Stable isotope (C, O, S) compositions of volatile-rich minerals in kimberlites: A review

Chemical Geology 374–375 (2014) 61–83

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Stable isotope (C, O, S) compositions of volatile-rich minerals in kimberlites: A review Andrea Giuliani a,⁎, David Phillips a, Vadim S. Kamenetsky b, Marco L. Fiorentini c, James Farquhar d, Mark A. Kendrick a,e a

School of Earth Sciences, The University of Melbourne, Parkville, 3010 Victoria, Australia School of Physical Sciences, University of Tasmania, Hobart, 7001 Tasmania, Australia Centre for Exploration Targeting, ARC Centre of Excellence for Core to Crust Fluid Systems, The University of Western Australia, 35 Stirling Highway, Crawley, 6009 WA, Australia d Department of Geology and ESSIC, University of Maryland, College Park, MD 20742, USA e Research School of Earth Sciences, The Australian National University, Canberra, 0200 ACT, Australia b c

a r t i c l e

i n f o

Article history: Received 15 August 2013 Received in revised form 28 February 2014 Accepted 3 March 2014 Available online 13 March 2014 Editor: David R. Hilton Keywords: Kimberlite Carbonate Serpentine Carbon-oxygen-sulphur isotopes

a b s t r a c t The composition of primary kimberlite melts and, in particular, the absolute and relative abundances of volatile components (mainly CO2 and H2O) are controversial issues, because kimberlite melts entrain and interact with abundant mantle and crustal xenoliths during ascent, react with wall rocks during emplacement, and lose some of their volatile inventory during pre- and syn-emplacement degassing. Compositional constraints are further complicated by the common alteration of kimberlitic rocks by post-emplacement fluids of various origin (e.g., deuteric, meteoric, hydrothermal). Consequently, the compositions of kimberlitic rocks may not be entirely representative of their parental melts. In kimberlitic rocks, CO2 is concentrated in carbonate minerals, whereas H2O is mainly stored in the secondary minerals serpentine and, to a lesser extent, chlorite and brucite, with minor contribution by primary magmatic phlogopite. This review focuses on utility of carbon, oxygen and sulphur stable isotopes to constrain the source of volatiles (i.e. magmatic vs non-magmatic) for carbonate, serpentine, sulphide and sulphate formation and the origin of fluids altering kimberlitic rocks. A global compilation of kimberlite carbonate data (δ13C = −11.9 to +0.2‰, median δ13C = −5.0‰, relative to VPDB; δ18O = 1.2–26.6‰, median δ18O = 13.2‰, relative to VSMOW) reveals that the majority of results (86%) plot within a range of δ13C ~ −2 to −8‰, which is considered representative of mantle carbon, but only 15% of analyses are in the field of oxygen isotopic values for mantle carbonates (δ18O ~ 6–9‰). Variations in kimberlite carbon isotopic compositions occur on regional scales, implying widespread mantle heterogeneity, possibly related to input of carbon from recycled crustal material and/or partial overprinting by secondary processes at the local scale. Carbonates in southern African Group I (or archetype) and Group II kimberlites (or orangeites) show different δ13C distributions (median values of −5.3‰ and −6.5‰, respectively). This is consistent with distinct mantle sources, as demonstrated previously by radiogenic isotope studies. Kimberlite breccia carbonates commonly have higher δ18O values than carbonates in massive and hypabyssal kimberlites, which suggests more extensive interaction of kimberlite rocks with hydrous fluids in the brecciated parts of kimberlite pipes. Modelling of the stable isotope compositions of carbonates from the Kimberley, Lac de Gras and Udachnaya-East kimberlites reveals that several processes are capable of modifying these compositions, including interaction with H2O-rich deuteric (i.e. late-stage magmatic) fluids, meteoric waters and/or hydrothermal fluids, and incorporation of sedimentary material. However, these processes can produce similar variations of the carbonate C–O isotopic compositions, which means that carbonate isotopes alone cannot provide tight constraints on the alteration of kimberlite rocks. Only few carbonates in hypabyssal kimberlites show isotopic compositions consistent with abundant CO2 degassing (i.e. increasing δ18O with decreasing δ13C values), thus implying that kimberlite magmas that are not emplaced explosively retain most of their CO2 concentrations prior to carbonate crystallisation. In kimberlitic rocks early-formed serpentine exhibits higher δ18O values (~+4–+6‰) than later serpentine rims and segregations (δ18O values as low as ~−2‰). These variations are consistent with serpentine crystallisation from hydrous fluids derived from mixing between deuteric fluids and meteoric/hydrothermal fluids, with progressive enrichment in the latter component. Serpentine is considered to have formed under hydrothermal conditions when externally derived hydrous fluids infiltrated the cooling kimberlite volcanic system. Only limited sulphur isotopic data are available for kimberlitic bulk rocks and sulphide and sulphate phases. Of these, relatively few sulphur isotopic ratios approach the δ34S values considered representative of the mantle

⁎ Corresponding author. Tel.: +61 3 90359873; fax: +61 3 83447761. E-mail address: [email protected] (A. Giuliani).

http://dx.doi.org/10.1016/j.chemgeo.2014.03.003 0009-2541/© 2014 Elsevier B.V. All rights reserved.

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A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

(0 ± 2‰, relative to VCDT). Elevated δ34S values (~ 14‰) characteristic of sulphates in the Udachnaya-East kimberlite are consistent with equilibration with sulphides (δ34S ~ 1–2‰) at temperatures of ~ 500–550 °C, after kimberlite melt outgassing under oxidising conditions. Conversely, the large δ34S range shown by some southern African and Yakutian kimberlites (−3–+12‰ and +15–+53‰, respectively) may be largely due to alteration and crustal contamination. In conclusion, the stable isotopic compositions of carbonates, serpentine and S-rich minerals in kimberlites, can be used in conjunction with detailed petrographic and geochemical analyses, to constrain processes affecting kimberlite magmas prior to, during, and subsequent to crystallisation. The available stable isotopic data indicate that externally derived (i.e. non-magmatic) hydrothermal fluids have affected the compositions of most kimberlites, including the hypabyssal varieties often used to reconstruct the compositions of primary kimberlite melts. This discrepancy remains a major obstacle in the quest for the primary composition of kimberlite melts. © 2014 Elsevier B.V. All rights reserved.

Contents 1.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1. Geology, mineralogy and geochemistry of kimberlites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2. Stable isotope geochemistry notation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.3. Oxygen, carbon and sulphur isotopes of mantle rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2. Carbon and oxygen isotopic geochemistry of carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Processes affecting the C–O isotopic compositions of carbonates in magmatic rocks . . . . . . . . . . . . . . . . . . . . . . . 2.2. Kimberlite carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.1. Regional variations of carbon isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.2. Regional variations of oxygen isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.3. Local variations of C–O isotopes: Kimberley (South Africa), Lac de Gras (Canada) and Udachnaya-East (Russia) kimberlites 3. Oxygen isotopes of serpentine minerals in kimberlites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4. Sulphur isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1. Processes affecting the sulphur isotopes of magmatic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2. Sulphur isotopes in kimberlitic rocks and minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5. Conclusions and future directions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Appendix A. Supplementary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction Kimberlites are enigmatic, rare, small volume igneous rocks that are important because they are the primary host rock to diamonds and because kimberlite parental melts originate from deep within the Earth (N150 km, i.e. within the diamond stability field — e.g., Haggerty, 1994; Ringwood et al., 1992; Tainton and McKenzie, 1994; Torsvik et al., 2010). In addition, kimberlite magmas have entrained abundant fragments of mantle and deep crust wall rocks en route to the surface, thus providing the major source of information about the petrology and geochemistry of the deep lithosphere in continental areas (e.g., Dawson, 1980; Menzies and Hawkesworth, 1987; Nixon, 1987; Pearson et al., 2003; Schmitz and Bowring, 2003a,b; Zartman et al., 2013; Giuliani et al., 2014). Due to their hybrid and volatile-rich nature and widespread alteration by deuteric (i.e. late-stage magmatic), meteoric and hydrothermal fluids, the primary composition of kimberlites has proven difficult to constrain. A number of studies have employed stable isotopes to provide insights into the evolution of kimberlite magmas upon emplacement, with particular emphasis on the source of fluids involved during alteration of kimberlite rocks (e.g., Vinogradov and Ilupin, 1972; Sheppard and Dawson, 1975; Kobelski et al., 1979; Ukhanov et al., 1986; Kirkley et al., 1989; Ustinov et al., 1994; Wilson et al., 2007; Mitchell, 2013). Carbon and oxygen stable isotopes have been widely used to constrain the source of CO2 and H2O, i.e. the most abundant volatile species, in kimberlites. This review critically evaluates the existing data for C and O isotopes in kimberlite carbonates, the O isotopes of serpentine and the S isotopes

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62 62 64 65 66 67 68 69 70 70 72 76 76 76 78 79 79 79

of sulphides and bulk kimberlitic rocks. A large number of analyses (N500) exist for carbonates, but limited data are available for serpentine and sulphide/sulphate minerals in kimberlites. After introducing the main geological, mineralogical and geochemical features of kimberlites, we will define the range of O–C–S stable isotopic compositions shown by mantle rocks. Detailed examination of carbonate data for kimberlite pipes and clusters reveal that kimberlite carbonates are ubiquitously affected by interaction with mixtures of deuteric and meteoric fluids and that the C and O isotopic compositions of only some carbonates reflect CO2 degassing. Conversely, the O isotopes of serpentine provide evidence that most of the water in kimberlites is of external derivation. Finally, we will summarise the most important processes that affect the S isotopes of magmatic rocks and discuss existing data for kimberlite rocks and the implications for kimberlite petrology. 1.1. Geology, mineralogy and geochemistry of kimberlites Kimberlites are silica-poor, volatile-rich igneous rocks of variable but broadly ultrabasic composition that mainly occur as volcanic pipes and hypabyssal intrusions in cratonic areas. Two distinct groups of kimberlites have been recognised based on mineralogy, major and minor elements and radiogenic isotope compositions. Kimberlites sensu stricto or Group I kimberlites (Smith, 1983) are hybrid olivine-rich rocks consisting of fragments of rocks and minerals of mantle and crustal origin intermixed in a matrix of carbonates, olivine, phlogopite, spinel, perovskite, apatite, serpentine and other minor phases, which crystallised from volatile-rich melts and fluids (Dawson, 1980; Mitchell, 1986,

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

2008; Skinner, 1989). Group I kimberlites show unradiogenic Sr and radiogenic Nd isotopic compositions similar to ocean island basalts (OIBs — Gaffney et al., 2007; le Roex, 1986; Nowell et al., 2004; Paton et al., 2009; Smith, 1983; Tappe et al., 2011, 2012; Woodhead et al., 2009). Orangeites or Group II kimberlites (Smith, 1983; Mitchell, 1995) differ from the archetypal Group I kimberlites because of higher phlogopite and therefore K2O contents, occurrence of tetraferriphlogopite and K–Ba titanates and, in more evolved orangeites, sanidine and Krichterite (Mitchell, 1986, 1995). Orangeites have more radiogenic Sr and less radiogenic Nd isotopes than archetypal kimberlites (e.g., Smith, 1983; Nowell et al., 2004; Woodhead et al., 2009). Because most of the results presented in this review are for archetypal Group I kimberlites, we will use “kimberlite” for Group I kimberlite rocks and “orangeite” for “Group II” kimberlites. The origin and composition of primary kimberlite melts are hotly debated issues. Kimberlite magmas have occasionally entrained diamonds hosting mineral inclusions thought to be typical of the mantle transition zone and lower mantle (e.g., magnesiowustite, ferropericlase, Ca-Si- and Mg–Si-perovskite, majoritic garnet, tetragonal almandinepyrope phase or TAPP, native Fe — Harte, 2010; Harte and Harris, 1994; Kaminsky, 2012; Kaminsky et al., 2001, 2009; Scott Smith et al., 1984; Stachel et al., 2000a,b; Walter et al., 2008, 2011; Wirth et al., 2009) and majorite-bearing mantle xenoliths of ultradeep (N300 km) derivation (e.g., Sautter et al., 1991). This evidence, coupled with radiogenic isotopes of kimberlites comparable to asthenospheric magmas (i.e. OIBs — le Roex, 1986; Smith, 1983), has been often employed to argue for kimberlite melt formation below the lithospheric mantle (e.g., Ringwood et al., 1992; Haggerty, 1994; Kesson et al., 1994; Tappe et al., 2011, 2012; Shirey et al., 2013), with the possible involvement of recycled oceanic lithosphere in the source (e.g., Nowell et al., 2004; Paton et al., 2009). The alternative view is that kimberlite melts are generated by partial melting of peridotites metasomatised by asthenospheric melts either at the base or within the lithospheric mantle (e.g., Tainton and McKenzie, 1994; le Roex et al., 2003; Keshav et al., 2005; Becker and le Roex, 2006; Gaffney et al., 2007; Brey et al., 2008). Metasomatism would impart an asthenospheric “flavour” to the source lithosphere (e.g., Gaffney et al., 2007) and transport of deep material in the lithosphere by metasomatic fluids shortly before kimberlite formation might also explain the occurrence of ultra-deep diamonds and xenoliths. Conversely, orangeites are generally thought to derive from partial melting of intensely metasomatised lithospheric mantle rocks (e.g., Smith, 1983; Mitchell, 1995; Ulmer and Sweeney, 2002; Nowell et al., 2004; Becker and le Roex, 2006). Once formed, kimberlite magmas ascend very rapidly (several m/s — Canil and Fedortchouk, 1999; Kelley and Wartho, 2000; Sparks et al., 2006) and, if they reach the surface, can erupt explosively or intrude into the upper crust (Sparks, 2013). The lack of evidence for contact metamorphism affecting kimberlite wall rocks (e.g., Barrett and Berg, 1975) suggests low temperature emplacement for kimberlite magmas, probably not exceeding 700 °C (Mitchell, 1975; Kamenetsky et al., 2012). The standard model of kimberlite pipe (Field and Scott Smith, 1999; Hawthorne, 1975; Mitchell, 1986; Skinner and Marsh, 2004; Sparks et al., 2006 — Fig. 1) includes i) a crater filled with pyroclastic and re-sedimented volcanoclastic material; ii) a diatreme, which is a carrot-shaped body with vertical axis that can extend to depths of 3 km or greater and is filled with volcanoclastic (or tuffisitic) kimberlite rocks (Fig. 2a) and minor wall rock breccias; and iii) a root zone, where thin feeder dykes and sills converge into an irregular (to elliptical in planar view) body of massive magmatic kimberlite (Figs. 2b and 3). Volcanological studies of kimberlites have shown that volcanism was probably initiated by magmatic and/or phreato-magmatic (i.e. with the involvement of water of external derivation) explosions, which progressively excavated kimberlite pipes (e.g., Sparks et al., 2006; Wilson et al., 2007; Kurszlaukis and Lorenz, 2008) and formed brecciated and pyroclastic rocks. The explosive activity was followed by, and/or alternated with, episodes of more quiescent activity forming shallow dykes

63

tuff ring

crater zone

diatreme zone

root zone

500 m

dykes

Fig. 1. Idealised model of a kimberlite pipe showing the crater, diatreme and root zones. Modified from Hawthorne (1975) and Mitchell (1986).

and sills and, eventually, lava eruptions (Dawson and Hawthorne, 1970; Sparks et al., 2006; Moss et al., 2009; Brown et al., 2012; Gernon et al., 2012; Sparks, 2013). Because of the porous and brecciated texture of crater and diatreme zones, the kimberlite pipe operated as a collector for groundwater in the surrounding area. This, coupled with the thermal anomaly produced by magma emplacement, favoured the development of a hydrothermal circulation in the pipe, which might be sustained for several years after kimberlite emplacement, thus promoting extensive interaction between hydrous fluids and kimberlite rocks (e.g., Mitchell, 1986; Ustinov et al., 1994; Stripp et al., 2006; Kurszlaukis and Lorenz, 2008; Buse et al., 2010; Sparks, 2013). Because kimberlites are devoid of quenched magmatic glasses, aphanitic (i.e. fine-grained, macrocryst-free) hypabyssal kimberlites have been widely considered to be the best proxies for primary kimberlite melts (e.g., Price et al., 2000; le Roex et al., 2003; Becker and le Roex, 2006; Kopylova et al., 2007; Kjarsgaard et al., 2009). However, a number of factors need to be considered when reconstructing kimberlite melt compositions using bulk samples: 1) kimberlitic rocks are ubiquitously altered either by deuteric (i.e. late-stage magmatic) or surface-derived hydrous fluids (or both). As kimberlites are enriched in olivine and may also contain water-soluble minerals, such as alkali-carbonates and chlorides (see Kamenetsky et al., 2004, 2007a, 2008, 2009a,b, 2012), interaction with hydrous fluids modifies the mineralogy and chemistry of kimberlites. 2) Volcaniclastic and even hypabyssal kimberlites never retain the volatile inventory of their parental melts because of degassing during ascent and emplacement. 3) Kimberlite melts have been shown to react with wall rocks and entrained xenoliths (e.g., Smith et al., 2004; Stripp et al., 2006; Mitchell, 2008; Brett et al., 2009; Kamenetsky et al., 2009b; Buse et al., 2010; Russell et al., 2012) before and during emplacement. 4) Finally, the hybrid nature of kimberlite rocks implies that xenogenic materials (e.g., xenocrystic cores of olivine grains — Arndt et al., 2010; Brett et al., 2009; Kamenetsky et al., 2008; Pilbeam et al., 2013) contribute to kimberlite bulk compositions. For these reasons, the composition of primary kimberlite melts remains enigmatic, with proposed compositions ranging from H2O ± CO2-rich ultramafic silicate melts (e.g., Price et al., 2000; le Roex et al., 2003; Kopylova et al., 2007; Kjarsgaard et al., 2009) to essentially anhydrous

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A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

a

b

CX

OM

2 cm

1 cm

Fig. 2. Hand-specimens of: (a) volcanoclastic kimberlite from the Misery-East pipe (Lac de Gras, Canada) and (b) hypabyssal kimberlite from the Koala pipe (Lac de Gras, Canada). CX: crustal xenolith; OM: olivine macrocryst.

Cl-rich alkali-carbonate magmas (Kamenetsky et al., 2004, 2008, 2009a, 2012). The concentration of volatile species (H2O, CO2, F, Cl, S) in pristine kimberlite melts is a particularly controversial issue because postemplacement alteration, syn-emplacement degassing and preemplacement interaction with mantle and crustal rocks and fluids have major impacts on the final volatile contents of kimberlites. Therefore, bulk analyses of even the least altered kimberlitic rocks are unlikely to provide reliable estimates of the original magmatic volatile contents. This has been demonstrated by the experiments of Sparks et al. (2009) and Brooker et al. (2011), who investigated the solubility of H2O and CO2 in putative melt compositions representing hypabyssal kimberlite rocks (corrected for contributions from crustal and mantle material). Brooker et al. (2011) found that the kimberlite compositions studied were unable to dissolve the concentrations of H2O and CO2 measured in kimberlite rocks. They concluded that the compositions of hypabyssal kimberlites could not represent pristine kimberlite melts. In kimberlitic rocks, CO2 is concentrated in carbonate minerals that are generally considered to be principally of primary magmatic origin or formed from deuteric (late-stage magmatic) fluids (Dawson and Hawthorne, 1970; Dawson, 1980; Mitchell, 1986; Armstrong et al., 2004; Wilson et al., 2007). However, trace element (Sr, Ba, REE) and Sr isotope studies of carbonates representing various textural types (e.g., groundmass grains, segregations, veins) have shown that some carbonates crystallised from post-magmatic fluids of external derivation (Brookins, 1967; Exley and Jones, 1983; Podvysotskiy, 1985; Armstrong et al., 2004). Moreover, in cases where kimberlite magmas are emplaced into carbonate-rich country rocks, such as limestones, significant enrichment in CO2 (carbonatisation) could have resulted from wall-rock assimilation (e.g., Kirkley et al., 1989).

5 mm

a

Cc

Serpentine is the major host of water in hypabyssal kimberlite rocks with minor contributions by phlogopite, chlorite and brucite. Serpentine occurs as a replacement mineral of olivine, carbonates and other phases, as groundmass minerals, and as segregations with variable amounts of carbonate (e.g., Dawson, 1980; Podvysotskiy, 1985; Mitchell, 1986, 2013; Mitchell et al., 2009; White et al., 2012). The significance of serpentine in kimberlitic rocks is a controversial issue, because the origin of the water in serpentine is unclear. One possibility is that the water has a deuteric (i.e. magmatic) origin; in this case, the high serpentine abundance (up to 50 vol.%; Skinner and Clement, 1979) in kimberlites would imply that kimberlite melts contain high water contents (Mitchell and Putnis, 1988; Kopylova et al., 2007; Mitchell, 2008, 2013; Kjarsgaard et al., 2009). Alternatively, the water derives from meteoric/ hydrothermal fluids of external derivation (Sheppard and Dawson, 1975; Ukhanov and Devirts, 1983; Ukhanov et al., 1986; Sparks et al., 2006, 2009; Stripp et al., 2006). From the above discussion, it is evident that resolving the origin of carbonate and serpentine minerals is of primary importance for understanding the sources of volatiles in kimberlitic rocks and, ultimately, the composition of kimberlitic melts. 1.2. Stable isotope geochemistry notation Stable isotopic variations are expressed conventionally using the delta notation in per mil (‰) units:
δa ð‰Þ ¼

 Ra −1  1000 Rst

where Ra and Rst values are 13C/12C, 18O/16O and 34S/32S ratios in the standard and the sample, respectively. The data presented in this review

2 mm

b

Ol

Ol

Fig. 3. Photomicrographs showing the texture of hypabyssal kimberlite samples from: (a) the Diavik mine (Lac de Gras, Canada), and (b) the Gollapalle pipe (Wajrakarur, India). Cc: calcite segregations; Ol: olivine.

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are limited to δ13C (relative to VPDB standard), δ18O (relative to VSMOW) and δ34S (relative to VCDT). When the stable isotopes of an element fractionate between two phases a and b, the following approximation is valid for differences in δ values b10‰

1986; Mattey et al., 1984; Palot et al., 2012; Pineau and Javoy, 1983; Taylor, 1986). Recycling of subducted oceanic crust hosting abundant organic carbon (δ13C ~ − 25‰) and carbonates (δ13C ~ 0‰) into the source region of some oceanic basalts, carbonatites, diamond-forming fluids and mantle metasomatic fluids/melts, has been claimed to produce shifts from the above canonical − 5‰ value (e.g., Mattey et al., 1984; Demeny et al., 1998, 2004a,b, 2010; Nishio et al., 1998; van Achterbergh et al., 2002; Ducea et al., 2005; Tappert et al., 2005; Ickert et al., 2013; Schulze et al., 2013). However, low δ13C values (≪−5‰) approaching and/or overlapping the δ13C values of organic carbon have been interpreted to reflect Rayleigh fractionation of mantle fluids (e.g., Cartigny et al., 1998) or carbon derivation from a primordial mantle reservoir with δ13C ~ −25‰ (Deines, 2002). Such low 13C/12C components are thought to reside in carbides, hydrocarbons, atomic C dissolved in silicate minerals and, more rarely, in diamond and graphite (Deines, 2002). CO2 degassing of C–O–H fluids is another process capable of reducing the δ13C values of carbonates crystallised from mantle fluids (e.g., Demeny and Harangi, 1996). The sulphur isotopic composition of the Earth's mantle has been assumed to be close to δ34S = 0‰ (e.g., Thode et al., 1961). Indeed, the compositions of mid-ocean ridge basalts (MORBs), which derive from the asthenospheric mantle, generally cluster around this canonical value (δ34S ~ −2–+2‰; Fig. 6 — Labidi et al., 2012, 2013; Sakai et al., 1984). However, the δ34S values of some ocean island basalts (OIBs) largely deviate from zero (Fig. 6), and have been attributed to processes such as melt degassing, sulphur fractionation under mantle conditions and/or recycling of crustal sulphur in the asthenospheric source (Chaussidon et al., 1989; Torssander, 1989; Cabral et al., 2013). The latter interpretation is supported by the recent discovery of sulphides included in Samoan OIB olivine phenocrysts with Δ33S ≠ 0‰ (Cabral et al., 2013). As Δ33S only deviates significantly from zero in crustal rocks older than ~2.45 Gyr (Farquhar and Wing, 2003; Farquhar et al., 2007; Philippot et al., 2012), the Samoan Δ33S values support recycling of subducted Archean crustal rocks into the mantle source (Cabral et al., 2013). Likewise, island arc basalts show high 34S/32S compared to MORBs, which is probably due to interaction of the mantle source with slab-derived fluids (e.g., de Hoog et al., 2001; de Moor et al., 2010). Sulphur isotope analyses of mantle xenoliths sampled by kimberlites and alkali-basalts (Muramatsu, 1983; Kyser, 1990; Ionov et al., 1992; Wilson et al., 1996; Giuliani et al., 2013), as well as sulphide inclusions in mantle

1000  ln αa–b ≈δa −δb ¼ Δa–b where α is the fractionation factor defined as αa–b ¼

Ra

65

. Rb

where Ra and Rb are the stable isotope ratios of phases a and b. Accordingly, fractionation factors will be expressed as Δ values throughout this review. 1.3. Oxygen, carbon and sulphur isotopes of mantle rocks Oxygen isotope studies of oceanic basaltic glasses and mantle peridotite xenoliths sampled by kimberlite and basalt magmas have shown a restricted δ18O distribution of +5 to +6‰ (Fig. 4 — Bindeman, 2008; Chazot et al., 1997; Eiler, 2001; Eiler et al., 2011; Mattey et al., 1994a; Rehfeldt et al., 2008; Schulze et al., 2001; Valley et al., 1998), which is considered representative of the mantle reservoir. Some basalts and olivine phenocrysts from intra-oceanic and island arc settings display δ18O values outside this range (Fig. 4), which has been attributed to contamination during magma residence in the crust (Bindeman, 2008 and references therein) or recycling of altered oceanic crust into the mantle source (Harmon and Hoefs, 1995; Eiler, 2001; Demeny et al., 2008). Likewise, eclogite xenoliths entrained by kimberlite magmas, and eclogite mineral inclusions in diamonds, often show fractionated oxygen isotopes (Fig. 4), which supports a subduction origin for some mantle eclogites (e.g., MacGregor and Manton, 1986; Neal et al., 1990; Mattey et al., 1994b; Schulze et al., 2003, 2013; Jacob, 2004; Ickert et al., 2013; Smart et al., 2014) — however, see Griffin and O'Reilly (2007) and Huang et al. (2012) for alternative views on eclogite formation in the mantle. Carbon isotope investigations of oceanic basalts, volcanic gases, diamonds and mantle xenoliths have identified a major carbon component in the mantle, characterised by δ13C ~ −5 ± 2‰ (Fig. 5 — e.g., Cartigny, 2005; Deines, 1989, 2002; Des Marais and Moore, 1984; Javoy et al.,

MORBs OIBs

to 28.5

carbonatites SCLM peridotite xenoliths kimberlite megacrysts

to 16.9

eclogite inclusions in diamonds SCLM eclogite xenoliths

to 26.6

kimberlite carbonates 0

4

8

12

δ 18 O (‰)

16

Fig. 4. Oxygen isotope (‰ δ18O relative to VSMOW) variations in mantle-derived magmas and mantle rocks and minerals from: mid-ocean ridge basalts (MORBs — Eiler, 2001 and references therein), ocean island basalts (OIB — Eiler, 2001; Bindeman, 2008, and references therein), carbonatites (Deines, 1989), peridotite and eclogite xenoliths entrained in the subcontinental lithospheric mantle (SCLM) by kimberlitic and alkali-basalt magmas (Chazot et al., 1997; Ickert et al., 2013 and references therein; Jacob, 2004 and references therein; Mattey et al., 1994a,b; Rehfeldt et al., 2008), megacrystic nodules in kimberlites (Valley et al., 1998; Schulze et al., 2001; Page et al., 2007), eclogitic inclusions in diamonds (Lowry et al., 1999; Schulze et al., 2003, 2013; Ickert et al., 2013) and kimberlite carbonates (including orangeite carbonates — Arima and Kerrien, 1988; Deines and Gold, 1973; Fedortchouk and Canil, 2004; Galimov and Ukhanov, 1989; Ito, 1986; Kamenetsky et al., 2012; Khar'kiv et al., 1991; Kirkley et al., 1989; Kobelski et al., 1979; Kuleshov and Ilupin, 1982; Price et al., 2000; Sheppard and Dawson, 1975; Tappe et al., 2011; Ustinov et al., 1994; Wilson et al., 2007). ‘SCLM peridotite xenoliths’ include analyses of olivine, orthopyroxene, clinopyroxene and garnet; ‘SCLM eclogite xenoliths’ comprise analyses of garnet and clinopyroxene; ‘kimberlite megacrysts’ include analyses of garnet and zircon.

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80

n.

70

30

δ 13 C (‰)

δ 13 C (‰)

25

Mantle xenoliths (n. 408)

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Oceanic basalts (n. 58) 20

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carbonates

40

reduced C + bulk rocks

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Peridotitic diamonds (n. 1874)

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Eclogitic diamonds (n. 1239)

150

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Kimberlite carbonates (n. 538)

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Carbonatites (n. 580)

60

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60 40

20

20

0

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-30

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Fig. 5. Compilation of carbon isotopic compositions (‰ δ13C relative to VPDB) for mantle-derived magmas, and mantle rocks and minerals. ‘Mantle xenoliths’ refer to rocks entrained by kimberlitic and basaltic magmas from continental and oceanic settings, and include analyses of bulk rocks, hydrocarbons, carbonaceous material, graphite and diamonds in xenoliths (‘reduced C + bulk rocks’ — Deines, 2002 and references therein; Demeny et al., 2010), plus carbonate minerals in xenoliths (Lee et al., 2000; van Achterbergh et al., 2002; Demeny et al., 2004a, 2010; Ducea et al., 2005; Scambelluri et al., 2009). ‘Oceanic basalts’ include heating (T N 600–700 °C) and in vacuo crushing analyses of MORB and OIB glasses (Pineau and Javoy, 1983; Des Marais and Moore, 1984; Mattey et al., 1984, 1989; Sakai et al., 1984; Exley et al., 1986; Javoy and Pineau, 1991; Marty and Zimmermann, 1999; MacPherson et al., 2005). The values for peridotitic and eclogitic diamonds are from the compilation of Cartigny (2005) with additional analyses from Cartigny et al. (2009), Deines et al. (2009), De Stefano et al. (2009), Hunt et al. (2012), Melton et al. (2013), Schulze et al. (2013), Smart et al. (2011), Tappert et al. (2005), and Thomassot et al. (2007, 2009). ‘Carbonatites’ include analyses of intrusive, extrusive (lavas and tuffs), veins and dykes (Deines, 1989 and references therein). See Fig. 4 caption for data sources for ‘kimberlite carbonates’.

minerals and diamonds (Chaussidon et al., 1987, 1989; Eldridge et al., 1991; Rudnick et al., 1993; Farquhar et al., 2002; Cartigny et al., 2009; Thomassot et al., 2009), show considerable variability, suggesting heterogeneity in the sub-continental lithospheric mantle (Fig. 6). Moreover, some of the analysed sulphide inclusions in diamonds of eclogitic paragenesis have Δ33S ≠ 0‰, which is consistent with a recycled crustal origin for the sulphur (Farquhar et al., 2002; Thomassot et al., 2009). 2. Carbon and oxygen isotopic geochemistry of carbonates In order to provide stable isotope constraints on the processes affecting carbonates crystallised from mantle magmas, it is essential to define a range of expected δ13C and δ18O values for magmatic carbonates crystallised from pristine mantle melts. The C isotopic composition of most unaltered carbonatite rocks, which are thought to crystallise from mantle-derived magmas, range between −2 and −7‰ (Fig. 5 — e.g., Deines, 1989; Deines and Gold, 1973; Santos and Clayton, 1995; Taylor et al., 1967). This range overlaps with the compositions of diamonds belonging to the peridotitic paragenesis (Fig. 5). Kimberlitic

carbonates have a similar range of δ13C to other mantle materials, which cluster between −3 and − 9‰ (Fig. 5), demonstrating that the mantle is the dominant source of carbon in carbonate phases within kimberlites (e.g., Deines and Gold, 1973; Kirkley et al., 1989; Kobelski et al., 1979; Wilson et al., 2007 — see Section 2.2). Although kimberlites are produced by low-degree melting of mantle rocks (e.g., Wyllie and Huang, 1975; Dalton and Presnall, 1998; Brey et al., 2008; Safonov et al., 2011), carbon occurs in trace abundance in the mantle (~ 50– 400 ppm — Deines, 2002; Gerlach et al., 2002; Javoy et al., 1986; McDonough and Rudnick, 1998; O'Nions and Oxburgh, 1988; Trull et al., 1993) and is extremely incompatible. Therefore, it is likely that carbon is entirely partitioned in the melt phase and carbon isotopes do not fractionate during kimberlite melt formation (Deines, 1989). Deines (1989) estimated a δ18O difference of ~ + 2‰ between carbonate and bulk peridotite compositions under mantle conditions, which suggests that the δ18O values of mantle carbonates range between 7 and 8‰ (δ18Omantle ~ 5–6‰; e.g., Kyser, 1986, 1990). Similarly, calcite in equilibrium with mantle olivine (δ18O ~ 5.2‰; Mattey et al., 1994a) at T = 800–1300 °C would have δ18O values of 6.5 to 8.0‰ (fractionation factor from Chiba et al., 1989). This range is consistent

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

67

OIBs MORBs SCLM peridotite xenoliths sulphides in kimberlite megacrysts peridotitic sulphides in diamonds eclogitic sulphides in diamonds sulphides in SCLM eclogites carbonatite sulphides

δ 34 S (‰)

kimberlite "magmatic" sulphides -15

-5

5

15

25

Fig. 6. Sulphur isotope (‰ δ34S relative to VCDT) variations in mantle-derived magmas, and mantle rocks and minerals: mid-ocean ridge basalts (MORBs — Labidi et al., 2012, 2013), ocean island basalts (OIB — Cabral et al., 2013; Chaussidon et al., 1989; Sakai et al., 1984; Torssander, 1989), peridotite xenoliths entrained in the sub-continental lithospheric mantle (SCLM) by kimberlitic and alkali-basaltic magmas (Muramatsu, 1983; Kyser, 1990; Ionov et al., 1992; Giuliani et al., 2013), sulphides inclusions in kimberlite megacrysts (Chaussidon et al., 1989) and diamonds of peridotitic and eclogitic paragenesis (Chaussidon et al., 1987; Eldridge et al., 1991; Farquhar et al., 2002; Cartigny et al., 2009; Thomassot et al., 2009), sulphides in SCLM eclogite xenoliths sampled by kimberlite magmas (Vinogradov and Ilupin, 1972; Tsai et al., 1979; Chaussidon et al., 1989), sulphides in carbonatites (Gomide et al., 2013 and references therein) and “magmatic” sulphides in kimberlites and orangeites (i.e. with δ34S values close to zero — Chaussidon et al., 1989; Dawson, 1980; Kamenetsky, 2013). ‘SCLM peridotite xenoliths’ include analyses of bulk rocks and sulphate veins.

with the values (6–9‰) commonly observed in intrusive carbonatites (Deines, 1989). 2.1. Processes affecting the C–O isotopic compositions of carbonates in magmatic rocks From the above discussion it is apparent that primary, unaltered carbonates crystallised from pristine, uncontaminated, mantle-derived melts should plot in a restricted δ18O–δ13C field (probably δ18O ~ 6– 9‰; δ13C ~ −2 to –8‰ or even smaller; Fig. 7). Detailed investigations of carbonates in carbonatites, kimberlites, lamprophyres and alkalibasalts have revealed a much broader range of δ18O and δ13C values (Figs. 4 and 5 — e.g., Deines, 1989; Demeny and Harangi, 1996; Kobelski et al., 1979; Santos and Clayton, 1995). Carbonate C–O isotopic compositions that plot outside the mantle field may indicate non-magmatic processes for carbonate formation (e.g., hydrothermal carbonates, sedimentary xenoliths) or crystallisation from magmas that have experienced (Fig. 7) i) magmatic assimilation of, and/or interaction with, carbonate-rich sedimentary rocks 0

δ 13C (‰)

-2 high-T meteoric fluids

-4 -6

hydrothermal fluids

“mantle carbonate box”

low-T meteoric fluids

-8 incorporation of organic carbon

-10

δ 18 O (‰) -12 0

5

10

15

20

Fig. 7. Schematic δ13C (‰ relative to VPDB)–δ18O (‰ relative to VSMOW) diagram showing expected C–O isotopic compositions of mantle carbonates (“mantle carbonate box”) and syn- and post-magmatic processes capable of modifying the compositions of carbonates crystallised from mantle-derived magmas. The arrows show, qualitatively, how each process could affect the C and O isotopes of magmatic carbonates. Modified from Demeny et al. (1998).

(e.g., marine limestone xenoliths and wall rocks) or fluids derived from such lithologies (e.g., Kirkley et al., 1989; Demeny et al., 1994, 1998; Santos and Clayton, 1995; Demeny and Harangi, 1996); ii) magmatic assimilation of organic carbon residing in entrained crustal xenoliths or country rocks; iii) degassing or exsolution of a fluid phase carrying significant amounts of CO2 (e.g., Kobelski et al., 1979; Gerlach and Taylor, 1990; Zheng, 1990a; Demeny et al., 1994; Demeny and Harangi, 1996; Scambelluri et al., 2009); and iv) crystallisation from and/or alteration by deuteric (i.e. late-stage magmatic) fluids, surface waters or hydrothermal fluids (Figs. 7 and 8; e.g., Sheppard and Dawson, 1975; Ustinov et al., 1994; Wilson et al., 2007). Interaction with deuteric fluids has the capacity to change the O isotopic signature of magmatic carbonates towards lighter or heavier compositions, depending on fluid CO2/H2O ratios (Wilson et al., 2007). However, deuteric fluids circulating in kimberlite pipes also carry dissolved CO2 with δ13C values fractionated with respect to the parental melt (Mattey et al., 1990). Therefore, these fluids could alter the C isotopic composition of magmatic carbonates. In order to quantify the effects of deuteric fluids on the composition of kimberlite carbonates, we have calculated the C–O isotopic compositions of carbonates crystallised from, and altered by (or that exchange with), deuteric fluids (Supplementary Material, Appendix 1). Our model shows that calcite crystallisation from deuteric fluids would significantly fractionate not only O, but also C isotopes, particularly at T ≤ 300 °C, and independently from the fluid CO2/H2O ratios (Fig. 8b). As discussed by Wilson et al. (2007), high-CO2/H2O deuteric fluids would crystallise calcite which is progressively more depleted in 18O over 16O with decreasing T, whereas low CO2/H2O fluids would precipitate high-18O/16O calcite. Alteration of magmatic carbonates by deuteric fluids at high fluid/rock ratios (~10) would produce similar C–O isotopic compositions to carbonates crystallised from deuteric fluids (Fig. 8a and c). However, carbonate crystallisation from H2O-rich, CO2-bearing deuteric fluids during cooling is unlikely, because calcite solubility increases with decreasing temperature at low pressure (Fein and Walther, 1987, 1989; Newton and Manning, 2002). Carbonate crystallisation from C–O–H fluids in the upper crust can be triggered by mixing deuteric fluids with low-T, CO2-poor fluids (Zheng, 1990a; Zheng and Hoefs, 1993) such as surface waters and hydrothermal fluids (see model details in Appendix 1). The various combinations of fluids and thermal conditions can produce a range of stable isotope compositions. For example, calcite crystallised from mixed fluids that are dominated by hydrothermal

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a

b

c Fig. 8. Modelled δ13C (‰ relative to VPDB)–δ18O (‰ relative to VSMOW) values for: (a) magmatic kimberlitic calcite (δ13 C = − 6‰; δ18 O = 6‰) altered by kimberlitic deuteric fluids (δ 13 C = − 4‰; δ18 O = 9‰) and meteoric/ground-waters (δ 18 O = − 9‰); and (b) calcite crystallised from kimberlitic deuteric fluids or mixed deuteric–meteoric/ hydrothermal fluids. Panel (c) compares the modelled compositions presented in (a) and (b) and shows the effects of magmatic calcite alteration by hydrothermal fluids (e.g., surface waters enriched in 18O over 16O during circulation in hot kimberlite pipe; δ18O = 0‰) and mixed deuteric–meteoric/hydrothermal fluids. Model parameters: CO2/H2O = 9 (molar ratio) for CO2-rich deuteric fluids; H2O/CO2 = 9 (molar ratio) for H2O-rich deuteric fluids; high fluid/rock ratio = 9; low fluid/rock ratio = 1. The shaded area is the mantle carbonate box. The calculations to produce the curves are presented in Supplementary Material, Appendix 1.

components with lesser contributions from deuteric fluids, would have very similar isotopic compositions to calcite precipitated from equal mixtures of H2O-dominated deuteric fluids and surface waters (Fig. 8b). From the above it is apparent that very high δ18O values (N15‰) in carbonates from magmatic rocks can result from low-T crystallisation (or re-crystallisation) and/or alteration of carbonates by H2O-rich deuteric fluids and/or hydrothermal fluids such as ground-water with high δ18O values resulting from isotope exchange during circulation in hot kimberlite pipes. Surface waters can only lower the δ18O values of magmatic carbonates, unless they react at very low temperatures (≤100 °C); however, in the latter case, evidence of weathering should

be apparent (e.g., Sheppard and Dawson, 1975). Crystallisation from, or alteration by, deuteric fluids or crystallisation from mixtures of deuteric and meteoric fluids, can produce carbonates with extremely variable C–O isotopic compositions, particularly at low temperatures. 2.2. Kimberlite carbonates Carbonates are major constituents of kimberlite rocks and can exceed 50 vol.% in hypabyssal kimberlites (Skinner and Clement, 1979). Carbonates occur as phenocrysts and micro-phenocrysts, segregations with variable amounts of serpentine, and small anhedral to acicular

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

grains in the groundmass (Fig. 9 — Mitchell, 1986; Armstrong et al., 2004). Carbonates may be replaced by serpentine, but can also occur as veins and amygdales of secondary hydrothermal origin (Brookins, 1967; Barrett and Berg, 1975; Dawson, 1980; Exley and Jones, 1983; Podvysotskiy, 1985; Mitchell, 1986; Stripp et al., 2006). Calcite is invariably the most abundant carbonate mineral, whereas dolomite is a subordinate component in most kimberlites. Alkali-rich carbonates are only important constituents in a few kimberlites (e.g., Watkinson and Chao, 1973; Mitchell, 1994; Kamenetsky et al., 2004, 2007b, 2012), although their absence in other kimberlites might be due to alteration and leaching by hydrous fluids (Kamenetsky et al., 2009a). The C and O isotopes of more than 500 kimberlite carbonates mainly from southern African, North American and Russian localities are shown in Figs. 10a and 11. Unfortunately, the majority of these analyses represent bulk rock carbonate measurements, whereas relatively few analyses of separated grains have been reported. The C–O isotopic compositions of different carbonate phases in the same samples can vary by up to several ‰ units, with phenocrysts more likely to preserve pristine mantle signatures than carbonates in groundmass and late-stage segregations (Kobelski et al., 1979; Wilson et al., 2007). Therefore, wholerock carbonate analyses represent the weighted mean compositions of different carbonate phases. Consequently, the range in δ18O–δ13C values for the compiled dataset is large: δ13C = − 11.9–+ 0.2; δ18O = 1.2– 26.6‰ (Table 1; Fig. 10a). 2.2.1. Regional variations of carbon isotopes The majority of kimberlite carbonates (86% of analyses) show δ13C values within the mantle carbonate range, between −2 and −8‰, indicating that secondary processes (Fig. 8) have had a relatively minor impact on the C isotopic signatures. On a regional scale, the southern African kimberlites (including orangeites or Group II kimberlites) have

0.4 mm

a

69

lower δ13C values (median δ13C = −5.6‰) than the Russian kimberlites (−3.9‰ — Table 1; Fig. 11). The North American kimberlites appear to have intermediate δ13C values (median δ13C = −4.9‰); however, this might be due to sampling bias because the majority of North American samples (67%) originate from the Lac de Gras kimberlites. Where several analyses of carbonates from a single kimberlite pipe or cluster of pipes exist, most samples plot within restricted ranges of δ13C values (b3–4‰ units; Fig. 12; Kobelski et al., 1979). Deines and Gold (1973) and Kobelski et al. (1979) identified statistically different average δ13C values for carbonates from different kimberlite pipes, although all values are still within the range of mantle carbonates. These regional and local variations can be attributed to different isotopic compositions of the mantle source region. Alternative interpretations, particularly for the local variations amongst pipes from the same region, include variable alteration processes and/or emplacement conditions. Carbonates in southern Africa orangeites show lighter C isotopes (median δ13C = − 6.5‰) than Group I kimberlites in southern Africa (–5.3‰), North America and Russia (Fig. 11). Given that orangeites are considered to derive from partial melting of highly metasomatised lithospheric rocks (e.g., Smith, 1983; Mitchell, 1995; Becker and le Roex, 2006), it might be inferred that lighter carbon was introduced during metasomatism of the lithospheric source to orangeites. A comparison between the C isotopes of hypabyssal/massive kimberlites and (volcanoclastic) kimberlite breccias is only possible for some Russian kimberlites, where sufficient measurements (n = 33) of kimberlite breccia are available. Fig. 11 shows that there are no systematic variations between samples from these two facies. This is consistent with previous observations of Kobelski et al. (1979) for southern African kimberlites (Kobelski et al., 1979 only reported median values for massive and fragmental kimberlites for each pipe studied).

0.1 mm

b

Ol

Cc Cc

Ol

Phl

0.1 mm

Ol Srp

c

0.2 mm

d Ol

Srp Srp

Cc Pvk

Cc

Spl

Phl Fig. 9. Transmitted-light photomicrographs (a, b, c) and SEM back-scattered electron image (d) showing various textural occurrences of carbonates in the Bultfontein kimberlite (Kimberley, South Africa). (a) Large calcite (Cc) segregation with inclusions of phlogopite (Phl; brown phase) and oxide minerals (opaque phases). (b) Groundmass calcite associated with small grains of brown, serpentinised olivine, occurring interstitially to large olivine (Ol) crystals. (c) Euhedral grains and overgrowths of calcite with polygonal faces protruding into a serpentine (Srp) serpophitic segregation (Pvk: perovskite; Spl: spinel-group minerals). (d) Calcite segregation (or composite phenocryst), partially replaced by serpentine, and calcite micro-phenocryst (red outline).

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1

δ 13 C (‰)

a

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kimberlites worldwide

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δ 18 O (‰)

-13 0 1

3

6

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12

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18

21

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δ 13 C (‰)

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δ 18 O (‰) -13 0

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10

15

20

25

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Fig. 10. (a) Compilation of δ13C (‰ relative to VPDB)–δ18O (‰ relative to VSMOW) values for kimberlitic carbonates worldwide (including orangeite carbonates). Most of data refer to bulk rock carbonate analyses with fewer measurements of carbonate segregations, phenocrysts and aggregates physically extracted from kimberlitic rocks. The shaded area is the mantle carbonate box. See Fig. 4 caption for data source. The full dataset is available in the Supplementary Material, Table S1. (b) Comparison between the δ13C and δ18O values of carbonates from the Kimberley (South Africa), Lac de Gras (Canada) and Udachnaya-East (Russia) kimberlites (see also Fig. 12).

2.2.2. Regional variations of oxygen isotopes In contrast to the C isotopic compositions, most carbonates from kimberlitic rocks do not plot within the mantle carbonate field (δ18O ~ 6–9‰; Figs. 10a and 11). The frequency distribution for O isotopes of kimberlitic carbonates peaks at δ18O ~ 12–13‰ (Fig. 11 — median 13.2‰; Table 1), which is similar to that of extrusive carbonatites (Deines, 1989). North American kimberlites show a larger number of δ18Ocarbonate values within the mantle range (Fig. 11 — median 10.3‰), because several samples from the Lac de Gras kimberlites have preserved carbonates with mantle-like O isotopes (Fig. 12b; Wilson et al., 2007). Conversely, very few Russian kimberlites display δ18Ocarbonate values in the mantle range (median 17.6‰). The trend towards elevated δ18O values, with minor variations in 13 δ C (Fig. 10) exhibited by kimberlite carbonates is probably due to interaction with H2O-rich deuteric and/or meteoric/hydrothermal fluids (Figs. 8 and 12), an inference supported by the different O isotopic values in hypabyssal/massive versus brecciated kimberlitic rocks from Russian pipes (Fig. 11). Carbonates in breccias have higher δ18O values (median 19.4‰) than those in hypabyssal/massive kimberlites (median 16.4‰), which can be explained by a larger influx of hydrous fluids in the brecciated parts of kimberlite pipes. Weathering might be another important factor. Sheppard and Dawson (1975) analysed a highly altered kimberlite sample from the Benfontein kimberlite (South Africa) and found exceptionally high δ18Ocarbonate values (26‰), comparable to the high δ18O values (28–31‰) of South African calcrete carbonates (Potts et al., 2009). However, most kimberlite samples selected for isotopic analyses are usually free of evident weathering features, such as oxidised, clay-rich portions.

Nonetheless, even limited exposure to air has been shown to modify the stable isotopic composition of water-soluble carbonates. For example, Keller and Hoefs (1995) measured the C–O isotopic compositions of lavas erupted at Oldoinyo Lengai (Tanzania), the only active (natro) carbonatitic volcano on Earth. They found considerable compositional and isotopic variations between lavas erupted within few weeks of one another, with the older lavas having been altered by meteoric waters and air humidity (Zaitsev and Keller, 2006; Zaitsev et al., 2008). The altered lavas showed higher δ 18 O (up to ~ 11‰ units) and, where weathering was intense, higher δ13C values (up to ~ 3‰ units — Hay, 1989; Keller and Hoefs, 1995). These studies demonstrate that the carbonate isotopic compositions of rocks hosting abundant alkali-carbonates (e.g., some kimberlites) can be modified within a short period of time after exposure to air.

2.2.3. Local variations of C–O isotopes: Kimberley (South Africa), Lac de Gras (Canada) and Udachnaya-East (Russia) kimberlites A comprehensive evaluation of the processes affecting kimberlite carbonates worldwide cannot be undertaken by examining the global dataset because variable local conditions (e.g., mantle source and wall rock compositions, local hydrogeology and rain water isotopic compositions) affect the carbonate C–O isotopic compositions for each kimberlite locality. Nonetheless, detailed examination of the C–O isotopes of carbonates from single kimberlite pipes and clusters can provide constraints on the syn- and post-magmatic processes affecting carbonate isotopic signatures. Figs. 10b and 12 show the C–O isotopic compositions of bulk rock and separated calcite and dolomite samples from the Kimberley kimberlite cluster, Lac de Gras kimberlite field and the Udachnaya-East kimberlite pipe. In Fig. 12 various vectors and curves representing possible isotopic modifications from mantle values by selected processes are also plotted. Although well-constrained trends cannot be identified, all three localities show sub-sets of samples which scatter towards decreasing δ13C with increasing δ18O values; these isotopic values seem consistent with carbonate crystallisation after extensive CO2 degassing. However, the elevated δ18O values displayed by the Kimberley and Udachnaya-East samples may indicate that carbonate crystallisation was followed by hydrothermal alteration. For the Kimberley and Lac de Gras samples, carbonate compositions with higher δ18O and similar or higher δ13C values than primary magmatic carbonates (‘mantle box’) can be modelled assuming crystallisation from, and/or alteration by, H2O-rich deuteric and/or meteoric/hydrothermal fluids (lines A2, A4, A5, AM6, AM7, C2, C3, CM7 in Fig. 12). However, it should be noted that other processes are capable of producing similar isotopic variations. For example, incorporation of sedimentary carbonates could produce similar elevated δ18O and δ13C values to crystallisation of carbonates from mixed deuteric/hydrothermal fluids dominated by hydrothermal waters (line CM7; Fig. 12a). Wilson et al. (2007) modelled the non-mantle O isotopic compositions of Lac de Gras carbonates and suggested carbonate crystallisation/ alteration by deuteric fluids with variable H2O/CO2 ratios. Our model combines C and O isotopes and reveals that: i) low-δ18O values in carbonates are not consistent with the involvement of CO2-rich deuteric fluids (cf. Wilson et al., 2007); these values are best explained by alteration of magmatic carbonates [δ13C = −4 (dolomite-rich samples) to − 6‰ (calcite-rich samples)] by low-T meteoric/ground-waters (Fig. 12b). Only a few analyses show low δ13C (b−6‰) and δ18O (b−6‰) values, which might be consistent with crystallisation from high-CO2/H2O deuteric fluids (Fig. 12b); however, similar compositions could also be acquired by carbonates that formed after CO2 degassing and subsequently interacted with ground-waters. ii) High-δ18O values coupled with ‘mantle’ δ13C values can be attributed to interactions with either H2O-rich deuteric fluids (Wilson et al., 2007) or mixed deuteric–meteoric/ hydrothermal fluids (Fig. 12b).

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

50

50

δ 13C (‰)

n.

30

δ 18O (‰)

n.

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71

40 orangeite S. Afr. Grou group kimberlite S. Afr. IGroup I

southern Africa (n. 239)

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mantle range

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north America (n. 140)

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Russia (n. 145)

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facies not reported

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mantle range

10

massive/hypabyssal

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Russia north America S. Afr. Group II S. Afr. Group I

worldwide

50

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δ 18O (‰)

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10

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-10

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15

Fig. 11. Compilation of carbon (as ‰ δ13C relative to VPDB) and oxygen (as ‰ δ18O relative to VSMOW) isotopic data for carbonates in kimberlites from southern Africa (Group I or archetypal kimberlites and Group II or orangeites), North America, Russia and worldwide. The Russian kimberlite data shows different oxygen isotopic distributions for carbonates in kimberlitic breccias and in massive and hypabyssal kimberlites. See Fig. 4 caption for data sources.

In the case of the Udachnaya-East samples, all carbonate isotopic compositions are shifted to higher δ18O values, above the mantle range (Fig. 12c). This observation has been interpreted as reflecting contamination of the Udachnaya-East kimberlite by crustal material (Kjarsgaard et al., 2009; Kopylova et al., 2013). However, increased 18 O/16O ratios may be partly due to post-sampling interaction between alkali-carbonates, which are abundant in the kimberlite groundmass (e.g., Kamenetsky et al., 2012), and humid air or meteoric waters. After removing these late-stage effects (estimated at ≥ + 3–+ 4‰ δ18O units based on replicate analyses of samples exposed to air — Kamenetsky V.S., unpublished), it is apparent that some Udachnaya-

East carbonate compositions plot in the mantle range, while other compositions appear to reflect similar processes to those inferred for the Kimberley and Lac de Gras carbonates (i.e. alteration by mixed deuteric–meteoric fluids, CO2 degassing, incorporation of sedimentary carbonates, etc. Fig. 12c). Importantly, the majority of samples showing very high δ18Ocarbonate values (N18–19‰) are kimberlite breccias, i.e. rocks more affected by exchange with hydrous fluids. Recently, Kopylova et al. (2013) suggested that carbonates crystallised from modern basinal brines with low δ13C values (b −10‰; i.e. enriched in C derived from organic material) are abundant in the Udachnaya-East kimberlite. However, modern brines in the Udachnaya country rocks

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Table 1 Statistical values for the C and O isotopes of kimberlite carbonates worldwide. δ13C (‰ relative to VPDB)

No.

Median

Min

Max

Average

Southern Africa Group I kimberlite Southern Africa orangeite Southern Africa total North America Russia hypabyssal/massive Russia breccia Russia total Worldwide

190 49 239 140 82 33 145 538

−5.3 −6.5 −5.6 −4.9 −3.8 −3.5 −3.9 −5.0

−11.8 −11.9 −11.9 −10.0 −9.9 −9.9 −9.9 −16.6

0.2 −1.3 0.2 −1.2 −1.3 0.1 0.1 0.9

−5.7 −6.4 −5.8 −5.0 −4.2 −3.6 −4.3 −5.2

δ18O (‰ relative to VSMOW)

No.

Median

Min

Max

Average

Southern Africa Group I kimberlite Southern Africa orangeite Southern Africa total North America Russia hypabyssal/massive Russia breccia Russia total Worldwide

190 49 239 140 82 33 145 538

12.5 13.4 12.5 10.3 16.4 19.4 17.6 13.2

6.6 7.9 6.6 1.2 7.0 13.8 7.0 1.2

26.5 21.5 26.5 26.6 23.0 25.5 25.5 26.6

13.0 13.6 13.1 10.5 16.3 19.7 17.5 13.7

are characterised by low δ 13 C coupled with low δ 18 O values (mainly −16.5 to −2.0‰; Alexeev et al., 2007) and Fig. 12c shows no evidence for carbonate crystallisation from (and/or interaction with) brines of this isotopic composition. In summary stable isotope studies of kimberlite carbonates can provide important, though not conclusive, constraints on the processes affecting kimberlite magmas during and after emplacement. Coupling stable isotope measurements with detailed textural observations and radiogenic (i.e. Sr) isotope determinations (e.g., Demeny et al., 2004a) can provide constraints to discriminate the origin of carbonates in kimberlite rocks. Primary magmatic values are commonly preserved in calcite phenocrysts and serpentine-free segregations (Wilson et al., 2007). Conversely, analyses of bulk rock carbonates, carbonate– serpentine segregations and carbonate aggregates have returned highly variable compositions, particularly with respect to oxygen isotopes (Kobelski et al., 1979; Wilson et al., 2007). These isotopic signatures can be produced by one or more of several processes (Fig. 12): interaction with H2O-rich deuteric fluids mixed with variable amounts of meteoric waters and/or hydrothermal fluids, alteration by surface waters and hydrothermal fluids, incorporation of sedimentary material or related fluids and fractionation of CO2-bearing gaseous or fluid phases. However, only a limited number of samples (~10–15%) show isotopic signatures consistent with extensive magmatic degassing or fractionation of CO2-rich fluids prior to carbonate crystallisation (Fig. 12). This evidence suggests that most CO2 is retained in the kimberlite magma before carbonate crystallisation and is consistent with the majority of measured carbonates being in hypabyssal kimberlites, which were emplaced non-explosively. Models of kimberlite magma ascent invoking abundant exsolution of CO2 as the principal mechanism to trigger fast and turbulent magma ascent (e.g., Wilson et al., 2007; Russell et al., 2012) appear therefore to be applicable only to a limited number of kimberlite pulses. In addition, there is limited evidence for the involvement of CO2-dominated fluids in the crystallisation/ alteration of kimberlite carbonates, which agrees with petrological investigations of volcanoclastic kimberlites (Stripp et al., 2006; Buse et al., 2010) and with studies of the resorption features of diamonds in kimberlites (Fedortchouk et al., 2010). 3. Oxygen isotopes of serpentine minerals in kimberlites In many kimberlite rocks serpentine comprises 20 to 50% of the groundmass mineralogy (Skinner and Clement, 1979). As summarised by Mitchell (1986, 2013), serpentine in kimberlites occurs as i) pseudomorphic lizardite replacing olivine, carbonates and other magmatic minerals (Fig. 13); ii) discrete segregations of serpophite

(i.e. cryptocrystalline serpentine; Fig. 9c) with variable amounts of calcite or dolomite; and iii) prograde serpophitic serpentine replacing preexisting lizardite. Lizardite, chrysotile and polygonal serpentine (mixture of lizardite and variable amounts of chrysotile) are the typical serpentine polytypes in kimberlites (Mitchell and Putnis, 1988; Stripp et al., 2006). Lizardite has been inferred to crystallise at T b 380 °C, and chrysotile at T b 260 °C (Stripp et al., 2006). The origin of serpentine in kimberlites is an important, but controversial issue for kimberlite petrology, because serpentine accommodates most of the water present in bulk kimberlitic rocks. The formation of serpentine from deuteric (i.e. late-stage magmatic) fluids implies a high H2O content in kimberlite melts (e.g., Kopylova et al., 2007; Mitchell, 2008, 2013; Kjarsgaard et al., 2009; Mitchell et al., 2009). For example, it has been argued that calcite–serpentine segregations in kimberlitic rocks crystallised from deuteric fluids (Mitchell, 1986, 2008), with stable isotope data (i.e. δ18Ocalcite higher than mantle values, δ13C values within the mantle range) used to support this inference (Wilson et al., 2007). However, our calculations (Figs. 8 and 12) demonstrate that interaction with mixed deuteric–meteoric/hydrothermal fluids can induce similar isotopic signatures in carbonates. Moreover, Sparks (2013) has questioned the capacity of kimberlite magmas to retain sufficient dissolved water (several % or more) in the upper crust, to form the large amounts of serpentine found in kimberlitic rocks. Detailed petrographic studies of volcanoclastic kimberlites have revealed that the extensive serpentinisation of kimberlites may be related to interactions with late-stage, low-CO 2 hydrothermal fluids (Stripp et al., 2006; Hayman et al., 2009; Buse et al., 2010; Porritt et al., 2012). During their upward migration to surface, kimberlitic magmas are thought to evolve towards carbonatitic compositions (e.g., Dawson and Hawthorne, 1970; Zurevinski and Mitchell, 2011; White et al., 2012) and should ultimately exsolve CO2-rich deuteric fluids (see Sparks et al., 2006, 2009). This contrasts with evidence for extensive serpentinisation of kimberlites by CO2-poor hydrous fluids. However, C–O isotopic models developed by Santos and Clayton (1995) and the current study, suggest that deuteric fluids released by carbonatitic and kimberlitic magmas are generally poor in CO2. This is supported by fluid inclusion studies of carbonatitic minerals, which have revealed the presence of aqueous fluids trapped in primary inclusions hosted by calcite and apatite (Samson et al., 1995). Therefore, the evidence supporting the direct crystallisation of serpentine from deuteric or non-magmatic fluids is largely equivocal. Kyser et al. (1999) showed that H isotopes in lizardite and chrysotile are reset rapidly by low-T meteoric fluids. Kyser et al. (1999) analysed the H isotopic composition of serpentine minerals from ultramafic complexes of different ages and found that the δD values of the parental hydrous fluids were consistent with those expected from modern meteoric waters, regardless of the age of serpentinisation. Lizardite and chrysotile-rich serpentinites were more affected than antigorite-rich rocks because of larger grain sizes and lower permeabilities of the latter, which prevented extensive H exchange with meteoric water. Conversely, O isotopes were shown to be more resistant to late-stage re-equilibration (Kyser et al., 1999), and can provide useful constraints on the source(s) of serpentinising fluids in kimberlites. Wenner and Taylor (1974) were the first to analyse the O isotopes of lizardite–chrysotile segregations in the Moses Rock kimberlite (Utah, USA). They found uniform δ18O values of between 6.7 and 7.8‰ (Fig. 14), and attributed this isotopic signature to kimberlite serpentinisation by meteoric waters that exchanged with country rocks. A similar conclusion was reached by Sheppard and Dawson (1975), Ukhanov and Devirts (1983) and Ukhanov et al. (1986), who measured the H and O isotopes of serpentine from southern African (Monastery, Kao; δ18O = 1.7–1.9‰) and Yakutian kimberlites (δ18O = 3.7–10.0‰; Fig. 14). Sheppard and Dawson (1975) calculated the H–O isotopic compositions of hydrous fluids in equilibrium with serpentine and found that these fluids were significantly different from the “magmatic waters” (i.e. deuteric fluids) in equilibrium with phlogopite macrocrysts, which Sheppard and Dawson assumed to

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

73

a

b

c

Fig. 12. δ13C (‰ relative to VPDB)–δ18O (‰ relative to VSMOW) diagrams for carbonates from: a) Kimberley kimberlites (Deines and Gold, 1973; Sheppard and Dawson, 1975; Kobelski et al., 1979; Kirkley et al., 1989), b) Lac de Gras kimberlites (Fedortchouk and Canil, 2004; Wilson et al., 2007) and c) Udachnaya-East kimberlite (Galimov and Ukhanov, 1989; Khar'kiv et al., 1991; Kamenetsky et al., 2012). The red arrows indicate isotopic variations from magmatic values (i.e. mantle carbonate box) due to various processes that may have affected the kimberlites. The red dotted and continuous lines represent possible isotopic compositions of carbonates crystallised from, and altered by, deuteric and/or meteoric-hydrothermal fluids (using the same compositions of magmatic calcite and deuteric fluids as shown in Fig. 9). In (b) the blue dotted line (labelled A4) assumes magmatic calcite δ13C of −4‰; ‘Wilson_calcite’ and ‘Wilson_dolomite’ refer to calcite-rich and dolomite-rich kimberlitic rocks analysed by Wilson et al. (2007). In (c) the modelled carbonate compositions (blue dotted and continuous lines) are traced by assuming magmatic calcite δ13C of −4‰ and, consequently, deuteric fluid δ13C of −2‰; 3‰ δ18O units have been added to calculated δ18O values to account for postsampling reactions with air humidity and meteoric waters (‘post-sampling’ arrow; see text for explanation); ‘breccia’ indicates carbonate analyses in kimberlite breccias; ‘country rock’ field from Khar'kiv et al. (1991).

have equilibrated with kimberlite magmas. Mitchell (2013) performed in situ analyses of the O isotopic composition of different textural types of fresh serpentine in southern African and north-American kimberlites, using an ion probe. Overall, the serpentine grains showed δ18O values of between −1.8 and +9.1‰ (Fig. 14), with no apparent differences between volcanoclastic and hypabyssal kimberlites. In most kimberlite pipes, early-formed serpentine (e.g., cores of pseudomorphic lizardite)

exhibits higher δ18O values than co-existing later serpentine (e.g., rims of pseudomorphic lizardite and serpophitic segregations; Fig. 14). Mitchell (2013) developed a model whereby the above serpentine isotopic compositions were achieved by reacting pre-existing olivine with meteoric fluids. According to this model, limited amounts of meteoric water were required to produce the observed serpentine O isotopic compositions (water/rock ratio ≤ 0.8 molar); therefore

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A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

0.2 mm

0.1 mm

a

b

Ol Srp Ol

Ol

0.1 mm

Ol

Srp

0.4 mm

c

Ol

d

Ol

Srp’

Srp’’

Srp

Fig. 13. Transmitted-light photomicrographs (a, b) and SEM back-scattered electron image (c,d) showing various textural occurrences of serpentine in the Bultfontein kimberlite, South Africa (a, b, c) and Pipe 1, eastern Finland (d). (a) Serpentine (Srp) pseudomorph of a pre-existing euhedral grain of olivine (Ol). (b) Aggregate of serpentine grains replacing olivine and other groundmass minerals. (c) Two distinct generations of serpentine (Srp′ and Srp″) replacing olivine. (d) Serpentine replacing olivine macrocryst along fractures.

crystallise interstitially to olivine (and the other pre-existing phases) and may inherit negligible oxygen from the olivine. In order to further test the role of deuteric and meteoric waters in the generation of kimberlite serpentine, we have calculated the

Mitchell (2013) concluded that most water was of deuteric origin. Although Mitchell's (2013) calculations can explain the isotopic compositions of pseudomorphic lizardite replacing olivine, its applicability to serpophitic segregations is questionable, because these serpentine types

kimberlite serpentine

volcanoclastic

Kao (Lesotho), segregation Kao (Lesotho), pseudomorphic rim Kao (Lesotho), pseudomorphic core Letseng-la-terae (Lesotho), segregation Letseng-la-terae (Lesotho), pseudomorphic rim Letseng-la-terae (Lesotho), pseudomorphic core Lethlekane (Botswana)

hypabyssal

Lac de Gras, 93T33 kimberlite (Canada) Ham West (Canada), segregation Ham West (Canada), psudomorphic Wesselton (South Africa), pseudomorphic rim Wesselton (South Africa), pseudomorphic core Iron Mountain (USA) Yakutian kimberlites (Russia) Monastery (South Africa) 18

Moses Rock (USA) -2

0

2

4

6

8

10

Fig. 14. Oxygen isotopic (as ‰ δ18O relative to VSMOW) variations in serpentine from worldwide kimberlites. ‘Moses Rock’ (Wenner and Taylor, 1974), ‘Monastery’ (Sheppard and Dawson, 1975) and ‘Yakutian kimberlites’ (Ukhanov et al., 1986) represent bulk serpentine analyses. Serpentine grains from Iron Mountain, Wesselton, Ham West, Lac de Gras, Lethlekane, Letseng-la-terae and Kao were analysed in situ with an ion probe (Mitchell, 2013).

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

O isotopic compositions of serpentine that: i) crystallises directly from hydrous fluids (i.e. serpophitic segregations), and ii) replaces olivine (δ18Oolivine = 5.6‰; Kamenetsky et al., 2008). This modelling employs the serpentine–water isotopic fractionation factors from Zheng

a

b

c

Fig. 15. Modelled δ 18 O Srp (‰ relative to SMOW)–T (°C) equilibrium relations for: (a) serpentine crystallised from kimberlitic deuteric fluids (δ 18 O = 9‰), meteoric/ground-waters (δ18O = − 9‰) and hydrothermal fluid (e.g., evolved meteoric/ground-waters) or mixtures of deuteric and meteoric fluids (δ18O = −2 and 0‰); (b) serpentine produced from alteration of olivine (δ18O = −5.6‰) by deuteric and/or meteoric fluids (with same δ18O values as for panel a) at variable fluid/rock ratios (F/R). In (c) δ18OSrp–δ18OH2O (‰ relative to SMOW) relations are shown for serpentine produced from alteration of olivine (δ18O = − 5.6‰) by hydrous fluids at constant fluid/rock ratio (5 molar) and variable temperatures (200 to 400 °C). ‘δ18OSrp min’ and ‘δ18OSrp max’ indicate the lowest and highest δ18O values reported by Mitchell (2013) for in situ analyses of serpentine in kimberlitic rocks; the high-δ18O serpentine from the Lethlekane kimberlite is exceptional and this data is reported separately (blue band labelled ‘δ18OLethlekane Srp’). The calculations used to construct the curves are presented in Supplementary Material, Appendix 2.

75

(1993) and the equations of Zheng and Hoefs (1993; Supplementary Material, Appendix 2). Distinct O isotopic compositions have been assumed to simulate deuteric fluids (δ18O = 9‰; Sheppard and Dawson, 1975; Wilson et al., 2007), meteoric/ground-water (δ18O = −9‰, arbitrary value) and hydrothermal fluids (e.g., ground-water enriched in 18O over 16O from circulating in the hot kimberlite pipe; δ18O = 0‰, arbitrary value). Mixing of deuteric and meteoric fluids would produce intermediate isotopic compositions, similar to the δ18O values modelled for hydrothermal fluids. The aim of the exercise is to reproduce the O isotopic compositions measured by Mitchell (2013). The first important result is that deuteric fluids alone cannot crystallise serpentine with isotopic compositions compatible with kimberlite serpentine, with the possible exception of serpentine in the Lethlekane volcanoclastic kimberlite (Fig. 15a). Even assuming serpentine formation after interaction with olivine, deuteric fluids can only crystallise serpentine with δ18O b 6.5‰ at extremely low fluid/rock ratio (≤0.1 molar; Fig. 15b). This is clearly incompatible with the large amounts of water required to produce the 20 to 50% serpentine generally observed in kimberlitic rocks (Skinner and Clement, 1979). The formation of lowδ18O (e.g., b 1–2‰) serporphitic segregations can be modelled by direct crystallisation from evolved (i.e. enriched in 18O over 16O) groundwater or mixed deuteric/meteoric fluids (0 b δ18O b −2‰) at temperatures between 200 and 400 °C (Fig. 15a). The range of serpentine isotopic compositions observed in individual kimberlite pipes cannot be modelled in terms of serpentine formation at decreasing temperatures from a fluid of constant O isotopic composition, because at lower temperatures highδ18O serpentine would crystallise, which contradicts petrographic evidence for the early formation of high-δ18O serpentine. Therefore, the higher δ18O values (N2–3‰) of serpentine, which are typical of early pseudomorphic serpentine, are better explained by olivine interaction with evolved ground water mixed with variable amounts of deuteric fluids (δ18Ofluid = 0–+6‰; Fig. 15c), at temperatures of 200–400 °C. A higher contribution by surface water, possibly coupled with a minor contribution of mantle oxygen from replaced olivine (and carbonates), seems the most likely explanation for the low δ18O values of later serpentine (e.g., rims of pseudomorphic lizardite). The exact amount of deuteric fluids in the mixtures cannot be quantified because: i) the O isotopic values of meteoric/ground-water are highly variable (at least −20–0‰; Craig, 1961; Dansgaard, 1964) and are likely to be modified toward higher 18O/16O ratios during water circulation in the hot kimberlite pipe; ii) contributions by other fluids (e.g., from deep hydrothermal circuits) are possible; and iii) other processes such as degassing and isotopic exchange with xenoliths and country rocks can change the O isotopes of deuteric fluids. In conclusion, our modelling predicts that kimberlite serpentinisation requires the influx of large amounts of externally derived fluids. Contributions by deuteric fluids might be significant in the early stages of serpentinisation (e.g., Lethlekane serpentine — Fig. 15), where the high δ18O values of pseudomorphic lizardite require interaction with hydrous fluids characterised by high 18O/16O ratios (see Podvysotskiy, 1985). On the other hand, hydrothermal fluids such as evolved meteoric/groundwater affected by evaporation and isotopic exchange with heated kimberlitic or country rocks could be sufficiently enriched in 18O over 16 O to generate high-δ18O serpentine, with minimal contributions from deuteric fluids. In support to this view, we note that Thakurta et al. (2009) found a similar δ18O range (−3–+10‰) to kimberlite serpentine in serpentine grains from Alaskan ultramafic massifs and ascribed this variation to serpentinisation by evolved meteoric fluids. Upon interaction with country rocks and crustal xenoliths, hydrothermal fluids would also become enriched in Si, which is also important for the formation of serpentine in kimberlites (e.g., Stripp et al., 2006). In our view the formation of serpentine in kimberlites marks the transition to the post-magmatic stage of kimberlites, where surface waters play an increasingly prominent role in altering the composition of kimberlitic rocks. Our model confirms previous inferences that most H2O stored in serpentine is essentially nonmagmatic in origin (Sheppard and Dawson, 1975; Ukhanov and Devirts,

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1983; Ukhanov et al., 1986; Sparks et al., 2006, 2009; Stripp et al., 2006; Buse et al., 2010). 4. Sulphur isotopes 4.1. Processes affecting the sulphur isotopes of magmatic rocks Archetypal (or Group I) kimberlite magmas present radiogenic isotopic compositions (Sr, Nd, Hf) similar to asthenospheric melts like OIBs (e.g., Smith, 1983; le Roex, 1986; Nowell et al., 2004; Tappe et al., 2011, 2012) and therefore might be expected to have asthenospheric δ34S values (i.e. within 1–2‰ units of zero). However, the lithospheric mantle sampled by kimberlite and alkali basalt magmas shows considerable variation (Fig. 6), which is interpreted as being due to the incorporation of subducted crustal material during continent formation (Eldridge et al., 1991; Farquhar et al., 2002; Thomassot et al., 2009) and prolonged enrichment by metasomatic fluids (Ionov et al., 1992; Wilson et al., 1996; Giuliani et al., 2013). Intra-continental basaltic and ultrabasic magmas show δ34S values away from zero, which have been attributed to mantle metasomatism and/or crustal contamination (e.g., Bekker et al., 2009; Fiorentini et al., 2012a,b; Godlevskii and Grinenko, 1963; Harmon et al., 1987; Penniston-Dorland et al., 2012; Schneider, 1970). Contamination is an important consideration for kimberlite geochemistry because kimberlitic magmas traverse thick (at least 150–250 km) sections of mantle and crustal rocks, and interact extensively with wall rocks and entrained xenoliths en-route to the surface (e.g., Mitchell, 2008; Kamenetsky et al., 2009b). The continental crust displays even greater variation than the lithospheric mantle (e.g., Seal et al., 2000; Seal, 2006; Johnston, 2011); thus a detailed knowledge of the crustal section traversed by kimberlite magmas is critical for understanding the possible effects of crustal contamination. The highly variable S isotopic compositions of intra-continental carbonatites (Fig. 6) only partly reflect mantle heterogeneity (Farrell et al., 2010), because other magmatic processes (e.g., degassing, crystallisation) modify the isotopic compositions of sulphides and sulphates in carbonatites (Mitchell and Krouse, 1975), and other magmatic rocks. Fractionation of S-bearing phases will enrich the residual melt of 34 S over 32S under oxidising conditions (i.e. high SO2/H2S ratios in melt) and vice versa (Zheng, 1990b; de Moor et al., 2010; Marini et al., 2011). Nonetheless the progressive decrease of δ34S values in sulphide minerals from evolved carbonatitic magmas is attributed to increasing oxidation of carbonatite magmas undergoing cooling, which leads to crystallisation of sulphates and concomitant 34S-depletion of residual melts (Mitchell and Krouse, 1975; Deines, 1989; Gomide et al., 2013). This means that phase relations need to be considered during modelling of S isotope evolution in crystallising magmas. The transition from sulphide-dominated to sulphate-dominated system occurs within 1–2 log units of the Ni–NiO buffer (Marini et al., 2011), but determining the exact conditions can be complicated because of dependency of fO2 on melt composition and temperature. Kimberlite magmas cover a wide range of oxygen fugacity conditions (ΔNNO − 5 to ΔNNO + 6, where NNO stands for the Ni–NiO buffer) and can evolve towards more oxidising or reducing conditions with cooling and crystallisation (Canil and Bellis, 2007). Therefore, the evolution of δ34S values during the crystallisation and degassing of kimberlite magmas is difficult to predict, and detailed knowledge of local emplacement conditions is required. 4.2. Sulphur isotopes in kimberlitic rocks and minerals Sulphides and sulphates are minor constituents of kimberlitic rocks. Common primary magmatic sulphides include pyrrhotite, pentlandite, chalcopyrite, pyrite and possibly, heazlewoodite, and occur as single crystals or polycrystalline globules disseminated in the groundmass (Dawson, 1980; Mitchell, 1986; Sharygin et al., 2007). Except for pyrite, which appears to precipitate from late carbonate-rich kimberlitic melts, sulphide grains are often partially resorbed. Potassium-rich sulphides

(e.g., djerfisherite, rasvumite) have been identified in rare cases, and appear to occur in equilibrium with other groundmass phases (Clarke, 1979; Sharygin et al., 2007, 2008). Serpentinisation of olivine produces excess Ni, which is usually accommodated in the Ni-rich sulphides (heazlewoodite, millerite) that often accompany serpentine minerals in kimberlites. Among the sulphates, barite with variable concentrations of Sr, is the most common sulphate in the groundmass of kimberlites (e.g., Mitchell, 1986, 1994). Celestine and alkali-rich sulphates are not widely documented (Kamenetsky et al., 2007a), but are common in secondary inclusions hosted by kimberlitic olivine (Golovin et al., 2007; Mernagh et al., 2011). Sulphate anions can also be accommodated in alkali-carbonates (Kamenetsky et al., 2007a), which have been occasionally reported in kimberlites (e.g., Watkinson and Chao, 1973; Kamenetsky et al., 2004, 2012). Only limited S isotopic data are available for kimberlites (Table 2). Muramatsu (1983) reported a δ34S range of − 3.0–+ 9.6‰ for five bulk rock samples from the Kimberley kimberlites, which contained sulphides (pyrite, heazlewoodite, pentlandite) and barite in variable, but unspecified abundances. However, the selected samples included volcanic breccias, altered surface rocks and anomalously S-rich samples, which suggests that degassing, alteration and crustal contamination may have modified primary magmatic compositions. Chaussidon et al. (1989) reported in situ S isotope analyses of pyrrhotite grains from the Bultfontein (Group I) and Roberts Victor (orangeite/Group II) kimberlites (South Africa). The measured δ34S values of between −0.1‰ and +2.5‰ indicate mantle-like S isotopic compositions for the parental melt. Similarly, Dawson (1980) reported δ34S values of 1.2–2.5‰ for pyrite grains extracted from the Saltpeterpan (Group I) and Bellsbank (orangeite/Group II) kimberlites (South Africa), suggesting crystallisation from pristine kimberlite magmas. Dawson (1980) and Tsai et al. (1979) also found high-δ34S grains of pyrite (δ34S = 6.6–12.1‰) from the Bellsbank and Roberts Victor orangeites. However, unambiguous interpretation of these results is not possible, because the origin and parageneses of these grains is unknown. The sulphur isotopic compositions of sulphide and sulphate minerals (pyrite, galena, sphalerite, gypsum, anhydrite) separated from some Yakutian kimberlites (including the Udachnaya-East and Mir pipes), have been documented by Vinogradov and Ilupin (1972). These sulphides and sulphates exhibit δ34S values that range from 15 to 38 and from 16 to 52‰, respectively, except for one pyrite analysis, which has a δ34S value of − 14‰. Crustal country rocks surrounding the Yakutian kimberlites show broadly similar isotopic compositions (5 to 44‰), whereas a single sample of hydrogen sulphide gas from a drill-hole in country rocks has low δ34S of −16.7‰ (Vinogradov and Ilupin, 1972). The above S isotopic compositions suggest that the Yakutian kimberlite sulphides and sulphates are of xenocrystic origin and/or crystallised from fluids derived from the wall rocks (Distler et al., 1987). This interpretation is consistent with the absence of minerals such as sphalerite, galena and gypsum (i.e., some of the phases analysed for S isotopes) in the groundmass of fresh Yakutian kimberlites (e.g., Kamenetsky et al., 2004, 2007b, 2012; Sharygin et al., 2007, 2008). Kamenetsky (2013) analysed S isotopes (32S, 34S) in 4 bulk rock samples from the Udachnaya-East pipe (Table 2; Fig. 16). Three of these samples were fresh and lacked any obvious alteration (e.g., serpentine was very scarce), whereas a fourth sample was included to monitor the effects of minor alteration (presence of secondary gypsum and serpentine-related sulphides). The fresh samples show δ34S values of 1.0–2.2 and 14.3–14.5‰ for the sulphide (monosulphide + disulphide) and sulphate components, respectively. In contrast, the partially altered sample is characterised by higher δ34S values for sulphides and sulphates (4.5 and 19.7‰, respectively). Although the analysed bulk sulphate component also includes S from alkalicarbonates and sodalite in the samples, the occurrence of disseminated grains of aphthitalite, celestine, barite, anhydrite and, in the altered sample, gypsum suggests that sulphate sulphur was largely dominated by the sulphate minerals.

A. Giuliani et al. / Chemical Geology 374–375 (2014) 61–83

77

Table 2 Sulphur isotopes (‰ δ34S relative to VCDT) of kimberlite sulphides, sulphates and bulk rocks. Kimberlite Sulphides Mir (Yakutia, Russia)

Udachnaya (Yakutia, Russia) Geofizcheskaya (Yakutia, Russia) Novinka (Yakutia, Russia) Premier (South Africa) Saltpeterpan (South Africa) Bultfontein (Kimberley, South Africa) Roberts Victor (South Africa)a

Bellsbank (South Africa)a

Sulphates Mir (Yakutia, Russia)

Udachnaya (Yakutia, Russia)

Novinka (Yakutia, Russia) Komsomol'skaya (Yakutia, Russia) Bulk rocks Benfontein (Kimberley, South Africa) De Beers (Kimberley, South Africa) Klipfontein (South Africa) Koffiefontein (South Africa) Loxtondal (South Africa)a Udachnaya-East (Yakutia, Russia)

Phase

δ34S (‰)

Data source

Comment

Pyrite Pyrite Pyrite Sphalerite Galena Galena Pyrite Galena Pyrite Pyrite Pyrite Pyrite Pyrrhotite Pyrite Pyrrhotite Pyrite Pyrite Pyrite Pyrite

52.2 24.1 29.1 22.8 15.0 15.0 38.2 14.5 −14.0 1.3 1.9 2.5 −0.1 2.5 0.0 6.6 1.2 7.2 16.1

(1) (1) (1) (1) (1) (1) (1) (1) (1) (2) (3) (3) (4) (4) (4) (2) (3) (3) (3)

Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase

16.6 25.4 16.1 17.5 43.8 36.0 18.7 23.2 19.0

(1) (1) (1) (1) (1) (1) (1) (1) (1)

Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase Xenocryst/alteration phase

2.0 4.8 9.6 1.8 −3.0 1.0 14.5 1.0 14.3 4.5 19.7 2.2 14.4

(5) (5) (5) (5) (5) (6) (6) (6) (6) (6) (6) (6) (6)

Sill Mica-rich dyke Surface outcrop Breccia

Anhydrite Gypsum Gypsum Celestine Celestine Celestine Gypsum Thaumasite Thaumasite

Total Total Total Total Total Sulphide Sulphate Sulphide Sulphate Sulphide Sulphate Sulphide Sulphate

Xenocryst (?) Xenocryst (?) Xenocryst (?)

Fresh sample (YBK-1) Fresh sample (YBK-1) Fresh sample (YBK-2) Fresh sample (YBK-2) Altered sample (YBK-3) Altered sample (YBK-3) Fresh sample (YBK-4) Fresh sample (YBK-4)

(1) Vinogradov and Ilupin (1972); (2) Tsai et al. (1979); (3) Dawson (1980); (4) Chaussidon et al. (1989); (5) Muramatsu (1983); (6) Kamenetsky (2013). a Orangeite or Group II kimberlite.

Fig. 16. Δ34Ssulphate–sulphide (‰ relative to VCDT) vs T (°C) equilibrium relations for coexisting sulphides and sulphates following the theoretical formulation of Ohmoto and Rye (1979) and the experimental formulation of Miyoshi et al. (1984). ‘UE_fresh’ and ‘UE_altered’ indicate the Δ34Ssulphate–sulphide values of fresh and altered samples of hypabyssal kimberlite from the Udachnaya-East kimberlite (Yakutia, Russia; Table 2; Kamenetsky, 2013).

The low δ34Ssulphide values of the fresh Udachnaya-East samples are consistent with derivation of S from a mantle source, with minimal contributions from local crustal contaminants that contain heavy δ34S values (Vinogradov and Ilupin, 1972). For these samples the Δ34Ssulphide–sulphate values (12.2–13.5‰) correspond to equilibration between sulphides and sulphates at ~ 500–550 °C (Fig. 16; Miyoshi et al., 1984; Ohmoto and Rye, 1979). Microthermometric studies of melt + solid inclusions hosted by olivine grains from the UdachnayaEast kimberlite indicate that these sulphides probably crystallised at temperatures above 800 °C (Golovin et al., 2007) and that at the temperatures recorded by sulphide–sulphate equilibrium (~ 500–550 °C) the residual kimberlite melts were still partially molten (Kamenetsky et al., 2004). We therefore infer that the sulphate minerals crystallised at temperatures N500 °C from residual melts in isotopic equilibrium with the sulphides. The high δ34S values of the Udachnaya-East sulphates could be ascribed to melt contamination by crustal wall rocks prior to sulphate crystallisation (e.g., Kopylova et al., 2013). However, Sr–Nd–Pb isotopic studies of the investigated Udachnaya-East kimberlitic rocks have produced typical kimberlitic values (Maas et al., 2005), which confirms the pristine mantle nature of the parental melt. We propose that the 34S/32S ratio of the residual melt, which crystallised

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the sulphates, increased because the Udachnaya-East kimberlite magma underwent degassing under oxidising open-system conditions after sulphide crystallisation. Nonetheless, sulphides and sulphates were able to achieve isotopic equilibrium after S loss from the system. This process have been previously defined “selective flux of sulphur” (Zheng, 1990b) and is widespread in volcanic rocks (e.g., Hawaiian sea-floor basalts — Sakai et al., 1984; Japanese arc volcanic rocks — Ueda and Sakai, 1984). The enrichment of SO2− in residual kimberlite 4 melts is indicated by the preservation of secondary sulphate inclusions in olivine (Golovin et al., 2007; Mernagh et al., 2011) and supports evolution of the Udachnaya-East kimberlite magma towards oxidising conditions. Finally, the partially altered sample returned slightly heavier isotopic compositions, which is consistent with petrographic evidence of partial contamination by crustal fluids. The higher δ34S values of both sulphide and sulphate components in this sample, indicates that sulphur was introduced both as reduced and oxidised species. In summary, the very few available analyses of fresh samples and individual magmatic sulphide grains (Dawson, 1980; Chaussidon et al., 1989; Kamenetsky, 2013) indicate that: i) kimberlite melts originate from mantle domains with S isotopes similar to, or with slightly higher 34S/32S compared to, the MORB source; ii) this signature is largely unmodified during ascent through the lithospheric mantle; and iii) there are no apparent differences in δ34S values between Group I kimberlites and orangeites (or Group II kimberlites). 5. Conclusions and future directions Stable isotopes represent a powerful tool for investigating the lowtemperature evolution of magmatic rocks and the extent to which parental magmas and crystallised minerals are contaminated by wall rock and external fluid interactions (e.g., Sheppard and Dawson, 1975; Taylor and Sheppard, 1986; Demeny et al., 1994, 1998; Santos and Clayton, 1995; Demeny and Harangi, 1996; Wilson et al., 2007; Fiorentini et al., 2012a,b; Gomide et al., 2013). These are particularly important issues for kimberlites, because kimberlitic magmas entrain abundant xenoliths and are known to interact with mantle and crustal rocks en route to the surface (e.g., Smith et al., 2004; Mitchell, 2008; Brett et al., 2009; Kamenetsky et al., 2009b). In addition, the brecciated and porous nature of kimberlite pipes favours infiltration of surface waters into the hot pipe, and the generation of circulating hydrothermal systems capable of altering kimberlitic rocks after emplacement (e.g., Sparks et al., 2006; Stripp et al., 2006; Kurszlaukis and Lorenz, 2008; Buse et al., 2010). Alternatively, the alteration process might be driven by hydrous fluids of magmatic (i.e. deuteric) origin that exsolved from the crystallising magma (Mitchell et al., 2009; Mitchell, 2013). These processes, together with volatile loss from the magma, prior to and during crystallisation, produce a significant shift in the composition of kimberlitic rocks from their parental melt compositions (Mitchell, 2008; Kamenetsky et al., 2009a, 2012; Sparks et al., 2009), particularly in the composition and concentrations of volatile species (Brooker et al., 2011). The C–O isotopic compositions of kimberlitic carbonates have been investigated extensively, and a global data compilation (Figs. 10 and 11) shows that the O isotopic composition of most carbonates have been severely affected by post-emplacement interaction with hydrous fluids. However, the majority of available analyses represent wholerock carbonates, which means that the reported isotopic data represent mixtures of different textural types of carbonates (e.g., phenocrysts, groundmass grains, veins). Regional and local variations of the C isotopes of kimberlite carbonates (Fig. 11) might represent mantle heterogeneities on different scales. Southern Africa Group I kimberlites and orangeites are characterised by distinct δ13C values, in accord with radiogenic isotopic data (Sr, Nd, Hf). These results suggest that these two groups originated from compositionally distinct mantle sources (e.g., Smith, 1983; Mitchell, 1995; Nowell et al., 2004). The C isotopes of carbonates in hypabyssal and brecciated rocks from some Russian

kimberlites are indistinguishable, whereas the kimberlite breccia carbonates are more enriched in 18O over 16O (Fig. 11). This variation, previously documented for southern African kimberlites (Kobelski et al., 1979), is consistent with more intense circulation of hydrous fluids in the brecciated parts of kimberlite pipes. Detailed studies of carbonate stable isotopic compositions for single kimberlite pipes and clusters (Fig. 12) reveals that the main process deviating δ18O and, to a lesser extent, δ13C values away from the mantle range, is low-T alteration by H2O-rich deuteric and/or meteoric–hydrothermal fluids. Few samples show isotopic signature consistent with CO2 degassing/fluid exsolution, which suggests that most CO2 is retained in the kimberlite magma prior to carbonate crystallisation, consistent with most of analysed carbonates being from hypabyssal kimberlites. The O isotopes of serpentine minerals in kimberlites indicate that large amounts of externally derived ground waters were introduced into kimberlite rocks. The O isotopic values of serpentine decrease during the post-emplacement evolution of kimberlitic rocks, with the cores of early-formed pseudomorphic lizardite commonly showing higher δ18O values than later rims of pseudomorphic lizardite and serpophitic segregations (Fig. 14 — Mitchell, 2013). These variations can be explained by progressive dilution of initial mixed deuteric– hydrothermal fluids with surface waters. Although the exact amount of deuteric fluid involved in the process cannot be quantified, it is likely to be low. Importantly, introduction of external ground waters modified the composition of kimberlitic rocks by replacing carbonates with serpentine (i.e. addition of H2O plus Si and Mg from xenoliths and wall rocks, and removal of CO2 and Ca — Sparks et al., 2009) and dissolving water-soluble minerals, such as alkali-carbonates, halides and alkalisulphates (i.e. removal of alkalis, Cl, S — Kamenetsky et al., 2009a, 2012). Limited S isotopic data is available for kimberlitic rocks. For example, analyses of fresh samples from the Udachnaya-East kimberlite show that sulphides preserve pristine asthenospheric mantle S signatures whereas sulphates crystallised after magma degassing under oxidising conditions. Conversely, an altered Udachnaya-East sample, as well as sulphide and sulphate grains from S-rich samples from this kimberlite (Vinogradov and Ilupin, 1972) provide compelling evidence that some portions of the pipe were altered by crustal fluids. Further studies to characterise the S isotopes (including 33S and 36S) of kimberlitic rocks and minerals are required, in conjunction with detailed petrographic and geochemical investigations, to provide constraints on the magmatic evolution of kimberlitic magmas, their emplacement history and mantle source compositions. Such studies may also elucidate whether ancient subducted crust is present in the source of kimberlitic melts (cf. Nowell et al., 2004; Paton et al., 2009), because mass-independent fractionation of S isotopes (i.e. Δ33S ≠ 0‰) was common only during the first ~2 Gyr of Earth history (Farquhar and Wing, 2003). Additional constraints on the processes affecting the volatile compositions of kimberlites may be obtained from the application of O isotopes to other magmatic minerals and by the use of alternative stable isotope systems (e.g., N). The former has been attempted by Sarkar et al. (2011), who analysed the O isotopes of kimberlitic perovskite. They found that perovskite is a robust phase capable of preserving the original magmatic O isotopic composition. Sarkar et al. (2011) interpreted the distinct isotopic signatures shown by different generations of perovskite in the Orapa and Wesselton pipes (southern Africa) to reflect the isotopic compositions of parental kimberlite melts and other processes affecting kimberlite magmas prior to perovskite crystallisation (e.g., degassing, assimilation of crustal xenolith, fractional crystallisation). Modern ion-probe analyses allow precise in situ measurements of the O isotopic composition of olivine (e.g., Bindeman, 2008; Eiler et al., 2011). Kimberlitic olivines generally present xenocrystic cores surrounded by one or more rims each with a distinct chemical composition (Kamenetsky et al., 2008; Brett et al., 2009; Arndt et al., 2010; Pilbeam et al., 2013). In situ analyses of the O isotopic compositions of olivine rims may therefore provide constraints on the evolution of kimberlite magmas. Similarly, detailed information on the emplacement

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conditions of kimberlite magmas can be obtained by in situ analyses of the stable isotopic compositions of carbonates (e.g., core–rim variations of phenocrysts and segregations, variations between different textural types) coupled with radiogenic isotope (e.g., Sr) measurements of carbonates. Investigation of N isotopes of minerals like phlogopite may provide evidence of crustal contamination, because crustal rocks generally display positive δ15N values, whereas mantle rocks, diamonds and mantle-derived melts show negative δ15N values (e.g., Peters et al., 1978; Sakai et al., 1984; Haendel et al., 1986; Boyd et al., 1993; Marty and Humbert, 1997; Dauphas and Marty, 1999; Fischer et al., 2002, 2005; Cartigny, 2005; Yokochi et al., 2009; Mitchell et al., 2010; Palot et al., 2012; Cartigny and Marty, 2013). In conclusion, the nature of primary kimberlite melt compositions remains enigmatic. Stable isotope geochemistry has provided key evidence that the volatile concentrations of kimberlitic rocks largely differ from their parental melts. This discrepancy remains a major obstacle in the quest for the primary composition of kimberlite melts. The application of in situ analytical techniques and the investigation of unexplored stable isotopes (e.g., 33S, 36S, 14N, 15N) represent exciting future directions for kimberlite geochemists, and will certainly provide further insights into the evolution and emplacement conditions of kimberlite magmas. Acknowledgments This work is dedicated to the memory of one of the pioneers of kimberlite research, Barry Dawson, whose life's work has inspired many to follow his quest to understand these fascinating, but enigmatic rocks. We thank Graham Hutchinson for support with SEM imaging at the University of Melbourne, and De Beers Consolidated Mines for providing access to some of the reported samples. We are also grateful to Bruce Wyatt for assisting in the identification of some of the more “obscure” kimberlite samples reported in the literature. Very detailed reviews by Chris Harris and Attila Demeny, and editorial handling by David Hilton considerably improved the original manuscript. AG's PhD research was supported by an International Australian Postgraduate Award and the Albert Shimmins Memorial Fund. This is contribution 433 from the ARC Centre of Excellence for Core to Crust Fluid Systems (http://www. ccfs.mq.edu.au). Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.chemgeo.2014.03.003. References Alexeev, S.V., Alexeeva, L.P., Borisov, V.N., Shouakar-Stash, O., Frape, S.K., Chabaux, F., Kononov, A.M., 2007. Isotopic composition (H, O, Cl, Sr) of ground brines of the Siberian Platform. Russ. Geol. Geophys. 48, 225–236. Arima, M., Kerrien, R., 1988. Jurassic kimberlites from Picton and Varty Lake, Ontario: geochemical and stable isotopic characteristics. Contrib. Mineral. Petrol. 99, 385–391. Armstrong, J.P., Wilson, M., Barnett, R.L., Nowicki, T., Kjarsgaard, B.A., 2004. Mineralogy of primary carbonate-bearing hypabyssal kimberlite, Lac de Gras, Slave Province, Northwest Territories, Canada. Lithos 76, 415–433. Arndt, N.T., Guitreau, M., Boullier, A.M., le Roex, A., Tommasi, A., Cordier, P., Sobolev, A., 2010. Olivine, and the origin of kimberlite. J. Petrol. 51, 573–602. Barrett, D.R., Berg, G.W., 1975. Complementary petrographic and strotium-isotope ratio studies of South African kimberlite. Phys. Chem. Earth 9, 619–635. Becker, M., le Roex, A.P., 2006. Geochemistry of South African on- and off-craton, Group I and Group II kimberlites: petrogenesis and source region evolution. J. Petrol. 47, 673–703. Bekker, A., Barley, M.E., Fiorentini, M.L., Rouxel, O.J., Rumble, D., Beresford, S.W., 2009. Atmospheric Sulfur in Archean Komatiite-Hosted Nickel Deposits. Science 326, 1086–1089. Bindeman, I., 2008. Oxygen isotopes in mantle and crustal magmas as revelaed by single crystal analysis. Rev. Mineral. Geochem. 69, 445–478. Boyd, S.R., Hall, A., Pillinger, C.T., 1993. The measurement of δ15N in crustal rocks by static vacuum mass spectrometry: application to the origin of the ammonium in the Cornubian batholith, southwest England. Geochim. Cosmochim. Acta 57, 1339–1347.

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