Abstract

We investigate the time scales of magma genesis, melt evolution, crystal growth rates and magma degassing in the Erebus volcano magmatic system using measurements of 238U–230Th–226Ra–210Pb–210Po, 232Th–228Ra–228Th and 235U–231Pa–227Ac. These are the first measurements of 231Pa–227Ac in volcanic samples and represent the first set of data in a volcanic system to examine the entire suite of relevant 238U, 235U and 232Th decay series nuclides. Our sample suite consists of 22 phonolite volcanic bombs, erupted between 1972 and 2005, and five anorthoclase megacrysts separated from bombs erupted in 1984, 1989, 1993, 2004 and 2005. The 238U–230Th, 230Th–226Ra and 235U–231Pa systems are uniform over the 34 years examined. The anorthoclase megacrysts and phonolite glasses show complementary 226Ra/230Th disequilibria with (226Ra/230Th) ∼40 in the anorthoclase and ∼0·75 in the phonolite glass. In all samples, (210Pb/226Ra) is in radioactive equilibrium for both phases. In two phonolite glass samples (227Ac/231Pa) is unity. For the phonolite glasses (228Ra/232Th) is in equilibrium, whereas in the anorthoclase megacrysts it is significantly greater than unity. Instantaneous crystal fractionation, with magma residence times greater than 100 years and less than 10 kyr, can account for the measured 238U–230Th–226Ra–210Pb and 235U–231Pa–227Ac. However, the significant 228Ra/232Th disequilibria in the anorthoclase megacrysts preclude this simple interpretation. To account for this apparent discrepancy we therefore developed an open-system, continuous crystallization model that incorporates both nuclide ingrowth and decay during crystallization. This open-system model successfully reproduces all of the measured 238U and 232Th disequilibria and suggests that the shallow magma reservoir at Erebus is growing. The implication of this modeling is that when the time scale of crystallization is comparable with the half-life of the daughter nuclide of interest (e.g. 226Ra) the simple isochron techniques typically used in most U-series studies can provide erroneous ages. The observation that (210Pb/226Ra) and (227Ac/231Pa) are in radioactive equilibrium suggests that the residence time of the magmas is >100 years. When considering the effect of 222Rn degassing on 210Pb/226Ra, the data indicate that the majority of magma degassing is deep and long before eruption, consistent with melt inclusion data. Additionally, for the 2005 lava bomb, whose eruption date (16 December 2005) is known explicitly, 210Po was not completely degassed from the magma at the time of eruption. Incomplete degassing of 210Po is atypical for subaerially erupted lavas and suggests that the Erebus shallow magma degasses about 1% of its Po per day. The combined 238U and 232Th data further indicate that the pyroclasts ejected by Strombolian eruptions at Erebus have compositions that are close to what would be expected for a near-steady-state system, reflecting inmixing of degassed magmas, crystal fractionation, and aging.

INTRODUCTION

The persistent lava lake at Erebus volcano, Antarctica, provides an unparalleled opportunity to study the time scales and rates of crystallization and magma degassing in an active alkaline magmatic system. Determining these time scales is fundamental to our physical understanding of eruption dynamics and critical for hazard assessment.

The Erebus lava lake is the top of a convecting magma conduit, which is continuously and quiescently degassing. Infrequent Strombolian eruptions often eject phonolite bombs onto the crater rim where they can be sampled. These bombs have provided an almost annual record of the magma composition from the early 1970s to the present. The highly evolved phonolite magma contains ∼30% anorthoclase megacrysts reaching up to 10 cm in length, which makes it a rarity among alkaline systems. A wealth of observational, geophysical, petrological and geochemical data provides a comprehensive and detailed perspective of the genesis and evolution of the Erebus volcanic system and the dynamics of the current lava lake (Giggenbach et al., 1973; Kyle & Cole, 1974; Kyle et al., 1982, 1992; Kyle, 1990a, 1990b; Reagan et al., 1992; Dunbar et al., 1994; Zreda-Gostynska et al., 1997; Esser et al., 2004; Harpel et al., 2004; Sims & Hart, 2006; Oppenheimer & Kyle, 2008, and papers therein; Oppenheimer et al., 2011).

The magmatic evolution of the Erebus lavas has been studied extensively and is convincingly interpreted in terms of simple magma differentiation processes (Kyle et al., 1992). In this contribution, we investigate the time scales of magma genesis, melt evolution, crystal growth rates and magma degassing in the Erebus magmatic system using measurements of 238U–230Th–226Ra–210Pb–210Po, 232Th–228Ra–228Th and 235U–231Pa–227Ac. These different nuclides have starkly contrasting chemistries resulting in significant elemental fractionations during a variety of magmatic processes, and their different half-lives allow the investigation of time scales ranging from less than 1 year to 105 years. This work builds on the dataset presented by Reagan et al. (1992), which investigated a limited range of U-series nuclides and was based on samples from only two eruptions. We note that the present study presents the first measurements of 231Pa–227Ac in volcanic samples and hence is the first time all the relevant nuclides from the 238U, 235U, and 232Th decay series have been measured in the same lavas of a volcanic system.

BACKGROUND ON EREBUS VOLCANO

Erebus (Fig. 1) is an active composite volcano and the largest of four volcanic centers forming Ross Island: Mt. Erebus (3794 m elevation, 2170 km3), Mt. Terror (3262 m, 1700 km3), Mt. Bird (1800 m, 470 km3), and Hut Point Peninsula (100 km3). About 4520 km3 of volcanic material has been erupted on Ross Island over the last ∼4 Myr (Esser et al., 2004). Ross Island is emplaced on thin (17–25 km) rifted continental crust at the southern boundary of the Terror Rift, which is a major graben located within the Victoria Land Basin on the western margin of the West Antarctic rift system (Cooper et al., 1987; Behrendt et al., 1991; Behrendt, 1999; Bannister et al., 2000; Finotello et al., 2011). Late Cenozoic, intraplate, silica-undersaturated, alkaline volcanic rocks erupted on the western margin of the Ross Embayment belong to the McMurdo Volcanic Group (MVG) (Kyle, 1990a). In the southern Ross Sea and McMurdo Sound the MVG is referred to as the Erebus volcanic province (Kyle, 1990b).

Fig. 1

Map of Mt. Erebus showing its location and the location and 40Ar/39Ar ages of its lavas. Map modified from Kelly et al. (2008a). Data taken from Harpel et al. (2004), Esser et al. (2004) and Kelly et al. (2008b).

Fig. 1

Map of Mt. Erebus showing its location and the location and 40Ar/39Ar ages of its lavas. Map modified from Kelly et al. (2008a). Data taken from Harpel et al. (2004), Esser et al. (2004) and Kelly et al. (2008b).

The eruptive history of Erebus volcano has been divided into three distinct phases using high-precision 40Ar/39Ar dates (Esser et al., 2004; Harpel et al., 2004, 2008; Kelly et al., 2008a): (1) a proto-Erebus shield building phase (1·3–1·0 Ma), during which basanites were erupted; (2) a proto-Erebus cone building phase (1·0 Ma–250 ka), dominated by more evolved phonotephrite lavas forming the steep slopes of the volcano; (3) the modern-Erebus cone building phase (250 ka–present), when activity increased and large volumes of anorthoclase-phyric tephriphonolite and phonolite lavas were extruded. Minor trachyte was erupted at about 170 ka during the third phase of activity (Kelly et al., 2008a).

Erebus volcano has hosted a persistent convecting and degassing lava lake of anorthoclase-phyric phonolite magma for nearly 50 years. The discovery and first observations of Mt. Erebus were made in 1841 when James Ross reported it to be in a state of vigorous eruption. Subsequent, but infrequent observations made between 1841 and 1956 indicated that Erebus was actively degassing and presumably had a lava lake as the gas plume was seen to glow during the winter months. A lava lake was present in 1963 based on aerial photographs, and was first observed directly by a scientific party in December 1972 (Giggenbach et al., 1973; Kyle et al., 1982).

All historical eruptive volcanic activity has originated from the phonolite lava lake and adjacent vents. Variations in the style and magnitude of volcanic activity include a period of larger and more frequent Strombolian eruptions in 1984 (Kyle, 1986; Caldwell & Kyle, 1994), a phreatic eruption in 1993, and a period of almost no eruptions from 2002 to 2004. Frequent bomb ejecting Strombolian eruptions occurred through 2006 and 2007, but eruptive activity has since been quiet up to the end of 2011.

Lavas on Ross Island show two major magmatic lineages, the DVDP lineage (Kyle, 1981), named after samples from the Dry Valley Drilling Project, and the Erebus Lineage (EL) (Kyle et al., 1992). The DVDP lineage lavas are older and occur at the volcanic centers surrounding Erebus volcano (Mt. Terror, Mt. Bird, and Hut Point Peninsula). They consist predominantly of basanite with minor microporphyritic kaersutite-bearing intermediate differentiates and small-volume phonolite domes. The Erebus lineage constitutes a surprisingly simple (Kyle et al., 1992) and coherent fractionation trend, defined by a single liquid line of descent from basanite to phonolite with a complete sequence of intermediate (phonotephrite, tephriphonolite) eruptive products. Minor volumes of more iron-rich and less silica-undersaturated benmoreite and trachyte, termed the enriched Fe series (EFS), occur as isolated outcrops on the flanks of Erebus and adjacent islands in Erebus Bay. The EFS lavas follow a different liquid line of descent and the trachytes are interpreted to have undergone both assimilation and fractional crystallization during their evolution.

Erebus lava samples have radiogenic 206Pb/204Pb (Sun & Hansen, 1975; Sims & Hart, 2006; Sims et al., 2008a), unradiogenic 87Sr/86Sr (Kyle et al., 1992; Sims & Hart, 2006; Sims et al., 2008a), and intermediate 143Nd/144Nd and 176Hf/177Hf (Sims & Hart, 2006; Sims et al., 2008a), and lie along a mixing trajectory between the two end-member mantle components DMM and HIMU (Sims et al., 2008a). The Erebus data show a marked distinction between the early-stage basanites and phonotephrites, whose Nd, Hf, Sr, and Pb isotope compositions are variable (particularly Pb), and the later, evolved phonolitic lavas and bombs, whose Nd, Hf, Sr, and Pb isotope compositions are essentially invariant (Sims et al., 2008a). Taken together, the Erebus lineage lavas define a trend of decreasing isotopic variability with increasing extent of differentiation, indicating that magma homogenization has played a fundamental role in establishing the isotopic and compositional uniformity of the more recent phonolites (Sims et al., 2008a).

A unique feature of the Erebus phonolite magma is the presence of abundant large anorthoclase feldspar megacrysts. These anorthoclase feldspar crystals are striking because of their large size, complex internal zoning and abundance of melt inclusions as well as inclusions of other crystalline phases such as pyroxene, apatite, magnetite, and pyrrhotite (Dunbar et al., 1994; Kelly et al., 2008b). The crystals can be up to 10 cm in length (Kyle, 1977), and can contain as much as 30 vol. % of melt inclusions trapped during crystal growth (Fig. 2). The melt inclusions are over 1 mm in length, and such large inclusions are typically irregular (Dunbar et al., 1994), but the crystals also contain a population of negative crystal shape inclusions that are 10–40 µm in diameter. Populations of melt inclusions are trapped along what appear to be growth zones in the crystals, producing a banded appearance in back-scattered electron (BSE) images. The anorthoclase crystals are complexly compositionally zoned at a number of scales and also show evidence of periods of resorption during crystal growth (Dunbar et al., 1994; Kelly et al., 2008b; Sumner, 2007). The compositional range of anorthoclase is An10–23 Ab62–68 Or11–27 (Kelly et al., 2008b). The coarsest scale of compositional zoning is defined by high- and low-Ca zones in the crystal, as illustrated by the low-Ca core and high-Ca rim of the crystal shown in Fig. 2. This zoning is attributed to either convective processes (Kelly et al., 2008b) or boundary layer effects during crystallization (Sumner, 2007). The observed skeletal crystal growth can be related to rapid cooling, large degrees of undercooling, or lack of available nucleation sites (Lofgren, 1980). The same effects that cause skeletal growth also tend to cause rapid crystal growth, suggesting that the Erebus anorthoclase crystals grew rapidly. Crystal size distributions, internal crystal structures and geochemical compositions (Dunbar et al., 1994) suggest that the crystals have low nucleation rates but underwent rapid growth in two or more stages. The initial growth stage produced spongy inclusion-rich cores, which were then overgrown by finely laminated rims. The composition of melt inclusions in the anorthoclase is very similar to that of the glass that forms the matrix of the samples, suggesting that the anorthoclase crystals formed in a fully evolved phonolitic melt. Assuming that the growth rate of crystals was similar to that of normal plagioclase, the age of an average Erebus anorthoclase crystal would be 100–300 years (Dunbar et al., 1994). This age span would be less if the crystals grew faster, but the presence of dissolution within the crystals suggests that a simple size distribution–age calculation may underestimate the true crystal age.

Fig. 2

Backscattered electron image (BSE) and Ca element map of a single anorthoclase crystal. The darkest areas in the BSE image are the anorthoclase host crystal. The abundant irregular areas, slightly brighter than the anorthoclase, are melt inclusions. Other included phases are pyroxene (light grey phase on Ca map), apatite (next brightest in BSE, highlighted in red on Ca map) and magnetite (brightest in BSE, dark on Ca map). The size of the apatite crystals has been slightly increased to improve visibility.

Fig. 2

Backscattered electron image (BSE) and Ca element map of a single anorthoclase crystal. The darkest areas in the BSE image are the anorthoclase host crystal. The abundant irregular areas, slightly brighter than the anorthoclase, are melt inclusions. Other included phases are pyroxene (light grey phase on Ca map), apatite (next brightest in BSE, highlighted in red on Ca map) and magnetite (brightest in BSE, dark on Ca map). The size of the apatite crystals has been slightly increased to improve visibility.

Melt inclusions found in the anorthoclase crystals contain lower concentrations of H2O, CO2 and S (Oppenheimer et al., 2011) than the assumed parental basanites. Oppenheimer et al. (2011) suggested that CO2 contents in the anorthoclase melt inclusions are consistent with entrapment at pressures of up to 3 kbar. This contrasts with earlier interpretations (Dunbar et al., 1994) based on H2O contents of melt inclusions that implied that crystals growth took place either in a degassed magma in the upper (<400 m) part of the magmatic system or in degassed magma that had circulated to depth by convection in the magma conduit. Seaman et al. (2006) reported variable H2O abundances in anorthoclase melt inclusions and showed ∼50 µm thick zones of elevated H2O contents on the boundary of crystals, and suggested that water diffused into crystals from melt inclusions.

In addition to containing large anorthoclase feldspar crystals, the Erebus magma also contains titanomagnetite, clinopyroxene, apatite, pyrrhotite and olivine (Kyle, 1977; Kelly et al., 2008b). The textural relationships between these phases suggest that the pyroxene, titanomagnetite and olivine crystallized first. The H2O contents of melt inclusions in pyroxene (n = 4, H2O = 0·1 ± 0·05 wt %) led Dunbar et al. (1994) to suggest that they crystallized high in the magmatic system.

There have been two prior studies measuring U-series disequilibria in Erebus lavas. Reagan et al. (1992) used 238U–230Th–226Ra disequilibria to constrain the magma residence times for anorthoclase–glass separates from bombs erupted in 1984 and 1988. Those workers showed that the anorthoclase crystals were strongly enriched in 226Ra over 230Th, whereas glass separates had 226Ra deficits. Based on the assumption that during crystal growth DRa = DBa for anorthoclase and glass, they constructed (226Ra)/Ba vs (230Th)/Ba isochrons for the two samples and obtained two-point isochron ages of 2520 years for the 1984 sample and 2225 years for the 1988 sample. Using the activity of 228Th as a proxy for 228Ra (which is reasonable given the short half-life of 228Th), Reagan et al. (1992) also measured significant (228Th/232Th) disequilibrium in the anorthoclase [∼2·16, whereas the glasses were in equilibrium with (228Th)/(232Th) of 1·0]. Because of the short half-life of 228Ra (t1/2 = 5·77 years) this result would seem to suggest that the anorthoclase crystals grew recently and rapidly. To reconcile this conundrum, Reagan et al. (1992) proposed that the 228Th disequilibria resulted from the presence of young rims with 228Ra excesses, whereas the majority of the crystals were older and had 228Th–228Ra–232Th equilibria.

The second U-series study was by Sims & Hart (2006), who measured 238U–230Th disequilibria in four historical Erebus bombs as part of a global study evaluating U-series disequilibria and Nd, Sr, and Pb isotope systematics in oceanic basalts. Their results showed that the (230Th/232Th) and (238U/232Th) values of Erebus bombs are intermediate relative to those of other ocean island and mid-ocean ridge basalts, and form the end-member HIMU mantle component on plots of Pb isotopes versus (230Th/232Th) and (238U/232Th). Although Erebus does not constitute the HIMU end-member mantle component, there are no samples young enough for 238U–230Th disequilibria studies from Mangaia–Tubuai (the HIMU end-member); hence, for the 238U–230Th isotope system, Erebus, by necessity, represents the best end-member approximation of the HIMU source. Finally, it is important to note that in both studies (230Th/238U) is >1, indicating that garnet was a residual phase during the mantle partial melting that led to the parental basanites, which in turn fractionated to form the phonolite.

SAMPLE COLLECTION AND PREPARATION

This study examines phonolite lava bombs erupted during Strombolian eruptions over a 34 year period, from 1972 to 2005. The bombs were collected on the summit crater rim or from the floor of the Main Crater.

The exact date and time of eruptions are known for some bomb samples (Table 1), whereas the young ages for other samples were determined based on their ‘fresh appearance’ at the time of collection. Recently erupted bombs have a distinctive metallic to iridescent vitreous luster that is quickly lost (∼1–2 weeks) upon exposure to the acidic gases emitted from the lava lake and surrounding fumaroles. Therefore, for samples without eruption dates, their fresh appearance suggested they had been erupted within less than 1 or 2 months prior to collection.

Table 1

230Th/232Th, 234U/238U and U, Th, 226Ra, 231Pa and Ba concentrations measured by mass spectrometry

Sample Eruption date [Th] (µg g–1§ [U] (µg g–1§ (238U/232Th) § (230Th/232Th)* § 
Glass          
ER 25724G (A1) 12/26/72 30·87 0·18 8·95 0·03 0·879 0·008 0·984 0·006 
ER 2E2G (A1) Dec-74 31·03 0·20 8·97 0·04 0·877 0·009 0·987 0·006 
ER 77016G (A1) Nov-77 31·82 0·19 9·23 0·04 0·880 0·009 0·985 0·006 
ER 79302G (A1) 12/26/79 30·91 0·17 8·96 0·04 0·879 0·009 0·985 0·006 
ER 82416G (A1) Dec-82 31·73 0·17 9·16 0·03 0·875 0·008 0·986 0·006 
ER 83220G (A1) Dec-83 30·64 0·18 9·01 0·03 0·892 0·008 0·985 0·006 
ER 83220G (B1)        0·987 0·009 
ER 83220G (C1)        0·987 0·005 
ER 84501G (A1) Dec-84 29·14 0·19 8·50 0·04 0·885 0·009 0·979 0·006 
ER 84501G (C1)          
ER 84505G (A1) Dec-84 29·35 0·15 8·49 0·04 0·877 0·008 0·976 0·006 
ER 84505G (B1)        0·973 0·009 
ER 84505G (B2)        0·979 0·009 
ER 84505G (B3)        0·975 0·010 
ER 84505G (B4)        0·979 0·009 
ER 84505G (B5)        0·976 0·010 
ER 85010G (A1) Dec-85 29·46 0·18 8·54 0·04 0·879 0·009 0·987 0·007 
ER 85010G (B1)        0·983 0·010 
ER 85010G (C1)        0·985 0·006 
ER 86022-1G (A1) 12/22/86 30·46 0·17 8·84 0·04 0·880 0·009 0·986 0·007 
ER 86022-1G (B1)        0·986 0·010 
ER 88104-G (A1) Dec-88 29·37 0·17 8·56 0·03 0·884 0·008 0·987 0·006 
ER 89001G (A1) Dec-89 29·28 0·17 8·49 0·04 0·880 0·009 0·989 0·006 
ER 89001G (B1)        0·989 0·009 
ER 89001G (D1)  26·28 0·14 7·77 0·03 0·897 0·008 0·976 0·003 
ER 89001G (D2)  26·39 0·14 7·78 0·03 0·894 0·008 0·976 0·004 
ER 89001G (D3)  27·23 0·14 8·09 0·03 0·901 0·008 0·984 0·003 
ER 89001G (D4)  26·47 0·14 7·81 0·06 0·895 0·011 0·977 0·003 
ER 91101G (A1) Jan-91 29·04 0·19 8·46 0·04 0·884 0·010 0·985 0·006 
ER 91101G (A2)  30·70 0·16 8·81 0·03 0·871 0·008 0·989 0·006 
Er 92KSG (A1) 12/10/92 31·16 0·16 9·00 0·04 0·876 0·008 0·988 0·006 
Er 92KSG (C1)        0·987 0·006 
ER 93 JCG (A1) Jan-93 29·53 0·17 8·51 0·04 0·874 0·009 0·987 0·006 
ER 96 G (A1) Dec-96 30·46 0·17 8·83 0·04 0·880 0·009 0·988 0·006 
Er 97G (A1) Dec-97 30·78 0·16 8·817 0·03 0·869 0·007 0·985 0·006 
Er 99G (A1) Dec-99 26·58 0·18 7·657 0·04 0·874 0·010 0·980 0·006 
Er 99G (A2)  26·96 0·18 7·698 0·03 0·866 0·009 0·983 0·006 
Er 2000 ‘Y2K’G (A1) Dec-00 29·01 0·18 8·298 0·03 0·868 0·008 0·984 0·006 
ER 2001G (A1) Dec-01 29·52 0·15 8·526 0·03 0·876 0·008 0·983 0·006 
ER Jan2004G (A1) Jan-04 30·46 0·17 8·768 0·04 0·873 0·008 0·983 0·006 
ER Jan2004G (C1)        0·983 0·005 
ER Dec2005G (A1) 12/16/05 29·07 0·19 8·40 0·04 0·877 0·010 0·981 0·006 
ER Dec2005G (C1)        0·980 0·005 
ER Dec2005G (D1)  22·20 0·12 6·419 0·021 0·877 0·008 0·970 0·006 
Anorthoclase megacrysts          
84 xyll (A1) 1984 0·28 0·003 0·07 0·001 0·765 0·015 0·977 0·006 
88 xyll (A1) 1988 0·22 0·002 0·06 0·001 0·771 0·015 0·980 0·006 
93 xyll (A1) 1993 0·19 0·002 0·05 0·000 0·778 0·016 0·981 0·006 
04 xyll (A1) 2004 0·23 0·002 0·06 0·001 0·782 0·016 0·981 0·006 
05 xyll (A1) 2005 0·27 0·003 0·07 0·001 0·780 0·016 0·982 0·006 
Quality assurance standards||          
TML (A) (n = 22; #dis. = 7)  30·50 0·71 10·79 0·20 1·073 0·045 1·071 0·006 
ATHO (A) (n = 11; #dis. = 3)  7·44 0·03 2·26 0·05 0·920 0·022 1·017 0·005 
BCR-2 (A) (n = 6; #dis. = 3)  5·89 0·03 1·71 0·02 0·879 0·015 0·874 0·005 
KL-31-KWWS-92 (A) (n = 1)    0·292 0·004     
HK-04-KWWS-92 (A) (n = 1)    0·665 0·010     
Kilauea_Skylight (D) (n = 1)    0·260 0·003     
Sample Eruption date [Th] (µg g–1§ [U] (µg g–1§ (238U/232Th) § (230Th/232Th)* § 
Glass          
ER 25724G (A1) 12/26/72 30·87 0·18 8·95 0·03 0·879 0·008 0·984 0·006 
ER 2E2G (A1) Dec-74 31·03 0·20 8·97 0·04 0·877 0·009 0·987 0·006 
ER 77016G (A1) Nov-77 31·82 0·19 9·23 0·04 0·880 0·009 0·985 0·006 
ER 79302G (A1) 12/26/79 30·91 0·17 8·96 0·04 0·879 0·009 0·985 0·006 
ER 82416G (A1) Dec-82 31·73 0·17 9·16 0·03 0·875 0·008 0·986 0·006 
ER 83220G (A1) Dec-83 30·64 0·18 9·01 0·03 0·892 0·008 0·985 0·006 
ER 83220G (B1)        0·987 0·009 
ER 83220G (C1)        0·987 0·005 
ER 84501G (A1) Dec-84 29·14 0·19 8·50 0·04 0·885 0·009 0·979 0·006 
ER 84501G (C1)          
ER 84505G (A1) Dec-84 29·35 0·15 8·49 0·04 0·877 0·008 0·976 0·006 
ER 84505G (B1)        0·973 0·009 
ER 84505G (B2)        0·979 0·009 
ER 84505G (B3)        0·975 0·010 
ER 84505G (B4)        0·979 0·009 
ER 84505G (B5)        0·976 0·010 
ER 85010G (A1) Dec-85 29·46 0·18 8·54 0·04 0·879 0·009 0·987 0·007 
ER 85010G (B1)        0·983 0·010 
ER 85010G (C1)        0·985 0·006 
ER 86022-1G (A1) 12/22/86 30·46 0·17 8·84 0·04 0·880 0·009 0·986 0·007 
ER 86022-1G (B1)        0·986 0·010 
ER 88104-G (A1) Dec-88 29·37 0·17 8·56 0·03 0·884 0·008 0·987 0·006 
ER 89001G (A1) Dec-89 29·28 0·17 8·49 0·04 0·880 0·009 0·989 0·006 
ER 89001G (B1)        0·989 0·009 
ER 89001G (D1)  26·28 0·14 7·77 0·03 0·897 0·008 0·976 0·003 
ER 89001G (D2)  26·39 0·14 7·78 0·03 0·894 0·008 0·976 0·004 
ER 89001G (D3)  27·23 0·14 8·09 0·03 0·901 0·008 0·984 0·003 
ER 89001G (D4)  26·47 0·14 7·81 0·06 0·895 0·011 0·977 0·003 
ER 91101G (A1) Jan-91 29·04 0·19 8·46 0·04 0·884 0·010 0·985 0·006 
ER 91101G (A2)  30·70 0·16 8·81 0·03 0·871 0·008 0·989 0·006 
Er 92KSG (A1) 12/10/92 31·16 0·16 9·00 0·04 0·876 0·008 0·988 0·006 
Er 92KSG (C1)        0·987 0·006 
ER 93 JCG (A1) Jan-93 29·53 0·17 8·51 0·04 0·874 0·009 0·987 0·006 
ER 96 G (A1) Dec-96 30·46 0·17 8·83 0·04 0·880 0·009 0·988 0·006 
Er 97G (A1) Dec-97 30·78 0·16 8·817 0·03 0·869 0·007 0·985 0·006 
Er 99G (A1) Dec-99 26·58 0·18 7·657 0·04 0·874 0·010 0·980 0·006 
Er 99G (A2)  26·96 0·18 7·698 0·03 0·866 0·009 0·983 0·006 
Er 2000 ‘Y2K’G (A1) Dec-00 29·01 0·18 8·298 0·03 0·868 0·008 0·984 0·006 
ER 2001G (A1) Dec-01 29·52 0·15 8·526 0·03 0·876 0·008 0·983 0·006 
ER Jan2004G (A1) Jan-04 30·46 0·17 8·768 0·04 0·873 0·008 0·983 0·006 
ER Jan2004G (C1)        0·983 0·005 
ER Dec2005G (A1) 12/16/05 29·07 0·19 8·40 0·04 0·877 0·010 0·981 0·006 
ER Dec2005G (C1)        0·980 0·005 
ER Dec2005G (D1)  22·20 0·12 6·419 0·021 0·877 0·008 0·970 0·006 
Anorthoclase megacrysts          
84 xyll (A1) 1984 0·28 0·003 0·07 0·001 0·765 0·015 0·977 0·006 
88 xyll (A1) 1988 0·22 0·002 0·06 0·001 0·771 0·015 0·980 0·006 
93 xyll (A1) 1993 0·19 0·002 0·05 0·000 0·778 0·016 0·981 0·006 
04 xyll (A1) 2004 0·23 0·002 0·06 0·001 0·782 0·016 0·981 0·006 
05 xyll (A1) 2005 0·27 0·003 0·07 0·001 0·780 0·016 0·982 0·006 
Quality assurance standards||          
TML (A) (n = 22; #dis. = 7)  30·50 0·71 10·79 0·20 1·073 0·045 1·071 0·006 
ATHO (A) (n = 11; #dis. = 3)  7·44 0·03 2·26 0·05 0·920 0·022 1·017 0·005 
BCR-2 (A) (n = 6; #dis. = 3)  5·89 0·03 1·71 0·02 0·879 0·015 0·874 0·005 
KL-31-KWWS-92 (A) (n = 1)    0·292 0·004     
HK-04-KWWS-92 (A) (n = 1)    0·665 0·010     
Kilauea_Skylight (D) (n = 1)    0·260 0·003     
Sample Eruption date [226Ra] (fg g–1§ (226Ra/230Th) § [231Pa] (fg g–1§ (231Pa/235U) § 
Glass          
ER 25724G (A1) 12/26/72 2469 37 0·73 0·01     
ER 2E2G (A1) Dec-74         
ER 77016G (A1) Nov-77 2560 38 0·73 0·01     
ER 79302G (A1) 12/26/79 2480 37 0·73 0·01     
ER 82416G (A1) Dec-82         
ER 83220G (A1) Dec-83         
ER 83220G (B1)          
ER 83220G (C1)          
ER 84501G (A1) Dec-84         
ER 84501G (C1)          
ER 84505G (A1) Dec-84 2388 36 0·75 0·01     
ER 84505G (B1)          
ER 84505G (B2)          
ER 84505G (B3)          
ER 84505G (B4)          
ER 84505G (B5)          
ER 85010G (A1) Dec-85         
ER 85010G (B1)          
ER 85010G (C1)          
ER 86022-1G (A1) 12/22/86         
ER 86022-1G (B1)          
ER 88104-G (A1) Dec-88 2413 36 0·75 0·01     
ER 89001G (A1) Dec-89 2306 35 0·72 0·01 3369 40 1·220 0·02 
ER 89001G (B1)          
ER 89001G (D1)      2735 29 1·083 0·02 
ER 89001G (D2)      3202 34 1·266 0·02 
ER 89001G (D3)      3120 31 1·186 0·01 
ER 89001G (D4)      2493 61 0·981 0·03 
ER 91101G (A1) Jan-91 2329 35 0·73 0·01 3481 42 1·265 0·02 
ER 91101G (A2)  2369 36 0·70 0·01     
Er 92KSG (A1) 12/10/92 2300 34 0·67 0·01 3585 43 1·225 0·02 
Er 92KSG (C1)          
ER 93 JCG (A1) Jan-93 2324 35 0·72 0·01 3376 41 1·220 0·02 
ER 96G (A1) Dec-96 2332 35 0·70 0·01     
Er 97G (A1) Dec-97 2298 34 0·68 0·01     
Er 99G (A1) Dec-99 2283 34 0·79 0·01 3300 40 1·325 0·02 
Er 99G (A2)  2300 34 0·78 0·01     
Er 2000 ‘Y2K’G (A1) Dec-00 2306 35 0·73 0·01 3283 39 1·216 0·02 
ER 2001G (A1) Dec-01         
ER Jan2004G (A1) Jan-04 2306 35 0·69 0·01 3449 41 1·209 0·02 
ER Jan2004G (C1)          
ER Dec2005G (A1) 12/16/05 2296 34 0·724 0·01     
ER Dec2005G (C1)          
ER Dec2005G (D1)      2623 28 1·256 0·02 
Anorthoclase megacrysts          
84 xyll (A1) 1984 917 14 30·66 0·46     
88 xyll (A1) 1988 950 14 39·93 0·60     
93 xyll (A1) 1993 939 14 44·70 0·67     
04 xyll (A1) 2004 1000 15 39·77 0·60     
05 xyll (A1) 2005 1105 17 37·25 0·56     
Quality assurance standards||          
TML (A) (n = 22; #Dis. = 7)  3675 110 1·01 0·05     
ATHO (A) (n = 11; #Dis. = 3)  847 56 1·01 0·03     
BCR-2 (A) (n = 6; #Dis. = 3)  565 47 0·99 0·02     
KL-31-KWWS-92 (A) (n = 1)      103·00 1·55 1·08 0·03 
HK-04-KWWS-92 (A) (n = 1)      389·20 4·67 1·79 0·05 
Kilauea_Skylight (D) (n = 1)      92·28 0·59 1·09 0·02 
Sample Eruption date [226Ra] (fg g–1§ (226Ra/230Th) § [231Pa] (fg g–1§ (231Pa/235U) § 
Glass          
ER 25724G (A1) 12/26/72 2469 37 0·73 0·01     
ER 2E2G (A1) Dec-74         
ER 77016G (A1) Nov-77 2560 38 0·73 0·01     
ER 79302G (A1) 12/26/79 2480 37 0·73 0·01     
ER 82416G (A1) Dec-82         
ER 83220G (A1) Dec-83         
ER 83220G (B1)          
ER 83220G (C1)          
ER 84501G (A1) Dec-84         
ER 84501G (C1)          
ER 84505G (A1) Dec-84 2388 36 0·75 0·01     
ER 84505G (B1)          
ER 84505G (B2)          
ER 84505G (B3)          
ER 84505G (B4)          
ER 84505G (B5)          
ER 85010G (A1) Dec-85         
ER 85010G (B1)          
ER 85010G (C1)          
ER 86022-1G (A1) 12/22/86         
ER 86022-1G (B1)          
ER 88104-G (A1) Dec-88 2413 36 0·75 0·01     
ER 89001G (A1) Dec-89 2306 35 0·72 0·01 3369 40 1·220 0·02 
ER 89001G (B1)          
ER 89001G (D1)      2735 29 1·083 0·02 
ER 89001G (D2)      3202 34 1·266 0·02 
ER 89001G (D3)      3120 31 1·186 0·01 
ER 89001G (D4)      2493 61 0·981 0·03 
ER 91101G (A1) Jan-91 2329 35 0·73 0·01 3481 42 1·265 0·02 
ER 91101G (A2)  2369 36 0·70 0·01     
Er 92KSG (A1) 12/10/92 2300 34 0·67 0·01 3585 43 1·225 0·02 
Er 92KSG (C1)          
ER 93 JCG (A1) Jan-93 2324 35 0·72 0·01 3376 41 1·220 0·02 
ER 96G (A1) Dec-96 2332 35 0·70 0·01     
Er 97G (A1) Dec-97 2298 34 0·68 0·01     
Er 99G (A1) Dec-99 2283 34 0·79 0·01 3300 40 1·325 0·02 
Er 99G (A2)  2300 34 0·78 0·01     
Er 2000 ‘Y2K’G (A1) Dec-00 2306 35 0·73 0·01 3283 39 1·216 0·02 
ER 2001G (A1) Dec-01         
ER Jan2004G (A1) Jan-04 2306 35 0·69 0·01 3449 41 1·209 0·02 
ER Jan2004G (C1)          
ER Dec2005G (A1) 12/16/05 2296 34 0·724 0·01     
ER Dec2005G (C1)          
ER Dec2005G (D1)      2623 28 1·256 0·02 
Anorthoclase megacrysts          
84 xyll (A1) 1984 917 14 30·66 0·46     
88 xyll (A1) 1988 950 14 39·93 0·60     
93 xyll (A1) 1993 939 14 44·70 0·67     
04 xyll (A1) 2004 1000 15 39·77 0·60     
05 xyll (A1) 2005 1105 17 37·25 0·56     
Quality assurance standards||          
TML (A) (n = 22; #Dis. = 7)  3675 110 1·01 0·05     
ATHO (A) (n = 11; #Dis. = 3)  847 56 1·01 0·03     
BCR-2 (A) (n = 6; #Dis. = 3)  565 47 0·99 0·02     
KL-31-KWWS-92 (A) (n = 1)      103·00 1·55 1·08 0·03 
HK-04-KWWS-92 (A) (n = 1)      389·20 4·67 1·79 0·05 
Kilauea_Skylight (D) (n = 1)      92·28 0·59 1·09 0·02 
Sample Eruption date (234U/238U)* § Ba# (µg g–1Ba-ID (µg g–1§ 
Glass       
ER 25724G (A1) 12/26/72 0·996 0·004 435   
ER 2E2G (A1) Dec-74 1·002 0·004    
ER 77016G (A1) Nov-77 1·001 0·003 461   
ER 79302G (A1) 12/26/79 1·003 0·004 454   
ER 82416G (A1) Dec-82 0·998 0·004 445   
ER 83220G (A1) Dec-83 1·001 0·004 463   
ER 83220G (B1)       
ER 83220G (C1)  1·001 0·003    
ER 84501G (A1) Dec-84 1·000 0·003 455   
ER 84501G (C1)  0·997 0·003    
ER 84505G (A1) Dec-84 1·001 0·002 465 473·5 7·1 
ER 84505G (B1)       
ER 84505G (B2)       
ER 84505G (B3)       
ER 84505G (B4)       
ER 84505G (B5)       
ER 85010G (A1) Dec-85 1·002 0·003 404   
ER 85010G (B1)       
ER 85010G (C1)  1·001 0·002    
ER 86022-1G (A1) 12/22/86 1·001 0·003 438   
ER 86022-1G (B1)       
ER 88104-G (A1) Dec-88 1·003 0·003  512·4 7·7 
ER 89001G (A1) Dec-89 1·002 0·003 463   
ER 89001G (B1)       
ER 89001G (D1)  1·002 0·002    
ER 89001G (D2)  1·005 0·002    
ER 89001G (D3)  1·002 0·002    
ER 89001G (D4)  1·001 0·002    
ER 91101G (A1) Jan-91 0·999 0·003 449   
ER 91101G (A2)       
Er 92KSG (A1) 12/10/92 1·002 0·002 463   
Er 92KSG (C1)  1·003 0·003    
ER 93 JCG (A1) Jan-93 1·002 0·002 454 502·38 7·5 
ER 96G (A1) Dec-96 1·000 0·002 436   
Er 97G (A1) Dec-97   442   
Er 99G (A1) Dec-99 1·001 0·002 429   
Er 99G (A2)       
Er 2000 ‘Y2K’G (A1) Dec-00 0·999 0·002 405   
ER 2001G (A1) Dec-01   475   
ER Jan2004G (A1) Jan-04 1·002 0·002 404 486·58 7·3 
ER Jan2004G (C1)  1·001 0·002    
ER Dec2005G (A1) 12/16/05 1·002 0·002  495·640 7·4 
ER Dec2005G (C1)  1·001 0·002    
ER Dec2005G (D1)  1·002 0·002    
Anorthoclase megacrysts       
84 xyll (A1) 1984 1·001 0·003  2428·6 36 
88 xyll (A1) 1988 1·002 0·003  2327·1 35 
93 xyll (A1) 1993 0·998 0·003  2568·3 39 
04 xyll (A1) 2004 1·003 0·003  2295·7 34 
05 xyll (A1) 2005 1·001 0·003  2389·3 36 
Quality assurance standards||       
TML (A) (n = 22; #dis. = 7)  1·0007 0·003    
ATHO (A) (n = 11; #dis. = 3)  0·9997 0·003    
BCR-2 (A) (n = 6; #dis. = 3)  1·0031 0·004    
KL-31-KWWS-92 (A) (n = 1)       
HK-04-KWWS-92 (A) (n = 1)       
Kilauea_Skylight (D) (n = 1)       
Sample Eruption date (234U/238U)* § Ba# (µg g–1Ba-ID (µg g–1§ 
Glass       
ER 25724G (A1) 12/26/72 0·996 0·004 435   
ER 2E2G (A1) Dec-74 1·002 0·004    
ER 77016G (A1) Nov-77 1·001 0·003 461   
ER 79302G (A1) 12/26/79 1·003 0·004 454   
ER 82416G (A1) Dec-82 0·998 0·004 445   
ER 83220G (A1) Dec-83 1·001 0·004 463   
ER 83220G (B1)       
ER 83220G (C1)  1·001 0·003    
ER 84501G (A1) Dec-84 1·000 0·003 455   
ER 84501G (C1)  0·997 0·003    
ER 84505G (A1) Dec-84 1·001 0·002 465 473·5 7·1 
ER 84505G (B1)       
ER 84505G (B2)       
ER 84505G (B3)       
ER 84505G (B4)       
ER 84505G (B5)       
ER 85010G (A1) Dec-85 1·002 0·003 404   
ER 85010G (B1)       
ER 85010G (C1)  1·001 0·002    
ER 86022-1G (A1) 12/22/86 1·001 0·003 438   
ER 86022-1G (B1)       
ER 88104-G (A1) Dec-88 1·003 0·003  512·4 7·7 
ER 89001G (A1) Dec-89 1·002 0·003 463   
ER 89001G (B1)       
ER 89001G (D1)  1·002 0·002    
ER 89001G (D2)  1·005 0·002    
ER 89001G (D3)  1·002 0·002    
ER 89001G (D4)  1·001 0·002    
ER 91101G (A1) Jan-91 0·999 0·003 449   
ER 91101G (A2)       
Er 92KSG (A1) 12/10/92 1·002 0·002 463   
Er 92KSG (C1)  1·003 0·003    
ER 93 JCG (A1) Jan-93 1·002 0·002 454 502·38 7·5 
ER 96G (A1) Dec-96 1·000 0·002 436   
Er 97G (A1) Dec-97   442   
Er 99G (A1) Dec-99 1·001 0·002 429   
Er 99G (A2)       
Er 2000 ‘Y2K’G (A1) Dec-00 0·999 0·002 405   
ER 2001G (A1) Dec-01   475   
ER Jan2004G (A1) Jan-04 1·002 0·002 404 486·58 7·3 
ER Jan2004G (C1)  1·001 0·002    
ER Dec2005G (A1) 12/16/05 1·002 0·002  495·640 7·4 
ER Dec2005G (C1)  1·001 0·002    
ER Dec2005G (D1)  1·002 0·002    
Anorthoclase megacrysts       
84 xyll (A1) 1984 1·001 0·003  2428·6 36 
88 xyll (A1) 1988 1·002 0·003  2327·1 35 
93 xyll (A1) 1993 0·998 0·003  2568·3 39 
04 xyll (A1) 2004 1·003 0·003  2295·7 34 
05 xyll (A1) 2005 1·001 0·003  2389·3 36 
Quality assurance standards||       
TML (A) (n = 22; #dis. = 7)  1·0007 0·003    
ATHO (A) (n = 11; #dis. = 3)  0·9997 0·003    
BCR-2 (A) (n = 6; #dis. = 3)  1·0031 0·004    
KL-31-KWWS-92 (A) (n = 1)       
HK-04-KWWS-92 (A) (n = 1)       
Kilauea_Skylight (D) (n = 1)       

*U and Th isotopic compositions measured by: (A) MC-ICP-MS at WHOI using the ThermoFisher NEPTUNE (Ball et al., 2008; Sims et al., 2008b, 2008c); (B) SIMS at WHOI using the Cameca IMS 1270 (Layne & Sims, 2000); (C) MC-ICP-MS at Wyoming High Precision Isotope Laboratory using the ThermoFisher NEPTUNE Plus; (D) MC-ICP-MS at Bristol Isotope Group using the ThermoFisher NEPTUNE (Hoffman et al., 2007).

For concentration measurements: (A) [U], [Th], [226Ra], [Ba] and [231Pa] measured by ID-ICP-MS at WHOI using the ThermoFisher ELEMENT2 (Choi et al., 2001; Pichat et al., 2004; Sims et al., 2008b, 2008c); (D) [U], [Th] and [231Pa], [Ba] measured by ID-ICP-MS at Bristol using the ThermoFisher NEPTUNE (Hoffman et al., 2007; Prytulak et al., 2007, 2008).

() denotes activity. Activity ratios were calculated using λ230 = 9·1577 × 10–6 yr–1, λ234 = 2·8262 × 10–6 yr–1, λ238 = 1·551 × 10–10 yr–1; λ232 = 4·948 × 10–11 yr–1; λ226 = 4·331 × 10–4 yr–1; λ231 = 2·1158 × 10–5 yr–1; λ235 = 9·8485 × 10–10 yr–1 (Le Roux & Glendenin, 1963; Robert et al., 1969; Jaffey et al., 1971; Holden, 1990; Cheng et al., 2000; Tuli, 2000).

§Errors (2σ) are calculated using standard error propagation methods and include uncertainties in: (1) the decay constants, λ230 (0·3%), λ232 (0·5%), λ238 (o·2%), λ238 (0·07%), λ232 (0·5%), λ226 (0·4%), λ231 (0·4%) or λ235 (0·07%) (Le Roux & Glendenin, 1963; Robert et al., 1969; Jaffey et al., 1971; Holden, 1990; Cheng et al., 2000; Tuli, 2000); (2) the time-averaged uncertainty in 233U (0·7%) 229Th (1%), 228Ra (1·3%) spikes used for isotope dilution; (3) the instrument parameters, including the uncertainty in determining the tailing of 232Th on 230Th (∼0·1–0·2%); (4) the weighing errors (<0·001%); (5) measurement precision for the samples and bracketing standards (0·03–0·4%).

Inter-laboratory (i.e. WHOI, Bristol, University of Wyoming) and inter-technique (MC-ICP-MS vs SIMS) replicates come from separate powder dissolutions. Intra-laboratory and intra-technique replicates come from different chemical processing of a single dissolution.

||TML, ATHO and BCR2 were measured for quality assurance. Values reported here represent the range of values measured for (238U/232Th), (230Th/232Th) (230Th/238U), [226Ra] and (226Ra/230Th) over the period of this study from ∼1999 to now. The table lists both the number of analyses and dissolutions incorporated into these numbers. Replicates for synthetic rock and standards run over the same time interval and their comparison with results from other laboratories has been reported by Layne & Sims (2000), Ball et al. (2007) and Sims et al. (2008b, 2008c). For the WHOI (231Pa/235U) measurements two Hawaiian samples, KL-31-KWWS-92 and HK-04-KWWS-92, were measured for quality assurance during the same time as these analyses; these measurements agree, within error, with the reported values for these same samples given by Sims et al. (1999) [see also Sims et al. (2008c)]. For Bristol’s (231Pa/235U) measurements, Hawaiian sample Kilauea Skylight (collected from a skylight of a Pu’u O’o flow by Dave Sherrod on 16 May 1996 was measured and gives a (231Pa/235U) very similar to other historical Kilauea Pu’u O’o samples measured by Sims et al. (1999) and KL-31-KWWS-92 measured here.

#Ba concentrations from standard ICP-MS come from Kelly et al. (2008b).

Our sample suite consists of 22 bombs and five anorthoclase megacrysts separated from bombs erupted in 1984, 1989, 1993, 2004 and 2005. All samples are purified glass or anorthoclase separates. Lava bombs were disaggregated and phenocrysts and glasses were separated by handpicking to obtain ∼20–30 g of each phase and subsequently purified using magnetic separation. The anorthoclase megacrysts were abraded by hand grinding to remove exterior glass. The handpicked separates (both anorthoclase and phonolite glass) were then ground to <0·04 mm and magnetically separated using a Frantz magnetic separator. Magnetic separation was repeated until visually clean separates of anorthoclase (melt inclusion free) and purified glass were obtained.

All glass samples have previously been characterized for their major and trace element abundances (Kelly et al., 2008b) and long-lived radiogenic Nd, Sr, Hf and Pb isotope compositions (Sims & Hart, 2006; Sims et al., 2008a).

ANALYTICAL METHODS

238U, 232Th and 235U decay series

The anorthoclase separates and glasses were dissolved in a series of digestions using concentrated HF and HNO3, followed by HNO3 + H3BO3 and HClO4 to break down all fluorides. Care was taken to ensure complete digestion and elimination of all fluorides.

238U, 232Th, 235U, 227Ac, 226Ra, 210Pb, 210Po, and 228Ra concentrations and 234U/238U and 230Th/232Th isotopic ratios were measured by a combination of multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS) (234U/238U and 230Th/232Th), secondary ionization mass spectrometry (SIMS) (230Th/232Th), isotope dilution (ID) mass spectrometry (232Th, 238U, 231 Pa, 226Ra and Ba concentrations), gamma spectrometry (214Pb and 214Bi as proxies for 226Ra, and 228Ac and 208Tl as proxies for 228Ra), and alpha spectrometry (210Po as a proxy for 210Pb for older samples and as a direct measure of 210Po, and 228Th as a proxy for 228Ra and 227Ac). Details of the methods for these measurements have been given in published analytical and interpretative papers (Choi et al., 2001; Layne & Sims, 2000; Pichat et al., 2004; Reagan et al., 2006, 2008; Ball et al., 2008; Hoffman et al., 2007; Prytulak et al., 2008; Prytulak & Elliott, 2009; Sims et al. 2008b, 2008c). Details of the 227Ac method are given in Appendix A and were published by Dulaiova et al. (2012). The 228Th method is also described in Appendix A. Tables 1–3 tabulate the data and provide information on relevant details such as the decay constants used and the method of error propagation.

Table 2

Short-lived gamma counting data 226Ra proxies (214Pb and 214Bi), 210Pb, and 228Ra proxies (228Ac and 208Tl)

  ID-ICP-MS ID-ICP-MS       
Sample Eruption date 226Ra 232Th (214Pb) 1σ RSD (214Pb/226Ra) (214Bi) 1σ RSD (214Bi/226Ra) 
  (d.p.m. g–1(d.p.m. g–1(d.p.m. g–1(%)  (d.p.m. g–1(%)  
Glass          
84505G 1984 5·24 7·17 5·26 0·01 1·00 5·13 0·02 0·98 
91101G 1991 5·15 7·09 4·99 0·01 0·97 4·92 0·02 0·96 
Er 92KS-G 1992 5·05 7·61 5·05 0·02 1·00 5·13 0·03 1·02 
ER 93 JC 1993 5·10 7·21 5·06 0·01 0·99 4·94 0·02 0·97 
Er 96G 1996 5·12 7·44 5·06 0·01 0·99 5·12 0·02 1·00 
Er 97G 1997 5·04 7·51 5·20 0·01 1·03 5·11 0·02 1·01 
Er 99G 1999 5·03 6·54 5·02 0·01 1·00 5·05 0·02 1·00 
Er 2000 ‘Y2K’G 2000 5·06 7·08 5·11 0·01 1·01 5·06 0·02 1·00 
JAN2004G 2004 5·06 7·44 5·17 0·01 1·02 5·08 0·02 0·98 
Dec-05 2005 5·04 7·10 5·16 0·01 1·02 5·09 0·02 0·98 
Anorthoclase megacrysts 
84 xyll 1984 2·01 0·067 1·94 0·03 0·96 b.d.   
93 xyll 1993 2·06 0·047 2·12 0·03 1·03 b.d.   
04 xyll 2004 2·19 0·056 2·22 0·03 1·01 2·081 0·04 0·98 
05 xyll 2005 2·42 0·066 2·60 0·02 1·07 2·470 0·02 0·98 
  ID-ICP-MS ID-ICP-MS       
Sample Eruption date 226Ra 232Th (214Pb) 1σ RSD (214Pb/226Ra) (214Bi) 1σ RSD (214Bi/226Ra) 
  (d.p.m. g–1(d.p.m. g–1(d.p.m. g–1(%)  (d.p.m. g–1(%)  
Glass          
84505G 1984 5·24 7·17 5·26 0·01 1·00 5·13 0·02 0·98 
91101G 1991 5·15 7·09 4·99 0·01 0·97 4·92 0·02 0·96 
Er 92KS-G 1992 5·05 7·61 5·05 0·02 1·00 5·13 0·03 1·02 
ER 93 JC 1993 5·10 7·21 5·06 0·01 0·99 4·94 0·02 0·97 
Er 96G 1996 5·12 7·44 5·06 0·01 0·99 5·12 0·02 1·00 
Er 97G 1997 5·04 7·51 5·20 0·01 1·03 5·11 0·02 1·01 
Er 99G 1999 5·03 6·54 5·02 0·01 1·00 5·05 0·02 1·00 
Er 2000 ‘Y2K’G 2000 5·06 7·08 5·11 0·01 1·01 5·06 0·02 1·00 
JAN2004G 2004 5·06 7·44 5·17 0·01 1·02 5·08 0·02 0·98 
Dec-05 2005 5·04 7·10 5·16 0·01 1·02 5·09 0·02 0·98 
Anorthoclase megacrysts 
84 xyll 1984 2·01 0·067 1·94 0·03 0·96 b.d.   
93 xyll 1993 2·06 0·047 2·12 0·03 1·03 b.d.   
04 xyll 2004 2·19 0·056 2·22 0·03 1·01 2·081 0·04 0·98 
05 xyll 2005 2·42 0·066 2·60 0·02 1·07 2·470 0·02 0·98 
Sample Eruption date (210Pb) 1σ RSD (210Pb/226Ra) (228Ac) 1σ RSD (228Ac/232Th) (208Tl) 1σ RSD (208Tl/232Th) 
  (d.p.m. g–1(%)  (d.p.m. g–1(%)  (d.p.m. g–1(%)  
Glass           
84505G 1984 4·84 0·06 0·92 7·20 0·02 1·00 6·98 0·02 0·97 
91101G 1991 5·26 0·04 1·02 6·99 0·02 0·99 6·89 0·02 0·97 
Er 92KS-G 1992 4·96 0·08 0·98 7·50 0·04 0·99 7·46 0·03 0·98 
ER 93 JC 1993 4·94 0·05 0·97 7·15 0·02 0·99 7·00 0·02 0·97 
Er 96G 1996 5·30 0·05 1·04 7·18 0·02 0·97 7·16 0·02 0·96 
Er 97G 1997  0·04  7·30 0·02 0·97 7·19 0·02 0·96 
Er 99G 1999 4·86 0·05 0·97 6·40 0·02 0·98 6·45 0·02 0·99 
Er 2000 ‘Y2K’G 2000 4·82 0·05 0·95 6·86 0·02 0·97 6·90 0·02 0·97 
JAN2004G 2004 4·91 0·05 0·97 7·14 0·02 0·96 7·12 0·02 0·96 
Dec-05 2005 4·80 0·05 0·95 6·88 0·02 0·97 6·87 0·02 0·97 
Anorthoclase megacrysts 
84 xyll 1984 2·05 0·12 1·02 b.d.   b.d.   
93 xyll 1993 2·27 0·11 1·10 b.d.   b.d.   
04 xyll 2004 2·35 0·12 1·07 b.d.   b.d.   
05 xyll 2005 2·65 0·09 1·10 b.d.   b.d.   
Sample Eruption date (210Pb) 1σ RSD (210Pb/226Ra) (228Ac) 1σ RSD (228Ac/232Th) (208Tl) 1σ RSD (208Tl/232Th) 
  (d.p.m. g–1(%)  (d.p.m. g–1(%)  (d.p.m. g–1(%)  
Glass           
84505G 1984 4·84 0·06 0·92 7·20 0·02 1·00 6·98 0·02 0·97 
91101G 1991 5·26 0·04 1·02 6·99 0·02 0·99 6·89 0·02 0·97 
Er 92KS-G 1992 4·96 0·08 0·98 7·50 0·04 0·99 7·46 0·03 0·98 
ER 93 JC 1993 4·94 0·05 0·97 7·15 0·02 0·99 7·00 0·02 0·97 
Er 96G 1996 5·30 0·05 1·04 7·18 0·02 0·97 7·16 0·02 0·96 
Er 97G 1997  0·04  7·30 0·02 0·97 7·19 0·02 0·96 
Er 99G 1999 4·86 0·05 0·97 6·40 0·02 0·98 6·45 0·02 0·99 
Er 2000 ‘Y2K’G 2000 4·82 0·05 0·95 6·86 0·02 0·97 6·90 0·02 0·97 
JAN2004G 2004 4·91 0·05 0·97 7·14 0·02 0·96 7·12 0·02 0·96 
Dec-05 2005 4·80 0·05 0·95 6·88 0·02 0·97 6·87 0·02 0·97 
Anorthoclase megacrysts 
84 xyll 1984 2·05 0·12 1·02 b.d.   b.d.   
93 xyll 1993 2·27 0·11 1·10 b.d.   b.d.   
04 xyll 2004 2·35 0·12 1·07 b.d.   b.d.   
05 xyll 2005 2·65 0·09 1·10 b.d.   b.d.   

ATHO is used to calibrate efficiencies for 46·52 keV (210Pb), 338·4 keV (228Ac), 351·99 keV (214Pb), 583·14 keV (208Tl), and 609·32 keV (214Bi).

Branching ratios used are 0·0405 for 46·52 keV (210Pb), 0·1136 for 338·4 keV (228Ac), 0·372 for 351·99 keV (214Pb), 0·8423 for 583·14 keV (208Tl), and 0·4628 for 609·32 keV (214Bi).

Comparing samples sealed in epoxy with samples that were not checked for Rn degassing. The two approaches give equivalent concentrations.

(228Ra) and (232Th) activities calculated from mass spectrometry data (Table 1).

b.d., below detection limit.

Table 3

Short-lived alpha counting data for 210Po, 227Ac and 228Th

Sample Date of analysis (226Ra) ID-ICP-MS* (d.p.m. g−1(210Po) (d.p.m. g−12σ (210Po/226Ra)0 Date of analysis (230Th/232Th) 2σ (228Th/232Th)° 2σ 
ER 84505 (An) 8/4/06 2·01 1·99 0·08 0·989      
ER 91101 (Gl) 6/27/06 5·15 5·26 0·2 1·021      
ER 92KS (Gl) 6/27/06 5·05 5·27 0·2 1·044      
ER 93 JC (An) 8/4/06 2·06 2·09 0·08 1·014      
ER 96 (Gl) 6/27/06 5·12 5·27 0·22 1·030      
ER 99 (Gl) 6/16/06 5·03 5·26 0·21 1·046      
ER 2000 (Gl) 6/16/06 5·06 5·12 0·19 1·012      
ER Jan2004 (Gl) 2/7/07 5·06 5·06 0·16 1·000      
 3/23/07 5·06 5·34 0·2 1·055      
ER Jan2004 (An) 5/26/06 2·19 2·22 0·09 1·012      
ER Dec2005 (Gl) 5/26/06  3·65 0·14       
 2/7/07  4·82 0·16       
 4/4/08 5·04 5·3 0·16 1·052      
ER Dec2005 (An) 5/26/06 2·42 2·42 0·1 0·998 4/27/2008 0·91 0·13 1·41 0·2 
Quality assurance           
BCR-2           
BCR-2 (Bristol)‡           
SAV-B6           
(1901 Samoa Flow)           
Sample Date of analysis (226Ra) ID-ICP-MS* (d.p.m. g−1(210Po) (d.p.m. g−12σ (210Po/226Ra)0 Date of analysis (230Th/232Th) 2σ (228Th/232Th)° 2σ 
ER 84505 (An) 8/4/06 2·01 1·99 0·08 0·989      
ER 91101 (Gl) 6/27/06 5·15 5·26 0·2 1·021      
ER 92KS (Gl) 6/27/06 5·05 5·27 0·2 1·044      
ER 93 JC (An) 8/4/06 2·06 2·09 0·08 1·014      
ER 96 (Gl) 6/27/06 5·12 5·27 0·22 1·030      
ER 99 (Gl) 6/16/06 5·03 5·26 0·21 1·046      
ER 2000 (Gl) 6/16/06 5·06 5·12 0·19 1·012      
ER Jan2004 (Gl) 2/7/07 5·06 5·06 0·16 1·000      
 3/23/07 5·06 5·34 0·2 1·055      
ER Jan2004 (An) 5/26/06 2·19 2·22 0·09 1·012      
ER Dec2005 (Gl) 5/26/06  3·65 0·14       
 2/7/07  4·82 0·16       
 4/4/08 5·04 5·3 0·16 1·052      
ER Dec2005 (An) 5/26/06 2·42 2·42 0·1 0·998 4/27/2008 0·91 0·13 1·41 0·2 
Quality assurance           
BCR-2           
BCR-2 (Bristol)‡           
SAV-B6           
(1901 Samoa Flow)           
Sample (231Pa) ID-ICPMS* (d.p.m. g−12σ Date of analysis (227Ac) (d.p.m. g−12σ (227Ac/231Pa) 2σ 
ER 84505 (An)        
ER 91101 (Gl)        
ER 92KS (Gl)        
ER 93 JC (An)        
ER 96 (Gl)        
ER 99 (Gl) 0·346 0·01 6/10/00 0·3498 0·02 1·01 0·06 
ER 2000 (Gl)        
ER Jan2004 (Gl)        
ER Jan2004 (An) 0·362 0·01 9/9/07 0·36815 0·02 1·02 0·06 
ER Dec2005 (Gl)        
ER Dec2005 (An)        
Quality assurance        
BCR-2 0·057 0·002  0·058 0·003 1·02 0·06 
BCR-2 (Bristol)‡ 0·058 0·0006  0·058 0·003 1·03 0·05 
SAV-B6 (1901 Samoa Flow) 0·041 0·001  0·0406 0·002 1·00 0·06 
Sample (231Pa) ID-ICPMS* (d.p.m. g−12σ Date of analysis (227Ac) (d.p.m. g−12σ (227Ac/231Pa) 2σ 
ER 84505 (An)        
ER 91101 (Gl)        
ER 92KS (Gl)        
ER 93 JC (An)        
ER 96 (Gl)        
ER 99 (Gl) 0·346 0·01 6/10/00 0·3498 0·02 1·01 0·06 
ER 2000 (Gl)        
ER Jan2004 (Gl)        
ER Jan2004 (An) 0·362 0·01 9/9/07 0·36815 0·02 1·02 0·06 
ER Dec2005 (Gl)        
ER Dec2005 (An)        
Quality assurance        
BCR-2 0·057 0·002  0·058 0·003 1·02 0·06 
BCR-2 (Bristol)‡ 0·058 0·0006  0·058 0·003 1·03 0·05 
SAV-B6 (1901 Samoa Flow) 0·041 0·001  0·0406 0·002 1·00 0·06 

*(226Ra) and (231Pa) are measured by ID-ICP-MS (Table 1).

210Po alpha counts are decay corrected to the date of 210Pb–210Po separation.

231Pa activity comes from average BCR-2 concentration of 550·3 ± 7·2 (2σ standard deviation; n = 15) from Prytulak et al. (2008).

In situ analysis of anorthoclase megacrysts, apatite and trapped melt inclusions

We measured Ba and Th concentrations in two 1984 and one 1988 Erebus anorthoclase megacrysts and their melt inclusions by SIMS (IMS-3f, Cameca) at the Northeast National Ion Microprobe Facility at the Woods Hole Oceanographic Institution (Table 4). The megacrysts were cut perpendicularly to their elongated axis. The polished sections of the samples were Au-coated. The analyses were made using a negatively charged oxygen ion (O) beam with a net voltage of 12·5 kV. The beam diameter was 10 μm. Molecular ion interferences were suppressed using an energy filtering technique (Shimizu & Hart, 1982a, 1982b) by offsetting the secondary accelerating voltage by 100 V. The 138Ba and 232Th counts were normalized to 30Si. We used four reference glasses to calibrate the ion microprobe (NIST610, NIST612, KL2G and ML3G) to obtain the empirical relationships between the concentrations of Ba and Th and the element/30Si ratio (e.g. Shimizu & Hart, 1982a, 1982b). The Ba and Th concentrations used for these glass standards are from the GeoREM database (Jochum et al., 2005). The calculated concentrations were within 7% and 13% of the certified values, respectively. Major element compositions were measured on the same crystal sections using a Cameca SX100 electron microprobe at New Mexico Tech. For the feldspar analyses, we used an accelerating voltage of 15 kV and a probe current of 20 nA, with the beam broadened to 10 μm to avoid Na loss. Count times for all elements were 20 s on peak and 10 s on background, with the exception of Sr and Ba, for which peak count times of 60 s were used, and 30 s background count times. We used the average SiO2 concentration measured by electron microprobe in the melt inclusions or in the anorthoclase to calculate the Th and Ba concentrations. The precision on SiO2 determinations, based on replicate analysis of standard reference materials, averaged ±0·35 wt %.

Table 4

Average in situ Ba and Th concentrations in anorthoclase megacrysts and their melt inclusions measured by SIMS (IMS-3f Cameca) at WHOI

 Ba
 
Th
 
 (µg g–11σ (µg g–11σ 
ER84-01     
Melt inclusion 490 90 31·7 4·4 
Anorthoclase 2185 291 0·10 0·04 
ER84-04     
Melt inclusion 549 31 28·0 3·3 
Anorthoclase 2279 47 0·03 0·02 
ER88-01     
Melt inclusion 410 18 24·66 2·36 
Anorthoclase 1840 54 0·13 0·07 
 Ba
 
Th
 
 (µg g–11σ (µg g–11σ 
ER84-01     
Melt inclusion 490 90 31·7 4·4 
Anorthoclase 2185 291 0·10 0·04 
ER84-04     
Melt inclusion 549 31 28·0 3·3 
Anorthoclase 2279 47 0·03 0·02 
ER88-01     
Melt inclusion 410 18 24·66 2·36 
Anorthoclase 1840 54 0·13 0·07 

These data are averages from multiple spot measurements (see Table A1).

Quantitative analyses of the Th (and U and Pb) contents of apatite inclusions in the anorthoclases were also carried out using the Cameca SX100 electron microprobe at New Mexico Tech. A standard accelerating voltage of 15 kV was used, but because of the low abundances of these elements in apatite, high probe currents (200 nA) and long peak count times (300–400 s) were used. Uranium and Pb abundances in apatite inclusions in the anorthoclases were below instrumental detection limits.

RESULTS

238U, 232Th, 235U, 227Ac, 226Ra, 210Pb, 210Po, and 228Ra concentrations and (234U/238U), (230Th/232Th), (230Th/238U), (226Ra/230Th), (235U/231Pa), (231Pa/227Ac), and (228Ra/232Th) are reported in Tables 1–3 and shown in Figs 3–8.

Fig. 3

(a) (230Th/232Th) vs (238U/232Th) isochron plot showing the Erebus samples compared with the global database for mid-ocean ridge basalt (MORB) and ocean island basalt (OIB). (b) Erebus data enlarged. The shift of anorthoclase megacrysts to lower U/Th exceeds what would be predicted from experimental uncertainties and is interpreted to be due to apatite inclusions (see text for discussion).

Fig. 3

(a) (230Th/232Th) vs (238U/232Th) isochron plot showing the Erebus samples compared with the global database for mid-ocean ridge basalt (MORB) and ocean island basalt (OIB). (b) Erebus data enlarged. The shift of anorthoclase megacrysts to lower U/Th exceeds what would be predicted from experimental uncertainties and is interpreted to be due to apatite inclusions (see text for discussion).

Fig. 4

Time-series data for (a) (230Th/232Th), (b) (238U/232Th), and (c) (230Th/238U) over the 34 year historical record from 1972 to 2005. Time is since 2012.

Fig. 4

Time-series data for (a) (230Th/232Th), (b) (238U/232Th), and (c) (230Th/238U) over the 34 year historical record from 1972 to 2005. Time is since 2012.

Fig. 5

Time-series data for (226Ra/230Th) over the 34 year historical record from 1972 to 2005. Time is since 2012.

Fig. 5

Time-series data for (226Ra/230Th) over the 34 year historical record from 1972 to 2005. Time is since 2012.

Fig. 6

(210Pb/226Ra) vs (226Ra/230Th) for Erebus samples compared with a global compilation. Data from Gauthier & Condomines (1999), Turner et al. (2004), Rubin et al. (2005), Reagan et al. (2006, 2008) and Berlo & Turner (2010).

Fig. 6

(210Pb/226Ra) vs (226Ra/230Th) for Erebus samples compared with a global compilation. Data from Gauthier & Condomines (1999), Turner et al. (2004), Rubin et al. (2005), Reagan et al. (2006, 2008) and Berlo & Turner (2010).

Fig. 7

(210Po) measurements for the sample erupted on 16 December 2005 showing ingrowth curve from 210Pb.

Fig. 7

(210Po) measurements for the sample erupted on 16 December 2005 showing ingrowth curve from 210Pb.

Fig. 8

Time-series data for 228Ra over the 34 year historical record from 1972 to 2005. Time is since 2012. *(228Ra/232Th) was determined using 228Th/232Th measured by alpha spectroscopy as a proxy. The half-life of 228Th is 1·9 years, so for this 2005 sample (measured in 2008) the (228Ra/232Th) inferred from (228Th/232Th) is a minimum value as the system had not yet attained full equilibrium. Reagan et al. (1992) data are shown for comparison.

Fig. 8

Time-series data for 228Ra over the 34 year historical record from 1972 to 2005. Time is since 2012. *(228Ra/232Th) was determined using 228Th/232Th measured by alpha spectroscopy as a proxy. The half-life of 228Th is 1·9 years, so for this 2005 sample (measured in 2008) the (228Ra/232Th) inferred from (228Th/232Th) is a minimum value as the system had not yet attained full equilibrium. Reagan et al. (1992) data are shown for comparison.

For quality assurance, replicate measurements of (238U/232Th), (230Th/232Th), (230Th/238U), (234U/238U), (226Ra/230Th), (235U/231Pa), and (231Pa/227Ac) were conducted for the rock standards ATHO and TML (Table 1); these analyses are consistent, within analytical uncertainties, with expectations of equilibrium, and also with previously reported MC-ICP-MS, thermal ionization mass spectrometry (TIMS) and SIMS measurements (see Sims et al., 2008a).

238U decay series

238U–234U–230Th

U and Th concentrations and (234U/238U) and (230Th/232Th) (234U t1/2 = 245 250 years and 230Th t1/2 = 75 690 years) were measured in the 22 historical phonolite bombs and five anorthoclase megacrysts (Table 1 and Figs 3 and 4). Replicate analyses for the mass spectrometric techniques (SIMS versus MC-ICP-MS) and the different laboratories (WHOI, University of Wyoming, Bristol) are reported in Table 1. All of the samples analyzed in this study have (234U/238U) activity ratios of unity ( ± 4 per mil) (Table 1). (230Th/232Th) and (238U/232Th) for all of the historical Erebus bombs are uniform within analytical uncertainties and (230Th/238U) is greater than one, indicating that daughter 230Th has been enriched relative to its parent 238U [hereafter (230Th/238U) >1 is referred to as 230Th excess]. For the anorthoclase megacrysts, 230Th/232Th is identical to the glass separates from the same bombs, but (238U/232Th) is lower and (230Th/238U) is higher (Table 1; Figs 3 and 4).

For the 230Th/232Th measurements, inter- and intra-technique reproducibility (SIMS and MC-ICP-MS) and inter-laboratory reproducibility (University of Wyoming, Bristol and WHOI) are within analytical error, which is less than 1%. For (238U/232Th), inter-laboratory (Bristol versus WHOI) reproducibility is also within analytical error, again less than 1%; however, the Bristol measurements give significantly lower U and Th concentrations than the WHOI results. We interpret the lower U and Th concentrations in the Bristol Isotope Group measurements as variable amounts of admixed anorthoclase megacrysts, as the Bristol laboratory received whole-rock phonolite bomb samples. Anorthoclase has extremely low U and Th concentrations (Table 1). Simple mixing calculations between average phonolite glass and anorthoclase (Table 5) show that incorporation of 20% crystals in the phonolite glass aliquots (which is less than their modal abundances of ∼30%) will lower the concentrations of U and Th by about 20%. This is comparable with the Bristol measurements and changes the (238U/232Th) of the samples by only 0·02%. We further note that U and Th concentrations also have been measured in purified glass separates by standard ICP-MS (Kelly et al., 2008b) and are similar, within analytical uncertainties, to the WHOI isotope dilution measurements.

Table 5

Mean, standard deviation, median, kurtosis and skewness of the distribution, and range of Erebus lava bombs, with and without the 1984 samples

 [Th] (µg g−1[U] (µg g−1Th/U (238U/232Th) [230Th/232Th]atm (ppm) (230Th/232Th) 
Phonolite glasses       
1984 bombs included       
n 22 22 22 22 22 22 
Average 30·031 8·687 3·457 0·878 5·322 0·985 
2σ SD 2·271 0·685 0·041 0·011 0·028 0·006 
Median 30·162 8·701 3·459 0·877 5·322 0·985 
Maximum 31·824 9·228 3·496 0·892 5·342 0·989 
Minimum 26·771 7·677 3·403 0·868 5·284 0·978 
Skewness −0·843 −0·992 −0·458 0·509 −1·049 −0·940 
Kurtosis 1·840 2·262 1·437 1·521 1·504 0·483 
Without 1984 bombs       
n 20 20 20 20 20 20 
Average 30·110 8·707 3·459 0·877 5·324 0·985 
2σ SD 2·326 0·708 0·041 0·011 0·023 0·004 
Median 30·459 8·792 3·460 0·877 5·323 0·985 
Maximum 31·824 9·228 3·496 0·892 5·342 0·989 
Minimum 26·771 7·677 3·403 0·868 5·301 0·981 
Skewness −1·064 −1·168 −0·574 0·630 −0·418 −0·418 
Kurtosis 2·254 2·482 1·941 2·056 −0·417 −0·417 
Anorthoclase megacrysts       
n 
Average 0·238 0·061 3·915 0·775 5·296 0·980 
2σ SD 0·035 0·009 0·036 0·007 0·011 0·002 
 [Th] (µg g−1[U] (µg g−1Th/U (238U/232Th) [230Th/232Th]atm (ppm) (230Th/232Th) 
Phonolite glasses       
1984 bombs included       
n 22 22 22 22 22 22 
Average 30·031 8·687 3·457 0·878 5·322 0·985 
2σ SD 2·271 0·685 0·041 0·011 0·028 0·006 
Median 30·162 8·701 3·459 0·877 5·322 0·985 
Maximum 31·824 9·228 3·496 0·892 5·342 0·989 
Minimum 26·771 7·677 3·403 0·868 5·284 0·978 
Skewness −0·843 −0·992 −0·458 0·509 −1·049 −0·940 
Kurtosis 1·840 2·262 1·437 1·521 1·504 0·483 
Without 1984 bombs       
n 20 20 20 20 20 20 
Average 30·110 8·707 3·459 0·877 5·324 0·985 
2σ SD 2·326 0·708 0·041 0·011 0·023 0·004 
Median 30·459 8·792 3·460 0·877 5·323 0·985 
Maximum 31·824 9·228 3·496 0·892 5·342 0·989 
Minimum 26·771 7·677 3·403 0·868 5·301 0·981 
Skewness −1·064 −1·168 −0·574 0·630 −0·418 −0·418 
Kurtosis 2·254 2·482 1·941 2·056 −0·417 −0·417 
Anorthoclase megacrysts       
n 
Average 0·238 0·061 3·915 0·775 5·296 0·980 
2σ SD 0·035 0·009 0·036 0·007 0·011 0·002 
 (230Th/238U) [226Ra] (fg g−1(226Ra/230Th) Ba ICP-MS* (µg g−1Ba-ID (µg g−1
Phonolite glasses      
1984 bombs included      
n 22 15 15 20 
Average 1·122 2361·2 0·72 446·95 494·10 
2σ SD 0·014 166·3 0·06 48·34 29·78 
Median 1·123 2324·1 0·73 451·5 495·64 
Maximum 1·134 2559·9 0·78 499 512·4 
Minimum 1·106 2291·6 0·67 404 473·49 
Skewness −0·737 1·3 0·13 −0·234 −0·31 
Kurtosis 0·866 0·8 0·57 0·344 −0·36 
Without 1984 bombs      
n 20 14 14 18 
Average 1·123 2359·3 0·72 445·5 499·25 
2σ SD 0·012 172·0 0·06 50·12 21·79 
Median 1·124 2315·2 0·72 447 499·01 
Maximum 1·134 2559·9 0·78 499 512·4 
Minimum 1·106 2291·6 0·67 404 486·58 
Skewness −0·780 1·4 0·30 −0·069 0·115 
Kurtosis 2·169 0·8 0·88 0·219 −0·280 
Anorthoclase megacrysts      
n  
Average 1·265 982·2 38·46  2401·8 
2σ SD 0·009 75·0 5·13  106·6 
 (230Th/238U) [226Ra] (fg g−1(226Ra/230Th) Ba ICP-MS* (µg g−1Ba-ID (µg g−1
Phonolite glasses      
1984 bombs included      
n 22 15 15 20 
Average 1·122 2361·2 0·72 446·95 494·10 
2σ SD 0·014 166·3 0·06 48·34 29·78 
Median 1·123 2324·1 0·73 451·5 495·64 
Maximum 1·134 2559·9 0·78 499 512·4 
Minimum 1·106 2291·6 0·67 404 473·49 
Skewness −0·737 1·3 0·13 −0·234 −0·31 
Kurtosis 0·866 0·8 0·57 0·344 −0·36 
Without 1984 bombs      
n 20 14 14 18 
Average 1·123 2359·3 0·72 445·5 499·25 
2σ SD 0·012 172·0 0·06 50·12 21·79 
Median 1·124 2315·2 0·72 447 499·01 
Maximum 1·134 2559·9 0·78 499 512·4 
Minimum 1·106 2291·6 0·67 404 486·58 
Skewness −0·780 1·4 0·30 −0·069 0·115 
Kurtosis 2·169 0·8 0·88 0·219 −0·280 
Anorthoclase megacrysts      
n  
Average 1·265 982·2 38·46  2401·8 
2σ SD 0·009 75·0 5·13  106·6 

For the 1984 bomb, (238U/232Th) and (230Th/232Th) are similar, within analytical uncertainty, to the values reported by Reagan et al. (1992). However, for the 1988 bomb the (230Th/232Th) reported by Reagan et al. (1992) is higher than our values by 2% for the glass and 5% for the anorthoclase, and the (238U/232Th) is higher than our values by 4% for the glass and 10% for the anorthoclase. Because our data give horizontal isochrons, are remarkably uniform for the 34 year sample suite, and have been measured in multiplicity by newer mass spectrometric methods and by different laboratories, we consider these newer data to be the most reliable.

230Th–226Ra

226Ra (t1/2 = 1600 years) concentrations were measured in glass separates from 15 historical phonolite bombs and five anorthoclase megacryst separates (Table 1, Fig. 5). For the bombs (226Ra/230Th) is <1 and relatively uniform for the glass separates (0·722 ± 0·057, 2σ), whereas for the anorthoclase separates (226Ra/230Th) ranges from 30·6 to 40·7. These values are consistent with the (226Ra/230Th) measured for the glass and anorthoclase in 1984 and 1988 bombs by Reagan et al. (1992).

226Ra was measured by both isotope dilution mass spectrometry using MC-ICP-MS and gamma counting of the short-lived 226Ra daughters 214Pb and 214Bi (Sims et al., 2008b). This latter approach assumes that (226Ra) is in radioactive equilibrium with (214Pb) and (214Bi), which is reasonable given the extremely short half-lives of these daughter nuclides (35 ms to 3·83 days). Within their respective analytical uncertainties, both daughter activities give concurrent results amongst themselves and with the mass spectrometric 226Ra measurements (Tables 1 and 2).

226Ra–210Pb

210Pb (t1/2 = 22·6 years) was measured by gamma spectroscopy and by 210Po ingrowth using alpha spectrometry (Figs 6 and 7; Tables 2 and 3). Although the 210Po alpha counting data are more precise and accurate, both methods give equivalent results within their respective analytical uncertainties. All samples measured in this study are in equilibrium with regard to 210Pb–226Ra (Fig. 6).

210Pb–210Po

For the 2005 sample (erupted on 16 December 2005) 210Po (t1/2 = 138·4 days) was measured by alpha spectrometry in three aliquots over a 1 year period (Table 3; Fig. 7). In most magmas, complete polonium volatilization occurs during and prior to eruption, creating an initial 210Po deficit relative to the 210Pb grandparent (e.g. Gill et al., 1985; Reagan et al., 2006). In this sample, (210Po/210Pb)I = 0·28, and thus Po was not entirely degassed upon eruption.

232Th decay series

232Th–228Ra–228Th

As a proxy for 228Ra (t1/2 = 5·77 years), we measured 228Ac and 208Tl by gamma spectroscopy in 10 glass samples (Fig. 8; Table 2). The 228Ac proxy assumes that (228Ra) is in radioactive equilibrium with its daughter (228Ac), which is a reasonable assumption given the extremely short half-life of 228Ac (t1/2 = 6·13 h). The 208Tl proxy assumes that (228Ra) is in radioactive equilibrium all the way down the chain to its distant granddaughter (208Tl), which is also reasonable as the half-lives of these different nuclei range from less than 1 µs to 3·6 days. All samples have (228Ac/232Th) and (208Tl/232Th) in radioactive equilibrium to within 1σ counting uncertainties (Fig. 8), and thus, by inference, they are in equilibrium with respect to (228Ra/232Th).

The count rates were below detection limits for 228Ac and 208Tl gamma spectrometry of anorthoclase (Table 2). For the 2005 anorthoclase separate, (228Ra/232Th) was determined using 228Th/232Th measured by alpha spectroscopy as a proxy. The half-life of 228Th is 1·9 years, so for this 3-year-old sample the (228Ra/232Th) inferred from (228Th/232Th) is a minimum estimate as the system had not attained full equilibrium.

235U decay series

235U–231Pa

231Pa (t1/2 = 32 757 years) was measured on eight samples (1989, 1991, 1992, 1993, 1999, 2000, 2004, 2005). The WHOI (231Pa/235U) data range from 1·22 to 1·33, with an average of 1·24 ± 0·04 (1σ), whereas the Bristol laboratory found lower U and Pa concentrations and (231Pa/235U) ranging from 0·98 to 1·26. The higher value is similar to that measured at WHOI. This variability in the Bristol data and its deviation from the WHOI data is attributed to the Bristol samples being whole-rock aliquots containing significant, but variable amounts of anorthoclase and other crystalline phases. We note that because the Pa does not vary systematically with U and Th concentrations, it is likely that anorthoclase is not the only phase affecting the variability in (231Pa/235U) measurements.

231Pa–227Ac

227Ac (t1/2 = 21·77 years) was measured on two phonolite glass samples (1999 and 2004). In both samples (227Ac/231Pa) is unity within analytical uncertainty. The absence of 231Pa–227Ac disequilibria suggests that either no magmatic process, such as melting or crystallization, has chemically fractionated 231Pa from 227Ac or if they were chemically fractionated then this process occurred more than 100 years ago.

In situ analysis of anorthoclase and melt inclusions

Average Ba and Th concentrations in two 1984 (ER84-01 and ER84-04) and one 1988 (ER88-01) anorthoclase megacrysts are presented in Table 4 (see Appendix Table A1 for raw data). The Ba concentrations in melt inclusions measured by ion microprobe are highly variable, but overall comparable with the ID values obtained by MC-ICP-MS in this study. The Th concentrations measured by ion microprobe in the anorthoclase are significantly lower than those obtained on bulk mineral separates using ID-ICP-MS.

For the electron microprobe measurements, Pb and U concentrations in Erebus apatite were always below the detection limit of 100 and 40 ppm, respectively, whereas the Th concentration ranges from slightly above detection limit (40 ppm) up to around 100 ppm.

DISCUSSION

Phonolites ejected by the Strombolian activity of the Erebus lava lake have been sampled on an almost yearly basis from 1972 to 2005 (Figs 4, 5 and 8). This sample suite has allowed for comprehensive geochemical analysis (Kyle et al., 1992; Caldwell & Kyle, 1994; Kelly et al., 2008b; Sims et al., 2008a) and provides a unique opportunity to constrain the temporal evolution of these highly differentiated lavas. These constraints are most robust with regard to the shallow, most differentiated part of the Erebus system because of the great petrological gap between the differentiated phonolites and their mantle-derived parental basanites.

Temporal variations in U-series

Anorthoclase and glass separates each have (230Th/232Th), (238U/232Th), (230Th/238U) and (226Ra/230Th) activity ratios that are uniform and normally distributed within analytical uncertainties. The one possible exception to this temporal uniformity is a bomb erupted in 1984 during a 4 month period of sustained Strombolian eruptions. Both ER 84501G and ER 84505G (which were replicated six times on separate powder aliquots) have lower (230Th/232Th) and skew the distribution slightly (Table 5). The average (230Th/232Th) for these two 1984 bombs, including replicates, is 0·9767 ± 0·0045 (2σ), whereas the average (230Th/232Th) for all the bombs from other years, including replicates, is 0·9854 ± 0·0046 (2σ). Admittedly, this difference is small (8·9 per mil) and overlaps at the 95% confidence level. However, in light of the fact that (1) 1984 was an exceptional year with larger and more frequent Strombolian eruptions (Kyle, 1986; Caldwell & Kyle, 1994) than previous and following years, and (2) the 1984 bombs have slightly lower 206Pb/204Pb and 208Pb/204Pb and higher 87Sr/86Sr (Sims et al., 2008a), we consider this difference in (230Th/232Th) as possibly real. For the long-lived isotopes we hypothesized that this difference for the 1984 bombs reflected a change in the magma conduit resulting in assimilation of shallow-level trachytes (Sims et al., 2008a).

Anorthoclase/phonolite glass partitioning

The (230Th/238U) and (226Ra/230Th) measurements of coexisting anorthoclase and phonolite glass provide a relative sense of U/Th and Th/Ra partitioning during anorthoclase crystallization. (230Th/232Th) is the same in both the anorthoclase and phonolite glass, whereas (238U/232Th) is fractionated (lower in the anorthoclase), suggesting that Th is more compatible than U. (226Ra/230Th) is much greater in the anorthoclase than in the phonolite glass, indicating that Ra is more compatible than Th. The relative partitioning of Ra and Pb cannot be determined from (210Pb/226Ra), as this activity ratio is unity (i.e. in radioactive equilibrium) in both phases.

Using our measured 238U, 232Th, 226Ra and Ba concentrations for the phonolite glasses and anorthoclase separates (Table 1), we can calculate ‘effective’ anorthoclase/melt partition coefficients for Th, U, Ra and Ba (Table 6). The calculated order of compatibility is DU < DThforumlaDRaforumlaDBa. These calculated partition coefficients (D values) represent values averaged over the whole of several crystals (10 g) and are limited by the purity of the anorthoclase and phonolite glass separates. The observation that the D values are very different but relatively uniform for all five anorthoclase–glass pairs suggests that the separation methods were effective. DRa is determined from 226Ra, which has been decaying (within the anorthoclase) and ingrowing (within the glass) toward an ‘equilibrium’ value with 230Th over the lifetime of the crystal. Thus, the calculated DRa is a minimum value.

Table 6

Calculated anorthoclase/phonolite glass partition coefficients for U, Th, Ra and Ba based on isotope dilution measurements (Table 1) and for Th and Ba based on in situ measurements (Table 4)

 DTh DU DRa DBa 
Bulk analysis of separated phases 
Average 0·0080 0·0071 0·42 4·86 
SD (1σ) 0·0013 0·0011 0·04 0·25 
In situ measurements     
Average 0·0032   4·37 
SD (1σ) 0·0022   0·19 
 DTh DU DRa DBa 
Bulk analysis of separated phases 
Average 0·0080 0·0071 0·42 4·86 
SD (1σ) 0·0013 0·0011 0·04 0·25 
In situ measurements     
Average 0·0032   4·37 
SD (1σ) 0·0022   0·19 

The average values are calculated from the individual effective anorthoclase/phonolite glass partition coefficients.

We can also calculate ‘effective’ anorthoclase/phonolite melt partition coefficients (Table 6) using our in situ Th and Ba measurements for anorthoclase and included melt. The in situ Ba concentrations show a large range of ‘intra’-crystal variability reflecting both true compositional variations within the crystal and uncertainties in the measurement. This variance is in contrast to the bulk separates, which show rather limited ‘inter’-crystal differences over five separate samples from different years. At the 1σ confidence level, the average DBa calculated from the in situ SIMS measurements overlaps the DBa calculated from the bulk anorthoclase. This similarity in the two different calculated values for DBa suggests that the ‘bulk’ anorthoclase and glass separates represent a homogenized average relatively clean from mutual cross-contamination in terms of their Ba (and by inference Ra) budget(s). The in situ Th concentrations of the melt inclusions measured by SIMS are also similar to the ID-ICP-MS measurements for the bulk glass separates. However, for the anorthoclase megacrysts, the in situ Th concentrations are much lower than the Th concentrations measured in the bulk separates. The resulting DTh values calculated from the in situ measurements are lower than determined from the bulk separates.

The large difference in U/Th between the anorthoclase and phonolite glass exceeds the extent of elemental fractionation predicted for either alkali or plagioclase feldspar crystallization (Blundy & Wood, 2003). Because the Th concentrations for anorthoclase measured in situ are much lower than measured in the bulk ‘anorthoclase’ separates we hypothesize that mineral separation failed to fully remove small mineral inclusions, such as apatite, from the anorthoclase and that these accessory mineral inclusions are influencing Th/U fractionation. The Erebus anorthoclase crystals contain abundant mineral inclusions of pyroxene, apatite, titanomagnetite, olivine and pyrrhotite trapped along apparent growth zones in the anorthoclase crystals (Fig. 2). Detailed BSE imaging of a set of 23 variably sized anorthoclase crystals indicates that pyroxene is the most abundant included phase, and also is the largest of the included phases, with sizes commonly up to 1 mm or greater. Apatite inclusions are also present in every crystal studied, and apatite is estimated to be the second most abundant included crystal. Many apatite crystals are small (of the order of 100 µm), but some larger crystals (up to 1 mm in length) are observed. Using image processing of Ca X-ray maps, the determined surface abundance of apatite ranges from 0·03 to 0·38% with an average of 0·14%. Most of the values for surface per cent of apatite exposed fall in a relatively small range between 0·07 and 0·15%. Accessory minerals, such as apatite, contain large amounts of U and Th (this work; Dawson & Hinton, 2003). Apatite/silicate melt U and Th partitioning experiments by Prowatke & Klemme (2006) showed that DTh and DU range from nearly equal to DTh > DU (with DTh/DU ranging from 0·7 to 12·78). Our electron probe measurements of Erebus apatite also suggest that Th is more compatible than U in apatite, the Th in the apatite ranging from 40 to 100 ppm, whereas U was below the detection limit (40 ppm). Given the extent and size of apatite in the Erebus anorthoclase and its potential to fractionate U from Th we suggest that our mineral separation methods did not adequately remove the apatite from the anorthoclase and that these small apatite inclusions lowered the (238U/232Th) of the anorthoclase crystals. Thus, for anorthoclase–phonolite melt partitioning the DTh determined from the in situ measurements is preferred over the DTh determined from measurements of bulk anorthoclase and glass separates.

The complementary (226Ra/230Th) in the anorthoclase and phonolite glass provides strong evidence that anorthoclase crystallization chemically fractionates Ra from Th. Based on measured concentrations we calculate DBa ∼ 4·5 (i.e compatible), DTh/DBa ∼ 7e – 4 and DRa/DBa ∼0·1. The observation that Ra and Ba are much more compatible than Th during feldspar crystallization is predicted theoretically (both Ra and Ba are alkali earths), and is consistent with partitioning experiments (Miller et al., 2007; Fabbrizio et al., 2009) and other measurements (Volpe & Hammond, 1991; Reagan et al., 1992; Cooper et al., 2001, 2003; Cooper & Reid, 2003; Zellmer et al., 2008; Cooper, 2009; Rubin & Zellmer, 2009). Our value for DRa/DBa of ∼0·1 is a lower limit (because of potential 226Ra decay) and has important implications for using Ba as a stable proxy for Ra when evaluating the time scales of magma evolution and anorthoclase growth. Previous work on Erebus samples (Reagan et al., 1992) assumed DRa/DBa = 1 during anorthoclase crystallization. However, Ba2+ and Ra2+ have different ionic radii (1·42 Å and 1·48 Å, respectively) in VIII-fold coordination (Blundy & Wood, 2003) and both theoretical predictions and experimental evidence indicate that DRa/DBa is significantly less than one during plagioclase and alkaline feldspar crystallization (Miller et al., 2007; Fabbrizio et al. 2009). Because of the unusual composition of Erebus anorthoclase (An16 Ab64·5 Or18·7), the absolute and relative compatibilities of Ba and Ra are not easily predicted from theoretical models of elastic strain moduli. These models are for the end-member plagioclase or alkali feldspar series and assume single substitution, whereas during anorthoclase crystallization Ba and Ra partitioning would probably involve a coupled substitution. Nevertheless, using the parameterizations of, for example, Blundy & Wood (2003), for an alkali feldspar with an orthoclase composition of Or18, DBa is predicted to be ∼6·3 and DRa/DBa ∼0·2, whereas for plagioclase with An17 (at ∼1000°C) DBa is predicted to be ∼1·4 and DRa/DBa ∼0·22. These values are similar to those calculated by Fabbrizio et al. (2009), who used an orthoclase composition of Or19 and a temperature of 1000°C to calculate a DRa/DBa of 0·29. These predictions are qualitatively consistent with our measured results, which indicate that during anorthoclase crystallization Ba is highly compatible and DRa/DBa is between 0·1 and one, with the partitioning ratio probably being closer to 0·1.

Crystal growth rates and the time scales of magma differentiation

Our measurements of 238U–230Th, 230Th–226Ra, 226Ra–210Pb and 232Th–228Ra–228Th disequilibria provide four salient observations pertinent to determining anorthoclase crystal growth rates and magma residence times in the shallow Erebus system.

  1. On a (230Th/232Th) vs (238U/232Th) isochron diagram both the anorthoclase and phonolite glass show significant (230Th/238U) disequilibria and cluster tightly in two distinct groups forming horizontal two-point isochrons (Fig. 3).

  2. On a (226Ra)/Ba vs (230Th)/Ba isochron diagram, the anorthoclase and glass lie on opposite sides of the ‘equiline’ forming steeply inclined two-point isochrons (Fig. 9).

  3. (210Pb/226Ra) values are within error of equilibrium for both the anorthoclase and phonolite glass (Fig. 6).

  4. (228Ra/232Th) values are within error of equilibrium in the phonolite glasses, but (228Ra/232Th) is >1 for the anorthoclase (Fig. 8).

Fig. 9

Ra/Ba isochrons. 230Th–226Ra–Ba data for anorthoclase and glass separates plotted on a conventional Ba-normalized 230Th–226Ra isochron diagram

Fig. 9

Ra/Ba isochrons. 230Th–226Ra–Ba data for anorthoclase and glass separates plotted on a conventional Ba-normalized 230Th–226Ra isochron diagram

Instantaneous crystal fractionation with a magma residence time of 102–103 years is consistent with the first three of these four constraints. The 238U–230Th data form horizontal ‘zero-age’ isochrons having an age resolution of 5–10 ka; 230Th–226Ra data form inclined two-point isochrons giving ages of ∼2500 years for DRa/DBa = 1, down to hundreds of years for DRa/DBa = 0·1 (Fig. 10); and the anorthoclase and glass separates both have equilibrium (226Ra/210Pb) values suggesting that they are older than 100 years. The equilibrium (227Ac/231Pa) in the phonolite glasses is also consistent with a magma residence time greater than 100 years, but without measurements of these nuclides in the anorthoclase these data have no direct bearing on the time scales of anorthoclase crystallization.

Fig. 10

Ra age vs DRa/DBa showing how the age of crystallization changes as a function of DRa/DBa. Calculated DRa/DBa values for Or18 and An17 are taken from Blundy & Wood (2003). The value of DRa/DBa of ∼0·1 is a lower limit (because of potential 226Ra decay).

Fig. 10

Ra age vs DRa/DBa showing how the age of crystallization changes as a function of DRa/DBa. Calculated DRa/DBa values for Or18 and An17 are taken from Blundy & Wood (2003). The value of DRa/DBa of ∼0·1 is a lower limit (because of potential 226Ra decay).

That being said, however, the observation that (228Ra/232Th) significantly exceeds unity in several young (1984, 1988) anorthoclase separates (Reagan et al., 1992, 2006; this study) is inconsistent with simple instantaneous crystal growth and requires some portion of the crystal to have grown less than 25 years prior to eruption. This is not surprising as numerous petrographic observations, in situ measurements, and determinations of crystal size distributions (e.g. Dunbar et al., 1994; Sumner, 2007; Kelly et al., 2008b) suggest that anorthoclase crystal growth was episodic and probably continued until eruption. Additionally, there are considerable data indicating that crystallization of magmas occurs over significant periods of time (tens to thousands of years) as heat is transferred away from the magma-wall–rock interface (see, e.g. Albarède, 1993; Charlier & Zellmer, 2000; Vazquez & Reid, 2002; Turner et al., 2003; Hawksworth et al., 2004; Zellmer & Clavero, 2006; Reagan et al., 2008). In light of this understanding, it is essential to explicitly consider the time scales of magma differentiation processes in the context of the half-lives of the shorter-lived U and Th decay series nuclides (e.g. 226Ra, 210Pb, 210Po, 222Rn, 227Ac, 228Ra, 228Th, etc.).

To account for all of the abundances of short-lived nuclides in the 238U and 232Th decay series in a manner consistent with our understanding of mineral abundances, partition coefficients, and geochemical variations in the Erebus system, we have developed a finite-element, continuous crystallization model (Figs 11–13). Although our initial model is similar in many respects to the continuous constant rate crystallization model presented by Snyder et al. (2004, 2007) for 238U–230Th–226Ra, we have the distinctive added requirement to be able to explain all of the U and Th decay time-series fractionations, including the very short-lived 232Th–228Ra system.

Fig. 11

Schematic illustration of the Erebus shallow magmatic system, detailing compositional constraints, model parameters and assumptions.

Fig. 11

Schematic illustration of the Erebus shallow magmatic system, detailing compositional constraints, model parameters and assumptions.

Fig. 12

Continuous closed-system crystallization model as described in the text.

Fig. 12

Continuous closed-system crystallization model as described in the text.

Fig. 13

(a) Goodness of fit of measured data to continuous closed-system crystallization model. It should be noted that the closed-system crystallization model (no recharge) cannot reproduce the compositions of Erebus anorthoclase and glass separates without setting the DBa anorthoclase/phonolite glass to a value of 2·8, indicating that the Ba concentrations must have been maintained at higher levels than allowed for by anorthoclase crystallization, implying that magma recharge was accompanying crystallization. (b) Goodness of fit of measured data to open-system crystallization model. Magma recharge to crystallization rate ratios set to a value of two. This uses the measured DBa anorthoclase/phonolite glass value of 4·5. For the crystals the Th misfit is interpreted to result from both residual apatite and glass that was not completely separated.

Fig. 13

(a) Goodness of fit of measured data to continuous closed-system crystallization model. It should be noted that the closed-system crystallization model (no recharge) cannot reproduce the compositions of Erebus anorthoclase and glass separates without setting the DBa anorthoclase/phonolite glass to a value of 2·8, indicating that the Ba concentrations must have been maintained at higher levels than allowed for by anorthoclase crystallization, implying that magma recharge was accompanying crystallization. (b) Goodness of fit of measured data to open-system crystallization model. Magma recharge to crystallization rate ratios set to a value of two. This uses the measured DBa anorthoclase/phonolite glass value of 4·5. For the crystals the Th misfit is interpreted to result from both residual apatite and glass that was not completely separated.

The objective of our modeling is to reproduce all of the observed U and Th decay series disequilibria, as well as Th and Ba abundances in both the phonolite glass and the anorthoclase. As starting compositions we use a tephriphonolite (SiO2 = 55·93 wt%; MgO = 1·23 wt%) with compositions near the inflection point in the ‘Erebus lineage’ (Kyle et al., 1992; Kelly et al., 2008a) that marks the change from plagioclase to anorthoclase crystallization, causing Ba to switch from being incompatible to compatible. Because of the complementary (226Ra/230Th) values in anorthoclase–glass pairs, we assume that the Erebus phonolitic magmatic system was closed and as the anorthoclase grew the residual melt or magma was effectively fractionated. The assumed crystallizing assemblage contains 90% anorthoclase, which is typical of Erebus whole phonolite bombs (Kyle et al., 1992; Kelly et al., 2008b). The fractionation extent per step was chosen to produce a best-fit model at ∼10 000 iterations (Appendix B).

The finite-element technique utilized here allows for a variety of crystallization models to be explored, including closed-system continuous crystallization (Fig. 12), variable-rate crystallization, even dissolution, and dissolution and recrystallization. These methods can also model open systems with magma recharge (Appendix B, Fig. 13). Nevertheless, because potential variations in crystallization rates are essentially unconstrained in this system, our models assume a constant rate of crystal fractionation (i.e. dF/dt = k) and therefore an approximately constant volumetric growth rate of anorthoclase (Fig. 12). Actual crystal growth consistent with this model does not necessarily have to be constant. Indeed, crystal size distribution studies (Dunbar et al., 1994) at Erebus suggest that the anorthoclases have undergone multiple episodes of growth and periods of resorption. Nevertheless, as long as the rates of the processes that control parameters such as magma cooling and degassing rates remain roughly uniform, then the overall growth rate will be roughly constant.

Bulk partition coefficients for Ba and Th employed in these models are constrained by the in situ analyses, but as will be shown below, the effective partition coefficient for Th appears to differ slightly from measured values owing to the entrapment of melt and apatite during crystallization. DRa values are resolved in best-fit models by finding solutions that match (226Ra/230Th) and (228Ra/232Th) in the anorthoclase and glass separates. DPb values are not well constrained for these models because of a poor understanding of Pb concentrations in the starting magmas and the equilibrium (210Pb/226Ra) values measured in the young phonolites. We use bulk DPb = 0·2, which is the value given for K-feldspar in trachyte by Larsen (1979), and is consistent with the positive correlations between Pb and other incompatible trace elements in differentiated alkaline lavas from Erebus (Kyle et al., 1992; Kelly et al., 2008b).

We first explored a variety of closed-system crystallization models (no recharge) to reproduce the compositions of Erebus anorthoclase and glass separates (Fig. 12). All the crystallization models using partition coefficients close to those measured did not fractionate Th and Ba to the level required. For example, the amount of crystallization required to reproduce the Th concentrations in the phonolite glasses removed too much Ba from the melt (Table 4). The Ba concentrations must have been maintained at higher levels than allowed for by anorthoclase crystallization, which implies that magma recharge was accompanying crystallization.

We then employed a numerical approach similar to the closed-system models, but in this model we allow magma recharge to accompany crystallization. Our initial open-system models utilized the measured partition coefficients from Table 6 for Th and Ba in anorthoclase, and assumed partition coefficients for both elements in other crystals in the crystallizing assemblage to be zero. To model both Th and Ba concentrations, magma recharge to crystallization rate ratios had to be above two. Whereas Th and Ba concentrations and the ratios between U-series nuclides can be reproduced utilizing the measured partition coefficients, the (226Ra/230Th) values were always higher in the models than in the glasses. This reflects the relatively low DRa required for modeling the (226Ra/230Th) values for the anorthoclase when the DTh is within the range of measured values for the anorthoclase–glass pairs (<0·008). The models that fit the trace element and U-series data best utilize DTh in anorthoclase of 0·012–0·016 (Table 6). An effective partition coefficient for Th in this range could result from the presence of 0·5–1% residual melt inclusion glass (and apatite) within the measured anorthoclase crystals. This level of melt entrapment would have little influence on the effective partition coefficients for compatible elements such as Ba. It should be noted that the DRa/DBa values of 0·2–0·24 in anorthoclase from these best-fit models are consistent with the theoretical feldspar values cited above, and having some residual glass during mineral purification is probable given that the anorthoclase megacrysts originally contained over 30% melt inclusions before purification.

As noted by O’Hara & Mathews (1981), trace elements eventually reach steady-state concentrations in magma chambers with constant recharge. The high recharge rates and degrees of crystallization modeled here are close to the values required for steady-state compositions to be attained, even for incompatible elements such as Th. Thus, the young anorthoclase phonolites from Erebus have major and trace element compositions that are close to those of a steady-state system.

The high recharge rates required by the best-fit models imply that the volume of phonolite in the shallow magma chamber and lava lake system at Erebus has been growing. This makes sense from a heat-balance point of view, as it explains how a lava lake and shallow magma reservoir can remain liquid (see Calkins et al., 2008). Our modeled duration of reservoir growth is ∼2000 years. This age coincides within error with the ages of the youngest two lavas from the flanks of Erebus (the Northwest and Upper Ice Tower Ridge flows; Harpel et al., 2004). This permits speculation that the Erebus summit magma chamber system began to grow after these eruptive events and is at present amassing magma in advance of another significant lava eruption from a flank vent.

The open-system process required to produce Erebus phonolites and the calculated millennium-scale magma accumulation time contrast significantly with the mode and time scale of production of the phonolite recently erupted subaqueously from the flanks of Nightingale Island near Tristan da Cunha (Reagan et al., 2008). At Tristan da Cunha, the phonolite had trace element compositions consistent with its generation by simple crystal fractionation of more mafic parents. The Nightingale phonolite also had significant 210Pb excesses over 226Ra and small 228Ra deficits with respect to 232Th, implying that the crystal fractionation lasted one to three centuries. The contrast in short-lived radionuclide ratios between Nightingale and Erebus demonstrates the power of using these ratios in conjunction with other geochemical data to determine the mechanisms and time scales of magma differentiation. At Nightingale, the data are consistent with a small volume of isolated alkaline magma that rapidly fractionated to phonolite before eruption. In contrast, the compositions of Erebus phonolites suggest that the shallow magma body beneath its lava lakes is growing and has been doing so for about two millennia.

Ra–Th geochronometery

It has been debated how the 226Ra–230Th disequilibrium can be used as a geochronometer for petrological events in magmas (Cooper, 2009; Rubin & Zellmer, 2009). Rubin & Zellmer (2009) showed that if DRa < DTh in minerals (e.g. magnetite) in equilibrium with a melt, and the equilibrium process is terminated rapidly (e.g. by removal of the melt or rapid cooling), then the Ra–Ba isochron method (e.g. Fig. 9) can be used to mark the time since the equilibrium was maintained because the differences in DRa and DBa affect the calculated ages less than the precision of the technique. However, ages calculated using Ra–Ba isochrons and minerals with DRa > DTh, such as the anorthoclase in Erebus phonolites, are erroneous because differences in DRa and DBa strongly affect the calculated ages (Fig. 10). In most cases, DRa < DBa because of the greater size of Ra2+ relative to Ba2+, which would lead to an overestimation of the age of any instantaneous event (Cooper et al., 2001; Cooper & Reid, 2003; Cooper, 2009).

Another issue with using any isochron method on large and long-lived systems such as Erebus is that crystals can grow over significant amounts of time (e.g. Reagan et al., 1992; this study) or can be recycled from older magmas (e.g. Pyle et al., 1998; Reagan et al., 2003; Sims et al., 2007). Thus, bulk samples can contain crystals grown over a variety of time scales and conditions, and their chemical compositions represent integrated averages. To obtain quantitative geochronological information from Ra–Th disequilibria in this circumstance, a continuous long-duration model, such as the one used here, is required. We should note that the high recharge rates used in our open-system models extend the calculated magma accumulation time significantly farther than the time frame calculated by simple continuous crystal fractionation models.

Implications for the processes and time scales of magma degassing from 210Pb/226Ra and 210Po

The measured equilibrium values for (210Pb/226Ra) and (227Ac/231Pa) and our open-system modeling for (226Ra/230Th) and (228Ra/232Th) suggest that the shallow magma residence time of the Erebus lavas is greater than 100 years. However, given that the Erebus lava lake is continuously degassing, both by quiescent emissions and Strombolian eruptions, the observation that 210Pb is in equilibrium with 226Ra (Fig. 8) is rather surprising. One would expect that removal of 222Rn, a noble gas that lies between 226Ra and 210Pb, by magma degassing would have a significant impact on the (210Pb/226Ra) of the magma (see, e.g. Gauthier et al., 2000; Berlo et al., 2010; Reagan et al., 2006; Kayzer et al., 2009). Devolatilization generally has a negligible effect on the concentrations of 210Pb or 226Ra themselves, as Ra is not volatile, and only about 1% of Pb typically degasses from magmas (Lambert et al., 1976; Rubin, 1997; Gauthier et al., 2000). In contrast, Rn strongly partitions into the gas phase from magma (Sato & Sato, 1977; Sato et al., 1980; Gill et al., 1985; Sato, 2003; Sims & Gauthier, 2007). So, if magma degassing endures for years to decades, (210Pb/226Ra) values are lowered as a function of the half-life of 210Pb and the 222Rn degassing efficiency (Gauthier & Condomines, 1999). Similarly, if 222Rn concentrates in the gas phase of a magma that has stalled somewhere, then 210Pb generated by decay of 222Rn can lead to a 210Pb excess in the magma (e.g. Turner et al., 2004; Berlo et al., 2010; Reagan et al., 2006; Kayzar et al., 2009). However, it should be noted that tens to thousands of times as much magma must degas 222Rn compared with the amount of magma that collects it because of the large differences in the half-lives of 222Rn (3·8 days) and 210Pb (22 years) (Reagan et al., 2006; Kayzar et al., 2009; Condomines et al., 2010).

Given the observed persistent degassing of the Erebus lava lake (both quiescent and Strombolian), four scenarios can explain the measured (210Pb/226Ra) of unity in the young, known age Erebus phonolite bombs, as follows.

Hypothesis 1: 222Rn is soluble in the Erebus phonolite melt and thus is not degassing from the magma. We consider this ‘soluble 222Rn scenario’ highly unlikely as Rn is a ‘noble gas’ and therefore should not be soluble in the Erebus phonolite melt. Although concentrations of 222Rn are too low to become saturated in the phonolite magma and produce a separate gas phase, it is generally assumed that 222Rn degassing is governed by Henry’s law partitioning of 222Rn between magma and a gas phase resulting from saturation of the major volatile species H2O, CO2, and SO2 (see, e.g. Gauthier & Condomines, 1999). H2O, CO2, and SO2 are significantly degassing from the Erebus lava lake (Oppenheimer and Kyle, 2008; Oppenheimer et al., 2009; Sweeney et al., 2008) and thus a ‘soluble 222Rn scenario’ would require the chemical behavior of the 222Rn to be in marked contrast to that of the other gaseous species H2O, CO2, and SO2. In addition, as will be discussed below, 210Po has been degassed relative to 210Pb, suggesting that the volatile, short-lived radionuclides are also being degassed.

Hypothesis 2: Mixing occurs between a surface magma, which has (210Pb/226Ra) < 1 from continuous degassing of 222Rn, with a deep magma that has (210Pb/226Ra) > 1 from continuous 222Rn accumulation. It is not possible to rule out this scenario, as the lava lake is convecting and in continuous motion, suggesting stirring of the system at the shallowest levels. Nonetheless, the perfect mass balance required to produce magma with (210Pb/226Ra) = 1 in all samples seems fortuitous.

Hypothesis 3: The anorthoclase-rich phonolite circulating in the upper part of the magmatic system is made of old, degassed magma that is not communicating with gas emissions that are coming from deeper parts of the magma chamber. This old, degassed magma hypothesis is also consistent with the observation that some melt inclusions in anorthoclase are highly degassed compared with those in olivines from parental basanites from around the flanks (Oppenheimer et al., 2011). That said, there is clear persistent degassing of the Erebus lake and this degassing is hypothesized, based on Fourier transform infrared (FTIR) spectroscopic measurements, to have three components: (1) a passively emitted water-rich gas exolved near the lake surface when magma episodically enters the lake (mass CO2/H2O ∼0·1); (2) a passively emitted deeper-sourced gas that percolates though a permeable magma-filled conduit (mass CO2/H2O ∼1·8) (Oppenheimer et al., 2009); (3) a deeper-sourced CO2-rich gas (mass CO2/H2O ∼4·3) observed during Strombolian eruptions (Oppenheimer et al., 2011). For this third hypothesis to be correct, the gas sourced from deeper in the system would need to ascend without interacting with the shallow magma system.

Hypothesis 4: Internal redistribution of radon occurs between a deep degassing magma and a shallow previously degassed magma. In this ‘internal radon redistribution’ scenario degassing is occurring by two distinct processes: (1) weak shallow degassing from the lava lake; (2) more intense degassing from deeper levels. Melt inclusions in some anorthoclase crystals (Oppenheimer et al., 2011) suggest that some lava lake phonolite is extensively degassed and thus could have a (210Pb)/(226Ra) activity ratio significantly lower than one. If deeper degassing events, which are contributing significantly to the Erebus gas (CO2) flux, supply radon to the shallowest level where it could decay and produce 210Pb atoms, then this would counterbalance the 210Pb loss owing to shallow 222Rn degassing. Essentially, in this scenario the 222Rn flux into the bottom of a shallow reservoir is equal to the 222Rn flux out of the magma system.

Relevant to degassing at Erebus is the observation that a lava bomb erupted on 16 December 2005, has 210Po not completely degassed from the lava at the time of eruption (Fig. 7). Incomplete degassing of 210Po is atypical for subaerially erupted lavas (e.g. Gill et al., 1985; Rubin et al., 1994; Rubin, 1997; Reagan et al., 2006, 2008). This incomplete degassing observed for the Erebus phonolite indicates that either the solubility of Po is higher than assumed in the Erebus phonolite (again unlikely given what we have seen in other volcanic systems), or that the average Erebus magma represented in the erupted lava bombs was on average degassed up to 80 days prior to eruption. Degassing could not have started two or more years before eruption or we would observe this in the 210Pb/226Ra.

To quantitatively address the equilibrium (210Pb/226Ra) (i.e. the effect of the intermediate volatile nuclide 222Rn on this ratio) and the incomplete degassing of 210Po, we have incorporated these volatile nuclides into our finite element model of magma evolution and anorthoclase crystallization discussed above. In this model, we assume that the intruding parental magma from depth was completely degassed and thus had activities of 210Po = 0 and 222Rn = 0 (210Pb in the intruding magma can be set to almost any value from unity to zero without much effect on the model glass 210Pb/226Ra). We set (1) the melt/gas partition coefficients of Rn and Po to zero, (2) the time steps to 7 days, which is about 5% of the half-life of 210Po, and (3) 210Po and 222Rn to degass at 1·3% per day. Because of the short half-life of 222Rn (t1/2 = 3·85 days), 222Rn ingrows rapidly, and thus the effect on 210Pb (and 226Ra/210Pb) is negligible. The 210Po/210Pb ratio, however, quickly moves to a steady-state value at 0·25, similar to the measured value (Fig. 7). Although this model fits the data well, it is probably not completely realistic as it assumes that the entire shallow magma chamber is stirred on a short enough time scale to keep 210Po/210Pb (and all other geochemical parameters) homogeneous throughout the chamber. Nevertheless, it shows that Erebus with its shallow stagnated magma does not lose all of its 210Po all of the time. The model also implies that our samples resided only in the shallow chamber before being thrown out by the Strombolian explosions. This is in contrast to most lavas and tephra from other volcanoes that are tapped from deeper levels and degas over long paths to the surface (Reagan et al., 2008).

U and Th decay series constraints on magma genesis and melt transport

In basaltic systems, it is well established that the longer-lived U-series isotopes (i.e. 230Th, 226Ra, 231Pa) can provide important constraints on the processes and time scales of magma genesis, namely magma production rates, time scales of melt extraction and the lithology (garnet/cpx) of the mantle source (McKenzie, 1985; Speigelman & Elliott, 1993; Lundstrom et al., 1995; Elliott, 1997; Sims et al., 1995, 1999, 2002; Bourdon & Sims, 2003; Stracke et al., 2006; Prytulak & Elliott, 2009; Waters et al., 2011). However, in more evolved volcanic systems, such as at Erebus, shallow-level magmatic processes can overprint and thus obfuscate much of the information on the initial melting processes.

In the Erebus magma system, shallow crustal processes and time scales have clearly and significantly influenced the relative activities of 230Th–226Ra, 210Pb–226Ra, 210Pb–210Po, 232Th–228Ra and 235Pa–227Ac and, as a result, these parent–daughter nuclides cannot provide direct information on the time scales and nature of the initial melting processes. Even so, we argue that the U–Th isotope system can still provide constraints on the lithology (cpx/gt) and long-term nature of the Erebus mantle source.

With regard to the lithology (cpx/gt) of the Erebus mantle source, based on experimentally determined U and Th mineral–melt partition coefficients the large 238U–230Th disequilibria of the Erebus phonolites indicate that melting began deep and most probably in the presence of garnet (Beattie et al., 1993a, b; LaTourrette et al., 1993; Landwehr et al., 2001; Elkins et al., 2008). The (230Th/238U) disequilibria measured in these phonolites are minimum values, as magma residence time would decrease the (230Th/232Th) of the lava and apatite fractionation would increase its (238U/232Th), which in both instances would decrease the (230Th/238U) of the lava. Our modeling of the anorthoclase crystallization ages and magma residence times suggests that magma residence times in the upper part of the Erebus magma system are short relative to the half-life of 230Th; as such, the decrease in (230Th/232Th), because of 230Th decay, will be insignificant compared with the analytical uncertainty. Although apatite fractionation appears to have lowered the (238U/232Th) and attendant (230Th/238U), it could not have produced the observed 238U–230Th disequilibria, for if it had, the Erebus lavas would be depleted in 230Th relative to 238U, rather than enriched as observed. Furthermore, because of (1) the small amount of apatite in the Erebus lavas and phonolites, (2) the coherent behavior of U and Th as highly incompatible elements and (3) the lack of correlation between Th/U and P2O5 through the entire liquid line of descent seen in the Mt. Erebus lavas (basanite to phonolite), we conclude that it is unlikely that apatite has significantly altered the (230Th/238U) disequilibrium originally imparted by the melting process.

With regard to the mantle sources of the Erebus lavas, as noted by Sims & Hart (2006) and discussed in the background section above, the 230Th/232Th and U/Th of Erebus bombs are intermediate relative to other ocean island and mid-ocean ridge basalts, and form the end-member HIMU mantle component on plots of Pb isotopes versus (230Th/232Th) and U/Th. Although Erebus does not represent the end-member mantle component HIMU, there are no samples young enough for U–Th disequilibria studies from Mangaia–Tubaui (the HIMU end-member). Hence, for the U–Th isotope system, Erebus, by necessity, represents the best available end-member approximation for the HIMU source.

CONCLUSIONS

This is the first study in which all the relevant nuclides from the 238U, 235U, and 232Th decay series have been measured in the lavas of a volcanic system. The contrasting chemistries and half-lives of these different nuclides have allowed us to investigate a variety of shallow magmatic processes over a wide range of time scales (from days to 105 years), thereby providing a better understanding of the rates and processes of magma differentiation than has been established in earlier studies. Furthermore, this work provides some perspective on the long-range hazard posed by the Erebus volcano.

The main conclusions from this study are as follows.

(1) U–Th–Ra disequilibria are uniform over 34 years of history. In the phonolite glass, from 1972 to 2005, (238U/232Th) and (230Th/232Th) are uniform, with one possible exception being the 1984 samples, which were erupted during a time of significantly increased Strombolian activity. (226Ra/230Th) is also uniform over the 34 year historical record. This uniformity suggests that the Erebus magma system, over the 34 year sample span, is in a state of dynamic equilibrium.

(2) For the Erebus lavas the equilibrium (210Pb/226Ra) and (227Ac/231Pa) suggest that the magma residence time is long compared with the half-lives of 210Pb (t1/2 = 22·6 years) and 227Ac (t1/2 = 21·77 years). However, the observation that (228Ra/230Th) (t1/2 = 5·77 years) is out of equilibrium in the anorthoclase suggests that the Erebus magmatic differentiation and crystallization processes are continuous and continuing. With this observation in mind we have developed an open-system, finite-element, continuous-crystallization model that incorporates ingrowth and decay of the different nuclides in the continuously growing anorthoclase crystals and associated phonolitic melt. Because we have measured Ba and Th concentrations in both the anorthoclase crystals and the melt we can constrain the anorthoclase/phonolite melt partitioning of Ba and Th and have found that the only way to replicate the observed Ba and Th concentrations and 238U, 235U and 232Th decay series data is to incorporate magma recharge into the shallow system. To successfully model the present dataset the recharge rate has to exceed the crystallization rate, which implies that the shallow magma reservoir within Erebus is growing. This has important implications for hazard assessment. Our modeled duration of reservoir growth is ∼2000 years, which coincides, within error, with the ages of the two youngest lavas from the flanks of Erebus (the Northwest and Upper Ice Tower Ridge flows; Harpel et al., 2004). This observed coincidence indicates that the Erebus summit magma chamber system began to grow after these eruptive events and is at present amassing magma in advance of another significant lava eruption from a flank vent. An important cautionary implication of this modeling is that when the time scale of crystallization is comparable with the half-life of the daughter nuclide of interest, the simple isochron techniques typically used in most U-series studies can provide erroneous ages.

(3) (210Pb/226Ra) is in equilibrium despite the observed persistent degassing of the Erebus lava lake. This observation requires that either (1) 222Rn is soluble in the Erebus phonolite melt and thus is not degassing from the magma, (2) mixing is taking place between a surface magma, which has (210Pb/226Ra) < 1 from continuous degassing of 222Rn, and a deep magma that has (210Pb/226Ra) > 1 from continuous 222Rn accumulation, (3) the anorthoclase-rich phonolite circulating in the upper part of the magmatic system and made of old, degassed magma is not communicating with gas emissions coming from deeper parts of the magma chamber, or (4) internal redistribution of radon is occurring between a deep degassing magma and a shallow previously degassed magma. The observed degassing of the Erebus lava has also partially removed 210Po from the magma. As a result, we prefer the fourth model, in which the flux of the intermediate 222Rn into the shallow system equals the flux of 222Rn degassing out of the system.

(4) Although shallow-level magmatic processes at Erebus have obfuscated much of the information on the initial melting processes, the significant and uniform 230Th excesses measured in the Erebus phonolites suggest that the parental basanites had residual garnet in their source.

(5) Finally, Erebus lavas form the end-member HIMU mantle component on plots of Pb isotopes versus (230Th/232Th) and U/Th.

FUNDING

This work was supported by NSF grant OPP-0126269 to K.W.W.S. and grants OPP-0125744 and ANT-0838817 to P.R.K., and French Institut National des Sciences de l’Univers and Agence Nationale de la Recherche grants (M&Ms) to J.B.-T. Fieldwork at Erebus volcano and operation of the Mount Erebus Volcano Observatory (erebus.nmt.edu) was supported by the Office of Polar Programs, NSF.

ACKNOWLEDGEMENTS

Field work has been greatly facilitated by NSF-directed civilian contractors Raytheon Polar Services Company. Special thanks go to the helicopter crews from PHI and Helicopters New Zealand. Reviews by Georg Zellmer, John Hora and Ken Rubin, and editorial handling by Simon Turner are gratefully acknowledged.

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APPENDIX A

ANALYTICAL METHODS FOR MEASUREMENT OF 227Ac AND 228Th

227Ac

The dissolved rock samples were spiked using ∼0·5 d.p.m. g–1 of rock of NIST certified 229Th/225Ac tracer in equilibrium. Samples were re-heated to allow for better tracer equilibration. Actinides from the sample solution were pre-concentrated via lead sulfate co-precipitation (Martin et al., 1995), which was then dissolved in 20 ml of 4 M HCl. This solution was loaded on a commercially available extraction chromatographic column containing N,N,N',N'-tetra-n-octyldiglycolamide [DGA column manufactured by Eichrom (Horwitz et al., 2005)]. In 4 M HCl, the actinides are well retained on the DGA column whereas common ions and alkaline earth elements have no affinity and pass through the column without retention. Several column volumes of 3 M HNO3 then rinse the leftover alkaline earth elements and iron from the column. In the next step we eluted actinium with 2 M HCl, which recovers Ac but leaves Th and other actinides retained on the DGA. The actinium fraction was prepared for alpha spectrometric measurement via cerium fluoride micro-precipitation (Dulaiova et al., 2001). The samples were counted immediately for 225Ac, from which we determined chemical recoveries, and then again in about 60 days after the ingrowth of 227Ac daughters 227Th + 223Ra. The 225Ac chemical recoveries were 67 ± 11% (n = 23). The minimum detectable activities (MDA) of 227Ac using alpha spectrometry were 0·001 d.p.m. per sample using an acquisition time of 1 week.

Table A1

Individual in situ Ba and Th concentrations in anorthoclase megacrysts (AN) and their melt inclusions (MI) and for selected pyroxene (PX) and adhering glass spots (GL) measured by SIMS (IMS-3f Cameca) at WHOI

Name Type Location Ba (ppm) Th (ppm) 
ER84-01      
ER84-01 MI core 453·3 37·37 
ER84-01 MI core 572·4 31·02 
ER84-01 MI core 654·6 34·68 
ER84-01 14 MI rim 402·2 26·81 
ER84-01 18a MI core 471·4 34·26 
ER84-01 18b MI core 434·7 32·51 
ER84-01 22a MI core 442·1 25·19 
Av.    490 31·7 
1σ SD    90 4·4 
ER84-01 11 GL   37·16 
ER84-01 16a GL  520·4 27·95 
ER84-01 16b GL  478·0 25·84 
ER84-01 AN core 2583·2 0·092 
ER84-01 AN core 2239·3  
ER84-01 AN core 2399·7 0·110 
ER84-01 AN core/rim 2255·8 0·092 
ER84-01 AN core/rim 2374·7  
ER84-01 12 AN core 2742·0  
ER84-01 15a AN rim 2234·8  
ER84-01 15b AN rim 2093·0 0·141 
ER84-01 19a AN core 1935·7 0·100 
ER84-01 19b AN core 1775·4 0·131 
ER84-01 20a AN core 1835·8 0·141 
ER84-01 20b AN core  0·012 
ER84-01 23b AN rim 1952·9 0·092 
ER84-01 26b AN rim 1984·5 0·093 
Av.    2185 0·10 
1σ SD    291 0·04 
ER84-01 PX core 2·9 0·194 
ER84-01 17–1 PX core 2·5 0·097 
ER84-01 17–2 PX core 1·4 0·012 
ER84-04      
ER84-04 MI rim 552·5 31·27 
ER84-04 MI core/rim 510·8 24·72 
ER84-04 MI core 588·5  
ER84-04 MI core 517·6 30·45 
ER84-04 10 MI core 546·9 25·70 
ER84-04 12 MI rim 576·0  
Av.    549 28·0 
1σ SD    31 3·3 
ER84-04 GL  499·7 24·44 
ER84-04 14 GL  529·7 22·08 
ER84-04 AN core/rim 2288·2 0·02 
ER84-04 AN rim 2250·1 0·00 
ER84-04 AN core 2327·7 0·03 
ER84-04 AN core 2213·2 0·05 
ER84-04 13 AN rim 2313·7 0·03 
Av.    2279 0·026 
1σ SD    47 0·021 
ER88-01      
ER88-01 MI core 413·7 26·11 
ER88-01 MI core 409·4 26·08 
ER88-01 MI core 413·2 26·98 
ER88-01 MI core 424·1 25·09 
ER88-01 MI core 399·0 23·47 
ER88-01 10 MI core 419·8 25·39 
ER88-01 12 MI core 398·7 22·63 
ER88-01 16 MI core/rim 392·2 22·72 
ER88-01 17 MI rim 388·9 21·56 
ER88-01 18 MI rim 416·1 22·12 
ER88-01 20 MI rim 450·4  
ER88-01 27 MI core/rim 389·2 29·17 
Av.    409·6 24·66 
1σ SD    17·7 2·36 
ER88-01 21 GL  389·3 22·24 
ER88-01 24 GL  386·1 21·99 
ER88-01 AN core 1883 0·192 
ER88-01 AN core 2051 0·012 
ER88-01 AN core 2031 0·215 
ER88-01 9b AN core 1829 0·012 
ER88-01 11 AN core 1781 0·153 
ER88-01 13 AN core 1829 0·109 
ER88-01 14b AN core/rim 1846 0·282 
ER88-01 15 AN core/rim 1881  
ER88-01 19 AN rim 1893 0·125 
ER88-01 23 AN rim 1793 0·114 
ER88-01 25 AN rim 1942 0·123 
ER88-01 26 AN rim 1769 0·082 
ER88-01 28 AN rim 1840  
ER88-01 29 AN rim  0·156 
ER88-01 30 AN core/rim  0·150 
Av.    1840 0·13 
1σ SD    54 0·07 
Name Type Location Ba (ppm) Th (ppm) 
ER84-01      
ER84-01 MI core 453·3 37·37 
ER84-01 MI core 572·4 31·02 
ER84-01 MI core 654·6 34·68 
ER84-01 14 MI rim 402·2 26·81 
ER84-01 18a MI core 471·4 34·26 
ER84-01 18b MI core 434·7 32·51 
ER84-01 22a MI core 442·1 25·19 
Av.    490 31·7 
1σ SD    90 4·4 
ER84-01 11 GL   37·16 
ER84-01 16a GL  520·4 27·95 
ER84-01 16b GL  478·0 25·84 
ER84-01 AN core 2583·2 0·092 
ER84-01 AN core 2239·3  
ER84-01 AN core 2399·7 0·110 
ER84-01 AN core/rim 2255·8 0·092 
ER84-01 AN core/rim 2374·7  
ER84-01 12 AN core 2742·0  
ER84-01 15a AN rim 2234·8  
ER84-01 15b AN rim 2093·0 0·141 
ER84-01 19a AN core 1935·7 0·100 
ER84-01 19b AN core 1775·4 0·131 
ER84-01 20a AN core 1835·8 0·141 
ER84-01 20b AN core  0·012 
ER84-01 23b AN rim 1952·9 0·092 
ER84-01 26b AN rim 1984·5 0·093 
Av.    2185 0·10 
1σ SD    291 0·04 
ER84-01 PX core 2·9 0·194 
ER84-01 17–1 PX core 2·5 0·097 
ER84-01 17–2 PX core 1·4 0·012 
ER84-04      
ER84-04 MI rim 552·5 31·27 
ER84-04 MI core/rim 510·8 24·72 
ER84-04 MI core 588·5  
ER84-04 MI core 517·6 30·45 
ER84-04 10 MI core 546·9 25·70 
ER84-04 12 MI rim 576·0  
Av.    549 28·0 
1σ SD    31 3·3 
ER84-04 GL  499·7 24·44 
ER84-04 14 GL  529·7 22·08 
ER84-04 AN core/rim 2288·2 0·02 
ER84-04 AN rim 2250·1 0·00 
ER84-04 AN core 2327·7 0·03 
ER84-04 AN core 2213·2 0·05 
ER84-04 13 AN rim 2313·7 0·03 
Av.    2279 0·026 
1σ SD    47 0·021 
ER88-01      
ER88-01 MI core 413·7 26·11 
ER88-01 MI core 409·4 26·08 
ER88-01 MI core 413·2 26·98 
ER88-01 MI core 424·1 25·09 
ER88-01 MI core 399·0 23·47 
ER88-01 10 MI core 419·8 25·39 
ER88-01 12 MI core 398·7 22·63 
ER88-01 16 MI core/rim 392·2 22·72 
ER88-01 17 MI rim 388·9 21·56 
ER88-01 18 MI rim 416·1 22·12 
ER88-01 20 MI rim 450·4  
ER88-01 27 MI core/rim 389·2 29·17 
Av.    409·6 24·66 
1σ SD    17·7 2·36 
ER88-01 21 GL  389·3 22·24 
ER88-01 24 GL  386·1 21·99 
ER88-01 AN core 1883 0·192 
ER88-01 AN core 2051 0·012 
ER88-01 AN core 2031 0·215 
ER88-01 9b AN core 1829 0·012 
ER88-01 11 AN core 1781 0·153 
ER88-01 13 AN core 1829 0·109 
ER88-01 14b AN core/rim 1846 0·282 
ER88-01 15 AN core/rim 1881  
ER88-01 19 AN rim 1893 0·125 
ER88-01 23 AN rim 1793 0·114 
ER88-01 25 AN rim 1942 0·123 
ER88-01 26 AN rim 1769 0·082 
ER88-01 28 AN rim 1840  
ER88-01 29 AN rim  0·156 
ER88-01 30 AN core/rim  0·150 
Av.    1840 0·13 
1σ SD    54 0·07 
228Th

The (228Th/232Th) value for an anorthoclase erupted in December 2005 was determined by alpha spectrometry in April 2008. Five grams of sample were digested in a mixture of concentrated HF and HNO3. After evaporation to dryness, the sample was completely dissolved in about 125 ml of 1 N HCl with 5 ml of saturated boric acid. A few drops of an FeCl3 solution were added to this solution, followed by concentrated NH3OH, until iron oxy-hydroxides ceased precipitating. The Th-bearing iron precipitate was then separated from the supernate by centrifuging. The supernate was discarded and the precipitate was dissolved in 2 ml concentrated HNO3, dried, re-dissolved in 7·5 N HNO3 and run through an anion exchange column charged with nitric acid. The Th was washed off with 6 N HCl and dried. The Th was then dissolved in 5 ml of a 2 M ammonium chloride solution adjusted to pH = 2·1 using HCl plus 1 ml of a saturated ammonium oxalate solution (pH = 2·0), and electroplated on a stainless steel disk using 2 amps of current over 20 min. The plated sample was alpha-counted for 12 days. Data reduction included subtracting background counts, tails of higher energy peaks into lower energy peaks, and the portion of the 224Ra peak (a short-lived daughter of 228Th) that is within the energy range counted for 228Th. The total adjustments to the raw counts were 2–4% for 230Th and 232Th and 17% for 228Th.

APPENDIX B

MODEL CALCULATIONS FOR CONTINUOUSLY RECHARGED OPEN-SYSTEM MODEL

Our open-system model algorithm is the following. It should be noted that recharge (step 5) is omitted for closed-system crystallization modeling.

For each iteration:

  1. one increment of anorthoclase by volume is added onto the pre-existing crystals using a mass-balance equation, and the composition of the entire crystal mass is recalculated. The degree of fractionation for each step is chosen so the model approximates measured compositions in 10 000 steps.

  2. The composition of the melt is then recalculated for the growth of the entire fractionating assemblage.

  3. 210Po is degassed from the magma 1·3% per day to best fit the model to the data. 222Rn is presumed to be in equilibrium with 226Ra based on the equilibrium (210Pb/226Ra) measured in both the anorthoclase crystals and glass.

  4. The activities of all radioactive nuclides in both the bulk anorthoclase and remaining melt are readjusted for radioactive ingrowth and decay. These calculations include any intermediate daughters of the decay series (e.g. 228Ra, where the measured nuclide is actually 228Th).

  5. New magma is added proportionally to the amount of crystallization based on a constant rate of recharge/crystallization [the ‘r’ value of DePaolo (1981)].

This is iterated for 10 000 steps.