Magma Evolution in the Primitive, Intra-oceanic Tonga Arc: Rapid ...

Magma Evolution in the Primitive, Intra-oceanicTonga Arc: Rapid Petrogenesis of Dacites atFonualeiVolcano

SIMON TURNER1*, JOHN CAULFIELD1, TRACY RUSHMER1,MICHAELTURNER1, SHANE CRONIN2, IAN SMITH3 ANDHEATHER HANDLEY1

1GEMOC NATIONAL KEY CENTRE, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY,

SYDNEY, NSW 2109, AUSTRALIA2INSTITUTE OF NATURAL RESOURCES, MASSEY UNIVERSITY, PALMERSTON NORTH, NEW ZEALAND3DEPARTMENT OF GEOLOGY, AUCKLAND UNIVERSITY, AUCKLAND, PB92019, NEW ZEALAND

RECEIVED MAY 19, 2011; ACCEPTED JANUARY 10, 2012ADVANCE ACCESS PUBLICATION FEBRUARY 22, 2012

Fonualei is unusual amongst subaerial volcanoes in theTonga arc be-

cause it has erupted dacitic vesicular lavas, tuffs and phreomagmatic

deposits for the last 165 years. The total volume of dacite may ap-

proach 5 km3 and overlies basal basaltic andesite and andesite

lavas that are constrained to be less than a few millennia in age. All

of the products are crystal-poor and formed from relatively

low-viscosity magmas inferred to have had temperatures of 1100^

10008C, 2^4 wt % H2O and oxygen fugacities 1^2 log units above

the quartz^fayalite^magnetite buffer. Major and trace element

data, along with Sr^Nd^Pb and U^Th^Ra isotope data, are used

to assess competing models for the origin of the dacites. Positive cor-

relations between Sc and Zr and Sr rule out evolution of the

within-dacite compositional array by closed-system crystal fraction-

ation of a single magma batch. An origin by partial melting of

lower crustal amphibolites cannot reproduce these data trends or, ar-

guably, any of the dacites either. Instead, we develop a model in

which the dacites reflect mixing between two dacitic magmas, each

the product of fractional crystallization of basaltic andesite

magmas formed by different degrees of partial melting. Mixing was

efficient because the two magmas had similar temperatures and visc-

osities. This is inferred to have occurred at shallow (2^6 km)

depths beneath the volcano. U^Th^Ra disequilibria in the basaltic

andesite and andesite indicate that the parental magmas had fluids

added to their mantle source regions less than 8 kyr ago and that frac-

tionation to the dacitic compositions took less than a few millennia.

The 165 year eruption period for the dacites implies that mixing

occurred on a similar timescale, possibly during ascent in conduits.

The composition of the dacites renders them unsuitable candidates

as contributors to average continental crust.

KEY WORDS: arcs; dacite; fractional crystallization; amphibolite

melting; timescales

I NTRODUCTIONJust how magmas, especially silicic ones, are generated inthe intra-oceanic (island arc) environment is crucial toour understanding of the volcanic, thermal and fluid re-gimes in this setting. Moreover, it is often inferred thatthis material eventually accretes to active continental mar-gins to form the continental masses. However, the originof felsic magmas (460% SiO2) in intra-oceanic arc settingsis a matter of much current debate (e.g. Wade et al., 2005;Smith et al., 2006, 2009; Brophy, 2008; Reubi & Blundy,2009). In essence, two very different processes are cur-rently invoked: fractional crystallization of basalticmagma and partial melting of lower crustal amphibolite.The latter has recently been given the term ‘hot-zonemodel’ as a result of a number of numerical simulations

*Corresponding author. Telephone: (61) 2 9850 8363. Fax: (61) 2 98506904. E-mail: [email protected]

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(e.g. Petford & Gallagher, 2001; Dufek & Bergantz, 2005;Annen et al., 2006). Unfortunately, fractional crystalliza-tion of basalt and partial melting of amphibolite can oftenlead to very similar major element and trace element char-acteristics in felsic magmas. Although the presence of a bi-modal silica distribution has frequently been interpretedto favor a partial melting origin for the silicic rocks, nu-merous studies have shown this need not be the case (e.g.Turner & Rushmer, 2010). Here we present the results of ageochemical investigation into the origin of young dacitesat Fonualei volcano in the primitive, intra-oceanic Tongaarc. This provides a companion to a similar study of bas-altic andesite evolution on nearby Tofua volcano(Caulfield et al., 2011, 2012).

GEOLOGICAL SETT INGThe Tonga^Kermadec arc has formed in response to sub-duction of the Pacific plate beneath the Australian plateand extends for several thousand kilometers northward ofNew Zealand. Convergence rates reach a global maximumto the north of Tonga and decrease southwards (Beviset al., 1995). The lavas erupted along the arc belong to thelow-K, tholeiitic series and have been studied in detail byEwart and co-workers (Ewart et al., 1973, 1977; Ewart,1976). They are interpreted to have evolved from primarymagmas derived from a mantle wedge that has becomehighly depleted, owing to back-arc melt extraction (Ewart& Hawkesworth, 1987; Caulfield et al., 2008; Cooper et al.,2010), and in which signatures from both aqueous fluids,derived from altered oceanic crust, and sediment meltsare recognizable (e.g. Regelous et al., 1997, 2010; Turneret al., 1997; Ewart et al., 1998; George et al., 2005). For along time the dominant products were thought to be bas-altic andesites and andesites. However, recent dredging bythe Australian research vessel Southern Surveyor has shownthat the majority of volcanoes are submerged calderastructures that are in fact often dominated by silicic (gener-ally dacitic) rocks (Graham et al., 2008). A similar observa-tion has been made in the Izu arc (Tamura et al., 2009)and so there is an emerging realization that the amount ofsilicic magma emplaced in primitive oceanic arcs (andthus their potential volcanic hazard) has been significantlyunderestimated. Additionally, the long-standing notionthat theTonga^Kermadec silicic magmas are the productsof fractional crystallization (e.g. Ewart, 1976) has recentlybeen challenged by Smith et al. (2003, 2006, 2009) and sothis arc lies at the heart of a broader debate.

ERUPT IVE H I STORY OFFONUALEI I SLANDFonualei Island is the small emergent summit of a largesubmarine stratovolcano that rises above the Tofua Ridge,70 km NW of Vava’U in Tonga (Fig. 1a). The submarine

arc front volcanoes to the north of Fonaulei are apparentlyinactive until the latitude of Niuatoputapu Island and thismay be due to the proximity of the obliquely orientedFonualei Spreading Centre causing a reduction inmantle-derived magma flux beneath this part of the arc(Keller et al., 2008). Roughly circular in plan, the island is2 km in diameter and rises to 200m at its centre (Fig. 1b).It forms one of the 7^9 historically active, subaerial volca-noes of the arc, but is distinguished from the others byhaving a predominance of dacite and an uncharacteristic,almost complete absence of basaltic andesite lavas. It isknown to have erupted six times since 1846 (Simkin &Siebert, 1994).The work here is based on 28 samples collected from

Fonualei during a field campaign conducted in 2008. Brieffield descriptions are given in Table 1. The oldest units onthe island are exposed in the south (see Fig.1b), where rela-tively mafic lava flows (e.g. samples F0813 and F0811) areoverlain by thick pyroclastic sequences representing twoor more overlapping pyroclastic cones (Fig. 2). These lavashave the least evolved compositions on the island, rangingfrom basaltic andesite to andesite. The overlying pyroclas-tic units are similar to proximal deposits associated withdevelopment of low-volume, typically basaltic phreato-magmatic pyroclastic cones (e.g. Sohn, 1996), although atFonualei the pyroclastic material is andesitic in compos-ition. The deposit consists of multiple, thin (50·1^0·5m)beds of ash-dominated, poorly sorted lapilli tuff, inter-spersed with well-sorted vesiculated lapilli. The sedi-mentary features, including cross- and ripple-bedding,accretionary lapilli, bomb-sag deformation and load struc-tures, are consistent with their emplacement via cool^moist pyroclastic surge (base surge) and air-fall deposition(see Dellino et al., 1990). There is a general trend toward agreater proportion of air-fall beds with coarser pumiceand lithic lapilli higher up in the sequence. The geometryof these deposits forms a remnant flank of a pyroclasticcone and suggests an earlier eruptive centre slightly SWofthe more recent locus of activity.Two other pyroclastic sequences form segments of a

pyroclastic cone(s) that cap lava-flow units in the northand western sectors of the island (Fig. 1b). These depositsare similar in depositional character to the southern pyro-clastic sequence, but comprise material that is primarilydacitic in composition (e.g. F0817). These cone structuresshow less surface erosion than the southern remnant.However, it could be that this observation is due to thenorthwestern sector being capped by very recent tephra ofsimilar depositional character and composition (e.g.F0814). Site F0817 (Fig. 1b) exposes a series of thinlybedded (50·1^0·3m) pyroclastic units that dip gentlyaway from the centre of the island. These have sedimenta-tion features that indicate deposition from repeated pyro-clastic surges formed during a phreatomagmatic eruption.

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Near the top of the sequence (F0816; Fig. 1b) there is atransition to thicker pyroclastic surge and flow unitswith intercalated well-sorted, vesiculated lapilli horizons.The inner wall of the pyroclastic cone at this location istypically steep with subsequent lava flows impoundedagainst it.The ages of the basaltic andesite and andesites are un-

known, although by analogy with Tofua may be less thana thousand years (Caulfield et al., 2011). The pyroclasticcone(s) remnants of the western and northern portions ofthe island are interpreted here to represent deposits of thelarge explosive eruption of Fonuelei documented in June1846. Reports suggest that the central crater following thisevent was below sea level, with breaches in the NW andeast (Bailey, 1846; Lawry, 1850; Erskine, 1853). Ashfall wasreported in Vava’u, 70 km to the SE of Fonualei, and bythe ship Massachussets some 80 km to the NE of Fonualei.

By late June 1846, an �50m high pyroclastic cone hadalready formed within the central crater area (Bailey,1846), and by August 1847 large volumes of lava had beenerupted (Lawry, 1850), refilling much of the crater. Thesmall dimensions of this crater (�1km diameter) and thethin bedding and fine grain-size of the deposits imply asmall to moderate-sized phreatomagmatic explosive erup-tion, driven by shallow subsurface (and in this case sub-marine) magma^water interaction (see Self et al, 1980).Collapse or excavation of the central crater area is there-fore envisaged to be the result of a maar-forming event,rather than from a caldera collapse that would be asso-ciated with much larger-scale, ignimbrite-forming erup-tions that are documented elsewhere at nearby volcanoes(e.g. Caulfield et al., 2011). This age-interpretation impliesthat the dacitic lavas stratigraphically abutting these pyro-clastic sequences are younger than 1846.

Fig. 1. (a) Regional map of theTonga arc. (b) Map of Fonualei island showing the general geology and sample locations. Black numbers indi-cate samples from lava flows; grey numbers indicate samples of pyroclastic material. All sample numbers have a prefix F08. Sample numberslie near the actual sample locations. Ages of lava flows and pyroclastic cones are approximate. These ages are based on the historical records,as well as the stratigraphy and morphology of each deposit. For example, a tephra overlies lava F0809 but not lava F0812. Therefore, F0809 isolder than F0812. However, the age difference could be many hours or many years. The andesitic lava flows (samples F0811, F0813) are theoldest lava flow preserved on the island.

TURNER et al. DACITE PETROGENESIS, FONUALEI


Table 1: Fonualei whole-rock major and trace element analyses

Sample no. Field description SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 H2O LOI Total

F0813 Lava flow beneath pyroclastic cone 57·14 0·65 14·67 11·94 0·21 3·07 7·14 2·14 0·77 0·15 0·42 1·81 100·10

F0811 Lava flow 60·41 0·65 14·37 10·99 0·20 2·46 7·39 2·40 0·86 0·17 0·02 �0·54 99·39

F0817 Lava flow 64·18 0·58 13·46 8·57 0·18 1·39 5·64 2·95 1·16 0·23 0·08 �0·34 98·08

F0804 Lava flow 64·46 0·59 13·50 8·64 0·18 1·41 5·68 2·99 1·16 0·21 0·11 �0·33 98·61

F0810 Lava flow beneath pyroclastic cone 64·62 0·59 13·64 8·67 0·18 1·46 5·73 2·95 1·18 0·22 0·07 �0·22 99·10

F0806C-2 Lava flow 64·96 0·57 13·77 8·58 0·18 1·43 5·76 2·97 1·19 0·21 0·03 �0·29 99·36

F0806D-2 Pumice from pyroclastic cone 65·00 0·55 13·74 8·57 0·18 1·44 5·73 2·99 1·09 0·22 0·09 �0·26 99·32

F0806C-1 Pumice from pyroclastic cone 65·04 0·57 13·71 8·54 0·18 1·41 5·72 2·99 1·16 0·21 0·07 �0·14 99·47

F0816 Obsidian bomb 65·11 0·57 13·66 8·55 0·18 1·42 5·71 3·01 1·15 0·22 0·03 �0·30 99·32

F0806B-2 Pumiceous bomb 65·14 0·57 13·70 8·55 0·18 1·41 5·70 2·99 1·20 0·21 0·04 0·08 99·77

F0823 Grey scoriaceous clast from pyroclastic cone 65·15 0·56 13·79 8·61 0·18 1·44 5·76 2·98 1·12 0·21 0·02 �0·41 99·42

F0818 Lava flow 65·21 0·59 13·63 8·66 0·18 1·44 5·71 3·00 1·15 0·22 0·03 �0·47 99·34

F0806D-1 Lava flow beneath eroded pyroclastic cone 65·31 0·57 13·83 8·62 0·18 1·43 5·77 2·98 1·22 0·22 0·03 �0·23 99·93

F0819 Black scoriaceous bomb 65·33 0·55 13·98 8·63 0·17 1·46 5·81 2·97 1·13 0·21 0·04 �0·41 99·88

F0807 Lava flow 65·34 0·58 13·83 8·60 0·18 1·43 5·49 3·03 1·17 0·23 0·11 �0·03 99·95

F0808 Lava flow 65·37 0·56 14·00 8·70 0·18 1·51 5·87 2·98 1·12 0·21 0·07 �0·43 100·13

F0815A Scoria clast from pyroclastic cone 65·44 0·56 13·66 8·54 0·18 1·41 5·67 3·03 1·16 0·21 0·02 �0·27 99·62

F0812 Pyroclastic clast 65·45 0·59 13·72 8·70 0·18 1·46 5·76 3·04 1·14 0·22 0·07 �0·40 99·93




F0801A Pyroclastic clast 65·78 0·58 13·76 8·63 0·18 1·44 5·70 3·03 1·15 0·22 0·02 �0·42 100·09

F0822 Lava flow 65·83 0·58 13·62 8·60 0·18 1·41 5·63 3·05 1·18 0·22 0·07 �0·40 99·97

F0801 Lava flow beneath pyroclastic cone 65·86 0·58 13·78 8·59 0·18 1·44 5·71 3·03 1·16 0·24 0·00 �0·46 100·09

F0802 Lava flow 65·90 0·58 13·83 8·72 0·18 1·49 5·74 3·06 1·14 0·22 0·03 �0·45 100·45

F0805 Lava flow 66·03 0·59 13·51 8·67 0·19 1·41 5·56 3·07 1·19 0·27 0·06 �0·39 100·16

F0814 Lava flow 66·05 0·54 13·28 8·00 0·18 1·21 5·21 2·97 1·22 0·21 0·04 �0·09 98·82

F0821 Lava flow 66·09 0·58 13·80 8·59 0·18 1·40 5·64 3·06 1·17 0·23 0·06 �0·41 100·39

JA1 64·80 0·84 15·27 7·04 0·16 1·54 5·68 3·87 0·78 0·16 – – 100·14

BHVO-2 – – – – – – – – – – – – –

Sample no. Li Be Sc V Cr Co Ni Cu Zn Ga Rb Sr Y Zr Nb Cs Ba La

F0813 5·16 0·59 43·61 338·0 2·83 30·74 3·84 98·7 112·84 10·70 10·64 311·45 18·08 32·25 0·625 0·367 153·26 2·774

F0811 2·78 0·61 38·41 221·9 2·37 23·21 1·91 19·3 103·18 10·65 10·56 302·23 19·64 36·24 0·702 0·137 166·85 2·915

F0817 8·14 0·75 28·60 81·1 2·66 15·47 1·42 28·4 105·72 10·52 15·65 290·71 24·11 47·49 0·923 0·553 220·29 4·006

F0804 6·47 0·62 24·92 69·9 2·51 11·81 0·64 23·5 92·94 9·43 13·58 256·19 21·10 41·06 0·746 0·465 190·02 3·500

F0810 7·71 0·71 28·33 87·6 2·58 14·57 1·44 26·5 107·23 11·04 15·59 293·05 24·11 46·59 0·877 0·568 227·18 4·157

F0806C-2 7·71 0·64 26·21 75·7 2·43 12·30 0·65 24·7 97·27 9·85 14·17 271·92 21·99 42·09 0·779 0·531 221·70 3·952

F0806D-2 4·20 0·69 30·13 100·2 5·26 14·90 0·92 18·3 111·12 11·11 15·68 308·52 24·63 48·64 0·950 0·540 254·61 4·486

F0806C-1 8·22 0·73 29·52 84·4 3·74 15·10 0·76 33·7 111·27 11·36 16·12 309·71 24·86 48·41 0·893 0·613 253·02 4·538

F0816 8·14 0·74 27·97 85·3 2·56 13·55 1·29 21·0 108·44 11·04 15·80 292·57 24·20 46·54 0·888 0·567 227·51 4·135

F0806B-2 7·17 0·61 24·30 69·7 2·26 11·57 0·58 23·6 92·02 9·29 13·43 256·58 20·63 40·33 0·744 0·509 210·43 3·759

F0823 6·25 0·75 31·16 102·8 3·83 15·78 1·03 24·4 112·90 11·38 15·20 321·03 25·86 49·45 0·963 0·256 257·73 4·710

F0818 7·25 0·64 27·57 75·6 2·48 13·05 0·84 26·7 98·23 10·06 14·24 279·99 23·07 44·92 0·855 0·383 232·68 4·246

F0806D-1 7·42 0·61 25·75 76·6 2·41 12·26 0·65 26·2 99·15 9·91 14·53 275·63 22·48 43·83 0·823 0·544 230·15 4·097

F0819 5·86 0·61 26·44 88·0 3·80 12·84 0·85 30·8 91·68 9·83 13·68 275·60 21·65 42·97 0·803 0·316 223·94 3·969

F0807 7·38 0·61 26·74 72·6 2·58 13·70 0·73 30·8 98·27 9·78 14·64 267·88 21·85 44·47 0·825 0·551 225·65 4·018

F0808 8·61 0·73 31·49 108·4 3·33 15·76 0·99 41·6 112·53 11·57 15·61 323·83 25·93 50·06 0·946 0·230 263·94 4·734

(continued)



Table 1: Continued

Sample no. Li Be Sc V Cr Co Ni Cu Zn Ga Rb Sr Y Zr Nb Cs Ba La

F0815A 8·21 0·73 27·74 84·5 4·09 13·33 1·34 20·7 104·29 10·35 15·48 288·35 23·80 46·57 0·893 0·550 223·70 4·012

F0812 7·40 0·75 28·49 86·5 2·89 13·70 1·38 20·0 104·38 10·40 15·40 289·14 23·78 46·50 0·889 0·540 223·98 4·020

F0803 7·56 0·75 28·31 81·9 3·12 13·54 1·36 21·3 103·18 11·11 15·64 286·74 23·90 46·22 0·883 0·562 228·80 4·170

F0809 6·81 0·72 28·37 81·1 3·17 13·58 1·29 20·7 104·40 10·80 15·46 287·68 23·89 46·16 0·874 0·533 220·47 4·002

F0820 6·54 0·64 26·77 72·2 2·51 12·37 0·72 20·0 98·26 8·86 14·11 272·02 22·77 43·80 0·829 0·487 227·31 4·111

F0801A 9·07 0·79 31·95 86·7 3·55 14·61 1·01 11·7 113·38 12·07 16·65 333·04 26·62 52·52 0·975 0·518 277·50 4·877

F0822 7·07 0·72 28·34 81·3 2·73 13·46 1·34 19·7 105·92 10·67 16·09 292·94 24·54 47·62 0·899 0·556 221·87 4·011

F0801 8·11 0·75 28·45 82·4 2·78 13·30 1·53 7·8 100·37 10·50 14·81 289·83 23·53 46·60 0·899 0·431 227·14 4·011

F0802 4·31 0·48 20·10 57·4 2·29 9·41 0·72 16·3 70·71 7·41 9·70 202·48 16·69 31·87 0·593 0·228 167·19 3·001

F0805 8·06 0·72 28·69 73·3 2·76 13·20 0·71 28·1 108·07 10·74 15·99 292·90 24·99 47·78 0·887 0·612 254·87 4·580

F0814 8·27 0·77 26·74 54·9 2·82 10·79 1·21 13·9 107·49 10·50 15·85 291·78 24·56 47·19 0·910 0·571 229·14 4·137

F0821 6·75 0·71 27·68 76·7 2·80 13·24 1·26 12·8 101·91 10·55 15·58 290·15 24·18 46·46 0·899 0·512 222·16 4·054

JA1 – – – – – – – – – – – – – – – – – –

BHVO-2 4·47 1·38 34·57 332·5 267·5 46·14 135·1 119·5 104·22 13·10 9·39 397·02 28·96 183·99 19·540 0·100 131·82 15·447

Sample no. Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Hf Ta Pb Th U

F0813 6·858 1·119 5·457 1·703 0·624 1·988 0·360 2·345 0·521 1·563 1·570 0·240 0·852 0·033 2·464 0·246 0·201

F0811 7·246 1·179 5·827 1·859 0·668 2·153 0·387 2·499 0·558 1·651 1·660 0·253 0·915 0·035 1·815 0·256 0·204

F0817 9·818 1·583 7·719 2·415 0·816 2·765 0·502 3·207 0·720 2·147 2·153 0·331 1·258 0·045 3·620 0·370 0·304

F0804 8·477 1·373 6·780 2·137 0·737 2·574 0·462 2·974 0·679 2·037 2·048 0·328 1·194 0·030 3·798 0·439 0·358

F0810 10·191 1·631 8·027 2·523 0·856 2·931 0·525 3·434 0·759 2·248 2·300 0·355 1·339 0·053 4·140 0·429 0·357

F0806C-2 9·664 1·562 7·630 2·406 0·833 2·865 0·514 3·280 0·747 2·237 2·235 0·352 1·291 0·049 3·910 0·432 0·361

F0806D-2 10·769 1·751 8·680 2·768 0·945 3·214 0·576 3·721 0·835 2·499 2·509 0·389 1·496 0·068 4·020 0·496 0·406

F0806C-1 11·062 1·791 8·759 2·745 0·941 3·269 0·588 3·760 0·848 2·562 2·546 0·401 1·480 0·050 4·516 0·500 0·409

F0816 10·222 1·616 7·972 2·489 0·848 2·902 0·517 3·384 0·753 2·239 2·251 0·352 1·319 0·053 4·106 0·419 0·350

F0806B-2 9·140 1·484 7·228 2·306 0·779 2·711 0·488 3·126 0·706 2·125 2·114 0·334 1·235 0·031 3·741 0·411 0·342

F0823 11·164 1·844 8·990 2·838 0·967 3·354 0·596 3·833 0·868 2·573 2·619 0·406 1·486 0·035 3·821 0·505 0·425

F0818 10·127 1·652 8·166 2·568 0·873 3·026 0·538 3·484 0·783 2·337 2·354 0·367 1·360 0·028 4·452 0·453 0·375

F0806D-1 9·803 1·610 7·912 2·472 0·849 2·946 0·522 3·370 0·758 2·262 2·304 0·358 1·332 0·032 4·023 0·445 0·369

F0819 9·535 1·565 7·735 2·407 0·838 2·849 0·505 3·261 0·728 2·185 2·227 0·345 1·292 0·017 3·100 0·435 0·362

F0807 9·635 1·574 7·737 2·423 0·830 2·856 0·507 3·297 0·742 2·222 2·237 0·349 1·321 0·035 3·871 0·437 0·360

F0808 11·330 1·845 9·096 2·833 0·991 3·365 0·600 3·880 0·869 2·596 2·631 0·411 1·519 0·045 3·144 0·506 0·419

F0815A 9·834 1·587 7·750 2·447 0·820 2·813 0·507 3·252 0·740 2·168 2·178 0·335 1·269 0·045 3·853 0·385 0·322

F0812 9·836 1·575 7·752 2·422 0·817 2·807 0·505 3·246 0·733 2·166 2·182 0·338 1·270 0·046 3·898 0·385 0·322

F0803 10·261 1·642 8·121 2·539 0·862 2·967 0·532 3·475 0·775 2·305 2·321 0·359 1·358 0·050 4·241 0·441 0·365

F0809 9·882 1·572 7·703 2·422 0·821 2·782 0·501 3·243 0·721 2·147 2·148 0·336 1·266 0·048 3·891 0·390 0·328

F0820 9·862 1·615 8·006 2·494 0·846 2·964 0·524 3·384 0·759 2·277 2·307 0·362 1·333 0·032 3·353 0·445 0·368

F0801A 12·017 1·943 9·517 3·001 1·029 3·548 0·642 4·121 0·934 2·755 2·772 0·436 1·610 0·045 3·380 0·544 0·429

F0822 9·885 1·571 7·711 2·409 0·820 2·771 0·496 3·213 0·709 2·107 2·111 0·324 1·223 0·046 3·592 0·360 0·294

F0801 9·881 1·596 7·845 2·460 0·840 2·843 0·518 3·299 0·746 2·206 2·225 0·343 1·295 0·048 2·736 0·411 0·326

F0802 7·307 1·185 5·811 1·846 0·628 2·186 0·393 2·501 0·568 1·712 1·707 0·268 0·992 0·030 2·809 0·332 0·275

F0805 11·144 1·810 8·876 2·819 0·943 3·328 0·597 3·764 0·856 2·571 2·535 0·402 1·464 0·032 4·399 0·494 0·407

F0814 10·168 1·621 7·975 2·525 0·838 2·886 0·519 3·359 0·747 2·222 2·230 0·348 1·289 0·052 4·023 0·403 0·334

F0821 9·952 1·584 7·778 2·448 0·825 2·794 0·501 3·235 0·722 2·169 2·161 0·333 1·265 0·045 3·476 0·383 0·321

JA1 – – – – – – – – – – – – – – – – –

BHVO-2 37·552 5·422 24·484 6·177 2·011 6·206 0·959 5·245 1·000 2·523 1·957 0·271 4·319 1·116 1·589 1·221 0·407

LOI, loss on ignition.



There have been at least 10 lava flows erupted since 1846,many originating from the summit and NW flanks of themain cone, which have built up much of the central partof the island and reached the sea through breaches in theeast, NW and SW flanks. Many of these lava flows wereprobably erupted soon after the 1846 event. A black, ves-icular lava flow in the SW that extends out to the coastthrough a narrow breach in the crater wall has been attrib-uted to an eruption in 1939 (Bryan et al., 1972). A central4100m high pyroclastic cone forms the current summit ofthe island and is blanketed by highly hydrothermallyaltered pyroclastic material. The crater is �100m in diam-eter and is breached to the east.There are several explosionpits located within this crater, which could be from thelast explosive volcanism from Fonualei in 1943 (Bryanet al., 1972). Site F0806 (Fig. 1b) marks the edge of a pyro-clastic cone constructed between the current summit coneand the western pyroclastic wall. This feature may alsohave been associated with the 1943 eruptions. The mostrecent eruptive activity known from Fonualei occurred inthe west, where a small lava flow (F0805; Fig. 1b) wasemplaced some time between 1979 and 1990, as evidenced

by successive aerial photographs. The approximatevolume of dacitic rock is estimated to be �5 km3, which, ifall erupted within the last 165 years (i.e. since 1846),implies an eruption rate of �0·03 km3a�1.

ANALYT ICAL TECHNIQUESFor mineral analysis, polished thin sections were carboncoated and analysed on a Cameca SX100 electron micro-probe at the Geochemical Analysis Unit (GAU) atMacquarie University. An accelerating voltage of 15 keVproduced a focused beam current of 20 nA. A 10 mm beamwas used and a counting time of 10 s for both peak andbackground measurements. Spectrometer calibration wasachieved using the following standards: albite (Na), hema-tite (Fe), kyanite (Al), olivine (Mg), chromium (Cr), spes-sartine garnet (Mn), orthoclase (K), wollastonite (Ca, Si)and rutile (Ti).Because of the potential for seawater contamination,

fresh cores were sawn out of the samples and crushed,then rinsed in Milli-Q water several times, dried andfinally powdered in an agate mill. Major element

Fig. 2. Over 100m of dacitic tephra or tuff beds, inferred to have erupted within the last 165 years, overlie more mafic lava flows on the southernquadrant of Fonualei. Typical, well-bedded tuff and lapilli tuff beds are seen at the contact (dashed white line) with underlying basaltic andesiteand andesite lavas. Total cliff height �150m.



concentrations were determined by X-ray fluorescence(XRF) on a Seimens SRS300 instrument at theUniversity of Auckland, following standard techniques(Norrish & Hutton, 1969). Trace element concentrationswere analysed at the GAU by inductively coupled plasmamass spectrometry (ICP-MS) using an Agilent 7500CSsystem. The samples were dissolved in an HF^HNO3 mix,taken up in 2% HNO3, spiked and analysed followingsimilar procedures to those described by Eggins et al.(1997). Results for rock standards JA1 and BHVO-2 arelisted inTable 1.Samples for Sr^Nd isotopic analysis were dissolved using

an HF^HNO3 mix in heated Teflon beakers. Sr was iso-lated by a single pass through Teflon columns containingBiorad� AG50W-X8 (200^400 mesh) cationic exchangeresin. The Nd fraction was purified using EIChrom�

LN-spec resin according to the method of Pin &Zalduegui (1997). Samples were loaded onto outgassedsingle (Sr) and double (Nd) rhenium filaments using 2 mlof TaCl5þHFþH3PO4þH2O (Birck, 1986) and 5 ml of1N HCl:0·35N H3PO4 activator solutions, respectively.Analyses were performed by thermal ionization mass spec-trometry (TIMS) using aThermoFinniganTriton� systemin static mode at the GAU. Instrument mass fractionationwas accounted for by normalizing 87Sr/86Sr and143Nd/144Nd to 86Sr/88Sr¼ 0·1194 and 146Nd/144Nd¼0·7219, respectively. Analysis of SRM-987 yielded an87Sr/86Sr value of 0·710272�4. BHVO-2 served as a rockstandard and yielded 87Sr/86Sr of 0·703498�6 and143Nd/144Nd of 0·512975�3.For Pb-isotope compositions, the powders were leached

following the procedures of Weis et al. (2005). Powderswere placed in 7ml Savillex� beakers with 1ml of 6Nsub-boiled HCl and ultra-sonicated for 15min. This stepwas repeated four times with a further three rinses inMilli-Q water to remove any residual acid and the residuadried at 1208C on a hot plate. Digestion of 100^120mg ofleached sample was carried out in the same beakers usinga 1:1 mixture of concentrated SEASTAR� HFand HNO3.Samples were dried down at 1508C before being taken upin 0·5ml of HNO3 and 130 ml of HBr and dried. Finally6ml of SEASTAR� 6N HCl was added. Pb was separatedby anionic exchange using 100 ml of Biorad� AG1-X8 resinloaded in shrink-fit Teflon columns and HBr^HNO3 mix-tures according to the method of Lugmair & Galer (1992).Pb isotopes were measured on a Nu Instruments� multi-collector (MC)-ICP-MS system at the GAU. The collectorcup configuration was the same as that outlined byBelshaw et al. (1998). Sample solutions were diluted to10 ppb to match the standard concentration and spikedwith Tl. Relevant Pb isotopes were analysed simultan-eously (static collection mode), along with interference-free isotopes 205Tl, 203Tl and 202Hg. 202Hg was analysedto monitor interference of 204Hg on 204Pb using a mercury

factor of 0·229. Mass fractionation was corrected for usingTl isotopes, applying an exponential correction and a205Tl/203Tl value of 2·3875 (Belshaw et al., 1998). Totalprocedural blank levels were �60 pg, which is negligiblecompared with the total amount of Pb analysed.Periodic analysis of SRM981 (n¼12) yielded meanvalues of 206Pb/204Pb¼16·936, 207Pb/204Pb¼15·492 and208Pb/204Pb¼ 36·689 respectively, and BHVO-2 yielded206Pb/204Pb¼18·620, 207Pb/204Pb¼15·482 and208Pb/204Pb¼ 38·116 (n¼ 2), similar to the lower end of re-sults from leached samples reported byWeis et al. (2005).U, Th and Ra concentrations and isotope ratios were

determined on samples that were spiked with 226U^229Thand 228Ra tracers and dissolved using an HF^HNO3^HClmix in heatedTeflon pressure bombs.The product was con-verted to chloride using 6N HCl and then 6N HCl satu-rated with H3BO3 to drive off residual fluorides. The finalproduct was then converted to nitrate using 14N HNO3

and finally taken up in 7N HNO3. U and Th purificationwas achieved via a single pass through a 4ml anionicresin column using 7N HNO3, 6N HCl and 0·2N HNO3

as elutants. We purposefully avoided the use of ElChrom�

resins for the U^Th chemistry as these can bleed organicmolecules that lead to memory effects and interferencesduring MC-ICP-MS analysis. Concentrations and isotoperatios were measured in dynamic mode on the NuInstruments� MC-ICP-MS system at the GAU. 238U and235U were analysed on Faraday cups, using the 238U/235Uratio to determine the U mass bias, assuming238U/235U¼137·88, whereas 236U and 234U were alter-nately collected in the IC0 ion counter that is preceded byan energy filter. The IC0 gain was determined duringinterspersed dynamic analyses of CRM145 assuming a234U/238U ratio of 5·286�10�5 (Cheng et al., 2000).Methods forTh isotope measurements employed a dynam-ic routine with 232Th in Faraday cups and 230Th and229Th alternating on IC0 and using bracketing measure-ments of the Th‘U’ standard to obtain the Th mass bias,which is different from that for U. Measurements atmasses 230·5 and 229·5 were used to derive a correctionfor residual 232Th tail interference as described in detailby Sims et al. (2008, Appendix A) and Turner et al. (2011).Multiple analyses of the secular equilibrium rock standardTML-3 (n¼ 5) performed at the same time as the analysesof the samples yielded the results U¼10·315 ppm,Th¼ 29·034 ppm, (234U/238U)¼ 1·002, (230Th/232Th)¼1·082 and (230Th/238U)¼ 1·004, which are within error ofsecular equilibrium and published values for this rock(Sims et al., 2008). [See Beier et al. (2010) and Turner et al.(2011) for the results for other rock standards and a full dis-cussion of precision and accuracy in this laboratory.]The Ra analysis procedure follows that used by Turner

et al. (2000). Ra was taken from the first elution from theanionic column and converted to chloride using 6N HCl.



This was then loaded in 3N HCl onto an 8ml cationiccolumn and Ra eluted using 3·75MHNO3, and the processrepeated on a scaled-down 0·6ml column. The rare earthelements (REE) were then removed using a 150 ml columnof ElChrom� Ln-spec resinTM and 0·1N HNO3. Ra andBa were finally chromatographically separated usingElChrom� Sr-spec resinTM and 3N HNO3 as elutant in a150 ml procedure. Samples were loaded onto degassed Refilaments using aTa^HF^H3PO4 activator solution (Birck,1986) and 228Ra/226Ra ratios were measured by TIMS toa precision typically �0·5% in dynamic ion countingmode on a ThermoFinnigan Triton� system at the GAU.Organic interferences are often noted at low temperaturesduring TIMS analysis for Ra but were eliminated here byusing the instrument fitted with a dry scroll pump insteadof the standard rotary pump. This prevents leakage of or-ganic molecules into the source during venting. Accuracywas assessed via replicate analyses (n¼ 5) of TML-3,which yielded 226Ra¼ 3532 fg g�1 and (226Ra/230Th)¼1·02, values that are within error of secular equilibrium[see also Beier et al. (2010) andTurner et al. (2011)]. Becauseeither the ages of the samples (including Late Island) areunknown, in the case of the Fonualei basaltic andesite andandesite, or the samples are likely to have erupted since1846, in the case of the Fonualei dacites, no age correctionwas applied to any of the U-series data.

PETROGRAPHY AND MINERALCHEMISTRY RESULTSRepresentative photomicrographs are shown in Fig. 3. Allof the rocks from Fonualei are crystal-poor (4^9%) andoften highly vesicular (up to 60%),containing small(�1mm) phenocrysts of plagioclase (4^7%), clinopyroxene(�2%) and orthopyroxene (�1%) with trace amounts oftitanomagnetite (Ewart et al., 1973). These often form glo-meroporphyritic clusters (1^3mm) and are set in a glassyto microcrystalline groundmass that is often laminatedand/or flow banded (Fig. 3). The crystal-poor nature ofthe rocks, including the basaltic andesites and andesites,precludes these compositions reflecting mixtures of crystalsand dacitic groundmass, as suggested for some other an-desites (e.g. Price et al., 2005; Reubi & Blundy, 2009).Comparison of the microprobe analyses of groundmassglass in Supplementary Data Electronic Appendix 1(available for downloading at http://www.petrology.oxfordjournals.org) with the whole-rock analyses inTable 1 shows that the groundmass is typically dacitic[one analysis presented here and two from Ewart et al.(1973) are rhyolitic] and, in keeping with the aphyricnature of the rocks, the whole-rock analyses (exceptingF0813) provide a good approximation of the liquid com-positions (Ewart et al., 1973). Groundmass minerals includeplagioclase, alkali feldspar, hypersthene, pigeonite, iron

oxides, quartz and needles of apatite, but no zircon hasbeen observed, consistent with the very low Zr concentra-tions (552 ppm).Phenocryst compositions are included in Electronic

Appendix 1. Orthopyroxene compositions are Wo5^4^En60^50^Fs45^36 whereas clinopyroxene compositions areWo39^36^En42^36^Fs27^21. Comparison of the Fe2þ/Mgratios of the pyroxenes with those of their host rocks(Fig. 4a and b) indicates equilibrium KD values in therange 0·25^0·4, consistent with published experimentaldeterminations (e.g. Grove & Bryan, 1983; Johnston &Draper, 1992). The 0·5GPa pyroxene quadrilateral projec-tion of Lindsley (1983) indicates temperatures rangingfrom 11008C for the andesites to 10008C for the dacites(Fig. 5), broadly consistent with previous geothermometrystudies (Ewart et al., 1973; Ewart, 1976).As noted by Ewart et al. (1973) forTonga rocks in general,

the plagioclase phenocryst compositions are extremelycalcic, being An84^89 in the basaltic andesite and andesiteand ranging from An85 to An81 even in the dacites.Comparison of the molar Ca/Na ratios of the plagioclaseswith those of their whole-rocks suggests equilibrium KD

values in the range 1^5·5, although the majority lie be-tween 1 and 3·4 (Fig. 4c). At �0·2GPa this corresponds toH2O contents of 0^4wt % (Sisson & Grove, 1993) and da-citic liquids with 2^4wt % H2O would be saturated at0·025^0·05GPa, respectively (Burnham, 1979). Thus, weconsider 6 km a maximum depth of equilibration and themagmas may have resided within a few kilometers of thesurface. On the basis that plagioclase is the liquidusphase, Ewart (1976) combined the temperature estimateswith experimental data to infer PH2O¼Ptotal50·2GPaand used Eu partitioning into plagioclase to estimate logfO2 �^10.The titanomagnetites contain �15^32mol % ulvo« spinel

(Supplementary Data Electronic Appendix 1). By combin-ing this compositional range with the temperature rangeinferred from the pyroxenes we estimate log fO2 �^8[quartz^fayalite^magnetite (QFM)þ 1 to 2] based on therelationships determined by Buddington & Lindsley(1964). This is, again, similar to that estimated by Ewart(1976).Using the method of Shaw (1972) we estimate effective

viscosities in the range 107^108 Pa s for basaltic andesite,andesite and dacites, assuming the temperatures deter-mined from the pyroxene pairs, a minimum of 1wt %H2O and 7% crystallinity with an average crystal diam-eter of 1mm. These viscosities are fairly low for siliceousrocks, reflecting their high temperatures and low crystalcontent. However, they are consistent with the predomin-ance of lava flows, which, although in many cases areshort and stubby, and show strong development of ogiveridges, have developed over relatively low gradients. Insome cases, highly fluid flow cores have formed obvious



(a) (b)

(c) (d)

Fig. 3. Photomicrographs of selected Fonualei rocks. (a) F0815A. Flow banded dacite with trachytic texture. Image shows pyroxene and plagio-clase phenocrysts. (b) F0813. Mildly vesicular, flow banded andesite containing volcanic rock fragments. (c) F0806C. Highly vesicular (30%)flow banded, glassy dacite containing fine-grained plagioclase microlites. (d) F0806B. Flow banded dacite with trachytic texture. Scale for allsections: 10·5mm length along long axis.



drained and collapsed cavities in the central parts of flows,with breakouts occurring from carapace areas of the flowtoes.

WHOLE -ROCK GEOCHEMICALRESULTSWhole-rock major and trace element data for the 28Fonualei samples are given inTable 1, arranged in order ofincreasing silica content. In addition, major and traceelement data plus some U-series data from six new samplescollected from Late Island on a previous expedition in2006 are included in Supplementary Data ElectronicAppendix 2. It should be noted that two of the LateIsland samples have (234U/238U) more than 3% beyondsecular equilibrium and are therefore not used in any ofthe following discussion or diagrams. The ages of all ofthese samples are unknown but are likely to be less than afew millennia, by analogy with age constraints fromTofua(Caulfield et al., 2011) and the presence of 226Ra excesses(see below). Specifically, the volcanological record suggeststhat the dacites from Fonualei are largely less than 165years old (see above). These and new Tofua (Caulfieldet al., 2012) and Late Island data (Supplementary DataElectronic Appendix 2) are presented in selected figuresto provide a reference for the likely composition of localmafic magma input into this part of arc.

Major and trace elementsAs shown in Fig. 6, all but two of the Fonualei samplesare dacites with SiO2 contents that vary from 64·2 to66·1wt %. The remaining two samples comprise a transi-tional basaltic andesite and an andesite that both comefrom the lowermost lava flows and have SiO2 contents of57·1 and 60·4wt %, respectively. This broadens slightlythe 60·3^66·3wt % range reported by Ewart et al. (1973).It should be noted that the matrix compositions fromthree samples (not plotted) extend the overall compos-itional range as high as 71·2wt % SiO2 (Ewart et al., 1973;Supplementary Data Electronic Appendix 1). Most of the

0.0

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Fe2+/Mg whole rock

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/Mg

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opyr

oxen

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e2+/M

g cl

inop

yrox

ene

Ca/Na whole rock

Ca/

Na

plag

iocl

ase

0.4

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0.4

0.25

dry

2 wt. % H2O

6 wt.

% H

2O

4 wt. % H 2O

FO811FO813FO806 C2FO815 AFO821



(c)

(b)

(a)

Fig. 4. Plots of Fe2þ/Mg in (a) orthopyroxene and (b) clinopyroxenevs whole-rock Fe2þ/Mg (cation proportions) assuming a Fe2O3/Fe2O3total ratio of 0·24 based on average FeO determinations ofFonualei dacites (Ewart et al., 1973). The dashed lines show equilib-rium KD values. One anomalous analysis with Fe2þ/Mg¼1·8 hasbeen omitted for clarity. (c) Ca/Na in plagioclase vs whole-rock Ca/Na (cation proportions) contoured for H2O content at �0·2GPa(Sisson & Grove, 1993).

Fig. 5. Pyroxene analyses projected onto the Di^En^Fs^Hd quadri-lateral using the method of Lindsley (1983) indicating magmatic tem-peratures of �11008C for the andesites and 10008C for the dacites.



other major elements have a very restricted range of abun-dances in the dacites (Table 1). For example, whereasFe2O3 and CaO, as well as TiO2 (not plotted) and MgO,decrease from the andesites to the dacites, the dacite suiteitself shows little internal variation in any of these majorelement oxides (Fig. 6b and c). Al2O3 is scattered andNa2O (not plotted) and K2O increase with increasingSiO2 (see inset to Fig. 6c). When combined with the Tofuaand Late Island data (Caulfield et al., 2012), the dacitefromTofua plots amongst the Fonualei dacites whereas theFonualei basaltic andesite and andesite lie between the da-cites and the Tofua and Late Island basaltic andesites onthe Harker diagrams (e.g. Fig. 6a).Concentrations of compatible trace elements are very

low (e.g. Ni and Cr55 ppm) in all of the Fonualei rocksincluding the basaltic andesite and andesite. Incompatibletrace element patterns are typical of arc rocks in generaland characterized by negative Th, Ta^Nb and Ti anoma-lies, and positive Ba, U, Pb and Sr anomalies (Fig. 7).Nevertheless, incompatible trace element abundances arevery low (typically55� primitive mantle), and the rockshave flat REE patterns reflecting the highly depletednature of rocks from this part of the arc (see inset toFig. 7). Their concentrations increase consistently acrossthe basaltic andesite to dacite compositional range butshow little correlation with SiO2 (see inset to Fig. 7) orMgO within the dacites themselves. In contrast, the cor-relations between different incompatible trace elementswithin the dacites are excellent and certainly much betterthan those between major and trace elements. In particu-lar, it is notable that Sc and Sr both increase linearly withincreasing Zr concentration (see below). The presence ofsmall negative Eu anomalies (Eu/Eu*40·94) is consistentwith fractionation of plagioclase under relatively oxidizedconditions, as inferred for these lavas.

Radiogenic isotopesThe new radiogenic isotope data for 10 samples are pre-sented in Table 2 and Fig. 8. The Fonualei rocks are moreradiogenic in 87Sr/86Sr (0·70375^0·70389) and less radio-genic in 143Nd/144Nd (0·512929^0·512961) than basalticandesites from Late Island orTofua. These two isotope sys-tems are not well correlated.Pb isotopes are well correlated (206Pb/204Pb ranges from

18·558 to 18·593, 207Pb/204Pb from 15·505 to 15·551 and208Pb/204Pb from 38·067 to 38·187) and form an array par-allel to that of basaltic andesites from Tofua (Caulfieldet al., 2012). As discussed by Hergt & Woodhead (2007)and Escrig et al. (2009), it is now clear that there are signifi-cant discrepancies between older generationTIMS Pb iso-tope data, such as those reported by Turner et al. (1997),and newer MC-ICP-MS or multi-spike Pb isotope data.However, our new data from Fonualei are in good agree-ment with the data recently presented by Hergt &Woodhead (2007) and Escrig et al. (2009). It should be

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Fig. 6. Harker variation diagrams for (a) total alkalis, (b) Fe2O3 and(c) CaO, highlighting the restricted compositional range of theFonualei dacites, with insets showing MgO and K2O variation.Basaltic andesites from Late Island (black filled circles;Supplementary Data Electronic Appendix 2; Ewart et al., 1998) andTofua Island (grey field; Caulfield et al., 2012) are shown for compari-son in (a). It should be noted that the two Fonualei andesites projectback towards the Late Island and Tofua data and that the Tofuadacite (grey filled circle; data from Caulfield et al., 2012) plots amongstthe Fonualei dacites. Also shown in (b) and (c) are the fields for ex-perimental melts (with SiO2 in the range 56^68wt %), under waterundersaturated conditions, of arc amphibolites with basaltic andesitecompositions (data from Beard & Lofgren, 1991; Rapp et al., 1991;Rushmer, 1991; Rapp & Watson, 1995).



noted that our new within-volcano suites of data fallwithin, but form steeper arrays than, published along-arcdata (Haase et al., 2002; Hergt & Woodhead, 2007; Escriget al., 2009). In keeping with other studies, the Pb isotopedata are consistent with an Indian mid-ocean ridge basalt(MORB) source for the Tonga lavas (e.g. Hergt &Woodhead, 2007; Pearce et al., 2007).

The Fonualei rocks have essentially the same range inradiogenic isotope ratios as the Tofua rocks but are dis-placed, along with published data from Late Island, tolower 206Pb/204Pb. Hergt & Woodhead (2007) suggestedthat the Pb isotopic composition of the fluid componentchanges along the arc and Castillo et al. (2009) haveshown that this is consistent with changes in the Pb isotope

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Rb Ba Th U K Ta Nb La Ce Pb Sr Nd Sm Zr Hf Eu Ti Tb Y Yb Lu

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La Ce Pr Nd SmEu Gd Tb Dy Ho Er Yb Lu3

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Fig. 7. Primitive mantle normalized incompatible trace element diagram for the most primitive Fonualei andesite (F0813) and the most evolvedFonualei dacite (F0821) with a mafic Tofua basaltic andesite (06TF49A) shown for comparison. Insets show Zr vs SiO2 and achondrite-normalized rare earth element diagram. Symbols and data sources as for Fig. 6.

Table 2: Radiogenic and U-series analyses of Fonualei whole-rocks

Sample no. 87Sr/ 143Nd/ 206Pb/ 207Pb/ 208Pb/ U Th 226Ra (234U/ (238U/ (230Th/ (230Th/ (226Ra/

86Sr 144Nd 204Pb 204Pb 204Pb (ppm) (ppm) (fg g�1) 238U) 232Th) 232Th) 238U) 230Th)

FO813 0·703885 0·512947 18·569 15·521 38·116 0·242 0·277 138·0 1·008 2·653 1·607 0·611 2·74

FO811 0·703745 0·512943 18·576 15·541 38·187 0·276 0·322 140·4 1·022 2·603 1·64 0·631 2·37

FO804 0·703765 0·512944 18·578 15·531 38·144 0·362 0·426 225·7 1·014 2·583 1·649 0·644 2·85

FO806D-2 0·703779 0·512929 18·575 15·537 38·176 0·342 0·420 255·2 1·017 2·474 1·648 0·672 3·22

FO806C-1 0·703810 0·512958 18·558 15·527 38·125 0·359 0·423 268·8 1·023 2·572 1·657 0·650 3·44

FO823 0·703779 0·512948 18·566 15·515 38·097 0·355 0·418 301·7 1·007 2·572 1·659 0·651 3·85

FO818 0·703783 0·512940 18·558 15·505 38·067 0·359 0·427 277·8 1·019 2·550 1·640 0·649 3·51

FO820 0·703785 0·512961 18·564 15·551 38·094 0·353 0·418 305·5 1·020 2·565 1·644 0·647 3·94

FO801 0·703799 0·512945 18·593 15·542 38·167 0·336 0·414 224·0 1·011 2·466 1·649 0·674 2·91

FO805 0·703769 0·512948 18·589 15·544 38·180 0·370 0·440 236·0 1·016 2·548 1·656 0·656 2·87



composition of the altered Pacific oceanic crust enteringthe trench. Some of the least radiogenic Pb isotope valuesare observed in the subducting crust adjacent to Fonualeiand, although it is not our primary objective here, themixing trajectories in Fig. 8 show that addition of fluidsfrom this crust to an Indian MORB source that had beenmodified by addition of �0·25% sediment melt can rough-ly simulate the data. Finally, we note that there is no cor-relation between Sr, Nd or Pb isotopes and indices ofdifferentiation, such as SiO2 or Sc (not shown).

U-series isotopesThe new U-series data also show a number of importantfeatures (Table 2). First, despite all efforts to avoid the ef-fects of seawater alteration, most of the samples have(234U/238U) ratios that exceed secular equilibrium beyondanalytical error (estimated to be 0·08% for this period ofanalysis; Beier et al., 2010; Turner et al., 2011). Comparedwith previously published data from Fonualei (Regelouset al., 1997; Turner et al., 1997), our samples have the same(230Th/232Th) ratios as those reported by Turner et al.(1997) but lower and more consistent U/Th ratios (see alsoCaulfield et al., 2012). The (230Th/238U) ratios range from0·611 to 0·674 and (226Ra/230Th) from 2·37 to 3·94, similarto the data presented by Turner et al. (1997, 2000).However, we find no correlation between (234U/238U) andeither (230Th/238U) or (226Ra/230Th). Thus, although wecannot preclude the possibility that some of the sampleshave been affected by a degree of seawater alteration, andthe interpretations below are therefore subject to thatcaveat, it seems unlikely that this is the primary controlon the observed U-series disequilibria.On a U^Th equiline diagram (Fig. 9a) the samples form

a horizontal array overlapping with data from LateIsland (Supplementary Data Electronic Appendix 2) andTofua (Caulfield et al., 2012). The basaltic andesite and an-desite have the lowest (226Ra/230Th) ratios of the suite andthis may reflect decay since eruption as the age of thesesamples is unconstrained. Conversely, the preservation of226Ra disequilibria requires that these samples are58 kyrold. In contrast, the higher (226Ra/230Th) ratios of the da-cites will not have changed over the 165 years since theywere erupted and they form an array separate from theTofua and Late Island rocks on a plot of (226Ra/230Th) vsSiO2 (Fig. 9b). As discussed by Caulfield et al. (2012),assuming that both the 238U and 226Ra excesses reflectfluid addition from the slab (Turner et al., 1997, 2000), thecombined data suggest that this addition must haveoccurred58 kyr ago.

DISCUSS IONModels for the petrogenesis of Tonga basaltic andesites andtheir source regions are discussed in a companion paper(Caulfield et al., 2012). The primary objective of this study,

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Fig. 8. Variation of (a) 87Sr/86Sr vs 143Nd/144Nd, (b) 207Pb/204Pb vs206Pb/204Pb and (c) 208Pb/204Pb vs 206Pb/204Pb comparing the newdata for Fonualei and Tofua (Caulfield et al., 2012) with publisheddata from basaltic andesites on Late Island (Hergt & Woodhead,2007). Curves show the effect of addition of a pelagic sediment melt(72 ppm Sr, 65 ppm Nd, 26 ppm Pb, 87Sr/86Sr¼ 0·7095,143Nd/144Nd¼ 0·5123, 206Pb/204Pb¼18·8, 207Pb/204Pb¼15·67,208Pb/204Pb¼ 38·87) to an Indian MORB source mantle wedge(10 ppm Sr, 0·7 ppm Nd, 0·02 ppm Pb, 87Sr/86Sr¼ 0·70313,143Nd/144Nd¼ 0·51315, 206Pb/204Pb¼18·2, 207Pb/204Pb¼15·49,208Pb/204Pb¼ 37·9) (compositions based on Johnson & Plank, 1999;Stracke et al., 2003; Hergt & Woodhead, 2007; Regelous et al., 2010). Asecond set of mixing vectors show the effect of adding fluid(750 ppm Sr, 3 ppm Nd, 14·6 ppm Pb, 87Sr/86Sr¼ 0·7039,143Nd/144Nd¼ 0·5128, 206Pb/204Pb¼18·52, 207Pb/204Pb¼15·49,208Pb/204Pb¼ 37·67) derived from the subducting Pacific alteredoceanic crust (Pb isotope array from Castillo et al., 2009) to a mantlewedge^sediment melt mixture. Symbols as in Fig. 6.



therefore, is to investigate the origin of the Fonualei dacites.Accordingly we first explore the simplest model ofclosed-system, fractional crystallization (e.g. Ewart et al.,1973) and subsequently more complex scenarios as theyappear to be required by the data. An issue, outlinedabove and discussed in detail by Brophy (2008), is theextent to which geochemical data can readily distinguishbetween the effects of fractional crystallization and am-phibolite partial melting.

Fractional crystallizationCaulfield et al. (2012) have shown that the Tofua dacite,which is similar in many (but not all) aspects to the

Fonualei dacites, could have been produced by fractionalcrystallization from a parental basaltic andesite magma.Similarly, Ewart et al. (1973) presented least-squares calcu-lations showing that the Fonualei dacites could be derivedfrom parental magmas similar to the basaltic andesiteserupted on Late Island and they lie on a projection of theLate^Tofua data in Fig. 6a. Here we use the MELTS algo-rithm (Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998) tofurther investigate the evolution of dacitic magmas in thiscentral section of theTonga arc.In Fig. 10 we compare data from Fonualei, Late Island

andTofua with the results of MELTS models conducted atvarious H2O contents using one of the most mafic Tofuabasaltic andesites for the starting composition. We notethat this has some different traits from the Fonualei basalt-ic andesite both in trace element pattern and isotopic com-position that suggest less of a slab component. Figure 10aillustrates the variation of two-pyroxene temperature vsSiO2, showing that the Fonualei and Tofua samples, forwhich temperature constraints are available, are well mod-eled by the 0·5wt % initial H2O curve (see Caulfieldet al., 2012). In plots of TiO2 vs Fe2O3* and Al2O3 vsCaO, the Fonualei data lie closer to the 1wt % initialH2O curve (Fig. 10b and c). The Late Island samples andthe field for Tofua show more scatter but straddle the0·5^2wt % initial H2O curves at the more primitive com-positional end.The simulations were run at low pressure (0·1GPa) con-

sistent with the mineralogical constraints and previouswork (Ewart et al., 1973; Caulfield et al., 2012). By the timedacitic compositions are achieved in the models shown inFig. 10, the calculated H2O contents are in the range2^4wt %, consistent with the more calcic plagioclase com-positions (Fig. 4c); such magmas would be water saturatedat low pressure in keeping with their highly vesicularnature. Importantly, simulations performed at higher pres-sures produced results that are less consistent with the data.For example, SiO2 is higher at any given temperature andthe suppression of plagioclase andmagnetite crystallizationleads to greater initial increases in Al2O3 and TiO2 withdecreasing CaOand Fe2O3*, respectively (see Fig.10).Overall, the MELTS models confirm that the major

element characteristics of the Fonualei dacites could resultfrom closed-system, low-pressure, fractional crystallizationfrom mafic parental magmas falling within a composition-al range similar to those sampled on Late Island andTofua. This is consistent with the findings of Ewart et al.(1973), and the phases and modes calculated by MELTSare similar to those implied by their least-squares calcula-tions. Nevertheless, the trace element data tell a more com-plex story.As mentioned above, the Fonualei dacites form

well-defined positive arrays on plots of Zr vs Sc and Sr vsSc (Fig. 11). There are a number of important features in

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Tofuadacite

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line

(230

Th/

232 T

h)(2

26R

a/23

0 Th)

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Tofua

Tofuadacite

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rapidfractionation?

modelamphibolite

melts

Fig. 9. New Fonualei U-series isotope data shown on plots of(238U/232Th) vs (230Th/232Th) and (226Ra/230Th) vs SiO2. Basaltic an-desites from Late Island (Supplementary Data Electronic Appendix2; Turner et al., 1997, 2000) andTofua Island (Caulfield et al., 2012) areshown for comparison. All three suites lie within a broadly horizontalarray and have 226Ra excesses implying that the time elapsed sincefluid addition to the mantle wedge (and by inference partial melting)occurred less than 8 kyr ago. The range of likely (226Ra/230Th) ratioscalculated for amphibolite melts (the maximum assuming the pres-ence of residual zircon) shown in (b) is from Berlo et al. (2004).Symbols as in Fig. 6.



these diagrams; the positive correlation between Zr and Scis particularly unusual. Peridotite melting can result in apositive correlation between Sc and Zr (or Sc and Sr)that passes through the more primitive end of the basalticandesite array in Fig. 11. However, Sc behaves as a compat-ible element during subsequent crystallization in mostmagmatic systems owing to its compatibility in pyroxene,and direct measurement of Sc concentrations in mineralseparates led Ewart et al. (1973) to infer pyroxene partitioncoefficients in the range 1·4^8·7 for the Tonga rocks [theFonualei rocks analysed by Ewart et al. (1973) were pre-dominantly low-Sc types]. In contrast, Zr will be an in-compatible element in the absence of zircon (as in theTonga rocks). Thus, a model andesitic fractionation vectorhas a negative slope (see Fig. 11a). Interestingly, the highplagioclase mode in the dacites results in a bulk partitioncoefficient for Sc that is less than unity (Table 3) and so awithin-dacite fractionation vector does have a positiveslope in Fig. 11a. However, the Sr partition coefficient forplagioclase is predicted to increase from 1·36 at 12008C inAn90 to 1·61 in An80 at 10008C (Blundy & Wood, 1991)and this increase is consistent with the measured Sr parti-tion coefficients in the Tonga rocks (Ewart et al., 1973).Consequently, the within-dacite fractionation vector has anegative slope, orthogonal to the Fonualei dacite Sr^Scarray, as shown in Fig. 11b.Nevertheless, the model curves in Fig.11 show that evolu-

tion from the Fonualei basaltic andesite through the andes-ite to one end of the dacite array is well modeled by�50% fractional crystallization of a mineral assemblagesimilar to that observed in the basaltic andesite and asdetermined by Ewart et al. (1973) using the least-squaresmodeling approach. Furthermore, because the Fonualeidacite array is broadly parallel to that of theTofua basalticandesites in Fig. 11a, it could, in principle, be reproducedby a series of similar fractional crystallization curves ema-nating from a range of parental compositions within theTofua array. However, apart from this requiring a ratherfortuitous set of circumstances and many magma batchesin just 165 years, such a model fits less well with the rela-tionships in Fig. 11b, where the Fonualei dacite and Tofuaarrays are strongly inclined to each other. Thus, we con-clude that the positive trends exhibited by the Fonualei da-cites in both Fig. 11a and b cannot result from simple,closed-system fractional crystallization of a single batch ofbasaltic andesite magma.

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Fig. 10. Plots comparing data from the volcanic rocks from Fonualei,Late Island and Tofua with the results of 0·1GPa MELTS models(Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998) for various initialH2O contents (shown as black curves) and using a mafic Tofuasample (TF1278) for the starting composition with Fe2O3/FeO¼ 0·35(tick marks indicate 258C increments). (a) Temperatures, inferredfrom pyroxene thermometry, vs SiO2 for samples from Fonualei andTofua (data from Caulfield et al., 2012). (b, c) TiO2 vs Fe2O3* andAl2O3 vs CaO for Fonualei and Late Island rocks plus the field forTofua rocks (data from Caulfield et al., 2012). Grey dashed curvesshow the results of higher-pressure MELTS models at 0·3, 0·5 and1GPa (anhydrous). Also shown in (a) is a field for experimentalmelts of arc amphibolites with basaltic andesite compositions under

water undersaturated conditions (data from Beard & Lofgren, 1991;Rapp et al., 1991; Rushmer, 1991; Patin‹ o Douce & Johnston, 1995;Rapp & Watson, 1995). These do not even overlap the parameterspace shown in (b) and (c), providing evidence that the dacites didnot form by partial melting of amphibolites. Symbols as in Fig. 6.



0

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andesitefractionation

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wedgemelting

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Fig. 11. Plots of Zr and Sr vs Sc for Fonualei, Late Island andTofua rocks with batch partial melting and Rayleigh crystal fractionation vectors(seeTable 3 for model parameters; italic numbers alongside the curves indicate the per cent melting or fractionation).The mantle wedge meltingcurve passes through the more primitive basaltic andesites from Late Island andTofua. The curve labeled ‘andesite fractionation’ shows the tra-jectory of fractional crystallization of the Fonualei basaltic andesite, which passes through the andesite to reach the dacite array.The ‘dacite frac-tionation’ vector shows that fractional crystallization cannot replicate the positive sloped Zr^Sc and Sr^Sc arrays within the dacites. Theamphibolite melting curve shows that partial melting of an amphibolite with a composition similar to some of the Late Island andTofua basalticandesites can also reach the Fonualei dacite array but this is precluded for a number of reasons discussed in the text. Mixing between productsof fractional crystallization of a series of basaltic andesites of variable composition is the preferred interpretation for the dacite array (see textfor discussion). Symbols as in Fig. 6.



Partial melting of amphiboliteAn alternative to fractional crystallization is that theFonualei rocks result from partial melting of amphibolitesin the arc crust. Rocks from Eua Island suggest that theTonga arc is built upon rifted Pacific oceanic crust ofEocene age (Ewart & Hawkesworth, 1987) and so theIndian MORB Pb and mantle-like 187Os/186Os (Turneret al., 2009) observed in the Fonualei dacites precludestheir origin via melting of this material. However, theseisotope constraints would be mitigated if the dacites resultfrom melting of very young amphibolites that formedwithin the arc crust from basaltic andesites that wereintruded during earlier magmatic episodes (e.g. Petford &Gallagher, 2001; Smith et al., 2003, 2006). A model inFig. 11 indicates that �30^40% partial melting of an am-phibolite similar in composition to some Tofua basaltic

andesites could yield appropriate dacitic compositions(seeTable 3 for model details). However, layers with appro-priate seismic velocity for mafic amphibolite tend to belocated in the lower crust of arcs (e.g. Fliedner &Klemperer, 2000), yet crystallization at Fonualei is inferredto have occurred at shallow levels (see Fig. 10). In addition,Fig. 10a suggests that the temperatures inferred from pyr-oxene thermometry are significantly higher than thoserequired to form dacitic melts by dehydration melting ofarc amphibolites of basaltic andesite composition (Beard& Lofgren, 1991; Rapp et al., 1991; Rushmer, 1991; Rapp &Watson, 1995). Likewise, experimental amphibolite meltshave lower Fe2O3 and CaO than the dacites (Fig. 6band c) and do not even overlap the parameter spaceshown in Fig. 10a and b. Finally, because Sc will becompatible in residual pyroxene, partial melting of am-phibolite also results in negatively sloped curves in Fig. 11aand b that cannot replicate the Fonualei dacite array.Conceivably, each dacite could reflect partial melting of adifferent amphibolite of appropriate composition such asto generate the dacite arrays in Fig. 11, but we considerthis highly improbable.Brophy (2008) also explored means by which to distin-

guish the geochemical effects of fractional crystallizationfrom those of amphibolite partial melting. His approachshowed that changes in mineral modes and REE partition-ing with increasing SiO2 allow amphibolite partial meltingand low- and high-pressure fractional crystallization to bedistinguished on plots of La andYb vs SiO2. In particular,partial melting of amphibolite results in horizontal ornegative arrays on such diagrams (Brophy, 2008). As canbe seen in Fig. 12, although a couple of Fonualei sampleslie close or parallel to the model amphibolite meltingcurve, the majority of the dacites lie well above it and donot form a horizontal trend. Moreover, the low-REE sam-ples in Fig. 12 are not the low-Sc dacites, and althoughdecreasing Dy/Yb ratios have been suggested to be a hall-mark of amphibole (Davidson et al., 2007) there is no suchdecrease within the Fonualei dacites (not illustrated), evenrelative to the basaltic andesite.Finally, both low-Sc (F0820) and high-Sc (F0823) da-

cites have (226Ra/230Th) �3·9. Indeed, all of the daciteshave (226Ra/230Th) significantly higher than that predictedfor partial melts of amphibolite, even if these melts wereformed in the presence of residual zircon (Berlo et al.,2004). It might be envisaged that an amphibolitic sourcecould be young enough (58 kyr) to retain mantle-deriveddisequilibria. However, the 238U excesses of the dacitesare so high that even after 10 kyr the (230Th/232Th) ratiosof a putative precursor amphibolite would have risen to1·70^1·74. It is clear from Fig. 9a that all of the daciteshave uniform (230Th/232Th) ratios �1·65 and the basalticandesite and andesite have lower values (1·61^1·64). Thesecombined observations provide strong evidence that the

Table 3: Partition coefficients*, modesy and concentrations

used in geochemical modeling

Phase or concentration Mode

(%)

Sc Sr Zr Source or

sample no.

Mantle wedge melting

olivine 59 0·3 0·001 0·002

orthopyroxene 27 0·6 0·002 0·001

clinopyroxene 14 1·5 0·095 0·089

CO (ppm) 26 21 3 Turner et al. (1997)

Andesite fractionation



plagioclase 57 0·0 1·360 0·040

magnetite 7 2·0 0·520 4·760

CO (ppm) 44 311 32 F0813

Dacite fractionation



plagioclase 80 0·1 1·600 0·040

magnetite 4 3·9 0·520 4·760

CO (ppm) 32 333 53 F0801A

Amphibolite melting


amphibole 50 2·9 0·420 0·310

plagioclase 25 0·1 1·360 0·040

CO (ppm) 42 159 22 06TF49Az

*Based on Blundy & Wood (1991), Rollinson (1993) andMcDade et al. (2003).yBased on modal proportions and least-squares models re-ported by Ewart et al. (1973).zFrom Caulfield et al. (2012) with 42 ppm Sc.



Fonualei dacites do not owe their origin simply to variabledegrees of partial melting of amphibolite(s).

Magma mixingBecause neither crystal fractionation nor amphibolite melt-ing alone seem able to explain the Fonualei dacite arraywe suggest that open-system processes provide the best ex-planation of the data. This could occur via combined frac-tional crystallization and assimilation of amphibolitewall-rocks (e.g. DePaolo, 1981). However, Sc is not corre-lated with any of the radiogenic isotope ratios and thehighly linear nature of the dacite trace element trends sug-gests that mixing of two magmas may be more applicable.On a plot of Zr/Ce vs Th/Ce (on which mixing is linear)the Fonualei dacites again define a tight linear array(Fig. 13). What is also striking is that this array is nearly

perpendicular to that formed by the Late Island andTofua basaltic andesites, implying that a different processgoverns the dacite petrogenesis. The positively slopedarray of the basaltic andesites is inferred to largely reflectvariations in parental magma composition arising fromvariations in the extent of wedge depletion and theamounts of sediment melt added (e.g. Ewart &Hawkesworth, 1987; Turner et al., 1997); we develop amodel for the dacite array below.The model curves in Fig. 11 suggest that the dacite array

could have developed from mixing between (1) two subtlydifferent magmas resulting from fractional crystallization,(2) the products of partial melting of two subtly differentamphibolite melts, or (3) a fractionated magma and anamphibolite melt. In principle, all three models are pos-sible. However, the major element compositions (includinginferred H2O contents) of both the high- and low-Scdacite end-members are virtually identical and both typesoccur together in the same stratigraphic unit. It seems im-probable that two very different processes could yield simi-lar results so we do not favor model (3) and, for thereasons detailed above we consider amphibolite meltingand model (2) unlikely as well.Oxygen isotope data for Tonga^Kermadec arc rocks

cluster around mantle values (Turner et al., 2009) andHaase et al. (2011) have used this to argue for theKermadec silicic rocks being derived by fractional crystal-lization. Thus in our preferred model (1), the Fonualei da-cites reflect mixing between two different products offractional crystallization of parental basaltic andesite

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SiO2

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low-PFC

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high-P FC

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(a)

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Fig. 12. (a) Variation of La andYb vs SiO2 for Fonualei, Late Islandand Tofua volcanic rocks along with vectors calculated by Brophy(2008) for upper crustal (low-P) and lower crustal (high-P) crystalfractionation (FC) compared with amphibolite partial melting. Greylow-P fractional crystallization vector in (a) schematically illustratesthe effects of crystal fractionation of an assemblage dominated byplagioclase (as observed in the Fonualei dacites). Symbols as in Fig. 6.

2

3

4

5

6

7

0.02 0.03 0.04 0.05 0.06

Th/Ce

Zr/

Ce Tofua

basaltic-andesites

dacites

Fig. 13. Variation of Zr/Ce vs Th/Ce for Fonualei, Late Island andTofua volcanic rocks. The composition variation within the basalticandesite (Tofua and Late Island) array is inferred to largely reflectsource variation (from more depleted mantle to mantle more enrichedwith sediment melt). The Fonualei dacites form a perpendiculararray inferred to reflect crystal fractionation and mixing (see text fordetails). The Tofua dacite plots off scale at Th/Ce¼ 0·021 and Zr/Ce¼ 7·57. Symbols as in Fig. 6.



resulting from different degrees of partial melting (seeTamura et al., 2011). The variation in the basaltic andesitesthat underlie the dacites on Fonualei and those erupted onLate Island andTofua provides ample evidence that a suffi-cient range of mafic parent magma compositions is avail-able in the system. Magma mixing is efficient in thisscenario because both magmas will have similar tempera-tures (around 10008C; see Fig. 10a) and viscosities. Thecorollary is that little textural or chemical disequilibriumof the phenocrysts is predicted in these rocks (see Couchet al., 2001).

Timescale constraintsThe eruptive history of Fonualei Island, field relationshipsand isotope data allow us to place some time constraintson the proposed petrogenetic model. Turner et al. (2000)argued that the 226Ra excesses in the Tonga^Kermadecarc lavas reflect fluid addition rather than partial melting.Thus the U^Th^Ra isotope data indicate that the finalepisodes of fluid addition from the slab and, by inference,partial melting to produce the parental magmas occurredless than 8 kyr ago. Moreover, because all of the daciteshave (226Ra/230Th)¼ 2·9^3·9, fractional crystallizationfrom parental magmas similar to the Late Island andTofua basaltic andesites must have occurred in less than a

few millennia. The lower (226Ra/230Th) ratios of theFonualei basaltic andesite and andesite suggest either thatthese rocks are �1^2 kyr old, consistent with age data forthe lower basaltic andesites on Tofua (Caulfield et al.,2001), or else that they fractionated from parental basaltover a more protracted period than subsequent batches ofmagma.Most of the dacitic lavas are interpreted to have erupted

within the last 165 years, with an eruption periodicity of afew decades. Such a short time span suggests that theywere all erupted from a single batch of mixed magma; ahypothesis supported by the limited geochemical variationof the rocks erupted since AD 1846. In contrast, the two da-citic lavas that were probably erupted during or directlyafter AD 1846 lie at opposite ends of the dacitic mixingtrend identified by the Sc and Zr data (Fig. 11).With the ex-ception of the sampled pyroclastic material (i.e. F0806),these lavas, F0802 and F0808, have the lowest and highestZr and Sc values respectively. The samples of pyroclasticmaterial comprise both xenolithic (i.e. clasts) and primarymaterial and hence samples from the same event can pro-vide variable geochemical data for rocks not part of thecurrent magma batch. Notwithstanding this caveat, theevidence suggests that the two inferred dacitic magmasmixed on a timescale of �165 years.

daciteandesite

Fonualei Island

~ 2-6 km basaltic andesiteto dacite< few kyr

< 165 yearsparental magmas:< 8 kyr since fluid

addition andmantle melting

dacite ascent ~ decades

fractionalcrystallisation

and mixing

Fig. 14. Schematic illustration of the preferred interpretation of the evolution of Fonualei (not to scale). Parental basaltic magmas with a smallrange of initial compositions are inferred, from U^Th^Ra systematics, to have been derived from the mantle less than 8 kyr ago. These stallat a depth of �2^6 km where they undergo fractional crystallization and subsequent mixing and eruption over the last 165 years.



A WORKING MODEL FORFONUALEIIn Fig. 14 we illustrate our working model for FonualeiIsland. The initial stages of growth involved eruptions ofbasaltic andesite and andesite similar to those erupted onadjacent islands, such as Late Island and Tofua. The U^Th^Ra isotope data indicate that less than 8 kyr haselapsed since fluid addition to the source of thesemagmas. Sometime prior to AD 1846, at least two freshbatches of basaltic andesite were emplaced between 6 and2 km beneath the volcano. Because of a reduced mantleflux in this region of the arc (e.g. Keller et al., 2008), thesecooled to around 10008C, crystallized and fractionated toproduce the two end-member dacitic compositions.Eruption of these end-member dacitic lavas commencedin 1846. Subsequently erupted dacites represent variablemixtures of these end-members. As discussed above, the165 year eruptive episode may reflect the timescale ofascent and mixing of these two magma batches in the con-duit beneath the volcano.In applying his method of discrimination to intra-

oceanic arc lavas, Brophy (2008) concluded that fraction-ation of basalt, as opposed to amphibolite melting, is thedominant mechanism for producing silicic magmas. Wegenerally concur with this view, not least because theU-series disequilibria observed in most arc lavas wouldseem to require a mantle origin and rapid differentiation.However, the depleted nature of the Fonualei dacites andthe presence of 8·5wt % Fe2O3 yet only 1·1wt % K2O,45 ppm Zr, 51ppm Nb and 510 times chondritic REEneeds to be emphasized; dacites dredged from the numer-ous silicic submarine edifices along the arc share the samemajor element characteristics (Graham et al., 2008).Clearly this is not ‘continental crust’ whose origin musttherefore involve different processes and/or tectonicsettings.In addition to our study of Fonualei, a number of other

recent studies have reached the same conclusion, includingthose of Anatahan in the Marianas (Reagan et al., 2003;Wade et al., 2005) and Katmai in the Aleutians (Turneret al., 2010). The conclusion from these studies that daciticmagmas are produced by fractional crystallization appearsat odds with some aspects of the so-called ‘hot-zone’modelof Annen et al. (2006). Rapid differentiation will be favoredby cooler wall-rocks and the magmatic system that pro-duced the Fonualei dacites appears to be located in themid- to upper crust, rather than at a zone of underplatingat the base of the crust (estimated to be 12^20 km thick inTonga; D. Wiens, personal communication). Turner et al.(2010) reached a similar conclusion for evolution frombasalt to rhyolite at Katmai. In addition, the compositionof the erupted products at Fonualei has become more silicicover time whereas the hot-zone model predicts that themost silicic magmas would be formed early on when the

hot-zone is at its coolest. Finally, the switch from mafic tofelsic volcanism at Fonualei has occurred within the last 8kyr and potentially within the last few millennia (by ana-logy with ages on Tofua) and this is much more rapidthan predicted by numerical models of hot-zone evolution(e.g. Annen et al., 2006). Future work should be directed ataddressing these debates for other volcanic centres.

ACKNOWLEDGEMENTSWe thank Kelepi Mafi and theTongan Survey for samplingpermission and logistical assistance, and Jeff aboard theM.V. Hakula for safe transport to and from FonualeiIsland. James Cowlyn assisted with fieldwork and initialsample preparation. We thank Jim Gill, Erin Todd andYoshi Tamura for highly insightful reviews that helped sig-nificantly to improve the paper, and acknowledge the edi-torial guidance of John Gamble. This is contribution 783from the Australian Research Council National KeyCentre for the Geochemical Evolution and Metallogenyof Continents (http://www.gemoc.mq.edu.au).

FUNDINGThis study used instrumentation funded by AustralianResearch Council (ARC) LIEF and DEST SystemicInfrastructure Grants, Macquarie University andIndustry. This research was directly supported byAustralian Research Council grant DP110103284 to T.R.and S.T. and an Australian Research Council ProfessorialFellowship to S.T.

SUPPLEMENTARY DATASupplementary data for this paper are available at Journalof Petrology online.

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