The evolution and storage of primitive melts in the Eastern Volcanic Zone of Iceland: the 10 ka Grímsvötn tephra series (i.e. the Saksunarvatn ash)

  • David A. Neave
  • John Maclennan
  • Thorvaldur Thordarson
  • Margaret E. Hartley
Open Access
Original Paper


Major, trace and volatile elements were measured in a suite of primitive macrocrysts and melt inclusions from the thickest layer of the 10 ka Grímsvötn tephra series (i.e. Saksunarvatn ash) at Lake Hvítárvatn in central Iceland. In the absence of primitive tholeiitic eruptions (MgO > 7 wt%) within the Eastern Volcanic Zone (EVZ) of Iceland, these crystal and inclusion compositions provide an important insight into magmatic processes in this volcanically productive region. Matrix glass compositions show strong similarities with glass compositions from the AD 1783–1784 Laki eruption, confirming the affinity of the tephra series with the Grímsvötn volcanic system. Macrocrysts can be divided into a primitive assemblage of zoned macrocryst cores (An78–An92, Mg#cpx = 82–87, Fo79.5–Fo87) and an evolved assemblage consisting of unzoned macrocrysts and the rims of zoned macrocrysts (An60–An68, Mg#cpx = 71–78, Fo70–Fo76). Although the evolved assemblage is close to being in equilibrium with the matrix glass, trace element disequilibrium between primitive and evolved assemblages indicates that they were derived from different distributions of mantle melt compositions. Juxtaposition of disequilibrium assemblages probably occurred during disaggregation of incompatible trace element-depleted mushes (mean La/Ybmelt = 2.1) into aphyric and incompatible trace element-enriched liquids (La/Ybmelt = 3.6) shortly before the growth of the evolved macrocryst assemblage. Post-entrapment modification of plagioclase-hosted melt inclusions has been minimal and high-Mg# inclusions record differentiation and mixing of compositionally variable mantle melts that are amongst the most primitive liquids known from the EVZ. Coupled high-field strength element (HFSE) depletion and incompatible trace element enrichment in a subset of primitive plagioclase-hosted melt inclusions can be accounted for by inclusion formation following plagioclase dissolution driven by interaction with plagioclase-undersaturated melts. Thermobarometric calculations indicate that final crystal–melt equilibration within the evolved assemblage occurred at ~1140 °C and 0.0–1.5 kbar. Considering the large volume of the erupted tephra and textural evidence for rapid crystallisation of the evolved assemblage, 0.0–1.5 kbar is considered unlikely to represent a pressure of long-term magma accumulation and storage. Multiple thermometers indicate that the primitive assemblage crystallised at high temperatures of 1240–1300 °C. Different barometers, however, return markedly different crystallisation depth estimates. Raw clinopyroxene–melt pressures of 5.5–7.5 kbar conflict with apparent melt inclusion entrapment pressures of 1.4 kbar. After applying a correction derived from published experimental data, clinopyroxene–melt equilibria return mid-crustal pressures of 4 ± 1.5 kbar, which are consistent with pressures estimated from the major element content of primitive melt inclusions. Long-term storage of primitive magmas in the mid-crust implies that low CO2 concentrations measured in primitive plagioclase-hosted inclusions (262–800 ppm) result from post-entrapment CO2 loss during transport through the shallow crust. In order to reconstruct basaltic plumbing system geometries from petrological data with greater confidence, mineral–melt equilibrium models require refinement at pressures of magma storage in Iceland. Further basalt phase equilibria experiments are thus needed within the crucial 1–7 kbar range.


Basalt Saksunarvatn Iceland Plagioclase-hosted melt inclusions Thermobarometry 


Primitive melts in the Eastern Volcanic Zone of Iceland

Unmodified mantle melts rarely erupt at the Earth’s surface. Numerous processes including crystallisation, mixing, assimilation and volatile exsolution affect primary melts as they ascend through the crust, overprinting geochemical records of mantle melting and source composition (Dungan and Rhodes 1978; Sinton and Detrick 1992; Langmuir et al. 1992; Rubin et al. 2009). Given that they host numerous primitive basalts and picrites, it is unsurprising that the Northern Volcanic Zone (NVZ) and Reykjanes Peninsula of Iceland have been the foci of intense geochemical study for some decades (Jakobsson et al. 1978; Elliott et al. 1991; Gurenko and Chaussidon 1995; Slater et al. 2001; Stracke et al. 2003; Thirlwall et al. 2004). Despite being the most volcanically active region of the country, no eruptions of primitive tholeiite (MgO > 7 wt%) have been reported from within the Eastern Volcanic Zone (EVZ) of Iceland (Jakobsson 1979; Thordarson and Höskuldsson 2008). The evolution and storage of primitive melts in the EVZ must therefore be investigated using indirect methods such as studying primitive crystals and the melt inclusions they contain.

Studies of Icelandic melt inclusions have contributed to the understanding of numerous geological processes including mantle melting, magma mixing and volatile exsolution (e.g. Gurenko and Chaussidon 1995; Thordarson et al. 1996; Gurenko and Sobolev 2006; Maclennan 2008a, 2008b; Hartley et al. 2014). Most of these studies used olivine-hosted inclusions, the major element contents of which are susceptible to post-entrapment modification that obscures primitive melt compositions (Gaetani and Watson 2000; Danyushevsky et al. 2002). Here, we present data from a suite of predominantly plagioclase-hosted melt inclusions in which post-entrapment modification of major and trace elements has been minimal and primitive melt evolution processes can be discerned. Significant pre-eruptive modification of melt inclusion volatile contents has nevertheless taken place.

Estimating magma storage depths

In order to correlate geophysical records of volcanic unrest with petrological records of pre-eruptive processes (e.g. Saunders et al. 2012; Kilgour et al. 2014), it is important that magma storage depths can be estimated robustly from the erupted products of volcanoes. From amongst other methods, the depth of magmatic processes can be obtained from melt inclusion entrapment pressures (e.g. Dixon et al. 1995; Moore 2008), or from barometric calculations that exploit the sensitivity of mineral compositions to pressure (e.g. Nimis 1995; Putirka et al. 1996). However, petrologically determined magma storage depths are subject to various uncertainties and pitfalls that can result in different methods returning incoherent depth estimates (Métrich and Wallace 2008; Putirka 2008). The presence of clinopyroxene macrocrysts in equilibrium with glassy melt inclusions in the 10 ka Grímsvötn tephra series offers an excellent opportunity for comparing different barometric techniques within the products of a single eruption. We find that discrepancies between different barometric techniques are greater than quoted model uncertainties as a result of incomplete calibration of clinopyroxene–melt equilibrium at the conditions relevant to magma storage in Iceland and pre-eruptive CO2 leakage of from melt inclusions.

The 10 ka Grímsvötn tephra series (i.e. the Saksunarvatn ash)

The tephra horizon commonly referred to as the Saksunarvatn ash occurs in a broad swathe across the North Atlantic from the Greenland ice cap to northern Germany (Mangerud et al. 1986; Grönvold et al. 1995; Merkt et al. 1993; Haflidason et al. 2000; Bramham-Law et al. 2013; Fig. 1). As such, it forms an important regional tephrochronological marker in Holocene sediment and ice formations (e.g. Rasmussen et al. 2006; Thornalley et al. 2011). Although the type locality, Lake Saksunarvatn, is in the Faroe Islands (Waagstein and Jóhansen 1968; Jóhansen 1975, 1985), the composition of the tephra shows that it originated in Iceland, more precisely from the Grímsvötn volcanic system in south central Iceland (e.g. Grönvold et al. 1995; Jennings et al. 2014). Specifically, the Saksunarvatn ash bears a striking compositional resemblance to the ~15.1 km3 AD 1783–1784 Laki eruption products, which are slightly less evolved than typical magmas erupted at the Grímsvötn central volcano (Thordarson 2014).
Fig. 1

a Map showing the extent (red dashed line) of the 10 ka Grímsvötn tephra series (i.e. Saksunarvatn ash) according to Bramham-Law et al. (2013). Terrestrial deposits are shown as green circles, and marine deposits are shown as blue diamonds. Samples were collected from Lake Hvítárvatn in central Iceland, which is marked by a black star. The Grímsvötn central volcano is marked by a red triangle. b Map of Iceland showing the location of various localities discussed in the text including the Eastern Volcanic Zone (EVZ), Northern Volcanic Zone (NVZ) and Reykjanes Peninsula (REP). Volcanic systems are shaded in red

The Saksunarvatn ash tephra layer has a primary thickness of 5–10 mm in the Faroe Islands (Mangerud et al. 1986), which implies that the tephra layer is at least volumetrically equivalent to the largest explosive basalt eruptions in Iceland during historical time, namely the ~3.3 km3 AD 870 Vatnaöldur and ~6.7 km3 AD 1477 Veiðivötn basaltic phreatoplinian eruptions (Larsen, 1984, 2005). Furthermore, recent studies have shown that the Saksunarvatn ash horizon is composed of tephra fall from at least six eruptions that took place in the Grímsvötn volcanic system over the period 10.5 ka to 9.9 ka, each ranging in volume from >1–30 km3 (Jóhannsdóttir et al. 2005; Jennings et al. 2014; Thordarson 2014). Ash from five of these events was dispersed westwards from the source vent, whilst ash from at least one event (i.e. the Saksunarvatn ash proper) was dispersed towards the east. Whilst all of these tephra layers have identical major element compositions (Table 1), a more thorough petrological characterisation has yet to be undertaken. Hereafter, these tephra layers are referred to as the 10 ka Grímsvötn tephra series, in order to distinguish the series fro m the Saksunarvatn ash proper that is observed in the Faroe Islands (Jennings et al. 2014; Thordarson 2014).
Table 1

Major element compositions of the 10 ka Grímsvötn tephra series at Lake Hvítárvatn (Thordarson, unpublished data, 2013)

Sample numbera


Number of analyses (n)













12.2–29.7 cm

Saks-eq I HVT (10.4 ka)

55, mean =












1σ =












GLAD-HVT03 1A-7H-1

83.0 cm–GLAD-HVT03 1A-7H-2 6.5 cm

Saks-eq II HVT (10.35 ka)b

153, mean =












1σ =












GLAD-HVT03 1A-7H-1

67.0–79.5 cm

Saks-eq III HVT (10.3 ka)

78, mean =












1σ =












GLAD-HVT03 1A-7H-1

42.9–43.1 cm

Saks-eq IV HVT (10.1 ka)

10, mean =












1σ =












GLAD-HVT03 1A-7H-1

32.5–33.1 cm

Saks-eq V HVT (10 ka)

10, mean =












1σ =












GLAD-HVT03 1A-6H-2

127.8–128.4 cm

Grímsvötn HVT (8.85 ka)

11, mean =












1σ =












aRefers to section of sediment core retrieved from Lake Hvítárvatn and the depth in core section

bSaks-eq II HVT is the focus of this study, and is referred to as the Hvítárvatn tephra layer

Sample collection and preparation

Glass and crystal fragments 150–800 µm in size were selected from the thickest layer of the 10 ka Grímsvötn tephra series (Saks-eq II HVT) observed in the GLAD-HVT03-1A-7H core collected from Lake Hvítárvatn in 2003 (Fig. 1). Saks-eq II HVT is dated to approximately 10.35 ka (Jóhannsdóttir 2007; Rasmussen et al., 2006; Lohne et al. 2014) and is henceforth referred to as the Hvítárvatn tephra layer. Tephra consists mainly of highly vesicular, holohyaline tephra clasts, and contains conspicuous colourless plagioclase fragments. Fragments of dark green clinopyroxene and pale green olivine are also present. In order to avoid using genetic terms such as phenocryst, xenocryst and antecryst, all crystals larger than 150 μm are referred to as macrocrysts hereafter (c.f. Thomson and Maclennan 2013).

After being washed in distilled water and then dried, macrocrysts were mounted in Buehler EpoThin™ resin and polished to expose melt inclusions for analysis. Not all melt inclusions were exposed after the first round of polishing. In order to maximise the number of melt inclusion analyses in different host crystals given the small volume of sample material available, crystals were re-polished after a first round of analyses in order to expose a second plane of melt inclusions deeper within the samples.

Analytical methods

Backscattered electron (BSE) images were produced using a JEOL JSM-820 scanning electron microscope (SEM) in the Department of Earth Sciences at the University of Cambridge, UK. Trace elements in melt inclusions, matrix glasses and clinopyroxene macrocrysts were measured by secondary ion mass spectrometry (SIMS). CO2 and H2O were also measured in glasses by SIMS. Major elements in all samples were measured by electron probe microanalysis (EPMA). Inclusion-hosted bubbles were investigated by Raman spectroscopy in order to determine whether they contained CO2.

Secondary ion mass spectrometry (SIMS)

SIMS analyses were performed at the NERC Ion Microprobe Facility at the University of Edinburgh using a Cameca ims-4f instrument. Analyses of CO2 in glasses were carried out first in order to avoid introducing surface contamination from carbon coating. Analyses were carried out with the instrument configured for a high mass resolving power in order to ensure separation of 12C and 24Mg2+ peaks. Measurements were taken using a primary O ion beam with an accelerating voltage of 15 kV, a beam current of 6 nA, a secondary accelerating voltage of 4500 V minus a 50 V offset and a 25-µm image field. The ion beam was rastered over an area of approximately 25 µm2 for 120 s prior to analysis to remove surface contamination. The raster was then switched off, and a 10 µm beam centred in the middle of the rastered area was used to make measurements. The following mass positions and isotopes were measured for 16 cycles, with counting times in seconds in parentheses: 12C(5), 24Mg2+(10), 28Si2+(2) and 30Si(2). Counts from the final eight cycles were used to calculate CO2 concentrations using a calibration curve constructed from basaltic glass standards from Shishkina et al. (2010) containing 0–2000 ppm CO2. Precision was estimated as 1σ = ± 5.9 % using repeat analyses of both standards and samples.

Analyses of H2O and trace elements in glass were then undertaken using a primary O ion beam with an accelerating voltage of 15 kV, a beam current of 6 nA, a secondary accelerating voltage of 4500 V minus a 75 V offset and a 25-µm image field. Analyses were made with a spot size of ~15 µm centred in the pit formed during preceding CO2 analyses. The following isotopes were measured for eight cycles, with counting times in seconds in parentheses: atomic mass 0.7(3), 1H(5), 30Si(2), 39K(3), 47Ti(3), 88Sr(3), 89Y(3), 90Zr(3), 93Nb(3), atomic mass 130.5(3), 138Ba(3), 139La(3), 140Ce(3), 141Pr(5), 143Nd(5), 149Sm(8), 151Eu(10), 156Gd(5), 159Tb(10), 161Dy(10), 165Ho(10), 167Er(15), 169Tm(20), 174Yb(20) and 175Lu(20). Peak positions were verified before each analysis, and masses 0.7 and 130.5 were measured to determine backgrounds in each cycle, which were always sufficiently close to zero to be ignored.

H2O concentrations were calculated using a calibration curve constructed from basaltic glass standards from Shishkina et al. (2010) containing 0–2.96 wt% H2O. Precision was estimated as 1σ = ± 3.9 % using repeat analyses of both standards and samples. Absolute trace element concentrations were calculated using JCION-6 software, where corrections were made for oxide interferences of light REEs on heavy REEs and BaO on Eu. NIST-610 was used as the primary calibration standard (Jochum et al. 2011), and matrix corrections were applied using the synthetic basalt standard GSD-1G (Jochum et al. 2005). Accuracy was monitored using international standards ML3B-G, GOR128-G, BCR2-G and KL2-G (Jochum et al. 2006) and was generally better than ± 5 % for trace elements. Precision was estimated as better than 1σ = ± 5 % for trace elements of high abundance (e.g. Zr and Sr), and as better than 1σ = ± 10 % for trace elements of low abundance (e.g. Yb and Lu) using repeat analyses of samples.

Trace elements in clinopyroxenes were analysed using the same conditions as for trace element analyses of glasses. NIST-610 was used as the primary calibration standard (Jochum et al. 2011). NIST-610 has similar ion yield patterns and efficiencies to clinopyroxene, negating the requirement for matrix corrections (Winpenny and Maclennan 2011). The following isotopes were measured for eight cycles, with counting times in seconds in parentheses: 40Ca(5), 44Ca(5), 45Sc(5), 47Ti(3), 88Sr(3), 89Y(3), 90Zr(3), 93Nb(3), atomic mass 130.5(3), 138Ba(3), 139La(3), 140Ce(3), 141Pr(5), 143Nd(5), 149Sm(8), 151Eu(10), 156Gd(5), 159Tb(10), 161Dy(10), 165Ho(10), 167Er(15), 169Tm(20), 174Yb(20) and 175Lu(20). Corrections were made for 44CaH+ interferences on 45Sc. Precision was estimated as better than 1σ = ± 5–15 % depending on element abundance using repeat analyses of samples.

Electron probe microanalysis (EPMA)

EPMA was performed at the University of Cambridge using a Cameca SX100 instrument. Glass analyses were performed with an operating potential of 15 kV, a beam current of 10 nA and a spot size of 5 µm or 10 µm depending on sample size. Olivine, clinopyroxene and Mg–Al-chromite analyses were performed with an operating potential of 15 kV, a beam current of 20 nA and a focussed beam. Plagioclase analyses were performed with a beam current of 10 nA and a spot size of 5 µm in order to minimise Na mobilisation. Counting times were as follows: 20 s for major elements in glass and crystals, except for Ti which was counted for 60 s and Na which was counted for 10 s; 30 s for minor elements in crystals; and 60 s for P, Cr, Mn and Ni in glass. Calibration standards were as follows: jadeite for Na, periclase for Mg, Si glass for Si, K-feldspar for K, rutile for Ti, fayalite for Fe, corundum for Al, apatite for P, and pure metals for Cr, Mn and Ni.

Repeat analyses of standards indicate that plagioclase anorthite content (An = 100 × atomic Ca/(Ca + Na + K)) was determined with a precision of 1σ = ± 0.5 mol. % and that olivine forsterite content (Fo = 100 × atomic Mg/(Mg + Fe2+)) was determined with a precision of 1σ = ± 0.23 mol. %. Clinopyroxene magnesium number (Mg# = 100 × atomic Mg/(Mg + Fe2+)) was determined with a precision similar to that obtained for olivine forsterite content. Precision was estimated as better than 1σ = ± 5 % for major elements in glasses, but as poor as 1σ = ± 30–40 % for minor elements in glasses (MnO, K2O and P2O5) using repeat analyses of samples.

Raman spectroscopy

A subset of inclusion-hosted bubbles was analysed by Raman spectroscopy using a Horiba LabRam™ instrument at the University of Cambridge, UK. The detection limit of CO2 in inclusion-hosted bubbles is better than 0.04 g/cm3 based on analysis of CO2-bearing samples on the same instrument (Hartley et al. 2014).


Macrocryst compositions and textures

Plagioclase, clinopyroxene and olivine macrocrysts are found throughout the Hvítárvatn tephra layer. Apparent macrocryst sizes are as follows: plagioclase, 200–800 µm; clinopyroxene, 150–500 µm; and olivine, 150–400 µm. However, macrocrysts may have been significantly larger before syn-eruptive fragmentation. Occasional smaller crystals (10–150 µm) are present within the matrix glass, but were not analysed.

Plagioclase macrocrysts (n = 108) can be divided into two populations depending on whether they are compositionally zoned (Figs. 2a, 3a): zoned macrocrysts have primitive, high-anorthite cores (An83–An92) encased within evolved, low-anorthite rims (~An65), whereas unzoned macrocrysts have low-anorthite contents throughout (An60–An68). Two plagioclase macrocrysts have core compositions of An80. Trace elements also reflect core–rim zonation: anorthite content correlates negatively with TiO2 concentration (Fig. 3b). Mg–Al chromites (Cr# = 41–51, where Cr# = 100 × atomic Cr/(Cr + Al)) are found as rare inclusions within high-anorthite cores and are further characterised by low TiO2 contents (<0.65 wt%).
Fig. 2

a False-colour BSE image of a plagioclase macrocryst. Most of the crystal is composed of a high-anorthite core with low amplitude oscillatory zoning (An87–An92) that is encased within a low-anorthite (An65) rim. Note the sharp core/rim boundary and ragged skeletal nature of the rim. b BSE image showing the inclusion-rich nature of some plagioclase macrocrysts. Investigation of inclusion-rich macrocrysts by optical microscopy reveals that inclusions are discrete—they are not joined by thin melt tubes. However, they may have been joined at the time of formation. Circular features are pits formed during SIMS analyses. c False-colour BSE image of a low-Mg# clinopyroxene macrocryst highlighting sector zoning. Hourglass zoning in this clinopyroxene is primarily defined by the replacement of Al2O3 and TiO2 in the brighter, denser zones by FeO and MgO. Note the skeletal nature of the macrocryst rim. d Reflected and transmitted light images of a low-forsterite olivine macrocryst that is ramified with melt inclusions and melt channels

Fig. 3

Plots summarising the composition of macrocrysts from the Hvítárvatn tephra layer. 2σ error bars are shown in bd. a Major element composition of macrocryst cores and rims. Vertical red bars show plagioclase, olivine and clinopyroxene compositions in equilibrium with the matrix glass according to mineral–melt equilibrium models of Namur et al. (2011), Herzberg and O’Hara (2002) and Eq. 35 of Putirka (2008), respectively. b TiO2 content of plagioclase cores and rims plotted against anorthite content. c TiO2 content of clinopyroxene cores and rims plotted against clinopyroxene Mg#. d TiO2 content of clinopyroxene macrocrysts plotted against Al2O3 to enable compositionally distinct sector zones to be discriminated (Fig. 2c). High- and low-Al2O3 low-Mg# clinopyroxene zones are marked as well as high-Mg# clinopyroxenes. e La/Ybcpx in a subset of clinopyroxene cores analysed by SIMS plotted against clinopyroxene Mg#

BSE images reveal that anorthite content changes abruptly at the boundary between cores and rims and that low amplitude oscillatory zoning (An84–An92) occurs within some high-anorthite cores (Fig. 2a). Whilst some cores contain no melt inclusions (Fig. 2a), others contain abundant small (5–40 μm in length) melt inclusions (Fig. 2b) that are often arranged into crystallographically controlled chains and concentric bands. Rare large inclusions up to ~210 μm in length are also present. Zoned plagioclase macrocryst rims are thin (~10 μm) and often skeletal in nature (Fig. 2a). Elongate, unzoned plagioclase macrocrysts are often texturally associated with low-Mg# clinopyroxene macrocrysts, indicating that they represent the fragmented remains of ophitic glomerocrysts (c.f. AD 1783–84 Laki eruption; Guilbaud et al. 2007; Neave et al. 2013).

Clinopyroxene macrocrysts (n = 48) can also be divided into two populations based on their core compositions (Fig. 3a): a primitive population containing high-Mg# cores (Mg# = 82–87) and an evolved population containing low-Mg# cores (Mg# = 71–78). Rims of high-Mg# clinopyroxene macrocrysts are thin (~10 μm) and evolved (Mg# ~ 73). Compositional bimodality is also reflected in clinopyroxene minor and trace element contents (Fig. 3c–e): high-Mg# cores are characterised by low TiO2 (~0.45 wt%) and La/Ybcpx (0.25 ± 0.06(1σ)), whereas low-Mg# cores have higher TiO2 (>0.75 wt%) and La/Ybcpx (0.40 ± 0.04(1σ)).

Although high-Mg# clinopyroxenes show no evidence of sector zoning—they form a single population on a plot of Al2O3 versus TiO2 (Fig. 3d)—low-Mg# clinopyroxenes exhibit strong sector zoning (Fig. 2c). Low-Mg# clinopyroxene analyses can be divided into a high-Al2O3 group (>3 wt%) that maintained crystal–melt equilibrium during growth, and a disequilibrium low-Al2O3 group. Sector zoning of this type results from the greater uptake of Mg and Fe2+ and concurrent dilution of Ca, Al and Ti on {100} faces with respect to on {010} and {110} faces during rapid crystal growth (Nakamura 1973). Therefore, only analyses from equilibrium low-Mg# crystal zones (i.e. those with Al2O3 > 3 wt%) were used to evaluate clinopyroxene–melt trace element equilibrium and as the input for thermobarometric calculations. Skeletal textures around the rims of some clinopyroxenes provide further evidence of rapid growth.

Whilst most olivine macrocrysts (n = 32) have evolved core compositions (Fo70–Fo76), a small number (n = 7) contain primitive cores (Fo84–Fo87) and one has a core composition of Fo79.5 (Fig. 3a). Rims of primitive olivine macrocrysts are 10–20 μm thick and have low-forsterite contents (Fo72.5) adjacent to the matrix glass. Whilst high-forsterite olivines occur as rounded crystal fragments, low-forsterite olivines are often ramified with melt inclusions and exhibit skeletal textures (Fig. 2d).

Matrix glass and melt inclusion compositions

Major elements

Melt inclusion and matrix glass major element compositions are summarised in Fig. 4. Matrix glass compositions are homogeneous and similar to previously published values for the Saksunarvatn ash (Mangerud et al. 1986; Grönvold et al. 1995; Jóhannsdóttir et al. 2005; Jóhannsdóttir 2007; Bramham-Law et al. 2013).
Fig. 4

Plots summarising melt inclusion and matrix glass major element compositions from the Saksunarvatn tephra. 2σ error bars are shown throughout. Melt inclusion compositions have not been corrected for post-entrapment crystallisation (PEC). The first principal component (PC1) of primitive melt inclusion major element data (9.5 wt% < MgO < 12 wt%) is shown by the solid black line. The composition of a hypothetical near-primary melt (Mg# = 70) is shown by the large open diamond. Variation along PC1 can be accounted for by removing a troctolitic assemblage consisting of plg:ol:cpx in a 13:7:0 ratio from the hypothetical near-primary melt. PEC (black crosses mark 5 % PEC intervals) moves plagioclase-hosted melt inclusion compositions along the solid blue line. Al2O3 (c) and FeOt (d), which are particularly sensitive to PEC, show little evidence of post-entrapment modification. Natural Icelandic compositions that most closely match the hypothetical near-primary melt are shown as inverted black triangles (Slater et al. 2001; Maclennan et al. 2003)

Melt inclusions can be divided into two main populations based on the composition of their host crystals: primitive melt inclusions that are hosted within high-anorthite plagioclase and high-forsterite olivine and evolved melt inclusions that are hosted within low-anorthite plagioclase and low-forsterite olivine (Fig. 4). Primitive plagioclase-hosted melt inclusions (n = 113) are MgO-rich (MgO = 9.3–12.3 wt%; Mg# = 62–70), whereas primitive olivine-hosted melt inclusions have lower MgO contents (5.8–6.1 wt%), consistent with post-entrapment crystallisation and re-equilibration (Danyushevsky et al. 2002). Evolved plagioclase-hosted melt inclusions (n = 4) have low MgO contents (MgO ~ 5.4 wt%; Mg# ~ 43) and are compositionally indistinguishable from the matrix glass. Additionally, five melt inclusions with MgO = 8–8.5 wt% are hosted within ~An80 plagioclase.

Trace elements

Matrix glass and melt inclusion trace element compositions are summarised on plots of selected element pairs (Fig. 5) and on a multi-element diagram normalised to depleted MORB mantle (DMM; Workman and Hart 2005) (Fig. 6). A cross-correlation matrix of elements in primitive plagioclase-hosted inclusions is provided in the Supplementary Material. Variability in trace element abundance between matrix glass samples is comparable with estimates of analytical precision (1σ = ± 5–10 %).
Fig. 5

Plots summarising melt inclusion and matrix glass trace element compositions from the Saksunarvatn tephra. 2σ error bars are shown. a Plot of highly incompatible La versus moderately incompatible Yb. The mean La/Yb of matrix glass compositions is plotted as a solid red line. The mean matrix glass composition from the Laki eruption (Hartley et al. 2014) is shown as a dashed green line. Lines of constant ratio are shown as solid pale grey lines. The black box shows the region plotted in (b). b La plotted against Yb at an expanded scale to show primitive melt inclusion compositions with greater clarity. Low-HFSE melt inclusions are relatively incompatible trace element-enriched (La/Yb > 2). c Plot of more incompatible Zr versus less incompatible Y. Low-HFSE melt inclusions have distinctively low Zr/Y values in comparison with the rest of the primitive melt inclusion population. Evolved melt inclusions and the matrix glass have higher Zr and Y contents than shown on this figure. d Plot of Sr/Y versus La/Y. Primitive melt inclusions define a positively correlated array. Instantaneous fractional melt compositions calculated using pMELTS (Ghiorso et al. 2002) shown as a dashed black line. Melting calculations were performed using a depleted MORB mantle source (DMM; Workman and Hart 2005). Plagioclase PEC, shown as a solid pale blue line, affects Sr/Y, but not La/Y. Evolved melt inclusions and the matrix glass have low Sr/Y, consistent with substantial crystallisation of plagioclase during melt evolution

Fig. 6

Depleted MORB mantle (DMM; Workman and Hart 2005) normalised multi-element diagram for melt inclusions and the matrix glass from the Saksunarvatn tephra. Melt inclusion compositions fall into two populations: an evolved population that contains the matrix glass and a few melt inclusions hosted in plagioclase and olivine, and a primitive population that contains the majority of plagioclase-hosted melt inclusions and two olivine-hosted melt inclusions. 2σ errors are shown along the top of the plot and DMM normalisation values are shown along the bottom. Variability in primitive plagioclase-hosted melt inclusions is reported as signal-to-noise ratios (σ t/σ r) for each element measured along the top of the plot (e.g. Maclennan et al. 2003). All trace elements apart from Yb have σ t/σ r values that are significant at the 99 % confidence level (i.e. σ t/σ r > 1.71). A sub-population of plagioclase-hosted melt inclusions, plotted as dashed red lines, have pronounced high-field strength element (HFSE) depletions that are particularly visible as negative Nb and Zr anomalies; low-HFSE melt inclusions have Zr/Y < 2 (Fig. 5c)

Primitive melt inclusions have lower concentrations of both highly and moderately incompatible trace elements than evolved melt inclusions and the matrix glass (Figs. 5, 6). Evolved melt inclusions are compositionally similar to the matrix glass in terms of both trace element abundances and trace element ratios (Fig. 5a): the mean matrix glass La/Yb of 3.55 ± 0.10(1σ) is within analytical error of the mean evolved melt inclusion La/Yb of 3.60 ± 0.18(1σ). In contrast, primitive melt inclusions are more diverse: primitive melt inclusion La/Yb values span the range 1.08–3.41 (Fig. 5b). The matrix glass and evolved melt inclusions also have pronounced negative Sr anomalies (Sr/Sr* = 0.49 ± 0.02(1σ), where Sr/Sr* = Srn/(Cen × Ndn)0.5 and n = chondrite values from Arevalo and McDonough (2010)), whereas primitive melt inclusions have, on average, slightly positive anomalies (Sr/Sr* = 1.13 ± 0.15(1σ)). There is no relationship between La/Yb and Sr/Sr* in primitive melt inclusions (r = 0.06).

All trace elements apart from Yb are significantly variable at the 99 % confidence level within the primitive melt inclusion population, i.e. have signal-to-noise ratios (σ t/σ r, where σ t is the true variability within a sample set and σ r is an estimate of analytical error (Maclennan et al. 2003)) in excess of 1.71 (Fig. 6). La/Yb, La/Sm and Sr/Sr* are also significantly variable at the 99 % confidence level, and Sm/Yb values are significantly variable at the 95 % confidence level.

Some primitive melt inclusions have striking high-field strength element (HFSE) (Fig. 6). In contrast to the full primitive melt inclusion population, HFSE-depleted melt inclusions have lower mean concentrations of moderately incompatible trace elements (Fig. 5c) and are also slightly more incompatible trace element-enriched (HFSE-depleted mean La/Yb = 2.51 ± 0.58(1σ) versus La/Yb = 1.95 ± 0.61(1σ)).

H2O and CO2

Matrix glass and melt inclusion H2O–CO2 systematics are summarised in Fig. 7a. The CO2 content of primitive melt inclusions ranges from 262 to 2070 ppm. Evolved melt inclusions contain 78–405 ppm CO2. All values <200 ppm are found in olivine-hosted melt inclusions. The CO2 content of the matrix glass is below the detection limit of CO2 by SIMS (~25 ppm). The H2O content of primitive melt inclusions ranges from 0.19 to 0.64 wt%, though most inclusions contain >0.4 wt% H2O. Evolved melt inclusions are H2O-rich, containing 0.56–0.68 wt% H2O. Matrix glass H2O contents span the range 0.11–0.46 wt% H2O.
Fig. 7

a Plot summarising melt inclusion and matrix glass H2O and CO2 contents. 2σ error bars are shown. b Cumulative probability distributions of CO2 in bubble-bearing and bubble-free primitive plagioclase-hosted melt inclusion populations. Bubble-bearing and bubble-free inclusions have similar CO2 distributions, indicating that bubbles contain negligible CO2. Note that bubble-free inclusions are systematically offset to slightly higher values by a small population of low-CO2 (<300 ppm) bubble-bearing inclusions

Inclusion-hosted bubbles

If inclusion-hosted bubbles remain in equilibrium with coexisting trapped melts during cooling; then, CO2 may be transferred into inclusion-hosted bubbles (Roedder 1979; Anderson and Brown 1993; Steele-Macinnis et al. 2011; Bucholz et al. 2013; Hartley et al. 2014; Moore et al. 2015). However, assuming that bubble–melt equilibrium is maintained within cooling, inclusions may not always be valid: rapid quenching kinetically inhibits CO2 diffusion, resulting in reduced sequestration of CO2 into bubbles (Neave et al. 2014a; Wallace et al. 2015).

Inclusion-hosted bubbles are present in 49 of the 133 melt inclusions analysed from the Hvítárvatn tephra layer. The majority of inclusion-hosted bubbles make up 3.0 ± 1.5(1σ) vol. % of total melt inclusion volumes, consistent with formation by thermal contraction (Lowenstern 1995; Moore et al. 2015). In three inclusions, bubbles make up larger proportions (8–12 vol. %), consistent with simultaneous entrapment alongside the silicate melt. No Fermi doublets were detected in the subset of inclusion-hosted bubbles investigated by Raman spectroscopy, indicating that that minimal CO2 currently resides within inclusion-hosted bubbles (<0.04 g/cm3). That glasses from bubble-bearing and bubble-free melt inclusion populations have similar CO2 composition distributions further suggests that sequestration of CO2 into inclusion-hosted bubbles has not occurred and that these near-vacuum bubbles contain, at most, noble gases and trace amounts of CO2 (Fig. 7b; Neave et al. 2014a) Whilst we conclude that inclusion glasses record the total CO2 content of melt inclusions at the time of quenching, our findings are unaffected if bubble-bearing inclusions are removed from the data set.


Provenance of the tephra

Hvítárvatn tephra glass compositions overlap with those from the Grímsvötn volcanic system (Fig. 8; Sigmarsson et al. 2000, 2013; Jude-Eton et al. 2012; Streeter and Dugmore 2014), with strombolian tephra glasses from the AD 1783–1784 Laki eruption providing the closest match (Thordarson et al. 1996; Hartley et al. 2014). In contrast, glasses from the Bárðarbunga-Veiðivötn system are more primitive and incompatible trace element-depleted than the Hvítárvatn tephra layer (Hansen and Grönvold 2000; Lawson et al. 2007; Neave et al. 2014a; Streeter and Dugmore 2014), whereas those from the Katla system are more evolved and incompatible trace element-enriched (Óladóttir et al. 2007).
Fig. 8

Plot comparing the major element compositions of Hvítárvatn tephra layer glasses and melt inclusions with literature values for the Saksunarvatn ash and other EVZ eruptions according to their Mg# and K2O/TiO2 contents. Literature values for the Saksunarvatn ash were taken from the following sources representing the full geographic coverage: Mangerud et al. (1986); Sigurgeirsson and Leósson (1993); Grönvold et al. (1995); Björck et al. (1992); and Bramham-Law et al. (2013). Glass compositions from the Grímsvötn volcanic system were taken from Thordarson et al. (1996), Sigmarsson et al. (2000), Jude-Eton et al. (2012), Hartley et al. (2014) and Streeter and Dugmore (2014). Glass compositions from the Bárðarbunga-Veiðivötn system were taken from Hansen and Grönvold (2000), Lawson et al. (2007), Neave et al. (2014a) and Streeter and Dugmore (2014). Glass compositions from the Katla system were taken from Óladóttir et al. (2007)

The presence of blocky and modestly vesiculated sideromelane glass within the Hvítárvatn samples suggests that magma–ice interactions occurred during eruption and may have contributed to an explosive eruption style (Jóhannsdóttir et al. 2005). During subglacial eruptions, rapid quenching restricts the exsolution of moderately soluble volatile species such as H2O and S, resulting in subglacially erupted tephras retaining higher volatile contents than their subaerial counterparts (Moore and Calk 1991; Óladottir et al. 2007; Tuffen et al. 2010). The most evolved melt inclusions from the Hvítárvatn tephra layer contain 0.60 ± 0.04(1σ) wt% H2O and provide the best estimate of immediately pre-eruptive H2O content. Tephra glass H2O contents of 0.11–0.46 wt% (0.26–3.21 MPa using the model of Shishkina et al. 2014) are significantly higher than strombolian tephra glass H2O contents from the AD 1783–1784 Laki eruption (0.08 ± 0.01 wt%; Hartley et al. 2014) and indicate that H2O exsolution prior to solidification was incomplete, consistent with the subglacial, and possibly phreatoplinian, eruption style suggested by Jóhannsdóttir et al. (2005).

Crystal–melt relationships

Macrocryst rims (~An65, Mg#cpx ~ 73, ~Fo72.5) are in equilibrium with the matrix glass according to the following mineral–melt equilibrium models: Namur et al. (2011) for plagioclase, Herzberg and O’Hara (2002) for olivine and Eq. 35 of Putirka (2008) for clinopyroxene (Fig. 3). A ferric to total iron ratio (Fe3+/ΣFe) of 0.1 (i.e. fO2 ~ QFM − 1; Maclennan 2008a, b) was assumed on the basis that it lies within the range of measured Icelandic basalts (0.08–0132; Breddam 2002; Óskarsson et al. 1994). It is thus likely that evolved macrocrysts (An60–An68, Mg#cpx = 71–78, Fo70–Fo76) crystallised from melts directly parental to the matrix glass. Melt inclusions hosted within evolved plagioclase and olivine macrocrysts have very similar trace element ratios to the matrix glass, indicating that they are related by evolution along a single liquid line of descent (Fig. 5a).

Compositional relationships between clinopyroxene macrocrysts and the matrix glass were investigated using analyses of trace elements within clinopyroxene macrocrysts. Following the methods of Winpenny and Maclennan (2011) and Neave et al. (2014b), La/Ybcpx values were converted into La/Ybmelt values using appropriate pressure–temperature–composition (P–T–X)-dependent partition coefficients (Wood and Blundy 1997). Appropriate pressures and temperatures were selected using the results of thermobarometric calculations discussed in later sections. La/Ybeqm values were only calculated for clinopyroxene zones that maintained crystal–melt equilibrium during growth. By calculating equilibrium ratios instead of equilibrium abundances, most errors associated with calculating partition coefficients cancel out because of the form of the Brice equation for lattice strain (Brice 1975). Uncertainties are discussed further by Winpenny and Maclennan (2011) and Neave et al. (2014b).

La/Ybmelt values calculated from clinopyroxene macrocrysts are shown alongside melt inclusion and matrix glass compositions in Fig. 9. The trace element composition of low-Mg# clinopyroxenes is consistent with crystallisation alongside low-anorthite plagioclase and low-forsterite olivine from melts closely related to the matrix glass (La/Ybmelt ~ 3.6). Furthermore, all evolved macrocryst phases show either skeletal textures or sector zoning that are consistent with rapid cooling rates (Nakamura 1973; Lofgren 1974; Welsch et al. 2013). In contrast, La/Ybmelt values calculated from high-Mg# clinopyroxenes are significantly more depleted in incompatible trace elements (La/Ybmelt = 1.75–3.34) and lie within the La/Yb range of the primitive melt inclusions.
Fig. 9

La/Yb versus Mg#melt for melts calculated to be in equilibrium with clinopyroxene macrocryst cores, melt inclusions and the matrix glass. REE partition coefficients were calculated at appropriate P–T–X conditions using Wood and Blundy (1997). The Mg# of melts in equilibrium with clinopyroxene macrocrysts was determined using \(K_{{d\;{\text{Fe}}\text{-}{\text{Mg}}}}^{{{\text{cpx}}\text{-}{\text{liq}}}}\) values from Eq. 35 of Putirka (2008). The 2σ error of La/Ybmelt in equilibrium with clinopyroxene macrocrysts was estimated by calculating equilibrium melts for multiple analyses within an individual macrocryst zone. Dashed black lines show fractional crystallisation paths from typical primitive incompatible trace element-depleted and trace element-enriched melt inclusion compositions calculated using Petrolog (Danyushevsky and Plechov 2011) and appropriate partition coefficients (Hart and Dunn 1993; Beattie 1994; Aigner-Torres et al. 2007). Calculations were performed at a pressure of 3 kbar with fO2 buffered at QFM-1

Trace element disequilibrium between primitive melt inclusions (mean La/Yb = 2.07) and high-Mg# clinopyroxenes (mean La/Ybmelt = 2.22) on the one hand, and evolved melt inclusions, low-Mg# clinopyroxenes and the matrix glass on the other (mean La/Ybmelt = 3.55) indicates that most primitive macrocrysts cannot be related to the matrix glass by a single liquid line of descent. Fractional crystallisation paths calculated for representative incompatible trace element-depleted and enriched primitive melt inclusions using Petrolog (Danyushevsky and Plechov 2011) with appropriate partition coefficients (Hart and Dunn 1993; Beattie 1994; Aigner-Torres et al. 2007) demonstrate that, whilst the most enriched primitive melt inclusions could evolve to be in trace element equilibrium with the matrix glass, most inclusions would remain too depleted (Fig. 8). Trace element variability between primitive and evolved components of the Hvítárvatn tephra layer thus reflects the growth of these two assemblages from different distributions of compositionally diverse mantle melts (Gurenko and Chaussidon 1995; Slater et al. 2001). Similar relationships between incompatible trace element enrichment and degree of evolution have been reported in samples across the EVZ, where they have been interpreted as a signal of mixing of primitive crystal mushes into evolved melts shortly before eruption (Hansen and Grönvold 2000; Halldórsson et al. 2008; Passmore et al. 2012; Neave et al. 2014b).

Conversion of Mg#cpx to Mg#melt using \(K_{{d\;{\text{Fe}}\text{-}{\text{Mg}}}}^{{{\text{cpx}}\text{-}{\text{liq}}}}\) from Eq. 35 of Putirka (2008) shows that only some primitive melt inclusions were three-phase saturated at the time of entrapment: high-Mg# clinopyroxene macrocrysts (Mg#cpx = 82–87) are in equilibrium with Mg#melt = 57–65, implying that inclusions with Mg#melt > 65 were not in saturated in clinopyroxene at the time of their entrapment. The appearance of clinopyroxene on the liquidus after olivine and plagioclase is nevertheless consistent with observations from petrologically similar systems from the EVZ (Hansen and Grönvold 2000; Neave et al. 2013, 2014b).

Geochemical variability in primitive plagioclase-hosted melt inclusions

Post-entrapment crystallisation

Melt inclusions hosted in high-anorthite plagioclase crystals can preserve comparable records of geochemical variability to those hosted in high-forsterite olivine (Sours-Page et al. 2002). However, before interpreting melt inclusion compositions in terms of primary magmatic processes, it is important to determine, and if necessary correct for, the effects of post-entrapment processes on inclusion compositions (e.g. Danyushevsky et al. 2000; Cottrell et al. 2002).

Post-entrapment crystallisation (PEC) paths were calculated using a forward model by incrementally crystallising plagioclase from representative inclusion compositions (blue lines on Figs. 4, 5d). Plagioclase major and trace element contents were calculated using Eq. 33 of Namur et al. (2011) with T–X dependent partition coefficients from Blundy and Wood (1991, 1994) and Bédard (2006). Calculated PEC paths reproduce the coupled Al2O3–Sr depletion and TiO2–FeOt–REE enrichment observed in natural samples that have experienced PEC (Nielsen 1995; Nielsen 2011), but lie at high angles to melt inclusion trends from the Hvítárvatn tephra layer (Figs. 4, 5d). The effect of PEC on this inclusion suite has thus been minimal, and no PEC corrections were applied.

Trace element variability and high-field strength element depletion

It is well established that primitive melt inclusions record the supply of diverse primary melts to magma reservoirs (Sobolev and Shimizu 1993; Nielsen 1995; Maclennan 2008a). It follows that most variability in melt inclusion and clinopyroxene trace element ratios from the Hvítárvatn tephra layer reflects the supply of compositionally heterogeneous primitive melts to the base of the Grímsvötn volcanic system (Gurenko and Chaussidon 1995; Maclennan 2008b). For example, strong positive correlations between Sr and highly incompatible trace elements such as La and Ce (r ≥ 0.80) can be accounted for by melting of aluminous lherzolite within the garnet stability field (e.g. McKenzie and O’Nions 1991), as illustrated by the range of pMELTS-derived incremental fractional melt compositions shown in Fig. 5d (Ghiorso et al. 2002). However, HFSE depletions present in a subset of plagioclase-hosted melt inclusions cannot be explained by simple mantle melting processes alone.

Following Nakamura and Shimakita (1998), Michael et al. (2002) proposed that HFSE depletion in plagioclase-hosted melt inclusions can result from the slow diffusion of these elements into melt channels formed by plagioclase dissolution that subsequently seal to form melt inclusions. The inclusion-rich, sieve-like texture of plagioclase macrocrysts in the Hvítárvatn tephra layer that host HFSE-depleted inclusions is consistent with inclusion formation by dissolution (Fig. 2b).

Phase equilibria experiments demonstrate that high-anorthite plagioclase is only stable in H2O-poor basalts when melt Ca/Na contents are high (Duncan and Green 1987; Panjasawatwong et al. 1995). In their survey of primitive melts from Iceland, Shorttle and Maclennan (2011) showed that incompatible trace element enrichment correlates negatively with Ca/Na in basalts of constant MgO. It follows that the stability of high-anorthite plagioclase in primitive basalts also correlates negatively with the degree of incompatible trace element enrichment. Crystallisation calculations performed at 3 kbar with fO2 buffered at QFM-1 using Petrolog (Danyushevsky and Plechov, 2011) confirm that high-anorthite plagioclase (>An85) only crystallises from the most incompatible trace element-depleted primitive liquids (Fig. 10a). Starting compositions for these Petrolog calculations were generated by mixing the experimental melts from Walter (1998) and Kogiso et al. (1998) that Shorttle and Maclennan (2011) identified as potential depleted and enriched primary melt compositions for Iceland.
Fig. 10

a Plot demonstrating how the stability and composition of olivine, plagioclase and clinopyroxene vary during the crystallisation of different Icelandic primary melts. Clinopyroxene Mg# is shown as black contours within the clinopyroxene-present field. Crystallisation calculations were performed at a pressure of 3 kbar with fO2 buffered at QFM-1 using Petrolog (Danyushevsky and Plechov 2011). Major element starting compositions were generated by linear mixing of the experimental melts from Walter (1998) and Kogiso et al. (1998) that were identified by Shorttle and Maclennan (2011) as potential depleted and enriched primary melts for Iceland. Trace element ratios were taken from Shorttle and Maclennan (2011). High-anorthite plagioclase and high-Mg# clinopyroxene are only stable in low-La/Yb, high-Ca/Na melts. Transfer of high-anorthite macrocrysts into incompatible trace element-enriched and plagioclase-undersaturated melts by mixing, as indicated by the red arrow, may result in HFSE-depleted melt inclusion formation by crystal dissolution. b BSE image showing a string of aligned melt inclusions that may represent re-annealed melt channels within plagioclase macrocryst cores. Circular features are pits formed during SIMS analyses. A transmitted light photomicrograph of the same grain showing the discrete nature of melt inclusions is shown in the top left

We suggest that high-anorthite macrocrysts, which initially crystallised from high-Ca/Na, low-La/Yb melts, subsequently became unstable following juxtaposition against low-Ca/Na, high-La/Yb and plagioclase-undersaturated melts. Such a juxtaposition could occur via convective stirring in magma reservoirs (Maclennan 2008b), or by flushing melts of variable composition through porous crystal mushes (Danyushevsky et al. 2003). These melt channels formed by plagioclase dissolution would then start to re-equilibrate with these plagioclase-undersaturated liquids. Melt channels would ultimately seal once the surrounding liquid became plagioclase-saturated again, whether by mixing or differentiation, and low-HFSE, high-La/Yb melt inclusions would be formed as a result. Elongate, strung-out chains of melt inclusions in high-anorthite cores may thus represent partially re-annealed melt channels (Fig. 10b). It is important to note that HFSE-undepleted inclusions may have also formed by dissolution, given that complete re-equilibration of HFSEs prior to entrapment can be achieved within hours under magmatic conditions (LaTourrette et al. 1996). Thus, although inclusions may not represent the melts from which their high-anorthite hosts initially crystallised, most primitive plagioclase-hosted melt inclusions nevertheless preserve a record of mantle-derived trace element variability.

Differentiation of primitive melts recorded in melt inclusion major element compositions

Although much of the trace element (i.e. La/Yb) variability in primitive melt inclusions from the Hvítárvatn tephra layer can be accounted for by mantle melting and subsequent mixing (e.g. Maclennan 2008b), these processes are unable to account for all major element diversity. For example, significant variability in primitive plagioclase-hosted melt inclusion Mg# contents (σ t/σ r = 10.1) suggests that major element compositions have been modified by fractionation controlled by ferromagnesian minerals.

Fitting liquid lines of descent between primitive and evolved end-member melt compositions using least-squares methods are a well-established technique for assessing the relative importance of different phases during melt evolution (Wright and Doherty 1970). However, by fitting the first principal component axis (PC1) of a whole data set it is possible to constrain crystallising phase proportions without having to select primitive and evolved end-members a priori (Maclennan et al. 2001; Neave et al. 2012). PC1 was identified in the Hvítárvatn data set by performing principal component analysis (PCA) on inclusions with MgO > 9.5 wt% (Mg# > 62) in the system SiO2–TiO2–Al2O3–FeOt–MgO–CaO–Na2O–K2O (CFNKTMAS). Least-squares fits were then calculated using mean primitive plagioclase, olivine and clinopyroxene compositions.

The best-fitting average liquid line of descent (RMS error = 0.22) indicates that plagioclase, olivine and clinopyroxene crystallised in the ratio 13:7:0. However, PC1 only accounts for 44 % of the total variance in the data set. PC2, which accounts for 23 % of the data set’s variance, is fitted best by crystallisation of plagioclase, olivine and clinopyroxene in the ratio 5:4:11 (RMS error = 0.23). Thus, whilst melt inclusion compositions are dominantly controlled by crystallisation of a troctolitic assemblage, evidence of clinopyroxene crystallisation, most likely from the lowest Mg# primitive melts, is nevertheless recorded. That PC1 and PC2 only account for 67 % of the total variance reflects the combined effects of analytical error, primary melt heterogeneity and PEC on major element variation.

Primary melts from the EVZ

A hypothetical, near-primary melt composition, marked on Fig. 4 as a white diamond, was calculated by adding primitive plagioclase and olivine into the mean primitive plagioclase-hosted melt inclusion composition in a 13:7 ratio until Mg#melt reached 70. Mg# = 70 represents a commonly assumed minimum value for MORB primary melts (Kinzler and Grove 1992) and lies in within the Icelandic primary melt Mg# range estimated by Shorttle and Maclennan (2011): 66–73 at MgO = 11.1 wt%. Whilst this hypothetical melt does not represent a unique primary melt composition for either the 10 ka Grímsvötn tephra series or the EVZ, it does enable a previously elusive near-primary melt composition from EVZ to be compared quantitatively with magma compositions from elsewhere in Iceland.

Major elements in the hypothetical near-primary melt lie largely within the 2σ error of most primitive melt inclusion analyses (Fig. 4). The near-primary melt composition was compared with a large (n = 3480) database of Icelandic whole-rock and glass compositions (see Appendix 1 of Shorttle and Maclennan (2011), with further additions from Koornneef et al. (2012)). Absolute differences between oxides (in wt%) in the hypothetical primary melt (w x ) and database entries (w i ) were summed to give a total misfit (M) as follows:
$$M = \sum\limits_{i} {\left| {w_{i}^{x} - w_{i}^{0} } \right|}$$
where i refers to an oxide in the CFNKTMAS system. Whole-rock compositions from the Borgarhraun flow in the Theistareykir system of the NVZ (Fig. 1; Maclennan et al. 2003) consistently return the lowest misfits (M < 4.5) and are shown in Fig. 4. Whole-rock samples from Borgarhraun also have La/Yb values of ~1.2 that lie within the range defined by primitive inclusions from Hvítárvatn (Maclennan et al. 2003). Eruptions from within the Brennisteinsfjöll volcanic system on the Reykjanes Peninsula also provide close matches (Fig. 1). However, most natural compositions from the NVZ have Al2O3 and CaO contents that are 1–1.5 wt% higher than the hypothetical melt. Explanations for these discrepancies include plagioclase accumulation in Borgarhraun whole-rock samples or plagioclase crystallisation within primitive inclusions from Hvítárvatn. Nevertheless, similarities between this hypothetical near-primary melt and primitive magmas from the Reykjanes Peninsula and NVZ imply that, even if they never reach the surface without modification, high-Mg#, high-Ca/Na and low-La/Yb melts are supplied to base of the EVZ as they are elsewhere in Iceland.

Magma storage conditions of Hvítárvatn magmas

Crystallisation of high-forsterite olivine (~Fo87) and high-anorthite plagioclase (~An92) from compositionally variable but, on average, incompatible trace element-depleted melts, is the highest temperature magmatic process recorded directly in the Hvítárvatn tephra layer. High-Mg# clinopyroxene (Mg#cpx ~ 87) probably joined the liquidus at Mg#melt ~ 65. Disequilibrium between primitive macrocrysts and both evolved macrocrysts (An60–An68, Mg#cpx = 71–78, Fo70–Fo76) and the matrix glass indicates that a low-La/Yb cargo of primitive macrocrysts was mixed into a high-La/Yb, evolved carrier liquid shortly before eruption (c.f. Neave et al. 2014b). There is no petrological record of high-La/Yb primitive melts contributing to the liquids that ultimately formed the matrix glass. High-anorthite plagioclase and high-Mg# clinopyroxene macrocrysts only record the evolution of low-La/Yb melts, indicating that the evolved carrier liquid was aphyric at the time of mixing. Within this petrogenetic framework, there are two key questions that can be interrogated using thermobarometry: first, under what conditions did the evolved assemblage grow and equilibrate with the matrix glass and second, under what conditions did the primitive assemblage form?

Thermal evolution of Hvítárvatn magmas

Matrix glass and plagioclase-hosted melt inclusion temperatures calculated using Eq. 14 of Putirka (2008), a melt composition thermometer (standard error of estimate (SEE) = ± 60 °C), are shown in Fig. 11. Matrix glass and evolved melt inclusion compositions record temperatures of ~1140 °C, whereas primitive plagioclase-hosted melt inclusion compositions mainly record temperatures within the range 1240–1300 °C (Fig. 11).
Fig. 11

Plot summarising the thermal evolution of melt inclusions, clinopyroxenes and the matrix glass in the Hvítárvatn tephra layer. The grey region of the upper plot shows the area enlarged for clarity in the lower plot. Melt inclusion and matrix glass temperatures were calculated using Eq. 15 from Putirka (2008). Olivine–melt temperatures were calculated using Eq. 22 from Putirka (2008) and plagioclase-hosted melt inclusion compositions that have \(K_{{d\;{\text{Fe}}\text{-}{\text{Mg}}}}^{{{\text{ol}}\text{-}{\text{liq}}}}\) values within 10 % of equilibrium according to Herzberg and O’Hara (2002). Clinopyroxene–melt equilibrium temperatures were calculated in two ways using the model of Putirka et al. (1996). First, equilibrium temperatures and pressures were calculated iteratively for each clinopyroxene using melts following the method of Neave et al. (2013) (pale grey triangles). Compared with melt and olivine–melt thermometers, these iterative calculations return high temperatures for evolved macrocrysts. Temperatures were also calculated for evolved clinopyroxenes in equilibrium with the matrix glass by assuming a pressure of 1 kbar (Fig. 12a). Calculations with this assumed pressure return similar temperatures to melt and olivine–melt thermometers (dark grey triangles). Standard errors of estimate are shown

In H2O-poor systems, temperatures calculated using olivine–melt equilibria are more precise than temperatures calculated from melt composition alone (Ford et al. 1983; Beattie 1993; Putirka 2008). Olivine–melt temperatures were calculated with Eq. 22 of Putirka (2008) (SEE = ±27 °C) using plagioclase-hosted melt inclusion and matrix glass compositions that, according to Herzberg and O’Hara (2002), return \(K_{{{\text{d}}\;{\text{Fe}}\text{-}{\text{Mg}}}}^{{{\text{ol}}\text{-}{\text{liq}}}}\) values within 10 % of equilibrium. Olivine-melt temperatures are consistent with melt-only temperatures (Fig. 11): ~Fo85 olivines in equilibrium with Mg#melt = 62 record ~1250 °C and ~Fo72 olivines in equilibrium with Mg#melt = 43 record ~1140 °C.

Two approaches were used to estimate temperatures using clinopyroxene–melt equilibria (Fig. 11). Following the method of Neave et al. (2013), temperatures and pressures were first calculated iteratively using equilibrium melt compositions selected from a database of Icelandic melt compositions (Shorttle and Maclennan 2011). Temperatures and pressures were estimated using the Jd-DiHd thermobarometer of Putirka et al. (1996) (SEE = ±46 °C) because it is calibrated specifically for H2O-poor, mafic compositions. High-Mg# clinopyroxene macrocrysts return temperatures of 1200–1240 °C, consistent with estimates from melt inclusions and olivine macrocrysts. However, low-Mg# clinopyroxene macrocrysts return temperatures of 1160–1190 °C that are substantially higher than those obtained for the evolved assemblage using other thermometers (~1140 °C). This overestimation results primarily from the model’s convergence to overly high pressures (see below). In the second approach, clinopyroxenes in Fe–Mg equilibrium with the matrix glass return temperatures of ~1150 °C when an equilibration pressure of 1 kbar is assumed (see below).

Numerous thermometers thus indicate that high-anorthite plagioclase-hosted melt inclusions and high-forsterite olivine macrocrysts were derived from high-temperature, near-primary melts (≥1240 °C). High-Mg# clinopyroxene macrocryst temperatures are consistent with crystallisation alongside primitive olivine and plagioclase. Pre-eruptive temperatures of ~1140 °C recorded by the evolved assemblage are consistent with those calculated for the compositionally similar AD 1783–1784 Laki eruption (Guilbaud et al. 2007; Neave et al. 2013).

Magma storage depths and barometric inconsistencies

Melt inclusion CO2 contents were used to calculate entrapment pressures with the Π* parameter model of Shishkina et al. (2014) developed from the Π parameter model of Dixon (1997) with relative errors of <20 % based on uncertainties in model fits (Shishkina et al. 2014). Estimating entrapment pressures from melt volatile contents assumes that melts were fluid-saturated at the time of inclusion formation (Moore 2008). Whilst this assumption may not be valid in all circumstances (e.g. Saal et al. 2002), the low CO2/Nb content of HFSE-undepleted primitive melt inclusions from Hvítárvatn (mean CO2/Nb = 159 ± 74(1σ)) indicates that trapped melts achieved CO2 saturation before eruption: prior to volatile exsolution, Icelandic primary melt CO2/Nb contents are likely to be ≥ 300 (Hauri et al. 2002; Hartley et al. 2014; Neave et al. 2014a). Evolved macrocryst-hosted melt inclusions record entrapment pressures of 0.7–1.5 kbar (Fig. 12). Most primitive melt inclusions have apparent entrapment pressures of 0.5–2.0 kbar, with a mean of 1.4 kbar. Eight inclusions record notably higher pressures of 2.0–5.2 kbar (Fig. 12).
Fig. 12

a Plot summarising the results of barometric calculations carried out on different components of the Hvítárvatn tephra layer. Melt inclusion entrapment pressures were calculated using the CO2 solubility model of Shishkina et al. (2014) and are shown as dark red circles. The <20 % error in melt inclusion entrapment pressures is based on the model’s ability to fit experimental data (Shishkina et al. 2014). Equilibration pressures of three-phase saturated matrix glass compositions and Mg#melt < 65 melt inclusions were calculated using the olivine–plagioclase–augite–melt (OPAM) boundary barometer of Yang et al. (1996) and are shown as red and pale red boxes. Clinopyroxene–melt equilibrium pressures were calculated iteratively for all clinopyroxene analyses using the model of Putirka et al. (1996) and melts selected from a large database of Iceland melt compositions following the method of Neave et al. (2013) (pale grey triangles). Equilibrium pressures were also calculated for clinopyroxene macrocrysts in equilibrium with the matrix glass using an imposed temperature of 1140 °C (dark grey triangles). A correction accounting for the systematic overestimation of pressure in experimental clinopyroxene–melt pairs was applied to both iteratively calculated primitive clinopyroxene pressures and imposed temperature evolved clinopyroxene pressures (pale and dark grey triangles with black outlines, respectively). The standard error of estimate for the Putirka et al. (1996) model is shown. b Plot comparing calculated and observed pressures for equilibrium clinopyroxene–melt pairs from experiments carried out in the pressure range 1–8 kbar. Experimental data were collated from Grove et al. (1992), Feig et al. (2006, 2010) and Villiger et al. (2007) and are not used in the calibration of the Putirka et al. (1996) model. Corrected clinopyroxene pressures shown in (a) were calculated using the regression line shown as a dotted line. The crossed data point was excluded from the regression

As described above, clinopyroxene–melt pressures were calculated using the method of Neave et al. (2013) and the model of Putirka et al. (1996) (SEE = ±1.4 kbar). Calculations performed on high-Al zones of low-Mg# clinopyroxenes return pressures of 4.2–6.4 kbar, with a mean of 5.3 kbar. Calculations performed on high-Mg# clinopyroxenes return higher pressures of 5.5–7.5 kbar, with a mean of 6.4 kbar (Fig. 12a). Pressures calculated using clinopyroxene–melt equilibria are thus up to ~5 kbar higher than apparent melt inclusion entrapment pressures from the same macrocryst assemblage. In order to evaluate the depth at which magmas feeding the Hvítárvatn tephra eruption were stored and processed satisfactorily, this discrepancy needs to be resolved.

The pressure of last mineral–melt equilibration of three-phase-saturated basalts can be estimated with an approximate uncertainty of 1 kbar using the location of the olivine–plagioclase–augite–melt (OPAM) boundary (Yang et al. 1996). OPAM boundary locations for matrix glass compositions indicate that they last equilibrated with their crystal load at 1 ± 1 kbar (Fig. 12a), consistent with the 0.7–1.5 kbar pressure recorded by the volatile content of the evolved melt inclusions. The difference between clinopyroxene–melt and melt inclusion/OPAM pressure estimates is beyond quoted model uncertainties, suggesting that clinopyroxene–melt pressures are incorrect.

If clinopyroxene–melt pressures are recalculated using an assumed temperature of 1140 °C, pressures drop from 4.2–6.4 kbar to 3.6–4.1 kbar, which are still much greater than the 0.7–1.5 kbar pressures recorded by evolved melt inclusions. However, Putirka (2008) noted that the model of Putirka et al. (1996) generally returns pressures that are 2–3 kbar too high for most 1 atm experiments. Whilst Putirka (2008) attributed this offset to Na volatilisation in furnace experiments (e.g. Tormey et al. 1987), the offset is still present when the thermobarometer is applied to clinopyroxene–melt pairs from experiments carried out at 1–7 kbar (Fig. 12b; Grove et al. 1992; Feig et al. 2006; Villiger et al. 2007; Feig et al. 2010). Nevertheless, if a pressure-dependent correction factor is applied to equilibrium low-Mg# clinopyroxene-matrix glass pairs a posteriori (see regression in Fig. 12b), then the pressure range drops from 3.6–4.1 kbar to 0.0–0.7 kbar (Fig. 12a).

Assuming negligible ice cover and an average crustal density of 2700 kg.m−3, a pressure of ~0.7 kbar corresponds to a depth of ~2.6 km. However, we do not interpret this as a depth of long-term storage for the following reasons: first, the volume of each eruption in the 10 ka Grímsvötn tephra series (>1–30 km3; Thordarson 2014) is at least comparable with that of the currently active shallow melt lens (~10 km3; Alfaro et al. 2007). Assuming that Grímsvötn’s subsurface architecture has not changed substantially over the past ~10 ka, this comparison indicates that recharge from depth must have been efficient and continuous in order to avoid magma exhaustion on both intra- and inter-eruption timescales. Second, evolved macrocrysts textures—skeletal plagioclase rims (Fig. 2a) and clinopyroxene sector zoning (Fig. 2c)—are consistent with rapid crystallisation imminently before eruption. We therefore suggest that the evolved assemblage grew within a 0.0–1.5 kbar pressure interval following the mixing of primitive macrocrysts into an evolved, aphyric melt that differentiated deeper in the plumbing system (Neave et al. 2014b). Although this crystallisation may have taken place as magma stalled in ephemeral, small reservoirs whilst en route to the surface, it may also have taken place in transit.

Reconciling differences between primitive melt inclusion entrapment pressures and high-Mg# clinopyroxene–melt pressures are more difficult. Applying a correction to high-Mg# clinopyroxene–melt calculations yields mid-crustal pressures of 2.5–5.5 kbar (Fig. 12a). However, given the significant uncertainties in the correction factor, clinopyroxene–melt pressures alone cannot be considered reliable. Through careful application of the OPAM barometer to melt inclusion compositions that were three-phase saturated at the time of entrapment, it is possible to place additional constraints on magma storage depths. Low-Mg# primitive inclusions (Mg#melt < 65) return pressures of 4 ± 1.5 kbar which are consistent with corrected clinopyroxene–melt pressures (Fig. 12a) and the appearance of clinopyroxene on the liquidus after plagioclase (Presnall et al. 1978; Grove et al. 1992; Villiger et al. 2007).

The persistent discrepancy between melt inclusion entrapment pressures and clinopyroxene–melt pressures following the validation of corrected clinopyroxene–melt pressures with the OPAM barometer thus calls the validity of primitive melt inclusion entrapment pressures into question. Considering the convergence of mineral–melt equilibrium and melt composition barometers on 4 ± 1.5 kbar, we suggest that most melt inclusion entrapment pressures no longer record the pressure of initial entrapment. Whilst a few high-CO2 inclusions (800–2100 ppm) may preserve true entrapment pressures of up to 5.2 kbar, we suggest that the mean entrapment pressure of 1.4 kbar records pre-eruptive CO2 loss within the shallow plumbing system, possibly during formation of the evolved assemblage.

Whilst a number of studies have investigated the post-entrapment behaviour of volatiles in olivine-hosted melt inclusions (e.g. Qin et al. 1992; Portnyagin et al. 2008; Bucholz et al. 2013; Hartley et al. 2015), volatile behaviour in plagioclase-hosted melt inclusions is less well understood. Whilst the high H2O content of melt inclusions precludes CO2 loss during the final stages of eruption, it is not currently possible to evaluate how CO2 may be transferred between inclusions and the external melt through plagioclase hosts, though diffusion along lattice defects and cleavage planes is likely to play a role (Viti and Frezzotti 2001; Kress and Ghiorso 2004). Nevertheless, once experimentally quantified, volatile exchange through plagioclase macrocrysts has the chronometric potential to reveal rates of immediately pre-eruptive processes.

Although interpreting the raw results of our barometric calculations at face value leads to barometric estimates that span the entire Icelandic crust for a single equilibrium assemblage (Darbyshire et al. 2000), re-assessing the reliability of pressure estimates on a barometer-by-barometer basis enables an internally consist model of magma storage to be developed. Multiple barometers indicate that the primitive macrocryst assemblage formed in the mid-crust at 4 ± 1.5 kbar (13.6 ± 5.1 km), whereas the evolved assemblage grew within the shallow crust at 0.0–1.5 kbar (0–5.2 km). Mid-crustal pressures are consistent with the depths of recent seismogenic magmatism in Iceland (Sigmundsson et al. 2010; Tarasewicz et al. 2014; Guðmundsson et al. 2014), as well as with the depth at which ground deformation data suggest a deep reservoir may currently exist under Grímsvötn (Reverso et al. 2014). Uncertainties nevertheless remain in the calibration of clinopyroxene–melt equilibrium and the geological significance of the shallow entrapment pressures recorded by most plagioclase-hosted melt inclusions. In order to resolve the diverse pressure estimates obtained from different components of the Hvítárvatn tephra layer, further phase equilibria experiments within the crucial 1–7 kbar pressure interval are required to refine thermobarometric calibrations. The improvement of barometer calibrations is vital if petrological data are to be used to inform ongoing volcano monitoring strategies.


Matrix glass compositions from the ~10 ka Grímsvötn tephra series (i.e. Saksunarvatn ash) at Hvítárvatn, central Iceland are consistent with an affinity with the Grímsvötn volcanic system. Macrocrysts can be divided into two populations based on their degree of evolution and incompatible trace element enrichment. The evolved assemblage contains low-anorthite plagioclase, low-Mg# clinopyroxene, low-forsterite olivine macrocrysts and macrocryst rims that are close to being in equilibrium with the matrix glass. The primitive assemblage contains high-anorthite plagioclase, low-Mg# clinopyroxene and low-forsterite olivine macrocryst cores that are in equilibrium with primitive, incompatible trace element-depleted melt inclusions. Trace element disequilibrium between primitive and evolved assemblages probably resulted from the disaggregation of primitive crystal mushes into evolved and aphyric melts shortly before eruption. Although a subset of primitive plagioclase-hosted melt inclusions exhibits HFSE depletions that reflect their formation by macrocryst dissolution within plagioclase-undersaturated melts, post-entrapment modification of major and trace element compositions was otherwise minimal. Primitive plagioclase-hosted inclusions record the supply of compositionally diverse mantle melts to the base of the Grímsvötn plumbing system that subsequently underwent fractional crystallisation dominated by the removal of an initially troctolitic then gabbroic assemblage. Primitive eruptives from the NVZ and Reykjanes Peninsula provide the closest analogues to a hypothetical near-primary melt composition for the Hvítárvatn primitive assemblage.

Final pre-eruptive mineral–melt equilibration of evolved assemblage occurred at 1140 °C and 0.0–1.5 kbar. The large volume of the tephra-forming eruption and the presence of textures consistent with rapid crystallisation within the evolved macrocrysts are inconsistent with 0.0–1.5 kbar representing a depth of long-term magma storage. Numerous thermometers indicate that primitive macrocrysts and melt inclusions they contain formed at much higher temperatures than the evolved assemblage: 1240–1300 °C. Although most primitive melt inclusions record apparent entrapment pressures of 1.4 kbar based on their CO2 content, other barometers consistently return mid-crustal pressures of 4 ± 1.5 kbar when evaluated critically. Barometric convergence on mid-crustal pressures indicates that most primitive melt inclusions lost CO2 during transport through the shallow crust. Assessing the accuracy and precision of mid-crustal clinopyroxene–melt pressure estimates is nevertheless impaired by incomplete calibration of barometers under the conditions relevant to magmatic processes beneath Iceland. In order to improve the fidelity of petrological estimates of magma storage depths, continued exploration of basalt phase equilibria in the crucial 1–7 kbar pressure interval is required.



D.A.N. was supported by a Natural Environment Research Council studentship (NE/1528277/1) at the start of this project. SIMS analyses were supported by Natural Environment Research Council Ion Microprobe Facility award (IMF508/1013). We would like to thank Robin Clarke and Iris Buisman for their help with sample preparation and EPMA at the University of Cambridge. We also thank Richard Hinton and all the staff at the Ion Microprobe Facility at University of Edinburgh for their assistance with SIMS analyses. We are very grateful to two anonymous reviewers for their detailed and highly constructive reviews as well as to Othmar Müntener for his efficient editorial handling.

Supplementary material

410_2015_1170_MOESM1_ESM.xls (88 kb)
Supplementary material 1 (XLS 87 kb)
410_2015_1170_MOESM2_ESM.xls (154 kb)
Supplementary material 2 (XLS 154 kb)
410_2015_1170_MOESM3_ESM.eps (418 kb)
Supplementary material 3 (EPS 418 kb)


  1. Aigner-Torres M, Blundy JD, Ulmer P, Pettke T (2007) Laser ablation ICPMS study of trace element partitioning between plagioclase and basaltic melts: an experimental approach. Contrib Mineral Petrol 153:647–667. doi: 10.1007/s00410-006-0168-2 CrossRefGoogle Scholar
  2. Alfaro R, Brandsdóttir B, Rowlands DP et al (2007) Structure of the Grímsvötn central volcano under the Vatnajökull icecap, Iceland. Geophys J Int 168:863–876. doi: 10.1111/j.1365-246X.2006.03238.x CrossRefGoogle Scholar
  3. Anderson AT, Brown GG (1993) CO2 contents and formation pressures of some Kilauean melt inclusions. Am Mineral 78:794–803Google Scholar
  4. Arevalo R, McDonough WF (2010) Chemical variations and regional diversity observed in MORB. Chem Geol 271:70–85. doi: 10.1016/j.chemgeo.2009.12.013 CrossRefGoogle Scholar
  5. Beattie P (1993) Olivine-melt and orthopyroxene-melt equilibria. Contrib Mineral Petrol 115:103–111CrossRefGoogle Scholar
  6. Beattie P (1994) Systematics and energetics of trace-element partitioning between olivine and silicate melts: implications for the nature of mineral/melt partitioning. Chem Geol 117:57–71CrossRefGoogle Scholar
  7. Bédard JH (2006) Trace element partitioning in plagioclase feldspar. Geochim Cosmochim Acta 70:3717–3742. doi: 10.1016/j.gca.2006.05.003 CrossRefGoogle Scholar
  8. Björck S, Ingólfsson Ó, Haflidason H et al (1992) Lake Torfadalsvatn: a high resolution record of the North Atlantic ash zone I and the last glacial–interglacial environmental changes in Iceland. Boreas 21:15–22. doi: 10.1111/j.1502-3885.1992.tb00009.x CrossRefGoogle Scholar
  9. Blundy JD, Wood BJ (1991) Crystal-chemical controls on the partitioning of Sr and Ba between plagioclase feldspar, silicate melts, and hydrothermal solutions. Geochim Cosmochim Acta 55:193–209. doi: 10.1016/0016-7037(91)90411-W CrossRefGoogle Scholar
  10. Blundy JD, Wood BJ (1994) Prediction of crystal melt partition coefficients from elastic moduli. Nature 372:452–454CrossRefGoogle Scholar
  11. Bramham-Law CWF, Theuerkauf M, Lane CS, Mangerud J (2013) New findings regarding the Saksunarvatn Ash in Germany. J Quat Sci 28:248–257. doi: 10.1002/jqs.2615 CrossRefGoogle Scholar
  12. Breddam K (2002) Kistufell: primitive melt from the Iceland Mantle Plume. J Petrol 43:345–373. doi: 10.1093/petrology/43.2.345 CrossRefGoogle Scholar
  13. Brice JC (1975) Some thermodynamic aspects of the growth of strained crystals. J Cryst Growth 28:249–253CrossRefGoogle Scholar
  14. Bucholz CE, Gaetani GA, Behn MD, Shimizu N (2013) Post-entrapment modification of volatiles and oxygen fugacity in olivine-hosted melt inclusions. Earth Planet Sci Lett 374:145–155. doi: 10.1016/j.epsl.2013.05.033 CrossRefGoogle Scholar
  15. Cottrell E, Spiegelman M, Langmuir CH (2002) Consequences of diffusive reequilibration for the interpretation of melt inclusions. Geochem Geophys Geosyst 3:1–26. doi: 10.1029/2001GC000205 CrossRefGoogle Scholar
  16. Danyushevsky LV, Plechov P (2011) Petrolog 3: integrated software for modelling crystallization processes. Geochem Geophys Geosyst 12:Q07021. doi: 10.1029/2011GC003516 CrossRefGoogle Scholar
  17. Danyushevsky LV, Della-Pasqua FN, Sokolov S (2000) Re-equilibration of melt inclusions trapped by magnesian olivine phenocrysts from subduction-related magmas: petrological implications. Contrib Mineral Petrol 138:68–83CrossRefGoogle Scholar
  18. Danyushevsky LV, McNeill AW, Sobolev AV (2002) Experimental and petrological studies of melt inclusions in phenocrysts from mantle-derived magmas: an overview of techniques, advantages and complications. Chem Geol 183:5–24. doi: 10.1016/S0009-2541(01)00369-2 CrossRefGoogle Scholar
  19. Danyushevsky LV, Perfit MR, Eggins SM, Falloon TJ (2003) Crustal origin for coupled “ultra-depleted” and “plagioclase” signatures in MORB olivine-hosted melt inclusions: evidence from the Siqueiros Transform Fault, East Pacific Rise. Contrib Mineral Petrol 144:619–637. doi: 10.1007/s00410-002-0420-3 CrossRefGoogle Scholar
  20. Darbyshire FA, White RS, Priestley KF (2000) Structure of the crust and uppermost mantle of Iceland from a combined seismic and gravity study. Earth Planet Sci Lett 181:409–428. doi: 10.1016/S0012-821X(00)00206-5 CrossRefGoogle Scholar
  21. Dixon JE (1997) Degassing of alkalic basalts. Am Mineral 82:368–378Google Scholar
  22. Dixon JE, Stolper EM, Holloway JR (1995) An experimental study of water and carbon dioxide solubilities in mid-ocean ridge basaltic liquids. Part I: calibration and solubility models. J Petrol 36:1607–1631Google Scholar
  23. Duncan R, Green D (1987) The genesis of refractory melts in the formation of oceanic crust. Contrib Mineral Petrol 96:326–342CrossRefGoogle Scholar
  24. Dungan MA, Rhodes JM (1978) Residual glasses and melt inclusions in basalts from DSDP Legs 45 and 46: evidence for magma mixing. Contrib Mineral Petrol 431:417–431CrossRefGoogle Scholar
  25. Elliott TR, Hawkesworth CJ, Grönvold K (1991) Dynamic melting of the Iceland plume. Nature 351:201–206CrossRefGoogle Scholar
  26. Feig ST, Koepke J, Snow JE (2006) Effect of water on tholeiitic basalt phase equilibria: an experimental study under oxidizing conditions. Contrib Mineral Petrol 152:611–638. doi: 10.1007/s00410-006-0123-2 CrossRefGoogle Scholar
  27. Feig ST, Koepke J, Snow JE (2010) Effect of oxygen fugacity and water on phase equilibria of a hydrous tholeiitic basalt. Contrib Mineral Petrol 160:551–568. doi: 10.1007/s00410-010-0493-3 CrossRefGoogle Scholar
  28. Ford CE, Russell DG, Craven JA, Fisk MR (1983) Olivine-liquid equilibria: temperature, pressure and composition dependence of the crystal/liquid cation partition coefficients for Mg, Fe2+, Ca and Mn. J Petrol 24:256–265CrossRefGoogle Scholar
  29. Gaetani GA, Watson EB (2000) Open system behavior of olivine-hosted melt inclusions. Earth Planet Sci Lett 183:27–41. doi: 10.1016/S0012-821X(00)00260-0 CrossRefGoogle Scholar
  30. Ghiorso MS, Hirschmann MM, Reiners PW, Kress VC (2002) The pMELTS: a revision of MELTS for improved calculation of phase relations and major element partitioning related to partial melting of the mantle to 3 GPa. Geochem Geophys Geosyst 3:1030. doi: 10.1029/2001GC000217 CrossRefGoogle Scholar
  31. Grönvold K, Óskarsson N, Johnsen S et al (1995) Ash layers from Iceland in the Greenland GRIP ice core correlated with oceanic and land sediments. Earth Planet Sci Lett 35:149–155CrossRefGoogle Scholar
  32. Grove TL, Kinzler RJ, Bryan WB (1992) Fractionation of mid-ocean ridge basalt (MORB). Geophys Monogr Geophys Union 71:281–310Google Scholar
  33. Guðmundsson Á, Lecoeur N, Mohajeri N, Thordarson T (2014) Dike emplacement at Bardarbunga, Iceland, induces unusual stress changes, caldera deformation, and earthquakes. Bull Volcanol 76:869. doi: 10.1007/s00445-014-0869-8 CrossRefGoogle Scholar
  34. Guilbaud M-N, Blake S, Thordarson T, Self S (2007) Role of Syn-eruptive cooling and degassing on textures of lavas from the AD 1783–1784 Laki eruption, South Iceland. J Petrol 48:1265–1294. doi: 10.1093/petrology/egm017 CrossRefGoogle Scholar
  35. Gurenko AA, Chaussidon M (1995) Enriched and depleted primitive melts included in olivine from Icelandic tholeiites—origin by melting of a single mantle column. Geochim Cosmochim Acta 59:2905–2917CrossRefGoogle Scholar
  36. Gurenko AA, Sobolev AV (2006) Crust-primitive magma interaction beneath neovolcanic rift zone of Iceland recorded in gabbro xenoliths from Midfell, SW Iceland. Contrib Mineral Petrol 151:495–520. doi: 10.1007/s00410-006-0079-2 CrossRefGoogle Scholar
  37. Haflidason H, Eiriksson J, Van Kreveld S (2000) The tephrochronology of Iceland and the North Atlantic region during the middle and Late Quaternary: a review. J Quat Sci 15:3–22. doi: 10.1002/(SICI)1099-1417(200001)15:1<3:AID-JQS530>3.0.CO;2-W CrossRefGoogle Scholar
  38. Halldórsson SA, Óskarsson N, Grönvold K et al (2008) Isotopic-heterogeneity of the Thjorsa lava—implications for mantle sources and crustal processes within the Eastern Rift Zone, Iceland. Chem Geol 255:305–316. doi: 10.1016/j.chemgeo.2008.06.050 CrossRefGoogle Scholar
  39. Hansen H, Grönvold K (2000) Plagioclase ultraphyric basalts in Iceland: the mush of the rift. J Volcanol Geotherm Res 98:1–32. doi: 10.1016/S0377-0273(99)00189-4 CrossRefGoogle Scholar
  40. Hart SR, Dunn T (1993) Experimental cpx/melt partitioning of 24 trace elements. Contrib Mineral Petrol 113:1–8. doi: 10.1007/BF00320827 CrossRefGoogle Scholar
  41. Hartley ME, Maclennan J, Edmonds M, Thordarson T (2014) Reconstructing the deep CO2 degassing behaviour of large basaltic fissure eruptions. Earth Planet Sci Lett 393:120–131. doi: 10.1016/j.epsl.2014.02.031 CrossRefGoogle Scholar
  42. Hartley ME, Neave DA, Maclennan J et al (2015) Diffusive over-hydration of olivine-hosted melt inclusions. Earth Planet Sci Lett 425:168–178. doi: 10.1016/j.epsl.2015.06.008 CrossRefGoogle Scholar
  43. Hauri EH, Grönvold K, Óskarsson N, McKenzie D (2002) Abundance of carbon in the Icelandic mantle: constraints from melt inclusions. Am Geophys Union, Spring Meet, p V51D-03Google Scholar
  44. Herzberg C, O’Hara MJ (2002) Plume-associated ultramafic magmas of Phanerozoic Age. J Petrol 43:1857–1883. doi: 10.1093/petrology/43.10.1857 CrossRefGoogle Scholar
  45. Jakobsson SP (1979) Petrology of recent basalts of the Eastern Volcanic Zone, Iceland. Acta Nat Islandica 26:1–103Google Scholar
  46. Jakobsson SP, Jónsson J, Shido F (1978) Petrology of the western Reykjanes peninsula, Iceland. J Petrol 19:669–705CrossRefGoogle Scholar
  47. Jennings A, Thordarson T, Zalzal K et al (2014) Holocene tephra from Iceland and Alaska in SE Greenland Shelf Sediments. Mar Tephrochronol. doi: 10.1144/SP398.6 Google Scholar
  48. Jochum KP, Willbold M, Raczek I et al (2005) Chemical characterisation of the USGS reference glasses and BIR-1G Using EPMA, ID-TIMS, ID-ICP-MS and LA-ICP-MS. Geostand Geoanal Res 29:285–302CrossRefGoogle Scholar
  49. Jochum KP, Stoll B, Herwig K et al (2006) MPI-DING reference glasses for in situ microanalysis: new reference values for element concentrations and isotope ratios. Geochem Geophys Geosyst 7:Q02008CrossRefGoogle Scholar
  50. Jochum KP, Weis U, Stoll B et al (2011) Determination of reference values for NIST SRM 610-617 glasses following ISO guidelines. Geostand Geoanal Res 35:397–429. doi: 10.1111/j.1751-908X.2011.00120.x CrossRefGoogle Scholar
  51. Jóhannsdóttir GE (2007) Mid Holocene to late glacial tephrochronology in West Iceland as revealed in three lacustrine environments. Masters Thesis, University of IcelandGoogle Scholar
  52. Jóhansen J (1975) Pollen diagrams from the Shetland and Faroe Islands. New Phytol 75:369–387CrossRefGoogle Scholar
  53. Jóhansen J (1985) Studies in the vegatation history of the Faroe and Shetland Islands. Føroya Fróðskaparfelag, TórshavnGoogle Scholar
  54. Jóhannsdóttir GE, Thordarson T, Geirsdóttir A, Larsen G (2005) The widespread ~10 ka Saksunarvatn Tephra : a product of three large basaltic Phreatoplinian eruptions? Geophys Res Abstr 7:05991Google Scholar
  55. Jude-Eton TC, Thordarson T, Guðmundsson MT, Oddsson B (2012) Dynamics, stratigraphy and proximal dispersal of supraglacial tephra during the ice-confined 2004 eruption at Grímsvötn Volcano, Iceland. Bull Volcanol 74:1057–1082. doi: 10.1007/s00445-012-0583-3 CrossRefGoogle Scholar
  56. Kilgour GN, Saunders KE, Blundy JD et al (2014) Timescales of magmatic processes at Ruapehu volcano from diffusion chronometry and their comparison to monitoring data. J Volcanol Geotherm Res 288:62–75. doi: 10.1016/j.jvolgeores.2014.09.010 CrossRefGoogle Scholar
  57. Kinzler RJ, Grove TL (1992) Primary magmas of mid-ocean ridge basalts 1. Experiments and methods. J Geophys Res Solid Earth 97:6885–6906CrossRefGoogle Scholar
  58. Kogiso T, Hirose K, Takahashi E (1998) Melting experiments on homogeneous mixtures of peridotite and basalt: application to the genesis of ocean island basalts. Earth Planet Sci Lett 162:45–61CrossRefGoogle Scholar
  59. Koornneef JM, Stracke A, Bourdon B et al (2012) Melting of a two-component source beneath Iceland. J Petrol 53:127–157. doi: 10.1093/petrology/egr059 CrossRefGoogle Scholar
  60. Kress VC, Ghiorso MS (2004) Thermodynamic modeling of post-entrapment crystallization in igneous phases. J Volcanol Geotherm Res 137:247–260. doi: 10.1016/j.jvolgeores.2004.05.012 CrossRefGoogle Scholar
  61. Langmuir CH, Klein EM, Plank T (1992) Petrological systematics of mid-ocean ridge basalts: constraints on melt generation beneath ocean ridges. Mantle flow and melt generation at mid-ocean ridges, 71st edn. Geophys Monogr. American Geophysical Union, pp 183–280Google Scholar
  62. Larsen G (1984) Recent volcanic history of the Veidivötn fissure swarm, southern Iceland—an approach to volcanic risk assessment. J Volcanol Geotherm Res 22:33–58CrossRefGoogle Scholar
  63. Larsen G (2005) Explosive volcanism in Iceland: three examples of hydromagmatic basaltic eruptions on long volcanic fissures within the past 1200 years. Geophys Res Abstr 7:10158Google Scholar
  64. LaTourrette T, Wasserburg G, Fahey A (1996) Self diffusion of Mg, Ca, Ba, Nd, Yb, Ti, Zr, and U in haplobasaltic melt. Geochim Cosmochim Acta 60:1329–1340CrossRefGoogle Scholar
  65. Lawson I, Gathorne-Hardy F, Church M et al (2007) Environmental impacts of the Norse settlement: palaeoenvironmental data from Mývatnssveit, northern Iceland. Boreas 36:1–19. doi: 10.1080/03009480600827298 Google Scholar
  66. Lofgren GE (1974) An experimental study of plagioclase morphology: isothermal crystallisation. Am J Sci 274:243–273CrossRefGoogle Scholar
  67. Lohne ØS, Mangerud J, Birks HH (2014) IntCal13 calibrated ages of the Vedde and Saksunarvatn ashes and the Younger Dryas boundaries from Kråkenes, western Norway. J Quat Sci 29:506–507. doi: 10.1002/jqs.2722 CrossRefGoogle Scholar
  68. Lowenstern J (1995) Applications of silicate-melt inclusions to the study of magmatic volatiles. In: Thompson JFH (ed) Magmas, fluids and ore deposits. Mineral Association Canada Short Course Series 23, pp 71–99Google Scholar
  69. Maclennan J (2008a) Lead isotope variability in olivine-hosted melt inclusions from Iceland. Geochim Cosmochim Acta 72:4159–4176CrossRefGoogle Scholar
  70. Maclennan J (2008b) Concurrent mixing and cooling of melts under Iceland. J Petrol 49:1931–1953. doi: 10.1093/petrology/egn052 CrossRefGoogle Scholar
  71. Maclennan J, McKenzie D, Grönvold K (2001) Plume-driven upwelling under central Iceland. Earth Planet Sci Lett 194:67–82CrossRefGoogle Scholar
  72. Maclennan J, McKenzie D, Hilton F et al (2003) Geochemical variability in a single flow from northern Iceland. J Geophys Res Solid Earth 108:1–21. doi: 10.1029/2000JB000142 Google Scholar
  73. Mangerud J, Furnes H, Jóhansen J (1986) A 9000-year-old ash bed on the Faroe Islands. Quat Res 265:262–265CrossRefGoogle Scholar
  74. McKenzie D, O’Nions RK (1991) Partial melt distributions from inversion of rare earth element concentrations. J Petrol 32:1021CrossRefGoogle Scholar
  75. Merkt J, Müller H, Knabe W et al (1993) The early Holocene Saksunarvatn tephra found in lake sediments in NW Germany. Boreas 22:93–100CrossRefGoogle Scholar
  76. Métrich N, Wallace PJ (2008) Volatile abundances in basaltic magmas and their degassing paths tracked by melt inclusions. Rev Mineral Geochem 69:363–402. doi: 10.2138/rmg.2008.69.10 CrossRefGoogle Scholar
  77. Michael PJ, McDonough WF, Nielsen RL, Cornell WC (2002) Depleted melt inclusions in MORB plagioclase: messages from the mantle or mirages from the magma chamber? Chem Geol 183:43–61. doi: 10.1016/S0009-2541(01)00371-0 CrossRefGoogle Scholar
  78. Moore G (2008) Interpreting H2O and CO2 contents in melt inclusions: constraints from solubility experiments and modeling. Rev Mineral Geochem 69:333–362. doi: 10.2138/rmg.2008.69.9 CrossRefGoogle Scholar
  79. Moore JG, Calk LC (1991) Degassing and differentiation in subglacial volcanoes, Iceland. J Volcanol Geotherm Res 46:157–180. doi: 10.1016/0377-0273(91)90081-A CrossRefGoogle Scholar
  80. Moore LR, Gazel E, Tuohy R et al (2015) Bubbles matter: an assessment of the contribution of vapor bubbles to melt inclusion volatile budgets. Am Mineral 100:806–823CrossRefGoogle Scholar
  81. Nakamura Y (1973) Origin of sector-zoning of igneous clinopyroxenes. Am Mineral 58:986–990Google Scholar
  82. Nakamura M, Shimakita S (1998) Dissolution origin and syn-entrapment compositional change of melt inclusion in plagioclase. Earth Planet Sci Lett 161:119–133. doi: 10.1016/S0012-821X(98)001447 CrossRefGoogle Scholar
  83. Namur O, Charlier B, Toplis M, Vander Auwera J (2011) Prediction of plagioclase-melt equilibria in anhydrous silicate melts at 1-atm. Contrib Mineral Petrol 163:133–150. doi: 10.1007/s00410-011-0662-z CrossRefGoogle Scholar
  84. Neave DA, Fabbro G, Herd RA et al (2012) Melting, differentiation and degassing at the Pantelleria Volcano, Italy. J Petrol 53:637–663. doi: 10.1093/petrology/egr074 CrossRefGoogle Scholar
  85. Neave DA, Passmore E, Maclennan J et al (2013) Crystal-melt relationships and the record of deep mixing and crystallization in the AD 1783–84 Laki eruption, Iceland. J Petrol 54:1661–1690. doi: 10.1093/petrology/egt027 CrossRefGoogle Scholar
  86. Neave DA, Maclennan J, Edmonds M, Thordarson T (2014a) Melt mixing causes negative correlation of trace element enrichment and CO2 content prior to an Icelandic eruption. Earth Planet Sci Lett 400:272–283CrossRefGoogle Scholar
  87. Neave DA, Maclennan J, Hartley ME et al (2014b) Crystal storage and transfer in basaltic systems: the Skuggafjöll eruption, Iceland. J Petrol 55:2311–2356. doi: 10.1093/petrology/egu058 CrossRefGoogle Scholar
  88. Nielsen RL (1995) Local diversity of MORB parent magmas: evidence from melt inclusions in high—an feldspar from the Gorda Ridge. Contrib Mineral Petrol 122:34–50. doi: 10.1180/minmag.1994.58A.2.75 CrossRefGoogle Scholar
  89. Nielsen RL (2011) The effects of re-homogenization on plagioclase hosted melt inclusions. Geochem Geophys Geosyst 12:Q0AC17. doi: 10.1029/2011GC003822 CrossRefGoogle Scholar
  90. Nimis P (1995) A clinopyroxene geobarometer for basaltic systems based on crystal-structure modeling. Contrib Mineral Petrol 121:115–125CrossRefGoogle Scholar
  91. Óladottir BA, Thordarson T, Larsen G, Sigmarsson O (2007) Survival of the Myrdalsjökull ice cap through the Holocene thermal maximum: evidence from sulphur contents in Katla tephra layers (Iceland) from the last 8400 years. Ann Glaciol 45:183–188CrossRefGoogle Scholar
  92. Óladóttir BA, Sigmarsson O, Larsen G, Thordarson T (2007) Katla volcano, Iceland: magma composition, dynamics and eruption frequency as recorded by Holocene tephra layers. Bull Volcanol 70:475–493. doi: 10.1007/s00445-007-0150-5 CrossRefGoogle Scholar
  93. Óskarsson N, Helgason Ö, Steinthórsson S (1994) Oxidation state of iron in mantle-derived magmas of the Icelandic rift zone. Hyperfine Interact 91:733–737. doi: 10.1007/BF02064599 CrossRefGoogle Scholar
  94. Panjasawatwong Y, Danyushevsky LV, Crawford AJ, Harris KL (1995) An experimental study of the effects of melt composition on plagioclase-melt equilibria at 5 and 10 kbar: implications for the origin of magmatic high—an plagioclase. Contrib Mineral Petrol 118:420–432CrossRefGoogle Scholar
  95. Passmore E, Maclennan J, Fitton JG, Thordarson T (2012) Mush disaggregation in basaltic magma chambers: evidence from the AD 1783–84 Laki eruption. J Petrol 35:2593–2623CrossRefGoogle Scholar
  96. Portnyagin M, Almeev R, Matveev S, Holtz F (2008) Experimental evidence for rapid water exchange between melt inclusions in olivine and host magma. Earth Planet Sci Lett 272:541–552. doi: 10.1016/j.epsl.2008.05.020 CrossRefGoogle Scholar
  97. Presnall DC, Dixon SA, Dixon JR et al (1978) Liquidus phase relations on the join diopside-forsterite-anorthite from 1 atm to 20 kbar: their bearing on the generation and crystallization of basaltic magma. Contrib Mineral Petrol 220:203–220CrossRefGoogle Scholar
  98. Putirka KD (2008) Thermometers and barometers for volcanic systems. Rev Mineral Geochem 69:61–120CrossRefGoogle Scholar
  99. Putirka KD, Johnson M, Kinzler R et al (1996) Thermobarometry of mafic igneous rocks based on clinopyroxene-liquid equilibria, 0–30 kbar. Contrib Mineral Petrol 123:92–108. doi: 10.1007/s004100050145 CrossRefGoogle Scholar
  100. Qin Z, Lu F, Anderson T Jr (1992) Diffusive reequilibration of melt and fluid inclusions. Am Mineral 77:565–576Google Scholar
  101. Rasmussen SO, Andersen KK, Svensson AM et al (2006) A new Greenland ice core chronology for the last glacial termination. J Geophys Res 111:D06102. doi: 10.1029/2005JD006079 Google Scholar
  102. Reverso T, Vandemeulebrouck J, Jouanne F et al (2014) A two-magma chamber model as a source of deformation at Grímsvötn Volcano, Iceland. J Geophys Res Solid Earth 119:4666–4683. doi: 10.1002/2013JB010569 CrossRefGoogle Scholar
  103. Roedder E (1979) Origin and significance of magmatic inclusions. Bull Minéral 102:487–510Google Scholar
  104. Rubin KH, Sinton JM, Maclennan J, Hellebrand E (2009) Magmatic filtering of mantle compositions at mid-ocean-ridge volcanoes. Nat Geosci 2:321–328. doi: 10.1038/ngeo504 CrossRefGoogle Scholar
  105. Saal AE, Hauri EH, Langmuir CH, Perfit MR (2002) Vapour undersaturation in primitive mid-ocean-ridge basalt and the volatile content of Earth’s upper mantle. Nature 419:451–455. doi: 10.1038/nature01073 CrossRefGoogle Scholar
  106. Saunders K, Blundy JD, Dohmen R, Cashman K (2012) Linking petrology and seismology at an active volcano. Science 336:1023–1027. doi: 10.1126/science.1220066 CrossRefGoogle Scholar
  107. Shishkina TA, Botcharnikov RE, Holtz F et al (2010) Solubility of H2O- and CO2-bearing fluids in tholeiitic basalts at pressures up to 500 MPa. Chem Geol 277:115–125. doi: 10.1016/j.chemgeo.2010.07.014 CrossRefGoogle Scholar
  108. Shishkina TA, Botcharnikov RE, Holtz F et al (2014) Compositional and pressure effects on the solubility of H2O and CO2 in mafic melts. Chem Geol 388:112–129. doi: 10.1016/j.chemgeo.2014.09.001 CrossRefGoogle Scholar
  109. Shorttle O, Maclennan J (2011) Compositional trends of Icelandic basalts: implications for short-lengthscale lithological heterogeneity in mantle plumes. Geochem Geophys Geosyst 12:Q11008CrossRefGoogle Scholar
  110. Sigmarsson O, Karlsson H, Larsen G (2000) The 1996 and 1998 subglacial eruptions beneath the Vatnajökull ice sheet in Iceland: contrasting geochemical and geophysical inferences on magma migration. Bull Volcanol 61:468–476. doi: 10.1007/PL00008912 CrossRefGoogle Scholar
  111. Sigmarsson O, Haddadi B, Carn S et al (2013) The sulfur budget of the 2011 Grímsvötn eruption, Iceland. Geophys Res Lett 40:6095–6100CrossRefGoogle Scholar
  112. Sigmundsson F, Hreinsdóttir S, Hooper A et al (2010) Intrusion triggering of the 2010 Eyjafjallajökull explosive eruption. Nature 468:426–430. doi: 10.1038/nature09558 CrossRefGoogle Scholar
  113. Sigurgeirsson MÁ, Leósson MA (1993) Two early Holocene tephra layers in the Sogamyri peat deposit Reykjavik. Náttúrufræðingurinn 62:129–137Google Scholar
  114. Sinton JM, Detrick RS (1992) Mid-ocean ridge magma chambers. J Geophys Res Solid Earth 97:197–216. doi: 10.1029/91JB02508 CrossRefGoogle Scholar
  115. Slater L, McKenzie D, Grönvold K, Shimizu N (2001) Melt generation and movement beneath Theistareykir, NE Iceland. J Petrol 42:321–354CrossRefGoogle Scholar
  116. Sobolev AV, Shimizu N (1993) Ultra-depleted primary melt included in an olivine from the Mid-Atlantic Ridge. Nature 363:151–154CrossRefGoogle Scholar
  117. Sours-Page R, Nielsen RL, Batiza R (2002) Melt inclusions as indicators of parental magma diversity on the northern East Pacific Rise. Chem Geol 183:237–261. doi: 10.1016/S0009-2541(01)00384-9 CrossRefGoogle Scholar
  118. Steele-Macinnis MJ, Esposito R, Bodnar RJ (2011) Thermodynamic model for the effect of post-entrapment crystallization on the H2O–CO2 systematics of vapor-saturated, silicate melt inclusions. J Petrol 52:2461–2482. doi: 10.1093/petrology/egr052 CrossRefGoogle Scholar
  119. Stracke A, Zindler A, Salters VJM et al (2003) Theistareykir revisited. Geochem Geophys Geosyst 4:8507Google Scholar
  120. Streeter R, Dugmore A (2014) Late-Holocene land surface change in a coupled social–ecological system, southern Iceland: a cross-scale tephrochronology approach. Quat Sci Rev 86:99–114. doi: 10.1016/j.quascirev.2013.12.016 CrossRefGoogle Scholar
  121. Tarasewicz J, White RS, Brandsdóttir B, Schoonman CM (2014) Seismogenic magma intrusion before the 2010 eruption of Eyjafjallajokull volcano, Iceland. Geophys J Int 198:906–921. doi: 10.1093/gji/ggu169 CrossRefGoogle Scholar
  122. Thirlwall MF, Gee MAM, Taylor RN, Murton BJ (2004) Mantle components in Iceland and adjacent ridges investigated using double-spike Pb isotope ratios. Geochim Cosmochim Acta 68:361–386CrossRefGoogle Scholar
  123. Thomson A, Maclennan J (2013) The distribution of olivine compositions in Icelandic basalts and picrites. J Petrol 54:745–768. doi: 10.1093/petrology/egs083 CrossRefGoogle Scholar
  124. Thordarson T (2014) The widespread ~10 ka Saksunarvatn tephra is not a product single eruption. Am Geophys Union, Fall Meet, p V24B-04Google Scholar
  125. Thordarson T, Höskuldsson Á (2008) Postglacial volcanism in Iceland. Jökull 58:197–228Google Scholar
  126. Thordarson T, Self S, Óskarsson N, Hulsebosch T (1996) Sulfur, chlorine, and fluorine degassing and atmospheric loading by the 1783–1784 AD Laki (Skaftár Fires) eruption in Iceland. Bull Volcanol 58:205–225. doi: 10.1007/s004450050136 CrossRefGoogle Scholar
  127. Thornalley DJR, McCave IN, Elderfield H (2011) Tephra in deglacial ocean sediments south of Iceland: stratigraphy, geochemistry and oceanic reservoir ages. J Quat Sci 26:190–198. doi: 10.1002/jqs.1442 CrossRefGoogle Scholar
  128. Tormey DR, Grove TL, Bryan WB (1987) Experimental petrology of normal MORB near the Kane Fracture Zone: 20–25°N, mid-Atlantic ridge. Contrib Mineral Petrol 96:121–139CrossRefGoogle Scholar
  129. Tuffen H, Owen J, Denton J (2010) Magma degassing during subglacial eruptions and its use to reconstruct palaeo-ice thicknesses. Earth Sci Rev 99:1–18. doi: 10.1016/j.earscirev.2010.01.001 CrossRefGoogle Scholar
  130. Villiger S, Ulmer P, Müntener O (2007) Equilibrium and fractional crystallization experiments at 0.7 GPa; the Effect of pressure on phase relations and liquid compositions of Tholeiitic magmas. J Petrol 48:159–184. doi: 10.1093/petrology/egl058 CrossRefGoogle Scholar
  131. Viti C, Frezzotti ML (2001) Transmission electron microscopy applied to fluid inclusion investigations. Lithos 55:125–138. doi: 10.1016/S0024-4937(00)00042-6 CrossRefGoogle Scholar
  132. Waagstein R, Jóhansen J (1968) Tre vulkanske askelage fra Færø. Meddelelser fra Dansk Geol Foren 18:257–264Google Scholar
  133. Wallace PJ, Kamenetsky VS, Cervantes P (2015) Melt inclusion CO2 contents, pressures of olivine crystallization, and the problem of shrinkage bubbles. Am Mineral 100:787–794CrossRefGoogle Scholar
  134. Walter MJ (1998) Melting of Garnet Peridotite and the origin of Komatiite and depleted Lithosphere. J Petrol 39:29–60CrossRefGoogle Scholar
  135. Welsch B, Faure F, Famin V et al (2013) Dendritic crystallization: a single process for all the textures of olivine in basalts? J Petrol 54:539–574. doi: 10.1093/petrology/egs077 CrossRefGoogle Scholar
  136. Winpenny B, Maclennan J (2011) A partial record of mixing of mantle melts preserved in Icelandic Phenocrysts. J Petrol 52:1791–1812. doi: 10.1093/petrology/egr031 CrossRefGoogle Scholar
  137. Wood BJ, Blundy JD (1997) A predictive model for rare earth element partitioning between clinopyroxene and anhydrous silicate melt. Contrib Mineral Petrol 129:166–181. doi: 10.1007/s004100050330 CrossRefGoogle Scholar
  138. Workman RK, Hart SR (2005) Major and trace element composition of the depleted MORB mantle (DMM). Earth Planet Sci Lett 231:53–72CrossRefGoogle Scholar
  139. Wright TL, Doherty PC (1970) A linear programming and least squares computer method for petrologic mixing problems. Geol Soc Am Bull 81:1995–2008. doi: 10.1130/0016-7606(1970)81 CrossRefGoogle Scholar
  140. Yang H-J, Kinzler RJ, Grove TL (1996) Experiments and models of anhydrous, basaltic olivine-plagioclase-augite saturated melts from 0.001 to 10 kbar. Contrib Mineral Petrol 124:1–18CrossRefGoogle Scholar

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Authors and Affiliations

  • David A. Neave
    • 1
    • 4
  • John Maclennan
    • 1
  • Thorvaldur Thordarson
    • 2
  • Margaret E. Hartley
    • 1
    • 3
  1. 1.Department of Earth SciencesUniversity of CambridgeCambridgeUK
  2. 2.Faculty of Earth SciencesUniversity of IcelandReykjavíkIceland
  3. 3.School of Earth, Atmospheric and Environmental SciencesUniversity of ManchesterManchesterUK
  4. 4.Institut für MineralogieLeibnitz Universität HannoverHannoverGermany

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