Lithium isotope fractionation during magma degassing: Constraints from silicicdifferentiates and natural gas condensates from Piton de la Fournaisevolcano (Réunion Island)
I. Vlastélic a,b,c,⁎, T. Staudacher d, P. Bachèlery e, P. Télouk f, D. Neuville g, M. Benbakkar a,b,c
a Clermont Université, Université Blaise Pascal, Laboratoire Magmas et Volcans, BP 10448, F-63000 Clermont-Ferrand, Franceb CNRS, UMR 6524, LMV, F-63038 Clermont-Ferrand, Francec IRD, R 163, LMV, F-63038 Clermont-Ferrand, Franced Observatoire Volcanologique du Piton de la Fournaise, Institut de Physique du Globe de Paris, CNRS UMR 7154, 14 RN3, le 27èmekm, 97418, La Plaine des Cafres, La Réunion, Francee Laboratoire GéoSciences Réunion, Université de La Réunion, Institut de Physique du Globe de Paris, CNRS UMR 7154, 15 Avenue René Cassin, 97715 Saint-Denis cedex 09,La Réunion, France`f Laboratoire des Sciences de la Terre, Ecole Normale Supérieure de Lyon, CNRS UMR 5570 46 Allée d'Italie, 69364 Lyon cedex 07, Franceg Laboratoire de Géochimie et Cosmochimie, Université Paris Diderot, Institut de Physique du Globe de Paris, CNRS UMR 7154, 1 rue Jussieu, 75238 Paris cedex 05, France
a b s t r a c ta r t i c l e i n f o
Article history:Received 30 August 2010Received in revised form 31 January 2011Accepted 2 February 2011Available online 5 March 2011
Edited by: R.L. Rudnick
Keywords:Lithium isotopesIsotopic fractionationMagma degassingPiton de la Fournaise volcano, Réunion Island
Recent volcanic products from the Piton de la Fournaise Volcano, Reunion, show pronounced depletion orenrichment in lithium and significant isotopic fractionation related to degassing. (1) trachytic pumices fromthe April 2007 eruption show extreme Li depletion (90%) and isotopic fractionation (δ7Li of −21‰). Thedepletion of water and volatiles (Cl, F, B, Cs) in these samples suggests that Li loss occurred in response todegassing, which most likely occurred as the small, isolated volume of magma underwent extensivedifferentiation near the surface. Because the pre-degassing composition is relatively well known, thecomposition of the degassed pumice constrains the partition coefficient to 60bDV–Mb135 and the isotopicfractionation factor, αV–M, to 1.010 at magmatic temperatures. Unlike DV–M, αV–M does not depend onwhether crystallization and degassing occurred successively or concomitantly. (2) basaltic samples from theinterior wall of the long-lived 1998 Piton Kapor were extensively altered by acidic gas. They also showextreme Li depletion, but barely significant isotopic fractionation (δ7Li=+4.5‰), suggesting that high-temperature leaching of Li by volcanic gas does not significantly fractionate Li isotopes. (3) high-temperature(400–325 °C) gas condensates formed during degassing of the thick lava flow of April 2007 display high Licontents (50–100 ppm), which are consistent with Li being as volatile as Zn and Sn. Their isotopically light Lisignature (average of −1.7‰) is consistent with their derivation from isotopically heavy vapor (+13.5‰) ifthe factor of isotopic fractionation between condensate and vapor is less than 0.985. A degassing-crystallization model accounts for the evolution of trace species, which, like lithium, are volatile but alsomoderately incompatible.
Alkali metals are only moderately volatile, in the sense that only asmall proportion (b0.1%) of the initial budget of mantle derived meltsis ultimately lost by magmas (Rubin, 1997). On the other hand, alkaliare largely mobilized during exsolution of H2O-rich vapor phase inmagmatic systems (Sakuyama and Kushiro, 1979). Selective vaportransport of potassium has been suggested during the late stage ofcrystallization of oceanic magmas (Sinton and Byerly, 1980), and frombottom to top of lava lakes (Richter and Moore, 1966) although the
Kilauea data on which this idea is based were subsequentlyquestioned (Helz et al., 1994). There has been recently a growinginterest in lithium, the lightest alkali metal, but above all one of thefastest diffusing elements in silicates (Jambon and Semet, 1978;Giletti and Shanahan, 1997; Richter et al., 2003). Indeed, lithiumseems to be particularly sensitive to vapor transfer, either in shallowmagma conduit shortly before eruption (Berlo et al., 2004; Kent et al.,2007) or during post-eruptive degassing of thick lava flows (Kuritaniand Nakamura, 2006).
While the use of lithium abundance as tracer of degassingprocesses is still in its infancy, new perspectives arise from lithiumstable isotopes (6Li and 7Li) geochemistry. The main reason is that ourknowledge of the laws governing lithium isotopes fractionation innature has improved significantly in recent years. First, it is now wellestablished that heavy lithium partitions preferentially into aqueous
Chemical Geology 284 (2011) 26–34
⁎ Corresponding author at: Laboratoire Magmas et Volcans, Observatoire dePhysique du Globe de Clermont-Ferrand, UMR 6524, 5 Rue Kessler, 63038 Clermont-Ferrand, France. Tel.: +33 4 73 34 67 10; fax: +33 4 73 34 67 44.
fluids over silicate rocks, and that the magnitude of this phenomenondecreases with increasing temperature (e.g., Wunder et al., 2007;Millot et al., 2010). Second, as 6Li diffuses faster than 7Li, large isotopicfractionation arises during diffusion (e.g., Richter et al., 2003). Thesetwo types of isotopic fractionation, which, for convenience, will besimply referred to as “chemical” and “kinetic”, respectively, may occurduring magma degassing (Beck et al., 2004; Rowe et al., 2008; Schiaviet al., 2010), although little is yet known on this topic.
Taking advantage of the recent intense volcanic activity of Piton dela Fournaise (Réunion Island), a Li isotopic study of late-stagemagmatic processes was undertaken. This paper focuses on exten-sively devolatilized samples (silicic differentiates) and gas conden-sates (natural sublimates), which provide new constraints on thebehavior of Li isotopes during magma degassing.
2. Geological setting and samples
Piton de la Fournaise (Réunion Island, Indian Ocean) in anintraplate shield volcano that has erupted on average once a yearsince 1930, and seemingly more frequently since 1998 (see a reviewof the most recent activity in Peltier et al. (2009)). The volcanoproduces dominantly transitional basalts with little compositionalvariability (these basalts are referred to as Steady State Basalts (SSB)following Albarède et al., 1997) and, every 15–30 years and morefrequently since 2001, olivine-rich basalts (the most magnesian beingcommonly referred to as “oceanites”). Geophysical data support theexistence of a shallow magma reservoir near sea level (2.5 km depth)and a deeper reservoir at the crust-mantle interface (7.5 km depth)
(Peltier et al., 2008). The volume of the storage system, in the range of0.1–0.35 km3, and magma residence time of 15–30 years wereestimated based on geochemical data (Albarède, 1993; Sigmarssonet al., 2005; Vlastélic et al., 2009a).
Among the recent products of Piton de la Fournaise, this studyfocused on samples that underwent pronounced depletion orenrichment in lithium in relation with degassing processes (Fig. 1).They include trachytic pumices recovered near the main vent of April2007 eruption. The pumices are almost completely glassy, highlyvesicular (80–90%) and often coated by basaltic glass. These silicicdifferentiates, the only known at Piton de la Fournaise, are thought tooriginate from small batches of magmas trapped at shallow depthwithin the volcanic edifice (Bachèlery, in prep.). Comparing the Pbisotopic signature of the pumices to the detailed Pb isotope temporalrecord (Vlastélic et al., 2009a) suggests derivation from liquidserupted between 1977 and 1986. Theses pockets of differentiatedliquids may have been disrupted and entrained during the paroxys-mal phase of the eruption, on April 5th (see review of the eruption byStaudacher et al. (2009)). To confirm this hypothesis we also analyzedthe olivine-bearing glass coating the pumices. The second type ofsample is gas-altered rocks from the interior wall of the 1998 Kaporcrater. These rocks, which have been extensively leached by acidic gasduring the unusually long 1998 eruption (nearly 6 months), havebeen selected to evaluate how Li isotopes behave during gas-rockinteraction. The last type of sample is gas condensates that formedduring degassing of the April 2007 lava flow. Post-eruptive degassingprocesses, which include formation of segregation veins and gas-filterpressing (Martin and Sigmarsson, 2007), have been shown to
Fig. 1. Map of Piton de la Fournaise volcano showing the eruptions of March 1998 (Piton Kapor) and April 2007 (Piton Tremblet). Locations and photos of the studied samples areindicated by numbers: (1) gas altered basalt from the interior wall of the 1998 Kapor crater. (2) trachytic pumice and its basaltic coating (recovered in May 2007 near PitonTremblet) were probably emitted during the paroxysmal phase of April 2007 eruption (Bachèlery et al., in prep.) (3) gas condensates from a fumarole that was active between 2007and 2009 in the region where April 2007 lava flow is the thickest (ca. 50 m).
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efficiently mobilize lithium (Kuritani and Nakamura, 2006). SincePiton de la Fournaise volcano lacks long-lived fumaroles, thevoluminous (130 Mm3) and slowly cooling lava flow of April 2007(Staudacher et al., 2009) offers a rare possibility to sample magmaticgas, in particular in its thickest part where an active fumarole wasdiscovered. Natural gas condensates were sampled four times at thissite between August 2008 and November 2009 as the venttemperature decreased from ~400 to ~300 °C. In addition, a largenumber of steady-state basalts (n=31) and olivine-rich basalts(n=30) erupted between 1927 and 2007 were also analyzed forcomparison.
3. Analytical methods
The bulk major element chemistry was determined by ICP-OES(rocks) and ion chromatography (condensates). Scanning electronmicroscopy (SEM) was used to investigate the compositionalheterogeneity of the altered basalt and to determine the mineralogyof gas condensates. Water content was determined in the trachyticpumice by Raman spectroscopy (IPGP, Paris) using a recentlyimproved calibration curve (Le Losq et al., in press). The Ramanspectra were recorded using a T64000 Jobin-Yvon spectrometer, withan excitation line at 514.532 nm. Acquisition times and laser powerwere set to 900s and 20 mW, respectively. Data treatment andaccuracy of the method are detailed in Le Losq et al. (in press). Whole-rock analysis of Cl, F and B concentrations has been assigned to theService d'Analyse des Roches et Minéraux (CRPG-Nancy). Chlorinecontent was determined by a spectrophotometric method based onthe formation of ferrithiocyanate, while F content was determined bypotentiometry using an ion-selective electrode. The precision of F andCl determination is 10–20% (2σ) for the concentration range (100–500 ppm) relevant to our study (Vernet et al., 1987). Followingsodium carbonate fusion, B was purified on ion-exchange resin and itsconcentration was determined by the colorimetric method usingCarmin with a precision of ca. 25%. Other trace elements and Liisotopic compositions were analyzed on same dissolutions. Between50 and 100 mg of sample were dissolved in HF-HNO3 (rocks) ordiluted HNO3 (gas condensates). Once digested, solutions wereevaporated to near dryness and re-dissolved in 7M HNO3. Thedissolved samples were then split into two aliquots. The first wasintended for determination of trace element abundance by quadru-pole ICPMS (Agilent 7500, Laboratoire Magmas et Volcans). To thisend, the aliquots were evaporated to near dryness and subsequentlydiluted in HNO3 0.4 M to reach a total dilution factor ranging from5000 (rocks) to 20,000 (gas condensates). The reaction cell (Hemode)was used to reduce interferences on masses ranging from 45 (Sc) to75 (As). The signal was calibrated externally with a reference basalticstandard (BHVO-2) dissolved as samples (rocks) or with a syntheticstandard (gas condensates). One of these two standards and pureHNO3 0.4 M were measured every 4 samples. The external reproduc-ibility of the method, as estimated by running repeatedly differentstandards (BCR-2, BIR, BEN) is b5% for most lithophile elements andb15% for most chalcophile and siderophile elements. The secondaliquot of dissolved sample was used for measurement of lithiumisotopic composition. Following evaporation and conversion of ions tochloride form (with 1 ml of HCl 6 N), samples were successfully fullydissolved in 1 ml of HCl 0.5 N, and Li was purified on AG50W-X8cation exchange resin following the two-step method described inVlastélic et al. (2009b). The lithium chemistry blank measured byquadrupole ICP-MS is less than 0.1 ng. Given the amount of lithiumrequired for isotopic analysis (N50 ng) and the loading limit of thefirst-step column (the equivalent of 50 mg of rock), the Li depletedpumices were processed through two columns and Li fractions weremerged before further purification on the second-step column. Lithiumisotopic compositions were measured by MC-ICPMS (Nu 500) at theEcole Normale Supérieure de Lyon. Samples were introduced through a
desolvator (Nu DSN) at a rate of 100 μl/min, yielding a total Li beam of4 to 6 V for 70 ng/g Li solutions. Standard operating conditions wereused (RF power of 1350W, Ar cool gas flow of 13 L/min, Ar auxiliary gasflow of 1 L/min, Ar sample gas flow of 0.8 L/min, and accelerationvoltage of 4000 volts). Lithium isotopes were measured in static modein L5 and H6 faraday cups. Measurements (60 cycles of 10s) wereperformed following an uptake-stabilization time of 90s. The washoutprocedure (300s with 0.65 N HNO3 and 300s with 0.05 N HNO3)reduced Li signal by a factor of 104. Mass fractionation was monitoredexternally with IRMM-016 standard using a sample-standard bracket-ing technique. Depending on the amount of lithium extracted, multiplemeasurements of each sample were performed. Repeated analysisof the USGS BHVO-2 standard (batch 759) over 4 years yielded δ7Li=+4.2‰±0.5 (2σ, 17 dissolutions). Repeated measurement (n=12) ofa L-SVEC solution (not processed through separation columns) duringa single session yielded δ7Li=−0.20‰ ±0.04 (S.E.), raising the pos-sibility that L-SVEC is isotopically slightly lighter than IRMM-016, aspreviously proposed (Millot et al., 2004).
4. Results
Lithium isotopic compositions and trace element concentrations ofrocks and gas condensates are reported in Table 1, together with theaverage composition of steady-state basalts and oceanites, the mostcommon types of basalts erupted at Piton de la Fournaise (a completelithium data set is provided in Supplementary Table 1). Major elementchemistry of the studied samples is reported in SupplementaryTable 2.
4.1. Silicic differentiates
The pumice samples 0704-PS and 0704-PN have total alkali andSiO2 contents (ca. 11 wt.% and 62 wt.%, respectively) that plot withinthe field of trachytes (Bachelery, in prep.). Loss-on-ignition at 1000 °Cis barely significant, indicating that water content in these differen-tiated samples is probably less than 0.5 wt.% (i.e., the maximumweight gain due to oxidation of iron). Consistently, precise determi-nation of water content by Raman spectroscopy yielded 0.2±0.1 wt.%.By comparison to steady-state basalts, the 2007 trachytic pumices areenriched inmost incompatible trace elements (Fig. 2a,b). Thedegree ofenrichment is the highest for Nb and Ta (ca. 7) and decreasesprogressively from light REE (ca. 5) to heavy REE (ca. 2). The increaseof Zr and Hf concentration with digestion strength (Table 1) suggeststhe occurrence of zircons. There are also a number of negativeanomalies,most of them (Eu, Sr, Ba, and Pb) being consistentwith low-pressure fractionation of plagioclase feldspar. Lithium, B and Cs,together with Cl and F (not reported on the plots), all beingincompatible in plagioclase, show the most pronounced anomalies,however. The lithium depletion is accompanied by strong isotopicfractionation, the two pumice samples analyzed showing extremelylight lithium compositions (−21bδ7Lib−17‰) when compared tohistorical lavas of Piton de la Fournaise (+2.7bδ7Lib+4.9‰), andterrestrial rocks in general. The olivine-bearing glass (0704-BAS)coating the pumices has trace element pattern (Fig. 2c) and Liisotopic signature (δ7Li of +4.2‰) indistinguishable from typicaloceanites.
4.2. Gas-altered basalt
Whole rock analysis (ICP-OES) of the altered basalt 9809–301revealed that it is essentially made of SiO2 (58 wt.%), CaO (11 wt.%)and a volatile component lost during sample fusing. Compared tofresh basalts, this specimen shows a marked enrichment in silica anddepletion in Al2O3, Fe2O3, and MgO, which have concentrations thatdo not exceed 2 wt.%. Inspection of the sample by SEM revealed theoccurrence of a silica-rich phase (SiO2 up to 87%vol.%) and anhydrite
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(CaSO4). Silicification in fumarolic environment was shown to be theresult of extensive leaching of major cations by acidic gas condensates(Africano and Bernard, 2000). Replicate analysis of this sample pointsto heterogeneous distribution of trace elements, the two extremepatterns being reported in Fig. 3. The alteration resulted in no loss in
Ba, Ta and Nb, small to significant (up to 50% in sample 9809-301a)loss of rare earth elements, Th, Pb, Zr, Hf and Y, and major to extreme(N90%) loss of Cs, Rb, U, Be and Li. In contrast with the pumice, lithiumloss is accompanied here by barely significant isotopic fractionation(+4.1bδ7Lib+5.1‰).
Table 1Trace element concentrations and Li isotopic compositions of rocks and gas condensates.
Main rock types Devolatilized trachytic pumice Gas altered basalt Gas condensates
All concentration in g/g (ppm).n.d.: not determined.a SSB: Steady State Basalts. Mean composition inferred from 1931 to 2007 lavas with MgOb9%wt.b Océanites: Lavas rich in cumulative olivine. Mean composition inferred from 1931 to 2007 lavas with MgON20%wt.c PARR bomb dissolution.d A complete lithium data set is provided in Supplementary Table 1. Source of trace element data: Vlastélic et al. (2005, 2007) and unpublished data.e δ7Li=((7Li /6Li)sample / (7Li/6Li)IRMM−1)⁎1000.f number of sample analyzed.g Repeated analysis of the same dissolution.
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4.3. Gas condensate
Analysis of the gas condensates by SEM identified K–Na sulfateswith K/Na suggesting the main occurrence of aphthitalite ((K,Na)3Na(SO4)2). The distribution of trace elements in gas condensates isshown on Fig. 4 where the elements are sorted according to theirenrichment relative to lavas. Enrichment factor (EF) is defined as:
EFX = X=XRð Þgas condensate = X=XRð Þlava ð1Þ
where X is the element of interest and XR the refractory element ofreference. Beryllium was chosen here because of its low volatility andlow abundance in lavas (Moune et al., 2006). The gas condensates aredepleted (EFb1) in rare earth elements and high field strengthelements and enriched (EFN1) in alkali and chalcophile elements,which pattern is consistent with element volatility at mafic volcanoes(Toutain et al., 1990; Rubin, 1997). Lithium (5bEFb6), like Zn, Sn, Rb,Sb, and As, displays amoderately volatile behavior during degassing ofthe April 2007 lava flow. The condensates have light and ratheruniform Li isotopic signatures (−2.8bδ7Lib−1.3‰), the lowest δ7Li(−2.8‰) being measured in the sample (08.10.281) with the lowestLi content (52.7 ppm, against N70 ppm in other samples).
5. Discussion
5.1. Origin of lithium depletion and isotopic fractionation in silicicdifferentiates
The trachytic pumices display very low lithium contents andextremely light Li isotope compositions that have not yet beenreported in fresh volcanic rocks. Such a combination rules out diffusiveinflux of lithium,which best accounts for very low δ7Li signatures in Li-rich rocks only (Rudnick and Ionov, 2007; Marschall et al., 2007).Hydrothermal alteration,which preferentially removes 7Li from solids,could yield such light signatures, but only in low-temperatureenvironments (Wunder et al., 2007; Millot et al., 2010), which isinconsistentwith themagmatic origin of pumices. Independently, the
TaLa
CeBe
PrPb
NaNd
SrSm
ZrHf
EuGd
TbDy Y
Ho ErTm
LiYb
Lu
Pb Sr Eu Li
0704-PS
0704-PN
0704-BAS
0.1
1
10
0.01
0.1
1
10
0.1
1
10
CsRb
BaTh
UK
Nb
Con
cent
ratio
n / S
SB
Con
cent
ratio
n / S
SB
Con
cent
ratio
n / O
cean
ite
Cs Ba Ua
b
c
B
B
Fig. 2. Trace element patterns of the April 2007 trachytic pumice (a,b) and its basalticcoating (c). Concentrations are normalized to the average concentrations of steady-state basalts (a,b) and oceanites (c), both being given in Table 1. Vertical lines are forelements showing negative anomalies.
Con
cent
ratio
n / S
SB
0.01
0.
1
1
10
9809-301a
Con
cent
ratio
n / S
SB
0.01
1
0.1
10
CsRb
BaTh
UNb
TaLa
CeBe
PrPb
NdSr
SmZr
HfEu
GdTb
Dy YHo Er
TmLi
YbLu
9809-301b
a
b
Fig. 3. Trace element patterns of the gas-altered basalts from the 1998 Piton Kapor.Concentrations are normalized to the average concentrations of steady-state basaltsgiven in Table 1. Vertical lines indicate elements that underwent the largest depletion.Amongst the three samples analyzed (see Table 1), the two extreme patterns are shownhere.
0.001
0.01
0.1
1
10
100
1 000
10 000
Ce
Pr
Er
Ho
Y
Th
Dy
Tb
Sm
Eu
Gd
La
Nd
Yb
Tm
Lu
Sr
Nb
Ba
Ti
Sc
Ga
Hf
Zr
Co
Cr
UNi
V
Be
ZnSn
Rb
Sb
As
Cs
Cu
In
Pb
W
Bi
CdTl
Enr
ichm
ent F
acto
r
Li
Gas condensates
Fig. 4. Enrichment factors of trace elements in 2008–2009 gas condensates. Enrichmentfactors are calculated according to Eq. (1), using beryllium as reference. Normalizationis made to the average composition of steady-state basalts (see Table 1). Elements aresorted according to their mean enrichment factor. For each element, the range ofvariation of the enrichment factor within the five samples analyzed is shown.
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mantle-like oxygen isotope signature of pumices (Bachèlery et al.,in prep.) does not support input of meteoric water. Fluxing byvapors released by melt degassing deeper in the plumbing system(Blundy et al., 2010) is another possibility, which could alsoexplain the less dramatic loss of uranium, which is fluid-mobile whenoxidized but non-volatile. Although this possibility cannot be definitelyruled out, our attempt to test it suggests that high-temperature leachingof lithium by magmatic gas does not significantly fractionate 7Li/6Li. Ontheother hand, the concomitant loss of B, Cs, Cl andF suggests that Liwaslost during devolatilization of the magma, as it underwent extensivecrystallization at shallow depth. Indeed, the water content of pumicesamples (0.2 wt.%) is unexpectingly low given the largemass fraction ofanhydrous crystals removed. Both major and trace elements areconsistent with a crystallizing mineral assemblage made dominantlyof clinopyroxene, plagioclase, olivine, Fe–Ti oxides, apatite, with noevidence for hydrous minerals such as amphibole or phlogopite.Elements (e.g., Th, Nb) that are highly incompatible in these mineralsare strongly enriched (a factor of 6.2 on average) in the trachytic sampleswith respect to steady-state basalts. Assuming these elements do notenter the crystallizing phases (bulk DS–M ~0), the fraction ofmelt havingcrystallized during the basalt–trachyte differentiation is estimated at ca.84%. From the water content of primitive melt (0.7–1.0%wt% accordingto Bureau et al. (1999)), and assuming that H2O is as incompatible asCe (Michael, 1995), awater content of 2.9–4.1 wt.%wouldbeexpected inthe trachytic melt if it evolved in a closed system. The residual watercontent of pumices (0.2±0.1 wt.%) suggests that between 2.6 and4.0 wt.% of water left the system and entrained a major fraction of theinitial budget of lithium.
5.2. Modeling lithium loss and isotopic fractionation during coupleddegassing and crystallization
To explain the composition of the dehydrated silicic differentiates,we resorted to the open system degassing-crystallization modelpreviously used for chlorine (Villemant and Boudon, 1999; Villemantet al., 2008). However, the model has to be modified to take intoaccount that, unlike chlorine, lithium significantly partitions intocrystallizing phases. The mass balance equations in the system madeof melt (M), crystals (S) and vapor (V) are:
dmS + dmV = −dmM ð2Þ
and, for lithium,
CS:dmS + CV :dmV = −d CM :mMð Þ ð3Þ
where C is the concentration of lithium. Following Villemant andBoudon (1999), we assume that the masses of crystallizing melt andexsolving vapor are proportional:
dmS = k:dmV ð4Þ
and, from Eq. (2), dmV=−dmM/(1+k) and dmS=−dmM/(1+1/k).Substituting in Eq. (3) yields:
dCMCM
=dmM
mMβ−1ð Þ with β =
DV−M + kDS−M
k + 1ð5Þ
where DV–M and DS–M are the vapor–melt and crystal–melt partitioncoefficients of lithium, respectively. We then set dmM/mM=dF/F, Fbeing the mass fraction of residual melt expressed as:
F = 1− 1 + kð ÞΔH20 ð6Þ
where ΔH20 is the mass fraction of vapor lost. Integrating Eq. (5) yieldsthe expression of lithium concentration in the residual melt (CM) as afunction of the initial melt content (C0M):
CM = CM0 :F β−1ð Þ ð7Þ
Note that if DS–M~0, β=DV–M/(1+k) as inferred for chlorine(Villemant et al., 2008). Considering lithium isotopes as individualspecies, Eq. (7) may also be used to compute the evolution of lithiumisotopic ratio of the residual melt:
7Li6Li
!M
=7Li6Li
!M
0
:Fk DS−M
7 −DS−M6ð Þ + DV−M
7 −DV−M6
1 + k
� �ð8Þ
where (7Li /7Li)0 is the initial isotopic ratio, and subscripts 7 and 6denote for 7Li and 6Li species. If lithium isotopes partition equally intothe crystallizing phases (D7
S–M=D6S–M):
7Li6Li
!M
=7Li6Li
!M
0
:FDV−M6
1 + k: αV−M−1ð Þ� �
ð9Þ
with
αV−M =7Li=6Li� �V7Li=6Li� �M =
M6 =M7ð Þ CV7 = CV
6
� �M6 =M7ð Þ CM
7 = CM6
� � =DV−M7
DV−M6
and
DV−M6 = DV−M
:1 + M7 =M6ð Þ 7Li=6Li
� �M1 + αV−M M7 =M6ð Þ 7Li=6Li
� �M0B@
1CA
where αV–M is the coefficient of isotopic fractionation between vaporand melt and M the molar mass. When there is only degassing (k=0)and DV–MNN1, it is convenient to use the simplified Rayleigh equationdescribing the isotopic evolution as a function of C/C0 and α:
7Li6Li
!M
=7Li6Li
!M
0
:F DV−M−1ð Þ αV−M−1ð Þ =7Li6Li
!M
0
:CM
CM0
! αV−M−1ð Þð10Þ
We stress that Eq. (10) only applies when elements stronglypartition into the vapor phase, and is thus not valid when elementsenter the crystalline assemblage significantly. The evolution of Licontent and Li isotopic composition of the residual melt was modeledconsidering that crystallization and degassing occurred eithersuccessively (case A) or simultaneously (case B). In both cases, thestarting composition is assumed to be that of steady-state basalts(Li=5.85 ppm, δ7Li=+3.6‰, see Table 1). It is also assumed that Liisotopes do not fractionate significantly during crystallization(Tomascak et al., 1999; Schuessler et al., 2009). In case A, Li is assumedto behave like heavy rare earth elements during low-pressurecrystallization (Ryan and Langmuir, 1987). Its concentration in theundegassed trachyticmelt is thus calculatedby considering smooth SSB-normalized trace element pattern (Li=LiSSB.[Tm/TmSSB+Yb/YbSSB]/2=11.8 ppm), which, in turn, implies a solid-melt partition coefficientof ca. 0.6. As discussed above, the differentiated melt must have lostbetween 2.6 and 4.0 wt.% of water during degassing. In case B, the initialwater content is assumed to be that of primitivemelts (~1 wt.%). Thus, kon the order of 100 is inferred from the fraction of vapor lost (~0.8 wt.%)and the fraction of melt having crystallized (~84 wt.%). Lithium lossduring degassing was modeled using Eq. (7) (Fig. 5), setting k=0 (nocrystallization) in caseA. The low lithiumcontent of thepumice requires60bDV–Mb95 in case A and β~1.9 in case B (DV–M=135 for DS–M=0.6
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and k=100). Note that if β=1 (DV–M=k(1−DS–M)+1), thecompeting effects of crystallization and degassing cancel out andthe Li content of the melt remains constant. Taking 0.6 as an upperbound for DS–M, it can be estimated that lithium content in meltdecreases only if DV–MN~40. The values of DV–M required to explainLi depletion in pumice are larger than those estimated for thepartitioning of lithium between fluids and melts, although trulyrelevant data lack in the literature. For granitic–rhyolitic composi-tions, DV–M was shown to increase from 0 to 13 with decreasingpressure from 400 MPa to 50 MPa (London et al., 1988; Webster etal., 1989), while 1bDV–Mb30 was inferred by mass balancefollowing degassing experiments (Koga et al., 2008). The elevatedDV–M value inferred in this study could be a consequence of thebasalt–trachyte differentiation process, which generally involveshigh concentrations of water (Ringwood, 1959). However, thisdoes not explain the scarcity of the Li signature of the 2007trachyte. The most comparable case, even if Li loss (~20%) is lessextreme, is that of post-eruptive internal differentiation of thethick lava flow of Rishiri Volcano (Japan) (Kuritani and Nakamura,2006). Based on these observations, it is suggested that DV–M couldbe very elevated during low-pressure, vapor-saturated crystalliza-tion, in particular if there is a long interaction (possibly 20 years forthe 2007 trachyte) between exsolved vapor and melt. Anotherpossibility is that fluxing by vapors released by deeper meltsaccompanied degassing (Blundy et al., 2010) and enhanced lithium
entrainment into the vapor phase. In this scenario, our calculatedDV–M values are overestimated.
Fractionation of Li isotopes during degassing was modeledusing Eq. (10) in case A (DV–M≫1) and Eq. (9) in case B assuming(D6)V–M~DV–M (since αV–M ~1) (Fig. 6). The end-member pumicecomposition (Li=1 ppm and δ7Li=−21‰) is consistent with αV–M of1.0100 (case A) and 1.0099 (case B). The similarity ofαV–M in both casesis at first glance surprising given the difference in the startingcomposition. It is consistent with αV–M being dependent only on theamount of lithium lost, which is the same in cases A and B. These valuesof αV–M are unexpectingly large (isotopic fractionation of ~10‰) for ahigh-temperature process. Large fractionations (1.005bαV–Mb1.015)during magma degassing were also proposed to explain Li isotopesystematics in 2003 ashes from Stromboli, suggesting that degassingfractionates lithium isotopesmore efficiently than fluid-rock interaction(Schiavi et al., 2010).
5.3. Lithium isotopic composition of volcanic vapor
The isotopic composition of Piton de la Fournaise volcanic gas hasnot been measured, but can reasonably be inferred from αV–M
calculated for the dehydrated trachytic melt. The instantaneouscomposition of the vapor is given by:
7Li6Li
!V
= αV−M7Li6Li
!M
ð11Þ
bcrystallization + degassing
0
2
4
6
8
0 20 40 60 80
0 20 40 60 800 1 2 3 4
Li in pumice
0
2
4
6
8
10
12
14
Li (
pp
m)
Li (
pp
m)
crystallizationDs-m=0.62
degassing
Fcrystal (%)
Fcrystal (%)
Fvapor (%)
Li in pumic
a
e
Dv-m=60
Dv-m=95
β=1.9
β=1.0
β=0.8
β = Dv-m + k.Ds-m
1 + k
Fig. 5. Modeling Li content in residual melt during crystallization and degassing. Twocases are considered: (A) crystallization occurs first and degassing occurs subsequently,and (B) crystallization and degassing occur concomitantly. In both cases, the average Licontent of steady-state basalts (5.85 ppm) is used for the starting composition. Thetarget composition is that of the trachytic pumice (1 ppm Li). In case A, the Li content ofmelt (11.8 ppm) after differentiation and before degassing, is estimated assuming thatLi displays no anomaly along trace element patterns when it is positioned between Tmand Yb. The fraction of melt having crystallized (Fcrystal) is inferred from the enrichmentof the most incompatible elements, while the fraction of vapor lost (Fvapor) isconstrained from the initial (assumed) and final (measured) water content. DS–M andDV–M are the crystal-melt and vapor-melt partition coefficients, respectively. FollowingVillemant and Boudon (1999), k is the mass ratio of crystal removed to vapor exsolved.See text for details.
δδ7L
i (m
elt)
δ7L
i vap
or)
-15
-10
-5
0
5
10
15
0 2 4 6 8 10 12 14
Li in melt (ppm)
B
A
vapor instantaneous
vapor accumulated
-25
-20
-15
-10
-5
0
5
10
0 2 4 6 8 10 12 14
crystallization
degassing(α=1.0100)
Undegassedtrachytic melt
Pumice (basaltic glass coating)
Pumice (trachytic glass)
Common lavas (SSB + oceanites)
AverageSSB
α=(7Li/6Li)vapor/(7Li/6Li)melt
A
Bcrystallization+ degassing(α=1.0099)
Pumice (bulk analysis)
Fig. 6. Lithium isotopic composition of the melt (upper panel) and vapor (lower panel)plotted against Li content in the residual melt. A and B refer to the two crystallization-degassing scenarios described in Fig. 5. Upper panel: all points are data points, with theexception of the undegassed trachtytic composition which is calculated as explained intext. The straight dashed line, onwhich plot the pumice samples (0704-PS and 0704-PN),the coating glass (0704-BAS) as well as a bulk analysis not reported in Table 1(Li=3.36 ppm; δ7Li=−1.8‰±0.3), suggests that the rock ismade of two homogeneouscomponents (basaltic and trachytic). Curved lines showtheevolutionofδ7Li in the residualmelt during degassing (Eqs. (10) and (9) for cases A and B). Lower panel: vaporcomposition is calculated for cases A and B using Eq. (12) and Eq. (13), respectively).
32 I. Vlastélic et al. / Chemical Geology 284 (2011) 26–34
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The average isotopic ratio of the accumulated (ACC) vapor isobtained by integrating Eq. (11). Knowing that the average value of afunction f(x) that is continuous over the interval a≤x≤b is 1/(b–a).∫f(x)dx, we infer:
7Li6Li
!V
ACC
=7Li6Li
!M
0
:1− CM
=CM0
� �αV−M
1− CM = CM0
� � ð12Þ
and
7Li6Li
!V
ACC
=7Li6Li
!M
0
:αV−M 1−F ε + 1ð Þ
� �ε + 1ð Þ 1−Fð Þ with ε =
DV−M6 αV−M−1
� �1 + k
ð13Þ
for cases A and B, respectively. During degassing of a finite batch ofmagma, the lithium isotopic composition of the accumulated vaporvaries widely, from ~13.5‰ in the first vapors down to about ~5‰ inthe case of extreme degassing, as for the April 2007 pumice (Fig. 6).The 7Li /6Li ratio of the gas phase will subsequently increase duringpartial condensation, as suggested by the isotopically light Licomposition of the April 2007 fumarolic condensates. While thecoefficient of isotopic fractionation during condensation (αCond-V) isconstrained to be less than unity, it cannot be determined preciselybecause both the isotopic composition of the condensing gas and thefraction of gas condensed are poorly known. To explain thecomposition of the April 2007 gas condensates (average of −1.7‰),we calculated thatαCond-V≤0.99 is required if the condensing gas hadδ7LiN8‰, or if more than 50% of the vapor condensed. If the vapor hadδ7Li=13.5‰, then αCond-V≤0.985.
The trachytic pumice recorded an extreme process that is notrepresentative of the degassing pattern of Piton de la Fournaisemagmas,however. Steady state basalts display rather uniform Li content andisotopic composition (Supplementary Table 1), suggesting that lithiumloss in normal degassing mode is of only a few percents and,consequently, that those volcanic emanationshave very heavy Li isotopicsignatures. The slightly heavier Li signature of oceanites (average of+3.9‰, Table 1) is due to the incorporationof cumulative olivine crystalshaving distinctly heavy composition (δ7Li=+5.2±1.4‰(2σ, N=10),Supplementary Table 1), whichmost likely reflect mineral-melt isotopicfractionation (Seitz et al., 2004) rather thana lessdegassed signature.Ourattempt to find a distinctive lithium signature in naturally quenchedsamples (Pele's hairs) failed (Li=5.6±0.4 ppm, δ7Li=+3.4±0.3‰(2σ, N=4), Supplementary Table 1). However, amongst lavas free ofcumulative olivine, the highest δ7Li (+4.0 to 4.5‰) coincide with theeruption of more primitive, less degassed magmas (as during the 1998Hudson eruptionor during the paroxysmal phase of April 2007 eruption)suggesting that degassing may influence whole-rock Li signature to ameasurable degree. We calculated that decreasing δ7Li from +4.5 to+3.6‰ (average of SSB) by degassing isotopically heavy lithium (αV–M
of 1.010) is consistent with 9% loss of lithium with δ7Li of +14.1‰.
6. Concluding remarks
The preliminary results obtained show that 7Li partitions preferen-tially into the vapor phase during low-pressure, vapor-saturatedcrystallization. This strongly suggests that Li isotopic fractionation iscontrolled by chemical rather than by kinetic effects. The composition ofthe degassed trachytic pumice places strong constraints on themagnitude of the vapor-melt isotopic fractionation (αV–M of 1.010)because the pre-degassing composition is relatively well known. Sincethe fraction of Li volatilized in normal degassing mode is very small, afewpercent at themost, a very heavy Li isotopic signature (δ7LiN+13‰)is expected in volcanic emanations. In addition, the Li composition of thevapor may subsequently increase as isotopically light lithium conden-sates. These findings and expectations are summarized in Fig. 7.
This behavior of lithium isotopes is consistent with that inferredfor boron isotopes, but in magmatic systems only (You, 1994; Gillis etal., 2003; Kuritani and Nakamura, 2006). It is possible that Li, as with B(Rose-Koga et al., 2006), obeys somewhat different laws duringevaporation and condensation processes in the ocean–atmospheresystem. By comparison with heavier isotopic systems, Li isotopesseem to behave rather like Zn isotopes during degassing (preferentialevaporation of heavy isotopes (Toutain et al., 2008)), but rather followTl isotopes during condensation (preferential removal of lightisotopes (Baker et al., 2009)). However, because these isotopicsystems were studied independently in different volcanic settings, itis unclear whether the contrasting isotope behaviors are fundamentalor just reflect distinct degassing contexts. A more comprehensiveview of trace isotopes behavior during degassingwill probably requiremulti-isotopic approaches.
Supplementarymaterials related to this article can be found onlineat doi:10.1016/j.chemgeo.2011.02.002.
Acknowledgements
The authors are grateful to A.J.R. Kent and an anonymous reviewerfor their constructive comments, and to R. Rudnick for editorialhandling. E. Rose-Koga, G. Falco and K. Koga are thanked fordiscussions. Thanks also to N. Bolfan-Casanova, C. Bosq, K. David, J.-L.Devidal, J.-L. Piro (LMV, Clermont-Ferrand), C. Douchet (LST, Lyon) andC. Le Losq (IPGP, Paris) for technical assistance in the lab. Thanks toSARM (CRPG, Nancy) and L. Bouvier (LaMP, Clermont-Ferrand) forproviding major element analyses. This study benefited financialsupport from the Institut National des Sciences de l'Univers (ST,A02009 and AO2010).
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