# Insights into crustal assimilation by Icelandic basalts from boron isotopes in melt inclusions from the 1783-1784 Lakagígar eruption

Maryjo Brounce, Maureen Feineman, Peter Lafemina, Andrey Gurenko

► To cite this version:

Maryjo Brounce, Maureen Feineman, Peter Lafemina, Andrey Gurenko. Insights into crustal assimilation by Icelandic basalts from boron isotopes in melt inclusions from the 1783-1784 Lakagígar eruption. Geochimica et Cosmochimica Acta, 2012, 94, pp.164-180. 10.1016/j.gca.2012.07.002. insu-03619304

HAL Id: insu-03619304

https://insu.hal.science/insu-03619304v1

Submitted on 6 Mar 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
1 Insights into Crustal Assimilation by Icelandic Basalts from Boron Isotopes in Melt

2 Inclusions from the 1783-1784 Lakagígar Eruption

3 Maryjo Brounce¹†, Maureen Feineman¹*, Peter LaFemina¹ and Andrey Gurenko²‡

4 ¹The Pennsylvania State University, Department of Geosciences, University Park, PA,

5 16802, USA

6 ²Woods Hole Oceanographic Institute, Department of Geology and Geophysics, Woods

7 Hole, MA, 02543, USA

8 †Now at University of Rhode Island, Graduate School of Oceanography, Department of

9 Marine Geology and Geophysics, Narragansett, RI, 02882, USA

10 ‡Now at Centre de Recherches Pétrographiques et Géochimiques, CNRS-Nancy

11 Université, BP 20, F-54501 Vandoeuvre-lès-Nancy, France

12 * Corresponding author

13
14 ABSTRACT

15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
32
33
34
35
36
37
The boron isotope system has great potential for tracing alteration and assimilation in basaltic systems due to the very low concentrations of B in mantle-derived melts and the strong isotopic contrast between the mantle and surface materials. However, variability in B concentrations and isotope ratios in basalts can also be interpreted to reflect inputs from enriched regions of the mantle, as the extent of mantle heterogeneity with respect to boron remains poorly delineated. We have determined boron concentrations and isotope ratios in fresh, glassy, plagioclase-hosted melt inclusions and unaltered scoriaceous matrix glasses from four localities associated with the 1783-1784 Lakagígar (Laki Fissure) eruption, Iceland. Boron concentrations range from 0.59-1.25 ppm in the melt inclusions, and from 1.25-1.65 ppm in the matrix glasses, while $\delta^{11}B_{NBS-951}$ ranges from -7.8‰ to -16.5‰ in the melt inclusions and -10.5‰ to -16.9‰ in the matrix glasses. In contrast to previous studies of boron in basaltic melt inclusions from other fissure swarms in Iceland (Gurenko and Chaussidon, 1997, *Chem. Geol.* 135, 21-34), the Lakagígar melt inclusions display a significant range of boron concentrations and isotope ratios at constant $K_2O$ wt.%, which is more consistent with B addition by assimilation of altered basalt than it is with mixing between depleted and enriched mantle sources. Assimilation of freshwater-altered crustal materials at depth may impart a light $\delta^{11}B$ signature such as that observed in the Lakagígar melt inclusions and tephra host glasses. Considering boron concentrations and isotope ratios in the Lakagígar glasses and previously studied altered Icelandic basalts, together with volatile equilibration depths of the Lakagígar melt inclusions, we propose that a) mantle-derived magmas formed beneath Lakagígar assimilated ~5-20% altered crust at a depth of ~3-4 km or more, probably during magma accumulation in sills formed at the boundaries of low-density hyaloclastite layers; and b)
the magma subsequently underwent extensive mixing and homogenization prior to eruption, quite possibly within the magma chamber beneath the Grímsvötn central volcano, assimilating an additional ~10% of altered crust at that time, for a total of up to 30% crustal assimilation. We hypothesize that volatiles including H₂O, CO₂, S, F, and Cl, which were responsible for the majority of the considerable casualties attributed to the Lakagígar eruption, were added together with isotopically light B by assimilation of hydrothermally altered crustal materials.

45

## 1. INTRODUCTION AND GEOLOGICAL BACKGROUND

Explosive, large volume eruptions impact climate and environment at local, regional and global scales. The Mid-Atlantic Ridge is expressed in Iceland in the form of sub-aerial neovolcanic zones (Figure 1a). Magmatic systems in these neovolcanic zones consist of central volcanoes with crustal magma chambers and fissure swarms that accommodate plate spreading via lateral magma injection (Sigurdsson and Sparks, 1976). Holocene basaltic fissure eruptions in Iceland are among the largest historic eruptions on Earth and have had profound global impacts (Thorarinsson, 1979). The 1783-84 Lakagígar eruption (also known as Laki Fissure, or the “Skaftár Fires”) was one of the most voluminous single eruptions of historic times. The fissure starts 50 kilometers southwest of the Grímsvötn central volcano and presumably continues beneath Vatnajökull to intersect with Grímsvötn itself, bisecting the older Laki hyaloclastite mountain (Figure 1b). Starting in June 1783 and continuing into January 1784, the fissure first erupted to the southwest near the Skaftá River and proceeded to the northeast back toward Grímsvötn,
producing ~14.7 km³ of lava, ~0.4 km³ of tephra and ~10¹² tons of volatiles, including ~10¹¹ tons each of sulfur aerosols, H₂O, and CO₂, and smaller but significant amounts of HF and HCl (Thorarinsson, 1979; Thordarson and Self, 1993; Thordarson et al., 1996).

Famine, disease, and the harsh winter following the eruption are all attributed to the tephra and, more importantly, the aerosols released. All told, these factors claimed as many as 10,000 lives worldwide, making the 1783-84 Lakagígar eruption not only one of the most voluminous, but also one of the deadliest historic eruptions (Thorarinsson, 1979; Tilling, 1996).

The striking geochemical homogeneity of the large volume of lava erupted from Lakagígar (Sigmarsson and Condomines, 1991) suggests efficient mixing in a large magma chamber at some level beneath the surface. If magma accumulation does occur at one or more levels in the crust, then we might expect to find some record of crustal assimilation. Indeed, previous geochemical studies indicate variable amounts of crustal assimilation prior to eruption at Lakagígar. Sigmarsson and Condomines (1991) infer assimilation of metabasalts on the order of ~20%, accompanied by magma mixing at the crust-mantle boundary, based on U and Th concentrations and Sr, Th and O isotope ratios. Others suggest shallower assimilation of up to 60% hydrothermally altered crustal materials based on isotopic data from basalts and melt inclusions, as well as phenocryst-matrix oxygen isotopic disequilibrium in Lakagígar lavas (Hattori and Muehlenbachs, 1982; Hémond et al., 1988; Bindeman et al., 2006, 2008). Métrich et al. (1991) suggest a shallow level for magma mixing (~4 km), perhaps within the Grímsvötn magma chamber, based on compositional trends and volatile concentrations observed in melt
83

inclusions, while Bindeman et al. (2006, 2008) argue for extensive assimilation at depths
below the central magma chamber. Note that these interpretations are by no means
mutually exclusive.

84

85

86

Boron isotope ratios (¹¹B/¹⁰B, expressed as δ¹¹B = ((¹¹B/¹⁰B)unknown/¹¹B/¹⁰B)standard - 1) * 1000 in permil variation relative to the NBS-951 boric acid standard) have the potential to contribute uniquely to understanding crustal assimilation at basaltic fissures.

87

88

89

Boron isotopes are not strongly fractionated during crystallization of magma (Benton et al., 1999), and boron is compatible in basaltic melt relative to fluid at low pressures, such that boron loss during degassing of basalts is likely to be limited (Hervig et al., 2002).

90

91

92

Since boron is present in low concentrations in primitive basaltic magmas (~0.5-1.0 ppm, Chaussidon and Marty, 1995), δ¹¹B is very sensitive to alteration and assimilation of altered crustal materials. In addition to crustal inputs, however, we must consider the boron concentrations and isotope ratios in the parental magma, which is not at all straightforward.

93

94

95

96

97

98

Previous analyses of basaltic melt inclusions and tephra host glasses suggest that the mantle source region for Icelandic basalts is uniformly light with respect to δ¹¹B (-11.3, 2σ ± 3.8‰; Gurenko and Chaussidon, 1997). In glasses from the Reykjanes and Hengill fissure swarms (Figure 1a), B concentrations increase with increasing K₂O wt.%. It has been demonstrated that B and K are not strongly fractionated by mantle melting (Ryan and Langmuir, 1993), such that variations in melt fraction will cause B and K₂O concentrations in a suite of basalts to co-vary with a constant B/K₂O as long as they are

99

100

101

102

103

104

105
derived from a homogeneous mantle source. The Reykjanes melt inclusions, which are
sampled at an area relatively far from the center of the presumed Icelandic plume (where
the plume center is at a location approximately equivalent to the 40 km Moho contour
shown in Figure 1a), are characterized by mid-ocean ridge (MORB)-like values, with
average B/K₂O = 20 at K₂O <0.1%. The Hengill melt inclusions, which are sampled
from an area somewhat closer to the Icelandic plume, range from MORB-like values
similar to those at Reykjanes to more ocean island basalt (OIB)-like values at K₂O = 0.1-
0.4 wt.%. The more enriched melt inclusions from Hengill are characterized by a
relatively constant B/K₂O = ~4, consistent with B addition from a plume-related source.
Lakagígar, which lies still closer to the center of the presumed Icelandic plume (Figure
1a), should by the same logic extend to higher B and K₂O concentrations at roughly the
same B/K₂O, with similarly invariant light δ¹¹B.

118

The purpose of this study is to determine to what extent the boron concentrations and
isotope ratios at Lakagígar are controlled by the mantle source, which in this case would
include an enriched OIB-like contribution characterized by unusually light δ¹¹B, and to
what extent they are controlled by assimilation of altered crustal materials. To this end,
we present new boron concentration and isotope ratio data, as well as lithium
concentrations, volatile concentrations, and major element compositions of tephra matrix
glasses and plagioclase-hosted melt inclusions from the 1783-1784 Lakagígar eruption.
Modeling of these data help constrain the source of the parent magma and the magnitude
of crustal assimilation, and provide a tentative clue to the source of the volatiles that were
so destructive in the aftermath of this eruption.
129

## 2. METHODS

### 2.1 Sample Selection

We sampled four tephra units originating (presumably) from two tuff cones, representative of the phreatomagmatic stages of the 1783-84 eruption (Figure 1b). Loose plagioclase phenocrysts were hand-picked from the tephras under a binocular microscope. In some cases, melt inclusions were clearly visible within the transparent crystals. The plagioclase crystals were mounted in epoxy and polished to expose the melt inclusions at the surface (Figure 2a). Melt inclusions were chosen for analysis based on the following criteria: 1) lack of visible daughter crystals; 2) clean, sharp contacts with the host crystal; 3) lack of any cracks or breaches from the melt inclusion to the crystal surface. In addition to melt inclusions, several intact scoria lapilli from each tephra sample were mounted in epoxy and polished for analysis of the matrix glass. The scoria samples studied were selected to be free of secondary alteration products. Microlites of plagioclase were ubiquitous in the matrix glasses (Figure 2b), and spots for analysis were carefully chosen to avoid these crystals.

145

### 2.2 Major Elements

Major element compositions were obtained using a Cameca SX-50 Electron Probe Micro-Analyzer (EPMA) at Penn State. The instrument was operated at an accelerating voltage of 15 keV and a beam current of 12 nA. Glasses were analyzed using a defocused electron beam (~20µm) to minimize Na volatilization, while plagioclase phenocrysts were analyzed using a focused electron beam (~1µm). The x-ray intensity
data were subject to a PAP (phi-rho-z) matrix correction algorithm as described in Pouchou and Pichoir (1987), and converted to concentrations by comparison to natural and synthetic standards. All samples were carbon-coated prior to EPMA analysis.

155

## 2.3 Boron and Lithium Concentrations and Boron Isotope Ratios

Boron and lithium concentrations and boron isotope ratios were analyzed in the scoria and plagioclase-hosted melt inclusions by Secondary Ionization Mass Spectrometry (SIMS) using the Cameca IMS 3f (some Li and B concentrations) or Cameca IMS 1280 (most concentrations and all isotope ratios) at the Northeast National Ion Microprobe Facility at the Woods Hole Oceanographic Institute (WHOI), following the techniques described by Chaussidon and Jambon (1994) and Chaussidon et al. (1997).

163

### 2.3.1 *Cameca IMS 3f*

Standards and unknown samples were sputtered with a nominal 8.3 kV ¹⁶O⁻ primary beam focused to about 20–30 µm in diameter. Instrument conditions are provided in Supplemental Data (Table S1). Each analysis consisted of 5 sequential blocks of the masses ⁷Li, ¹¹B and ³⁰Si. Element concentrations relative to independently known SiO₂ concentrations were calculated using calibration curves obtained by linear regressions of a set of 3 reference glasses (NBS-610, 357 ppm B; NBS-612, 35 ppm B; and KL-2G, 3 ppm B; values from Jochum et al., 2000 and Rosner et al., 2008; calibration line shown in Supplemental Data, Figure S1a). Instrument drift was monitored and corrected using daily replicate analyses of KL2-G reference glass. The analytical precision was better than 15% relative standard deviation.
175

176 2.3.2 *Cameca IMS 1280*

177 Standards and unknown samples were sputtered with a nominal 13 kV <sup>16</sup>O primary beam
178 focused to approximately 20-25 µm in diameter. Instrument settings are given in
179 Supplemental Data Table S1. Samples were sputtered for 500 seconds before they were
180 analyzed, and each measurement included 50 cycles for element concentrations or 200-
181 250 cycles for boron isotope ratios. Boron and lithium concentrations were calculated
182 relative to independently known SiO<sub>2</sub> concentrations in the unknown samples of interest
183 using calibration lines obtained by linear regression of measurements made on a set of
184 five reference glasses (GOR128-G, 20 ppm B; StHs6/80-G, 11.6 ppm B; ATHO-G, 6
185 ppm B; KL-2G, 3 ppm B; and ML3B-G, 2 ppm B; values from Jochum et al., 2000 and
186 Rosner et al., 2008), as shown in Supplemental Data, Figure S1b. Analytical precision
187 was ~5-10% relative standard deviation for boron and less than 5% relative standard
188 deviation for lithium.

189

190 For isotope ratio measurements, the instrumental fractionation factor was determined
191 based on multiple measurements of the MPI-DING glass standard StHs6/80-G (11.6 ppm
192 B, δ<sup>11</sup>B = -4.39‰; Rosner et al., 2008). The suitability of this standard was verified by
193 analyzing its boron isotope ratios in the same session with five other standard reference
194 glasses (NBS-610, NBS-612, Rosner et al., 2008; GOR-128G, GOR-132G, GB4-CRPG,
195 Rosner and Meixner, 2004). The isotope ratio data for the six standards are linearly
196 correlated when plotted against the known δ<sup>11</sup>B values, with r<sup>2</sup> = 0.982 and α = 0.9714
197 for this session (Supplemental Data, Figure S1c). The stability of the instrument and
relative homogeneity of the glass standard were verified by comparing repeat analyses of StHs6/80-G over the course of the 3-day session, which yields an average ¹¹B/¹⁰B of 3.9315 ± 0.0111 (2σ external; Figure S1d). Reported uncertainties on unknowns (Table 1) are 2σ internal precision on individual measurements, which are greater than the external reproducibility of the standard. All isotope ratios are expressed as δ¹¹B calculated relative to the NBS-951 standard (¹¹B/¹⁰B = 4.04558 ± 0.00033; Spivack and Edmond, 1986).

## 2.4 Volatile Concentrations

Concentrations of the volatile species CO₂, H₂O, F, S, and Cl were determined using the Cameca IMS 1280 at WHOI, following the methods described by Hauri et al. (2002) and Shimizu et al. (2009). The samples were carefully polished to remove residual carbon from the carbon-coating process used in preparation for prior analysis by EPMA, then removed from the epoxy mount, cleaned, and remounted by pressing them into indium metal. Samples were sputtered with a 10 kV Cs⁺ primary beam focused to 10-20 µm diameter. Instrument settings are given in Supplemental Data Table S1. The measurements were started after a 200 second pre-sputtering, and some spots required multiple analytical cycles to remove unnaturally high concentrations of residual carbon from the carbon coating. Six reference glasses (GL07 51-3, ALV519-4-1, EN113 46D, ALV1649-3, ALV1654-3, FRND 6001; volatile concentration values as in Helo et al., 2011) were analyzed to create the calibration lines used to calculate absolute concentrations of volatile species in the unknown samples (Supplemental Data Figure S2).
221

222 **3. RESULTS**

223 Major element concentrations, boron and lithium concentrations, boron isotope ratios,
224 and volatile data are presented in Table 1, and selected data are presented in Figures 3
225 and 4.

226

227 **3.1 Analytical Results**

228 Although the Lakagígar glasses measured here are, overall, homogeneous tholeiitic
229 basalts (SiO₂ wt % ~48.5-50.5), subtle variations can be observed in the major and
230 incompatible element data. For example, the melt inclusions have higher Mg numbers
231 [Mg# = molar Mg/(Mg+Fe)×100], ranging from ~43-46 with an outlier at 51, relative to
232 the scoria host glasses, which have Mg numbers ~39-41 (Figure 3a and b). We observe a
233 moderately wide range of TiO₂ (~1.8-3.1 wt.%) over a very restricted range of Mg# in the
234 melt inclusions, which is not consistent with fractional crystallization from a common
235 parent magma. In contrast, correlation between TiO₂ and Mg# in the *host glasses* could
236 reflect a small amount of crystallization, likely occurring upon eruption. The rather good
237 correlation between CaO/Al₂O₃ and Mg# over the complete range of glasses analyzed
238 (Figure 3b) is consistent with fractional crystallization of clinopyroxene, although the
239 trend is not observed *within* the suite of melt inclusions with Mg# = 43-46.
240 Clinopyroxene is believed to be a late-crystallizing phase in the Lakagígar basalts (Bell
241 and Humphries, 1972). The matrix glasses have extremely restricted ranges of TiO₂ and
242 total alkalis (Na₂O+K₂O) relative to the melt inclusions (Figure 3c), with the matrix
243 glasses occupying the extreme high end of the range observed in TiO₂ wt % of the
244 inclusions. The matrix glasses and melt inclusions overlap in terms of B concentrations,
245 with the matrix glasses extending toward higher B concentrations and the melt inclusions
246 extending toward lower B concentrations (Figure 3d). While the major and minor
247 element correlations in the Lakagígar glasses are generally weak and characterized by
248 exceptions, the Li vs. B and δ¹¹B vs. B trends are relatively clear, and more or less
249 continuous from the inclusions to the hosts glasses (Figures 3e and f).
250

251 A few characteristic melt inclusion types can be identified from among the larger dataset.
252 Most notably, inclusion 17pl3a is characterized by significantly higher Mg# (51) and
253 CaO wt % (10.82) relative to all the other inclusions analyzed in this study (43-46 and
254 9.41-9.76, respectively). In contrast, inclusion 17pl3b, which is located in the same
255 phenocryst (slightly closer to the edge), has a lower Mg# (43) more typical of the
256 Lakagígar melt inclusions, and high TiO₂ wt. % (2.98) that overlaps with that of the host
257 glasses. Inclusions 12pl1a and c, both trapped in the same melt inclusion, are outliers in
258 terms of their high total alkali contents (Na₂O+K₂O wt.% = 3.6-3.8), low B and Li
259 concentrations (0.59 and 3.31 ppm, respectively), and relatively high δ¹¹B (-7.8‰).
260 Inclusion 8pl3a has the highest boron concentration (1.25 ppm) and lightest δ¹¹B (-
261 16.5‰) of any melt inclusion analyzed, but is in no way distinguishable from the
262 “typical” Lakagígar melt inclusions in terms of major and incompatible element
263 characteristics. Inclusion 14pl1a is most similar to the matrix glasses in terms of TiO₂
264 and total alkalis, and has relatively light δ¹¹B (-13.9‰). Relative to the melt inclusions,
265 the host glasses are more homogeneous, and are characterized by relatively high TiO₂,
266 low Mg#, high boron and lithium concentrations, and light δ¹¹B.
267

268

The volatile components ($H_2O$, $CO_2$, S, F, Cl) display a broadly positive correlation with B concentrations in the melt inclusions (Figure 4a-e). The host glasses, in contrast, have slightly higher B concentrations despite being strongly degassed with respect to $H_2O$, $CO_2$, and S (Figure 4a, b, and c). The halogens F and Cl are only partially degassed in the matrix glasses relative to the melt inclusions, and in fact some melt inclusions have lower F and Cl concentrations than the host glasses (Figure 4d and e). Despite the fact that both $H_2O$ and $CO_2$ seemingly correlate with boron concentrations, the correlation between $H_2O$ and $CO_2$ in the melt inclusions is somewhat ambiguous (Figure 4f), and will be discussed in more detail in Section 4.

269

270

271

272

273

274

275

276

277

278

## 3.2 Effects of Post-Entrapment Crystallization

279

Attempts to correct the melt inclusion compositions for post-entrapment crystallization of plagioclase using both the method of Kress and Ghiorso (2004) and that of Hamada and Fujii (2007) predicted unreasonably large degrees of post-entrapment crystallization (on the order of 25-60%). The method of Kress and Ghiorso (2004), which uses the MELTS software package, does not adequately account for the affect of $H_2O$ in the magma, which serves to increase the anorthite content of plagioclase in equilibrium with a given melt composition. The method of Hamada and Fujii (2007), which utilizes an empirically determined pressure-, temperature-, and composition- dependent partition coefficient, predicted unreasonably high degrees of crystallization for melt inclusions from Izu-Oshima with $H_2O$ contents 0.2-1.4%, comparable to the Lakagígar melt inclusions. Hamada and Fujii (2007) attribute this mismatch to hydrogen loss by diffusion, which

280

281

282

283

284

285

286

287

288

289
caused them to under-estimate the H₂O contents of the melt inclusions at the time of the final stage of post-entrapment crystallization. As will be discussed in Section 4.5, under-estimation of the initial H₂O contents is possible for the Lakagígar melt inclusions as well. It is also possible that the composition-dependent partition coefficient of Hamada and Fujii (2007) still does not adequately describe the effect of H₂O in the melt on the anorthite content of equilibrium plagioclase.

The boron and volatile concentrations in the melt inclusions presented here have not been corrected for post-entrapment crystallization. Careful visual inspection by transmitted light, reflected light, backscattered and secondary electron imaging suggests a lack of significant post-entrapment crystallization; however, visual detection of post-entrapment crystallization is not always possible. Boron is incompatible during crystallization of An-rich plagioclase from basalt (Bindeman et al., 1998). Métrich et al. (1991) calculated 2-11% post-entrapment crystallization of Lakagígar melt inclusions, mostly based on olivine-hosted inclusions. Thus we estimate that a limited amount of post-entrapment crystallization may have increased B and volatile concentrations in the melt inclusions by up to 10% relative to their values at the time of entrapment, which is within the uncertainty of the data. Boron isotopes are not fractionated by crystallization of basaltic magma (Benton et al., 1999), so no correction is needed.

# 4. DISCUSSION

## 4.1 Compositional Variation in Lakagígar Glasses

### 4.1.1 Major elements
Although the melt inclusions from Lakagígar are relatively homogeneous in terms of their major element compositions (SiO₂ = 48.4-50.6 wt.%; Na₂O+K₂O = 2.70-3.73; Mg# = 43-46, with one inclusion at 51), several compositional outliers can be identified. For example, inclusions 12pl1a and 12pl1c have high total alkalis, while 14pl1a and 17pl3b have high TiO₂, and 17pl3a has distinctively high Mg# and CaO, coupled with low TiO₂ (Figure 3). We interpret, based on the compositional contrast between inclusions 17pl3a and 17pl3b (hosted in the same phenocryst, but slightly closer to the rim), that inclusion 17pl3a is most likely trapped in a xenocrystic plagioclase core that was later overgrown by new plagioclase. The xenocrystic core represents either an earlier incarnation of the Lakagígar-Grímsvötn magmatic system or even another magmatic system entirely; such xenocrysts have been previously recognized in Lakagígar magmas on the basis of oxygen isotopic disequilibrium (Bindeman et al., 2008). We note that melt inclusions such as 17pl3a with Mg# > 50 were observed in greater abundance by Métrich et al. (1991); these more primitive melt inclusions may be under-represented in our study. No coherent incompatible element variation is observed to suggest variable degrees of partial melting or fractional crystallization from a single source – for example, TiO₂ does not co-vary with K₂O or total alkalis (Figure 4c). Likewise, B shows no coherent variation with K₂O or total alkalis (Figures 4d and 6a).

We propose that the apparently random variability in major element compositions is a result of stalling and assimilation of older basaltic units at various depths within the crust beneath Lakagígar. The Tertiary through Holocene basalts underlying Lakagígar may have formed by similar processes and from similar mantle sources, hence the general
336 homogeneity, but at different times and from different magmatic systems such that they
337 need not be identical in composition. Assimilation by the parental Lakagígar magma of
338 different proportions of crustal materials at different depths and with slightly different
339 compositions could explain the uncorrelated major element variations observed in both
340 the melt inclusions and the matrix glasses. The influence of crustal assimilation even
341 within the melt inclusions is supported by the range of Mg#s, which at 43-51 is
342 considerably lower than other OIB-associated melt inclusions (e.g., Mg# = 55-75 in
343 Hawaii, Kobayashi et al., 2004; Mg# up to 88 in Polynesia, Saal et al., 1998).
344

345 4.1.2 Boron concentrations and isotope ratios

346 We can compare the melt inclusions and host glasses from Lakagígar with previously
347 analyzed melt inclusions and host glasses from the Reykjanes and Hengill fissure swarms
348 (Gurenko and Chaussidon, 1997), also located in southern Iceland (Figure 1a), on the
349 basis of their B/K2O ratios and δ11B (Figure 5). Previous authors have shown that
350 B/K2O ratios vary between Mid-Ocean Ridge Basalts (MORB) and Ocean Island Basalts
351 (OIB), with lower B/K2O found in OIB relative to MORB (Figure 5a; Ryan et al., 1996;
352 Chaussidon and Marty, 1995; Chaussidon and Jambon, 1994). Because B and K are
353 believed to have similar partition coefficients during mantle melting (Ryan and
354 Langmuir, 1993), this variability in B/K2O has been interpreted to represent either
355 fundamental differences in the mantle sources sampled by these two basalt types (Ryan et
356 al., 1996) or addition of B to MORB by assimilation of hydrothermally altered oceanic
357 crust (Chaussidon and Marty, 1995). Geographically, Lakagígar lies in close proximity
358 to the center of the proposed plume (OIB source) beneath Iceland, while Reykjanes is far
359 from the plume, thus presumably more influenced by a MORB-like source (Figure 1a).

360 These assumptions regarding the nature of the mantle sources for magmas at the two
361 volcanic centers are consistent with the B/K₂O ratios, which are more MORB-like in melt
362 inclusions from Reykjanes and more OIB-like in inclusions from Lakagígar (Figure 5a).

363 Hengill, which lies geographically between Reykjanes and Lakagígar, has a range of
364 B/K₂O that spans the difference between the two. The Icelandic melt inclusions as a
365 whole contain some of the lightest boron ratios ever measured in fresh basalts (δ¹¹B = -8
366 to -14‰ at Reykjanes and Hengill, extending to less than -16‰ at Lakagígar, Figure 5b).

367 All but two of the Lakagígar glasses fall within 2σ of δ¹¹B values previously measured in
368 Icelandic melt inclusions (δ¹¹B = -11.3±3.8(2σ)‰; Gurenko and Chausidon 1997), and
369 all but three fall within 2σ of δ¹¹B values reported for OIB worldwide (δ¹¹B = -
370 9.9±2.6(2σ)‰; Chausidon and Marty, 1995). However, the negative correlation
371 between B concentration and δ¹¹B is a feature unique to Lakagígar.

372

373 On the basis of B/K₂O ratios and the range of δ¹¹B, it may seem at first glance that the
374 Lakagígar melt inclusions simply extend the previous dataset for Iceland to more strongly
375 OIB-influenced values, where the OIB in question is sourced from a mantle plume
376 particularly depleted in ¹¹B. However, the increase of B independently of K₂O is not
377 consistent with variable degrees of melting or crystallization from a single OIB mantle
378 source, nor is it consistent with mixing between MORB and OIB. On a plot of B vs.
379 B/K₂O (Figure 5a), the overall trend for OIB is toward increasing B at a relatively
380 constant B/K₂O of ~3, which reflects the similar partitioning behavior of B and K. The
381 Lakagígar glasses, although they exhibit a range of B/K₂O that is typical of OIB, show a distinct linear correlation between B and B/K₂O, which indicates that B was added independently of K₂O. On the plot of δ¹¹B vs. B (Figure 5b), the Lakagígar glasses show a trend of decreasing δ¹¹B with increasing boron concentration, a trend that is not observed in the glasses from Hengill or Reykjanes. The good correlation between B, Li, and δ¹¹B, in the absence of any correlation with K₂O, TiO₂, Mg# (Figures 3 and 5), or P₂O₅ (not shown) suggests that the process controlling B and Li abundances and δ¹¹B is independent of primary magmatic processes including melt fraction, fractional crystallization, and even mantle source variability. Assimilation of altered basaltic materials in the crust, in which the B and Li abundances are controlled by the extent of alteration rather than primary magmatic processes, would explain the observed correlation between the light elements and δ¹¹B in the absence of correlation with other incompatible elements.
394

395 It is additionally worth noting that despite the MORB-like B/K₂O in the Reykjanes and some Hengill melt inclusions, the only glasses from Iceland with MORB-like δ¹¹B are the *host glasses* from Reykjanes (Figure 5). That is, none of the previously analyzed basaltic melt inclusions from Iceland, including those from the presumably MORB-influenced Reykjanes fissure swarm, contain MORB-like δ¹¹B (Gurenko and Chaussidon, 1997).
396
397
398
399 Thus although the Lakagígar glasses seem to reflect a strongly plume-influenced source relative to Reykjanes and Hengill, the lack of MORB-like boron in any Icelandic melt
inclusions suggests that simple MORB-OIB mixing does not satisfactorily explain the
full range of boron behavior.

404

405 **4.1.3 Volatiles**

We observe a weak positive correlation between B and volatile concentrations in the melt
inclusions (Figure 4). The general correlation might be interpreted as suggesting that B
contents are controlled by B loss during degassing. However, several lines of evidence
seem to contradict this interpretation. First of all, the scoriaceous matrix glasses, which
are strongly degassed with respect to H₂O, CO₂, and S, have higher B contents than the
melt inclusions as a whole. In order to explain the difference in B versus volatile
behavior between the melt inclusions and matrix glasses, we would need to call upon
very different B partitioning behavior at depth relative to that accompanying shallow-
level or syn-eruptive degassing. Secondly, Hervig et al. (2002) report experimental
evidence that B is relatively compatible in basaltic liquids relative to aqueous fluid
(D<sub>B</sub><sup>fluid/melt</sup> = 0.33-0.54), and argue that B concentrations in melt inclusions from Cerro
Negro volcano show no evidence for significant B loss accompanying degassing of CO₂-
H₂O fluids. Finally, Hervig et al. (2002) also show experimentally that the isotopic
fractionation factor for B partitioning between melt and fluid is such that the fluid favors
heavier B (1000 lnα<sup>melt/fluid</sup> = -1.3 to -5.2), which means that the residual magma
following boron degassing should have *lower* δ<sup>11</sup>B compared to the relatively undegassed
parent magma. The trend observed in our data is just the opposite; the melt inclusions
with the lowest B (and volatile contents) have the highest δ<sup>11</sup>B.
424

Although we may be able to rule out degassing as a controlling factor, interpretation of the volatile concentrations is complicated by possible post-entrapment re-equilibration of volatiles by diffusion through the host phenocryst (this possibility is discussed in more detail in Section 4.5). And even if we take the volatile concentrations at face value (that is, if we assume the concentrations in the melt inclusions represent the concentrations in the magma at the time of entrapment), interpretation of the H₂O-CO₂ relationship is not straightforward. The general clustering of the H₂O contents around 0.75 wt% is consistent with open-system degassing, while the lower H₂O contents in inclusions 12pl1a and 12pl1c seem to suggest mixing of components with different volatile contents (Figure 4f). In fact, the two processes may be simultaneously competing for control of the H₂O and CO₂ contents in the magma. Thus we treat the volatile concentrations as interesting supporting evidence, but refrain from drawing hard conclusions based on volatile correlations alone.

438

An alternative explanation for the correlation between B (and Li) contents, δ¹¹B, and perhaps volatile contents is that isotopically light boron is added together with Li and volatiles by assimilation of hydrothermally altered materials such as palagonite-bearing hyaloclastites or altered basalt flows. The lack of correlation between B and Mg#, total alkalis, and CaO/Al₂O₃, all of which are indices of the degree of melting and/or fractional crystallization, suggests that the processes controlling B, Li, and volatile behavior are decoupled from primary magmatic processes. The overall homogeneity of Lakagígar
446 eruptive products (for example, the almost complete lack of variability in total alkalis vs.

447 silica) suggests that the assimilated material is broadly similar in major element

448 composition to the parental Lakagígar magma. The subtle variations in major element

449 compositions within the melt inclusion suite and between the melt inclusions and the host

450 glasses could reflect subtly different assimilants at different depths beneath the fissure.

451 Although the lower B/K₂O ratios at Lakagígar relative to Reykjanes and Hengill are

452 consistent with a more OIB-like source, the lack of correlation between B or δ¹¹B with

453 K₂O *within* the Lakagígar suite prevents us from calling upon variable plume input as a

454 clear source of isotopically light B. Assimilation of variably altered basaltic materials,

455 however, could provide a means of adding B, Li, and volatiles independently of any

456 variability (or lack thereof) in bulk composition.

457

458 **4.2 Boron in Icelandic Basalts and Geothermal Fluids**

459 Boron isotope ratios in volcanic glasses from Iceland have been used previously to

460 quantify the extent of assimilation of altered crustal materials in magma compositions

461 ranging from basalts to rhyolites. In the case of basaltic to rhyolitic tephras studied by

462 Rose-Koga and Sigmarsson (2008), crustal assimilation was associated with heavy boron

463 isotope ratios, in contrast to the light boron isotope ratios determined in this study.

464 Isotopically heavy boron is commonly observed in seawater-altered crustal materials. The

465 fractionation factor for boron isotopic distribution between basalt and fluid is such that

466 the rock strongly prefers ¹⁰B, while ¹¹B is more enriched in the fluid (α¹⁰rock/water = 0.968;

467 Spivack and Edmond, 1987). Thus basalts that have been altered by hydrothermal
interaction with seawater are isotopically much lighter than the seawater itself – but since seawater is so strongly positive (δ¹¹B = +39.5‰, Spivack and Edmond, 1987) relative to mantle-derived basalt (δ¹¹B = -10‰, Chaussidon and Marty, 1995), even with fractionation on the order of -30‰, the seawater-altered basalt will still be isotopically heavier than unaltered basalt (Figure 6; see also Spivack and Edmond, 1987; Smith et al., 1995). Chaussidon and Marty (1995) proposed that while low-temperature hydrothermal alteration of basalt by seawater drives boron isotope ratios toward heavier values, high-temperature hydrothermal alteration by meteoric water drives boron isotope ratios toward lighter values. Meteoric waters have much lighter δ¹¹B relative to seawater (e.g., δ¹¹B = -1.8 to +18.2‰ in Icelandic lake waters, Aggarwal et al., 2000). More importantly, during high-temperature alteration involving meteoric waters (for example, when a shallow magmatic system interacts with groundwater, or basaltic lavas are erupted beneath a glacier), B is very soluble in the fluid relative to the basalt and isotopically heavy boron will be stripped away from the rock, leaving the altered basalt with relatively light boron isotopic values. An unavoidable corollary of this conceptual model, however, is that the light boron isotope ratios will be associated with very low B concentrations.

Raffone et al. (2010) report isotopically light boron in altered basalts from the surface to 3 km depth in a geothermal well in the Reykjanes geothermal system (well RN-17 of the Icelandic Deep Drilling Project). They attribute the light boron at depth to hydrothermal alteration by meteoric waters early in the history of the geothermal system, while the currently active geothermal system at Reykjanes has evolved to low-temperature circulation of seawater. This interpretation is supported by oxygen and deuterium
isotopic variations observed in hydrothermal minerals from other wells at the Reykjanes
geothermal field (Pope et al., 2009). The whole rocks sampled at well RN-17 at 3000 m
depth contain 9.27 ppm B with δ¹¹B = -14.4‰. Altered basalts sampled at the surface
contained even lighter boron, with δ¹¹B = -18.3‰ at a concentration of 3.83 ppm.
Raffone et al. (2010) determined by *in situ* SIMS analysis that the primary magmatic
minerals in the rocks sampled from 650-3000 m depth within the drillhole contain very
little B (~0.01-0.50 ppm in clinopyroxene and plagioclase), and that the boron is
concentrated in hydrothermal minerals such as epidote (0.3-9.0 ppm) and amphibole
(0.14- 2.3 ppm). Nonetheless, there is insufficient B even in the hydrothermal minerals
to account for the high whole rock B concentrations. We infer that the isotopically light
boron may be concentrated on altered surfaces within the basalt, which we discuss in
detail in Section 4.3.

503

Assimilation of altered basalts enriched in isotopically light boron, such as those sampled
up to 3 km below the surface at Reykjanes, could explain the relationship between B and
δ¹¹B observed at Lakagígar, independent of major element variation. For example, the
altered basalts from below 650 m in well RN-17 contain only ~0.1 wt.% K₂O (Marks et
al., 2010), which would be consistent with the lack of correlation between B and K₂O
contents in the melt inclusions, if the B addition was by assimilation of similarly altered
basalts.

511

512

4.3 Modeling the Effects of Hydrothermal Alteration
Before we can assess the extent of crustal assimilation in Lakagígar magmas, we must first identify the mantle-derived and crustal end members. With respect to the mantle-derived endmembers, the range of basalts can be distilled down to two representative compositions: OIB ([B] = 1.1 (2σ±1.0) ppm, δ¹¹B = -9.9 (2σ±2.6)‰; Chaussidon and Marty, 1995) and N-MORB ([B] = 0.5 (2σ±0.2) ppm, δ¹¹B = -4.0 (2σ±3.2)‰; Chaussidon and Marty, 1995). A few of the melt inclusions and matrix glasses from Lakagígar have δ¹¹B values outside the range of these endmember basaltic compositions, leaving us in need of an isotopically light mixing endmember. Although one might argue that Iceland is simply characterized by an unusually light OIB source that is variably expressed in the Lakagígar glasses, the addition of B independently of K₂O within the Lakagígar suite (Figure 5a) suggests that the source of light boron may be something more exotic. Bindeman et al. (2006, 2008) suggested that low δ¹⁸O found in eruptive products from Lakagígar results from subsurface stoping, digestion, and devolatilization of hyaloclastites that have been hydrothermally altered by synglacial meltwaters. The case of boron is slightly more complicated than that of oxygen, because low-temperature meteoric waters do not contain enough B to contribute significantly to the water-rock system (e.g., meteoric waters from the Icelandic lakes Myvatn and Thingvallavatn contain <50 ppb B; Aggarwal et al., 2000). If these waters are incorporated into high-temperature hydrothermal systems, however, the boron-depleted fluids will scavenge boron from the basaltic materials through which they flow. The Cl/B ratios found in geothermal waters in the area (~30, on a molal basis; Arnórsson and Andresdottir, 1995) are similar to those measured in the melt inclusions and matrix glasses from Lakagígar, suggesting an intimate relationship between geothermal fluids and basalts with respect to
these elements. High-temperature hydrothermal waters from various locations in Iceland contain, on average, [B] = 3 ppm and δ¹¹B = -3.7‰, indicating that the extent of isotopic fractionation associated with boron scavenging during this sort of high-temperature fluid-rock interaction is limited to a few permil (Aggarwal et al., 2000).

Although high-temperature hydrothermal fluids from Iceland contain higher concentrations of isotopically light boron relative to low-temperature hydrothermal fluids and meteoric waters (Aggarwal et al., 2000), the partitioning of B between basalt and fluid at high temperatures is such that the fluid will remove B from the rock, rather than adding it. The altered basalts from the Reykjanes geothermal field (Raffone et al., 2010) contain isotopically light boron at uncharacteristically high concentrations for basaltic materials (up to ~12 ppm B). These lavas would have had to have tremendously high initial boron concentrations if the light δ¹¹B observed today simply reflected loss of heavy boron from an OIB-like basalt with (δ¹¹B ≈ -10‰) during high-temperature interaction with B-poor meteoric waters. We speculate that, as meteorically-derived hydrothermal fluids circulate, they first accumulate boron from interaction with basalt, then cool to a temperature at which the boron can adsorb onto mineral surfaces (e.g., the clay-rich matrix that results from palagonitization of hyaloclastite), with preferential adsorption of ¹⁰B relative to ¹¹B (Palmer et al., 1987). Thus, interaction with boron-bearing fluids can potentially drive hyaloclastite or other altered basaltic materials to high B concentrations characterized by low δ¹¹B. The fluids must cool in order for this hypothetical model to function, because if they were to remain at temperatures >200°C, all of the boron would remain in the fluid rather than adhere to mineral surfaces.
Alternatively, isotopically light boron could be taken up directly from meteoric waters during the process of palagonitization. Due to the very low concentration of B in meteoric water, large quantities of water would be required relative to the amount of basalt, which poses thermal limitations. The basaltic lava would likely be quenched before enough meteoric water can be processed to account for the high B concentrations in the altered basalts.

Let us explore, hypothetically, the B concentrations and δ¹¹B expected in hyaloclastites or other altered basalts onto which B has been adsorbed from cooling hydrothermal waters. Hyaloclastites are volcanic breccias composed of crystals and glass fragments in a hydrated matrix. They are formed when interaction between hot lava and cold water causes the lava to shatter into mm- to cm-sized shards and ash-like fragments on millisecond timescales in a cyclic process of contact-surface steam explosivity. The water subsequently reacts with the surface of the basaltic glass (sideromelane) to form an altered rind of palagonite. This reaction proceeds at a rate of 2.8 µm per year at 100°C (Jakobsson and Moore 1997), such that the tiny particles that make up the matrix of the hyaloclastite are altered entirely to palagonite in a relatively short time. Over time the unstable, amorphous palagonite evolves into more organized crystalline structures, ultimately transforming into the clay mineral smectite (Stroncik and Schminke, 2001).

Quantitatively, the B concentrations and δ¹¹B will be determined by the adsorption coefficient, the isotopic fractionation factor, and the cumulative water-rock ratio. To calculate the adsorption of boron onto clays in the hyaloclastite matrix, we use values for
the adsorption coefficient ($K_d$ = 2.6) and fractionation factor ($\alpha$ = 0.975) determined by Palmer et al. (1987) for adsorption of boron onto marine clays at 25°C and pH = 7.8. The primary clay mineral in marine clays is illite or mixed-layer illite-smectite, while the clay mineral in hyaloclastites and other altered basalts is smectite, thus we must assume for the purposes of this exercise that adsorption of B onto smectite surfaces is similar to adsorption onto illite surfaces. The isotope ratio in altered hyaloclastite can then be calculated as a function of water-rock ratio (W/R) using the approach of Spivack and Edmond (1987):

$$ \delta^{11}B_R = \alpha(\delta^{11}B_w + 10^3)\exp\left[\frac{K_d \cdot (1-\alpha)}{W/R}\right] - 10^3 \qquad \text{Eq. 1} $$

where $\delta^{11}B_R$ is the $\delta^{11}B$ in the rock, $\delta^{11}B_w$ is the $\delta^{11}B$ in the hydrothermal fluid. Analogously, the boron concentration in altered hyaloclastite is given by

$$ [B]_R = [B]_w \cdot K_d \cdot \exp\left(\frac{-K_d}{W/R}\right) \qquad \text{Eq. 2} $$

where $[B]_R$ is the boron concentration in the rock and $[B]_w$ is the boron concentration in the hydrothermal fluid. Model endmembers and results for adsorption of boron onto mineral surfaces in altered basaltic materials are presented in Table 2 and Figure 6. The approach described by Equations 1 and 2 above assumes that all of the boron in the altered rock is derived from the fluid, such that the starting composition of the basalt is inconsequential (that is, the fresh basalt composition could be MORB-like or OIB-like, and the result of this model would be the same). This assumption is valid assuming that the initial boron concentration in the rock is low. We find that modest water/rock ratios of 4-10 generate materials with boron concentrations and isotope ratios that are comparable to those of hydrothermally altered basalts from Reykjanes geothermal well RN-17
604 (Raffone et al., 2010), and make suitable endmembers for crustal assimilation modeling
605 (see Section 4.4). It should be noted that the lack of a modern hydrothermal system at
606 Lakagígar is of no consequence to this conceptual model, as we require simply that there
607 was hydrothermal activity at some time in the history of the Tertiary to Holocene basalts
608 underlying the fissure; this activity may well pre-date by quite some time the eruption of
609 1783-84.

610

611 **4.4 Crustal Assimilation Beneath Lakagígar**

612 Figure 7 presents the results of a parabolic mixing model, using as endmembers the
613 Lakagígar melt inclusion with the lowest boron content (12pl1a, δ¹¹B = -7.8‰, [B] =
614 0.59 ppm), two representative altered basalts from the Icelandic Deep Drilling Project
615 (one from the surface and one from 3000 m depth in well RN-17; Raffone et al., 2010),
616 and two model altered hyaloclastites onto which boron has been adsorbed from cooled
617 geothermal fluids at water/rock ratios of 4 and 10. Values of mixing endmembers are
618 given in Table 2. Given the existing data, we are unable to predict what the boron
619 concentrations and isotope ratios in a “pure” Icelandic plume melt might be. We have
620 chosen melt inclusion 12pl1a from Lakagígar to represent the background Icelandic
621 mantle source because it overlaps with the previous data from both Reykjanes and
622 Hengill (Gurenko and Chaussidon, 1997). Assuming this composition as representative
623 of the parental magma, the suite of melt inclusions sampled at Lakagígar reflect ~5-15%
624 assimilation of altered crustal materials. The matrix glasses, which represent the bulk of
625 the eruptive products, have assimilated ~15-30% hyaloclastite or altered basalt. These
626 values are consistent with previous estimates for crustal assimilation based on Sr and Th
627 isotope ratios (20%, Sigmarsson and Condomines, 1991), and are within the range of
628 estimates based on oxygen isotope ratios (10-60%; Bindeman et al., 2006, 2008).

629

630 **4.5 Depths of Entrapment**

631 Volatile equilibration pressures were determined using the solubility model of Papale et 632 al (2006), as shown in Figure 4f. The minimum equilibration pressures determined for 633 the melt inclusions on the basis of their measured H2O and CO2 concentrations range 634 from ~0.5 to ~1.5 kb. These pressures can be interpreted as entrapment pressures or, if 635 the magma storage time is sufficiently long for H2O and CO2 to re-equilibrate by 636 diffusion through the plagioclase host, as the pressure at which the magma was last stored 637 prior to eruption. When volatiles in melt inclusions re-equilibrate subsequent to 638 entrapment, they almost certainly do so at a shallower level (lower pressure) than that at 639 which they were originally trapped. Thus we can reasonably interpret the H2O-CO2 640 equilibration pressures as representing minimum depths of entrapment, recognizing that 641 the true depths of entrapment may be greater. The volatile equilibration pressures 642 determined here are equal to or less than cotectic (olivine-plagioclase-clinopyroxene) 643 crystallization pressures previously determined for Lakagígar (1.45 kbar; Kelley and 644 Barton, 2008), and considerably less than the maximum crystallization pressures (2.5 645 kbar, Guilbaud et al., 2007; 7.66 kbar, Kelley and Barton, 2008), suggesting that the H2O 646 and CO2 in at least some of the melt inclusions may have re-equilibrated by diffusion 647 during magma storage and transport. Alternatively, the H2O-CO2 variation in Figure 4f 648 could be interpreted to reflect mixing of relatively volatile-rich and volatile-poor 649 components at under-saturated conditions (in which case the depths of mixing would
have to be *greater* than the minimum entrapment depths). Overall, the model equilibration pressures correspond to minimum depths of ~3-4 km, assuming a density of 2.5-3.0 g/cm³ in the overlying crust. These depths are consistent with estimates of the depth to the magma chamber beneath Grímsvötn based on gravity and magnetic methods (1.5-4.0 km; Gudmundsson and Milsom, 1997), geodetic methods (>1.6 km; Sturkell et al., 2003), and teleseismic P-wave delays (3-4 km; Alfaro et al., 2007).

Geophysical studies of Icelandic magmatic systems over the last two decades have increased our understanding of the storage and migration of magma (e.g. Sturkell et al. 2006). These studies indicate that magma migrates from Moho depths (~20 km) to shallower levels (~5 km). The latter also correlates with geodetic estimates of the thickness of the brittle crust along the axes of the neovolcanic zones (e.g., LaFemina et al., 2005; Arnadottir et al., 2009). The March - April 2010 eruption of Eyjafjallakjokull volcano was the end result of approximately two decades of geophysical unrest marked by magmatic intrusion imaged by seismic and geodetic studies (Sigmundsson et al., 2010 and references therein). These studies indicate three occurrences of magma intrusion into sills at depths of 4.0 - 6.5 km, where density and strength variations of crustal layers - for example between basaltic lava flows and hyaloclastite or tillite strata - can disrupt the vertical migration of magma and cause sill intrusion (e.g., Taisne and Jaupart, 2009).

Our analysis of volatile saturation depths for our plagioclase-hosted melt inclusions indicates equilibration depths of ~4 km, in agreement with geodynamically determined magma chamber depths and upper crustal thicknesses in the Eastern Volcanic Zone. The stalling of magma and horizontal intrusion of sills at contacts between units at these or
greater depths, perhaps just beneath the primary magma chamber, would generate the necessary conditions for assimilation of various similar, yet distinct, basaltic compositions, consistent with our observation of subtle, random variability in the major element compositions of the Lakagígar melt inclusions. Variability in B and Li concentrations and δ¹¹B could be overprinted atop the initial magmatic compositions of these units as a result of hydrothermal alteration prior to assimilation. Figure 8 illustrates the potential magma pathways and accumulation sites in the middle to upper crust beneath Grímsvötn and Lakagígar.

681

## 4.6 A Recycled Subducted Component in the Plume Source?

We must also consider an alternative explanation – namely, that the source of very low δ¹¹B in Icelandic magmas is recycled, subducted oceanic crust that has passed through the deep mantle prior to being returned to the surface as a component of the mantle plume presumed to exist beneath Iceland. Some basalts from central Iceland have been interpreted as representing a mantle plume source containing subducted, altered oceanic crust on the basis of isotopic and incompatible trace element characteristics (e.g., Chauvel and Hémond, 2000; Skovgaard et al., 2001; Gurenko and Chaussidon, 2002; MacPherson et al., 2005). Correlation of moderately isotopically depleted boron (δ¹¹B as low as -8‰) and oxygen (δ¹⁸O as low as 4.9‰) with radiogenic isotope ratios in basalts from the Azores has been interpreted to reflect the influence of recycled subducted crust in OIB sources (Turner et al. 2007). The relatively low B/K₂O (=1-4) and moderate K₂O wt.% (0.36-0.50) in the Lakagígar glasses suggests a dominantly OIB-like magmatic source. However, the positive correlation of B concentrations with B/K₂O (Figure 5a) is
inconsistent with mixing between a MORB-like and an OIB-like source, or with variable melting of a single OIB-like source. In contrast, the B vs. B/K₂O observed in some of the melt inclusions from the Hengill fissure swarm (Gurenko and Chaussidon, 1997) might be explained by MORB-OIB mixing.

Previous analysis of basaltic melt inclusions from the Reykjanes and Hengill fissure swarms in southwestern Iceland found that the inclusions were characterized by a more or less invariant δ¹¹B of -11.3 (2σ ±3.8‰) over a range of B concentrations from 0.26-1.42 ppm (Gurenko and Chaussidon, 1997). If we were to exclude the melt inclusions 12pl1a and 12pl1c (both hosted in the same phenocryst) from this study, the range of B, K₂O, and δ¹¹B in the remaining Lakagígar melt inclusions would be generally consistent with the previous data for Reykjanes and Hengill, extending to slightly lower δ¹¹B (Figure 5), and we might be more inclined to consider MORB-OIB mixing. Given the small size of our data set, however, there is no good justification for rejecting two out of six melt inclusions for which we have boron data. As long as we consider all of our data at face value, the trends we observe strongly suggest coherent B, Li, and δ¹¹B variation that is decoupled from mantle source variation or magmatic processes, and we believe this is most consistent with assimilation of B- and Li- enriched altered crustal materials.

## 4.7 Origin of Volatiles in Lakagígar Magma

If we assume that assimilation of altered crust occurred prior to the 1783-1784 eruption, then it is quite likely that volatile components were added during assimilation of altered materials in addition to boron and lithium. As mentioned previously, we are reluctant to
attach too much significance to the observed correlation between boron concentrations and volatile concentrations, as this correlation could potentially be an artifact of post-entrapment effects. Nonetheless, some of the Lakagígar glasses extend to compositions that are clearly enriched in isotopically light boron relative to typical oceanic basalts, and these melt inclusions tend to have higher volatile contents relative to those with lower boron concentrations and more typical isotope ratios. It has been demonstrated previously that Lakagígar lavas and melt inclusions are also enriched in isotopically light oxygen (Hattori and Muehlenbachs, 1982; Bindeman et al., 2006; 2008). Palagonites that form the matrix of hyaloclastites such as those that make up Laki Mountain, through which Lakagígar cuts, have variable but high volatile contents (10-20% H₂O; up to 10% CO₂; and up to 1% SO₂; Stroncik and Schminke, 2001). We suggest that the massive quantities of released volatiles that made the 1783-84 Lakagígar eruption so deadly could have originated in altered crustal materials, such as palagonite-rich hyaloclastites, that were assimilated by the magma prior to eruption. The low concentrations of volatiles in the matrix glasses indicate that the volatile components were largely degassed at the time of eruption, leaving behind only the isotopic record of the assimilated materials.

735

## 5. CONCLUSIONS

Basaltic glasses from the 1783-84 eruption of Lakagígar are characterized by light boron isotope ratios relative to typical oceanic basalts. We have shown that isotopically light boron in fresh basaltic glasses can be explained by subsurface assimilation of very light boron in altered Icelandic basalts, such as those found at depths up to 3 km in well RN-17 from the Reykjanes geothermal field (Raffone et al., 2010). We propose a hypothetical
model to explain the generation of high concentrations of light boron in altered Icelandic basalts by adsorption onto mineral surfaces from meteorically-derived hydrothermal waters. Mixing models indicate that the range and trend observed in B concentrations and δ¹¹B at Lakagígar can be produced by assimilation of up to 30% hydrothermally altered crustal materials. This percentage of assimilated material is in agreement with the findings of Sigmarsson and Condomines (1991) based on ²³⁰Th/²³²Th, [Th], and Th/U in Lakagígar lavas (20% assimilation), and of Bindeman et al. (2008) based on δ¹⁸O in phenocrysts (10-60% assimilation). We propose that the altered crustal component is added during magma migration through sills in the deeper levels of the crust, becoming a ubiquitous feature in the final stages of evolution and mixing in a magma chamber located at ~3-4 km depth, most likely that of the Grímsvötn central volcano. We tentatively suggest that the large mass of volatiles released during the Lakagígar eruption may likewise have originated in part from the assimilation of hyaloclastite and other altered crustal materials at depths of several kilometers beneath the fissure.

Acknowledgements. The authors wish to thank Mark Angelone and John Cantolina of the PSU Materials Characterization Laboratory (MCL) for their assistance with analyses, and Kari Kristjansson of the Lakagígar Skaftafell National Park for granting access to the tephra deposits. This work was supported by a PSU-EMS Miller Award to MDF, NSF EAR 09117456 to PCL, and awards to MNB by the GSA Northeast Division Grant for Undergraduate Research, the Dept. of Geosciences Undergraduate Senior Thesis Grant, and PSU Undergraduate Research Discovery Grant. This project was inspired by the Center for Advancement of Undergraduate Studies and Experience (CAUSE) field course
765 in Iceland, 2008, sponsored by the College of Earth and Mineral Sciences at PSU. This
766 manuscript was improved by thorough reviews from Stefan Arnorsson and two
767 anonymous reviewers, and by helpful comments from the editor, Steve Shirey.

768

769 REFERENCES

770 Aggarwal J., Palmer M., Bullen T., Arnorsson S., and Ragnarsdottir K. (2000) The boron
771 isotope systematic of Icelandic geothermal waters: 1. Meteoric water charged systems.
772 *Geochim. Cosmochim. Acta* **64**, 579-585.

773

774 Alfaro R., Brandsdottir B., Rowlands D., and White R. (2007) Structure of the Grímsvötn
775 central volcano under the Vatnajokull icecap, Iceland. *Geophys. J. Int.* **168** 863-876.

776

777 Arnadottir T., Lund B., Jiang W., Geirsson H., Bjornsson H., Einarsson P., and
778 Sigurdsson T. (2009) Glacial rebound and plate spreading: Results from the first
779 countrywide GPS observations in Iceland. *Geophys. J. Int.* **177** 691-716.

780

781 Arnórsson S. and Andresdottir A. (1995) Processes controlling the distribution of boron
782 and chlorine in natural waters in Iceland. *Geochim. Cosmochim. Acta* **59**, 4125-4146.

783

784 Bell J.D. and Humphries D.J. (1972) Lakagígar fissure eruption. In: *Progress in
785 Experimental Petrology*, Nat. Environ. Res. Council, UK Publication Series D2, 110-112.

786
787 Benton L., Tomascak P., and Helz R. (1999) A study of boron isotopes in Kilauea Iki lava lake, Hawaii. *EOS Trans. AGU* 80, Fall Meeting Supplement, Abstract V52B-06.

789
790 Bindeman I., Davis A., and Drake M. (1998) Ion microprobe study of plagioclase-basalt partition experiments at natural concentration levels of trace elements. *Geochim. Cosmochim. Acta*, **62**, 1175-1193.

793
794 Bindeman I., Gurenko A., Sigmarsson O., and Chaussidon M. (2008) Oxygen isotope heterogeneity and disequilibria of olivine crystals in large volume Holocene basalts from Iceland: Evidence for magmatic digestion and erosion of Pleistocene hyaloclastites. *Geochim. Cosmochim. Acta* **72**, 4397-4420.

798
799 Bindeman I., Sigmarsson O., and Eiler J. (2006) Time constraints on the origin of large volume basalts derived from O-isotopes and trace element mineral zoning and U-series disequilibria in the Lakagígar and Grimsvötn volcanic system. *Earth Planet. Sci. Lett.* **245**, 245-249.

803
804 Chaussidon M. and Jambon A. (1994) Boron content and isotopic composition of oceanic basalts: Geochemical and cosmochemical implications. *Earth Planet. Sci. Lett.* **121**, 277-291.

807
808 Chaussidon M. and Marty B. (1995) Primitive boron isotope composition of the mantle. *Science* **269**, 383-386.
810

811

Chaussidon M., Robert F., Mangin D., Hanon P., and Rose E. (1997) Analytical procedures for the measurement of boron isotope compositions by ion microprobe in meteorites and mantle rocks. *J. Geostand. Geoanal. Res.* **21**, 7-17.

812

813

814

815

Chauvel C., and Hémond C. (2000) Melting of a complete section of recycled oceanic crust: trace element and Pb isotopic evidence from Iceland. *Geochem. Geophys. Geosys.* **1**, doi:10.1029/1999GC000002..

816

817

818

819

Darbyshire F., White R., Priestley K. (2000) Structure of the crust and uppermost mantle of Iceland from a combined seismic and gravity study. *Earth Planet. Sci. Lett.* **181**, 409-428.

820

821

822

823

Einarsson P. and Sæmundsson K. (1992) Earthquake epicenters 1982–1985 and volcanic systems in Iceland (map). In: *Í hlutarins eðli*. Festschrift for Þorbjörn Sigurgeirsson, ed. Þorsteinn Sigfússon, Menningarsjóður, Reykjavík.

824

825

826

827

Gudmundsson M. and Milsom J. (1997) Gravity and magnetic studies of the subglacial Grímsvötn Volcano, Iceland; implications for crustal and thermal structure. *J. Geophys. Res.* **102**.

828

829

830
831 Guilbaud M., Blake S., Thordarson T., and Self S. (2007) Role of syn-eruptive cooling and degassing on textures of lavas from the AD 1783-1784 Lakagígar eruption, south Iceland. *J. Petrol.* **48**, 1265-1294.

835 Gurenko A. and Chaussidon M. (1997) Boron concentrations and isotopic compositions of the Icelandic mantle: evidence from glass inclusions in olivine. *Chem. Geol.* **135**, 21-34.

839 Gurenko A. and Chaussidon M. (2002) Oxygen isotope variations in primitive tholeiites of Iceland: evidence from a SIMS study of glass inclusions, olivine phenocrysts and pillow rim glasses. *Earth Planet Sci. Lett.* **205**, 63-79.

843 Hamada M. and Fujii T. (2007) H2O-rich island arc low-K tholeiite magma inferred from Ca-rich plagioclase-melt inclusion equilibria. *Geochem. J.* **41**, 437-461.

846 Hattori K., and Muehlenbachs K. (1982) Oxygen Isotope Ratios of the Icelandic Crust. *J. Geophys. Res.* **87**, 6559-6565.

849 Hauri E., Wang J., Dixon J.E., King P.L, Mandeville C., and Newman S. (2002) SIMS analysis of volatiles in silicate glasses 1: Calibration, matrix effects and comparisons with FTIR. *Chem. Geol.* **183**, 99-114.

853 Helo C., Longpré M-A., Shimizu N., Clague D.A., and Stix J. (2011) Explosive eruptions at mid-ocean ridges driven by CO₂-rich magmas. *Nature Geosc.* **4**, 260-263.
855

856

Hémond C., Condomines M., Fourcade S., Allegre C., Oskarsson N., and Javoy M.

857

(1988) Thorium, strontium and oxygen isotopic geochemistry in recent tholeiites from

858

Iceland: crustal influence on mantle-derived magmas. *Earth Planet Sci. Lett.* **87**, 273-

859

285.

860

861

Hervig R., Moore G., Williams L., Peacock S., Holloway J., and Roggensack K. (2002)

862

Isotopic and elemental partitioning of boron between hydrous fluid and silicate melt. *Am.*

863

*Mineral.* **87**, 769-774.

864

865

Jakobsson S. and Moore J. (1997) Hydrothermal minerals and alteration rates at Surtsey

866

volcano, Iceland. *Geol. Soc. Am. Bull.* **97**, 648-659.

867

868

Jochum K.P., Dingwell D.B., Rocholl A., Stoll B., Hofmann A.W., Becker S., Besmehn

869

A., Bessett D., Dietze H.J., Dulski P., Erzinger J., Hellebrand E., Hoppe P., Horn I.,

870

Janssens K., Jenner G.A., Klein M., McDonough W.F., Maetz M., Mezger K., Münker

871

C., Nikogosian I.K., Pickhardt C., Raczek I., Rhede D., Seufert H.M., Simakin S.G.,

872

Sobolev A.V., Spettel B., Straub S., Vincze L., Wallianos A., Weckwerth G., Weyer S.,

873

Wolf D., and Zimmer M. (2000) The preparation and preliminary characterization of

874

eight geological MPI-DING reference glasses for in-situ microanalysis. *Geostand.*

875

*Newslett.* **24**, 87-133.

876
877 Kelley D. and Barton M. (2008) Pressures of crystallization of Icelandic magmas. *J. Petrol.* **49**, 465-492.

878 Kobayashi K., Tanaka R., Moriguti T., Shimizu K., Nakamura E. (2004) Lithium, boron, and lead isotope systematics of glass inclusions in olivines from Hawaiian lavas: evidence for recycled components in the Hawaiian plume. *Chem. Geol.* **212**, 143-161.

879
880
881
882
883
884
885
886
887
888
889
890
891
892
893
894
895
896
897
898
Kress V. and Ghiorso M. (2004) Thermodynamic modeling of post-entrapment crystallization in igneous phases. *J. Volcanol. Geotherm. Res.* **137**, 247-260.

LaFemina P.C., Dixon T. H., Malservisi R., Árnadóttir E., Sturkell E., Sigmundsson F., and Einarsson P. (2005) Geodetic GPS measurements in south Iceland: Strain accumulation and partitioning in a propagating ridge system. *J. Geophys. Res.* **110**, 1-21.

MacPherson C., Hilton D., Day J., Lowry D., and Gronvold K. (2005) High ³He/⁴He, depleted mantle and low δ¹⁸O recycled lithosphere in the source of central Iceland magmatism. *Earth Planet Sci. Lett.* **233**, 411-427.

Marks N., Schiffman P., Zierenberg R.A., Franzson H., Fridleifsson G.Ó. (2010) Hydrothermal alteration in the Reykjanes geothermal system: Insights from Iceland deep drilling program well RN-17. *J. Volcanol. Geotherm. Res.* **189**, 172-190.
899 Métrich N., Sigurdsson H., Meyer P., and Devine J. (1991) The 1783 Laki eruption in Iceland: geochemistry, CO₂ and sulfur degassing. *Contrib. Mineral. Petrol.* **107**, 435-447.

901
902 Papale P., Moretti R. and Barbato D. (2006) The compositional dependence of the
903 saturation surface of H₂O + CO₂ fluids in silicate melts. *Chem. Geol.* **229**, 78-95.

904
905 Palmer M., Spivack A., and Edmond J. (1987) Temperature and pH controls over isotopic
906 fractionation during adsorption onto marine clays: *Geochim. Cosmochim. Acta*, **51**, 2319-
907 2323.

908
909 Pope E., Bird D., Arnórsson S., Fridriksson T., Elders W.A., Fridleifsson G.Ó. (2009)
910 Isotopic constraints on ice age fluids in active geothermal systems: Reykjanes, Iceland.
911 *Geochim. Cosmochim. Acta* **73**, 4468-4488.

912
913 Pouchou JL. and Pichoir F. (1987) Basic expression of “PAP” computation for
914 quantitative EPMA. In: *Eleventh International Congress on X-Ray Optics and
915 Microanalysis*, Ontario University Press, 249-252.

916
917 Raffone N., Ottolini L.P., Tonarini S., Gianelli G., D-Orazio M. and Fridleifsson G.
918 (2010) An investigation of trace and isotope light elements in mineral phases from well
919 RN-17 (Reykjanes Peninsula, SW Iceland). In: *11<sup>th</sup> European Workshop on Modern
920 Developments and Applications in Microbeam Analysis, IOP Conf. Series: Materials
921 *Science and Engineering* **7**, 012026.
922

923

Rose-Koga E. and Sigmarsson O. (2008) B-O-Th isotope systematics in Icelandic tephra.

924

*Chem. Geol.* **255**, 454-462.

925

926

Rosner M. and Meixner A. (2004) Boron Isotopic Composition and Concentration

927

of Ten Geological Reference Materials. *Geostand. Geoanal. Res.* **28**,431-441.

928

929

Rosner M., Wiedenbeck M., and Ludwig T. (2008) Composition-induced variations in

930

SIMS instrumental mass fractionation during boron isotope ratio measurements of silicate

931

glasses. *Geostand. Geoanal. Res.* **32**, 27-38.

932

933

Ryan J. and Langmuir C. (1993) The systematics of boron abundances in young volcanic

934

rocks. *Geochim. Cosmochim. Acta* **57**, 1489-1498.

935

936

Ryan J., Leeman W., Morris J., Langmuir C. (1996) The boron systematics of intraplate

937

lavas: Implications for crust and mantle evolution. *Geochim. Cosmochim. Acta* **60**, 415-

938

422.

939

940

Saal A., Hart S., Shimizu N., Hauri E., Layne G. (1998) Pb isotopic variability in melt

941

inclusions from oceanic island basalts, Polynesia. *Science* **282**, 1481-1484.

942

943

Sella G., Dixon T., Mao A. (2002), REVEL: A model for recent plate velocities from

944

space geodesy, *J. Geophys. Res.* **107(B4)**, 2081, doi:10.1029/2000JB000033.
945

946

Shimizu K., Shimizu N., Komiya K., Suzuki, K., Maruyama S. and Tatsumi Y. (2009)

947

CO₂-rich komatiitic melt inclusions in Cr-spinels within beach sand from Gorgona Island,

948

Columbia. *Earth Planet. Sci. Lett.* **299**, 33-43.

949

950

Sigmarsson O., and Condomines M. (1991) Extreme magma homogeneity in the 1783-84

951

Lakagígar eruption: origin of large volume of evolved basalt in Iceland. *Geophys. Res.*

952

*Lett.* **18**, 2229-2232.

953

954

Sigmundsson F., Hreinsdottir S., Hooper A., Arnadottir T., Pedersen R., Roberts M.,

955

Oskarsson N., Auriac A., Decriem J., Einarsson P., Geirsson H., Hensch M., Ofeigsson

956

B., Sturkell E., Sveinbjornsson H., and Feigl K. (2010) Intrusion triggering of the 2010

957

Eyjafjallajokull explosive eruption. *Nature* **468**, 426-430.

958

959

Sigurdsson H. and Sparks S. (1978) Rifting Episode in North Iceland in 1874-1875 and

960

the Eruptions of Askja and Sveinagja. *Bull. Volcanol.* **41**, 149-167.

961

962

Skovgaard A.C., Storey M., Baker J., Blusztajn J., and Hart S.R. (2001) Osmium-oxygen

963

isotopic evidence for a recycled and strongly depleted component in the Iceland mantle

964

plume. *Earth Planet. Sci. Lett.* **194**, 259-275.

965

966

Smith H., Spivack A., Staudigel H., and Hart S. (1995) The boron isotopic composition

967

of altered oceanic crust. *Chem. Geol.* **126**, 119-135.
968

969

Spivack A. and Edmond J. (1986) Determination of boron isotope ratios by thermal ionization mass spectrometry of the dicesium metaborate cation. *Anal. Chem.* **58**, 31-35.

970

971

972

Spivack A. and Edmond J. (1987) Boron isotope exchange between seawater and the oceanic crust. *Geochim. Cosmochim. Acta*, **51**, 1033-1043.

973

974

975

Stroncik N., and Schminke H. (2001) Evolution of palagonite: Crystallization, chemical changes and elemental budget. *Geochem. Geophys. Geosys.* **2**, 1-43.

976

977

978

Sturkell E., Einarsson P., Sigmundsson F., Hreinsdottir F., and Geirsson H. (2003) Deformation of Grímsvötn volcano, Iceland: 1998 eruption and subsequent inflation. *Geophys. Res. Lett.* **30**, 1182.

979

980

981

982

Sturkell E., Einarsson P., Sigmundsson F., Geirsson H., Olafsson H., Pedersen R., de Zeeuw-van Dalfsen E., Linde A.T., Sacks I.S., and Stefansson R. (2006) Volcano geodesy and magma dynamics in Iceland. *J. Volcanol. Geotherm. Res.* **150**, 14-34.

983

984

985

986

Sun S-S. and McDonough W. (1995) Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes. In: Saunders A, Norry M (eds) Magmatism in ocean basins. *Geol. Soc. Spec. Publ.* **42**, 313-345.

987

988

989
990 Taisne B. and Jaupart C. (2009) Dike propagation through layered rocks. *J. Geophys.*

991 *Res.* **114**.

992

993 Thorarinsson S. (1979) On the damage caused by volcanic eruptions with special

994 reference to tephra and gases. In: *Volcanic Activity and Human Ecology*, 125-159.

995

996 Thordarson T., and Self S. (1993) The Laki (Skaftár Fires) and Grímsvötn eruptions in

997 1783-1785. *Bull. Volcanol.* **55**, 233-263.

998

999 Thordarson T., Self S., Óskarsson N., and Hulsebosch T. (1996) Sulfur, chlorine and

1000 fluorine degassing and atmospheric loading by the 1784-1784 AD Laki (Skaftár Fires)

1001 eruption in Iceland. *Bull. Volcanol.* **58**, 205-225.

1002

1003 Tilling R.I. (1996) Hazards and climatic impact of subduction zone volcanism: A global

1004 and historical perspective. In: *Subduction Top to Bottom. Geophysical Monograph* **96**,

1005 American Geophysical Union.

1006

1007 Turner S., Tonarini S., Bindeman I., Leeman W., and Schaefer B. (2007) Boron and

1008 oxygen isotope evidence for recycled components over the past 2.5 Gyr. *Nature* **447**,

1009 702-705.

1010

1011
1012

1013 **FIGURES**

1014 **Figure 1:** a) Map of Iceland. NA = North American Plate; EU = European Plate; EVZ = Eastern Volcanic Zone; WVZ = Western Volcanic Zone; NVZ = Northern Volcanic Zone. Contours are depth to Moho (km) from Darbyshire et al., 2000. The 40 km contour is assumed to correspond to the center of the Icelandic plume. Dashed box delineates the area expanded in Figure 1b. The Reykjanes, Hengill, and Lakagígar fissure swarms are marked with solid stars. Shaded gray areas are fissure swarm locations from Einarsson and Sæmundsson, 1992. Spreading rate of 19.8 mm/yr from Sella et al., 2002.

1021 b) Location of Lakagígar (Laki Fissure) in southeastern Iceland. The grey region marks the extent of lava flows from the 1783-1784 eruption. Inverted triangles indicate sample locations, and open circles indicate locations of two tuff cones presumed to be the sources of the phreatomagmatic tephra deposits.

1025 **Figure 2:** Examples of basalt glass types analyzed. a) Transmitted light photomicrograph of melt inclusions in crystal 14pl2; b) Secondary electron image of a tephra clast (LG-0808-8.2) in cross section, showing the locations of plagioclase microlites.

1028 **Figure 3:** Selected major and trace element concentrations and δ¹¹B in Lakagígar melt inclusions and host glasses. a) Mg/(Mg+Fe)×100 vs. TiO₂ wt.%; b) Mg/(Mg+Fe)×100 vs. CaO/Al₂O₃; c) (Na₂O+K₂O) wt.% vs. TiO₂ wt.%; d) (Na₂O+K₂O) wt.% vs. B ppm; e) Li ppm vs. B ppm; f) δ¹¹B vs. B ppm. Selected “end-member” melt inclusions for various elemental concentrations, ratios, or δ¹¹B values are labeled. Error bars are average 2σ external reproducibilities based on replicate analysis of unknowns; internal errors are comparable. Solid error bars apply to analyses performed using the Cameca
ims 1280, and dashed error bars apply to analyses performed using the Cameca ims 3f.

In general, melt inclusions were analyzed by 1280 and host glasses were analyzed by 3f.

See Table 1 for actual uncertainties for each data point.

**Figure 4:** Volatile concentrations plotted against boron concentrations in melt inclusions (closed symbols) and host glasses (open symbols). a) H₂O wt.% vs. B ppm; b) CO₂ ppm vs. B ppm; c) S ppm vs. B ppm; d) F ppm vs. B ppm; e) Cl ppm vs. B ppm; f) H₂O wt.% vs. CO₂ ppm. Isobars on the plot of H₂O vs. CO₂ represent minimum entrapment pressures, and are calculated using the volatile saturation model of Papale et al. (2006).

Error bars are average 2σ external reproducibilities based on replicate analysis of unknowns; internal errors are about an order of magnitude less. Solid error bars apply to analyses performed using the Cameca ims 1280, and dashed error bars apply to analyses performed using the Cameca ims 3f. In general, melt inclusions were analyzed by 1280 and host glasses were analyzed by 3f (for B concentrations only – all volatiles by 1280).

See Table 1 for actual uncertainties for each data point.

**Figure 5:** B/K₂O, B, and δ¹¹B in scoria matrix glasses (open symbols) and melt inclusions (closed symbols). a) B vs. B/K₂O; b) B vs. δ¹¹B. Lakagígar glasses (this study) are shown as circles, and glasses from the Hengill and Reykjanes fissure swarms (Gurenko and Chaussidon, 1997) are shown as squares and triangles, respectively, for comparison. Error bars represent 2σ uncertainty. When replicate analysis was possible, errors shown represent external reproducibility, otherwise errors represent internal precision. For boron concentrations and isotope ratios, internal and external uncertainties were essentially the same. Dashed error bars indicate boron concentrations determined by Cameca ims 3f, all other data collected by Cameca ims 1280. Medium gray field
represents range of OIB from the literature: Chaussidon and Jambon, 1994; Chaussidon and Marty, 1995; Ryan et al., 1996; Turner et al., 2007. Dark gray field represents range of N-MORB from Chaussidon and Jambon, 1994.

**Figure 6:** Effects of water-rock interaction on boron concentrations and isotope ratios in mid-ocean ridge basalts (MORB) and meteoric water. Solid lines with arrows indicate trend of low-temperature basalt alteration by seawater (upper trend) or adsorption of boron from cooled, glacially-derived, geothermal fluids (lower trend). Tick marks on lower trend indicate the water/rock ratio during adsorption. Alteration of OIB or “typical” Icelandic basalt would follow a similar trend starting from slightly higher [B] and lower $\delta^{11}$B. Dashed line with arrow depicts the trend of boron enrichment in meteoric (glacial) water as a result of high-temperature water-rock interaction. Closed symbols represent basalts, open symbols represent waters. Data sources given in Table 2.

**Figure 7:** Results of parabolic mixing model using the Lakagígar melt inclusion with the lowest B concentration (12pl1a) as a starting point, and either model altered hyaloclastite (this study) or actual altered Icelandic basalts as endmembers. Model altered hyaloclastite compositions are calculated assuming that B is adsorbed onto the surface of palagonite during hydrothermal circulation at water-rock ratios of 4 and 10. Actual basalt compositions are whole rock analyses of samples collected at the surface and at 3000m depth in geothermal well RN-17, Reykjanes geothermal field (Raffone et al., 2010). Endmember values for mixing model are given in Table 2. Tick marks indicate 5% or 10% increments of altered crust assimilation. Closed symbols represent melt inclusions, open symbols represent matrix glasses. Error bars are average $2\sigma$ external reproducibilities based on replicate analysis of unknowns; internal errors are comparable.
Solid error bars apply to analyses performed using the Cameca IMS 1280, and dashed error bars apply to analyses performed using the Cameca IMS 3f. In general, melt inclusions were analyzed by 1280 and host glasses were analyzed by 3f. See Table 1 or Figure 5b for actual uncertainties for each data point.

**Figure 8:** Illustration of the magma plumbing system beneath Lakagígar showing the development of sills at boundaries of density contrast (i.e., between lava flows and hyaloclastites) at mid- to deep-crustal levels beneath the Grímsvötn magma chamber. Assimilation of hydrothermally altered materials within these sills contributes isotopically light boron to the magma, accompanied by subtle variations in major element compositions. Mixing in the magma chamber beneath the central volcano homogenizes the magma prior to the fissure eruption.

1092
<figure><img src="image_2.png" /></figure>

<figure><img src="image_3.png" /><figcaption>Figure 1</figcaption></figure>
<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /><figcaption>Figure 2</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Figure 3</figcaption></figure>
<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /></figure>

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /><figcaption>Figure 4</figcaption></figure>
<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>
<figure><img src="image_15.png" /><figcaption>Figure 6</figcaption></figure>
<figure><img src="image_16.png" /><figcaption>Figure 7</figcaption></figure>
<figure><img src="image_17.png" /><figcaption>Figure 8</figcaption></figure>
<table><caption>**Table 1.** Major elements, B concentrations, δ¹¹B, and volatile contents in plagioclase-hosted melt inclusions and matrix glasses</caption><thead><tr><th rowspan="2">Sample #<br>Crystal/incl.</th><th colspan="3">Matrix Glass</th><th colspan="3">Melt Inclusions</th></tr><tr><th>LG-0808-8.2</th><th>LG-0808-8.4</th><th>LG-0808-12</th><th>LG-0808-14</th><th>LG-0808-17</th><th>LG-0808-8<br>pIa</th><th>LG-0808-12<br>pIa</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>50.01</td><td>50.01</td><td>49.32</td><td>49.11</td><td>49.86</td><td>48.42</td><td>49.15</td></tr><tr><td>TiO<sub>2</sub></td><td>2.93</td><td>2.89</td><td>3.02</td><td>2.94</td><td>3.00</td><td>2.45</td><td>2.27</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>12.84</td><td>12.93</td><td>12.51</td><td>12.90</td><td>12.92</td><td>12.53</td><td>12.31</td></tr><tr><td>FeO*</td><td>13.45</td><td>13.82</td><td>14.10</td><td>13.93</td><td>14.04</td><td>13.54</td><td>13.93</td></tr><tr><td>MnO</td><td>0.23</td><td>0.20</td><td>0.19</td><td>0.22</td><td>0.26</td><td>0.21</td><td>0.23</td></tr><tr><td>MgO</td><td>5.32</td><td>5.24</td><td>5.15</td><td>5.29</td><td>5.26</td><td>5.97</td><td>6.09</td></tr><tr><td>CaO</td><td>9.36</td><td>9.52</td><td>9.34</td><td>9.28</td><td>9.53</td><td>9.71</td><td>9.76</td></tr><tr><td>Na<sub>2</sub>O</td><td>2.61</td><td>2.63</td><td>2.65</td><td>2.60</td><td>2.69</td><td>2.58</td><td>3.01</td></tr><tr><td>K<sub>2</sub>O</td><td>0.42</td><td>0.41</td><td>0.45</td><td>0.41</td><td>0.42</td><td>0.45</td><td>0.50</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>0.25</td><td>0.27</td><td>0.31</td><td>0.26</td><td>0.25</td><td>0.28</td><td>0.12</td></tr><tr><td>Total</td><td>97.42</td><td>97.91</td><td>97.04</td><td>96.94</td><td>98.23</td><td>96.14</td><td>97.37</td></tr><tr><td>Mg# in glass</td><td>41</td><td>40</td><td>39</td><td>40</td><td>40</td><td>44</td><td>44</td></tr><tr><td>An% in host</td><td>---</td><td>---</td><td>---</td><td>---</td><td>---</td><td>83</td><td>85</td></tr><tr><td>B (ppm)</td><td>1.65<sup>a</sup></td><td>1.39<sup>a</sup></td><td>1.43</td><td>1.31<sup>a</sup></td><td>1.25<sup>a</sup></td><td>1.25</td><td>0.59</td></tr><tr><td>2σ<sup>b</sup></td><td>0.45 (n=1)</td><td>0.31 (n=2)</td><td>0.10 (n=2)</td><td>0.37 (n=1)</td><td>0.22 (n=2)</td><td>0.11 (n=1)</td><td>0.10 (n=1)</td></tr><tr><td>Li (ppm)</td><td>10.5<sup>a</sup></td><td>8.20<sup>a</sup></td><td>6.81</td><td>10.7<sup>a</sup></td><td>9.70<sup>a</sup></td><td>6.52</td><td>3.31</td></tr><tr><td>2σ<sup>b</sup></td><td>1.4 (n=1)</td><td>1.57 (n=2)</td><td>0.40 (n=2)</td><td>1.57 (n=1)</td><td>0.38 (n=2)</td><td>0.45 (n=1)</td><td>0.44 (n=1)</td></tr><tr><td>δ<sup>11</sup>B ‰</td><td>-12.9</td><td>nd</td><td>-10.5</td><td>-14.7</td><td>-16.9</td><td>-16.5</td><td>-7.8</td></tr><tr><td>2σ<sup>b</sup></td><td>3.0 (n=1)</td><td>---</td><td>3.4 (n=1)</td><td>2.5 (n=2)</td><td>4.1 (n=2)</td><td>3.6 (n=1)</td><td>4.2 (n=1)</td></tr><tr><td>H<sub>2</sub>O%</td><td>0.13</td><td>0.12</td><td>0.13</td><td>0.13</td><td>0.14</td><td>0.80</td><td>0.38</td></tr><tr><td>2σ<sup>b</sup></td><td>0.00 (n=1)</td><td>0.03 (n=2)</td><td>0.01 (n=2)</td><td>0.06 (n=2)</td><td>0.07 (n=4)</td><td>0.07 (n=2)</td><td>0.03 (n=2)</td></tr><tr><td>CO<sub>2</sub> (ppm)</td><td>5.9</td><td>5.2</td><td>4.2</td><td>5.1</td><td>5.2</td><td>504</td><td>298</td></tr><tr><td>2σ<sup>b</sup></td><td>0.8 (n=1)</td><td>2.2 (n=2)</td><td>1.2 (n=1)</td><td>2.4 (n=1)</td><td>5.2 (n=3)</td><td>44 (n=2)</td><td>93 (n=2)</td></tr><tr><td>S (ppm)</td><td>407</td><td>414</td><td>436</td><td>406</td><td>479</td><td>1759</td><td>1169</td></tr><tr><td>2σ<sup>b</sup></td><td>3 (n=1)</td><td>31 (n=2)</td><td>71 (n=2)</td><td>21 (n=2)</td><td>185 (n=4)</td><td>247 (n=2)</td><td>87 (n=2)</td></tr><tr><td>Cl (ppm)</td><td>155</td><td>161</td><td>172</td><td>162</td><td>152</td><td>205</td><td>147</td></tr><tr><td>2σ<sup>b</sup></td><td>1 (n=1)</td><td>6 (n=2)</td><td>26 (n=2)</td><td>1 (n=2)</td><td>27 (n=4)</td><td>32 (n=2)</td><td>11 (n=2)</td></tr><tr><td>F (ppm)</td><td>513</td><td>519</td><td>549</td><td>471</td><td>476</td><td>592</td><td>508</td></tr><tr><td>2σ<sup>b</sup></td><td>3 (n=1)</td><td>21 (n=2)</td><td>55 (n=2)</td><td>33 (n=2)</td><td>75 (n=4)</td><td>57 (n=2)</td><td>39 (n=2)</td></tr></tbody></table>
<table><caption>Table 1 cont'd. Major elements, B concentrations, δ<sup>11</sup>B, and volatile contents in plagioclase-hosted melt inclusions and matrix glasses</caption><thead><tr><th rowspan="2">Sample #<br>Crystal/incl.</th><th colspan="8">Melt Inclusions</th></tr><tr><th>LG-0808-12<br>p11c</th><th>LG-0808-14<br>p11a</th><th>LG-0808-14<br>p12a</th><th>LG-0808-14<br>p12c</th><th>LG-0808-14<br>p12d</th><th>LG-0808-17<br>p13a</th><th>LG-0808-17<br>p13b</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>50.60</td><td>49.20</td><td>48.74</td><td>50.08</td><td>49.90</td><td>49.00</td><td>49.61</td></tr><tr><td>TiO<sub>2</sub></td><td>1.95</td><td>2.85</td><td>2.14</td><td>1.85</td><td>2.38</td><td>1.73</td><td>2.98</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>12.68</td><td>13.01</td><td>12.87</td><td>12.86</td><td>12.76</td><td>12.68</td><td>12.34</td></tr><tr><td>FeO*</td><td>13.27</td><td>13.06</td><td>13.06</td><td>13.34</td><td>13.24</td><td>12.15</td><td>13.52</td></tr><tr><td>MnO</td><td>0.26</td><td>0.21</td><td>0.18</td><td>0.21</td><td>0.24</td><td>0.22</td><td>0.24</td></tr><tr><td>MgO</td><td>6.36</td><td>5.68</td><td>5.56</td><td>5.81</td><td>5.82</td><td>7.09</td><td>5.79</td></tr><tr><td>CaO</td><td>9.66</td><td>9.57</td><td>9.62</td><td>9.55</td><td>9.54</td><td>10.82</td><td>9.41</td></tr><tr><td>Na<sub>2</sub>O</td><td>3.29</td><td>2.67</td><td>2.51</td><td>2.81</td><td>2.70</td><td>2.34</td><td>2.59</td></tr><tr><td>K<sub>2</sub>O</td><td>0.44</td><td>0.42</td><td>0.42</td><td>0.49</td><td>0.49</td><td>0.36</td><td>0.47</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>0.17</td><td>0.25</td><td>0.22</td><td>0.14</td><td>0.24</td><td>0.18</td><td>0.35</td></tr><tr><td>Total</td><td>98.68</td><td>96.92</td><td>95.32</td><td>97.14</td><td>97.31</td><td>96.57</td><td>97.30</td></tr><tr><td>Mg# in glass</td><td>46</td><td>44</td><td>43</td><td>44</td><td>44</td><td>51</td><td>43</td></tr><tr><td>An% in host</td><td>85</td><td>76</td><td>84</td><td>84</td><td>84</td><td>88</td><td>88</td></tr><tr><td>B (ppm)</td><td>nd</td><td>1.02<sup>a</sup></td><td>1.11</td><td>nd</td><td>nd</td><td>0.96</td><td>0.79</td></tr><tr><td>2σ<sup>b</sup></td><td>---</td><td>0.34 (n=1)</td><td>0.07 (n=1)</td><td>---</td><td>---</td><td>0.14 (n=1)</td><td>0.08 (n=1)</td></tr><tr><td>Li (ppm)</td><td>nd</td><td>8.81<sup>a</sup></td><td>7.08</td><td>nd</td><td>nd</td><td>4.36</td><td>5.53</td></tr><tr><td>2σ<sup>b</sup></td><td>---</td><td>1.42 (n=1)</td><td>0.41 (n=1)</td><td>---</td><td>---</td><td>0.21 (n=1)</td><td>0.41 (n=1)</td></tr><tr><td>δ<sup>11</sup>B ‰</td><td>nd</td><td>-13.9</td><td>-10.6</td><td>nd</td><td>nd</td><td>-9.0</td><td>nd</td></tr><tr><td>2σ<sup>b</sup></td><td>---</td><td>4.9 (n=1)</td><td>4.6 (n=1)</td><td>---</td><td>---</td><td>5.4 (n=1)</td><td>---</td></tr><tr><td>H<sub>2</sub>O%</td><td>0.54</td><td>nd</td><td>0.83</td><td>0.80</td><td>0.84</td><td>0.71</td><td>0.67</td></tr><tr><td>2σ<sup>b</sup></td><td>0.01 (n=1)</td><td>---</td><td>0.01 (n=1)</td><td>0.01 (n=1)</td><td>0.01 (n=1)</td><td>0.02 (n=1)</td><td>0.04 (n=2)</td></tr><tr><td>CO<sub>2</sub> (ppm)</td><td>220</td><td>nd</td><td>333</td><td>303</td><td>178</td><td>469</td><td>247</td></tr><tr><td>2σ<sup>b</sup></td><td>28 (n=1)</td><td>---</td><td>10 (n=1)</td><td>8 (n=1)</td><td>6 (n=1)</td><td>12 (n=1)</td><td>102 (n=2)</td></tr><tr><td>S (ppm)</td><td>1497</td><td>nd</td><td>1713</td><td>1587</td><td>1668</td><td>1273</td><td>1261</td></tr><tr><td>2σ<sup>b</sup></td><td>14 (n=1)</td><td>---</td><td>14 (n=1)</td><td>12 (n=1)</td><td>16 (n=1)</td><td>17 (n=1)</td><td>86 (n=2)</td></tr><tr><td>Cl (ppm)</td><td>203</td><td>nd</td><td>195</td><td>189</td><td>188</td><td>138</td><td>138</td></tr><tr><td>2σ<sup>b</sup></td><td>7 (n=1)</td><td>---</td><td>2 (n=1)</td><td>2 (n=1)</td><td>3 (n=1)</td><td>3 (n=1)</td><td>11 (n=2)</td></tr><tr><td>F (ppm)</td><td>600</td><td>nd</td><td>624</td><td>580</td><td>579</td><td>619</td><td>467</td></tr><tr><td>2σ<sup>b</sup></td><td>4 (n=1)</td><td>---</td><td>4 (n=1)</td><td>5 (n=1)</td><td>5 (n=1)</td><td>9 (n=1)</td><td>45 (n=2)</td></tr></tbody></table>

Major elements presented as oxide weight percent; analyses conducted using the Cameca SX-50 in the Materials Characterization Laboratory, Penn State. Unless otherwise indicated, trace element, isotope ratio, and volatile analyses conducted using the Cameca IMS 1280 at NENIMF (WHOI). nd = not determined for this sample.

(a) Boron and lithium concentrations by Cameca IMS 3f at NENIMF (WHOI).

(b) When n=1, reported uncertainties are internal precision. When n>1, reported uncertainties are the greater of the external reproducibility and internal precision (in most cases this will be the external reproducibility). All uncertainties reported as 2 standard deviations of the mean (95% confidence).
<table><caption>**Table 2.** Mixing model endmembers and reference values</caption><thead><tr><th>Endmember</th><th>B (ppm)</th><th>d<sup>11</sup>B (‰)</th><th>Reference</th></tr></thead><tbody><tr><td colspan="4"><em>Mixing model endmembers</em></td></tr><tr><td>Low-B Lakagígar melt inclusion (12pl1a)</td><td>0.59</td><td>-7.8</td><td>This study</td></tr><tr><td>Altered Reykjanes basalt, surface flow</td><td>3.83</td><td>-18.3</td><td>Raffone et al., 2010</td></tr><tr><td>Altered Reykjanes basalt, 3000m depth (geothermal well)</td><td>9.27</td><td>-14.4</td><td>Raffone et al., 2010</td></tr><tr><td colspan="4"><em>Model adsorbed boron on altered basalt<sup>a</sup></em></td></tr><tr><td>Water/rock = 4</td><td>3.8</td><td>-12.7</td><td>This study</td></tr><tr><td>Water/rock = 10</td><td>5.6</td><td>-22.3</td><td>This study</td></tr><tr><td colspan="4"><em>Reference values</em></td></tr><tr><td>N-MORB</td><td>0.5</td><td>-4</td><td>Chaussidon and Marty, 1995</td></tr><tr><td>Seawater altered MORB</td><td>5.2</td><td>+3.4</td><td>Smith et al., 1995</td></tr><tr><td>Glacial meltwater (meteoric)<sup>b</sup></td><td>0.05</td><td>-1.8</td><td>Aggarwal et al., 2000</td></tr><tr><td>Icelandic geothermal water<sup>c</sup></td><td>2.8</td><td>-3.7</td><td>Aggarwal et al., 2000</td></tr><tr><td>Seawater</td><td>4.6</td><td>+39.5</td><td>Spivack and Edmond, 1987</td></tr></tbody></table>

<footnote><sup>a</sup> Model for boron absorbed onto clay-like alteration products using an adsorption coefficient of 2.6 and an isotopic fractionation factor (a) of 0.975 (Palmer et al., 1987). Fluid composition is that of “Icelandic geothermal water” given above. See text for model details.</footnote>

<footnote><sup>b</sup> from Myvatn</footnote>

<footnote><sup>c</sup> Average of several high-temperature geothermal systems sampled over multiple years</footnote>