Rapid timescales of differentiation and evidence for crustal contamination at intra-oceanic arcs: Geochemical and U–Th–Ra–Sr–Nd isotopic constraints from Lopevi Volcano, Vanuatu, SW Pacific

Earth and Planetary Science Letters 273 (2008) 184–194

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Rapid timescales of differentiation and evidence for crustal contamination at intra-oceanic arcs: Geochemical and U–Th–Ra–Sr–Nd isotopic constraints from Lopevi Volcano, Vanuatu, SW Pacific Heather K. Handley a,⁎, Simon P. Turner a, Ian E.M. Smith b, Robert B. Stewart c, Shane J. Cronin c a b c

GEMOC, Department of Earth and Planetary Sciences, Macquarie University, Sydney, NSW 2109, Australia Department of Geology, University of Auckland, Auckland, New Zealand Institute of Natural Resources, Massey University, Palmerston North, New Zealand

A R T I C L E

I N F O

Article history: Received 4 April 2008 Received in revised form 19 June 2008 Accepted 20 June 2008 Available online 2 July 2008 Editor: R.W. Carlson Keywords: Uranium series isotopes assimilation fractional crystallisation intra-oceanic arc geochemistry Vanuatu

A B S T R A C T The extent and geochemical impact of crustal contamination during magmatic evolution in intra-oceanic subduction zone settings is often assumed to be of minimal significance but remains poorly constrained. Acquiring such information is crucial if meaningful timescales of magma generation and crustal residence beneath volcanoes are to be determined. High-MgO basalts and differentiates erupted over the last 100 years at Lopevi volcano display a strong negative correlation between 87Sr/86Sr isotope ratio and indices of differentiation (e.g. SiO2) even though these lavas have a restricted Sr–Nd isotopic range compared to other Vanuatu arc lavas. This trend provides compelling evidence for the interaction of rising mafic magmas with ‘primitive’ sub-arc crust and so this volcano affords an excellent opportunity to investigate and ascertain timescales of crustal interaction using U-series data. Quantitative geochemical modelling of whole-rock trace element ratios, 87Sr/86Sr isotope compositions and U-series data shows that assimilation of a relatively small-degree partial melt of N 380 Ka mafic oceanic crust (similar to Pacific- or Indian-MORB in 87Sr/86Sr isotopic composition) during fractional crystallisation of magma exerts major control on the (230Th/232Th) and (226Ra/230Th) activity ratios of the lavas. The incorporation of higher (230Th/232Th) and lower (226Ra/230Th) assimilated material drives compositions closer to secular equilibrium than simple closed-system differentiation, reducing calculated apparent timescales of closed-system differentiation from Th isotope composition (104–105) by orders of magnitude. Modelling suggests that assimilation occurs extremely rapidly at Lopevi with timescales for magma generation, differentiation and eruption on the order of 102 years. © 2008 Elsevier B.V. All rights reserved.

1. Introduction Crystallisation during magmatic evolution is linked to the eruptive behaviour (e.g. style and frequency) of volcanoes because of the effect of crystallisation on volatile build-up (Tait et al., 1989). Therefore, obtaining information on the timescales of magmatic differentiation is essential for comprehensive assessment of volcanic hazards. Despite a growing number of magmatic timescale studies, there is still significant debate over whether differentiation from basaltic to andesitic magma compositions occurs relatively rapidly, within 102 or 103 years (e.g. Turner et al., 2000a,b, 2003; Blake and Rogers, 2005; Yokoyama et al., 2006; Garrison et al., 2006), or over considerably longer timescales 104-105 years (e.g. Turner et al., 1996; Heath et al., 1998; Asmerom et al., 2005; Touboul et al., 2007) in subduction zone settings. Disequilibria studies of U, Th and Ra isotopes of the Uraniumseries (U-series) chain provide extremely important tools for evaluat⁎ Corresponding author. E-mail address: [email protected] (H.K. Handley). 0012-821X/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2008.06.032

ing timescales as they provide unique insight into the timing and rates of magmatic processes operating within the last 100 to 380,000 years before present (see Condomines et al. (2003), Reid (2003) and Peate and Hawkesworth. (2005) for recent reviews). The timescales obtained are usually calculated assuming closed-system fractional crystallisation. However, even though this process is considered to be the major cause of magmatic differentiation (Bowen, 1928), a substantial number of volcanic systems provide geochemical, isotopic, petrographic or mineralogical evidence for the role of open-system processes such as crustal contamination (e.g. Briqueu et al., 1986; Davidson, 1987; Hildreth and Moorbath 1988; Vroon et al., 1993; Reagan et al., 2003; Price et al., 2007) and magma mixing or mingling (e.g. Nakamura, 1995; Tepley et al., 2000; Murphy et al., 2000; Defant et al., 2001; Reubi et al., 2003) during magmatic evolution. Such opensystem processes introduce complications for the correct interpretation of U-series data and determination of meaningful timescales (Zellmer et al., 2005; Yokoyama et al., 2006). Intra-oceanic arcs have often been selected for petrogenesis studies because the modification of mantle-derived magma through

H.K. Handley et al. / Earth and Planetary Science Letters 273 (2008) 184–194

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Fig. 1. (a) and (b) Tectonic setting and location of Lopevi volcano within the Vanuatu arc, SW Pacific. (c) Topographic sketch map of Lopevi showing the distribution of recent lava flow deposits.

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interaction with crust is assumed to be limited (e.g. Woodhead, 1989; Turner et al., 1997: Taylor and Nesbitt, 1998), but the degree of interaction is poorly constrained and may be difficult to detect due to the potentially similar compositional nature of both the arc crust and ascending magma using traditional discriminating tools (87Sr/86Sr, δ18O and trace element ratios). In contrast, U-series isotopes in arc magmas may be much more sensitive to assimilation of arc crust during magmatic evolution if the crust is old (i.e. in U-series secular equilibrium, N380 Ka) even if the assimilant is otherwise (e.g. geochemically and in terms of radiogenic isotopes) similar to the magma. Assimilation of this type of material is likely to lead to an overestimate of the timescales of differentiation (Dosseto et al., in press). However, despite the importance of AFC processes on the interpretation of time scale information from U-series data, only a few volcanic studies have tried to quantitatively assess the effect of these processes (e.g. Yokoyama et al., 2006; Price et al., 2007) and much focus has been placed on along arc studies (particularly true for the Vanuatu arc). Detailed, single-volcano-scale studies are therefore required to fully characterise magmatic differentiation, prior to constraining timescales. This study is the first detailed comprehensive geochemical and Useries study of an individual volcano in the Vanuatu arc. Whole-rock major and trace element abundances, Sr and Nd isotopic ratios and U, Th and Ra isotopic compositions have been determined for recent, temporally well-defined sequences of lava flow deposits at Lopevi volcano. The data are used to constrain the processes and timescales of magma genesis and evolution beneath the volcano and show strong evidence for significant magma-crust interaction which complicates the interpretation of U-series disequilibria. Lopevi volcano is ideal for this study because 1) it has a relatively simple crustal setting in a

youthful subduction zone, 2) young and temporally well defined sequence of samples are available, and 3) the eruption of relatively rare (in island-arcs) high MgO lavas will allow characterisation of the earliest stages of magmatic differentiation in subduction-zone volcanic systems. 2. Geological setting The presently active Vanuatu volcanic arc has developed above the eastward dipping subduction zone associated with the convergence of the Australian plate and the North Fiji Basin (Fig. 1a). The precursor to the present arc is believed to have dipped westwards and formed part of the now inactive Vitiaz Arc (remnants of this activity are represented by the islands of Espiritu Santo and Malakula, Fig. 1b). The reversal in arc polarity probably began in mid to late Miocene time (Meffre and Crawford 2001, Schellert et al., 2006) in an event likely to be related to the collision of the Ontong Java Plateau with the Solomon Islands segment of the trench to the north (Meffre and Crawford, 2001) and the Melanesian Platform (Schellert et al., 2006). The subaerial part of the arc is a 1200 km long chain of volcanoes extending from the Torres Islands in the north to Hunter Island (Fig. 1a and b). The chain is divisible into northern (Torres Islands-Aoba), central (Ambrym-Efate) and southern segments (Erromango-Anatom) (Fig. 1b). The tectonics of the Vanuatu arc are currently influenced by the ongoing collision of the now extinct D'Entrecasteaux subduction zone with the north-central part of the arc (Maillet et al., 1983; Greene and Collot, 1994; Taylor et al., 1995), a collision which started at about 2–3 Ma (Greene et al., 1988, 1994). The Vanuatu subduction system shows considerable variation in isotope, major and trace element composition along the arc. To the

Fig. 2. SiO2 variation with (a) MgO, (b) Al2O3, (c) Ba/Th and (d) 87Sr/86Sr for Lopevi volcanic rocks. Filled circles distinguish the subset of samples used for isotopic study. Open circles from unpublished data of R.B. Stewart. Ba/Th N-MORB data from Sun and McDonough, 1989. 87Sr/86Sr 2σ error is shown in lower left-hand corner of Fig. d. Model fractionation curves are shown in Figs. a and b for removal of the mineral assemblages suggested by least squares modelling from PV17 to PV38 (Stage 1; model 1 in supplementary data Table 5) and from PV38 to PV112 (Stage 2; model 5 in supplementary data Table 5). Mineral data used in modelling are given in supplementary data Tables 1–4. Fractionation curves are labelled with the percentage of crystallisation.

Table 1 Major element, trace element and Sr–Nd isotope whole rock data of Lopevi volcanic rock PV05

PV13

PV15

PV17

PV18

PV37

PV38

PV47

PV49

PV57

PV100

PV109

PV110

PV111

PV112

PV138

Eruption year

2003

2003

1964

1964

1964

1939

1939

1964

1964

2002

2003

pre 1900?

2001

2001

2000

Debris flow

Latitude

16° 29′ 53.2"S

16°29′ 50.9"S

16°29′ 50.9"S

16°29′ 50.9"S

16°29′ 50.9"S

16°29′ 44.6"S

16°29′ 57.5"S

16°30′ 22.4"S

16°30′ 43.9"S

16°31′ 38.05"S

16°29′ 42.1"S

16°30′ 05.6"S

16°30′26.9"S

16°30′31.2"S

16°30′46.3"S

16°29′05.6"S

Longitude

168°18′ 46.5"E

168°19′ 06.7"E

168°19′ 06.7"Ea

168°19′ 06.7"Ea

168°19′ 06.7"Ea

168°18′ 31.0"E

168°18′ 24.2"E

168°18′ 27.6"E

168°18′ 44.7"E

168°20′ 02.1"E

168°18′ 57.6"E

168°19′ 50.6"E

168°19'56.8"E 168°19'49.0"E 168°19'26.6"E 168°19'56.5"E

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 H2O LOI Total Cs Ba Rb Sr Pb Th U Zr Nb Hf Ta Y Sc V Cr Ni Cu Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 87 Sr/86Sr

51.54 0.74 19.93 9.26 0.15 4.03 11.09 2.61 0.78 0.17 0 −0.38 99.92 0.32 147 12.3 370 2.74 0.83 0.34 68 1.54 1.63 0.19 19.1 30.35 325 25 14 128 68 6.48 13.60 1.94 9.65 2.66 0.87 3.04 0.51 3.46 0.72 2.18 0.34 2.23 0.34 0.703980 ± 21 0.512967 ± 10

51.8 0.75 19.85 9.32 0.15 4.15 11.11 2.65 0.78 0.17 0 −0.36 100.37 0.33 149 12.5 369 2.79 0.82 0.32 70 1.58 1.76 0.15 19.4 30.98 326 30 16 130 68 6.40 13.71 1.94 9.45 2.72 0.89 3.14 0.50 3.49 0.74 2.24 0.34 2.23 0.33 0.704003 ± 11 0.512966 ± 11

49.95 0.67 16.12 10.53 0.17 8.09 12.3 2.02 0.65 0.15 0 − 0.4 100.25 0.30 123 10.3 352 2.38 0.61 0.24 50 1.06 1.19 0.11 15.3 45.44 346 248 64 114 67 7.16 10.73 1.57 7.76 2.12 0.76 2.56 0.42 2.85 0.60 1.75 0.26 1.78 0.27 0.704086 ± 6 0.512932 ± 26

49.79 0.65 15.77 10.5 0.17 8.49 12.64 1.9 0.64 0.14 0 −0.33 100.36 0.30 123 10.0 353 2.37 0.52 0.24 47 0.95 1.09 0.08 14.7 47.8 346 268 68 125 66 4.74 10.28 1.47 7.70 2.19 0.70 2.54 0.39 2.85 0.58 1.75 0.25 1.69 0.25 0.704070 ± 10 0.512958 ± 14

50.98 0.71 19.46 9.56 0.16 4.79 11.56 2.46 0.72 0.16 0 −0.34 100.22 0.29 143 11.1 373 2.49 0.73 0.29 59 1.32 1.46 0.12 17.4 33.42 333 51 22 122 67 7.02 12.23 1.78 8.82 2.52 0.83 2.75 0.47 3.08 0.67 2.03 0.30 2.07 0.30

51.25 0.69 19.89 9.00 0.15 4.71 11.43 2.5 0.54 0.13 0.02 − 0.25 100.06 0.22 96 8.3 274 1.93 0.49 0.21 56 1.29 1.35 0.15 17.2 33.66 309 45 21 98 62 5.20 9.74 1.45 7.43 2.17 0.76 2.67 0.44 3.09 0.66 1.96 0.29 2.05 0.31 0.704029 ± 8 0.512978 ± 24

51.22 0.68 20.2 8.84 0.15 4.44 11.53 2.49 0.53 0.13 0.04 − 0.2 100.05 0.24 97 8.4 279 1.87 0.5 0.2 56 1.31 1.37 0.12 17.5 33.03 304 44 21 109 61 4.89 9.91 1.49 7.40 2.13 0.81 2.62 0.44 3.12 0.68 1.96 0.30 2.07 0.32 0.703944 ± 9

49.62 0.65 15.84 10.46 0.18 8.38 12.64 1.98 0.64 0.14 0.04 −0.38 100.19 0.30 122 9.9 356 2.38 0.58 0.24 46 1.03 1.16 0.11 14.6 47.07 348 250 63 123 65 4.63 10.19 1.50 7.63 2.25 0.77 2.39 0.39 2.67 0.56 1.70 0.25 1.78 0.25 0.704078 ± 5 0.512953 ± 12

50.68 0.66 17.34 9.9 0.18 7.14 11.55 1.98 0.57 0.15 0.01 −0.22 99.94 0.26 108 9.6 288 2.14 0.62 0.24 53 1.16 1.28 0.11 16.3 39.46 323 178 61 90 65 5.34 11.64 1.67 8.49 2.32 0.83 2.78 0.45 3.00 0.64 1.86 0.28 1.90 0.30 0.704016 ± 10 0.512953 ± 23

51.39 0.75 19.14 9.47 0.16 4.57 11.11 2.64 0.79 0.17 0.01 −0.34 99.86 0.31 148 12.5 362 2.83 0.84 0.33 68 1.54 1.7 0.14 19.2 32.43 328 57 21 142 67 6.38 13.78 1.95 9.67 2.67 0.87 3.20 0.51 3.48 0.74 2.23 0.32 2.31 0.35

51.48 0.75 19.54 9.31 0.15 4.31 11.09 2.68 0.78 0.17 0.02 −0.24 100.04 0.32 146 12.3 362 2.73 0.79 0.32 67 1.54 1.63 0.12 19.0 31.04 322 33 16 128 66 6.48 13.43 1.94 9.40 2.57 0.87 3.04 0.51 3.55 0.74 2.23 0.31 2.21 0.33 0.703993 ± 22 0.512967 ± 8

50.9 0.69 16.93 10.14 0.17 6.56 11.54 2.35 0.67 0.15 0.06 −0.17 99.99 0.25 142 10.6 312 2.46 0.76 0.29 60 1.3 1.5 0.11 17.8 41.84 342 112 36 118 68 5.71 12.34 1.75 8.73 2.42 0.82 2.89 0.49 3.30 0.70 2.03 0.30 2.07 0.31 0.704008 ± 6 0.512954 ± 19

52.25 0.79 18.75 9.46 0.16 4.34 10.54 2.8 0.87 0.19 0.03 −0.19 99.99 0.34 160 14.0 331 2.96 0.96 0.38 77 1.8 1.97 0.14 21.6 32.23 321 43 17 130 65 6.98 15.25 2.19 10.80 3.01 0.92 3.48 0.56 3.91 0.83 2.51 0.38 2.50 0.38 0.703992 ± 5

143

Nd/ Nd

144

0.512973 ± 9

52.66 0.81 18.74 9.46 0.16 4.19 10.27 2.89 0.91 0.2 0.02 −0.23 100.08 0.35 163 14.7 327 3.1 0.99 0.38 83 1.89 2.05 0.16 22.5 31.15 318 56 16 134 69 9.90 15.87 2.30 11.00 3.18 0.93 3.61 0.59 4.12 0.87 2.53 0.38 2.67 0.39 0.703995 ± 8

51.46 0.74 19.08 9.46 0.15 4.48 10.94 2.58 0.75 0.17 0.08 −0.14 99.75 0.32 146 12.2 352 2.82 0.81 0.33 66 1.54 1.72 0.11 18.7 32.87 330 50 22 107 66 7.85 12.98 1.87 8.99 2.48 0.85 2.86 0.46 3.35 0.71 2.13 0.31 2.32 0.34 0.703964 ± 16

0.512977 ± 7

0.512958 ± 9

187

Errors on isotopes are within-run 2SE on the final quoted significant figure. a Sample collected 50 m west of this location.

0.512959 ± 8

51.24 0.74 19.82 9.25 0.15 4.01 11.14 2.61 0.77 0.17 0.07 −0.22 99.75 0.31 146 12.3 368 2.79 0.76 0.3 68 1.55 1.68 0.26 18.7 30.26 324 30 14 124 68 6.17 13.45 1.94 9.45 2.59 0.83 3.00 0.49 3.33 0.69 2.12 0.31 2.18 0.33

H.K. Handley et al. / Earth and Planetary Science Letters 273 (2008) 184–194

Sample

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H.K. Handley et al. / Earth and Planetary Science Letters 273 (2008) 184–194

lesser extent K2O and Al2O3 are observed with increasing SiO2 content in Lopevi lavas (e.g. Fig. 2a and b). Trace element characteristics are typical of island arc volcanoes with enrichment of the more mobile LILE and LREE relative to HFSE and HREE (e.g. Ba/Th) compared to midocean-ridge basalt (MORB) and decreasing Ba/Th ratios with increasing differentiation at Lopevi (Fig. 2c). 3. Analytical techniques

Fig. 3. Spatial variation in 87Sr/86Sr of volcanic rocks from the Vanuatu arc, moving from north to south along the arc. Lopevi data from this study is shown as filled circles. Vanuatu data from Dupuy et al. (1982) and Peate et al. (1997).

north and south of the D'Entrecasteaux collision zone the magma compositions are arc tholeiites with MORB-like characteristics while in front of the collision the magmas are calc-alkaline (Laporte et al., 1998). The southernmost segment of the arc that includes the Hunter and Matthew volcanoes are boninite-like high magnesium andesites (Monzier et al., 1993, 1997). The strongest influence on arc compositional variation is the collision of the Eocene D'Entrecasteaux subduction ridge. Opposite the collision zone the calc-alkaline lavas are Mg-, K-, LILE- rich and Al- and Si-poor compared with those to the north and south. Isotopically, the collision zone lavas are higher in 87 Sr/86Sr, 207Pb/204Pb and 208Pb/204Pb and lower in 206Pb/204Pb and 143 Nd/144Nd (Monzier et al., 1997; Peate et al., 1997). This has been variously ascribed to contamination by a DUPAL-type anomalous mantle (Crawford et al., 1995; Monzier et al., 1997; Laporte et al., 1998) or Indian-MORB-like mantle (Peate et al., 1997). Lopevi is a small, steep-sided island about 7-km in diameter, rising to 1413 m above sea level. The small summit crater is nested within at least two remnant cinder cones, the youngest of which is breached to the NW. The island is bisected by a NW–SE trending rift (Fig. 1c), which is the locus of flank fissure eruptions. Lopevi is one of the most active volcanoes in Vanuatu with eruptions occurring at the summit crater and at flank fissures primarily on the NW and SE sides. Historical eruptions date back to the mid-19th century. Observed eruption types have ranged from moderately explosive sub-plinian events with associated block and ash flows and hot debris flows from the summit to lava flow effusion from both summit and rift. Occasionally, as in 2006, the summit crater is occupied by a lava lake. In recent times the volcano has been active in 1939, during 1960–65 (Williams and Curtis, 1964; Warden, 1967), and since 2000 with intermittent eruptions continuing through 2008. The composition of erupted magma varies from magnesian basalts (up to 9 wt.% MgO) to low-Mg basaltic andesites (Fig. 2a, Table 1). Furthermore, at least in the most recent activity these two magma types appear to be related in a magmatic cycle that was initiated by the eruption of the high-Mg composition. Our focus in this paper is on the magmatic cycle which began with eruptions of high-Mg lavas in 1960 and is being continued by the most recent eruptions (sample ages are given in Table 1). Vesicular basalts and basaltic andesites erupted at Lopevi are seriate or glomeroporphyritic in texture, containing phenocrysts of olivine, clinopyroxene, plagioclase and oxides set within a glassy or fine-grained crystalline groundmass. Large, subhedral olivine crystals (up to 5 mm) are abundant in the basalts while tabular plagioclase dominates the phenocryst populations in the more evolved samples. Representative mineral data used in geochemical modelling are given in Tables 1–4 in the supplementary material section (Appendix A). Decreases of MgO, CaO and Fe2O3 and increases in Na2O, TiO2 and to a

Full details of the analytical techniques are given in Appendix A of the supplementary material. An outline of isotope data quality is given here. A subset of Lopevi samples (filled circles Fig. 2) covering a range in SiO2 content was selected for Sr–Nd–U–Th–Ra isotope analysis. Sr and Nd isotopes were determined on a ThermoFinnigan Triton® TIMS at Macquarie University. Mass fractionation was corrected for by normalising Sr to 86Sr/88Sr = 0.1194 and Nd to 146Nd/144Nd = 0.7219. Analyses of BHVO-2 yielded 87Sr/86Sr = 0.703495 ± 8 (n = 2) and 143Nd/ 144 Nd = 0.512971 ± 6 (n = 2). U and Th concentrations and isotope ratios were measured on a Nu Instruments® MC-ICP-MS at Macquarie University. The mass bias and gain were determined from measurements of natural U (CRM 145) assuming 238U/235U = 137.88. The IC0 gain was calculated assuming a 234 U/238U ratio of 5.286 × 10− 5 for CRM 145 (Cheng et al., 2000). Accuracy (b0.3%) and precision (b0.1%) were assessed by regular analyses of the U010 and ThA solution standards. Analysis of rock standard TML-3 yielded the results: U = 10.17 ppm, Th = 29.18 ppm, (234U/238U) = 0.999 = 0.003, (230Th/232Th) = 1.076 = 0.006, (230Th/238U) = 1.005 ± 0.004 (n = 2) (see Sims et al., 2008 for a recent compilation of TML-3 standard data). 228 Ra/226Ra ratios were measured to a precision typically~0.5% on a ThermoFinnigan Triton® TIMS at Macquarie University. Accuracy was assessed via replicate analyses of TML-3 that yielded 226Ra = 3524 fg/g and (226Ra/230Th) = 1.007 ± 0.006 (n= 3). Lopevi (230Th/232Th) and (226Ra/230Th) ratios have not been recalculated for differences in eruption age because all samples were erupted in the last 100 years and so post-eruption radioactive decay is insignificant compared to the half life of 230Th (75,690 years) and 226Ra (1599 years). 4. Results 4.1. Radiogenic isotope data 87

Sr/86Sr and 143Nd/144Nd isotopic compositions of Lopevi lavas range from 0.703944–0.704086 and 0.512932–0.512978, respectively Table 2 U-series data for whole rock samples of Lopevi volcano Sample Eruption Th (TIMS) PV05 PV13 PV15 PV17 PV18 PV37 PV38 PV47 PV49 PV57 PV100 PV109 PV110 PV111 PV112 PV138

U (230Th/ (TIMS) 232Th)

(238U/ Th)

(234U/ U)

232

238

1.337 ± 2 1.326 ± 2 1.354 ± 2 1.398 ± 3 1.317 ± 2 1.250 ± 1 1.425 ± 4 1.250 ± 1 1.308 ± 6 1.250 ± 1 1.353 ± 2 1.326 ± 4

0.994 ± 2 0.995 ± 3 0.996 ± 2 0.984 ± 5 0.981 ± 6 0.997 ± 2 0.977 ± 5 0.997 ± 2 0.977 ± 7 0.994 ± 2 0.998 ± 3 0.960 ± 5

226

Ra

(226Ra/ Th)

230

year

(mg/g)

(mg/g)

2003 2003 1964 1964 1964 1939 1939 1964 1964 2002 2003 Pre 1900s? 2001 2001 2000 Debris flow

0.720 0.761 0.548 0.468 0.664 0.516 0.463 0.556 0.537 0.784 0.749 0.707

0.317 0.332 0.245 0.216 0.288 0.213 0.218 0.229 0.231 0.323 0.334 0.309

1.216 ± 4 1.234 ± 3 1.178 ± 4 1.188 ± 4 1.195 ± 4 1.231 ± 3 1.233 ± 7 1.187 ± 3 1.169 ± 4 1.227 ± 3 1.226 ± 6 1.199 ± 4

(fg/g)

0.847 0.748 0.924 0.739

0.380 0.325 0.416 0.325

1.244 ± 6 1.361 ± 2 0.995 ± 2 434.56 3.67 ± 2 1.229 ± 9 1.319 ± 6 0.994 ± 5 378.40 3.66 ± 7 1.250 ± 5 1.367 ± 2 0.994 ± 2 418.43 3.22 ± 7 1.218 ± 5 1.336 ± 2 0.997 ± 2 – –

415.36 404.87 382.30 332.59 392.61 204.84 204.85 376.43 273.31 418.59 365.23 381.86

4.22 ± 7 3.83 ± 4 5.27 ± 6 5.30 ± 8 4.40 ± 4 2.84 ± 5 3.18 ± 4 5.07 ± 12 3.86 ± 8 3.87 ± 6 3.54 ± 3 4.00 ± 2

H.K. Handley et al. / Earth and Planetary Science Letters 273 (2008) 184–194

189

Fig. 4. (a) (230Th/232Th) versus (238U/232Th) equiline diagram for Lopevi whole rock samples. Other Vanuatu arc data from Turner et al. (1999). (b) Enlarged view of diagram (a) with wt.% SiO2 content labelled for all data points. Age labels display calculated timescales for closed-system magmatic differentiation for each of the vertical data series used (dashed arrows). Calculation details are described in the text. 2σ error shown in bottom left-hand corner.

(Table 1) and display negative (87Sr/86Sr) and positive (143Nd/144Nd) correlations with SiO2 (e.g. Fig. 2d). Spatial variations in isotopic ratios (Sr, Nd and Pb) along the arc were described by Peate et al. (1997) who showed that Sr isotope ratios increase from the northern to central section of the arc and then systematically decrease to the southernmost section of the arc. This along-arc isotopic variation is illustrated in Fig. 3 where 87Sr/86Sr isotope data for Vanuatu lavas (Dupuy et al., 1982; Peate et al., 1997) are plotted against geographic distance along the arc from north to south. 87Sr/86Sr ratios of Lopevi volcanic rocks are slightly higher than those reported for the southern and northernmost sections of the Vanuatu arc but are characteristic of volcanoes similarly located in the central Vanuatu arc, above where the D'Entrecasteaux Ridge is being subducted and accreted.

increasing (230Th/232Th) in the samples, which is an important observation for the focus of later discussion. Lopevi lavas are characterised by relatively large 226Ra excesses with (226Ra/230Th) activity ratios between 2.84–5.27 (Fig. 5), especially when compared with (226Ra/230Th) determined for lavas from Tanna and Matthew in southern section of the Vanuatu arc (1.25–1.67) (Turner et al., 2001). However, they have typical values for subduction zone lavas, which usually range between 1 and 7 (226Ra/230Th) (Peate and Hawkesworth, 2005). A negative correlation is observed between (226Ra/230Th) and SiO2 for Lopevi lavas with the highest 226Ra excesses found in the most mafic samples.

4.2. Uranium series isotopes

5.1. Fractional crystallisation

Lopevi whole-rock U–Th–Ra isotope data are listed in Table 2. Most samples have (234U/238U) = 1.00 within analytical error, suggesting that the samples are fresh and unlikely to have been affected by subsolidus alteration or interaction with seawater (cf. Yokoyama et al., 2003). Lopevi lavas display a much more limited range in (238U/232Th) and (230Th/232Th) activity ratios (1.25–1.43 and 1.17–1.25, respectively) than other Vanuatu arc lavas ((238U/232Th) = 0.991–2.134; Turner et al., 1999) and lie to the right of the equiline in Fig. 4a, with 238U excesses of 25–43%. The enlarged section of the equiline diagram (Fig. 4b) shows that the base of the Lopevi array is relatively horizontal and formed by the least evolved (lowest SiO2) samples. SiO2 increases with

5.1.1. Major element constraints The inflections in major element data trends observed in Fig. 2a and b suggest that a change in either the mineral assemblage or proportion of phases crystallising from the magma (or both) occurs at ~51 wt.% SiO2. The downward inflection observed for Al2O3 versus SiO2 may be attributed to the onset of plagioclase fractionation while the plateau and slight increase in MgO content above ~ 51 wt.% SiO2 indicates that the magnitude of mafic-phase fractionation decreases at this point. These observations are investigated quantitatively by least squares modelling of the major element data using the XLFRAC programme of Stormer and Nicholls (1978). Summaries of least squares modelling results and modelling parameters are presented in the supplementary data section (Table 5). The results suggest that initial fractionation from the least evolved (lowest SiO2) rocks (PV17 and PV47) to an intermediate composition (~ 51 wt.% SiO2) is controlled by the removal of clinopyroxene and olivine (models 1–3, Stage 1 in Fig. 2a and b). From intermediate to evolved compositions the fractionating phase assemblage is dominated by plagioclase and significantly less mafic mineral fractionation (models 4 and 5, Stage 2 in Fig. 2a and b). All models yield excellent results, with Σr2 values less than 0.09 suggesting that major element variation can be explained by 2-stage fractional crystallisation with an overall maximum degree of crystallisation (C) of ~ 40%. Nevertheless, the radiogenic isotope variations require a more complex evolution (see below).

Fig. 5. (226Ra/230Th) against SiO2 for Lopevi lavas. The time required for (226Ra/230Th) to decay and return to secular equilibrium is illustrated on the right-hand side of the diagram, assuming that the degree of disequilibria in the sample with the highest (226Ra/230Th), 5.27 represents that at the onset of differentiation. 2σ error is shown in top right-hand corner.

5. Discussion: magmatic differentiation

5.1.1. Th–Ra constraints Major element fractionation modelling suggests that Lopevi lavas can, in principle, be related by the process of fractional crystallisation alone. In such cases, time information on closed-system crystallisation of magma can be extracted from the U–Th equiline diagram (Fig. 4). The excess 238U measured in the Lopevi lavas indicates that an event

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fractionating U and Th occurred less than 380 Ka ago; the fractionation has typically been attributed to addition of slab fluid to the mantle wedge (Gill, 1981; Allègre and Condomines, 1982). The horizontal nature of the base of the Lopevi array suggests that at time (t) = 0 all samples had similar Th isotope ratios at different U/Th ratios and also indicates that little, if any Th, was transported by the slab-fluid added to the mantle wedge. SiO2 content increases with increasing (230Th/232Th) activity ratios in Lopevi lavas (Fig. 4b). This might record the time over which the more evolved samples resided in the crust during differentiation, such that they evolved closer to the equiline due to the radioactive decay of 230Th. Therefore, if their geochemical magmatic evolution is controlled by closed system fractional crystallisation at Lopevi (but see below), variation in (230Th/232Th) of lavas possessing comparable (238U/232Th) (a vertical rise on the equiline diagram) will reflect the time taken for magmatic differentiation. This time can be calculated from the equation: t = 1/λ ln(1 − m), where λ is the 230Th decay constant, m = ((230Th/232Th)top − (230Th/232Th)base)/((238U/232Th)base − (230Th/232Th)base) and top and base refer to the top and bottom samples of a vertical array, respectively. The schematic arrows on Fig. 4b show 3 vertical trends for which timescales have been calculated which range from 104 to 105 (58,000– 120,000) years. These timescales are comparable to the long differentiation timescales reported for the Lesser Antilles intraoceanic arc by Turner et al. (1996) and Heath et al. (1998). One problem with the long differentiation timescale constraints inferred from the U–Th isotope system is the presence of 226Ra excesses in all Lopevi lavas (Fig. 5) which implies that crustal residence times of magmas at Lopevi are, in fact, significantly 226 b8000 years (5.t226 Ra/230Th) to decay 1/2 Ra). The time required for ( through closed system magmatic evolution can be calculated using the decay equation (1 − eλt) and is displayed on the left-hand side in Fig. 5. Here we assume that the magnitude of disequilibria in the least evolved lava ((226Ra/230Th) = 5.27) represents that at the onset of differentiation. Such calculations implicate short timescales for magmatic differentiation at Lopevi: b1000 years for the basalts and b2000 years for the basaltic andesites (Fig. 5). The most probable explanation for the discrepancy between the timescale information from the Th and Ra isotope data is that the correlation of both (230Th/232Th) and (226Ra/230Th) with SiO2 reflects assimilation of old crustal material (in secular equilibrium) during magmatic evolution. This would lead to extended apparent timescales for differentiation due to incorporation of higher (230Th/232Th) and lower (226Ra/230Th) material by shifting the samples closer to equilibrium than would otherwise occur if only the effects of the time taken for differentiation were important. Therefore, in order to determine meaningful timescales of magmatic differentiation at Lopevi we need to constrain the geochemical influence of opensystem processes (e.g. crustal contamination and magma mixing) on lava composition that may have modified Th–Ra systematics (cf. Yokoyama et al., 2006; Price et al., 2007). Fig. 6. (a) Ba/Th (b) (230Th/232Th) (c) (226Ra/230Th) versus 87Sr/86Sr diagrams showing combined assimilation and fractional crystallisation (AFC) and bulk mixing models described in the text. Note closed-system fractional crystallisation (FC) shown in (a) is unable to explain the data. See supplementary data Table 6 for end member compositions. 2A and 10A refer to a 2% and 10% melt, respectively, of the assimilant using residual mineral assemblage A. 2B and 10B refer to a 2% and 10% melt, respectively, of the assimilant using residual mineral assemblage B (see below). Tick marks on curves indicate the percentage of liquid remaining in 10% increments. FC and AFC calculations use the mineral proportions suggested by least squares major element modelling of PV17 to PV112 (supplementary data Table 5, model 6): 0.15 olivine; 0.52 cpx; 0.28 plag; 0.04 spinel. Mineral assemblages and proportions used for the contaminants: basalt (assemblage A): 0.4 plag, 0.4 opx, 0.2 cpx; metabasalt (assemblage B): 0.3 amphibole, 0.3 plag, 0.3 opx, 0.1 cpx. The rate ratio (r) of mass assimilated to mass crystallised is 0.25.

5.2. Assimilation and fractional crystallisation Long-lived radiogenic isotope analysis provides a powerful tool to identify and constrain crustal interaction with magma providing that the assimilant is isotopically distinct to the magma (e.g. Davidson, 1987; Thirlwall et al., 1996). Despite the restricted range in Lopevi Sr isotope ratios compared to the broad range found in Vanuatu arc rocks in general (Fig. 3), a more detailed investigation of the data shows that 87Sr/86Sr isotope ratios correlate negatively with indices of differentiation e.g. SiO2 (Fig. 2d) requiring that assimilation occurred concomitantly with fractional crystallisation (AFC; DePaolo, 1981) and may have also modified the Th–Ra isotope systematics of the lavas. Continental-type crustal rocks commonly have elevated 87Sr/86Sr isotope ratios compared to arc volcanic rocks and mantle-derived magmas, and assimilation of this material would result in a positive

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correlation with silica. Thus, the negative correlation observed here suggests that the assimilant was isotopically more primitive than the Lopevi lavas. The exact nature of the crust beneath Lopevi is unknown, but is likely to be composed of oceanic crust and old volcanic material. 87 Sr/86Sr isotope ratios of Pacific- and Indian-MORB (representing the oceanic basement) are appreciably lower than those in Vanuatu volcanic rocks and therefore we consider such oceanic crust as a suitable assimilant. The low 87Sr/86Sr ratios required by the assimilant suggest that it is unlikely to have been significantly affected by hydrothermal alteration (cf. Nicholson et al., 1991; Staudigel et al., 1995), which is further supported by a lack of (234U/238U) disequilibrium in the samples. We now explore AFC modelling for Lopevi trace element ratios, Th and Ra isotope ratios against Sr isotope ratios to refine the models presented above. 5.2.1. AFC model 1: bulk assimilation First we consider the simplest case for AFC modelling of bulk assimilation of crust using the equation: Cm =C0m F− z +(r/r − 1) Ca (1 −F− z), where Cm and Ca are concentrations in the melt and the assimilant respectively, F is the fraction of melt remaining, r is the ratio of assimilation rate to crystallisation rate and z = (r +D − 1)/(r − 1), where D is the bulk partition coefficient (DePaolo, 1981). The fractionating mineral assemblage and phase proportions were taken from the results of least squares major element modelling from the least to most evolved Lopevi lavas (supplementary data Table 5, model 6). Bulk D values are given in supplementary data Table 6 and partition coefficients were taken from Berlo et al. (2004), Blundy and Wood (2003), Beattie (1993) and the Geochemical Earth Reference Model database (GERM). A r value of 0.25 was used. Sample PV17, with the lowest MgO content and highest (226Ra/ 230 Th) was taken as the starting end member. Compositional parameters used for the assimilant are given in supplementary data Table 6 and its 87 Sr/86Sr isotope ratio was inferred to be 0.7025, since this represents the region of overlap between Pacific and Indian MORB (Price et al., 1986; Ito et al., 1987; Rehkämper and Hoffmann, 1997; references within Chauvel and Blichert-Toft, 2001) and is therefore a typical value for oceanic crust beneath Lopevi. The (230Th/232Th) ratio of the oceanic basement is also unknown, but we assume that it is older than 380 Ka and will therefore plot along the equiline. Considering the range in (238U/232Th) of the Lopevi samples, this constrains the (230Th/232Th) of the assimilant to lie between 1.25–1.42 (see Fig. 4a), and we have selected an intermediate value of 1.32 to characterise this crust. The initial Ra concentration and (226Ra/230Th) ratio of the contaminant was calculated assuming that it is in 238U–230Th–226Ra secular equilibrium. The results are plotted in Fig. 6a–c and show that bulk AFC models starting from the least evolved Lopevi lava cannot produce curves that fit the data. Varying the r value from 0.1 to 0.9 does not improve the situation. However, the shape of the bulk AFC curve in Ba/Th–87Sr/86Sr space (Fig. 6a) is very similar to the shape of Lopevi data array. If we allow for heterogeneity in the basaltic starting composition and use PV15 as the starting magma, the bulk AFC model shows a good fit to the data. Nonetheless, this model fails to account for the data array in Th and Ra isotope versus Sr isotope space (Fig. 6b and c). For comparison, simple bulk mixing models between the least evolved Lopevi lava and the oceanic crustal end-member are shown in Fig. 6a–c, but these also are unable to fit the Lopevi data array. Furthermore, inflections of data trends in the major element variation diagrams (Fig. 2a and b) are inconsistent with variation in lava composition being controlled by simple binary mixing processes. 5.2.2. AFC model 2: assimilation of a crustal partial melt The bulk assimilation model above assumes that sub-volcanic lithosphere in contact with rising mafic magma will melt 100%. It may be more plausible that the crustal material encountered will only partially melt due to insufficient heating. Berlo et al. (2004) suggest that partial melting of mafic lower crust will only produce small disequilibria between 238U, 230Th and 226Ra. Therefore, we may

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not expect to see significantly different results between AFC trends where the assimilant is incorporated either as a bulk component or partial melt. However, the elemental concentrations and therefore relevant elemental ratios (e.g. Ba/Th) may be significantly altered as a result of melting, depending on the compatibility of elements in the residual mineral assemblage considered, and therefore produce distinctive AFC models. In order to calculate what the composition of a partial melt of the assimilant would be, a number of assumptions are required: the degree of melting, residual mineralogy, melt/solid partition coefficients and original starting composition. To determine Sr, Ba, Th and Ra element concentrations and (226Ra/230Th) activity ratios expected in partial melts of the assimilant we used a simple batch melting model for variable degrees of partial melting (2–10%) of the same assimilant end member used above. Solid/melt partition coefficients of Berlo et al. (2004), Blundy and Wood (2003), Beattie (1993) and the GERM database were used. Melt compositions were calculated for two different residual mineral assemblages: (A) plagioclase, orthopyroxene and clinopyroxene (tholeiitic basalt) and (B) amphibole, plagioclase, orthopyroxene and minor clinopyroxene (a stable metabasalt/amphibolite phase assemblage at relatively low pressure (0.8–1.0 GPa) and temperature (900–1000 °C) (Berlo et al. (2004) and references within). 87Sr/86Sr isotope ratios and (230Th/232Th) of the assimilant are assumed fixed. Partial melt end member compositions are given in supplementary data Table 6. An interesting observation is that (2–10%) partial melting of MORBtype oceanic crust (in secular equilibrium) containing amphibole, plagioclase, orthopyroxene and clinopyroxene has little effect on the (226Ra/230Th) of the created crustal melt due to the similar bulkdistribution coefficients of Ra and Th (supplementary data Table 6). Partial melting of this composition only produces minor 226Ra excess in the melts and thus confirms the findings of Berlo et al. (2004). However, the amphibole-free mineral assemblage yields slight 226Ra deficiencies (0.60-0.85 (226Ra/230Th)) in the partial melt produced due to the retention of Ra in plagioclase as this mineral phase has a higher affinity for Ra, than for Th, and is modally more abundant in the basaltic assemblage considered (see Fig. 6 caption and supplementary data Table 6 for details). Fig. 6a shows that AFC models in which the assimilated component is a partial melt can produce much better fits to the Lopevi data than the bulk AFC models. Most of the samples lie between the two curves for 2 and 10% partial melting of the assimilant, containing plagioclase, orthopyroxene and clinopyroxene in the residual mineralogy (solid lines labelled 2A and 10A joined by tie-lines of equal percentage of crystallisation). These models also fit well with the overall degree of crystallisation (~ 40%) suggested by fractional crystallisation modelling of major element data (supplementary data Table 5, model 6). Assimilation of a 10% partial melt of metabasalt (curve 10B) during fractional crystallisation also fits the Lopevi array, indicating that the solution is non-specific, with a range of suitable contaminant partial melt compositions capable of explaining the data. Similarly, good results are observed in Fig. 6b for partial melt AFC models considering Th-Sr isotope variations, confirming that it is highly likely that crustal contamination has modified (230Th/232Th) ratios of the lavas and that they do not provide (closed-system) crystallisation timescale information. When considering the same partial melt AFC modelling parameters in (226Ra/230Th)-Sr isotope space (Fig. 6c), we see a poorer fit of the 2A and 10A model curves to the data. This suggests that AFC processes may operate over a timescale similar to that which can be recorded by Ra-Th isotope disequilibria, and therefore the degree of displacement of the Lopevi data below the model curves may record the time involved for the combined AFC process. Due to the range in acceptable model solutions identified above, which are dependent on degree of partial melting and residual mineralogy, we believe it is impractical to calculate a specific timescale for AFC at Lopevi, but if the degree of displacement from the model curves in Fig. 6c do in fact represent time, this suggests no more than 1000 years for magma

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generation, differentiation and eruption at Lopevi (as exemplified by the 1 Ka aging curve of model 2A in Fig. 6c). Lopevi volcanic rocks do not display a progressive temporal transition from high- to low-MgO compositions; the highest MgO contents are found in lavas erupted during the 1960s (Table 1). This may indicate that injection of more mafic magma in to the plumbing system occurred within the 25 year interval since the preceding, more silicic eruptions in 1939. Numerical simulations indicate that fresh injections of magma can increase closed-system differentiation timescales (Dosseto et al., in press), producing the opposite effect on U-series disequilibria than that of AFC. However, as recharge periodicity approaches 500 years or less and/or recharge volume approaches ≥20% of the resident magma volume, both U–Th and Ra– Th disequilibria become buffered (Hughes and Hawkesworth, 1999). Therefore, the wide range of U-series disequilibria observed at Lopevi are inferred to indicate that magma recharge is not as significant as AFC processes. 6. Petrogenetic model In Fig. 7 we show a schematic model for magma petrogenesis at Lopevi volcano, highlighting the timescale constraints obtained from the new U-series data. Stage 1: Fluid addition from the subducting slab to the mantle wedge significantly b380 Ka ago produces fractionation between U and Th and creates 238U excesses in the primary basaltic magmas. Assuming that Ra is also preferentially fractionated from Th into the slab-fluid component, the 226Ra excesses indicate a much more recent time constraint for fluid addition to the mantle wedge of b8000 years. Stage 2: Early magmatic differentiation is dominated by fractionation of olivine and clinopyroxene during simultaneous assimilation of partially melted, old (N380 Ka), oceanic crust in U–Th–Ra secular equilibrium. The basaltic or metabasaltic assimilant envisaged has 87Sr/86Sr isotope ratios similar to Pacific- or Indian-MORB that has undergone minimal alteration by hydrothermal fluids. Stage 3: AFC continues as the magma ascends through the arc crust via conduits and possibly small-scale magma chambers. At around ~ 51 wt.% SiO2 in magmatic evolution, the role of ferromagnesian mineral fractionation decreases and plagioclase takes over as the dominant phase crystallising in the magma. It is feasible that assimilation of older volcanic material may also occur at this stage.

Fig. 7. Schematic representation of petrogenesis at Lopevi Volcano, following the description given in the text.

However, the relative effect of this on lava geochemistry may be subtle and insignificant in comparison to the effect from assimilation of older and more primitive oceanic crust. Ra-Th disequilibria in the erupted lavas suggests that assimilation of arc crust during differentiation from basalt to basaltic andesite occurs extremely rapidly over a time period of a few hundred years. 7. Concluding remarks A detailed geochemical, isotopic and U-series study of Lopevi volcanic rocks has revealed rapid timescales for magmatic differentiation from basalt to basaltic andesite (b1000 years) indicating that the relatively thin crust permits rapid ascent and magma differentiation. Such conclusions are consistent with geophysical studies indicating that crustal magma chambers are small or absent in the Vanuatu arc (Iyer, 1984) and also concur with the 102–103 year timescales predicted by some numerical models for cooling and crystallisation of magma chambers (Marsh, 1989) and AFC processes (e.g. Edwards and Russell, 1998). Eruptive episodes at Lopevi occur on the scale of decades and so are shorter than can be resolved with 226Ra–230Th data. Future study of U-series nuclides with shorter half lives e.g. 210Pb (22.6 years) or 228Ra (5.75 years) (Holden, 1990) may be able to place tighter constraints on timescales of AFC and magmatic cycles at Lopevi. Here we explain the discrepancy of long (104–105 years) and short 3 (10 years) timescale information given by closed-system analysis of the 230Th and 226Ra isotope data (respectively) (i.e. crystallisation timescales) by the operation of open-system processes. Assimilation of old (N380 Ka) crustal material in U–Th–Ra secular equilibrium exerts strong control on Th and Ra isotopes and does not permit truthful timescales of magmatic differentiation to be extracted from the U–Th equiline diagram. U-series isotopes are, therefore, extremely sensitive to assimilation of old crustal material, even in intra-oceanic arcs where magma rises rapidly. This study suggests that the early stage of magmatic differentiation in intra-oceanic arcs may be characterised by assimilation of ‘primitive’ oceanic crust during fractional crystallisation and that the high MgO arc lavas may form part of a genetically linked suite (cf. distinct magma types). The main implication from this study is that crustal interaction may be more common than previously suspected at intra-oceanic arcs, and than can be constrained by conventional isotopic detectors (e.g. Sr–Nd isotopes) if the arc crust is dominated by older volcanic material. Subsequently, we propose that timescales at intra-oceanic arcs may be significantly overestimated where ‘invisible’ (i.e. undetectable) AFC processes operate and are hard to discern. This is thought to be the case at Miyakejima volcano in Japan (Yokoyama et al., 2006) where interaction between older volcanic edifice and recent magmatism has not been recorded by the Sr–Nd–Pb isotopic systems due to the relatively short timescales involved. Long-times scales of magma residence at intra-oceanic arcs, for example, 60 kyr at Soufrière volcano on St. Vincent (Heath et al., 1998) and 30 to70 kyr at La Soufrière volcano in Guadeloupe (for closed and open-system i.e. magma injections, respectively) (Touboul et al., 2007) in the Lesser Antilles, could possibly result from some degree of ‘invisible’ crustal assimilation. Heath et al. (1998) take the lack of positive correlation between 87Sr/86Sr and SiO2 as evidence that the effects of crustal contamination are insignificant, despite the general positive correlation observed in the older pre-somma lavas on St. Vincent and evidence for AFC processes in other volcanoes of the central and northern Lesser Antilles (Davidson, 1987; Defant et al., 2001). This hypothesis is further supported by the presence of slight 226Ra excesses, (226Ra/230Th ~ 1.4–2.7) in recently erupted (20th century) andesites on St. Vincent (Chabaux et al., 1999; Turner et al., 2001) and recent lavas on Guadaloupe (Touboul et al., 2007) which conflict with long closed-system timescales from Th isotopes. ‘Invisible’ AFC is also thought to explain U–Th isotope variations in some continental arc

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magmas from the Aleutians (George et al., 2004) and Cascades (Reagan et al., 2003). On the basis of U-Th-Ra data Garrison et al. (2006) indicate the time taken for assimilation of mafic/juvenile material and ascent of andesite at Cotopaxi volcano in the Andean continental-arc was relatively fast, less than b8000 years (cf. 70– 100 kyr for the rhyolites). Finally, we point out that correlation of Ba/Th ratios with U-series data within consanguineous magmatic suites is not solely controlled by variation in fluid flux from slab (cf. Turner et al., 2001 and references within) but can also be created by AFC of low Ba/Th material (i.e. oceanic crust) and fractional crystallisation of a mineral assemblage including plagioclase. This study, therefore, emphasises the importance of detailed geochemical studies of individual volcanic centres prior to determination of meaningful timescales of magmatic differentiation and the undertaking of magmatic source investigations. Acknowledgements We thank Peter Wieland and Norm Pearson for invaluable assistance during analytical work at Macquarie University. We are grateful to Ken Sims and Georg Zellmer for constructive reviews which greatly improved the manuscript. Editorial handling by Rick Carlson is also acknowledged. This study used instrumentation funded by Australian Research Council LIEF and DEST Systemic Infrastructure Grants, Macquarie University and Industry. S.T. acknowledges the support of an Australian Research Council (ARC) Federation Fellowship. This project was supported by ARC discovery project 0771610 and is GEMOC publication #536. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2008.06.032. References Allègre, C.J., Condomines, M., 1982. Basalt genesis and mantle structure studied through Th isotopic geochemistry. Nature 299, 21–24. Asmerom, Y., DuFrane, S.A., Mukasa, S.B., Cheng, H., Edwards, R.L., 2005. Time scale of magma differentiation in arcs from protactinium-radium isotopic data. Geology 33, 633–636. Beattie, P., 1993. The generation of uranium series disequilibria by partial melting of spinel peridotite: constraints from partitioning studies. Earth Planet. Sci. Lett. 117, 379–391. Berlo, K., Turner, S., Blundy, J., Hawkesworth, C., 2004. The extent of U-series disequilibria produced during partial melting of the lower crust with implications for the formation of the Mount St. Helens dacite. Contrib. Mineral. Petrol. 148, 122–130. Blake, S., Rogers, N., 2005. Magma differentiation rates from 226Ra/230Th and the size and power output of magma chambers. Earth Planet. Sci. Lett. 236, 654–699. Blundy, J., Wood, B., 2003. Mineral-melt partitioning of uranium, thorium and their daughters. In: Bourdon, B., Henderson, G.M., Lundstrom, C.C., Turner, S.P. (Eds.), Uranium Series Geochemistry. Reviews in Mineralogy and Geochemistry, vol. 52. The Mineralogical Society of America, Washington, DC, pp. 59–123. Bowen, N.L., 1928. The Evolution of Igneous Rocks. Princeton University Press, Princeton, NJ. Briqueu, L., Javoy, M., Lancelot, J.R., Tatsumoto, M., 1986. Isotope geochemistry of recent magmatism in the Aegean arc: Sr, Nd, Hf and O isotopic ratios in the lavas of Milos and Santorini — geodynamic implications. Earth Planet. Sci. Lett. 80, 41–54. Chabaux, F., Hémond, C., Allègre, C.J., 1999. 238U–230Th–226Ra disequilibria in the Lesser Antilles arc: implications for mantle metasomatism. Chem. Geol. 153, 171–185. Chauvel, C., Blichert-Toft, J., 2001. A hafnium isotope and trace element perspective on melting of the depleted mantle. Earth Planet. Sci. Lett. 190, 137–151. Cheng, H., Edwards, R.L., Hoff, J., Gallup, C.D., Richards, D.A., Asmerom, Y., 2000. The half-lives of uranium-234 and thorium-230. Chem. Geol. 169, 17–33. Condomines, M., Gauthier, P.-J., Sigmarsson, O., 2003. Timescales of magma chamber processes and dating of young volcanic rocks. In: Bourdon, B., Henderson, G.M., Lundstrom, C.C., Turner, S.P. (Eds.), Uranium Series Geochemistry. Reviews in Mineralogy and Geochemistry. The Mineralogical Society of America, Washington, DC, pp. 125–174. Crawford, A.J., Briqueu, L., Laporte, C., Hasenaka, T., 1995. Coexistence of Indian and Pacific oceanic upper mantle reservoirs beneath the Central New Hebrides Island Arc. In: Taylor, B., Natlund, J. (Eds.), Active margins and marginal basins of the western Pacific. Geophysical Monograph, vol. 88. American Geophysical Union, pp. 199–217.

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