ORIGINAL PAPER

Iron isotope compositions of carbonatites recordmelt generation, crystallization, and late-stagevolatile-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-tope compositions yet measured for igneous rocks, record-ing significant isotopic fractionations between carbonate,oxide, and silicate minerals during generation in the mantleand subsequent differentiation. In contrast to the relativelyrestricted range in δ56Fe values for mantle-derived basalticmagmas (δ56Fe=0.0±0.1‰), calcite from carbonatites haveδ56Fe values between −1.0 and +0.8‰, similar to the rangedefined by whole-rock samples of carbonatites. Based onexpected carbonate-silicate fractionation factors at igneousor mantle temperatures, carbonatite magmas that havemodestly negative δ56Fe values of ~ −0.3‰ or lower canbe explained by equilibrium with a silicate mantle. Morenegative δ56Fe values were probably produced by differen-tiation processes, including crystal fractionation and liquidimmiscibility. Positive δ56Fe values for carbonatites are,however, unexpected, and such values seem to likely reflectinteraction between low-Fe carbonates and Fe3+-rich fluidsat igneous or near-igneous temperatures; the expected δ56Fevalues for Fe2+-bearing fluids are too low to produced theobserved positive δ56Fe values of some carbonatites,indicating that Fe isotopes may be a valuable tracer of redoxconditions in carbonatite complexes. Further evidence for

fluid-rock or fluid-magma interactions comes from thecommon occurrence of Fe isotope disequilibrium amongcarbonate, oxide, silicate, and sulfide minerals in themajority of the carbonatites studied. The common occurrenceof Fe isotope disequilibrium among minerals in carbonatitesmay also indicate mixing of phenocyrsts from distinctmagmas. Expulsion of Fe3+-rich brines into metasomaticaureols that surround carbonatite complexes are expected toproduce high-δ56Fe fenites, but this has yet to be tested.

Introduction

Stable and radiogenic isotope studies of carbonatites havebeen used to monitor the secular evolution of the sub-continental mantle (e.g. Bell and Rukhlov 2004), theevolution of carbonated melts as they migrate from mantleto crustal levels (e.g. Harmer 1999), and sub-soliduscooling and fluid/rock interaction (e.g. Deines 1989).Carbonatites range in age from Archean to present, arefound on all continents (Woolley and Kjarsgaard 2008), andhave distinct chemical compositions relative to silicateigneous rocks (e.g. Simonetti et al. 1997; Chakmouradian2006). Although volumetrically small compared to otherigneous rocks, carbonatites provide unique probes into themantle, and, because their ages extend back into theArchean, they can be used to monitor the chemical andisotopic evolution of the mantle over much of Earth’shistory. Radiogenic isotope (Sr, Nd, Pb) compositions ofcarbonatites have shown, unequivocally, that carbonatitemagmas are of mantle origin, that many have compositionsthat are similar to those found in OIBs, and that mixingbetween isotopically distinct, carbonatitic melts is common.Stable and radiogenic isotope disequilibrium among min-erals, even within the same sample, demonstrates the

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

Editorial handling: A. Simonetti

C. M. Johnson (*) : B. L. Beard :A. I. ShultisDepartment of Geology and Geophysics,Lewis G. Weeks Hall for Geological Sciences,1215 W. Dayton Street,Madison, WI 53706-1692, USAe-mail: clarkj@geology.wisc.eduURL: http://www.geology.wisc.edu

K. BellIsotope Geochemistry and Geochronology Research Centre,2117 Herzberg Laboratories, Carleton University,Ottawa, ON K1S 5B6, Canada

commonly cumulate nature of carbonatites, and is wellexplained by mixing within magma chambers, as well asthe effects of intrusion cooling and alteration (e.g. Simonettiand Bell 1994a; Bizzarro et al. 2003; Haynes et al. 2003).Most young carbonatites (<200 Ma) have isotopic compo-sitions that are typical of sub-oceanic mantle, pointing tosub-lithospheric sources, similar to the HIMU, EM I, andFOZO components defined by oceanic basalts (Bell andTilton 2001; Bell and Simonetti 2009). A substantial bodyof isotopic data now exists for carbonatites from EastAfrica, which suggests mixing between the HIMU and EMI mantle components (e.g. Bell and Dawson 1995; Bell andSimonetti 1996). Noble gas compositions indicate mantlesources (e.g. Marty et al. 1998; Tolstikhin et al. 2002), andsome carbonatites have Li, C, and O isotope compositionsthat are similar to those of oceanic basalts (e.g. Deines1989; Keller and Hoefs 1995; Halama et al. 2008). Amongthe models proposed for the sources of carbonated melts,isotopic data generally support one involving mantleupwelling such as plumes/hot spot activity, accompanied insome cases perhaps by interaction with the continentallithosphere (see discussion in Bell and Simonetti 2009, andreferences within).

Carbonatite complexes commonly contain a wide varietyof rock types, and their close spatial association with deep-crustal fracturing and rifting implies an intimate relationbetween intrusion and tectonism; this is particularly wellshown by alkalic-carbonatitic complexes in the EastAfrican Rift Valley System (e.g. Bailey 1993) and theTrans-Superior Tectonic and Kapuskasing Structural Zonesof the Superior Province, Canada (e.g. Sage 1991). Mostcarbonatite complexes take the form of circular- or tear-shaped plutons, many associated with silicate rocks ofmiaskitic affinity. Where related alkalic silicate rocks occur,these form large stocks or ring complexes, and thecarbonatites generally occur as plug-like intrusive bodiesthat have diameters <5 km. Carbonatite complexes mayconsist solely of carbonatite, usually dolomitic in compo-sition, whereas others are associated with silicate rocks,normally undersaturated with respect to silica. Silicaterocks commonly associated with carbonatites includesyenite, nepheline syenite, gabbro, melilitolites, and theirvolcanic equivalents, as well as pyroxenites (e.g. King andSutherland 1960; Le Bas 1977). Calciocarbonatites and/ormagnesiocarbonatite make up the bulk of the carbonatiteswithin a given complex, but late-stage carbonatites dooccur, where these generally comprise only a few percent ofthe total carbonatite volume. The most common, late-stagecarbonatites are composed of ankerite or form ankeriticdolomite-bearing carbonatites, enriched in REEs, fluorite,and incompatible trace elements such as U and Th (Le Bas1989). Carbonatite complexes are commonly surrounded byfenites, which are metasomatic aureoles produced by

expulsion of alkali-rich fluids from the carbonate and/orsilicate magmas into the surrounding country rocks.

In this study, we present the first Fe isotope study ofcarbonatites, including whole-rocks and mineral phases. Weshow here that the largest range in Fe isotope compositionsyet measured in igneous rocks is found in carbonatites. Therelatively large Fe isotope fractionations among carbonates,silicates, and oxides at igneous temperatures, coupled with thelarge contrasts in Fe contents among these mineral groups,makes Fe isotopes a particularly sensitive tracer of processesthat are commonly invoked in models for carbonatite genesisand evolution, includingmagmatic and fluid evolution, crystalfractionation, and liquid immiscibility, (for review, see Leeand Wyllie (1994)). The results of this Fe isotope survey ofcarbonatites suggest that Fe isotope fractionations amongsilicates, carbonates, oxides, and sulfides rarely recordisotopic equilibrium in carbonatites. Rather, the Fe isotopecompositions measured in these minerals record complexdifferentiation pathways, mixing of phenocrysts from distinctmagmas, and late-stage fluid interactions.

Fenitization and the role of fluids in carbonatiteevolution

The enormous capacity of carbonated mantle-derivedmagma to dissolve CO2 and H2O, along with other volatilessuch as Cl, F, and S, requires fluid phases to develop andevolve during carbonatite magma differentiation. Fluidcompositions can be estimated from fenites, fluid inclusionstudies of carbonatites themselves, and mineral chemistry,especially REEs abundances and their distribution patterns.Because CO2 and H2O were the principal volatiles used inmelting experiments, the concept of a “carbothermal fluid”of varied CO2 and H2O ratio was introduced in the literature,but it was suggested that other components, especially thehalogens, may be just as important as H2O (Gittins 1989).

Fenites, a term coined by Brögger (1921), generallyconsist of alkali feldspar, sodic pyroxene, and/or alkaliamphibole that formed at sub-igneous temperature. Fenitesmay be zoned relative to a carbonatite intrusion, with aninnermost part composed of amphibole and pyroxene, andan outer part rich in biotite (e.g. Le Bas 2008). Metasoma-tised rocks can be broadly divided into sodic or potassicvarieties, although other, more complicated classificationschemes have been developed (Morogan 1994). Bothsodium- and potassium-rich fenites can occur around asingle intrusion, and it has been suggested that thesedistinctions may be related to depth, where potassiumfenitization occurs at the upper levels of a carbonatitecomplex, and sodic fenitization occurs at greater depths (LeBas 1989). Fenitization is envisioned to occur by infiltrationof fluids from carbonatitic/silicate melts along distinct

92 C.M. Johnson et al.

pathways, producing a network of alteration minerals. Insome cases, the presence of disseminated fenitizationproducts implies diffusive processes, where precursor min-erals have been completely replaced, reflecting pervasivepenetration by fenitizing fluids (e.g. White-Pinilla 1996).Extreme degrees of fenitization has been invoked to explainformation of magmatic silicate rocks, such as syenites,nephelinites and ijolites, by palingenesis of crustal wall rocksafter high-grade fenitization (Kramm and Sindern 1998).

A great diversity of compositions has been inferred forfenitizing fluids. Wide variations in the inferred composi-tion of fenitizing fluids has been argued to reflect complexevolution of fluids (e.g. Andersen 1986; Kresten andMorogan 1986; Andersen 1989; Bühn and Rankin 1999;Rankin 2005). Variables that control fenitization by carbo-natitic fluids include XCO2 of the fluid, temperaturegradients, fO2, FeO/MgO ratio, and activity gradients ofSiO2, Al2O3, and CaO (Morogan 1994). At Alnö, forexample, two distinct fluids are thought to have beeninvolved in fenitization, derived from carbonatitic and ijoliticmagmatic sources (Morogan and Wooley 1988). Based onfenitization products, the carbonatitic-type fluid had XCO2 >

XH2O, high αCaO, possibly high aNa2O or aK2O (or both),very low aSiO2, and possibly high F and P contents.Temperatures and fO2 were initially high, but decreasedsharply with distance from the source. The occurrence offluorite, including in some low-grade fenites, suggestsinteraction with a F-rich fluid (Morogan 1989). Late-stage,selective removal of La and HREE was attributed to passageof a late, highly-oxidizing post-magmatic fluid.

Direct determination of fluid compositions involved incarbonatite/fluid interaction and country rock/fluid interac-tion have relied primarily on fluid inclusions in mineralssuch as apatite and fluorite. The complexities of fluid-mineral interaction are well shown by fluid inclusionstudies of apatite from the Jacupiranga carbonatite, Brazil(Costanzo et al. 2006), where magmatic evolution isinterpreted to have involved crystal fractionation andsettling of a carbonatite mineral assemblage in a fluid-stratified magma chamber. Studies of fluid inclusions inapatite from the Fen carbonatite, Norway, identified amagmatic fluid that was rich in CO2 and NaCl, whichevolved during magmatic differentiation to a CO2-free,water-dominated system that contained higher salinities anddensities as solidification progressed (Andersen 1986). Inaddition, marked changes in HF fugacity indicated thepresence of two or more independent or semi-independentlines of magmatic descent (Andersen and Austrheim 1991).Changes in fluid composition upon late-stage mixing offenitizing fluids and low-salinity meteoric waters weredocumented by Williams-Jones and Palmer (2002) based onfluid inclusions in the fenites from the Amba Dongarcomplex, India. Increases in fO2 during fenitization have

been inferred from S isotope compositions of sulfides fromferrocarbonatite from Swartbooisdrift, Namibia (Drueppelet al. 2006).

Pertinent to this survey of Fe isotope compositions ofcarbonatites, fluid inclusion studies have also demonstratedthe importance of Fe-rich chlorides in orthomagmaticcarbonatite fluids (Bühn et al. 2002; Rankin 2005). Thehigh Fe3+/Fe2+ ratios measured in many fenites, as well ascommon occurrence of disseminated iron oxides in fenites,indicates that Fe3+-bearing fluids may be expelled bycarbonatite intrusions during crystallization and solidification(e.g. Le Bas 2008). The presence of hematite, especially inthe low-grade fenites at Alnö, indicates equilibration with ahighly oxidizing fluid (Morogan 1989). Bühn and Rankins’(1999) fluid inclusion study of a fenite associated with theKalkfeld carbonatite complex, Namibia, showed that virtu-ally all alkali metals and Cl, and a major proportion F, Th, U,and Ti, were preferentially partitioned into the fluid. Inaddition, the fenitizing fluid at Kalkfeld was estimated tohave contained 3.0 to 4.1 wt. % total Fe.

Iron isotope geochemistry

Iron isotope fractionations among aqueous species andminerals are controlled by redox state and bonding environ-ments (e.g. Beard and Johnson 2004a; Schauble 2004).Geologic substances that contain Fe3+ tend to have higher56Fe/54Fe ratios relative to Fe2+ substances, with theimportant exception of pyrite, where Fe is covalentlybonded (Polyakov and Mineev 2000). Based on experi-mentally determined or predicted Fe isotope fractionationfactors, the relative order of increasing 56Fe/54Fe ratios isCa-Mg-Fe carbonate, Fe carbonate, Fe2+ silicate, aqueousFe2+, magnetite, aqueous Fe3+, hematite, and pyrite, wherethe relative order for aqueous species may be significantlychanged by chloride (Polyakov and Mineev 2000; Welch etal. 2003; Wiesli et al. 2004; Anbar et al. 2005; Polyakov etal. 2007; Hill and Schauble 2008; Shahar et al. 2008).

Although Fe isotope fractionations at igneous or mantletemperatures are relatively small, mantle-derived ultramaficrocks and minerals have a measureable range in Fe isotopecompositions, and their average 56Fe/54Fe ratio is about 0.1per mil (‰) lower than the average of basaltic rocks,possibly reflecting the effects of melting and/or metasoma-tism (Beard and Johnson 2004b; Poitrasson et al. 2004;Williams et al. 2004; Weyer et al. 2005; Williams et al.2005; Weyer and Ionov 2007). Within typical analyticaluncertainties of ±0.05 to 0.1‰ for 56Fe/54Fe ratios, allbasalts appear to be isotopically homogenous. Evolvedigneous rocks, however, may, but not always, haverelatively elevated 56Fe/54Fe ratios that are ~0.1 to 0.3‰higher than primitive basaltic rocks, and these isotopic

Iron isotopes in carbonatites 93

compositions have been interpreted to reflect magmaticdifferentiation processes, including crystal fractionation,assimilation, and/or fluid exsolution (Poitrasson and Freydier2005; Schoenberg and Von Blanckenburg 2006; Heimann etal. 2008; Teng et al. 2008).

Samples

We have analyzed carbonatites and their minerals fromeleven different countries covering three continents. Most

of the 35 samples studied are from Africa. Table 1summarizes the seventeen complexes studied. Ages rangefrom the Archean to present, and most of the carbonatitesare sövites (coarse-grained calcite carbonatite). The major-ity of the complexes are associated with well-developedzones of fenitization. Almost all of the samples are eitherplutonic or hypabyssal, other than those from OldoinyoLengai; these latter samples are chemically dissimilar to theother carbonatites analyzed in that the Oldoinyo Lengaisamples are natrocarbonatites, which have combined alkalisof 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,ijolite, pyroxenite, hornblendite, nepheline syenite.

Yes (K-rich)

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,phonolite, nephelinite melteigite, ijolite, pyroxenite, hornblendite,nepheline syenite.

Yes (K-rich)

Homa Bay, Kenya 12−1.3 Sövite, alvikite and ferrocarbonatite. Silicate rocks: ijolite,phonolite, nephelinite.

Yes

Oldoinyo Lengai, Tanzania Still active Natrocarbonatite flows and tuffs. Silicate rocks: phonolitic andnephelinitic tuffs and rare lavas including rare combeite- andmelilite-bearing types. Blocks of pyroxenite, ijolite series rocks,nepheline syenite.

Yes(metasomatizedblocks)

Panda Hill, Tanzania 113±6 Mainly sövite but areas of ferrocarbonaite and dikes ofdolomite carbonatite.

Yes

Sengeri Hill, Tanzania Same as Panda? Dolomitic dikes. Yes

Dicker Willem, Namibia 49±1 Sövite and alvikite. Silicate rocks: Ijolite xenoliths incarbonatite, trachyte dikes.

Yes

NORTH AMERICA

Borden, Canada 1872±13 Sövite, silicocarbonatite, beforsite dikes. Yes

Oka, Canada 110 Sövite and dolomite carbonatite. Silicate rocks: Okaite-jacupirangite(melilitolite-pyroxenite), and ijolite series rocks; lamprophyre dikes.

Yes

St. Honoré, Canada ca. 650 Sövite, dolomite and ankerite carbonatite. Silicate rocks: Syenite,nepheline syenite, and ijolite.

Yes

Magnet Cove, USA ca 100 Sövite. Silicate rocks: Pyroxenite, gabbro, jacupirangite, melteigite,ijolite, syenite, trachyte, phonolite.

Yes

SOUTH AMERICA

Jacupiranga, Brazil 130 Sövites, with a zone of dolomitization, and dikes of alvikite andbeforsite. Silicate rocks: peridotite, pyroxenite, jacupirangite, ijolite,nepheline syenite, essexite, tinguaite, monchiquite.

Yes

EUROPE

Kaiserstuhl, Germany 16.0 Sövite, alvikites. Silicate rocks: limburgite, phonolite and leucitetephrite, shonkinite, bergalite, nephelinite, phonolite, monchiquite,essexite and theralite.

Unknown

Kovdor, Russia 365±8 Sövite and dolomite carbonatite. Silicate rocks: Olivinite,pyroxenite, nepheline pyroxenite, melteigite, ijolite, melilite-and monticellite-bearing rocks, nepheline syenite.

Yes

Siilinjarvi, Finland 2617±10 Sövite, silicocarbonatite. Silicate rocks: Glimmerite,syenite, lamprophyre.

Yes

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 havebeen used to assess some of the ages. Included among these are those for Kovdor and Siilinjarvi (Rhuklov and Bell, this volume)

94 C.M. Johnson et al.

to analyses of whole-rock samples, Fe isotope measurementswere made of calcite, dolomite, magnetite, spinel, olivine,pyroxene, biotite, phlogopite, melilite, nepheline, perovskite,pyrite, monticellite, and actinolite. Because carbonatites areknown to contain a great variety of minerals, this studypermitted the first Fe isotope analyses of many of theseminerals from samples that formed at igneous temperatures.

Analytical methods and nomenclature

Mineral separates were obtained by hand picking, withspecial care given to the carbonate fraction of carbonatitesamples. The very low Fe contents of carbonate minerals incarbonatites required mineral separate purity in excess of99%. Mineral separates were washed in doubly distilledwater (2X H2O) prior to dissolution. Iron isotope analysisof carbonatites presents special challenges due to very lowratios of Fe to alkali or alkali-earth elements, as well as thegreat sensitivity of low-Fe carbonate minerals to smallamounts of contamination by Fe-rich silicate or oxideminerals, and the new methods we developed to addressthese issues are described below.

Partial dissolution experiments and application to carbonateiron isotope analyses

Despite a mineral separate purity of >99%, the very low Fecontents of the carbonate mineral separates raises the

possibility that their Fe isotope compositions may havebeen modified by small amounts of oxide and/or silicatemineral contamination. Total dissolution of several hand-picked, >99% pure calcite mineral separates produced Fecontents that significantly exceeded those determined byelectron microprobe analysis by previous studies of thesame samples (Haynes et al. 2003), suggesting the presenceof small amounts of oxide or silicate contamination thatwas not visible under a binocular microscope. We thereforedeveloped a partial dissolution protocol that completelydissolved carbonate but did not significantly dissolve oxideor silicate that may have existed in the carbonate mineralseparates (Table 2). Partial dissolution experiments ofmagnetite and silicate (clinopyroxene) minerals involvedhigh (85°C) and low (25°C) temperatures, two sieved grainsizes, and three different molarities of HCl. Small amountsof Fe are dissolved from magnetite using 7 M HCl at highor low temperatures, at 1.6% and 1.1% dissolution at 85°Cand 25°C, respectively. Magnetite that was sized to 1 mmand 0.1 mm in diameter responded differently to partialdissolution, where hot 7 M HCl dissolved 1.6% and 1.2%,respectively, and cold 7 M HCl dissolved 1.1% and 0.7%,respectively. At lower HCl molarity the differences in theamount dissolved between grain sizes and temperature wasless pronounced. The lowest amount dissolved (0.01%) wasaccomplished using 1 mmmagnetite crystals in cold 1MHCl.

Parallel experiments were performed on clinopyroxene.The 85°C 7 M HCl treatment dissolved 0.06% of the Fefrom 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 electronmicroprobe analysis, are from Haynes et al. (2003). Dissolved Fe determined by Ferrozine assay

Iron isotopes in carbonatites 95

0.5 M HCl treatment dissolved 0.045% of the Fe from a0.5 mm clinopyroxene grain. Although the clinopyroxenehad a larger amount of Fe partially dissolved at 25°C using0.5 M HCl as compared to magnetite, the amounts of Fedissolved in both minerals would be too small to affect theFe isotope composition of the carbonate. Based on theseresults, dissolution of a 99% pure carbonate mineralseparate that contained 1% magnetite or silicate in cold0.5 M HCl should produce less than 1% Fe contaminationfrom magnetite or silicate in the dissolved Fe component,which would have no effect on the measured Fe isotopecomposition of the carbonate.

It is important to note that we did not use acetic acid(HAc) for carbonate dissolution because our previous testsshowed that magnetite may undergo incongruent dissolu-tion using HAc (Valaas-Hyslop et al. 2008), which mayproduce spurious Fe isotope compositions if HAc is used toselectively dissolve carbonate from a mixture of carbonateand magnetite. In addition to producing anomalous δ56Fevalues during partial dissolution of magnetite using HAc,Valaas-Hyslop et al. (2008) noted that Fe(II)/FeTotal ratioschanged in the solutions relative to those of magnetite,reflecting redox changes in solution through interactionwith acetate, and likely formation of a new surface phase.These results were challenged by Von Blanckenburg et al.(2008), who observed no anomalous Fe isotope effectswhen comparing dissolution of natural whole-rock carbo-nates using HCl and HAc, although we note that their testsdid not involve minerals of known isotopic compositions, nordid they measure Fe(II)/FeTotal ratios of solution produced bypartial dissolution of magnetite. We contend that proton-promoted carbonate dissolution without the presence oforganic ligands such as acetate remains the safest approachfor dissolving carbonate components in the presence ofminor mineral contaminants for Fe isotope analysis.

Sample processing and mass analysis

Carbonate mineral separates were digested in 0.5 M HCl,where the sample was left at room temperature for ~1 h,followed by ~20 min in an ultrasonic bath, then centri-fuged. Visual inspection showed that no residue waspresent after centrifugation. This approach was used basedon the partial dissolution experiments noted above. Silicateminerals were digested in ~5 ml of 29 M HF and ~500μl of7 M HNO3 in Teflon containers on hotplates for 24 h. Thesolution was then dried down and brought up in 5 ml of8 M HCl. The solutions were heated on hotplates in Tefloncontainers for at least 12 h, and once more dried down.Magnetite mineral separates were digested in ~5 ml of 8 MHCl in closed Teflon containers on hotplates for 24 h, followedby centrifugation to remove any residual silicate minerals orother oxides that are more refractory, such as rutile.

Whole-rock powders of carbonatites required specialdissolution procedures due to their very high Ca contents,which would produce extensive formation of Ca fluoridesusing traditional HF dissolution methods. Magnetite, ifpresent, was removed from the samples using a hand magnetand set aside. The remaining powder was dissolved in 0.5 MHCl for ~1 h at room temperature, followed by ~20 min in anultrasonic bath to remove the calcite fraction. Followingcentrifugation, the dissolved calcite fraction was set aside, andthe remaining “silicate” fraction was washed in 2X H2O,followed by digestion in ~5 ml of 29 M HF and ~500μl of7 M HNO3 in closed Teflon containers on hotplates for 24 h.The HF-HNO3 solution was then dried down and brought upin 5 ml of 8 M HCl. The solution was heated on hotplates inclosed Teflon containers for at least 12 h and once moredried down. The magnetite fraction was digested as notedabove, then brought up in 0.5 M HCl. The calcite fractionand the magnetite fraction were then recombined with the“silicate” fraction to provide a whole-rock analysis.

Iron was separated by anion-exchange chromatography.Samples were passed through ion-exchange columns 2–4times, where the greatest number of passes were requiredfor high Ca/Fe ratio samples; previous work has shown thatmultiple passes do not fractionate Fe isotopes compositionsbecause yields are ~100% (Beard et al. 2003). Isotopiccompositions were determined on a Micromass IsoProbe, asingle focusing, multi-collector inductively-coupled-plasmamass spectrometer with a magnetic sector mass analyzer,equipped with a Cetac Aridus desolvating micro-concentricnebulizer and Elemental Scientific spray chamber. Samplesolutions contained 300 ppb Fe, were aspirated at 50μL/minfor eight minutes, resulting in consumption of 120 ng of Feper analysis. Instrumental mass bias and drift was correctedusing a standard-sample-standard approach; details can befound in Beard et al. (2003) and Albarède and Beard (2004).

The accuracy and precision of the Fe isotope measure-ments were assessed using analyses of standards, as well asmultiple analysis of samples. Separated Fe solutions for 35samples were analyze 2 or 3 times under different runningconditions, and the average difference in the analyses is0.04‰ in 56Fe/54Fe ratios. Five samples were dissolved twoor three times, and the average difference between themultiple dissolutions, which included separate processingthrough ion-exchange columns, was 0.06‰ in 56Fe/54Feratios. These assessments of precision and accuracy matchthose obtained through analysis of ultra-pure Fe standards,and we estimate the average 2σ reproducibility (~2SD) tobe 0.06‰ for 56Fe/54Fe ratios.

Nomenclature

Iron has four naturally occurring stable isotopes, 54, 56, 57,and 58, and isotopic compositions have been generally

96 C.M. Johnson et al.

reported using the three major isotopes, either as 56Fe/54Fe or57Fe/54Fe ratios. Data are reported here in standard δnotation, where:

d56Fe ¼ 56Fe�54Fe

� �sample

.56Fe

�54Fe

� �standard

�1h i

� 103

ð1Þin units of per mil (‰), where (56Fe/54Fe)standard is taken asthe average of terrestrial igneous rocks (Beard et al. 2003).Inter-laboratory comparison of Fe isotope ratios can be madeby comparison to the measured iron isotope composition ofthe certified reference material IRMM-014, which has aδ56Fe value of –0.09‰ on the igneous rock scale. The δ57Fevalue may be defined in an analogous manner using the57Fe/54Fe ratio, and δ57Fe and δ56Fe values should be relatedin a mass-dependent manner.

We discuss Fe isotope fractionations between twophases, A and B, as the difference in the measured δ56Fevalues:

$56FeA�B ¼ d56FeA � d56FeB ð2ÞThis is an approximation to the isotope fractionation

factor αA-B, which is related to ∆56FeA-B through:

103ln aA�B � $ 56FeA�B ð3Þfollowing standard practice. For isotopic fractionations onthe order of 1–3‰, use of ∆56FeA-B introduces negligibleerror relative to analytical uncertainty. Finally, we comparemeasured Fe isotope fractionations with those predictedfrom theory or measured in experiment using a self-consistent set of reduced partition function ratios for56Fe/54Fe, defined as β56/54, following standard practice.To a very good approximation, these can be related to∆56FeA-B by:

$56FeA�B � 103ln "A56=54 � 103ln "B

56=54 ð4ÞIt is important to note, however, that Eqs. (2) and (4) do

not imply that δ56Fei is equal to 103ln βi56/54.

Results

The δ56Fe values for carbonate minerals (calcite anddolomite) from the carbonatites studied here, as well aswhole-rock samples, span a significantly greater range thanthat defined by other samples considered to reflect mantlecompositions (Fig. 1). Basaltic rocks have δ56Fe=0.0±0.05‰ (1SD) (e.g. Beard et al. 2003; Poitrasson et al. 2004;Weyer et al. 2005; Weyer and Ionov 2007). Mantle-derivedxenoliths and Alpine peridotites define a slightly largerrange in δ56Fe values, although most analyses lie within0.2‰ of δ56Fe=0.0 (Fig. 1). In contrast, the average δ56Fe

value for carbonatites (carbonate minerals or whole-rocks)is clearly less than the average of basaltic or ultramaficrocks (Fig. 1), and the range in δ56Fe values spans thelargest range yet measured for igneous rocks, from δ56Fe=−1.0 to +0.8‰. Remarkably, this range spans that com-monly measured in low-temperature aqueous environments(e.g. Johnson et al. 2008).

The Fe isotope compositions of different minerals in thecarbonatites studied vary greatly (Fig. 2). The low-Fecontent minerals such as calcite and dolomite have thegreatest range in δ56Fe values, where most samples havenegative δ56Fe values, but some are positive. The δ56Fevalues for silicate minerals and magnetite cluster aboutδ56Fe=0.0‰, although there is considerable spread relativeto basaltic magmas and silicate minerals from ultramaficrocks (Figs. 1 and 2). Based on temperatures calculatedusing magnetite-calcite O isotope thermometry for some ofthe complexes we have investigated (Haynes et al. 2003),there are no clear correlations between crystallizationtemperature and δ56Fe values for carbonate, silicates, ormagnetite. Moreover, the range in δ56Fe values for carbo-nates, silicates, and magnetite is larger than the rangeexpected to be in equilibrium with the mantle, based on theFe isotope fractionations predicted from theory or measured

Fig. 1 Histograms of δ56Fe values for (a) carbonate minerals andwhole-rocks from this study, (b) whole-rock samples of ultramaficrocks, and (c) minerals from ultramafic rocks. Data for (b) and (c)from Zhu et al. (2002), Beard and Johnson (2004b), Poitrasson et al.(2004), Williams et al. (2004), (2005), Weyer et al. (2005), Shultis(2006), and Weyer and Ionov (2007)

Iron isotopes in carbonatites 97

in experiment (Fig. 2). Although the relative order of δ56Fevalues for carbonates, silicates, and magnetite that isexpected based on isotopic fractionation factors is generallyreflected in the average δ56Fe values measured, whereδ56FeCarbonate < δ56FeSilicate < δ56FeMagnetite, the range in themeasured isotopic compositions is considerably larger anddoes not correlate with crystallization temperatures (Fig. 2).

Discussion

Below we touch on several results of our initial Fe isotopesurvey of carbonatites. First, Fe isotope disequilibriumamong minerals is evaluated relative to predicted orexperimentally determined Fe isotope fractionation factors.Second, we compare the measured Fe isotope compositionswith Li, C, O, and Sr isotope compositions determined onthe same samples. Third, the Fe isotope evidence for ex-pulsion of Fe3+-bearing fluids is discussed. Fourth, we bringthe Fe isotope variations determined in this study into amodel for carbonatite genesis and evolution that is consistentwith current petrogenetic models for carbonatites.

Iron isotope fractionations at igneous temperatures

The wide range in measured Fe isotope fractionationsamong the minerals in carbonatites illustrated in Fig. 2 isexplored on a sample-by-sample basis in Fig. 3. Data forminerals are cast relative to magnetite in traditional “δ-δ”plots, and isotopic fractionation lines are shown formagnetite-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 twominerals wereanalyzed, ten samples have Fe isotope data for two or moremineral pairs that could be used to evaluate internal isotopicequilibrium. The greatest confidence in evaluating equilibri-um is placed in the magnetite-olivine Fe isotope fractionationfactors, where experimental determinations using the “three-isotope-method” (Shahar et al. 2008) have confirmedpredicted fractionations based on theory (Polyakov andMineev 2000; Polyakov et al. 2007). Based on Fe isotopeanalyses of minerals from a wide variety of silicate-dominated igneous rocks, there does not appear to be asignificant Fe isotope fractionation among the varioussilicate minerals, at least in the case of minerals where Feis largely Fe2+ (Polyakov and Mineev 2000; Beard andJohnson 2004b; Polyakov et al. 2007; Heimann et al. 2008),and so the magnetite-olivine fractionation curve of Shahar etal. (2008) is taken to be applicable to all fractionationsbetween magnetite and Fe2+ silicate minerals.

Of the ten samples where two or more mineral pairswere analyzed, only three samples have Fe isotopefractionations that indicate magnetite-silicate Fe isotopeequilibrium (samples P2-670, STH-08, and P10-208), andthese are plotted together in Fig. 3a. Six samples that havebeen measured for magnetite-carbonate and magnetite-silicate fractionations do not record equilibrium magnetite-silicate fractionations (samples 5–15, BO-203, P2-680,P10-202, KO-101, and SIL 106), and this group is plottedin Fig. 3b. An additional sample is plotted in this group,MC-1, where spinel, calcite, mica, and montecellite wereanalyzed, based on the fact that spinel-silicate fractionations

Fig. 2 δ56Fe-crystallization temperature variations for carbonatitesfrom this study (Table 3). Crystallization temperatures from O isotopethermometry (Haynes et al. 2003), except those that have error bars;these samples are arbitrarily plotted at a temperature of 600±100°C.Shaded curved fields show predicted δ56Fe values as a function oftemperature, using the β56/54 factors from Table 4. Fields defined bytwo end member Fe isotope compositions for magnetite or silicateδ56Fe=0, which spans the major Fe repositories in the samplesanalyzed that may have been in equilibrium with mantle peridotite andhence control the Fe isotope composition of the samples and constituentminerals. Cc=calcite, Ank-ankerite, Fa=fayalite, Mt=magnetite

98 C.M. Johnson et al.

are predicted to be significantly smaller than magnetite-silicate fractionations (Polyakov and Mineev 2000; Polyakovet al. 2007), a conclusion supported by comparing spinel-olivine fractionations measured in ultramafic rocks (Williamset al. 2004; Williams et al. 2005) with the magnetite-olivine

fractionations experimentally measured by Shahar et al.(2008). These considerations suggest that sample MC-1 isnot in Fe isotope equilibrium. Four samples analyzed formagnetite-calcite or magnetite-dolomite fractionations werenot analyzed for magnetite-silicate fractionations (samplesBD-1499, BD-1485, BD-724, and Phalaborwa), and hencedo not have an independent check on Fe isotope equilibrium;this group is plotted in Fig. 3c.

Several samples that are not in magnetite-silicate Feisotope equilibrium (Fig. 3b) appear to have equilibriumsilicate-silicate or magnetite-pyrite fractionations. Sample5–15 has similar magnetite-pyroxene and magnetite-nepheline fractionations, suggesting that pyroxene andnepheline are in Fe isotope equilibrium. Samples P2-680and BO-203 have similar magnetite-pyroxene andmagnetite-mica Fe isotope fractionations, indicating thatthe two silicates are in isotopic equilibrium. Sample MC-1has similar spinel-mica and spinel-olivine fractionations,suggesting that although spinel is not in Fe isotopeequilibrium with the silicates, the two analyzed silicatesare 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 Feisotope equilibrium in sample SIL 106 (Fig. 3c).

The Fe isotope fractionations among minerals are castrelative 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 andminimizes the effects of sub-solidus equilibration. It isimportant to note, however, that for some of the samplesstudied by Haynes et al. (2003), calcite-biotite O isotopetemperatures were up to 270°C lower than magnetite-calcitetemperatures, indicating sub-solidus cooling and isotopicexchange did occur. If independent crystallization temper-atures were not available, a temperature of 600±100°C wasassumed in Fig. 4.

When viewed relative to calculated or estimated crystal-lization temperatures, the three samples that have equilib-rium magnetite-silicate fractionations have relatively highmagnetite-calcite fractionations (Fig. 4a), exceeding pre-dicted magnetite-siderite or magnetite-ankerite fractiona-tions (Table 4). We therefore propose an empiricalmagnetite-calcite Fe isotope fractionation curve using theβ56/54 values in Table 4, which has not been measured inexperiments nor explicitly calculated from theory, based ona) the observation that the magnetite-calcite Fe isotopefractionations for samples that are in magnetite-silicateequilibrium are generally higher than the predictedmagnetite-ankerite or magnetite-siderite fractionations, and

Fig. 3 Variations in δ56Fe values for magnetite (Mt) relative to otherco-existing minerals (Cc=calcite, Dol=dolomite, Px=pyroxene,Amph=amphibole, Neph=nepheline, Py=pyrite, Ank-ankerite, Ol=oli-vine, Sid=siderite). Minerals from same sample connected by tie lines.Samples grouped relative to those that have equilibrium magnetite-silicate Fe isotope fractionations (a), those that do not haveequilibrium magnetite-silicate fractionations (b), and those that haveno independent control on Fe isotope equilibrium (c). Note in (b) thatone sample analyzed is spinel (sp), not magnetite. Magnetite-olivineand magnetite-pyrite fractionation lines calculated for a temperature of600°C, using the β56/54 factors in Table 4. Isotopic data from Table 3

Iron isotopes in carbonatites 99

b) the predicted and experimentally measured trend ofdecreasing β56/54 in carbonates with decreasing Fe content(Polyakov and Mineev 2000; Johnson et al. 2005), whichwould produce increasing magnetite-carbonate fractiona-tions with decreasing carbonate Fe content. When theisotopic data are viewed in the context of our proposedmagnetite-calcite fractionation curve, all but one of thesamples (BO-23) that do not have equilibrium magnetite-silicate fractionations also have anomalously low magnetite-calcite fractionations (Fig. 4b) Two of the four samples thatdo 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 inFig. 4c), whereas two samples do not (BD-1499 andPhalaborwa in Fig. 4c). Finally, we note that the twopyrite-bearing samples, which crystallized at markedlydifferent temperatures, have magnetite-pyrite fractionationsthat generally lie along those predicted by Polyakov andMineev (2000) and Polyakov et al. (2007), despite theirdisequilibrium 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). Calciumcarbonate magma in equilibrium with silicate mantle shouldhave a δ56Fe value of ~ −0.3 to −0.8‰, assuming an averagesilicate mantle δ56Fe value of 0.0 to −0.1‰ (Beard andJohnson 2004b; Williams et al. 2004; Weyer et al. 2005;Williams et al. 2005; Schoenberg and Von Blanckenburg2006; 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 anyother mineral (Table 4) indicates that crystal fractionation ofnon-carbonate, Fe-bearing minerals (silicates, oxides, sul-fides) will move a carbonatite magma toward more negativeδ56Fe values, lower than the −0.3 to −0.8‰ values that areexpected for carbonate magmas in equilibrium with thesilicate mantle. In Fig. 5, we compare the Fe, Li, O, and Srisotope compositions of the carbonatites studied here withthose expected for primary carbonate melts, as well aschanges that may occur through magmatic differentiationand fluid interaction.

Based on Sr, Nd, and Pb isotopes in East Africancarbonatites, Bell and Tilton (2001) interpreted the isotopiccompositions to largely reflect mixing between HIMU andEM I mantle components. Our sample suite includes thecarbonatite complexes that lie closest to these end members(EM I: Homa Bay; HIMU: Sukulu), and data for the othercomplexes analyzed here scatter between these mantlecomponents (Fig. 5a). Similar trends are observed for Nd

Fig. 4 Variations in Fe isotope fractionations between magnetite andco-existing minerals relative to temperature (106/T2). Symbols andabbreviations as in Fig. 3. Sample grouping also follows that of Fig. 3:samples that have equilibrium magnetite-silicate Fe isotope fractiona-tions (a), those that do not have equilibrium magnetite-silicatefractionations (b), and those that have no independent control on Feisotope equilibrium (c). Note in (b) that one sample analyzed is spinel(sp), not magnetite. Temperatures based on O isotope thermometrydetermined on the same samples. For samples that do not haveindependent temperature constraints, temperatures were arbitrarilyplotted as 600±100°C. Isotopic data from Table 3. Temperature-dependent isotope fractionation curves calculated using β56/54 factorsfrom Table 4

100 C.M. Johnson et al.

and Pb isotope variations (not shown). For the carbonatitesthat have δ56Fe values that reflect equilibrium with themantle (δ56Fe < −0.3‰), there is no correlation between Feisotope compositions and Sr, Nd, or Pb isotope ratios,suggesting that Fe isotopes are not a sensitive indictor ofEM I or HIMU reservoirs. This conclusion is consistent withthe observation that δ56Fe values for basaltic rocks areessentially homogeneous regardless of their radiogenic

isotope compositions (e.g. Beard et al. 2003; Weyer andIonov 2007).

A larger suite of the samples analyzed here has beenanalyzed 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 δ18Ovalues for the samples analyzed for Fe isotopes fall within arestricted range that is consistent with derivation from themantle. Sample DU-365 from Toror has an elevated δ18Ovalue (Fig. 5b), suggesting fluid loss, which will tend toincrease δ18O values (e.g. Deines 1989), although it is notclear that fluid loss would produce a sufficient increase inδ18O values. The fact that this sample has a relatively highδ56Fe value is consistent with passage of Fe3+-bearingfluids, as will be discussed below. The δ13C values of thesamples analyzed for Fe isotopes vary greatly (Fig. 5c),scattering beyond the range that reflects equilibrium withthe mantle. Relatively high δ13C values in carbonatites aregenerally interpreted to reflect extensive crystal fraction-ation, whereas loss of CO2 should decrease δ13C values(e.g. Deines 1989). Extensive crystal fractionation, however,should also produce an increase in δ18O values (Deines1989), which is not generally observed in the samplesstudied. An alternative explanation for an increase in δ13Cvalues but not δ18O values, is passage of CO2-bearingfluids that produced a net addition of high-δ13C carbon,rather than CO2 loss. As will be discussed below, samplesthat have the highest δ56Fe values are interpreted to reflectpassage of Fe-bearing solutions, and if such solutions wereCO2 bearing, this may be the explanation for the tendencyof high δ56Fe values to be associated with high δ13C values.

Eight samples analyzed by Halama et al. (2008) for Liisotopes were analyzed in this study, and the data suggest ageneral negative correlation between δ7Li and δ56Fe values(Fig. 5d), although the data base is somewhat limited.Halama et al. (2008) interpreted decreasing δ7Li values,relative to mantle values, to reflect expulsion of Li-bearingfluids upon extensive crystal fractionation, and they noted abroad negative correlation between δ7Li and δ13C valuesthat supports this interpretation. Because fluids rich in Liand Fe would most likely be chloride bearing, it isanticipated that further studies of carbonatite complexesthat experienced fluid loss would have strong correlationsbetween Li and Fe isotopes, as well as fluid inclusionevidence for chloride-bearing brines.

Evidence for interactions with Fe3+-bearing fluids

Calciocarbonatite magmas in equilibrium with the mantlemust have negative δ56Fe values, most likely less than

Fig. 5 Variations in Sr, O, C, and Li isotope compositions, asavailable, for samples analyzed in this study for Fe isotopecompositions. For most samples, whole-rocks were analyzed, but forsome O and C isotope compositions, calcite was analyzed, and thesewere plotted against δ56Fe values determined on the same calcitemineral separate. Dark gray boxes denote fields for primary carbonatemagma compositions. Arrows denote various processes that maychange the isotopic compositions, as discussed in the text. Sr, O, C,and Li isotope data from Bell and Tilton (2001), Haynes et al. (2003),and Halama et al. (2008). Fe isotope data from Table 3

Iron isotopes in carbonatites 101

−0.3‰, posing a significant interpretive problem for thelarge number of carbonatites and calcites that have δ56Fe >−0.3‰. Crystal fractionation of silicates, oxides, andsulfides from carbonate magmas cannot explain δ56Fe >−0.3‰, as noted above. Although the modal abundances ofsilicates, oxides, and sulfides are significantly less thanthose of carbonate minerals in the samples studied, themajority of Fe in all of the samples lies in the non-carbonate minerals, confirming that crystal fractionation, byitself, should strongly decrease δ56Fe values in thecarbonatite complexes studied here. Similarly, generationof an immiscible silicate-carbonate melt (e.g. Kjarsgaardand Hamilton 1989; Brooker 1998; Kjarsgaard 1998; Leeand Wyllie 1998) should produce negative δ56Fe values inthe carbonate melt component, particularly if peralkalinesilicate magmas were involved; the high alkali contents insuch magmas produce high Fe3+/Fe2+ ratios in the melt dueto M2+-Fe2+ and M+-Fe3+ charge balance, without signifi-cant changes in fO2, and Fe3+-bearing silicates have δ56Fevalues ~0.2 to 0.4‰ higher than Fe2+-bearing silicates atigneous temperatures, based on calculated Fe isotopefractionation factors (Polyakov and Mineev 2000), as wellas experiments involving Fe3+-rich silicate melts (Schuessleret al. 2007). The general effect of increasing δ56Fe valueswith increasing Fe3+/Fe2+ ratios in silicate minerals has beenobserved in igneous complexes (Schoenberg et al. 2009).

The strong negative correlation between magnetite-calcite Fe isotope fractionations and the δ56Fe value ofcalcite (Fig. 6a) indicates that the non-equilibriummagnetite-calcite fractionations illustrated in Figs. 3 and 4largely reflect anomalously high δ56Fe values for calcite. Incontrast, δ56Fe values for magnetite are relatively constant

over the wide range of magnetite-calcite Fe isotopefractionations, although there is a weak tendency for the mostdisequilibrium magnetite-calcite fractionations to be associat-ed with the lowest δ56Fe values for magnetite (Fig. 6b). Wepropose that passage of Fe-bearing fluids through carbonatemagmas or crystallized minerals is the most likely explana-tion for the anomalously high δ56Fe values for calcite. Thevery low Fe contents of calcite makes this mineralparticularly sensitive to modification by Fe-bearing fluids,and hence can explain the strong correlation between theδ56Fe values for calcite and the magnetite-calcite fractiona-tions (Fig. 6). Based on calculated β56/54 factors for Fe- andFe-bearing fluids (Table 4), however, the δ56 Fe values ofcalcite should decrease upon loss of a Fe-bearing fluid.Although Heimann et al. (2008) documented evidence forincreasing the δ56 Fe values of evolved, metaluminoussilicate magmas through loss of an Fe-bearing fluid, thismodel cannot explain the increase in δ56Fe values for calcitein carbonatites because, under equilibrium conditions, theδ56Fe values of calcite are lower than those of Fe2+- orFe3+-bearing fluids, silicates, or oxides (Table 4).

A simple fluid-rock mixing model (e.g. Criss 1999) mayexplain the trend of decreasing magnetite-carbonate fractio-nations with increasing δ56Fe values for calcite (Fig. 6),reflecting net addition of Fe that had high-δ56Fe values,rather than fluid loss. A mixing model that incorporates arange of Fe contents in calcite and fluid, as well as isotopiccompositions (Table 5), can produce the spread in δ56Fevalues for calcite. Note that the much higher Fe contents ofmagnetite, relative to calcite, results in no significantchange in the δ56Fe values of magnetite in the fluid-rockmixing model (Fig. 6b). A fluid-rock mixing model that

Fig. 6 Variation in magnetite-calcite Fe isotope fractionations(Δ56FeMt-Cc) with δ56Fe values for calcite (a) and magnetite (b). Graybox indicates Δ56FeMt-Cc fractionations and δ56FeCc and δ56FeMt valuesthat would be in equilibrium with the mantle at igneous temperatures.Equilibration temperatures for calcite-magnetite fractionation shown asdashed lines, as calculated using β56/54 factors from Table 4. Note thatone sample, spinel (sp) was analyzed instead of magnetite. Symbolsindicate groupings of data according to magnetite-silicate Fe isotope

equilibrium (samples in Figs. 3a and 4a), magnetite-silicate disequilib-rium (samples in Figs. 3b and 4b), or no independent control on Feisotope equilibrium (samples in Figs. 3c and 4c). The strong negativecorrelation indicates that the Δ56FeMt-Cc fractionations are controlled bythe δ56FeCc values. Three models for fluid-rock interaction shown,where arrows indicate direction of increasing fluid/rock ratios from zero(in gray box) to 10 (end of arrow) for Models A, B, and C of Table 5

102 C.M. Johnson et al.

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±0.04SU 103 CC A-1 2996 −0.24±0.03 −0.27±0.05

A-2 −0.18±0.03 −0.37±0.07WR −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±0.06WR −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±0.03Dol A-1 −0.85±0.07 −1.25±0.03

A-2 −0.96±0.04 −1.49±0.05TO-100B WR −0.37±0.03 −0.58±0.05Toror, Uganda

DU-365 WR A-1 −0.04±0.06 −0.08±0.03A-2 −0.06±0.03 −0.03±0.05

Homa Bay, Kenya

Homa Bay CC A-1 4385 −0.72±0.05 −1.07±0.03A-2 −0.70±0.04 −1.04±0.05

WR −0.45±0.04 −0.57±0.06Oldoinyo Lengai, Tanzania

OL-2 1993 Mt −0.18±0.05 −0.21±0.04OL-10 1993 Mt −0.20±0.06 −0.22±0.04BD-114 WR A-1 −0.49±0.04 −0.78±0.04

A-2 −0.42±0.04 −0.71±0.07BD-118 WR −0.62±0.02 −0.99±0.03Panda 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±0.06A-2 −0.26±0.04 −0.28±0.09

Tan-213 WR −0.30±0.05 −0.42±0.06BD-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±0.03Sengeri 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±0.04A-2 −0.15±0.08 −0.24±0.03A-3 −0.10±0.05 −0.19±0.07

Iron isotopes in carbonatites 103

Table 3 (continued)

Sample Mineral Dissol. ppm Fe δ56Fe δ57Fe

CC A-1 12348 −0.36±0.08 −0.57±0.03A-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±0.03CC A-1 2405 −0.15±0.04 −0.25±0.03

A-2 −0.15±0.03 −0.20±0.03NORTH AMERICA

Borden, Canada

BO-203 Mt A-1 −0.12±0.07 −0.16±0.04A-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±0.08A-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±0.04CC A 596 −0.73±0.04 −1.13±0.04

B −0.77±0.08 −1.17±0.04Mica −0.30±0.04 −0.57±0.05cpx −0.28±0.05 −0.31±0.03

P2–680 Mt −0.19±0.04 −0.32±0.07CC 176 −0.34±0.02 −0.53±0.03Mica −0.05±0.04 0.01±0.04

cpx A-1 −0.18±0.03 −0.24±0.04A-2 −0.10±0.04 −0.20±0.05

OC 203 mica −0.09±0.06 −0.16±0.06melilite A-1 0.06±0.04 0.01±0.04

A-2 −0.02±0.04 −0.04±0.04OC 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±0.04A-2 −0.54±0.02 −0.81±0.04

OC 307 Bio −0.11±0.03 −0.13±0.04OC 310 CC −0.34±0.05 −0.45±0.04OC 318 px −0.10±0.04 −0.16±0.05OC 320 CC −0.57±0.10 −0.85±0.06

Mica A-1 −0.07±0.04 −0.11±0.05A-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±0.03A-2 −0.72±0.05 −1.08±0.03

Mica A-1 −0.12±0.03 −0.25±0.04A-2 −0.05±0.03 −0.01±0.03

104 C.M. Johnson et al.

Table 3 (continued)

Sample Mineral Dissol. ppm Fe δ56Fe δ57Fe

WR −0.16±0.04 −0.16±0.04Magnet Cove, USA

MC-1 Spinel −0.02±0.06 −0.02±0.03CC A 100 −0.87±0.04 −1.09±0.03

B −1.12±0.03 −1.46±0.05Mica −0.48±0.07 −0.59±0.04Mont A-1 −0.37±0.05 −0.48±0.06

A-2 −0.38±0.03 −0.55±0.03SOUTH AMERICA

Jacupiranga, Brazil

P10–202 Mt A −0.04±0.04 0.07±0.05

B 0.01±0.03 −0.03±0.04CC A-1 169 −0.47±0.04 −0.68±0.04

A-2 −0.49±0.03 −0.60±0.03B −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±0.04A-2 −0.30±0.03 −0.38±0.04B −0.35±0.07 −0.48±0.10C −0.34±0.06 −0.52±0.04

Mica 0.06±0.05 0.13±0.05

cpx −0.07±0.06 −0.13±0.05EUROPE-ASIA

Kaiserstuhl, Germany

K-102 CC A-1 −0.65±0.05 −0.94±0.03A-2 −0.60±0.05 −0.91±0.03

Mica A-1 −0.19±0.04 −0.27±0.05A-2 −0.13±0.03 −0.22±0.03

Kovdor, Russia

KO-101 Mt A-1 −0.43±0.05 −0.60±0.03A-2 −0.37±0.03 −0.51±0.04

CC 632 −0.47±0.03 −0.65±0.03Mica 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±0.04A-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 massanalysis. Analyses noted with “-1”, “-2”, etc. indicate repeat analysis of same solution obtained during ion-exchange chromatography, but underdifferent days.

Iron isotopes in carbonatites 105

reflects a net addition of Fe through passage of Fe-bearingfluids is consistent with a weak correlation between calciteFe contents and δ56Fe values (not shown; data in Table 3).The fluid-rock mixing model that most successfullyexplains the data uses a relatively high δ56Fe value for theFe-bearing fluid (model C in Fig. 6 and Table 5), which, atigneous temperatures, would most likely be an Fe3+-bearingfluid based on β56/54 values (Table 4). In this model, anFe2+-bearing fluid cannot explain the wide range in δ56Fevalues of calcite, even at very high fluid/rock ratios, giventhe lower β56/54 factors relative to those of Fe3+-bearingfluids. Invoking an Fe3+-bearing fluid is consistent with theperalkaline nature of carbonatite-silicate systems, whereFe3+/Fe2+ ratios are relatively high as compared to metal-uminous silicate magmas, the common occurrence of Fe3+

oxides (magnetite, hematite) in carbonatites, and the abun-dance of Fe3+-oxides in the fenites that surround intrusivecarbonatite complexes.

The slight decrease in δ56Fe values for magnetite forsamples that have high δ56Fe values for calcite (Fig. 6)remains a puzzle. This weak trend cannot be explained by anet addition of a high-δ56Fe, Fe3+-bearing fluid. Crystalli-zation of magnetite from a high-δ56Fe carbonate magma, ifin isotopic equilibrium with calcite, would produce highδ56Fe values for magnetite. One possible explanation is thatmagnetite in the high-δ56 Fe calcite samples equilibratedwith an Fe3+-rich fluid at low temperatures, which couldpotentially produce a decrease in δ56Fe values for magne-tite, given the relative β56/54 factors (Table 4). In such amodel, however, calcite must maintain isotopic disequilib-rium with the fluid and magnetite. It is also possible that thecalcite and magnetite reflect, in part, physical mixtures ofphenocrysts from unrelated magmas. A deeper understand-ing of the Fe isotope exchange kinetics of carbonate, oxide,and fluid is required to test these models, as well asexperimental confirmation of the Fe isotope fractionations

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

Olivinea Magnetiteb Sideritec Ankerited Calcitee Pyritef [FeIICl4]2-g FeIIICl3

h

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 This study; calculated using magnetite β56/54 from b and 22 assuming ∆56 FeMt-CC =∆56 FeMt-Ank +∆56 FeSid-Ank, which is equivalent to assuming∆56 FeMt-Carbonate varies linearly with Fe content, and that calcite contains very minor Fef Polyakov et al. (2007)g Schauble et al. (2001)h Hill and Schauble (2008)

Table 5 Parameters for fluid-rock interaction model

Model Initial δ56FeCalcite Initial Fe concentrationcalcite (ppm)

Initial δ56FeFeCl3 Initial Fe concentrationfluid (ppm)

δ56FeMagnetite

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

Magnetite Fe concentration assumed to be stoichiometric, at 724,138 ppm. Initial δ56 FeFeCl3=+0.27 reflects predicted composition at 800°C, andinitial δ56 FeFeCl3=+0.56 reflects predicted composition at 600°C, using β56/54 factors from Table 4

106 C.M. Johnson et al.

between Fe3+- and Fe2+-bearing fluids, oxides, and carbo-nates at igneous temperatures.

Fe isotope constraints on carbonatite genesis and evolution

We bring the above discussion into a summary model forcarbonatite genesis and evolution in terms of Fe isotopevariations in Fig. 7. This model is consistent with currentpetrogenetic models for carbonatites. Generation ofcarbonate magmas in the mantle, either in a plume or thelithosphere, will produce negative δ56Fe values,certainly < −0.3‰ at mantle temperatures, but probablynot lower than −0.8‰ at relatively low temperatures, asdiscussed above. It is possible that slightly lower δ56Fevalues for carbonatite magma may be produced if meltingof a previously carbonated peridotite (e.g. Dalton andPresnall 1998; Wyllie and Lee 1998; Yaxley et al. 1998)occurred, assuming such a peridotite had slightly lowerδ56Fe values due to addition of low-δ56Fe value carbonate.If mafic silicate magmas were associated with carbonatitegenesis, these should have δ56Fe values of ~0.0±0.1‰,given the relatively homogenous nature of basaltic

magmas in terms of Fe isotope compositions (e.g. Beardet al. 2003; Schoenberg and Von Blanckenburg 2006;Weyer and Ionov 2007). Crystal fractionation in carbo-natite magmas (e.g. King and Sutherland 1960; Le Bas1989; Simonetti and Bell 1994b) will produce a decreasein δ56Fe values for carbonate, and this is a likelyexplanation for the samples that have the most negativeδ56Fe values. Generation or evolution of carbonatitemagmas through liquid immiscibility (e.g. Lee andWyllie 1998) will also decrease δ56Fe values in carbo-natites, given the positive silicate-carbonate fractionationfactors at igneous temperatures, and this mechanism willexert the greatest Fe isotope effect at low temperatures,where the fractionation factors are relatively large, or if aperalkaline, Fe3+-bearing silicate magma is involved. In amagmatic system where carbonate and silicate magmascoexist, the Fe isotope compositions of the carbonatemagma should be shifted greatest because of its relativelylow Fe contents, whereas the δ56Fe values of the silicatemagmas would likely remain close to zero to the degreethat the majority of the Fe is contained in the silicatecomponent of the system.

Fig. 7 Cartoon illustrating our preferred model for Fe isotopevariations in carbonatites. Silicate magmas in equilibrium with themantle will have δ56Fe values near zero, whereas carbonatites inequilibrium with the mantle will have δ56Fe values ≤ −0.3‰; thesecompositions apply to generation of carbonatites in plume environ-ments or the lithospheric mantle. Crystal fractionation or liquidimmiscibility will tend to move carbonatite magmas toward morenegative δ56Fe values, reaching values as low as ~ −1.0‰, whereasthese processes will not significantly change the δ56Fe values ofrelated silicate magmas. Residence of crystallizing carbonatite

magmas in the upper crust will release fluids through the outerportions of the intrusive complexes, as well as the surroundingcountry rocks, producing fenite zones. Relatively high δ56Fe valuesfor carbonates in carbonatites require addition of a high-δ56Fe fluid,which would most likely be an Fe3+-bearing fluid, given the relativelyhigh β56/54 factors for Fe3+ fluids (Table 4). Based on fluid-rockmodeling, and currently available β56/54 factors, the δ56Fe values ofcalcite-rich carbonatites might be increased up to +0.4‰ by passageof Fe3+-rich fluids

Iron isotopes in carbonatites 107

At high levels in the crust, continued crystallization ofcarbonatite magmas will produce a free fluid phase, which,based on the discussion above, is likely to be chloride- andFe-rich. At Oldoinyo Lengai, for example, the natrocarbo-natites show that Cl, F, and S are probably constituents ofcarbonatite fluids, and the mineralogy of the fenites indicatesthat the fluid must have contained Ca, Mg, K, Na, and Fe3+

(e.g. Morogan and Martin 1983; Gittins 1989). An NaClbrine was considered to be the fenitizing fluid at CallanderBay, Ontario, Canada (Currie and Ferguson 1971). Theabundance of acmitic pyroxene, riebeckite, and hematite infenites suggests that fenitization occurred at a depth wherethe fenitizing fluids had a relatively high f O2, probably closeto the hematite-magnetite buffer (Gittins 1989). Fenitizationof the surrounding granite terrane at the Iron Hillscarbonatite, Colorado, was attributed to an early- and alate-stage fluid. Fluid evolution, based mainly on pyroxeneand amphibole, began with high Mg/Fe ratios and high totalFe and Fe 3+, relative to the unaltered country rock, and asthe Mg/Fe ratio decreased, high Fe 3+ aegerine-augite andaegerine were formed (White-Pinilla 1996). All fluids fromIron Hills had overlapping temperatures between 510 and560°C (Lowers 2005).

Loss of Fe-rich fluids provides an additional means forproducing low-δ56Fe values in the remaining carbonatemagma because Fe-bearing fluids, particularly those thatare Fe3+-rich, will have relatively positive δ56Fe values(Table 4). We propose, however, that the high δ56Fe valuesin calcite were produced not by fluid loss but by netaddition of high-δ56Fe iron through passage of Fe3+-richbrines, increasing δ56Fe values in calcite up to +0.4‰(circular inset in Fig. 7). These brines must have passedthrough the carbonatites at relatively high temperatures toretain their high O isotope temperatures (Haynes et al.2003), as there is no correlation between δ56FeMt-CC

fractionations and O isotope temperatures (Fig. 4). Assum-ing a simple intrusion geometry, our proposal predicts aconcentrically zoned carbonatite complex, where the innerzones have relatively low δ56Fe values, and the outer partsof the carbonatites, as well as the fenites surrounding thecarbonatites, have relatively high δ56Fe values. In our broadsurvey, there is insufficient sampling at any single complexto test this proposal, nor are Fe isotope data available forfenites. Such studies would be fruitful lines of futureresearch that would provide a full understanding of the Feisotope mass balance of carbonatite complexes and theirassociated alteration and mineralization zones.

Conclusions

Previous stable isotope studies of carbonatites, including Li,C, and O, have provided constraints on crystallization

temperatures and fluid-rock interactions. Iron isotopevariations in carbonatites suggest that this relatively newstable isotope system can also provide a tracer of fluidinteractions and cooling history. Many minerals in thecarbonatites studied are out of Fe isotope equilibrium atigneous temperatures, and, considering the large contrastsin Fe contents among carbonate, silicate, oxide, and sulfideminerals, isotopic disequilibrium likely reflects the effectsof cooling and fluid/rock interaction. Iron isotope disequi-librium may also record mixing of phenocrysts fromdistinct magmas. Iron isotopes, therefore, provide addition-al evidence for isotopic disequilibrium in carbonatitemagmas, which previous stable (e.g. Haynes et al. 2003)and radiogenic (e.g. Simonetti and Bell 1993) isotopestudies have demonstrated. Because Fe isotope fractiona-tions between fluids and minerals/magmas are sensitive tooxidation state, Fe isotopes represent a stable isotopesystem that is uniquely poised to investigate the redoxstate of carbonatite magmas and their fluids. Moreover, thevery large range in Fe contents of carbonates, oxides,silicates, and sulfides provides a means to investigatedifferential mass-balance responses to differentiation andfluid-interaction processes. As our understanding of Feisotope fractionation factors improves, as well as Fediffusion rates in minerals, the exact mechanisms respon-sible for producing Fe isotope distributions among carbo-natite minerals and whole-rocks will be better understood.

Acknowledgments C.M.J., B.L.B., and A.I.S. thank the organizersof this special volume in honor of our co-author Keith Bell, includingguest editor Antonio Simonetti. This work was supported by theDepartment of Geology and Geophysics (U.W. Madison), theGeological Society of America, and the National Science Foundation(grant EAR-0525417). In addition to samples in the collection of K.B.,samples were provided by J.B. Dawson, D. Moecher, E. Haynes, andM. Spicuzza. Journal reviews were provided by R. Schoenberg and ananonymous reviewer, whose comments helped to improve themanuscript. We thank A. Simonetti for editorial handling of the paper.

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