Download Boron and boron isotope systematics in the peralkaline Ilímaussaq

Lithos 125 (2011) 51–64
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Lithos
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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. Ilona Fin and
Oliver Wienand are thanked for preparing the thin sections for the
SIMS and microprobe measurements. Katalin Gméling and Zsolt
Kasztovszky, Budapest, are thanked for the PGNA analyses and Sigrid
Hirth-Walther, Freiburg, is thanked for the AAS analyses. Comments
by Ed Grew and Henning Sørensen on an earlier version of the
manuscript are highly appreciated. The paper gained from careful
reviews by Ian Coulson and an anonymous reviewer, and from
editorial handling by Nelson Eby, which is gratefully acknowledged.
Appendix A. Supplementary data
Supplementary data to this article can be found online at
doi:10.1016/j.lithos.2011.01.006.
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