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Lithos 125 (2011) 51–64 Contents lists available at ScienceDirect Lithos j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / l i t h o s Boron and boron isotope systematics in the peralkaline Ilímaussaq intrusion (South Greenland) and its granitic country rocks: A record of magmatic and hydrothermal processes Melanie Kaliwoda a, Horst R. Marschall b,c,⁎, Michael A.W. Marks d, Thomas Ludwig e, Rainer Altherr e, Gregor Markl d a Mineralogische Staatssammlung, LMU, Theresienstrasse 41, 80333 München, Germany University of Bristol, Department of Earth Sciences, Wills Memorial Building, Queens Road, Bristol, BS8 1RJ, United Kingdom c Woods Hole Oceanographic Institution, Department of Geology and Geophysics, Woods Hole, MA 02543, USA d Universität Tübingen, Institut für Geowissenschaften, Wilhelmstrasse 56, 72074 Tübingen, Germany e Universität Heidelberg, Institut für Geowissenschaften, INF 236, 69120 Heidelberg, Germany b a r t i c l e i n f o Article history: Received 8 November 2010 Accepted 23 January 2011 Available online 31 January 2011 Keywords: Boron isotopes Ilímaussaq Peralkaline intrusion Magmatic–hydrothermal SIMS PGNAA a b s t r a c t Concentrations of boron in whole rocks and minerals of the peralkaline, 1.16 Ga Ilímaussaq intrusion and its granitic country rocks (South Greenland) were analysed using secondary ion mass spectrometry (SIMS) and prompt gamma neutron activation (PGNAA) analysis. The intrusion consists of an early augite–syenite shell, a later alkali-granite sheet and still later nepheline syenites, which dominate the Complex. Boron concentrations are high (250–280 μg/g) in all rocks containing fresh sodalite, whereas boron is constantly low in the sodalite-free augite syenites (4–6 μg/g) and in the alkali granites (7–22 μg/g). Rocks with sodalite altered to analcime contain only low amounts of boron (2–7 μg/g), which records boron extraction by latemagmatic fluids. Concentration profiles of B in the analysed minerals (olivine, amphibole, clinopyroxene, aenigmatite, eudialyte, biotite, feldspar, nepheline and sodalite) record magmatic fractionation to various extents, latemagmatic to hydrothermal fluid/rock interaction, and sub-solidus diffusion. Whole-rock concentration data cannot be directly translated into the geochemical evolution of the peralkaline melts, since they are largely affected by cumulate fractionation of sodalite and amphibole and furthermore by late-stage hydrothermal alteration processes resulting in B loss. However, trace-element concentrations of mineral zones representing equilibrium fractionation from magmatic liquids can be used in combination with mineral–melt partition coefficients to unravel the enrichment processes of elements in the melt. Boron isotope values of minerals from the intrusion and the country rocks resemble the trend observed for Li isotopes in an earlier study. Amphibole and feldspar display a clear trend from light boron in the inner nepheline syenitic part of the intrusion (δ11B = −20‰ and −17‰ for amphibole and feldspar, respectively) through intermediate values in the outer augite syenites (δ11B = −10‰ and −6‰ for amphibole and feldspar, respectively) to heavy boron with δ11B = + 2‰ for amphibole and + 4‰ for feldspar in the country rock granites close to the contact with the intrusion. The values are interpreted to reflect the entry of meteoric fluids with heavy B along the intrusive contact. © 2011 Elsevier B.V. All rights reserved. 1. Introduction Boron and its isotopes are important tracers for mass transfer processes in terrestrial systems (e.g. Brenan et al., 1998a; Kaliwoda et al., 2008; Kessel et al., 2005; Leeman and Sisson, 2002; Marschall et al., 2006, 2007; Moriguti et al., 2004; Peacock and Hervig, 1999). Hence, it is ⁎ Corresponding author. E-mail address: [email protected] (H.R. Marschall). 0024-4937/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2011.01.006 important to characterise the geochemical reservoirs and to understand the compositional variations in terrestrial systems. Boron data are available for primitive mantle rocks (Chaussidon, 1995; Chaussidon and Jambon, 1994; Kaliwoda et al., 2008; McDonough and Sun, 1995; Ottolini et al., 2004; Palme and O'Neill, 2003), arc volcanic rocks and subduction-related rocks (e.g. Marschall et al., 2006; Ryan and Langmuir, 1993), but there is still only limited information on granitic to rhyolitic and on phonolitic to syenitic systems (e.g., Bailey, 2006; Sauerer and Troll, 1990; Tonarini et al., 2003). Boron data for minerals in syenitic rocks are so far only reported by Bailey (2006) and detailed 52 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 information or studies on B mineral–mineral, mineral–fluid and/or mineral–melt partitioning in magmatic systems of variable composition is still lacking. Whole-rock abundances of boron have been investigated for granitic to rhyolitic and for phonolitic to syenitic systems. They range from a few to a few hundred μg/g in most plutonic and volcanic rocks, but can reach several percent in tourmaline-bearing peraluminous rhyolites (e.g., Bailey, 2006; London et al., 2002; Pichavant et al., 1987). Tourmaline, however, is not stable in peralkaline melts and thus, B concentrations have been traditionally assumed to increase with progressive fractionation in tourmaline-absent systems (Morgan and London, 1989). Mineral/melt partition coefficients (DBmin/melt) in basaltic to andesitic systems are generally very low with B being highly incompatible (Brenan et al., 1998b; Klemme et al., 2002). Mineral/fluid partition coefficients (DBmin/fluid) are equally low, and B is considered to be highly mobile in fluids (Brenan et al., 1998a; Marschall et al., 2007a; Wunder et al., 2005). Minerals investigated in earlier studies include clinopyroxene, orthopyroxene, plagioclase, amphibole and garnet. Some of these minerals are relevant for granitic and syenitic rocks the present study deals with. However, it remains to be tested, whether compatibilities in peralkaline magmatic systems are similar, or if they are significantly influenced by additional liquid phases. Boron has two stable isotopes, 10B and 11B, occurring at an abundance ratio of ~ 1:4 in nature. Their large relative mass difference leads to an isotopic fractionation of N100‰ during geological processes (Barth, 1993). Boron isotope ratios are given in the delta notation with δ11B defined as deviation from standard NIST-SRM 951 (Catanzaro et al., 1970): 11 δ B= 11 10 B= B sample = 11 10 B= B NIST−SRM−951 3 −1 × 10 : Temperature-dependent equilibrium B isotope fractionation between two different phases is primarily driven by differences in coordination with oxygen (e.g., Palmer and Swihart, 2002). Threefold coordinated sites preferentially incorporate the heavier isotope (11B), while fourfold coordinated sites prefer the lighter isotope (10B) (Kakihana et al., 1977, 1982). Most silicates (e.g., amphibole, mica, pyroxene, feldspar) were, until recently, considered to incorporate B in tetrahedral coordination, substituting for Si, and would, therefore, fractionate 10B (Hervig et al., 2002; Tonarini et al., 2003; Werding and Schreyer, 2002). However, a recent study on B-doped synthetic diopside employing various spectroscopic methods demonstrated that B incorporation into this chain silicate reduces the tetrahedral sites to a trigonal coordination, where B replaces Si (Hålenius et al., 2010). Another important exception, beside some borates and boro-silicates, is sodalite, in which B is incorporated in both threefold and fourfold coordination as demonstrated by spectroscopic methods (Kaliwoda, unpublished data). The structural coordination of B in melts and fluids and the B isotopic fractionation involving these phases is furthermore a function of temperature, pressure, pH and composition (e.g., Hervig et al., 2002; Morgan and London, 1989; Schmidt et al., 2005; Thomas, 2002). The study of abundances and isotopic compositions of light lithophile elements with their contrasting behaviour in melt–solid– fluid systems provides an interesting tool to unravel the processes in magmatic–hydrothermal systems. However, the literature review above clearly shows that very little information exists on boron in intrusive complexes as the frozen equivalents of modern magma chambers (Bailey, 2006; Coradossi and Martini, 1981; Tonarini et al., 2001). In view of this aspect, we undertook a combined in-situ mineral and whole-rock study of the alkaline to peralkaline [=(Na + K)/ Al N 1.0] intrusive Ilímaussaq complex in South Greenland and its granitic country rocks. Indicators of differentiation in whole rocks and minerals (e.g., Fe/Mg, (Na + K)/Al, Ca/(Na + K), Rb/Sr, and Mg/Li) indicate that Ilímaussaq represents the most differentiated alkaline igneous rock suite yet documented (Bailey et al., 2001; Sørensen, 2001). Based on earlier work, the Ilímaussaq complex can be regarded as a textbook example of a mantle-derived peralkaline plutonic system that evolved largely as a closed system (e.g., Larsen and Sørensen, 1987; Marks et al., 2004, 2007). We derive insight into fluid–melt–mineral partitioning of boron in this well-studied peralkaline intrusive system. Boron concentrations and boron isotope compositions of various minerals were analysed in situ by secondary ion mass spectrometry (SIMS), and boron whole-rock analyses were completed by PGNAA (prompt gamma neutron activation analysis). In combination with published data on Li isotopes (Marks et al., 2007), Fe isotopes (Schoenberg et al., 2008), and Nd, O and H isotopes (Marks et al., 2004) for the same samples, our data provide a base for the investigation of B behaviour in complex alkaline magmatic systems and of the geochemical evolution of highly fractionated granitic to syenitic systems in general. 2. Geology The 1.16 Ga Ilímaussaq intrusion (Krumrei et al., 2006; Fig. 1) is part of the Mid-Proterozoic Gardar rift province in South Greenland (e.g., Emeleus and Upton, 1976; Upton et al., 2003). Ten major plutons of gabbroic, nepheline syenitic and granitic composition intruded into a basement of Early Proterozoic granites and gneisses (i.e. Julianehåb batholith; Garde et al., 2002) that are overlain by a succession of sandstones and basalts (i.e. Eriksfjord Formation; Poulsen, 1964). The Ilímaussaq Complex comprises four magmatic pulses (phases I, II, IIIa and IIIb) containing both nepheline- and quartz-normative rocks (e.g., Ferguson, 1964; Markl et al., 2001; Sørensen, 2001). The northern part of the complex (interpreted as the roof region) intrudes sandstones and volcanic units of the Eriksfjord Formation. In the southern part (hence the lower part), the Julianehåb batholith forms the country rock of the Complex (Fig. 1). The fractionation process of the Ilímaussaq magmas took place in a deep magma chamber and continued during transport to a shallow emplacement level of around 4 km depth (100 MPa; Konnerup-Madsen and Rose-Hansen, 1984). The initial magma was characterised by very low fO2 (ΔlogFMQ= −3 to −5; Marks and Markl, 2001), resulting in methane (CH4) as the main stable species in the fluid phase (e.g., Krumrei et al., 2007). Increasing fractionation in the third and fourth magma batch led to the oxidation of the magma, and eventually to the stabilisation of an H2O-dominated fluid phase (e.g., Krumrei et al., 2007; Markl and Baumgartner, 2002; Markl et al., 2001). Prolonged hydrothermal activity led the Ilímaussaq rocks to crystallise over an extended temperature range from 1000 °C down to ~300 °C (Larsen, 1976; Markl, 2001; Markl et al., 2001; Marks and Markl, 2001; Marks et al., 2004). In general, the Ilímaussaq complex can be regarded as a largely closed system during most of its orthomagmatic evolution (Graser and Markl, 2008; Larsen and Sørensen, 1987; Marks et al., 2004, 2007). The initial magmatic pulse (phase I) formed a SiO2-saturated to weakly under-saturated augite syenite, with alkali feldspar, augite, olivine and Fe–Ti oxides as major phases, plus minor amounts of nepheline, Ca-amphibole and biotite. Today, these rocks occur as the outer rim of the intrusion and as xenoliths within the subsequently intruded nepheline syenites (Fig. 1). The second magmatic pulse (phase II) crystallised a peralkaline granite, which is interpreted as an evolved equivalent of the augite syenite (phase I) contaminated with lower crustal material (Marks et al., 2004). Major minerals are alkali feldspar, quartz, Na-amphibole and minor aegirine. The main volume of the complex consists of a series of peralkaline nepheline syenites, related to two successive, but independent magma batches (Larsen and Sørensen, 1987; Markl et al., 2001; Sørensen, 1997, 2001; Sørensen et al., 2006). These nepheline syenites M. Kaliwoda et al. / Lithos 125 (2011) 51–64 53 lik rmi Se GM 1212 GG02 GM 1257 GM 1369 GM 1370 GM 1843 GM 1303 q GM 1342 se Ta ik iarf gl unu GM 1246 GM 1294 T ILM124 GM 1272 Narssaq Intrusion Julianehåb batholith Eriksfjord Formation GM 1219 GM 1214 GM 1337 GM 1335/36 augite syenite Phase I alkali granite Phase II sodalite foyaite Phase IIIa lujavrite Phase IIIb kakortokite agpaites naujaite k su rs agpaitic dike lua d er ng fault Ka sample locality with sample number GM 1857 GM 1858 JG5-JG13 0 2 4 km magmatic pulses augite syenite alkali granite Phase II Phase I pulaskite sodalite foyaite naujaite Phase IIIa lujavrite Phase IIIb kakortokite 2 km Fig. 1. Geological map of the Ilímaussaq intrusion (South Greenland) after Ferguson (1964). Sample localities marked with black circles. Lower panel: schematic vertical cross section of the Ilímaussaq intrusion. consist of variable amounts of sodalite, feldspar, nepheline, Naamphibole, clinopyroxene, eudialyte (we use the term “eudialyte” for all eudialyte-group minerals) and minor aenigmatite. The latter two minerals classify these highly evolved rocks as agpaites. Based on different textures and modal compositions, the whole series is subdivided into (1) a roof series (phase IIIa) and (2) the sandwich and floor series (phase IIIb). Rocks of the roof series were formed by downward crystallisation and flotation of minerals (mainly sodalite) less dense than the melt. From top to bottom they mainly consist of pulaskite (nepheline-bearing syenite), sodalite foyaite (sodalite-bearing syenite with a foyaitic texture) and naujaite (nepheline sodalite syenite with a poikilitic texture). Within this series, naujaite represent a more than 600 m thick weakly layered body (Krumrei et al., 2006, 2007; Rose-Hansen and Sørensen, 2002) forming approximately 40% of the complex. Rocks of the floor and sandwich series comprise kakortokites (nepheline syenites with pronounced cumulate textures and igneous layering with a repetition of layers enriched in feldspar, Na-amphibole and eudialyte, respectively) and lujavrites (melanocratic and finegrained nepheline syenites with pronounced igneous lamination), respectively. Kakortokites form an approximately 300 m thick strongly layered sequence. Within one of these layers, autoliths of augite syenite, sodalite foyaite and naujaite are present. The overlying lujavrites are close to a fluid-rich kakortokitic residual liquid and intrude the overlying, already solidified naujaites (Ferguson, 1964; Larsen and Sørensen, 1987; Pfaff et al., 2008; Sørensen, 2001). Late-stage activity at Ilímaussaq is represented by pegmatites (e.g., Müller-Lorch et al., 2007) and late-magmatic to hydrothermal veins (e.g., Graser and Markl, 2008; Markl and Baumgartner, 2002). Most hydrothermal veins consist of variable amounts of albite, nepheline, sodalite, analcime, sodic clinopyroxene (aegirine) and Na-amphibole (arfvedsonite) and may contain a wealth of unusual accessories (e.g., astrophyllite, neptunite and others). Some hydrothermal veins record fluids with very high pH (Markl and Baumgartner, 2002) or with high Be concentrations, evident from significant modes of Be minerals (e.g., tugtupite; Markl, 2001). 54 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 3. Petrography of sample material Augite syenites (phase I: sample ILM193, GM1858, 1857, 1330 and 1332) are medium to coarse grained. Early magmatic minerals are perthitic alkali feldspar, olivine, augite, Fe–Ti-oxides and apatite. Latemagmatic phases are nepheline and Ca-amphibole. A second generation of fine-grained amphibole intergrown with biotite, surrounding olivine and magnetite are probably of hydrothermal origin (Marks and Markl, 2001). Peralkaline granites (phase II: GM1303, 1342) are coarse grained and consist of euhedral Na-amphibole, alkali feldspar, abundant interstitial quartz and minor zircon. Late-stage aegirine overgrows occasionally magmatic Na-amphibole. Rocks of the roof series (phase IIIa) include pulaskite, sodalite foyaite and naujaite. In pulaskite (sample P3–5), early magmatic phases are olivine, clinopyroxene, nepheline, Fe–Ti oxides and alkali feldspar. The feldspar is almost completely altered to zeolite, olivine and Fe–Ti oxides are partly overgrown by aenigmatite, and late orthomagmatic Na-amphibole replaces former pyroxenes. In sodalite foyaites (samples GM1214, 1219), orthomagmatic phases are alkali feldspar, sodalite, Na-amphibole, eudialyte and minor olivine. Most amphiboles are resorbed along their rims and overgrown by aegirine. Sodalite is altered to analcime, and eudialyte shows only tiny relics with unaltered composition. Naujaites (samples GM1369, 1370) contain large euhedral sodalite grains (1–4 mm), and interstitial alkali feldspar, nepheline and amphibole. Na-amphibole is partly replaced by aegirine. In sample GM1370, large amounts of interstitial aenigmatite with inclusions of former Na-amphibole and aegirine occur. In contrast to the sodalite foyaite samples, alkali feldspar and eudialyte in the naujaites are not altered. Phase IIIb consists of kakortokites (floor series) and overlying lujavrites (sandwich horizon). The generally coarse-grained kakortokites (red eudialyte-rich sample: GM1335; white feldspar-rich sample: GM1336 and black amphibole-rich sample: GM1337) consist of euhedral eudialyte and perthitic alkali feldspar and subhedral amphibole and clinopyroxene intergrown with each other. Minor minerals are sodalite, nepheline, aenigmatite and late-stage albite laths, which occur along grain boundaries. Lujavrites (samples GM1294 and 1843) are fine-grained and show a fluidal texture. They mainly consist of almost pure albite and microcline along with eudialyte, sodalite, nepheline, Na-amphibole and minor aegirine. A late-stage agpaitic dyke (sample GM1212) with a fluidal texture exhibits macroscopically visible ocelli-like features produced by exsolution due to liquid immiscibility (Markl, 2001). The sample investigated consists of a fine-grained mixture of aegirine and Naamphibole with very minor feldspar in the dark portions. Lightercoloured ocelli are dominated by feldspar, analcime and small amounts of Na-amphibole. Three different types of hydrothermal veins were investigated: Sample GG02 consists of tugtupite in a matrix of fine-grained albite, Na-amphibole, aegirine and minor neptunite. Sample GM1257 is taken from a typical albite–aegirine vein crosscutting the naujaite. It comprises a fluidal texture and consists of fine-grained albite and needle-shaped aegirine. Sample GM1246 consists of large amounts of ussingite and large aegirine grains (up to 1 cm), partly containing Na-amphibole inclusions. A second aegirine generation forms tiny (b100 μm) aggregates accumulated in clusters. This sample also contains accessory feldspar. From the host Julianehåb granite, samples JG5–JG13 were taken along a traverse towards the marginal augite syenite at distances of between 0.2 and 225 m from the contact (Fig. 1). All samples contain quartz, altered alkali feldspar, plagioclase and minor amounts of biotite, which is intergrown with magnetite of variable grain size. Some of the biotite grains are marginally altered or completely transformed to chlorite. Samples from closer distance (≤45 m) to the contact with the Ilímaussaq rocks (JG10, 11, 12 and 13) contain additional Na-pyroxene ± Na-amphibole. In these samples, pyroxene is in contact with quartz and feldspar. Na-rich amphibole and clinopyroxene in the granite samples are interpreted to have precipitated from fluids released from the Ilímaussaq intrusion (Marks et al., 2007). However, this type of alkali metasomatism in the country rocks has yet to be investigated in detail. 4. Methods Whole rock B concentrations were determined using prompt gamma neutron activation analysis (PGNAA, Budapest research reactor, Hungary). The reactor unit is equipped with a cold neutron source (20 K), with a neutron flux of ~ 5 × 107 cm− 2 s− 1 at the target position. The beam area was set to 4 cm2 with an exposure time of 1–4 h for the samples measured. Gamma radiation in the energy range of 30 keV to 11 MeV was detected using a high-purity germanium semiconductor (HPGe) bismuth germanate (BGO) scintillator detector system in Compton-suppressed mode. A Canberra S100 multi-channel analyser performed the data capture. Gamma spectra were evaluated using the programme Hypermet PC (Révay et al., 2004). The accuracy of the boron gauging was controlled by analyses of reference materials provided by the Budapest Neutron Center (BNC) and was ~ 10% relative (see also Gméling et al., 2005, 2007). The relative precision is ~ 1.5% for concentrations N5 μg/g and ~ 1.7% for concentrations in the range of 1.9 to 5 μg/g (Gméling et al., 2005, 2007; Marschall et al., 2005). Using the standard setup, the detection limit for boron in natural samples and standard materials is ~ 0.3 μg/g. We also determined whole-rock concentrations of Li and Be by atomic absorption spectrometry (AAS). 0.5 g of powdered rock sample was mixed with water and 25 ml of HF–HClO4 (HF 40%–HClO4 70%). This mixture was heated twice, first to 80 °C within a platinum cup on a sand-bath and fumed off overnight and then to 130 °C until a crystal mush was produced. This sample cake was spiked with 6 ml HCl, heated up and diluted with distilled water to 100 ml. Solutions used for calibration were diluted from 1000 μg/g standard solutions from Merck® to 5, 10 and 20 ng/g for Be and to 1, 2 and 3 μg/g for Li, respectively. The sample solutions to be analysed were diluted to a concentration falling within the calibrated concentration range. Beryllium was analysed using the graphite tube (Perkin-Elmer Zeeman 4110), lithium by the use of flame AAS (AAS Vario Analytik Jena) using the Be-line at 243.9 nm and the Li-line at 670.8 nm, respectively. The detection limits within the solutions are b1 ng/g for Be and 0.05 μg/g for Li, respectively. The overall relative precision (1 standard deviation) for Be is ≤8%, while that for Li is ≤0.8%. Concentrations of boron (and Li and Be) in minerals (olivine, clinopyroxene, amphibole, biotite, aenigmatite, eudialyte, feldspar, nepheline, sodalite, and quartz) and B isotope ratios of amphibole, feldspar and sodalite were measured by secondary ion mass spectrometry (SIMS) using a Cameca ims 3f ion microprobe at the Institut für Geowissenschaften (Universität Heidelberg, Germany). All analyses were performed using a 14.5 keV 16O− primary ion beam. 4.5 keV positive secondary ions were counted using a single electron multiplier. Samples were cleaned as described in Marschall and Ludwig (2004). Pre-sputtering time (including peak calibration) was 4–5 min. For the concentration measurements, the energy window was set to 40 eV and the energy filtering technique was applied with an offset of 75 eV at a mass resolution of m/Δm ~1000 (at 10%) in order to suppress interfering molecules and to minimise matrix effects (Ottolini et al., 1993). The primary beam current was 20 nA resulting in a spot size of ~20 μm. The imaged field was limited to a diameter of ~12 μm by using a 700 μm field aperture (imaged field mode 25 μm, FA #2), so that only secondary ions originating from the centre of the sputtered crater contributed to the analyses. This technique reduces the influence of surface contamination to an apparent B concentration below the detection limit of ~2 ng/g (Marschall and Ludwig, 2004). The results have not been corrected for M. Kaliwoda et al. / Lithos 125 (2011) 51–64 contents are high in rocks with unaltered sodalite, which is the main boron carrier (Fig. 3). The very low boron content of the altered sodalite foyaite samples (2–7 μg/g) may be explained by the lack of sodalite, now replaced by analcime. Feldspar and nepheline are main B carriers in all rocks lacking sodalite (Fig. 3). 5.2. Boron abundances and zoning patterns in individual mineral grains Fresh sodalite is the major boron carrier (up to 173 μg/g; Figs. 3–5, Table 1). Boron content decreases in sodalite altered to analcime (e.g., in the sodalite foyaites; phase IIIa), which contains only 0.9–2.3 μg/g B (Fig. 3; Table 1). In contrast to its behaviour in sodalite, B concentration in nepheline increases from kakortokite to lujavrite (Fig. 4; Table 1). Amphibole (up to 4.6 μg/g) and pyroxene (up to 1.9 μg/g) from all rock types display no correlation between their average boron content and magmatic fractionation (Fig. 4; Table 1). Interestingly, the B content in pyroxenes from the granitic country rocks is considerably higher (6.2–7.2 μg/g; Fig. 4; Table 1). The same is observed for quartz, which shows higher B contents in granitic country rocks (0.1–4.2 μg/g) compared to quartz in the alkali granites of the intrusion (0.04–0.6 μg/g; Fig. 4; Table 1). Biotite in the granitic country rocks and in the augite syenites (phase I) have identical compositional ranges (0.5–5.5 μg/g; Fig. 4; Table 1). B abundances in eudialyte decrease from pulaskite to sodalite foyaite and naujaite from about 6.1 to 1.1 μg/g, but show an increase in the final melt batch, kakortokite (1.9 μg/g) to lujavrite (12.3 μg/g). Nearly all minerals display complex and highly unsystematic zoning patterns in B (as well as Li and Be). Zoning in the minerals investigated points to a composite record of fractionation processes during mineral growth and sub-solidus diffusive redistribution. Typically, the core of a mineral shows the older magmatic signature, while the rim records diffusive re-equilibration with adjacent minerals. In the following, we will present some characteristic boron zoning features in order to discuss the processes that may have been involved in their generation. A more complete record of mineral profiles across the grains and a discussion of Li and Be zoning is available in Appendix A. Boron concentrations in amphibole either increase or decrease in contact with feldspar or quartz (Fig. 5). But they are constant if amphibole is in contact with magnetite. Some amphibole grains in the sodalite foyaites and naujaites (phase IIIa) show a complex W-shaped pattern, i.e., a decrease from core to rim and a subsequent increase in sodalite-free samples sodalite altered to analcite sodalite-bearing samples data from Bailey et al. (2001) and Bailey (2006) increasing fractionation 1000 Li B whole rock (µ µg/g) the combined background of the mass spectrometer and the counting system of 0.02 ± 0.01 s− 1, corresponding to concentrations of ≤1 ng/g. Each concentration analysis comprises N = 10 cycles with total integration times of 80 s for Li, 160 s for Be and B and 20 s for Si. The average ratios (Li/Si, Be/Si, and B/Si) are then used to calculate pffiffiffiffi the concentration and the standard deviation of the mean (1σ = N) is reported as internal precision, which is dominated by counting statistics at trace element concentrations. For the setup chosen and assuming a concentration of 1 μg/g in the sample, the internal precision is ~ 2% for Li, ~ 1.5% for Be and ~ 3% for B (see Appendix A). For higher concentrations the internal precision improves until it is dominated by other effects, e.g. the stability of the primary ion source and the stability of the mass spectrometer's magnet. Prior to each analytical session 5 analyses were performed on the NIST SRM610 glass with a typical precision (1 RSD) of ≤1% (see Appendix A for a compilation of all Li, Be and B calibration data from 2003 to 2009 of the Heidelberg SIMS). Secondary ion yields of Li, Be and B relative to Si (RIY) were then calculated from the average ratios of these analyses using concentrations (preferred averages) taken from Pearce et al. (1997). Regarding the accuracy of Li, Be and B, SIMS analyses we refer to Ottolini et al. (1993) who reported ±20% relative accuracy for Li and ±10% for Be and B. For Li see Marks et al. (2008), where SIMS analyses (same setup, calibration method, reference material and ionprobe as in this work) were compared to MC ICP-MS analyses of (i) an amphibole (638± 18 μg/g SIMS vs 667 ± 35 μg/g MC ICP-MS) and (ii) a Na-pyroxene (46.9 ± 1.7 μg/g SIMS vs 43.0± 1.8 μg/g MC ICP-MS). In case of B, we compared our SIMS result of 203 ± 1 μg/g for sample B6 (obsidian) to the mean value of 203.8 ± 8.9 μg/g reported in Gonfianti et al. (2003). Furthermore, we compared the Li, Be and B RIYs of the reference minerals described in Dyar et al. (2001) to the RIYs obtained with the NIST SRM610 glass. The maximum RIY deviations relative to SRM610 were +10.2% for Li, −10.2% for Be and −13.7% for B (see Appendix A). Although it would certainly be worthwhile to improve the accuracy of ~10% for trace element analyses, we still prefer to base all analyses on the widely distributed and well determined reference glass SRM610 instead of using sample-matched in-house reference materials of very limited distribution, which are typically not well-determined (e.g. only by one method and one lab). The SiO2 content of the minerals was determined by electron probe micro analyser (EPMA; Universität Tübingen). As the composition of the minerals analysed in our samples closely conform to earlier analyses of Larsen (1976), Markl et al. (2001) and Marks et al. (2004), they are not repeated here. Boron isotope ratios were analysed at m/Δm=1200, sufficient to separate 10B1H+ interference on 11B+, and 9Be1H+ interference on 10B+. The imaged field diameter was 150 μm (150 μm imaged field mode and FA #1). The energy window was set to 100 eV and no offset was applied. Integration times per acquisition cycle were 3.32 s for 10B and 1.66 s for 11 B. The primary beam current was 30 nA with a spot size of ~30 μm. Instrumental mass fractionation was determined using the obsidian glass B6 with δ11B= −1.8‰ (average of values taken from Tonarini et al., 2003 and of P-TIMS values in Gonfianti et al., 2003). The internal precision (1σm of the mean ratio over all acquisition cycles, dominated by counting statistics) and the external reproducibility of the analyses on B6 with N=100 cycles both were 0.6‰ (1σ). For a discussion of possible matrix effects see Rosner et al. (2008). For the analyses of the unknown samples N ranged from 100 to 400, depending on the B concentration. The internal precision of single analyses is again given as 1σm. 55 100 Be 10 5. Results 1 5.1. Boron whole rock contents Boron whole-rock concentrations increase from 3 μg/g in phase I (augite syenites) to 22 μg/g in phase II (alkali granites). In phase IIIa and IIIb, B concentrations vary between 2 and 40 μg/g (Fig. 2). Boron phase I phase II phase III a phase III b late stage/ veins Fig. 2. Boron whole rock concentrations within the four melt batches (phase I–IIIa, b) and the late stage fluid rich veins. Error bars are smaller than symbols. Li and Be contents are given as shaded areas for comparison (and are given in Appendix A). 56 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 mass fraction of B (% of whole rock) 100 Eud+Am+Cpx Eud+Am+Cpx 80 Eud+Am+Cpx Ne Ne 60 Sdl Sdl (altered) 40 20 Akf 0 Akf Akf pulaskite sodalite foyaite kakortokite Fig. 3. Boron budgets for three different rock types: pulaskite, sodalite foyaite and kakortokite. Budgets are calculated by comparing the B concentrations in the minerals multiplied by their modal abundance with the measured whole-rock boron abundances. contact with feldspar and pyroxene. The B zoning profile of amphibole in kakortokites and lujavrites is rather homogenous with some outliers related to cracks in the analysed minerals. Olivine, measured in augite syenites, are unzoned in B, whereas B contents of most pyroxene grains in augite syenites, sodalite foyaites, naujaites and kakortokites increase in contact with feldspar and with olivine. All other pyroxene grains – even pyroxene inclusions – show homoge- 1000 hydrothermal veins lujavrite Phase IIIb kakortokite naujaite Sdl foyaite alk. granite Phase II Phase IIIa pulaskite Phase I augite syenite Julianehåb granite Ilímaussaq intrusion: increasing fractionation Cpx feldspar B (µg/g) 100 amphibole 10 1 0.1 0.01 1000 B (µg/g) 100 nepheline sodalite quartz eudialyte biotite 10 1 0.1 neous boron concentrations. Biotite in augite syenites and host granites is homogeneous in boron, except for biotite in granites in contact with feldspar. Eudialyte grains in sodalite foyaites, kakortokites and lujavrites are also homogeneous in B, with an occasional sharp increase next to cracks and directly at the grain margins. Most feldspars show no systematic zoning in B. However, some alkali feldspar grains from augite syenites, alkali granites and sodalite foyaites show an increase in B in contact with amphibole. In contrast, boron in albite in lujavrites decreases from core to rim next to pyroxene and amphibole. Boron in some feldspar crystals in the country rocks rises in contact with quartz and with biotite. Nepheline is generally homogeneous in B with the exception of one grain in a kakortokite, which displays a core-to-rim decrease in B in contact with amphibole. As mentioned before, the highest boron contents were found in sodalite, with different boron concentrations in different growth zones of the grains, deduced from variably luminescent zones in CL images (Figs. 5 and 6). Cores have higher B concentrations than subsequently grown parts and rims of crystals (Figs. 5 and 6). Superimposed on this apparent growth zoning, B concentrations of sodalite generally increase in contact with amphibole, nepheline, feldspar and pyroxene. 5.3. Boron isotope zoning patterns of different minerals Boron isotopic compositions (expressed as δ11B) were determined for sodalite, amphibole and feldspar (Table 2; Figs. 6 and 7). Boron abundances of other minerals, such as eudialyte, nepheline, biotite and pyroxene are too low for B isotope analyses, and tiny mineral inclusions, cracks and fluid inclusions additionally complicate the analysis of these minerals. Boron isotope values in the cores of sodalite grains vary in a small range between − 15‰ and − 10‰ (Table 2; Fig. 6). δ11B values of amphibole and feldspar scatter around −20‰ and − 17‰, respectively, in rocks from the inner part of the intrusion (phases II, IIIa and IIIb; Table 2; Fig. 7), but increase to higher and even positive values at the rim of the intrusive complex (phase I) and in the granitic country rocks (Table 2; Fig. 7), the same was observed for δ7Li in amphibole (Marks et al., 2007, Fig. 7). As for B concentration profiles, we observed a core to rim variation of δ11B values in the different minerals. Again, δ11B values in the rim of a particular mineral appear to be affected by the adjacent minerals. For example, sodalite shows commonly decreasing δ11B values from −13.2‰ in the core to −17.4‰ at the rim in contact with feldspar (Fig. 6) and to −13.6‰ in contact with pyroxene (Fig. 6). Invariably, the decrease of δ11B from core to rim in sodalite is more pronounced in contact with feldspar than in contact with pyroxene. Surprisingly, cores of sodalite grains have significantly higher B contents than the subsequent growth zones (discriminated by CL), but their δ11B values vary only between −10.0 and −12.0‰ (Fig. 6). In contrast, δ11B in amphibole in all samples (with the exception of one pulaskite and one naujaite) is lower in the cores than in the rims of crystals (Fig. 6), if adjacent to feldspar. In contact with pyroxene or magnetite, however, only minor core–rim variations are observed. Feldspar grains only show a change at the contact with amphibole (core: −15.9‰, rim: −19.8‰) and pyroxene (from − 13.4 to −20.9‰, Fig. 6). 6. Discussion 1246 1257 1212 GG02 1294 1843 1336 1337 1369 1370 1214 1219 P3-5 1303 1342 1858 1330 1332 1857 ILM-193 JG5 JG6 JG7 JG8 JG9 JG10 JG11 JG12 JG13 0.01 Sample Fig. 4. Boron contents of amphibole, feldspar, pyroxene, nepheline, sodalite, quartz, eudialyte and biotite within the different rocks of Phase I to IIIb, the hydrothermal veins and the Julianehåb granite. 6.1. Boron budget of whole rocks Boron concentrations in minerals decrease in the order BSdl ≫BNe N BFsp N B Eud N BAm N BCpx N BBt N BOl. Sodalite is the main mineral carrier for boron in all sodalite-bearing rocks, hosting generally N70% of the total B budget (Fig. 3). The alteration of sodalite to analcime, however, is accompanied by a loss of boron, which explains the low M. Kaliwoda et al. / Lithos 125 (2011) 51–64 57 Fig. 5. Boron profiles measured by SIMS across amphibole and sodalite grains from four different rock types (i.e., augite syenite, alkali granite, naujaite and kakortokite). Boron zoning is partly dependent on the neighbouring mineral grains. Furthermore, the different growth zones are visible in cathodoluminescence images, and the oldest zone shows the highest B content, compared to the younger ones. (Error bars are mostly smaller than symbol size). B concentrations found in sodalite foyaites. Interestingly, lujavrites, which were derived from the final-stage melts, show a slight increase in boron, which could point to an additional B-rich melt batch (a hypothetical phase IV), or preferably to the input of B by fluids. Boron-rich fluids may have been generated, for instance, by the interaction of late-stage fluids with magmatic sodalite. In sodalite-free rocks, nepheline and feldspar are the major boron carriers (Fig. 3) and, to a lesser extent, amphibole, eudialyte and pyroxene. In general, the B concentrations in the Ilímaussaq whole rocks do not record clear differentiation trends, but are largely influenced by the accumulation of sodalite and its late-stage alteration. In summary, B (as well as Li and Be; see Appendix A) whole-rock data do not directly reflect the geochemical evolution of peralkaline melts, since the concentrations analysed are largely affected by cumulate formation (in particular sodalite and amphibole) and, in the case of B, also by late-stage hydrothermal alteration processes resulting in a loss or redistribution of B. 6.2. Significance of boron zoning in individual minerals The strong dependence of core-to-rim profiles on the respective adjacent minerals allows us to draw two important conclusions: (i) zonations formed while the adjacent minerals were already in place and the minerals were no longer enclosed by melt, and (ii) the transport of B along grain boundaries was inefficient and solid-state diffusion was dominant at this stage. These two restrictions together demonstrate that the asymmetric core-to-rim zonations were generated at fluid-absent sub-solidus conditions. Redistribution of B was probably driven by cooling in combination with the temperature dependence of inter-mineral partition coefficients. Boron concentrations of pyroxene and amphibole increase at the rim in contact with feldspar, indicating B diffusion from feldspar into pyroxene and amphibole. Apart from one exception, the sodalite foyaite, where sodalite is altered, B and Be show a similar behaviour in all rocks investigated. The lack of a significant fractionation of B from 58 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 Table 1 Boron content (μg/g) of minerals in different rock types of the Ilímaussaq intrusion and the Julianehåb country-rock granite. Magma pulse Rock type Sample Amphibole Clinopyroxene Olivine Alk. feldspar Albite Sodalite Nepheline Quartz Eudialyte Biotite Phase I Augite syenite Phase II Alkali granite Phase III Pulaskite Sdl foyaite ILM193 GM1858 GM1330 GM1332 GM1857 GM1303 GM1342 P3–5 GM1214 GM1219 GM1369 GM1370 GM1336 GM1337 GM1294 GM1843 GM1212 GG02 GM1246 GM1257 0.48 (6) 1.60 (5) 0.44 (8) 0.40 (15) 0.75 (26) 2.26 (63) 4.2 (16) 0.87 (61) 1.73 (53) 0.99 (17) 1.07 (5) 4.55 (419) 1.49 (56) 1.46 (43) 0.77 (8) 0.57 (10) 0.77 (8) 0.57 (10) – – 0.22 (12) 0.51 (18) 0.30 (26) 0.14 (3) 0.15 (2) – – 1.07 (27) 0.23 (5) 0.36 (1) 1.93 (58) 0.23 (5) 0.42 (12) 0.36 (12) – – – – 0.10 (6) 0.77 (55) – 0.23 (13) 0.17 (17) – 0.13 (3) – – – – – – – – – – – – – – – 0.23 (10) 1.14 (41) 1.22 (73) 1.70 (89) 3.6 (17) 4.4 (8) 7.1 (17) 3.6 (14) 0.67 (19) 0.91 (8) 0.36 (5) – 1.7 (10) 5.0 (18) – – 0.17 (6) 0.04 (2) – 0.19 (7) – – – – – – – – – – – – – – 1.72 (26) 0.36 (7) – – – – – – – – – – – – 1.58 (55) 1.88 (31) 25.3 (63) 59.7 (48) 126 (36) 173 (95) 32.9 (24) 157 (120) – – – – – – – – – – – – – – – – 1.27 (4) 1.41 (14) 5.36 (80) 19 (15) – – – – – – – – – 0.05 (1) 0.61 (1) – – – – – – – – – – – – – – – – – – – – – 6.1 (19) – 1.11 (23) 0.97 (39) – 1.89 (10) 12.3 (36) – – – – – – 0.60 (5) 5.49 (4) – 1.08 (10) – – – – – – – – – – – – – – – – – – – – – – – 0.52 (22) 2.5 (12) 3.06 (86) 5.5 (17) 27 (17) 6.5 (18) 5.4 (28) 6.1 (12) – – – – – – – – – – – – – – – – – – – – – – – – 0.26 (10) 0.76 (18) 0.67 (48) 0.88 (47) 0.65 (18) 3.4 (12) 1.25 (20) 0.79 (24) – – – – – – – – 1.90 (60) 1.43 (52) 1.6 (11) 2.75 (13) – 4.63 (52) 2.39 (25) 1.39 (94) Naujaite Phase IV Kakortokite Lujavrite Hydr. veins Other minerals: Country rock Granite GM1370 aenigmatite: 1.72 (52) μg/g GG02 neptunite: 0.58 (42) μg/g JG5 – – JG7 – – JG8 – – JG9 – – JG10 – – JG11 – 6.2 (32) JG12 – – JG13 – 7.2 (62) All concentrations in (μg/g). Numbers in parentheses give 2 standard errors on the last given digit. Be implies that fluid exsolution did not play a major role in the orthomagmatic evolution of the complex. Otherwise, B abundances should be strongly affected, while Be would show a rather gentle evolution, based on the strong difference in the relative fluidmobilities of the two elements (e.g., Brenan et al., 1998a; Kaliwoda et al., 2008). The sodalite foyaites, which were altered by late-stage hydrous fluids show this effect. These sodalite grains are depleted in B, from ~ 70 μg/g in sodalite to 2 μg/g in the alteration products, whereas the decrease in Be is minor (i.e., from ~ 50 μg/g to ~30 μg/g). The different growth zones of individual sodalite grains discriminated in CL images (Fig. 6) display different B contents. The, old, cores consistently have higher B contents (80–300 μg/g) compared to the subsequently crystallised zones (70–80 μg/g). The final growth zones are lowest in B (50–70 μg/g). The link of B concentrations to sodalite growth zones demonstrates a limited intra-crystalline diffusivity of B in the sodalite, and decreasing concentrations are in agreement with a compatible behaviour of B. 6.3. Effects of hydrothermal alteration There are two important alteration effects in the Ilímaussaq minerals: (i) The alteration of sodalite to analcime causing boron loss of ~90%, which is related to late-magmatic to hydrothermal fluids in the temperature range of 500–200 °C (Markl, 2001). The fluid-mobile element B is depleted in rocks, in which abundances of Be, which is relatively immobile in hydrous fluids, are more or less constant. (ii) In contrast, the B contents of fresh and altered eudialyte are similar (but differ strongly in both Li and Be; see Appendix A). We have no conclusive explanation for this non-enrichment in B, but it is probably related to a crystal-chemical effect in eudialyte in comparison to its alteration products. 6.4. The possibility of kinetic fractionation of boron isotopes The margin of the Ilímaussaq intrusive complex and the adjacent Julianehåb granite country rocks were influenced by fluid–rock interaction during final cooling of the Ilímaussaq intrusion (Graser and Markl, 2008; Marks et al., 2007). As such, kinetic Li isotope fractionation caused by diffusion of Li from the Li-rich alkaline rocks into the country rock produced very low δ7Li values in the latter, and an increase of δ7Li values in the former (Marks et al., 2007). For boron isotopes one may speculate on the possibility of a similar mechanism. Kinetic fractionation of the two isotopes 10B and 11B needs to be evaluated for diffusion of B in three different types of phases that interacted in the intrusion, namely hydrous fluid, silicate melt and silicate minerals. Silicate minerals investigated in this study, such as amphibole, feldspar and sodalite, show asymmetric B zoning, depending on the minerals in contact with the analysed grains. This is taken as evidence for sub-solidus redistribution of B among different minerals in the rocks by solid-state diffusion. Disturbed B isotope ratios in the rim domains of the minerals, causing variations of up to 10‰, contrasting with homogeneous cores strongly suggest that B isotopes may indeed be fractionated kinetically during diffusion of 10B3+ and 11B3+ through the crystal structures. This, however, is restricted to intra-crystalline diffusion. Boron, in contrast to Li and other alkali metals, does not form ionic bonds with oxygen. Instead, B–O bonds have covalent character and B does not occur as hydrated B3+ ion in hydrous fluids, but forms charged and neutral B–O-bearing complexes. This is of great significance for the diffusion mechanism of B in fluids and, hence, for its possible isotopic fractionation. The two dominant B complexes 11B(OH)3 and 11B(OH)− 4 have masses of approximately 62 and 79 u, respectively. In high-pH 10 B fluids, only B(OH)− 4 occurs, the relative mass difference between the − − 11 (OH)4 and B(OH)4 complexes are small (~ 1.3%) and kinetic fractionation is, therefore, unlikely. In near-neutral pH fluids, B(OH)3 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 59 Fig. 6. (a, b, e) Boron isotopic composition of amphibole (Am) is lighter in the cores compared to the rims of the grains. (c, d, e) Alkali feldspar (Akf) and sodalite (Sdl) are lighter in the rim, compared to the core of the grains. The δ11B values depend on the neighbouring minerals: alkali feldspar displays a change in δ11B only where it is in contact with amphibole or pyroxene; sodalite δ11B varies where it is in contact with pyroxene or feldspar. The cathodoluminescence image (c) displays the different growth zones of sodalite. Whereas the B concentration decreases from the oldest zone to the youngest zone, the B isotopic composition is nearly constant throughout the grain. and B(OH)− 4 complexes, with a large relative mass difference (~28%) coexist. Apart from the mass difference, the larger size of the tetrahedral complex and the difference in charge should add to a significantly lower diffusivity of B(OH)− 4 compared to B(OH)3. Isotopic fractionation will enrich the lighter isotope 10B in the heavier and slower diffusing complex (B(OH)− 4 ). This leads to the remarkable prediction that the heavier B isotope will diffuse faster in near-neutral hydrous fluids, and kinetic fractionation will be opposite to that of Li. The magnitude of this fractionation, however, will be negligible, as the fractionation of the B isotopes between the two B complexes is in the permil range. Boron diffusion in silicate melt over a wide temperature range was studied by Chakraborty et al. (1993) who found that no isotopic fractionation occurred in their experiments, despite of strong concentration gradients. In conclusion, the intra-mineral B isotopic patterns are probably influenced by kinetic isotope fractionation, while diffusion of boron through melts or fluids is very unlikely to influence the B isotopic composition. Therefore, variations on the outcrop scale must have been caused by equilibrium fractionation processes and by mixing of B from different sources. 6.5. Boron isotope modelling The boron isotopic composition of a magma may be affected during differentiation. A number of processes are likely to influence the 11B/10B ratio as well as the B abundances of an evolving magma system, such as fractional crystallisation, contamination by country rock material or by fluids, and exsolution of fluid from the melt. In this M. Kaliwoda et al. / Lithos 125 (2011) 51–64 Amphibole Phase I Augite syenite ILM193 GM1857 GM1858 GM1303 P3–5 GM1214 GM1369 GM1370 GM1336 GM1337 GM1294 GM1843 JG10 JG12 JG13 − 19.1 −18.8 − 9.4 − 22.1 −22.1 − 24.3 −21.7 − 22.7 − 23.0 − 23.2 −19.8 − 16.5 − 32.7 −5.4 − 1.4 − 12.6 − 13.1 − 8.8 − 20.0 −9.8 −14.8 − 17.2 −17.2 −3.3 −17.1 − 13.8 −15.4 − 23.1 −0.7 +2.9 − 15.2 −16.9 − 9.1 − 20.7 −16.8 − 20.6 −20.0 − 19.9 −16.2 − 19.6 −16.8 − 16.0 − 28.2 − 2.9 + 1.4 9.2 6.5 0.6 2.3 9.5 6.2 4.9 8.5 22.4 6.4 8.4 1.6 6.9 4.7 4.9 Phase II Phase III Phase IV Alk. granite Pulaskite Sdl foyaite Naujaite Kakortokite Lujavrite Country rock Granite Julianehåb Granite Phase IIIa Phase IIIb lujavrite 2σ kakortokite Mean naujaite Max Sdl foyaite δ11B min pulaskite Sample augite syenite Rock type 20 whole rock amphibole Marks et al. (2007) δ7Li Magma pulse alk. granite Phase II increasing fractionation Table 2 Boron isotopic composition (‰ deviation from SRM951) of amphibole, feldspar and sodalite granite. Phase I 60 10 0 -9 amphibole Phase II Phase III Phase IV Augite syenite Alk. granite Pulaskite Sdl foyaite Naujaite Kakortokite Lujavrite Country rock Granite Sodalite Phase III Naujaite Phase IV Kakortokite Lujavrite ILM193 GM1857 GM1858 GM1303 P3-5 GM1214 GM1369 GM1336 GM1337 GM1294 GM1843 JG5 JG7 JG10 JG12 JG13 −11.0 − 8.8 − 11.0 − 19.3 − 19.8 − 20.9 − 20.4 −19.2 − 23.5 − 15.4 − 14.0 −12.2 − 4.8 − 13.5 − 7.7 − 1.8 +2.9 −6.8 +0.6 − 15.0 −15.9 −13.4 − 7.0 − 16.7 − 13.8 − 9.2 − 5.0 − 0.4 +4.4 −5.4 +5.6 + 7.4 − 6.2 − 7.8 − 5.8 − 17.5 − 17.4 − 17.1 − 15.0 −18.0 − 17.1 − 12.3 − 10.3 − 5.2 + 0.4 −10.5 − 1.6 + 3.5 9.9 1.9 9.8 4.5 4.3 10.6 12.3 3.5 11.2 8.8 9.3 9.6 8.0 12.6 11.1 6.0 GM1369 GM1370 GM1336 GM1337 GM1294 GM1843 − 17.7 − 15.0 − 13.6 − 13.1 −12.9 − 12.2 − 13.2 −11.7 −13.6 −10.3 −10.8 − 9.8 − 15.3 − 13.8 − 13.6 − 11.7 − 12.0 −11.2 4.5 3.3 0.0 2.8 2.1 2.3 paragraph, we employ Rayleigh fractionation models to evaluate the effects of these processes quantitatively. As discussed above, B isotopic fractionation between coexisting phases is primarily governed by their respective coordination of B to oxygen. The heavier isotope 11B is fractionated into the smaller, trigonally coordinated sites, while the larger, tetrahedral sites favour the light isotope 10B. The B isotopic fractionation between trigonal and tetrahedral sites can be calculated from the following formula 1000lna = 5:68–12290 = T ð1Þ determined by Hervig et al. (2002), where α is the isotope fractionation factor and T is the temperature in Kelvin. Similar temperature-dependent isotopic fractionations were found by other workers (e.g., Williams et al., 2001; Wunder et al., 2005) with the formulations being in very good agreement with Eq. (1) used here. Hence, an estimate on the coordination of B in the interacting phases (i.e., minerals, fluids, and melts) is essential for a first-order quantification of the B isotope fractionation caused by the interaction processes. Fractionation of early magmatic phases, such as olivine, pyroxene, amphibole and feldspar, led to the enrichment of incompatible elements in the more evolved magmas of the intrusion. These minerals generally show very low mineral/melt partition coefficients for B of 0 δ11B Feldspar Phase I feldspar -10 -20 -30 Fig. 7. Lithium (upper panel) and boron (lower panel) isotopic composition of minerals in the Ilímaussaq intrusion and the country-rock granite. δ11B values of amphibole and feldspar scatter around − 20‰ within the inner part of the intrusion, but increase to higher or even positive values within the granitic country rocks. δ7Li in amphibole and whole rocks (Marks et al., 2007) display a similar trend, i.e. constant values in the inner part of the intrusion and an increase towards the rim of the intrusion, which points to an infiltration of heavy B and Li into the margins of the Ilímaussaq intrusion. ~0.01 (Brenan et al., 1998b; Chaussidon and Jambon, 1994; Ryan and Langmuir, 1993; Tiepolo et al., 2004), and were, until recently, generally accepted to contain boron in tetrahedral coordination substituting for tetrahedral Al and Si (Grew, 2002; Hervig et al., 2002; Tonarini et al., 2003; Werding and Schreyer, 2002). The recent spectroscopic study by Hålenius et al. (2010) showed that B in clinopyroxene could be dominantly in trigonal coordination. The effects of crystal fractionation on the B isotopic composition of the melt were modelled with a constant partition coefficient DB(Am-melt) = 0.01 and two different endmember scenarios with B in trigonal and in tetrahedral coordination, respectively, in all minerals (with the exception of sodalite; see below). However, the B isotopic effect of mineral fractionation is negligible up to very large modes of fractionation (N80%) with such a low partition coefficient, and consequently, the same is true for the difference between the two end-member scenarios. Boron abundances in sodalite are much higher than in all other silicates in the rocks and it was argued that B is highly compatible in sodalite (Bailey, 2006; Bailey et al., 1981). The actual partition coefficient DB(Sdl-melt) can be estimated by combining amphibole–melt partition coefficients with the apparent sodalite–amphibole partition coefficients in samples GM1336 and GM1337 (kakortokites; phase IIIb). These kakortokites show petrographic evidence for contemporaneous crystallisation of sodalite and amphibole, with apparent partition coefficients DB(Am-Sdl) of 0.012 and 0.015 for GM1336 and GM1337, respectively, resulting in a best estimate for DB(Am-Sdl) =0.013±0.001. The partitioning of B between M. Kaliwoda et al. / Lithos 125 (2011) 51–64 (a) -10 sdl composition exsolution of a basic fluid 450 °C first D=5 15 last 10 D=2 15 -15 δ11B (‰) 5 initial melt fractionation of am, cpx, akf, ne, eud 10 5 30 40 5 50 last 25 fractionation of sdl (D = 1.63) fluid composition (D = 5) -20 last first fluid composition (D = 2) -25 0 100 200 first 300 400 500 B (µg/g) (b) -10 750 °C exsolution of a basic fluid sdl composition first D=5 15 -15 10 D=2 5 initial melt 5 δ11B (‰) amphibole and sodalite is thus in the same range as experimentally determined values for B partitioning between amphibole and silicate melt, i.e. DB(Am-melt) =0.004–0.022 (Brenan et al., 1998b; Tiepolo et al., 2004, 2007). The values for DB(Sdl-melt) deduced from these data range from 0.30 (i.e., slightly incompatible) to 1.63 (i.e., compatible). Consequently, it must be assumed that, if B is compatible in sodalite, it is definitely not highly compatible, but has a DB(Sdl-melt) value close to unity. Thus, sodalite fractionation will not strongly deplete the remaining melt in B, as previously assumed (Bailey, 2006; Bailey et al., 1981). A DB(Sdl-melt) value of 1.63 was employed in the model (Fig. 8). Bailey (2006) discussed boron coordination briefly. He argued that substitution of B for tetrahedral Al in the alumo-silicate framework cannot be the major mechanism for incorporation of B, as this would be operating in feldspar, nepheline and analcime in a similar way. These framework silicates, however, show one to two orders of magnitude lower B abundances compared to sodalite (Bailey, 2006, and our data). Alternatively, B may be incorporated in secondary anion complexes into the large cavities in the sodalite structure in replacement for Cl−. In a parallel study, we synthesised B-saturated sodalite in B2O3-rich hydrous fluids and analysed the run-products by IR spectroscopy. Sodalite with ~0.6 wt.% B2O3 showed significant B in trigonal coordination but only minor tetrahedral B (Kaliwoda, unpublished data). These results suggest that B substitutes for Cl− as B(OH)3 or B(OH)2O− in sodalite. The hypothesis of trigonal coordination of B in sodalite is consistent with the δ11B values of sodalite being on average 5‰ higher than those of coexisting amphibole in all investigated samples (if B in amphibole is tetrahedrally coordinated). Sodalite fractionation was modelled assuming exclusively trigonally coordinated B in this mineral (Fig. 8). Boron coordination in H2O at ambient P–T conditions is strongly dependent on pH. At pHN 9, B is predominantly tetrahedrally coordinated in B(OH)4– groups. In neutral and acidic (pHb/=7) hydrous fluids, it is trigonally coordinated in B(OH)3 units. Experiments have shown that this also holds true at high pressures and temperatures (Schmidt et al., 2005). Hydrous fluids were exsolved from the Ilímaussaq magmas at the naujaite, kakortokite and lujavrite stages, with the NaCl activity buffered by the mineral assemblage Sdl+Ab+Ne (Markl and Baumgartner, 2002). The pH of these fluids must have been between 8 and 10 independent of temperature (Markl and Baumgartner, 2002). Boron coordination in such highly alkaline fluids is almost exclusively tetrahedral. Thus, the model employs a BIV/(BIV +BIII) ratio of 1 for the fluid. Elemental partitioning of B between melt and hydrous fluid has been investigated in a few studies, consistently showing fractionation of B into the fluid with DB(fluid–melt) ranging from 1.2 to ~5.0 (Hervig et al., 2002; London et al., 1988; Pichavant, 1981; Schatz et al., 2004). The model presented here was calculated for two different values, i.e. DB(fluid–melt) = 2.0 and DB(fluid–melt) =5.0. Boron in boro-silicate melts varies between trigonally and tetrahedrally coordinated, depending on silica/B2O3 and alkali oxide/B2O3 ratios (Dell et al., 1983; Dingwell et al., 2002). For peraluminous silicate melts containing B at trace levels, Hervig et al. (2002) suggested that B occupies trigonally coordinated sites, arguing on the basis of experimentally observed fluid/melt B isotope fractionation. Spectroscopic studies by Tonarini et al. (2003) on natural rhyolitic glasses revealed that B at trace levels is predominantly in trigonal coordination, with only 8–26% of the B occupying tetrahedral sites. It has been argued that increasing alkali contents in silicate melts stabilise B in tetrahedral coordination, while high Al contents lead to a dominance of trigonally coordinated B (Dingwell et al., 2002). In the model presented below, we employed a BIV/(BIV + BIII) ratio of 0.5 for the melt. The initial B concentration of the melt is estimated to 100 μg/g, based on abundances of ~1 μg/g in amphibole combined with DB(Am-melt) = 0.01. The initial δ11B value is estimated to −17‰ based on the average δ11B value of amphibole in the early stage magmatic rocks (i.e., −20.3‰) combined with a Δ11B(Am-melt) = δ11BAm − δ11Bmelt = 3.3‰ (Eq. (1); BIV/(BIV + BIII)melt = 0.5). 61 last 30 40 25 last fractionation of sdl (D = 1.63) -20 -25 0 100 fractionation of am, cpx, akf, ne, eud 50 last fluid composition (D = 2) fluid composition (D = 5) first first 200 300 400 500 B (µg/g) Fig. 8. Results of B isotope modelling for two different temperatures of (a) 450 °C and (b) 750 °C. The initial composition of the parental melt ([B] = 100 μg/g; δ11B = −17‰) is marked by the star. The solid lines display the evolution of the melt in response to various fractionation processes, while the broken lines display the evolution of the fractionated sodalite, and of basic fluids with different B partition coefficients, as labelled. Fractionation of sodalite would drive the melt to lower δ11B values, while exsolution of basic fluids drive it towards a heavier composition. Fractionation of minerals with very low B compatibility has hardly any effect on the B isotopic composition of the melt. In the case of Ilímaussaq, both exsolution of basic fluids and sodalite fractionation occurred. Counter-balancing of the two processes may have kept the B isotopic composition relatively constant throughout magmatic evolution and explain the constant B isotopic composition throughout the intrusion (Fig. 7). The effect of temperature on the modelling results are relatively small. Mineral abbreviations are: am = amphibole, cpx = clinopyroxene, akf = alkali feldspar, ne = nepheline, eud = eudialyte, and sdl = sodalite. The effects of fractional crystallisation and of fluid exsolution were modelled using a Rayleigh formulation: D−1 c = ci = ð1−F Þ ð2Þ 11 11 a−1 δ B + 1000 = δ Bi + 1000 = ðc =ci Þ ð3Þ ci and c are the initial and final B concentrations of the melt, respectively, F is the mass fraction of crystals (or fluid) removed from the melt and δ11Bi and δ11B are the initial and final δ11B values of the 62 M. Kaliwoda et al. / Lithos 125 (2011) 51–64 melt, respectively, and D is the mineral–melt partition coefficient for B. In Fig. 8 the model detailed above is illustrated for temperatures of 450 °C and 750 °C. Starting at the star in the diagram (−17‰ at 100 μg/g B), fractionation of sodalite would create lower δ11B values (−20‰) in the subsequently crystallised rocks, while exsolution of a basic fluid would drive the crystallising rock towards more positive δ11B values (around − 12‰). The third possibility, a fractionation of amphibole, feldspar (pyroxene, nepheline and eudialyte) would increase the boron whole-rock content at constant δ11B (− 17‰), with the actual coordination of B in these minerals being insignificant. In our specific case, all three processes operated more or less simultaneously. Interestingly, the result obviously was that the initial B abundance and B isotopic composition of the Ilímaussaq magmas are very similar to the most differentiated portions, and hence that the starting point in the model is also the endpoint. The three possible reactions apparently compensated each other. Changing temperature does not alter this result significantly (Fig. 8). The lack of any significant increase in B abundances and the surprisingly constant B isotopic composition of the highly variable peralkaline intrusive complex of Ilímaussaq are in agreement with the Rayleigh fractionation model. This leads us to the conclusion that the very low δ11B value of −17‰ calculated for the Ilímaussaq magma is not a product of magmatic processes, but must represent the B isotopic composition of the parental magma(s). Magmas with δ11B values in the range of −17‰ are too 11B-depleted to be derived from primitive, depleted or subduction-zone-fluid-enriched mantle or from any typical continental crust: all these reservoirs range from approximately −12‰ to +5‰. The only known reservoir that is expected to introduce very light B into the upper mantle (the source region of the Ilímaussaq parental magmas) are high-pressure rocks subducted to mantle depths. Those rocks are typically driven to low δ11B values during low-temperature dehydration at the early stages of subduction (e.g., Marschall et al., 2007a; Peacock and Hervig, 1999; Rosner et al., 2003). 6.6. Entrance of heavy boron by meteoric fluids? The discussion above shows that processes such as crystal fractionation and fluid-exsolution cannot explain the isotopically heavy δ11B values for some of the marginal augite syenites and adjacent host granites. Similarly, kinetic boron isotope fractionation has no validity to account for this. We therefore suggest that these heavy δ11B values are caused by interaction with an external boron source. Based on O, H and Li isotopes, Marks et al. (2007) demonstrated that seawater (or possibly saline lake water) was entrained into the augite syenite zone and parts of the adjacent country rock granite. The very heavy Li of the augite syenite (δ7Li = + 14‰; Marks et al., 2007) requires the circulating fluid to have carried even heavier Li. Seawater, which has a very high δ7Li value is equally enriched in heavy B (modern seawater: δ11B = + 39.61‰; [B] = 4.4 μg/g; Foster et al., 2010; Spivack and Edmond, 1987) and would have, thus, also increased the δ11B value of the augite syenite and parts of the Julianehåb granite. The B isotopic composition of Proterozoic seawater is not well determined, but evidence suggests that it was also very heavy (Kasemann et al., 2010; Palmer and Slack, 1989). In contrast to the deep trough in δ7Li, no low-δ11B zone is expected, due to the lack of kinetic fractionation of B isotopes, as outlined above. 7. Summary and conclusions The concentration of B (as well as Li and Be) within the major rock types of the Ilímaussaq complex are mainly an effect of accumulation of minerals being compatible for the respective element. The major host for boron is sodalite, and whole rock concentrations are highest for sodalite-rich rocks. Consequently, the B whole-rock concentration does not allow for direct monitoring of magmatic differentiation processes. The intra-mineral B isotope patterns are interpreted to be influenced by kinetic isotope fractionation. On the other hand, relative differences in the diffusion of the two B isotopes through melts and fluids are probably insignificant and kinetic fractionation of B isotopes on the outcrop scale is hence considered unlikely. Equilibrium fractionation of sodalite would drive the δ11B of the remaining melt to lower values (by ~3‰ in our model). In contrast, exsolution of basic hydrous fluids would drive the δ11B values in the degassed melt to more positive values (by 5‰ in the modelled case). The fractionation of silicates with very low mineral–melt partition coefficients for B, such as amphibole, pyroxene, feldspar, nepheline and eudialyte, would have no significant effect on δ11B of the magma. In the case of Ilímaussaq, B concentrations and δ11B values of magmatic amphiboles throughout the three phases remain relatively constant at approximately 1 μg/g and − 20‰, respectively. This demonstrates that the effects of sodalite fractionation, fluid exsolution and low-B silicate fractionation on B abundances and B isotopic composition of the magmas compensated each other. The high δ11B values in the margin of the intrusive complex, i.e. the augite syenites and the adjacent country-rock granite is most probably influenced by entrained seawater, enriched in heavy boron (higher δ11B). Acknowledgements This work was financed by the DFG (Deutsche Forschungsgemeinschaft), MA 2135/8-1 and AL 166/18-1, which is appreciated. The authors are grateful to Thomas Wenzel and Julian Schilling for assistance with the microprobe analyses. We also thank Gesa Graser, Jasmin Köhler, Thomas Krumrei, Johannes Schönenberger, Thomas Wagner and Thomas Wenzel for fruitful discussions. 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