Miner Petrol (2010) 98:91–110 DOI 10.1007/s00710-009-0055-4

ORIGINAL PAPER

Iron isotope compositions of record melt generation, crystallization, and late-stage volatile-transport processes

Clark M. Johnson & Keith Bell & Brian L. Beard & Aaron I. Shultis

Received: 6 January 2009 /Accepted: 1 May 2009 /Published online: 30 May 2009 # Springer-Verlag 2009

Abstract Carbonatites define the largest range in Fe iso- fluid-rock or fluid- interactions comes from the tope compositions yet measured for igneous rocks, record- common occurrence of Fe isotope disequilibrium among ing significant isotopic fractionations between carbonate, carbonate, oxide, silicate, and sulfide minerals in the oxide, and silicate minerals during generation in the mantle majority of the carbonatites studied. The common occurrence and subsequent differentiation. In contrast to the relatively of Fe isotope disequilibrium among minerals in carbonatites restricted range in δ56Fe values for mantle-derived basaltic may also indicate mixing of phenocyrsts from distinct (δ56Fe=0.0±0.1‰), calcite from carbonatites have magmas. Expulsion of Fe3+-rich brines into metasomatic δ56Fe values between −1.0 and +0.8‰, similar to the range aureols that surround complexes are expected to defined by whole-rock samples of carbonatites. Based on produce high-δ56Fe fenites, but this has yet to be tested. expected carbonate-silicate fractionation factors at igneous or mantle temperatures, carbonatite magmas that have modestly negative δ56Fe values of ~ −0.3‰ or lower can Introduction be explained by equilibrium with a silicate mantle. More negative δ56Fe values were probably produced by differen- Stable and radiogenic isotope studies of carbonatites have tiation processes, including crystal fractionation and liquid been used to monitor the secular evolution of the sub- immiscibility. Positive δ56Fe values for carbonatites are, continental mantle (e.g. Bell and Rukhlov 2004), the however, unexpected, and such values seem to likely reflect evolution of carbonated melts as they migrate from mantle interaction between low-Fe carbonates and Fe3+-rich fluids to crustal levels (e.g. Harmer 1999), and sub-solidus at igneous or near-igneous temperatures; the expected δ56Fe cooling and fluid/rock interaction (e.g. Deines 1989). values for Fe2+-bearing fluids are too low to produced the Carbonatites range in age from Archean to present, are observed positive δ56Fe values of some carbonatites, found on all continents (Woolley and Kjarsgaard 2008), and indicating that Fe isotopes may be a valuable tracer of redox have distinct chemical compositions relative to silicate conditions in carbonatite complexes. Further evidence for igneous rocks (e.g. Simonetti et al. 1997; Chakmouradian 2006). Although volumetrically small compared to other Editorial handling: A. Simonetti igneous rocks, carbonatites provide unique probes into the C. M. Johnson (*) : B. L. Beard : A. I. Shultis mantle, and, because their ages extend back into the Department of Geology and Geophysics, Archean, they can be used to monitor the chemical and Lewis G. Weeks Hall for Geological Sciences, isotopic evolution of the mantle over much of Earth’s 1215 W. Dayton Street, history. Radiogenic isotope (Sr, Nd, Pb) compositions of Madison, WI 53706-1692, USA e-mail: [email protected] carbonatites have shown, unequivocally, that carbonatite URL: http://www.geology.wisc.edu magmas are of mantle origin, that many have compositions that are similar to those found in OIBs, and that mixing K. Bell between isotopically distinct, carbonatitic melts is common. Isotope Geochemistry and Geochronology Research Centre, 2117 Herzberg Laboratories, Carleton University, Stable and radiogenic isotope disequilibrium among min- Ottawa, ON K1S 5B6, Canada erals, even within the same sample, demonstrates the 92 C.M. Johnson et al. commonly cumulate nature of carbonatites, and is well expulsion of alkali-rich fluids from the carbonate and/or explained by mixing within magma chambers, as well as silicate magmas into the surrounding country rocks. the effects of intrusion cooling and alteration (e.g. Simonetti In this study, we present the first Fe isotope study of and Bell 1994a; Bizzarro et al. 2003; Haynes et al. 2003). carbonatites, including whole-rocks and mineral phases. We Most young carbonatites (<200 Ma) have isotopic compo- show here that the largest range in Fe isotope compositions sitions that are typical of sub-oceanic mantle, pointing to yet measured in igneous rocks is found in carbonatites. The sub-lithospheric sources, similar to the HIMU, EM I, and relatively large Fe isotope fractionations among carbonates, FOZO components defined by oceanic basalts (Bell and silicates, and oxides at igneous temperatures, coupled with the Tilton 2001; Bell and Simonetti 2009). A substantial body large contrasts in Fe contents among these mineral groups, of isotopic data now exists for carbonatites from East makes Fe isotopes a particularly sensitive tracer of processes Africa, which suggests mixing between the HIMU and EM that are commonly invoked in models for carbonatite genesis I mantle components (e.g. Bell and Dawson 1995; Bell and and evolution, including magmatic and fluid evolution, crystal Simonetti 1996). Noble gas compositions indicate mantle fractionation, and liquid immiscibility, (for review, see Lee sources (e.g. Marty et al. 1998; Tolstikhin et al. 2002), and and Wyllie (1994)). The results of this Fe isotope survey of some carbonatites have Li, C, and O isotope compositions carbonatites suggest that Fe isotope fractionations among that are similar to those of oceanic basalts (e.g. Deines silicates, carbonates, oxides, and sulfides rarely record 1989; Keller and Hoefs 1995; Halama et al. 2008). Among isotopic equilibrium in carbonatites. Rather, the Fe isotope the models proposed for the sources of carbonated melts, compositions measured in these minerals record complex isotopic data generally support one involving mantle differentiation pathways, mixing of phenocrysts from distinct upwelling such as plumes/hot spot activity, accompanied in magmas, and late-stage fluid interactions. some cases perhaps by interaction with the continental lithosphere (see discussion in Bell and Simonetti 2009,and references within). Fenitization and the role of fluids in carbonatite Carbonatite complexes commonly contain a wide variety evolution of rock types, and their close spatial association with deep- crustal fracturing and rifting implies an intimate relation The enormous capacity of carbonated mantle-derived between intrusion and tectonism; this is particularly well magma to dissolve CO2 and H2O, along with other volatiles shown by alkalic-carbonatitic complexes in the East such as Cl, F, and S, requires fluid phases to develop and African Rift Valley System (e.g. Bailey 1993) and the evolve during carbonatite magma differentiation. Fluid Trans-Superior Tectonic and Kapuskasing Structural Zones compositions can be estimated from fenites, fluid inclusion of the Superior Province, Canada (e.g. Sage 1991). Most studies of carbonatites themselves, and mineral chemistry, carbonatite complexes take the form of circular- or tear- especially REEs abundances and their distribution patterns. shaped plutons, many associated with silicate rocks of Because CO2 and H2O were the principal volatiles used in miaskitic affinity. Where related alkalic silicate rocks occur, melting experiments, the concept of a “carbothermal fluid” these form large stocks or ring complexes, and the of varied CO2 and H2O ratio was introduced in the literature, carbonatites generally occur as plug-like intrusive bodies but it was suggested that other components, especially the that have diameters <5 km. Carbonatite complexes may halogens, may be just as important as H2O (Gittins 1989). consist solely of carbonatite, usually dolomitic in compo- Fenites, a term coined by Brögger (1921), generally sition, whereas others are associated with silicate rocks, consist of alkali , sodic , and/or alkali normally undersaturated with respect to silica. Silicate amphibole that formed at sub-igneous temperature. Fenites rocks commonly associated with carbonatites include may be zoned relative to a carbonatite intrusion, with an syenite, syenite, , melilitolites, and their innermost part composed of amphibole and pyroxene, and volcanic equivalents, as well as pyroxenites (e.g. King and an outer part rich in biotite (e.g. Le Bas 2008). Metasoma- Sutherland 1960; Le Bas 1977). Calciocarbonatites and/or tised rocks can be broadly divided into sodic or potassic magnesiocarbonatite make up the bulk of the carbonatites varieties, although other, more complicated classification within a given complex, but late-stage carbonatites do schemes have been developed (Morogan 1994). Both occur, where these generally comprise only a few percent of sodium- and -rich fenites can occur around a the total carbonatite volume. The most common, late-stage single intrusion, and it has been suggested that these carbonatites are composed of ankerite or form ankeritic distinctions may be related to depth, where potassium dolomite-bearing carbonatites, enriched in REEs, fluorite, fenitization occurs at the upper levels of a carbonatite and incompatible trace elements such as U and Th (Le Bas complex, and sodic fenitization occurs at greater depths (Le 1989). Carbonatite complexes are commonly surrounded by Bas 1989). Fenitization is envisioned to occur by infiltration fenites, which are metasomatic aureoles produced by of fluids from carbonatitic/silicate melts along distinct Iron isotopes in carbonatites 93 pathways, producing a network of alteration minerals. In been inferred from S isotope compositions of sulfides from some cases, the presence of disseminated fenitization ferrocarbonatite from Swartbooisdrift, Namibia (Drueppel products implies diffusive processes, where precursor min- et al. 2006). erals have been completely replaced, reflecting pervasive Pertinent to this survey of Fe isotope compositions of penetration by fenitizing fluids (e.g. White-Pinilla 1996). carbonatites, fluid inclusion studies have also demonstrated Extreme degrees of fenitization has been invoked to explain the importance of Fe-rich chlorides in orthomagmatic formation of magmatic silicate rocks, such as syenites, carbonatite fluids (Bühn et al. 2002; Rankin 2005). The nephelinites and ijolites, by palingenesis of crustal wall rocks high Fe3+/Fe2+ ratios measured in many fenites, as well as after high-grade fenitization (Kramm and Sindern 1998). common occurrence of disseminated iron oxides in fenites, A great diversity of compositions has been inferred for indicates that Fe3+-bearing fluids may be expelled by fenitizing fluids. Wide variations in the inferred composi- carbonatite intrusions during crystallization and solidification tion of fenitizing fluids has been argued to reflect complex (e.g. Le Bas 2008). The presence of hematite, especially in evolution of fluids (e.g. Andersen 1986; Kresten and the low-grade fenites at Alnö, indicates equilibration with a Morogan 1986;Andersen1989; Bühn and Rankin 1999; highly oxidizing fluid (Morogan 1989). Bühn and Rankins’ Rankin 2005). Variables that control fenitization by carbo- (1999) fluid inclusion study of a fenite associated with the natitic fluids include XCO2 of the fluid, temperature Kalkfeld carbonatite complex, Namibia, showed that virtu- gradients, f O2, FeO/MgO ratio, and activity gradients of ally all alkali metals and Cl, and a major proportion F, Th, U, SiO2,Al2O3, and CaO (Morogan 1994). At Alnö, for and Ti, were preferentially partitioned into the fluid. In example, two distinct fluids are thought to have been addition, the fenitizing fluid at Kalkfeld was estimated to involved in fenitization, derived from carbonatitic and ijolitic have contained 3.0 to 4.1 wt. % total Fe. magmatic sources (Morogan and Wooley 1988). Based on > fenitization products, the carbonatitic-type fluid had XCO2 α a a XH2O,high CaO, possibly high Na2O or K2O (or both), Iron isotope geochemistry a very low SiO2, and possibly high F and P contents. Temperatures and f O2 were initially high, but decreased Iron isotope fractionations among aqueous species and sharply with distance from the source. The occurrence of minerals are controlled by redox state and bonding environ- fluorite, including in some low-grade fenites, suggests ments (e.g. Beard and Johnson 2004a; Schauble 2004). interaction with a F-rich fluid (Morogan 1989). Late-stage, Geologic substances that contain Fe3+ tend to have higher selective removal of La and HREE was attributed to passage 56Fe/54Fe ratios relative to Fe2+ substances, with the of a late, highly-oxidizing post-magmatic fluid. important exception of pyrite, where Fe is covalently Direct determination of fluid compositions involved in bonded (Polyakov and Mineev 2000). Based on experi- carbonatite/fluid interaction and country rock/fluid interac- mentally determined or predicted Fe isotope fractionation tion have relied primarily on fluid inclusions in minerals factors, the relative order of increasing 56Fe/54Fe ratios is such as apatite and fluorite. The complexities of fluid- Ca-Mg-Fe carbonate, Fe carbonate, Fe2+ silicate, aqueous mineral interaction are well shown by fluid inclusion Fe2+, , aqueous Fe3+, hematite, and pyrite, where studies of apatite from the Jacupiranga carbonatite, Brazil the relative order for aqueous species may be significantly (Costanzo et al. 2006),wheremagmaticevolutionis changed by chloride (Polyakov and Mineev 2000; Welch et interpreted to have involved crystal fractionation and al. 2003; Wiesli et al. 2004; Anbar et al. 2005; Polyakov et settling of a carbonatite mineral assemblage in a fluid- al. 2007; Hill and Schauble 2008; Shahar et al. 2008). stratified magma chamber. Studies of fluid inclusions in Although Fe isotope fractionations at igneous or mantle apatite from the Fen carbonatite, Norway, identified a temperatures are relatively small, mantle-derived ultramafic magmatic fluid that was rich in CO2 and NaCl, which rocks and minerals have a measureable range in Fe isotope 56 54 evolved during magmatic differentiation to a CO2-free, compositions, and their average Fe/ Fe ratio is about 0.1 water-dominated system that contained higher salinities and per mil (‰) lower than the average of basaltic rocks, densities as solidification progressed (Andersen 1986). In possibly reflecting the effects of melting and/or metasoma- addition, marked changes in HF fugacity indicated the tism (Beard and Johnson 2004b; Poitrasson et al. 2004; presence of two or more independent or semi-independent Williams et al. 2004; Weyer et al. 2005; Williams et al. lines of magmatic descent (Andersen and Austrheim 1991). 2005; Weyer and Ionov 2007). Within typical analytical Changes in fluid composition upon late-stage mixing of uncertainties of ±0.05 to 0.1‰ for 56Fe/54Fe ratios, all fenitizing fluids and low-salinity meteoric waters were basalts appear to be isotopically homogenous. Evolved documented by Williams-Jones and Palmer (2002) based on igneous rocks, however, may, but not always, have fluid inclusions in the fenites from the Amba Dongar relatively elevated 56Fe/54Fe ratios that are ~0.1 to 0.3‰ complex, India. Increases in f O2 during fenitization have higher than primitive basaltic rocks, and these isotopic 94 C.M. Johnson et al. compositions have been interpreted to reflect magmatic of the 35 samples studied are from Africa. Table 1 differentiation processes, including crystal fractionation, summarizes the seventeen complexes studied. Ages range assimilation, and/or fluid exsolution (Poitrasson and Freydier from the Archean to present, and most of the carbonatites 2005; Schoenberg and Von Blanckenburg 2006;Heimannet are sövites (coarse-grained calcite carbonatite). The major- al. 2008;Tengetal.2008). ity of the complexes are associated with well-developed zones of fenitization. Almost all of the samples are either plutonic or hypabyssal, other than those from Oldoinyo Samples Lengai; these latter samples are chemically dissimilar to the other carbonatites analyzed in that the Oldoinyo Lengai We have analyzed carbonatites and their minerals from samples are natrocarbonatites, which have combined alkalis eleven different countries covering three continents. Most of about 40 wt. %, of which ~31 wt. % is Na2O. In addition

Table 1 Carbonatite complexes analyzed for Fe isotope compositions

Name of complex, country Age (Ma) Petrologic details Fenites

AFRICA Bukusu, Uganda 25±2.5 Sövite, ankerite sovite, and rauhagite. Silicate rocks: melteigite, Yes (K-rich) ijolite, pyroxenite, hornblendite, . Sukulu, Uganda <40 Sövite. Silicate rocks: rim of syenite, phonolite dikes. Probably Tororo, Uganda 40 Sövite. Silicate rocks: Pyroxenite, melteigite, ijolite, nepheline syenite. Yes Toror, Uganda 15.5±6 Sövite, ferrocarbonatite, dolomitic sovite. Silicate rocks: trachyte, Yes (K-rich) phonolite, nephelinite melteigite, ijolite, pyroxenite, hornblendite, nepheline syenite. Homa Bay, Kenya 12−1.3 Sövite, alvikite and ferrocarbonatite. Silicate rocks: ijolite, Yes phonolite, nephelinite. Oldoinyo Lengai, Tanzania Still active Natrocarbonatite flows and tuffs. Silicate rocks: phonolitic and Yes nephelinitic tuffs and rare lavas including rare combeite- and (metasomatized melilite-bearing types. Blocks of pyroxenite, ijolite series rocks, blocks) nepheline syenite. Panda Hill, Tanzania 113±6 Mainly sövite but areas of ferrocarbonaite and dikes of Yes dolomite carbonatite. Sengeri Hill, Tanzania Same as Panda? Dolomitic dikes. Yes Dicker Willem, Namibia 49±1 Sövite and alvikite. Silicate rocks: Ijolite xenoliths in Yes carbonatite, trachyte dikes. NORTH AMERICA Borden, Canada 1872±13 Sövite, silicocarbonatite, beforsite dikes. Yes Oka, Canada 110 Sövite and dolomite carbonatite. Silicate rocks: Okaite-jacupirangite Yes (melilitolite-pyroxenite), and ijolite series rocks; lamprophyre dikes. St. Honoré, Canada ca. 650 Sövite, dolomite and ankerite carbonatite. Silicate rocks: Syenite, Yes nepheline syenite, and ijolite. Magnet Cove, USA ca 100 Sövite. Silicate rocks: Pyroxenite, gabbro, jacupirangite, melteigite, Yes ijolite, syenite, trachyte, phonolite. SOUTH AMERICA Jacupiranga, Brazil 130 Sövites, with a zone of dolomitization, and dikes of alvikite and Yes beforsite. Silicate rocks: peridotite, pyroxenite, jacupirangite, ijolite, nepheline syenite, essexite, tinguaite, monchiquite. EUROPE Kaiserstuhl, Germany 16.0 Sövite, alvikites. Silicate rocks: limburgite, phonolite and leucite Unknown tephrite, shonkinite, bergalite, nephelinite, phonolite, monchiquite, essexite and . Kovdor, Russia 365±8 Sövite and dolomite carbonatite. Silicate rocks: Olivinite, Yes pyroxenite, nepheline pyroxenite, melteigite, ijolite, melilite- and monticellite-bearing rocks, nepheline syenite. Siilinjarvi, Finland 2617±10 Sövite, silicocarbonatite. Silicate rocks: Glimmerite, Yes syenite, lamprophyre.

Geological summary of the complexes arranged by country and continent. Most of the petrological details taken from Woolley and Kjarsgaard (2008). Details of the fenites not given, other than Toror. Some of the analytical uncertainties for the ages are unavailable. Some recent data have been used to assess some of the ages. Included among these are those for Kovdor and Siilinjarvi (Rhuklov and Bell, this volume) Iron isotopes in carbonatites 95 to analyses of whole-rock samples, Fe isotope measurements possibility that their Fe isotope compositions may have were made of calcite, dolomite, magnetite, spinel, , been modified by small amounts of oxide and/or silicate pyroxene, biotite, phlogopite, melilite, nepheline, perovskite, mineral contamination. Total dissolution of several hand- pyrite, monticellite, and actinolite. Because carbonatites are picked, >99% pure calcite mineral separates produced Fe known to contain a great variety of minerals, this study contents that significantly exceeded those determined by permitted the first Fe isotope analyses of many of these electron microprobe analysis by previous studies of the minerals from samples that formed at igneous temperatures. same samples (Haynes et al. 2003), suggesting the presence of small amounts of oxide or silicate contamination that was not visible under a binocular microscope. We therefore Analytical methods and nomenclature developed a partial dissolution protocol that completely dissolved carbonate but did not significantly dissolve oxide Mineral separates were obtained by hand picking, with or silicate that may have existed in the carbonate mineral special care given to the carbonate fraction of carbonatite separates (Table 2). Partial dissolution experiments of samples. The very low Fe contents of carbonate minerals in magnetite and silicate (clinopyroxene) minerals involved carbonatites required mineral separate purity in excess of high (85°C) and low (25°C) temperatures, two sieved grain 99%. Mineral separates were washed in doubly distilled sizes, and three different molarities of HCl. Small amounts water (2X H2O) prior to dissolution. Iron isotope analysis of Fe are dissolved from magnetite using 7 M HCl at high of carbonatites presents special challenges due to very low or low temperatures, at 1.6% and 1.1% dissolution at 85°C ratios of Fe to alkali or alkali-earth elements, as well as the and 25°C, respectively. Magnetite that was sized to 1 mm great sensitivity of low-Fe carbonate minerals to small and 0.1 mm in diameter responded differently to partial amounts of contamination by Fe-rich silicate or oxide dissolution, where hot 7 M HCl dissolved 1.6% and 1.2%, minerals, and the new methods we developed to address respectively, and cold 7 M HCl dissolved 1.1% and 0.7%, these issues are described below. respectively. At lower HCl molarity the differences in the amount dissolved between grain sizes and temperature was Partial dissolution experiments and application to carbonate less pronounced. The lowest amount dissolved (0.01%) was iron isotope analyses accomplished using 1 mm magnetite crystals in cold 1 M HCl. Parallel experiments were performed on clinopyroxene. Despite a mineral separate purity of >99%, the very low Fe The 85°C 7 M HCl treatment dissolved 0.06% of the Fe contents of the carbonate mineral separates raises the from a 0.5 mm clinopyroxene grain, whereas the 25°C

Table 2 HCl partial dissolution experiments

Total initial Fe (g) Fe dissolved (µg) Temperature (°C) Acid Grain size (mm) Mineral % Fe dissolved

0.01224 199.8 85 7 M HCl 1 MT 1.63 0.02568 308.4 85 7 M HCl 0.1 MT 1.20 0.01464 8.8 85 7 M HCl 0.5 CPX 0.06 0.01385 150.4 25 7 M HCl 1 MT 1.09 0.01397 100.9 25 7 M HCl 0.1 MT 0.72 0.01201 17.6 85 1 M HCl 1 MT 0.15 0.01173 35.5 85 1 M HCl 0.1 MT 0.30 0.00503 1.2 85 1 M HCl 0.5 CPX 0.02 0.02612 2.7 25 1 M HCl 1 MT 0.01 0.01723 11.3 25 1 M HCl 0.1 MT 0.07 0.00935 21.1 85 0.5 M HCl 1 MT 0.23 0.00956 11.6 85 0.5 M HCl 0.1 MT 0.12 0.00295 1.5 85 0.5 M HCl 0.5 CPX 0.05 0.00377 2.2 25 0.5 M HCl 1 MT 0.06 0.01066 3.8 25 0.5 M HCl 0.1 MT 0.04 0.00943 4.3 25 0.5 M HCl 0.5 CPX 0.05

Magnetite (MT) and clinopyroxene (CPX) from samples P10–208 and P2–670, respectively, and initial Fe contents, as determined by electron microprobe analysis, are from Haynes et al. (2003). Dissolved Fe determined by Ferrozine assay 96 C.M. Johnson et al.

0.5 M HCl treatment dissolved 0.045% of the Fe from a Whole-rock powders of carbonatites required special 0.5 mm clinopyroxene grain. Although the clinopyroxene dissolution procedures due to their very high Ca contents, had a larger amount of Fe partially dissolved at 25°C using which would produce extensive formation of Ca fluorides 0.5 M HCl as compared to magnetite, the amounts of Fe using traditional HF dissolution methods. Magnetite, if dissolved in both minerals would be too small to affect the present, was removed from the samples using a hand magnet Fe isotope composition of the carbonate. Based on these and set aside. The remaining powder was dissolved in 0.5 M results, dissolution of a 99% pure carbonate mineral HCl for ~1 h at room temperature, followed by ~20 min in an separate that contained 1% magnetite or silicate in cold ultrasonic bath to remove the calcite fraction. Following 0.5 M HCl should produce less than 1% Fe contamination centrifugation, the dissolved calcite fraction was set aside, and from magnetite or silicate in the dissolved Fe component, the remaining “silicate” fraction was washed in 2X H2O, which would have no effect on the measured Fe isotope followed by digestion in ~5 ml of 29 M HF and ~500μlof composition of the carbonate. 7 M HNO3 in closed Teflon containers on hotplates for 24 h. It is important to note that we did not use acetic acid The HF-HNO3 solution was then dried down and brought up (HAc) for carbonate dissolution because our previous tests in 5 ml of 8 M HCl. The solution was heated on hotplates in showed that magnetite may undergo incongruent dissolu- closed Teflon containers for at least 12 h and once more tion using HAc (Valaas-Hyslop et al. 2008), which may dried down. The magnetite fraction was digested as noted produce spurious Fe isotope compositions if HAc is used to above, then brought up in 0.5 M HCl. The calcite fraction selectively dissolve carbonate from a mixture of carbonate and the magnetite fraction were then recombined with the and magnetite. In addition to producing anomalous δ56Fe “silicate” fraction to provide a whole-rock analysis. values during partial dissolution of magnetite using HAc, Iron was separated by anion-exchange chromatography.

Valaas-Hyslop et al. (2008) noted that Fe(II)/FeTotal ratios Samples were passed through ion-exchange columns 2–4 changed in the solutions relative to those of magnetite, times, where the greatest number of passes were required reflecting redox changes in solution through interaction for high Ca/Fe ratio samples; previous work has shown that with acetate, and likely formation of a new surface phase. multiple passes do not fractionate Fe isotopes compositions These results were challenged by Von Blanckenburg et al. because yields are ~100% (Beard et al. 2003). Isotopic (2008), who observed no anomalous Fe isotope effects compositions were determined on a Micromass IsoProbe,a when comparing dissolution of natural whole-rock carbo- single focusing, multi-collector inductively-coupled-plasma nates using HCl and HAc, although we note that their tests mass spectrometer with a magnetic sector mass analyzer, did not involve minerals of known isotopic compositions, nor equipped with a Cetac Aridus desolvating micro-concentric did they measure Fe(II)/FeTotal ratios of solution produced by nebulizer and Elemental Scientific spray chamber. Sample partial dissolution of magnetite. We contend that proton- solutions contained 300 ppb Fe, were aspirated at 50μL/min promoted carbonate dissolution without the presence of for eight minutes, resulting in consumption of 120 ng of Fe organic ligands such as acetate remains the safest approach per analysis. Instrumental mass bias and drift was corrected for dissolving carbonate components in the presence of using a standard-sample-standard approach; details can be minor mineral contaminants for Fe isotope analysis. found in Beard et al. (2003) and Albarède and Beard (2004). The accuracy and precision of the Fe isotope measure- Sample processing and mass analysis ments were assessed using analyses of standards, as well as multiple analysis of samples. Separated Fe solutions for 35 Carbonate mineral separates were digested in 0.5 M HCl, samples were analyze 2 or 3 times under different running where the sample was left at room temperature for ~1 h, conditions, and the average difference in the analyses is followed by ~20 min in an ultrasonic bath, then centri- 0.04‰ in 56Fe/54Fe ratios. Five samples were dissolved two fuged. Visual inspection showed that no residue was or three times, and the average difference between the present after centrifugation. This approach was used based multiple dissolutions, which included separate processing on the partial dissolution experiments noted above. Silicate through ion-exchange columns, was 0.06‰ in 56Fe/54Fe minerals were digested in ~5 ml of 29 M HF and ~500μlof ratios. These assessments of precision and accuracy match

7 M HNO3 in Teflon containers on hotplates for 24 h. The those obtained through analysis of ultra-pure Fe standards, solution was then dried down and brought up in 5 ml of and we estimate the average 2σ reproducibility (~2SD) to 8 M HCl. The solutions were heated on hotplates in Teflon be 0.06‰ for 56Fe/54Fe ratios. containers for at least 12 h, and once more dried down. Magnetite mineral separates were digested in ~5 ml of 8 M Nomenclature HCl in closed Teflon containers on hotplates for 24 h, followed by centrifugation to remove any residual silicate minerals or Iron has four naturally occurring stable isotopes, 54, 56, 57, other oxides that are more refractory, such as rutile. and 58, and isotopic compositions have been generally Iron isotopes in carbonatites 97 reported using the three major isotopes, either as 56Fe/54Fe or 57Fe/54Fe ratios. Data are reported here in standard δ notation, where: hi . d56 ¼ 56 54 56 54 3 Fe Fe Fe sample Fe Fe standard 1 10 ð1Þ

56 54 in units of per mil (‰), where ( Fe/ Fe)standard is taken as the average of terrestrial igneous rocks (Beard et al. 2003). Inter-laboratory comparison of Fe isotope ratios can be made by comparison to the measured iron isotope composition of the certified reference material IRMM-014, which has a δ56Fe value of –0.09‰ on the scale. The δ57Fe value may be defined in an analogous manner using the 57Fe/54Fe ratio, and δ57Fe and δ56Fe values should be related in a mass-dependent manner. We discuss Fe isotope fractionations between two phases, A and B, as the difference in the measured δ56Fe values:

56 56 56 $ FeAB ¼ d FeA d FeB ð2Þ

This is an approximation to the isotope fractionation δ56 α ∆56 Fig. 1 Histograms of Fe values for (a) carbonate minerals and factor A-B, which is related to FeA-B through: whole-rocks from this study, (b) whole-rock samples of ultramafic 3 a $ 56 ð Þ rocks, and (c) minerals from ultramafic rocks. Data for (b) and (c) 10 ln AB FeAB 3 from Zhu et al. (2002), Beard and Johnson (2004b), Poitrasson et al. (2004), Williams et al. (2004), (2005), Weyer et al. (2005), Shultis following standard practice. For isotopic fractionations on (2006), and Weyer and Ionov (2007) 56 the order of 1–3‰, use of ∆ FeA-B introduces negligible error relative to analytical uncertainty. Finally, we compare measured Fe isotope fractionations with those predicted value for carbonatites (carbonate minerals or whole-rocks) from theory or measured in experiment using a self- is clearly less than the average of basaltic or ultramafic consistent set of reduced partition function ratios for 56 rocks (Fig. 1), and the range in δ Fe values spans the 56Fe/54Fe, defined as β56/54, following standard practice. largest range yet measured for igneous rocks, from δ56Fe= To a very good approximation, these can be related to −1.0 to +0.8‰. Remarkably, this range spans that com- ∆56Fe by: A-B monly measured in low-temperature aqueous environments $56 3 " 56=54 3 " 56=54 ð Þ (e.g. Johnson et al. 2008). FeAB 10 ln A 10 ln B 4 The Fe isotope compositions of different minerals in the It is important to note, however, that Eqs. (2) and (4)do carbonatites studied vary greatly (Fig. 2). The low-Fe δ56 3 β 56/54 not imply that Fei is equal to 10 ln i . content minerals such as calcite and dolomite have the greatest range in δ56Fe values, where most samples have negative δ56Fe values, but some are positive. The δ56Fe Results values for silicate minerals and magnetite cluster about δ56Fe=0.0‰, although there is considerable spread relative 56 The δ Fe values for carbonate minerals (calcite and to basaltic magmas and silicate minerals from ultramafic dolomite) from the carbonatites studied here, as well as rocks (Figs. 1 and 2). Based on temperatures calculated whole-rock samples, span a significantly greater range than using magnetite-calcite O isotope thermometry for some of that defined by other samples considered to reflect mantle the complexes we have investigated (Haynes et al. 2003), 56 compositions (Fig. 1). Basaltic rocks have δ Fe=0.0± there are no clear correlations between crystallization 0.05‰ (1SD) (e.g. Beard et al. 2003; Poitrasson et al. 2004; temperature and δ56Fe values for carbonate, silicates, or Weyer et al. 2005; Weyer and Ionov 2007). Mantle-derived magnetite. Moreover, the range in δ56Fe values for carbo- xenoliths and Alpine peridotites define a slightly larger nates, silicates, and magnetite is larger than the range 56 range in δ Fe values, although most analyses lie within expected to be in equilibrium with the mantle, based on the 56 56 0.2‰ of δ Fe=0.0 (Fig. 1). In contrast, the average δ Fe Fe isotope fractionations predicted from theory or measured 98 C.M. Johnson et al.

Discussion

Below we touch on several results of our initial Fe isotope survey of carbonatites. First, Fe isotope disequilibrium among minerals is evaluated relative to predicted or experimentally determined Fe isotope fractionation factors. Second, we compare the measured Fe isotope compositions with Li, C, O, and Sr isotope compositions determined on the same samples. Third, the Fe isotope evidence for ex- pulsion of Fe3+-bearing fluids is discussed. Fourth, we bring the Fe isotope variations determined in this study into a model for carbonatite genesis and evolution that is consistent with current petrogenetic models for carbonatites.

Iron isotope fractionations at igneous temperatures

The wide range in measured Fe isotope fractionations among the minerals in carbonatites illustrated in Fig. 2 is explored on a sample-by-sample basis in Fig. 3. Data for minerals are cast relative to magnetite in traditional “δ-δ” plots, and isotopic fractionation lines are shown for magnetite-pyrite and magnetite-olivine, using the β56/54 factors from Table 4 and a reference temperature of 600°C. Of the 14 samples analyzed where at least two minerals were analyzed, ten samples have Fe isotope data for two or more mineral pairs that could be used to evaluate internal isotopic equilibrium. The greatest confidence in evaluating equilibri- um is placed in the magnetite-olivine Fe isotope fractionation factors, where experimental determinations using the “three- isotope-method” (Shahar et al. 2008) have confirmed predicted fractionations based on theory (Polyakov and Mineev 2000; Polyakov et al. 2007). Based on Fe isotope analyses of minerals from a wide variety of silicate- dominated igneous rocks, there does not appear to be a significant Fe isotope fractionation among the various Fig. 2 δ56Fe-crystallization temperature variations for carbonatites silicate minerals, at least in the case of minerals where Fe from this study (Table 3). Crystallization temperatures from O isotope is largely Fe2+ (Polyakov and Mineev 2000; Beard and thermometry (Haynes et al. 2003), except those that have error bars; Johnson 2004b;Polyakovetal.2007; Heimann et al. 2008), these samples are arbitrarily plotted at a temperature of 600±100°C. Shaded curved fields show predicted δ56Fe values as a function of and so the magnetite-olivine fractionation curve of Shahar et temperature, using the β56/54 factors from Table 4. Fields defined by al. (2008) is taken to be applicable to all fractionations two end member Fe isotope compositions for magnetite or silicate between magnetite and Fe2+ silicate minerals. 56 δ Fe=0, which spans the major Fe repositories in the samples Of the ten samples where two or more mineral pairs analyzed that may have been in equilibrium with mantle peridotite and hence control the Fe isotope composition of the samples and constituent were analyzed, only three samples have Fe isotope minerals. Cc=calcite, Ank-ankerite, Fa=fayalite, Mt=magnetite fractionations that indicate magnetite-silicate Fe isotope equilibrium (samples P2-670, STH-08, and P10-208), and these are plotted together in Fig. 3a. Six samples that have in experiment (Fig. 2). Although the relative order of δ56Fe been measured for magnetite-carbonate and magnetite- values for carbonates, silicates, and magnetite that is silicate fractionations do not record equilibrium magnetite- expected based on isotopic fractionation factors is generally silicate fractionations (samples 5–15, BO-203, P2-680, reflected in the average δ56Fe values measured, where P10-202, KO-101, and SIL 106), and this group is plotted 56 56 56 δ FeCarbonate < δ FeSilicate < δ FeMagnetite, the range in the in Fig. 3b. An additional sample is plotted in this group, measured isotopic compositions is considerably larger and MC-1, where spinel, calcite, mica, and montecellite were does not correlate with crystallization temperatures (Fig. 2). analyzed, based on the fact that spinel-silicate fractionations Iron isotopes in carbonatites 99

fractionations experimentally measured by Shahar et al. (2008). These considerations suggest that sample MC-1 is not in Fe isotope equilibrium. Four samples analyzed for magnetite-calcite or magnetite-dolomite fractionations were not analyzed for magnetite-silicate fractionations (samples BD-1499, BD-1485, BD-724, and Phalaborwa), and hence do not have an independent check on Fe isotope equilibrium; this group is plotted in Fig. 3c. Several samples that are not in magnetite-silicate Fe isotope equilibrium (Fig. 3b) appear to have equilibrium silicate-silicate or magnetite-pyrite fractionations. Sample 5–15 has similar magnetite-pyroxene and magnetite- nepheline fractionations, suggesting that pyroxene and nepheline are in Fe isotope equilibrium. Samples P2-680 and BO-203 have similar magnetite-pyroxene and magnetite-mica Fe isotope fractionations, indicating that the two silicates are in isotopic equilibrium. Sample MC-1 has similar spinel-mica and spinel-olivine fractionations, suggesting that although spinel is not in Fe isotope equilibrium with the silicates, the two analyzed silicates are in isotopic equilibrium with each other. Two pyrite- bearing samples analyzed (samples KO-101 and SIL 106) appear to have equilibrium or near-equilibrium magnetite- pyrite fractionations, despite the fact that magnetite, calcite, and mica are not in Fe isotope equilibrium in sample KO- 101, and magnetite, calcite, and amphibole are not in Fe isotope equilibrium in sample SIL 106 (Fig. 3c). The Fe isotope fractionations among minerals are cast relative to temperature in Fig. 4. Crystallization temper- atures for some of the complexes are taken from magnetite- calcite O isotope thermometry of Haynes et al. (2003), which produces the highest O isotope temperatures and minimizes the effects of sub-solidus equilibration. It is important to note, however, that for some of the samples studied by Haynes et al. (2003), calcite-biotite O isotope temperatures were up to 270°C lower than magnetite-calcite temperatures, indicating sub-solidus cooling and isotopic δ56 Fig. 3 Variations in Fe values for magnetite (Mt) relative to other exchange did occur. If independent crystallization temper- co-existing minerals (Cc=calcite, Dol=dolomite, Px=pyroxene, Amph=amphibole, Neph=nepheline, Py=pyrite, Ank-ankerite, Ol=oli- atures were not available, a temperature of 600±100°C was vine, Sid=siderite). Minerals from same sample connected by tie lines. assumed in Fig. 4. Samples grouped relative to those that have equilibrium magnetite- When viewed relative to calculated or estimated crystal- silicate Fe isotope fractionations (a), those that do not have lization temperatures, the three samples that have equilib- equilibrium magnetite-silicate fractionations (b), and those that have no independent control on Fe isotope equilibrium (c). Note in (b) that rium magnetite-silicate fractionations have relatively high one sample analyzed is spinel (sp), not magnetite. Magnetite-olivine magnetite-calcite fractionations (Fig. 4a), exceeding pre- and magnetite-pyrite fractionation lines calculated for a temperature of dicted magnetite-siderite or magnetite-ankerite fractiona- β56/54 600°C, using the factors in Table 4. Isotopic data from Table 3 tions (Table 4). We therefore propose an empirical magnetite-calcite Fe isotope fractionation curve using the β56/54 values in Table 4, which has not been measured in are predicted to be significantly smaller than magnetite- experiments nor explicitly calculated from theory, based on silicate fractionations (Polyakov and Mineev 2000;Polyakov a) the observation that the magnetite-calcite Fe isotope et al. 2007), a conclusion supported by comparing spinel- fractionations for samples that are in magnetite-silicate olivine fractionations measured in ultramafic rocks (Williams equilibrium are generally higher than the predicted et al. 2004; Williams et al. 2005) with the magnetite-olivine magnetite-ankerite or magnetite-siderite fractionations, and 100 C.M. Johnson et al.

b) the predicted and experimentally measured trend of decreasing β56/54 in carbonates with decreasing Fe content (Polyakov and Mineev 2000; Johnson et al. 2005), which would produce increasing magnetite-carbonate fractiona- tions with decreasing carbonate Fe content. When the isotopic data are viewed in the context of our proposed magnetite-calcite fractionation curve, all but one of the samples (BO-23) that do not have equilibrium magnetite- silicate fractionations also have anomalously low magnetite- calcite fractionations (Fig. 4b) Two of the four samples that do not have independent checks on Fe isotope equilibrium (group in Fig. 4c) appear to have equilibrium magnetite- calcite fractionations (samples BD-1485 and BD-724 in Fig. 4c), whereas two samples do not (BD-1499 and Phalaborwa in Fig. 4c). Finally, we note that the two pyrite-bearing samples, which crystallized at markedly different temperatures, have magnetite-pyrite fractionations that generally lie along those predicted by Polyakov and Mineev (2000) and Polyakov et al. (2007), despite their disequilibrium magnetite-silicate Fe isotope fractionations.

Co-variations in Fe, Li, C, O, and Sr isotopes

Our survey of Fe isotopes in carbonatites includes samples (same powders) previously analyzed for Li, C, O, Sr, Nd, and Pb isotopes from the studies of Bell and Tilton (2001), Haynes et al. (2003), and Halama et al. (2008). Calcium carbonate magma in equilibrium with silicate mantle should have a δ56Fe value of ~ −0.3 to −0.8‰, assuming an average silicate mantle δ56Fe value of 0.0 to −0.1‰ (Beard and Johnson 2004b;Williamsetal.2004; Weyer et al. 2005; Williams et al. 2005; Schoenberg and Von Blanckenburg 2006; Weyer and Ionov 2007), and equilibration temper- atures of 400 to 1,000°C, using the β56/54 values in Table 4. That the β56/54 values for carbonates are lower than any other mineral (Table 4) indicates that crystal fractionation of non-carbonate, Fe-bearing minerals (silicates, oxides, sul- fides) will move a carbonatite magma toward more negative δ56 − − ‰ Fig. 4 Variations in Fe isotope fractionations between magnetite and Fe values, lower than the 0.3 to 0.8 values that are co-existing minerals relative to temperature (106/T2). Symbols and expected for carbonate magmas in equilibrium with the abbreviations as in Fig. 3. Sample grouping also follows that of Fig. 3: silicate mantle. In Fig. 5, we compare the Fe, Li, O, and Sr samples that have equilibrium magnetite-silicate Fe isotope fractiona- isotope compositions of the carbonatites studied here with tions (a), those that do not have equilibrium magnetite-silicate fractionations (b), and those that have no independent control on Fe those expected for primary carbonate melts, as well as isotope equilibrium (c). Note in (b) that one sample analyzed is spinel changes that may occur through magmatic differentiation (sp), not magnetite. Temperatures based on O isotope thermometry and fluid interaction. determined on the same samples. For samples that do not have Based on Sr, Nd, and Pb isotopes in East African independent temperature constraints, temperatures were arbitrarily plotted as 600±100°C. Isotopic data from Table 3. Temperature- carbonatites, Bell and Tilton (2001) interpreted the isotopic dependent isotope fractionation curves calculated using β56/54 factors compositions to largely reflect mixing between HIMU and from Table 4 EM I mantle components. Our sample suite includes the carbonatite complexes that lie closest to these end members (EM I: Homa Bay; HIMU: Sukulu), and data for the other complexes analyzed here scatter between these mantle components (Fig. 5a). Similar trends are observed for Nd Iron isotopes in carbonatites 101

isotope compositions (e.g. Beard et al. 2003; Weyer and Ionov 2007). A larger suite of the samples analyzed here has been analyzed for C and O isotopes, including from Oka (Canada), Magnet Cove (USA), and Jacupiranga (Brazil) from Haynes et al. (2003), as well as Sukulu (Uganda), Toror (Uganda), Homa Bay (Kenya), Oldoinyo Lengai (Tanzania), Panda Hill (Tanzania), and Borden (Canada) from Halama et al. (2008). With one exception, all δ18O values for the samples analyzed for Fe isotopes fall within a restricted range that is consistent with derivation from the mantle. Sample DU-365 from Toror has an elevated δ18O value (Fig. 5b), suggesting fluid loss, which will tend to increase δ18O values (e.g. Deines 1989), although it is not clear that fluid loss would produce a sufficient increase in δ18O values. The fact that this sample has a relatively high δ56Fe value is consistent with passage of Fe3+-bearing fluids, as will be discussed below. The δ13C values of the samples analyzed for Fe isotopes vary greatly (Fig. 5c), scattering beyond the range that reflects equilibrium with the mantle. Relatively high δ13C values in carbonatites are generally interpreted to reflect extensive crystal fraction- 13 ation, whereas loss of CO2 should decrease δ C values (e.g. Deines 1989). Extensive crystal fractionation, however, should also produce an increase in δ18O values (Deines 1989), which is not generally observed in the samples studied. An alternative explanation for an increase in δ13C 18 values but not δ O values, is passage of CO2-bearing fluids that produced a net addition of high-δ13C carbon,

rather than CO2 loss. As will be discussed below, samples that have the highest δ56Fe values are interpreted to reflect passage of Fe-bearing solutions, and if such solutions were

CO2 bearing, this may be the explanation for the tendency of high δ56Fe values to be associated with high δ13C values. Eight samples analyzed by Halama et al. (2008) for Li Fig. 5 Variations in Sr, O, C, and Li isotope compositions, as isotopes were analyzed in this study, and the data suggest a available, for samples analyzed in this study for Fe isotope general negative correlation between δ7Li and δ56Fe values compositions. For most samples, whole-rocks were analyzed, but for some O and C isotope compositions, calcite was analyzed, and these (Fig. 5d), although the data base is somewhat limited. 7 were plotted against δ56Fe values determined on the same calcite Halama et al. (2008) interpreted decreasing δ Li values, mineral separate. Dark gray boxes denote fields for primary carbonate relative to mantle values, to reflect expulsion of Li-bearing magma compositions. Arrows denote various processes that may fluids upon extensive crystal fractionation, and they noted a change the isotopic compositions, as discussed in the text. Sr, O, C, δ7 δ13 and Li isotope data from Bell and Tilton (2001), Haynes et al. (2003), broad negative correlation between Li and C values and Halama et al. (2008). Fe isotope data from Table 3 that supports this interpretation. Because fluids rich in Li and Fe would most likely be chloride bearing, it is anticipated that further studies of carbonatite complexes and Pb isotope variations (not shown). For the carbonatites that experienced fluid loss would have strong correlations that have δ56Fe values that reflect equilibrium with the between Li and Fe isotopes, as well as fluid inclusion mantle (δ56Fe < −0.3‰), there is no correlation between Fe evidence for chloride-bearing brines. isotope compositions and Sr, Nd, or Pb isotope ratios, suggesting that Fe isotopes are not a sensitive indictor of Evidence for interactions with Fe3+-bearing fluids EM I or HIMU reservoirs. This conclusion is consistent with the observation that δ56Fe values for basaltic rocks are Calciocarbonatite magmas in equilibrium with the mantle essentially homogeneous regardless of their radiogenic must have negative δ56Fe values, most likely less than 102 C.M. Johnson et al.

−0.3‰, posing a significant interpretive problem for the over the wide range of magnetite-calcite Fe isotope large number of carbonatites and calcites that have δ56Fe > fractionations, although there is a weak tendency for the most −0.3‰. Crystal fractionation of silicates, oxides, and disequilibrium magnetite-calcite fractionations to be associat- sulfides from carbonate magmas cannot explain δ56Fe > ed with the lowest δ56Fe values for magnetite (Fig. 6b). We −0.3‰, as noted above. Although the modal abundances of propose that passage of Fe-bearing fluids through carbonate silicates, oxides, and sulfides are significantly less than magmas or crystallized minerals is the most likely explana- those of carbonate minerals in the samples studied, the tion for the anomalously high δ56Fe values for calcite. The majority of Fe in all of the samples lies in the non- very low Fe contents of calcite makes this mineral carbonate minerals, confirming that crystal fractionation, by particularly sensitive to modification by Fe-bearing fluids, itself, should strongly decrease δ56Fe values in the and hence can explain the strong correlation between the carbonatite complexes studied here. Similarly, generation δ56Fe values for calcite and the magnetite-calcite fractiona- of an immiscible silicate-carbonate melt (e.g. Kjarsgaard tions (Fig. 6). Based on calculated β56/54 factors for Fe- and and Hamilton 1989; Brooker 1998; Kjarsgaard 1998; Lee Fe-bearing fluids (Table 4), however, the δ56 Fe values of and Wyllie 1998) should produce negative δ56Fe values in calcite should decrease upon loss of a Fe-bearing fluid. the carbonate melt component, particularly if peralkaline Although Heimann et al. (2008) documented evidence for silicate magmas were involved; the high alkali contents in increasing the δ56 Fe values of evolved, metaluminous such magmas produce high Fe3+/Fe2+ ratios in the melt due silicate magmas through loss of an Fe-bearing fluid, this to M2+-Fe2+ and M+-Fe3+ charge balance, without signifi- model cannot explain the increase in δ56Fe values for calcite 3+ 56 cant changes in f O2, and Fe -bearing silicates have δ Fe in carbonatites because, under equilibrium conditions, the values ~0.2 to 0.4‰ higher than Fe2+-bearing silicates at δ56Fe values of calcite are lower than those of Fe2+-or igneous temperatures, based on calculated Fe isotope Fe3+-bearing fluids, silicates, or oxides (Table 4). fractionation factors (Polyakov and Mineev 2000), as well A simple fluid-rock mixing model (e.g. Criss 1999) may as experiments involving Fe3+-rich silicate melts (Schuessler explain the trend of decreasing magnetite-carbonate fractio- et al. 2007). The general effect of increasing δ56Fe values nations with increasing δ56Fe values for calcite (Fig. 6), with increasing Fe3+/Fe2+ ratios in silicate minerals has been reflecting net addition of Fe that had high-δ56Fe values, observed in igneous complexes (Schoenberg et al. 2009). rather than fluid loss. A mixing model that incorporates a The strong negative correlation between magnetite- range of Fe contents in calcite and fluid, as well as isotopic calcite Fe isotope fractionations and the δ56Fe value of compositions (Table 5), can produce the spread in δ56Fe calcite (Fig. 6a) indicates that the non-equilibrium values for calcite. Note that the much higher Fe contents of magnetite-calcite fractionations illustrated in Figs. 3 and 4 magnetite, relative to calcite, results in no significant largely reflect anomalously high δ56Fe values for calcite. In change in the δ56Fe values of magnetite in the fluid-rock contrast, δ56Fe values for magnetite are relatively constant mixing model (Fig. 6b). A fluid-rock mixing model that

Fig. 6 Variation in magnetite-calcite Fe isotope fractionations equilibrium (samples in Figs. 3aand4a), magnetite-silicate disequilib- 56 56 (Δ FeMt-Cc)withδ Fe values for calcite (a) and magnetite (b). Gray rium (samples in Figs. 3band4b), or no independent control on Fe 56 56 56 box indicates Δ FeMt-Cc fractionations and δ FeCc and δ FeMt values isotope equilibrium (samples in Figs. 3cand4c). The strong negative 56 that would be in equilibrium with the mantle at igneous temperatures. correlation indicates that the Δ FeMt-Cc fractionations are controlled by 56 Equilibration temperatures for calcite-magnetite fractionation shown as the δ FeCc values. Three models for fluid-rock interaction shown, dashed lines, as calculated using β56/54 factors from Table 4.Notethat where arrows indicate direction of increasing fluid/rock ratios from zero one sample, spinel (sp) was analyzed instead of magnetite. Symbols (in gray box)to10(end of arrow) for Models A, B, and C of Table 5 indicate groupings of data according to magnetite-silicate Fe isotope Iron isotopes in carbonatites 103

Table 3 Iron isotope data for minerals and whole-rocks from carbonatites

Sample Mineral Dissol. ppm Fe δ56Fe δ57Fe

AFRICA Bukusu, Uganda BD-1477 Mt A-1 0.01±0.03 0.02±0.04 A-2 0.01±0.02 0.02±0.03 Sukulu, Uganda SU-100 WR −0.52±0.06 −0.76±0.04 SU 103 CC A-1 2996 −0.24±0.03 −0.27±0.05 A-2 −0.18±0.03 −0.37±0.07 WR −0.55±0.05 −0.79±0.03 BD-1499 Mt A-1 0.02±0.03 0.00±0.03 A-2 0.06±0.06 0.03±0.03 CC A-1 14088 0.25±0.03 0.38±0.03 A-2 0.34±0.05 0.54±0.07 Dol −0.14±0.03 −0.11±0.06 WR −0.07±0.03 −0.17±0.04 Tororo, Uganda BD-1485 Mt 0.01±0.02 0.07±0.03 CC 3175 −0.67±0.03 −0.96±0.03 Dol A-1 −0.85±0.07 −1.25±0.03 A-2 −0.96±0.04 −1.49±0.05 TO-100B WR −0.37±0.03 −0.58±0.05 Toror, Uganda DU-365 WR A-1 −0.04±0.06 −0.08±0.03 A-2 −0.06±0.03 −0.03±0.05 Homa Bay, Kenya Homa Bay CC A-1 4385 −0.72±0.05 −1.07±0.03 A-2 −0.70±0.04 −1.04±0.05 WR −0.45±0.04 −0.57±0.06 Oldoinyo Lengai, Tanzania OL-2 1993 Mt −0.18±0.05 −0.21±0.04 OL-10 1993 Mt −0.20±0.06 −0.22±0.04 BD-114 WR A-1 −0.49±0.04 −0.78±0.04 A-2 −0.42±0.04 −0.71±0.07 BD-118 WR −0.62±0.02 −0.99±0.03 Panda Hill, Tanzania Tan-210 CC A-1 0.38±0.04 0.55±0.04 A-2 0.30±0.04 0.33±0.03 Tan-212 WR A-1 −0.18±0.09 −0.23±0.06 A-2 −0.26±0.04 −0.28±0.09 Tan-213 WR −0.30±0.05 −0.42±0.06 BD-724 Mt A-1 0.15±0.05 0.25±0.04 A-2 0.15±0.04 0.26±0.04 CC 5053 −0.58±0.03 −0.82±0.03 Sengeri Hill, Tanzania Sengeri Hill WR 0.80±0.07 1.17±0.04 Dicker Willem, Namibia 5–15 Mt A-1 −0.22±0.02 −0.29±0.04 A-2 −0.15±0.08 −0.24±0.03 A-3 −0.10±0.05 −0.19±0.07 104 C.M. Johnson et al.

Table 3 (continued)

Sample Mineral Dissol. ppm Fe δ56Fe δ57Fe

CC A-1 12348 −0.36±0.08 −0.57±0.03 A-2 −0.38±0.05 −0.54±0.03 Px 0.02±0.07 0.03±0.06 Neph 0.20±0.03 0.28±0.04 Phalaborwa, South Africa Phalaborwa Mt −0.16±0.04 −0.34±0.03 CC A-1 2405 −0.15±0.04 −0.25±0.03 A-2 −0.15±0.03 −0.20±0.03 NORTH AMERICA Borden, Canada BO-203 Mt A-1 −0.12±0.07 −0.16±0.04 A-2 −0.17±0.03 −0.27±0.03 CC A-1 2232 0.58±0.04 0.89±0.09 A-2 0.52±0.04 0.66±0.04 cpx A-1 0.30±0.07 0.50±0.05 A-2 0.30±0.03 0.44±0.03 Ol A-1 −0.17±0.03 −0.19±0.08 A-2 −0.11±0.03 −0.15±0.05 Phlog 0.25±0.03 0.31±0.04 Oka, Canada P2–670 Mt −0.04±0.03 −0.09±0.04 CC A 596 −0.73±0.04 −1.13±0.04 B −0.77±0.08 −1.17±0.04 Mica −0.30±0.04 −0.57±0.05 cpx −0.28±0.05 −0.31±0.03 P2–680 Mt −0.19±0.04 −0.32±0.07 CC 176 −0.34±0.02 −0.53±0.03 Mica −0.05±0.04 0.01±0.04 cpx A-1 −0.18±0.03 −0.24±0.04 A-2 −0.10±0.04 −0.20±0.05 OC 203 mica −0.09±0.06 −0.16±0.06 melilite A-1 0.06±0.04 0.01±0.04 A-2 −0.02±0.04 −0.04±0.04 OC 302 CC −0.10±0.07 −0.03±0.06 perovskite −0.05±0.07 0.03±0.04 OC 305 Mt A-1 −0.54±0.05 −0.88±0.04 A-2 −0.54±0.02 −0.81±0.04 OC 307 Bio −0.11±0.03 −0.13±0.04 OC 310 CC −0.34±0.05 −0.45±0.04 OC 318 px −0.10±0.04 −0.16±0.05 OC 320 CC −0.57±0.10 −0.85±0.06 Mica A-1 −0.07±0.04 −0.11±0.05 A-2 −0.08±0.04 0.04±0.06 St. Honore, Canada STH-08 Mt 0.08±0.05 0.07±0.06 CC A-1 10151 −0.65±0.05 −0.94±0.03 A-2 −0.72±0.05 −1.08±0.03 Mica A-1 −0.12±0.03 −0.25±0.04 A-2 −0.05±0.03 −0.01±0.03 Iron isotopes in carbonatites 105

Table 3 (continued)

Sample Mineral Dissol. ppm Fe δ56Fe δ57Fe

WR −0.16±0.04 −0.16±0.04 Magnet Cove, USA MC-1 Spinel −0.02±0.06 −0.02±0.03 CC A 100 −0.87±0.04 −1.09±0.03 B −1.12±0.03 −1.46±0.05 Mica −0.48±0.07 −0.59±0.04 Mont A-1 −0.37±0.05 −0.48±0.06 A-2 −0.38±0.03 −0.55±0.03 SOUTH AMERICA Jacupiranga, Brazil P10–202 Mt A −0.04±0.04 0.07±0.05 B 0.01±0.03 −0.03±0.04 CC A-1 169 −0.47±0.04 −0.68±0.04 A-2 −0.49±0.03 −0.60±0.03 B −0.53±0.04 −0.75±0.06 Mica A-1 0.07±0.03 0.10±0.04 A-2 0.11±0.03 0.13±0.04 B 0.12±0.03 0.26±0.03 P10–208 Mt A 0.08±0.03 0.13±0.03 CC A-1 216 −0.35±0.05 −0.52±0.04 A-2 −0.30±0.03 −0.38±0.04 B −0.35±0.07 −0.48±0.10 C −0.34±0.06 −0.52±0.04 Mica 0.06±0.05 0.13±0.05 cpx −0.07±0.06 −0.13±0.05 EUROPE-ASIA Kaiserstuhl, Germany K-102 CC A-1 −0.65±0.05 −0.94±0.03 A-2 −0.60±0.05 −0.91±0.03 Mica A-1 −0.19±0.04 −0.27±0.05 A-2 −0.13±0.03 −0.22±0.03 Kovdor, Russia KO-101 Mt A-1 −0.43±0.05 −0.60±0.03 A-2 −0.37±0.03 −0.51±0.04 CC 632 −0.47±0.03 −0.65±0.03 Mica 0.13±0.04 0.28±0.04 Pyrite A-1 0.08±0.04 0.17±0.04 A-2 0.09±0.03 0.20±0.03 Sillinjarvis, Finland SIL 106 Mt A-1 −0.35±0.06 −0.61±0.04 A-2 −0.33±0.04 −0.51±0.04 CC A-1 4583 0.62±0.03 0.88±0.03 A-2 0.63±0.03 0.97±0.04 Pyrite 0.34±0.04 0.41±0.04 Act A-1 0.14±0.05 0.21±0.04 A-2 0.14±0.04 0.19±0.03

Dissolutions A, B, C indicate different dissolutions of same sample, including separate processing through ion-exchange columns and mass analysis. Analyses noted with “-1”, “-2”, etc. indicate repeat analysis of same solution obtained during ion-exchange chromatography, but under different days. 106 C.M. Johnson et al.

Table 4 Set of self-consistent β56/54 factors

T°C 103ln β56/54 103ln β56/54 103ln β56/54 103ln β56/54 103ln β56/54 103ln β56/54 103ln β56/54 103ln β56/54 a b c d e f II 2-g III h Olivine Magnetite Siderite Ankerite Calcite Pyrite [Fe Cl4] Fe Cl3

300 1.15 1.76 1.12 0.72 0.32 3.20 1.09 2.17 400 0.84 1.28 0.81 0.52 0.23 2.33 0.79 1.63 500 0.63 0.97 0.62 0.39 0.16 1.77 0.60 1.29 600 0.50 0.76 0.48 0.31 0.14 1.39 0.47 1.06 700 0.40 0.61 0.39 0.25 0.11 1.12 0.38 0.89 800 0.33 0.50 0.32 0.21 0.10 0.92 0.32 0.77 900 0.28 0.42 0.27 0.17 0.07 0.77 0.27 0.68 1000 0.23 0.36 0.23 0.15 0.07 0.66 0.23 0.61 a Polyakov and Mineev (2000) b Calculated from Shahar et al. (2008) using β56/54 from a c Polyakov and Mineev (2000) d Polyakov and Mineev (2000) e 56/54 b 56 56 56 This study; calculated using magnetite β from and 22 assuming ∆ FeMt-CC =∆ FeMt-Ank +∆ FeSid-Ank, which is equivalent to assuming 56 ∆ FeMt-Carbonate varies linearly with Fe content, and that calcite contains very minor Fe f Polyakov et al. (2007) g Schauble et al. (2001) h Hill and Schauble (2008) reflects a net addition of Fe through passage of Fe-bearing The slight decrease in δ56Fe values for magnetite for fluids is consistent with a weak correlation between calcite samples that have high δ56Fe values for calcite (Fig. 6) Fe contents and δ56Fe values (not shown; data in Table 3). remains a puzzle. This weak trend cannot be explained by a The fluid-rock mixing model that most successfully net addition of a high-δ56Fe, Fe3+-bearing fluid. Crystalli- explains the data uses a relatively high δ56Fe value for the zation of magnetite from a high-δ56Fe carbonate magma, if Fe-bearing fluid (model C in Fig. 6 and Table 5), which, at in isotopic equilibrium with calcite, would produce high igneous temperatures, would most likely be an Fe3+-bearing δ56Fe values for magnetite. One possible explanation is that fluid based on β56/54 values (Table 4). In this model, an magnetite in the high-δ56 Fe calcite samples equilibrated Fe2+-bearing fluid cannot explain the wide range in δ56Fe with an Fe3+-rich fluid at low temperatures, which could values of calcite, even at very high fluid/rock ratios, given potentially produce a decrease in δ56Fe values for magne- the lower β56/54 factors relative to those of Fe3+-bearing tite, given the relative β56/54 factors (Table 4). In such a fluids. Invoking an Fe3+-bearing fluid is consistent with the model, however, calcite must maintain isotopic disequilib- peralkaline nature of carbonatite-silicate systems, where rium with the fluid and magnetite. It is also possible that the Fe3+/Fe2+ ratios are relatively high as compared to metal- calcite and magnetite reflect, in part, physical mixtures of uminous silicate magmas, the common occurrence of Fe3+ phenocrysts from unrelated magmas. A deeper understand- oxides (magnetite, hematite) in carbonatites, and the abun- ing of the Fe isotope exchange kinetics of carbonate, oxide, dance of Fe3+-oxides in the fenites that surround intrusive and fluid is required to test these models, as well as carbonatite complexes. experimental confirmation of the Fe isotope fractionations

Table 5 Parameters for fluid-rock interaction model

56 56 56 Model Initial δ FeCalcite Initial Fe concentration Initial δ FeFeCl3 Initial Fe concentration δ FeMagnetite calcite (ppm) fluid (ppm)

A −0.75 10,000 0.27 400 0.1 B1 −0.75 100 0.27 400 0.0 C −0.75 100 0.56 10,000 −0.1

56 Magnetite Fe concentration assumed to be stoichiometric, at 724,138 ppm. Initial δ FeFeCl3=+0.27 reflects predicted composition at 800°C, and 56 56/54 initial δ FeFeCl3=+0.56 reflects predicted composition at 600°C, using β factors from Table 4 Iron isotopes in carbonatites 107 between Fe3+- and Fe2+-bearing fluids, oxides, and carbo- magmas in terms of Fe isotope compositions (e.g. Beard nates at igneous temperatures. et al. 2003; Schoenberg and Von Blanckenburg 2006; Weyer and Ionov 2007). Crystal fractionation in carbo- Fe isotope constraints on carbonatite genesis and evolution natite magmas (e.g. King and Sutherland 1960;LeBas 1989; Simonetti and Bell 1994b) will produce a decrease We bring the above discussion into a summary model for in δ56Fe values for carbonate, and this is a likely carbonatite genesis and evolution in terms of Fe isotope explanation for the samples that have the most negative variations in Fig. 7. This model is consistent with current δ56Fe values. Generation or evolution of carbonatite petrogenetic models for carbonatites. Generation of magmas through liquid immiscibility (e.g. Lee and carbonate magmas in the mantle, either in a plume or the Wyllie 1998) will also decrease δ56Fe values in carbo- lithosphere, will produce negative δ56Fe values, natites, given the positive silicate-carbonate fractionation certainly < −0.3‰ at mantle temperatures, but probably factors at igneous temperatures, and this mechanism will not lower than −0.8‰ at relatively low temperatures, as exert the greatest Fe isotope effect at low temperatures, discussed above. It is possible that slightly lower δ56Fe where the fractionation factors are relatively large, or if a values for carbonatite magma may be produced if melting peralkaline, Fe3+-bearing silicate magma is involved. In a of a previously carbonated peridotite (e.g. Dalton and magmatic system where carbonate and silicate magmas Presnall 1998;WyllieandLee1998;Yaxleyetal.1998) coexist, the Fe isotope compositions of the carbonate occurred, assuming such a peridotite had slightly lower magma should be shifted greatest because of its relatively δ56Fe values due to addition of low-δ56Fe value carbonate. low Fe contents, whereas the δ56Fe values of the silicate If mafic silicate magmas were associated with carbonatite magmas would likely remain close to zero to the degree genesis, these should have δ56Fe values of ~0.0±0.1‰, that the majority of the Fe is contained in the silicate given the relatively homogenous nature of basaltic component of the system.

Fig. 7 Cartoon illustrating our preferred model for Fe isotope magmas in the upper crust will release fluids through the outer variations in carbonatites. Silicate magmas in equilibrium with the portions of the intrusive complexes, as well as the surrounding mantle will have δ56Fe values near zero, whereas carbonatites in country rocks, producing fenite zones. Relatively high δ56Fe values equilibrium with the mantle will have δ56Fe values ≤−0.3‰; these for carbonates in carbonatites require addition of a high-δ56Fe fluid, compositions apply to generation of carbonatites in plume environ- which would most likely be an Fe3+-bearing fluid, given the relatively ments or the lithospheric mantle. Crystal fractionation or liquid high β56/54 factors for Fe3+ fluids (Table 4). Based on fluid-rock immiscibility will tend to move carbonatite magmas toward more modeling, and currently available β56/54 factors, the δ56Fe values of negative δ56Fe values, reaching values as low as ~ −1.0‰, whereas calcite-rich carbonatites might be increased up to +0.4‰ by passage these processes will not significantly change the δ56Fe values of of Fe3+-rich fluids related silicate magmas. Residence of crystallizing carbonatite 108 C.M. Johnson et al.

At high levels in the crust, continued crystallization of temperatures and fluid-rock interactions. Iron isotope carbonatite magmas will produce a free fluid phase, which, variations in carbonatites suggest that this relatively new based on the discussion above, is likely to be chloride- and stable isotope system can also provide a tracer of fluid Fe-rich. At Oldoinyo Lengai, for example, the natrocarbo- interactions and cooling history. Many minerals in the natites show that Cl, F, and S are probably constituents of carbonatites studied are out of Fe isotope equilibrium at carbonatite fluids, and the mineralogy of the fenites indicates igneous temperatures, and, considering the large contrasts that the fluid must have contained Ca, Mg, K, Na, and Fe3+ in Fe contents among carbonate, silicate, oxide, and sulfide (e.g. Morogan and Martin 1983;Gittins1989). An NaCl minerals, isotopic disequilibrium likely reflects the effects brine was considered to be the fenitizing fluid at Callander of cooling and fluid/rock interaction. Iron isotope disequi- Bay, Ontario, Canada (Currie and Ferguson 1971). The librium may also record mixing of phenocrysts from abundance of acmitic pyroxene, riebeckite, and hematite in distinct magmas. Iron isotopes, therefore, provide addition- fenites suggests that fenitization occurred at a depth where al evidence for isotopic disequilibrium in carbonatite thefenitizingfluidshadarelativelyhighf O2,probablyclose magmas, which previous stable (e.g. Haynes et al. 2003) to the hematite-magnetite buffer (Gittins 1989). Fenitization and radiogenic (e.g. Simonetti and Bell 1993) isotope of the surrounding granite terrane at the Iron Hills studies have demonstrated. Because Fe isotope fractiona- carbonatite, Colorado, was attributed to an early- and a tions between fluids and minerals/magmas are sensitive to late-stage fluid. Fluid evolution, based mainly on pyroxene oxidation state, Fe isotopes represent a stable isotope and amphibole, began with high Mg/Fe ratios and high total system that is uniquely poised to investigate the redox Fe and Fe 3+, relative to the unaltered country rock, and as state of carbonatite magmas and their fluids. Moreover, the the Mg/Fe ratio decreased, high Fe 3+ aegerine- and very large range in Fe contents of carbonates, oxides, aegerine were formed (White-Pinilla 1996). All fluids from silicates, and sulfides provides a means to investigate Iron Hills had overlapping temperatures between 510 and differential mass-balance responses to differentiation and 560°C (Lowers 2005). fluid-interaction processes. As our understanding of Fe Loss of Fe-rich fluids provides an additional means for isotope fractionation factors improves, as well as Fe producing low-δ56Fe values in the remaining carbonate diffusion rates in minerals, the exact mechanisms respon- magma because Fe-bearing fluids, particularly those that sible for producing Fe isotope distributions among carbo- are Fe3+-rich, will have relatively positive δ56Fe values natite minerals and whole-rocks will be better understood. (Table 4). We propose, however, that the high δ56Fe values in calcite were produced not by fluid loss but by net Acknowledgments C.M.J., B.L.B., and A.I.S. thank the organizers addition of high-δ56Fe iron through passage of Fe3+-rich of this special volume in honor of our co-author Keith Bell, including guest editor Antonio Simonetti. This work was supported by the δ56 ‰ brines, increasing Fe values in calcite up to +0.4 Department of Geology and Geophysics (U.W. Madison), the (circular inset in Fig. 7). These brines must have passed Geological Society of America, and the National Science Foundation through the carbonatites at relatively high temperatures to (grant EAR-0525417). In addition to samples in the collection of K.B., retain their high O isotope temperatures (Haynes et al. samples were provided by J.B. Dawson, D. Moecher, E. Haynes, and M. Spicuzza. Journal reviews were provided by R. Schoenberg and an δ56 2003), as there is no correlation between FeMt-CC anonymous reviewer, whose comments helped to improve the fractionations and O isotope temperatures (Fig. 4). Assum- manuscript. We thank A. Simonetti for editorial handling of the paper. ing a simple intrusion geometry, our proposal predicts a concentrically zoned carbonatite complex, where the inner zones have relatively low δ56Fe values, and the outer parts References of the carbonatites, as well as the fenites surrounding the carbonatites, have relatively high δ56Fe values. 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