 0.0718 ppm, sample: T3-72D254-16-4, Keays and Scott 1976) and

377 a Hawaiian basalt from the Kilauea Iki lava lake (퐶퐴푔= 0.09 ppm, sample: 67-3-6.8, Greaney et

378 al. 2017) under sulfide-free conditions. In this example, Kilauea Iki is thought be analogous to a

379 sulfide-under saturated degassing, cumulate pile.

380 As illustrated in Fig. 9, Ag increases in both the volatile and liquid phase as the weight

381 fraction of silicate liquid decreases. Also plotted is a dashed line corresponding to Ag solubility

382 in an S-free basaltic liquid (~ 5.5 ppm, Zajacz et al. 2013). For MORB liquid, extreme

383 fractionation (퐹 = 0.002) is required before MORB liquid reaches Ag saturation values of ~ 5.5

푣푎푝/푙푖푞 384 ppm. Because 퐷퐴푔 is 10, coevolving MORB vapor saturates in Ag much sooner during the

385 evolution of the liquid (퐹 = 0.02). The evolution of the Hawaiian basalt is identical in behavior

386 only saturation of Ag in the liquid and vapor occurs earlier in the evolution of the liquid (퐹 =

387 0.015 and 0.15, respectively). In the absence of sulfide, Ag (and other metals with a similar

388 affinity for Cl) will preferentially concentrate in the coevolving vapor phase. This observation is

389 confirmed by numerous studies of gas condensates from volcanic fumeroles (Yudovskaya et al.

390 2006; Yudovskaya et al. 2008; Chaplygin et al 2015)

391 In the following, the partitioning of Ag into a volatile phase in both sulfide-free (Case 1)

392 and sulfide-bearing (Case 2) systems is examined. We follow the methods of Boudreau and

393 McCallum (1992) where fractional vaporization from an interstitial liquid is assumed similar to

394 fractional melting and is described by equation (4);

395 1 퐶푟 (퐷−1) 396 = (1 – 퐹푣푎푝) (4) 퐶표 397

398 Where, 퐶푟 = the residual concentration of the element of interest, 퐶표 = the initial concentration of

399 the element of interest, 퐹푣푎푝 = the weight fraction of vapor exsolving from the silicate liquid and

400 퐷 = bulk partition coefficient given by equation (5);

푋푙푖푞 퐷푠푢푙푓 401 퐷 = ( ) + 푋푠푢푙푓 ∗ ( ) (5) 퐷푙푖푞 퐷푓

402 Where, 푋푙푖푞 = the weight fraction of silicate liquid remaining in the system, 푋푠푢푙푓 = the weight

403 fraction of sulfide in the system, 퐷푓 = the fluid/silicate liquid partition coefficient while 퐷푠푢푙푓 =

404 the sulfide/silicate liquid partition coefficient for the element of interest. Finally, the

405 concentration of the element of interest in the separated fluid at any interval is given by:

퐶 1 퐶 406 푓 = ∗ 푟 (6) 퐶표 퐷 퐶표

407 where, 퐶푓 = the concentration in the exsolved fluid.

408 Case 1. The partitioning of Ag into an exsolved fluid in a sulfide-free system (Fig. 10a,

409 b). Here, it is assumed that 75% silicate liquid has crystallized to form silicate minerals while

410 25% interstitial silicate liquid remains. The liquid is considered to be undersaturated in volatile at

411 25% interstitial silicate liquid. The liquid is considered to saturate in Cl-bearing vapor at 5%

412 such that for every 1% silicate liquid crystallized, 5% of that is Cl-bearing vapor. Initial Cl

413 concentrations of 1, 5, 10, and 20 mol are assumed corresponding to modest values of 12.5, 62.5

3 푓푙푢푖푑/푠푖푙푖푐푎푡푒 414 125, and 250 ppm (ρ = 2.8 g/cm ) or 퐷퐴푔 values of 4, 20, 40, and 80, respectively

415 (Zajacz et al., 2008). Cl is considered to be perfectly incompatible in the solids prior to vapor

416 saturation so that concentrations increase exponentially as crystallization takes place. As shown

417 in Fig. 9a and 9b, residual concentrations are most affected by larger Ag fluid/silicate liquid

418 partition coefficients where residual Ag concentrations are half the original value after ~ 20%

419 silicate liquid has crystallized and the bulk weight fraction H2O has evolved to ~ 0.012 (Fig. 10a,

420 b). Similarly, Ag in the exsolved fluid is highest in the early stages of fluid evolution before

421 values diminish due to dilution (Fig. 10b).

422 Case 2. Fig.10c and 10d show a similar calculation for a sulfide-bearing system

423 composed of 74.9% crystals, 25% silicate liquid, and .1% sulfide liquid. Cl molarities remain

424 unchanged from Case 1. It is assumed that the liquid is sulfide saturated at 400 ppm S. Each

425 fraction of water evolved contains 5% S which is subtracted from the starting S content before

426 each subsequent calculation. The Ag sulfide/silicate liquid partition coefficient is taken to be

427 1000 (Jenner 2017) while Ag fluid/silicate liquid partition coefficient values are as in the S-free

428 calculation. Calculations suggest that minimal change is observed in residual concentration prior

429 to sulfide being completely exhausted (vertical dashed grey line, Fig. 10c). This is in large due to

430 the high Ag sulfide/silicate liquid partition coefficient values suggesting the partitioning of silver

431 into a fluid in an sulfide-bearing system is controlled by the amount of sulfide present. Indeed,

푓푙푢푖푑/푠푖푙푖푐푎푡푒 432 for each 퐷퐴푔 value, concentration in the exsolved fluid is maximized after sulfide is

433 entirely exhausted (Fig. 10d).

434 In either scenario, calculations suggest that a volatile phase efficiently removes Ag from

435 the residual silicate liquid. In the sulfide-free system, the earliest exsolved fluid (푋푤푎푡푒푟 < .002)

436 experiences the highest enrichment in Ag that decreases as Ag is removed from the system. The

437 residual silicate liquid experiences a gradual reduction in initial Ag content as water evolves. In

438 the sulfide-bearing system, Ag in the residual silicate liquid is most affected after sulfide is lost

439 to the volatile phase (푋푤푎푡푒푟 ~ .0094). This suggests that if sulfide is present, Ag will be

440 preferentially partitioned into sulfide unless large fractions of volatiles evolve from the

푓푙푢푖푑/푠푖푙푖푐푎푡푒 441 interstitial silicate liquid or 퐷퐴푔 values are sufficiently high.

442

443 5.5 A Model

444 A scenario is envisioned, that is not unlike that proposed by Meurer et al. (1999), where

445 Ag (with other Cl soluble elements having low 퐷푚푟푛푙/푚푒푙푡 partition coefficients) becomes

446 increasingly concentrated in the residual liquid during crystallization. At high silicate

447 mineral/liquid mass fraction (퐹) values, Ag is concentrated to quantities where saturation in Ag

448 becomes favorable. At this point, or shortly thereafter, enough silicate liquid volume has

449 crystallized to allow for the separation of H2O, CO2, Cl and other volatile components with low

450 퐷푚푛푟푙/푚푒푙푡partition coefficients. Owing to its relatively high 퐷푓푙푢푖푑/푠푖푙푖푐푎푡푒 partition coefficient,

451 Ag partitions into the exsolved fluid as Ag-chloride or perhaps Na-Ag-chloride. The exsolved

452 fluid transporting Ag will travel vertically along grain boundary networks, and fluid pathways,

453 interacting with other interstitial phases. Heating of the fluid as it moves into hotter rocks can

454 lead to S and SiO2 loss from the host assemblage while leaving behind a small portion of the Ag-

455 bearing volatile phase. Eventual cooling of the fluid can cause Ag precipitation or, when a fluid

456 under-saturated interstial liquid is encountered, re-dissolution of the fluid initiates the precipitation

457 of Ag and perhaps an increased volatile content for that particular section of the intrusion. Minor

458 amounts of metal + volatile are trapped in the interstitial liquid and are presumed to have

459 generated textures seen in Fig. 5a, and Fig. 8. The cumulative effect of this process most affected

460 the MZ where the silicate liquid became saturated in fluid + vapor early enough to preserve fluid

461 inclusions in olivine, clinopyroxene, and plagioclase.

462

463 6. Conclusions

464 Evidence for sulfur loss is present throughout the Skaergaard stratigraphy, with the

465 relative proportion of low-Ti magnetite replacing sulfide being most extensive in the Middle

466 Zone. Native Ag is regularly associated with trace quantities of Cl in and below the MZ,

467 consistent with a coexisting metal + brine assemblage. Ag can be readily mobilized by an

468 exsolved volatile Cl-bearing phase in both sulfide-free and sulfide-bearing liquids. It is suggested

469 that a late fluid rich in halides transported silver and perhaps other precious metals vertically

470 through the Skaergaard cumulate pile. It is proposed that the Kilaeua Iki lava lake may be

471 analogous to a S-free degassing cumulate pile. Finally, this contribution presents new evidence

472 supportive of late hydrothermal processes in LMIs and suggests that similar work may be helpful

473 in understanding late stage evolution of other intrusions.

474

475 Acknowledgements

476 Discussions with J. Cann, E. Klein, and T.L. Nielsen were instructive and improved the

477 quality of this manuscript. Alex Hammerstrom is thanked for his work compiling sample

478 location information.

479

480 References

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630 631 Fig. 1 A) Simplified geologic map of the Skaergaard Intrusion with the approximate location of 632 samples used in this study denoted. Geologic map modified after McBirney (1989).

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640 641 Fig. 2 a) Stratigraphy of the Skaergaard Intrusion with approximate location of samples used in 642 this study denoted by white stars. Grey bars denote extent of occurrence for low-Ti magnetite, 643 biotite, calcite, and apatite. b) Sulfide assemblage proportions by oxides and sulfides (left 644 column) and other accessory minerals (right column). c) Distribution and occurrence of silver 645 (left) and percentage of silver associated with trace chlorine (right).

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651 652 Fig. 3 Area and elongation measurements for individual sulfide assemblages. Elongation is unit 푃2 653 less and is equal to . Spherical objects have elongation values of 1. 4∗휋∗퐴 654

655

656 657 Fig. 4 a) Sample SK90-5 (LZa) Sulfide assemblage readily altered to magnetite. Note native 658 copper occurring where much of the sulfide has been lost and that plagioclase immediately 659 surrounding the assemblage is more calcic (brighter orange, white dashed line) b) Sample SK90- 660 5 (LZa) Polysulfide assemblage cross-cut by magnetite and biotite. Note that plagioclase 661 immediately surrounding the assemblage is more calcic (brighter orange, white dashed line). c) 662 Sample SK90-5 (LZa) Sulfide assemblage truncated by biotite, orthopyroxene, and low-Ti 663 magnetite. Note that the area immediately surrounding the sulfide is more calcic plagioclase 664 (brighter orange, white dashed line). d) Sample SK90-50 (MZ) Interstitial sulfide from the 665 Middle Zone that is entirely oxidized. Note the minor Ag occurrence. e) Sample SK90-32A 666 (UZa) Rounded sulfide with areas of high and low S concentrations with no sign of host silicate 667 alteration. f) SK90-32A (UZa) Interstitial bornite replaced by low-Ti magnetite. Note near 668 sulfide high-Ca plagioclase (brighter orange, white dashed line). All figures are EDS 669 composition maps.

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671

672 673 Fig. 5 a) Sample SK90-5 (LZa) Silver interstitial to plagioclase, apatite, and clinopyroxene. b) 674 Sample SK90-5 (LZa) Native silver in plagioclase near grain boundary. c) Sample SK90-5 (LZa) 675 Native silver with associated silver halides at olivine-clinopyroxene grain boundary. d) Sample 676 SK90-50 (MZ) Native silver in plagioclase with calcite and Cl-rich inclusion. e) Sample SK90- 677 11-1b (MZ) Silver – gold intergrowth in plagioclase. Note associated Cu and calcite. f) Sample 678 SK90-50 (MZ) Ag-Zn alloy with associated native copper in a chlorine rich inclusion in 679 clinopyroxene. All figures are EDS composition maps.

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689 690 Fig. 6 a) Back-scattered electron image of an apparent fluid inclusion in clinopyroxene 691 containing native silver with trace quantities of quartz. b) Semi-quantitative X-Ray Spectra for 692 the field outlined in (a). Note the abundance of Ca, Na, Si, Cl, and K. c) EDS map for Ag. Same 693 field of view as (a). d) Semi-quantitative X-Ray Spectra for the field outlined in (c). Note the 694 composition is almost entirely Ag with trace Cl + K.

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701 702 Fig. 7 Plain polarized transmitted light photomicrography of fluid inclusions in olivine. Sample 703 SK90-50 is from the MZ.

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713 714 Fig. 8 Sample SK90-5 (LZa). EDS composition map of native gold in plagioclase with 715 associated halite and sylvite. Blue regions that are not labeled represent areas of high K 716 concentration while purple regions that are not labeled represent areas of high Na concentration. 717 The right most occurrence is interpreted as a solvus texture indicative of cooling.

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734 735 Fig. 9 Numerical calculation of Ag enrichment in endmember basaltic liquids and evolving vapor as a 736 function of fraction of silicate liquid remaining. Ag concentrations for the liquid (red) and co-existing 737 vapor (blue) are shown for Hawaiian basalt and MORB. The horizontal red dashed line indicates the 738 saturation level of Ag in a S-free basaltic liquid (~5.5 ppm, Zajacz et al., 2013). See text for further 739 discussion. 740

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748 749 Fig. 10 Numerical calculations simulating the dissolution of silver in sulfide-free (top) and 750 sulfide-bearing (bottom) systems. See text for further discussion.

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