North Tongan high-Ca boninite petrogenesis: The role of samoan plume and subduction zone-transform...

e J. Geodynamics Vol. 20, No . 3, pp . 219-241 ,1995 Copyright © 1995 Elsevier Science Ltd

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e NORTH TONGAN' HIGH-Ca BONINITE PETROGENESIS: THE ROLE OF SAMOAN PLUME AND SUBDUCTION ZONE-TRANSFORM FAULT TRANSITION

LEONID V. DANYUSHEVSKY,' ALEXANDER V. SOBOLEV2 and TREVOR J. FALLOON' 'Department of Geology, University of Tasmania, GPO Box 252C, Hobart , Tasmania , 7001, Australia and 2Yernadsky Institute of Geochemistry and Analytical Chemistry, Kosygin Street 19, Moscow 117975, Russia

(Received 14 October 1994; accepted in revised form 11 April 1995)

Abstract-High-Ca boninites are characterized by high CaO/Al203 values ( >0.75). In the North Tonga forearc , they occur close to the northern termination of the trench , where it swings west into a major transform fault. Contemporaneous OIB-like lavas have been found close to the boninites on the forearc slope into the trench . The Tongan boninites are characterized by: olivine up to F094 ;

a range of (La/Yb)N 0.4-13 in rocks with very similar major element compositions; isotopic composition similar to Samoan OIB regardless of La/Yb values ; and a correlation of H20 content of primary melts with the degree of enrichment, the most depleted varieties being most H20 enriched with H20/K20 > 30. The most enriched varieties have olivine up to F0 93 and are identical to adjacent OIB-like lavas in terms of incompatible element and isotopic ratios. The primary Tongan high-Ca boninite melts had MgO contents - 24 wt% and very high temperatures ( -1470°C at the depth of magma generation).We argue that the open edge of the mantle wedge beneath the transform fault has allowed the Samoan plume to intrude abov e the subducted slab . The intruding plume has been depleted by extraction of OIB magmas, leaving a hot, dry, refractory peridotite residue . Slab-derived fluids react with the plume residue , producing boninite primary magmas at 15-17 kbar. During ascent to the surface , some of these magmas mix with earlier formed OIB-like melts , producing a range of variably-enriched boninite melts . As a result , more enriched boninites have lower H 20 contents and Mg# values, but the same isotopic composition as the more depleted boninites in this magma series .

INTRODUCTION

Boninites, a subduction-related magmatic suite, are characterized by relatively high Si02, H20 and low Ti02 contents compared to tholeiitic suites (Le Maitre, 1989). They are also characterized by the absence of plagioclase in rocks more mafic than andesites, and by the presence of very magnesian olivine phenocrysts (up to F094 , Crawford et al., 1989). Where they occur in modern island-arc systems, boninites have been found exclusively in the forearc regions. It has been proposed that these rocks originate from mantle which has been variably

219

220 Leonid V. Danyushevsky et al.

depleted in the basaltic component compared to a typical mid-ocean ridge basalt (N-MORB) source, but is always extremely depleted in incompatible elements (Crawford et al., 1989). Prior to or during boninite petrogenesis these depleted mantle sources are believed to be metasomatized by subduction-related components. Variations in the degree of depletion of these boninite source peridotites in the basaltic component are reflected by a wide range of CaOIAlz03 values (0.4-0.85, Crawford et al., 1989) in primitive boninites (those unaffected by clinopyroxene fractionation). On the basis of CaOI Alz03 , boninites have been divided into two groups: low-Ca and high-Ca boninites (Crawford et al., 1989), with a boundary set at CaOIAlz0 3 ~0.75. As was demonstrated by Falloon et al. (1989), the CaO/Alz03 values of boninites reflect the presence or absence of clinopyroxene in their mantle source. The low CaOI Alz03 values of low-Ca boninites imply that clinopyroxene was not present in their mantle sources and therefore they are interpreted as being produced from relatively more depleted sources than high-Ca boninites. Variations in CaOIAlz0 3 values and TiO z contents recorded in several boninite suites (e.g. Howqua, Australia (Crawford and Cameron, 1985), Izu-Bonin forearc (Umino, 1986; Pearce et al. , 1992; Taylor et al. , 1992 and 1994)) have been explained as a result of progressive source depletion during boninite magma genesis.

The nature of depleted mantle sources involved in boninite petrogenesis has been a subject of debate. Most workers (Crawford et al. , 1989) propose a depleted mantle wedge residual from prior extraction of MORB. However , it was also suggested (Sharaskin et al., 1983; Falloon and Crawford, 1991) that at least some high-Ca boninites can originate from a refractory ocean island basalt (OIB) mantle source. A possible involvement of hot-spot-related melts as enriching components in boninite petrogenesis has been proposed in several studies (Hickey and Frey, 1982; Rogers et al., 1989; Falloon and Crawford, 1991 ; Stern et al. , 1991; Kostopoulos and Murton, 1992).

High-Ca boninite suites which are not spatially associated with low-Ca boninites have so far been described from three regions: in the late Cretaceous Troodos ophiolite , Cyprus (Cameron, 1985; Sobolev et al., 1993), and in two modern island-arc systems: at the northern termination of the Tonga Trench (Sharaskin et al., 1983a; Falloon et al., 1987) and at the southern termination of the Vanuatu Trench (Sigurdsson et al. , 1993). There is no evidence for involvement of mantle sources severely depleted in basaltic component (i.e. clinopyroxene absent in the mantle source) in the petrogenesis of high-Ca boninites which are not associated with low-Ca boninites. In addition to the typical boninite compositional features, high-Ca boninite suites from the above three localities a~e characterized by:

-Abundant olivine phenocryst-rich rocks (Sobolev et al. , 1993; Sigurdsson et al., 1993; Sobolev and Danyushevsky , 1994).

-A large temperature interval of olivine-only crystallization (up to 150°C) (Sobolev et al., 1993; Sobolev and Danyushevsky, 1994).

-----~--------- - - -

North Tongan high-Ca boninite petrogenesis 221

-A correspondingly large difference between compositions of the most magnesian olivines and low-Ca pyroxenes: no orthopyroxenes more magnesian than Mg# = 90 (Mg# = Mg/(Mg + Fe2+)) have been found in high-Ca boninites (Sobolev et al., 1993; Sigurdsson et al., 1993; Sobolev and Danyushevsky, 1994), whereas orthopyroxenes and clinoenstatites as magnesian as Mg# = 95 are common in low-Ca boninites (Crawford, 1980; Walker and Cameron, 1983; Cameron, 1989; Ohnenstetter and Brown, 1992; Taylor et al., 1994).

-By the common presence of very magnesian clinopyroxenes (Mg# up to 92) , which reflects the high CaO content of these melts (Sobolev et al., 1993; Sobolev and Danyushevsky, 1994).

In all three locations, high-Ca boninite lavas overlie thick sequences of peridotites, gabbros and dykes which, in Cyprus, form the Troodos ophiolite.

In the two modern settings peridotite-gabbro-dyke associations comprise the forearc basement (Sharaskin et al., 1983a; Sigurdsson et al., 1993). The tectonic environment of both occurrences is characterized by two major features:

-Termination of the trench and a transition from subduction to transform tectonics.

-Propagation of a back-arc spreading system into the forearc (Falloon et al., 1987; Falloon and Crawford, 1991; Sigurdsson et al., 1993).

On the basis of experimental studies (Duncan and Green, 1987; Crawford et al., 1989; Umino and Kushiro, 1989; and references therein), it is now widely accepted (Crawford et al. , 1989; Pearce et al., 1992; Hawkins, 1994) that boninite petrogenesis requires temperatures of 120o-1350°C and pressures below 10 kbar. Such temperatures are several hundred degrees higher than those postulated in geophysical models for the thermal structure of the mantle wedge under the modern forearc regions (Crawford et al., 1989). Because of this, various competing models of boninite petrogenesis (Pearce et al., 1992) suggest either a source for extra heat in the forearc region , or a specific tectonic setting during boninite petrogenesis, such as initia tion of subduction. A model proposed to explain the petrogenesis of modern high-Ca boninite suites (Crawford et al., 1989; Falloon and Crawford, 1991), invokes ascending diapirs caused by the propagation of the back-arc spreading system into the forearc, as a source of heat required to produce these melts from the depleted , cold, Il .Ovbearing, metasomatized shallow mantle wedge .

Recent experimental studies of melt inclusions in olivine phenocrysts from Tongan (Sobolev and Danyushevsky, 1994) and Troodos (Sobolev et al., 1993) high-Ca boninites, devoted to quantitatively establishing primary melt compositions and crystallization temperatures, have demonstrated that these primary melts were characterized by high MgO (21-24 wt%) and moderate H20 (- 2 wt%) contents, and crystallization temperatures of 135O-1380°C. A highpressure experimental study (Danyushevsky et al., 1994) has demonstrated that such melts are in equilibrium with a harzburgite mantle residue at 15-17 kbar and 143O-1470°C. Such temperatures we believe are significantly higher than can

--- - - -


be achieved by re heating of the cold mantle wedge, as proposed in the current model of high-Ca boninite petrogenesis (Crawford et al., 1989; Falloon and Crawford, 1991). Therefore the aim of this paper is to present a new model for the origin of high-Ca boninites which avoids this problem. The high temperatures of the high-Ca boni nite primary melts in our proposed model are due to their derivation from hot OIB mantle sources. The transition from subduction to transform tectonics is proposed as a major factor in allowing the intrusion of hot OIB mantle in the mantle wedge and therefore the mantle source for high-Ca boninites was the depleted OIB diapir itself. Green et al. (1987) and Ewart and Hawkesworth (1987) have also suggested the possible role of subductiontransform transitions in arc and backarc magma genesis for the Tonga-Lau supra-subduction zone.

In this paper we also present new mineralogical, geochemical (including isotopes) and K/ Ar age data for the North Tongan boninites.

A SUMMARY OF THE NORTH TONGAN BONINITE GEOLOGY AND GEOCHEMISTRY

The geological setting, petrology, geochemistry and mineralogy of North Tongan high-Ca boninites have been presented in a number of papers (Sharaskin et al., 1983a,b; Falloon and Green, 1986; Danyushevsky and Sobolev, 1987; Falloon et al., 1987, 1989; Vysotsky, 1989; Falloon and Crawford, 1991; Danyushevsky et al., 1992; Sobolev and Danyushevsky, 1994). Some new mineralogical and chemical data are presented in Tables 1, 2, 3 and 4.

Geology and petrography The northern termination of the Tongan Ridge is a tectonically complex area.

Several tectonic processes such as propagation of back-arc spreading centers, subduction and large-scale strike-slip faulting have been identified (Tiffin, 1993). The bend at the northern end of the To nga tre nch reflects the intersection of the Tonga-Kermadec subduction zone with the transform fault which separates the Pacific plate from the Lau and the North Fiji basins (Tiffin, 1993). The subducted slab is presently located -100 km beneath the area of boninite occurrence (lsacks and Barazangi, 1977).

High-Ca boninites, arc tholeiites and OIB-like lavas were dredged from the northern termination of the Tonga trench during cruises of R/V "Kallisto" in 1982 (Falloon et al., 1987) and R/V "Natsushima" in 1984 (Sharaskin et al., 1983a) (Fig. 1). Boninites occupy a 50 km-long section of the upper inner part of the trench slope . Arc tholeiites were dredged 50 km to the east of the boninite outcrops (Sharaskin et al., 1983b) and alkaline OIB-like basalts, which have been shown to originate from the adjacent Samoan plume (Zlobin et al., 1991a), were recovered 60 km to the west of the boninites (Sharaskin et al., 1983a) at the same depths. Two boninitic lavas (sample 3-23 from station 21 and 5-27 from

Table 1. Compositions of spinel inclusions in olivine phenocrysts

Sample 5-25 5-25 5-25 5-25 5-28 5-28 5-28 6-3 6-3 7-18 7-18 3-44 3-44 3-44 3-44 5-4 5-24 No 30 44A 53 67 11 7B 8 23 4 35A 37 18 32 33 37 13 4A

Spinel inclusions Z ..,Sial 0.10 0.06 0.06 0.07 0.07 0.04 0.12 0.07 0.11 0.19 0.14 0.09 0.08 0.09 0.07 0.08 0.08 0

SoTi02 0.15 0.12 0.15 0.09 0.08 0.19 0.08 0.20 0.27 0.42 0.32 0.09 0.24 0.09 0.19 0.09 0.08 -'lAl203 7.03 4.88 5.08 6.62 6.03 8.97 6.43 6.85 7.32 8.78 7.33 6.80 10.56 6.16 9.03 5.94 6.17

::l 0

FeO* 15.38 25.17 21.75 14.16 15.85 24.54 14.70 17.17 19.73 22.06 20.19 19.14 28.98 19.30 28.91 15.44 14.36 (JQ co

MnO 0.14 0.19 0.26 0.18 0.21 0.19 0.13 0.15 0.16 0.30 0.20 0.22 0.15 0.17 0.19 0.13 0.17 e ::rMgO 14.77 9.96 11.23 15.35 13.95 11.48 14.45 13.68 12.39 12.77 12.50 13.03 11.50 13.15 10.19 14.26 14.05 tiQ.

Cr203 61.47 58.48 60.54 63.05 63.29 53.68 62.97 61.55 58.54 53.32 57.80 59.42 47.72 60.55 50.30 64.06 64.44 ::r nNiO 0.04 0.00 0.07 0.17 0.06 0.09 0.13 0.08 0.10 0.16 0.12 0.08 0.11 0.11 0.08 0.09 0.13 co

Cr# 85.43 88.94 88.88 86.47 87.56 80.06 86.79 85.77 84.29 80.29 84.10 85.43 7.20 86.83 78.89 87.86 87.51 00 ::l

Host olivine 2: Si02 41.34 40.03 40.47 41.56 40.93 40.22 40.71 40.66 40.69 40.18 40.68 40.97 39.86 40.94 39.63 40.99 41.23 <> FeO 6.43 12.92 10.75 6.16 6.95 12.02 6.39 7.84 8.59 9.50 8.13 8.34 11.67 8.38 14.01 6.52 6.59 '0

~MnO 0.06 0.28 0.15 0.09 0.09 0.20 0.14 0.08 0.11 0.15 0.09 0.11 0.16 0.14 0.22 0.04 0.04 '" (3 (JQMgO 52.18 46.82 48.54 52.25 50.88 46.90 51.01 50.45 50.00 49.05 50.57 50.46 47.48 50.52 45.69 50.76 51.17

CaO 0.13 0.12 0.17 0.13 0.13 0.19 0.16 0.13 0.13 0.18 0.17 0.17 0.21 0.16 0.19 0.13 0.11 '" e

Cr203 0.17 0.19 0.19 0.18 0.22 0.13 0.16 0.12 0.14 0.11 0.08 0.10 0.13 0.08 0.14 0.15 0.16 '"on''"NiO 0.24 0.19 0.18 0.32 0.23 0.18 0.40 0.28 0.21 0.29 0.29 0.19 0.16 0.26 0.22 0.35 0.34 Fo 93.54 86.60 88.95 93.80 92.88 87.43 93.44 91.98 91.21 90.20 91.73 91.52 87.88 91.49 85.33 93.28 93.26

FeO* = total iron as FeO . Cr# =Cr/(Cr+ AI) . Mineral analyses performed using a Cameca SX-50 microprobe at University of Tasmania using WDS . Operating conditions: 15 kV and 20 nA , counting time 10 sec. Analytical routine and standards are presented in (Sobolev and Danyushevsky, 1994).

N N W

224 Leonid V. Danyushevsky et at.

Table 2. Compositions of melt inclusions in olivine phenocrysts from the Eastern group rocks and estimated primary melt compositions

Melt inclusions Primary melts

Sample 3-44 5-28 5-28 5-28 PM1 PM2 TR No 1 21 7 7A

sro, 59.90 59.12 57.84 56.14 51.4 53.1 49.8 TiOz 0.28 0.22 0.28 0.25 0.1 0.4 0.2 Alz0 3 13.43 13.38 13.65 14.69 7.0 9.6 10.1 FeO* 6.00 5.87 8.10 8.54 9.2 8.7 8.7 MnO 0.14 0.17 0.14 0.13 0.2 0.2 MgO 3.17 4.42 3.87 4.25 24.0 19.2 20.6 CaO 12.84 13.36 12.98 12.50 6.4 7.7 8.4 NazO 1.20 0.97 1.16 1.25 0.6 1.3 0.5 KzO 0.34 0.19 0.15 0.17 0.1 0.6 0.07 HzO 1.93 2.31 1.90 1.91 2.0 1.1 1.7

Fo 89.9 91.6 87.1 86.7 94.0 93.0 94.0 Tcr (0C) 1390 1360 Tm (0C) - 1470 - 1430 Pm (kb) 15-17 15

Fo = composition of host olivines for the inclusions , and equilibrium olivine composition for primary melts ; PMI = Tongan high-Ca boninite primary melt ; PM2 = the most primitive most enriched Tongan high-Ca boninite melt (see text for discussion) ; TR = primary melt of Troodos high-Ca boninites, Cyprus; Tcr = temperature of beginning of crystallization; Tm = temperature of equilbrium with mantle source; Tp = pressure of equilibrium with mantle source. Compositions of Tongan and Troodos primary melts are from Sobolev and Danyushevsky (1994) and Sobolev et at. (1993) respectively. Temperature and pressure of melting are from (Danyushevsky et a/. , 1994). HzO content of melt inclusions was analyzed by ion microprobe technique on Cameca 3F at CRPG, Nancy, France . See Sobolev and Chaussidon, 1995 for the details of technique .

station 23) give 40Arj39Ar ages 1.4 ±0.03 Ma and 2.03 ±0.11 Ma correspondingly (R. A. Duncan, analyst), in agreement with K-Ar ages of Acland et al. (1993). Unpublished K-Ar data from lavas from the North Tongan island of Niuatoputapu give ages of 3 to 3.5 Ma (J. Gill, pers. comm.). Thus the age of the high-Ca boninites is younger than the oldest of the modern Tofua Arc volcanoes, and we can confidently relate the petrogenesis of high-Ca boninites to tectonic processes occurring during the development of the modern Tofua Arc-Lau Basin system.

Boninites and arc tholeiites overlie a forearc sequence consisting of peridotite , gabbro, diabase and plagiogranite (Sharaskin et al., 1983a). Although gabbros with both boninitic and tholeiitic affinities have been identified in the forearc sequence (Sharaskin et al., 1983a), the occurrence of tholeiitic gabbros which give K-Ar ages of 50 + 9 Ma , and other indirect evidence including degree of alteration, bathymetric position and the extent of tectonic reworking (Sharaskin et al., 1983a), all suggest that it is unlikely that boninite volcanic rocks belong to the forearc magmatic complex. Quenched glasses from the boninite suite have been shown to be saturated with H20 , and H20 concentrations correlate with the depth of dredging (Danyushevsky et al., 1993). Probably, boninites were

225 North Tongan high-Ca boninite petrogenesis

Table 3. Trace element contents in melt inclusions and rock samples from the

Eastern group

Sample 3-44 5-28 5-28 No 1 21 rock

Sr 248 61 Nb 1.50 0.3 Zr 14 5 Y 6.3 4.8 La 3.42 1.23 0.42 Ce 6.04 2.36 1.13 Nd 3.51 1.17 0.94 Sm 1.09 0.51 0.36 Eu 0.14 Tb 0.10 Dy 1.42 0.94 Ho 0.17 Er 1.06 0.65 Yb 0.82 0.85 0.54 Lu 0.09

Trace elements in melt inclusions were analyzed using ion microprobe Cameca 4F at MEKHANOBR, St. Petersburg, Russia. See Shimizu and Hart, 1992 and Sobolev et al. , 1993 for details of technique. REE in sample 5-28 were analyzed by standard INAA at Becquerel Laboratories, N .S.W., Australia.

recently erupted on the top of the earlier-formed forearc sequence, the age of which roughly corresponds to the initiation of subduction in this region (Acland et al., 1993; Hawkins, 1994).

The Tongan high-Ca boninite suite consists of a variety of rocks from strongly olivine-phyric (with up to 50% olivine phenocrysts) to olivine + orthopyroxene + clinopyroxene-phyric, two-pyroxene-phyric, two-pyroxene + plagioclase-phyric variants, boninitic andesites, dacites and aphyric rhyolites. All rocks are characterized by an essentially glassy groundmass with varying amounts of microlites. Chemical and mineralogical features of mainly olivinephyric samples containing more than 12 wt. % MgO are summarized in this paper.

Table 4. Isotopic composition of Tongan high-Ca boninites

Sample Group 87Sr!""Sr 143Nd/1-wNd z''''Pb/Z04Pb Z07Pb/z'llpb Z08Pb/Z04Pb

7-18 Western 0.704175 0.512708 19.035 15.622 38.951 6-3 Western 0.704515 0.512708 18.872 15.626 38.853 3-44 Eastern 0.704528 0.512787 18.823 15.586 38.766 5-24 Eastern 0.704422 0.512770 5-25 Eastern 0.704047 0.512961 18.529 15.569 38.355

Sr and Nd data from Falloon et al., 1989 and Falloon and Crawford, 1991. Lead isotopic composition determined at the Research School of Earth Sciences, ANU, Canberra. Dr. J. D. Woodhead , analyst.

---------------------------------- - - - - - -

226 leonid V. Danyushevsky et al.

Fig. I . (a) Schematic locality map of the Lau-Tonga supra-subduction zone, showing the regional setting of the north Tongan boninite localities. Rectangle labelled I covers the area where boninites were dredged, and is shown in more detail in Fig. I(b). Rectangles 2 and 3 correspond to the area of OIB-like and tholeiitic lavas (see text for description) . Thick black lines represent positions of currently active spreading centers in the Lau Basin after (Tiffin, 1993). KTJ = King's triple junction; ClSC and ElSC = Central and Eastern present day Lau spreading centers. Heavy shading is defined by the 500 m water-depth contour. Heavy black spots are islands of the Lau Ridge and Tofua Arc. (b) Northern termination of the Tonga trench after Sharaskin et al., 1983a and Falloon et al. , 1987 rectangle I, (Fig la). Black dots indicate the locations of dredge stations; numbers next to the dots indicate dredge station numbers. The Western group includes stations 24, 25, 26, and 28 (samples discussed in this paper: 6-3 (station 24) , 7-18 (25), 26-1 (26),26-2 (26» ; the Eastern group includes stations 21, 23, 51, 54, and 55 (samples discussed in this paper: 3-44 (21), 5-24 (23), 5-25 (23), 5-28 (23».

Depths are shown next to the isobars in kilometers.


14 0 S+-,------'-------+b)

Fig. 1. Continued.

Chemistry and mineralogy Despite strong similarities in major element (SiOz, Alz0 3 , CaO, MgO)

compositions between all Tongan high-Ca boninites, it has been shown (Sobolev and Danyushevsky, 1994) that they may be divided into two groups on the basis of total iron as FeO (FeO*) content: a Western group , including stations 24, 25, 26,28 (samples discussed in this paper: 6-3 (station 24), 7-18 (25) , 26-1 (26) ,26-2 (26)), and an Eastern group (stations 21,23 ,51 ,54,55, samples 3-44 (21) , 5-24 (23),5-25 (23) , 5-28 (23)). The Western group rocks, which are located closer to OIB-like lavas, are characterized by relatively lower FeO* content.

Another difference between the two groups is evident from compositions of olivine phenocrysts and their Cr-spinel inclusions (Fig. 2). Interestingly, the most magnesian olivine in the more Fe-rich Eastern group rocks is F094 , whereas in the Western group it is F093 • t This difference should indicate either more oxidized crystallization conditions of the Eastern group melts (Le. lower Fez+) , or higher MgO content of their primary melts. The latter is more probable, as spinel inclusions in olivine phenocrysts from both groups have similar Fez

+ /Fe3+

values (Table 1; Sobolev and Danyushevsky, 1994). However, spinels included in olivine phenocrysts from the Western group are systematically less Cr-rich compared to spinels in olivine phenocrysts of the same composition from the Eastern group rocks.

The Western group spinels are also characterized by higher TiOz and generally higher Alz0 3 contents than the Eastern group spinels (Fig. 3a) . TiOz and Alz0 3

contents in Cr-spinels have been shown to correlate with concentrations of these elements in the equilibrium melt (e.g. Maurel and Maurel, 1982; Zlobin et al.,

t Note that the accuracy of magnesian olivine electron-microprobe analyses is - 0.15mol.% Fo. More than 500 olivine phenocrysts were analyzed from the Western group and - 300 from the Eastern group .

- - - - - ----- - - - - - - --- -- - -


90

89 91 93 Fo

75 +-~-r_r--...,---r----r~-.---~

85 87

Fig. 2. Cr# [lOO*Cr/(Cr+ AI)] of spinel inclusions in olivine phenocrysts from the Western and Eastern groups of the Tongan high-Ca boninites versus the composition of their host crystals . Dashed line encircles spinel

inclusions in evolved olivine phenocrysts from sample 5-25.

1991; Arai, 1992). Spinels from both groups included in the most magnesian olivines (Fo > 90) form a tight trend (Fig. 3a), which cannot be accounted for by crystallization, as TiOz varies by an order of magnitude, and implies mixing with a Ti-rich component. This trend extends towards the field of spinel inclusions in magnesian olivines from OIB (Fig . 3b), and the compositions of spinel inclusions from OIB-like lavas dredged from the same area match well the extrapolation of this trend (Fig. 3b). Note that the Western group rocks are located closer to the OIB-like lavas.

Rare earth element (REE) concentrations in Tongan boninites (Fig. 4a) vary significantly. The Western group rocks and melt inclusions in their magnesian olivines have significantly higher light REE (LREE) concentrations compared to the Eastern group rocks and inclusions. This can be explained by different mixing proportions between two components. The depleted component, with (La/Yb)N-0.5 (component D of Sobolev and Danyushevsky, 1994), is believed to result from melting of a depleted mantle source. The enriched component with, (La/Yb)N -16 (component E2 of Sobolev and Danyushevsky, 1994) has been tentatively identified as an incompatible element-enriched silicate melt. The absolute level of heavy REE (HREE) concentrations in the depleted component may vary by a factor of 4.

Increasing proportions of mixing with E2, which can be evaluated by (La)N (Fig. 4), correlate with changes in the ratio values of REE versus large ion lithophile elements (LILE) and high-field strength elements (HFSE) (Fig. 5). The most depleted samples are characterized by the lowest LREE/LILE, Sm/Hf, Eu/Ti and K/Ti values and by the highest LREE/Nb and LREE/Zr values. This indicates that these melts had prominent negative Nb and Zr and positive LILE anomalies compared to REE of a similar degree of incompatibility, which is a typical geochemical feature of subduction-related magmas


0.6 ...----------,,---------,-,a)

o ...

N oE:: 0.2

......2.0 ,.... \ b) \\.. \ 1< TONGAN OIB

l>. HUNTER III 1.6

• TROODOS~ 1.2

\

'\" * \

" \ ... MORBN

"" ri • OIB .... ~/S 0.6 Eo<

0.4 a)

0.0 5 10 15 20 25 30 35 40 45 Al203 wt.5f>

Fig. 3. Ti02 versus A120J relationship of spinel inclusions in olivine phenocrysts. (a) Spinels from North Tongan high-Ca boninites. I and 2 = the Western group spinels in primitive (Fo > 90) and evolved (Fo < 90) olivines, respectively; 3 and 4 = same for the Eastern group spinels. Dashed line encircles spinel inclusions in evolved olivines from sample 5-25. (b) Spinel inclusions in primitive olivine phenocrysts from various magmatic suites. TONGAN OIB = spinel inclusion from OIB-like lavas adjacent to high-Ca boninites from the northern termination of the Tonga trench, F0 R6.5 (Zlobin et al., 199Ia); HUNTER = spinel inclusions from high-Ca boninites from the Hunter Ridge, along the southern edge of the North Fiji Basin, Fo >90 (Sigurdsson et al., 1993; Sigurdsson , unpubl.); TROODOS = spinel inclusions from high-Ca boninites from the Troodos ophio lite, Cyprus (Sobolev et al., 1993) Fo > 90; MORB = spinel inclusions from Atlantic MORB Fo > 89 (Sobolcv et al ., 1989; Dmitriev et al., 1991; Danyushevsky , unpubl.); OIB =spinel inclusions from OIB lavas from Hawaii, Reunion and Iceland, Fo >88 (Nikogosian, 1990; Gurenko, 1991). Solid line = the field of spinel

inclusions from Tongan high-Ca boninites from Fig. 3(a), Fo >90.

(Pearce et al., 1984; Saunders and Tarney, 1984) . In contrast, Ti and Hf concentrations are not anomalous compared to REE abundances.

The most enriched melts, in which incompatible element concentrations were ultimately controlled by the E2 component, had significant negative Ti, Zr and Hf anomalies, whereas Nb and LILE concentrations were approximately that expected from REE abundances. Such melts are characterized by the highest K/Ti values. The relative abundances of incompatible elements in these melts are practically indistinguishable from those in some OIB lavas dredged 60 km west from the Western group boninites (Fig. 6).

According to a study of primary fluid inclusions in phenocrysts (Sobolev and Danyushevsky, 1994), high-Ca Tongan boninites crystallized at low pressures « 1 kbar) in fluid-saturated conditions. The fluid composition was almost pure H20 . This implies that melt H20 content decreased during crystallization, resulting in significantly less H20 contents in quenched glasses than in primitive melts (Sobolev and Danyushevsky, 1994). Unlike quenched glasses, H20

230 Leonid V . Danyushevsky et al.

La Nd Sm Gd Dy Er Yb Ce Eu Tb Ho

E2 • - 7-18 • - 6-3 • - 5-25* * - 26-2 10

'" • - 26-1 o. - 3-44 D • - 5-28

• - 5-24

1 1

D La Nd Sm Gd Dy Er Yb

Ce Eu Tb Ho

Fig. 4. REE co nte nts in Tongan high- Ca boninites (CSAMP) ' Open symbols-me lt inclu sion s in primitive (Fo > 90) olivine phenocrysts, filled symbo ls- wholerock ana lyses . Dat a from Fallo on et al., 1989, Falloon and Crawfo rd, 1991 and Sob olev and Dan yushevsky, 1994 and T abl e 3. Normalization to primitive mantl e (C PM ,

Sun and McDonough, 1989). D and E2 show samples whose co mpositions were mainly determined by co rresponding components . T o enable dir ect comparison of the data , all original analyses were recalculated to the composition of th e primary melt (PMI , Table I). Ca lculations were performed by det ermining the amo unt o f o livine nece ssary to add to (or subtract from) each composition to obtain the primary melt composition. Onl y SiOz, A120 3, FeO , MgO and Ca O contents were used in calculations. Ca lculations were don e using a

mixing technique simila r to th at de scribed in Wright and D oh erty , 1970 .

displays a systematic behaviour in melt inclusions in magnesian olivine phenocrysts (Fo > 90) (Fig . 7). The most depleted primitive melts unaffected by E2 component (lowest KzO/TiOz) have the highest HzO/KzO, HzO/NazO, and HzO/Alz0 3 values which gradually decrease with increasing proportions of E2 component. An important feature of Tongan boninite geochemistry is that absolute HzO contents of the most depleted melts is higher than those of the most enriched (Sobolev and Danyushevsky, 1994). This implies the involvement of another independent component (component E1 of Sobolev and Danyushevsky, 1994) in Tongan boninite petrogenesis. Component E1 contained mainly HzO and LILE and was characterized by very high HzO/KzO values ( ~30) compared to component E2 (2-3). Such high HzO/KzO values , not known for any other mantle-derived primitive melts (Danyushevsky et al., 1993), imply that the E1 component was a fluid rather than a melt. The effect of E1 component on the composition of Tongan boni nites becomes progressively more obscure with the increasing proportion of E2 component , and cannot be seen in the most enriched varieties.

Melt inclusions in more evolved phenocrysts have lower values of HzO/(any element incompatible during olivine crystall ization, e.g. AI, K, Na) compared to more primitive melts with similar KzO/Ti Oz values (Fig . 7) , due to continuous degassing of melts during crystallization (Sobolev and Danyushevsky, 1994).

With the exception of sample 5-25, Tongan high-Ca boninites have similar 143Nd/144Nd, 87Sr/86Sr and Pb isotopic values regardless of their degree of LREE enrichment (Table 4; Fig. 8). The young age of these rocks implies that both 0 and E2 components had very similar isotopic values which are in the range of


~

~0.8 :!-

1.11 T""""---------, 0.8.,....----------, 1.2.,....----------,

0.0 -1-................--.--.-....-.................0.0 -l-.................-..--.-....-.................

.••

•lO-III ~o.• * :!-

•••••

•..: •

0-1-................--.--.-....-.................

8.,....--------,••• •• • 11••. _...

•*•....

0.2 -1=-.................-..--.-....-.................

1.' T""""------:;.--------,

3.,....-------lrO

•

18

• ••

18

E2.1.• .,------=c=--,

• ... 08 ~D+El

4 (fa>x 12

~ ... ~0.8 II

U-

~ ....= ........1

-~

.

Fig. 5. Variations in incompatible element contents in Tongan high-Ca boninites. Star = sample 5-25. Data from Falloon et al. , 1989, Falloon and Crawford, 1991 , Sobolev and Danyushevsky, 1994 and Tabl e 3. All data were recalculated as described in the caption to Fig. 4, and normalized to primitive mantle (Sun and McDonough, 1989). D , E 1 and E2 show samples whose compositions were mainly determined by corre spond

ing components. See text for discussion .

adjacent, contemporaneous alB Samoan rocks (Fig. 8). On the other hand , the isotopic composition of the high-Ca boninites differs significantly from those of (1) the present day Central Lau spreading ridges; (2) the Tofua Arc; and (3) the Valu Fa Ridge, which is located close to the active Tofua Arc significantly south from the northern Tongan area. However, it is significant that the isotopic compositions of basalts from the King's Triple Junction (Fig. 8), which is located in the north-eastern Lau Basin, appear to have been influenced by the Samoan plume. We believe that the isotopic compositions of the high-Ca boninites and King's Triple Junction basalts reflects the involvement of Samoan plume in magma generation processes in the northern Lau Basin. This has also been proposed by Poreda (1985) and Volpe et al. (1988). Note that another enriched mantle domain in the vicinity of the northern Tonga forearc, the remnant Louisville plume, has different isotopic composition (mainly lower 87Sr/86Sr values, Cheng et al. (1987)) and thus is not likely to be involved in the Tongan high-Ca boninite petrogenesis.

One sample, 5-25, is characterized by a number of odd features relative to the other primitive Tongan high-Ca boninites. Melt and spinel inclusions in the most

---- --- --- --


E~ lava

10

0 .1

K Zr Hf Ti

Nb La Ce P Nd SIn Eu Gd Dy Er Yb

Fig. 6. A comparison of incompatible element contents in high-Ca Tongan boninites and adjacent OIB-lik e lavas. Pattern of samples 7-18 represents the most enriched Tongan high-Ca boninites. Pattern 26-1 represents moderately enriched melts and is compiled from wholerock composition of sample 26-1 and compo sitions of melt inclusions in magnesian olivines from this sample. Pattern 5-24 repr esent s the most depleted melts and is compil ed from wholerock compositions of samples 5-24 and 5-28 and compositions of melt inclusions in magnesian olivines from these samples. OIB-like lava is sample 16-22/2 (Zlobin et al. , 199Ib) . Dashed patt ern repr esent s mantle xenoliths from Samo an OIB which were meta somatized by carbonatite melts, sample 90SAVI -I (Ha uri et al., 1993). D , EI and E2 show samples whose compositions were mainly det ermined by

corresponding components. Normali zat ion to prim itive mantle (Sun and McDon ough, 1989).

magnesian olivines from this sample are indistinguishable from those in other Tongan high-Ca boninites (Fig. 2, 3a and 7) , whereas spinel inclusions in the evolved olivine phenocrysts are strikingly different, having much lower Al20 3

content (Fig. 2 and 3a). The REE pattern of sample 5-25 is also different from the other Tongan boninites (Fig. 4b), being essentially flat. This sample has similar isotopic composition to Tofua Arc lavas (Fig. 8). Chemical features of this sample may reflect interaction of high-C a boninite melts or mantle sources with the modern Tofua Arc magmatic system, although we realize that such a hypothesis can hardly explain low Al20 3 contents of spinel inclusions in lowforsterite olivine phenocrysts from this rock . Unfortunately , the small size of the sample itself precludes any detailed chemical, mineralogical and experimental study of melt inclusions, which are necessary to clarify its origin.

The primary melt composition As different proportions of E2 component are recorded in melt inclusion s in

the most magnesian olivines , it can be concluded that mixing between this component and boninite melts took place before crystallization began. This mixing is likely to be responsible for a subtle increase in Al20 3 content in the Western group melts , which is recorded in spinel compositions but almost hidden in wholerock compositions, for a decrease of FeO* and H 20 contents , and for an apparent decrease in MgO content (reflected in lower Fo of the most magnesian olivines) . If the above suggestions are correct , they imply that the composition of the first olivine to crystallize decreased from F0 94 to - F093 with increasing


b)

1.60.8 1.2

0.35 ~-----------__:_l

0.05 .j--~_--.--_-...._-_ __._-_...J 0.4

'"00.25 N

~ ~0.15 ::z:=

a)

0.8 1.2

,,·N.,,' *, ~ \ 't \

, I.. , , I, ,, .'\ 94"• '\ * . ' ,

87 • ".02 9'0"87 "'" ' ... " • 00-,

" • Q'a ......

30 .,.----------------,

~O/TiOa KaO/TiOa

1 L_-------=.i~~--=:.:::....---.::Jl

6 c) 6 ,"'''I, d) e- 26-15 A- 26-20 O

N I

5,~ .- 3-44~4 as 4:I I

I 5-25I , * IIZ Z .- 5-28<, 3 ~~'3o 0 N \~'N 2 2 'Q <.

::z:= ::z:= "Q~~ 90-----02 - - - - ____E2 1 87~---- 91'\ "'

~ ----------Jc90 ~ o f:=::::==:::===~==:J 00.4 0.8 1.2 1.6 0.01 0.03 0.05 0.07

KaO/TiOa KaO/Ala0 3

Fig. 7. H 20/K 20 , H20 /N a20 and H20!AI203 values in melt inclusion s in olivin e phenocrysts from Tongan high-Ca boninites. Lab els show forsterite contents of host olivin es. Dashed field = compositions of melt inclusions in magn esian olivin es (Fo > 90) ; so lid field = composition s of natural glasses . El and E2 show samples which compositions were mainly determined by corresponding components. Note intermediate position of melt inclusions in evolved olivine phenocrysts which illustrates continuous degassing of Tongan

high -Ca boninite melts during crystalli zation . Data fro m Sobolev and Danyushevsky , 1994 and Table 2.

proportion of E2 component. On the basis of a melt inclusion study, and modeling of the reverse of olivine crystallization , Sobolev and Danyushevsky (1994) have determined the compositions and crystallization temperature of depleted Eastern group melts in equilibrium with F094 and enriched Western group melts in equilibrium with F093 • The melt in equilibrium with F094 , which is unaffected by mixing with E2 component (composition PM1, Table 2) , represents the primary melt of Tongan high-Ca boninites which was in equilibrium with their mantle source. The most primitive melt with the highest proportion of the E2 component (composition PM2 , Table 2) can be considered as a parental melt for the most enriched Tongan boninites. If mixing with the E2 component has taken place after separation of the primary melt from its source (e.g. en route to the surface), then composition PM2 never was in equilibrium with the mantle source. The data obtained indicate that there was a spectrum of parental melt compositions between the two extremes shown in Table 2. Demonstrated very high temperatures of the Tongan boninite mantle source ( - 1470°C, Table 2) preclude an opportunity for H 20 to be stored in the mantle and imply that the


0.51310

"C Z

: 0.51290 <, "C Z ~ ... 0.51270

{..LAU BASIN . \. VALU FA RIDGE

'\ 1\. TOFUA ARC\.'..:'*5-25, , , I'.

0.51310

0.51290

0.51270 INDIAN MORB

TONGAN BONINITES

0.704 0.706 87 Sr/86Sr

Fig. 8. Isotopic compositions of Tongan high-Ca boninit es compared to adjace nt contemporaneous lavas. Filled triangle-basalt ic lava from the King's Tripl e Junct ion . Data sources: Pacific MORB (Cohen et al., 1980; White et al. , 1987; Mahoney et al., 1994); Samoa (Wright and White, 1986; Haur i and Hart , 1993; Farley et al. , 1992); Tofua Arc (modern volcanic chain of the north Tonga Arc) (Ewart and Hawkeswor th , 1987); Indian MORB (Coh en et al. , 1980; Mahoney et al. , 1989; and references there in) ; Lau Basin (present day central Lau Basin spreading Ridges) (Loock et al. , 1990; Boespftug et al., 1990) ; Valu Fa Ridge (Loock et al. , 1990; Boespflug et al. , 1990; Vallier et al. , 1992); King's Tripl e Junction (Loo ck et al. , 1990). See text for discussion .

mixing of hot depleted mantle with an H20-rich fluid (component E1) has initiated the melting process.

DISCUSSION

Chemical features of Tongan high-Ca boninites As LREE are often considered to be more mobile than HFSE during various

metasomatic processes , the existence of significant negative Nb and Zr anomalies in the LREE depleted primary melt of Tongan boninites (Figs 5 and 6) indicates that LREE and MREE concentrations in this melt may be higher than can be expected as a result of melting of unmetasomatized depleted mantle (component D). HREE are probably unaffected , as there is a significantly smaller negative Ti anomaly. Note the obvious decoupling of Zr from Hf, as normalized Hf concentrations are approximately at the REE level of concentrations (Figs 5 and 6). This enrichment in LREE is not due to the presence of small amounts of the E2 component, as OIB lavas do not show negative Nb anomalies. We speculate that this enrichment may reflect some significant concentrations of LREE and MREE in the subduction-related fluid. We can further suggest that the proposed decoupling of Nb from Ta during slab degassing (Ta preferentially enters the fluid phase, Green, 1993) also applies to Zr and Hf. Slightly higher 143Nd/ I44Nd values of depleted Tongan boninites compared to enriched samples (Table 4) can then reflect the addition of some amounts of more radiogenic Nd from the slab. However , we believe that the data available are insufficient (especially bearing in mind large variations of 143Nd/144Nd in the Samoan OIB

20


and the lack of Ta data for depleted Tongan boninites) for any quantitative estimations of fluid (component E1) trace element composition.

An alternative explanation for the negative Nb, Zr and Ti anomalies is that they may reflect a complex history of interaction between mantle material of the Samoan plume and carbonatitic melts. The existence of such melts in the vicinity of the Samoan plume has recently been proposed by Hauri et al. (1993). Peridotites affected by carbonatitic metasomatism have prominent negative Nb , Zr, and Ti anomalies and very high La/Yb values (Fig. 6, Hauri et al. , 1993). If component D has been metasomatized pr ior to boninite genesis, this would result in higher LREE and MREE concentrations compared to Nb, Zr and to a lesser degree Ti. However, the overall REE pattern may still remain depleted.

We also emphasize that trace element features of metasomatized peridotites are similar to those of the E2 component, with Nb the only exception. Thus, it is possible that the unusual chemical features of some Samoan OIB magmas , and especially of those which are mixed with high-Ca boninitic melts (namely prominent negative Ti, Zr and Hf anomalies within an otherwise smooth enriched pattern (Figs 5 and 6», also result from metasomatic alteration of their sources by carbonatitic melts.

Demonstrated significant variations in the HREE contents of D component (Fig. 4) most likely reflect different degrees of depletion by previous melting events which produced OIB magmas.

H20 content ofprimary melts Demonstrated continuous degassing of Tongan high-Ca boninites during

crystallization implies that an important question to address is a possible degassing of boninite melts before crystallization began. Note that such degassing, if it occurred, will not have affected the inferred temperatures of the primary melts and mantle sources, as these were obtained independently from the melt inclusion studies. However, it can affect estimations of the amount of the E1 component involved in boninite petrogenesis.

The beginning of degassing should trigger crystallization of an olivinesaturated melt. As we have analyzed H 20 contents in the melt in equilibrium with the first olivine to crystallize (F094 ; Sobolev and Danyushevsky, 1994), the only possibility for degassing to occur is when the mel t ascends faster than the slope of olivine-saturated liquidus (5 "Czkbar, Ford et al., 1983). If we assume an adiabatic ascent of the melt ( - 3 "Czkbar, Nisbet, 1982), then the primary melts could be overheated by - 40 DC at shallow levels. Although the dependence of the olivine liquidus on the H 20 content of melts with MgO > 20 wt. % is not well constrained, a synthesis of available data (Green et al., 1975; Danyushevsky et al. , 1992) shows that 40 DC of cooling of an overheated melt may be responsible for the loss of -1 wt. % H 20 . Thus, the maximum possible H 20 content of primary melts in equilibrium with their source was 3 wt. %. However, we believe that the existence of well-understood correlation between the H20 contents of


Tongan boninites and the amount of the E2 component added (Fig. 7), implies that this "hidden" degassing did not occur.

A model of high-Ca boninite petrogenesis Any model which describes Tongan high-Ca boninite petrogenesis should

explain three major features: very high temperatures (-1470 °C) of the incompatible element-depleted mantle above the subducted slab (Table 2), similar isotopic compositions between components D and E2 (Fig. 8), and a striking similarity between relative concentrations of incompatible elements in the most incompatible element-enriched Tongan boninite, and adjacent Olb-Iike lavas (Fig. 6).

Transition from subduction to transform tectonics in the north Tongan region forms a "window" from the mantle wedge towards sub-Pacific mantle. We suggest that this "window" is used by the ascending Samoan plume to penetrate the mantle wedge above the subducted slab. As was shown in Poreda (1985) and Volpe et al. (1988), many seamounts in the northern part of the Lau Basin are isotopically similar to Samoan basalts. This implies that either the plume material itself, or the plume-derived melts have occupied a large area of the mantle wedge. Olfs-like magmas erupted on the northern Tongan fore-arc to the West of the boninite occurrences are a result of decompressional melting of this plume. These lavas have olivine-pyroxene-Cr-spinel inclusions of boninitic composition (Zlobin et al., 1991a), which indicates that they were erupted in situ, and are not seamount fragments accreted from the Pacific plate. This further supports the suggestion that the Samoan plume has penetrated into the mantle wedge. However, these inclusions cannot be used to reveal the relative timing of eruptions of high-Ca boninites and Olfs-like lavas, as these inclusions may represent gabbros with boninitic affinities from the forearc basement.

The demonstrated eastward "rollback" of the Tonga Trench (Carlson and Melia, 1984) may be important for the understanding of the Tongan high-Ca boninite petrogenesis. This rollback created an extensional tectonic regime in the Northern Lau Basin, which has caused part of the Samoan mantle plume to be drawn into the mantle wedge above the slab. Also, the gradual eastward shift of the rigid subducted lithosphere can result in the removal of the outer colder parts of the Samoan plume exposing the hotter inner parts above the slab . :

Once above the slab, and having been depleted during previous formation of alB melts, hot residual plume lherzolite (component D) interacts immediately with H20-rich fluids derived from the subducted slab (component E1), producing primary high-Ca boninite melts. The estimated depth of origin of these melts ( - 45-55 km, Danyushevsky et al., 1994) is some 50 km less than the depth of the slab under this region. These high pressures of primary melt generation are supported by a large temperature interval of olivine-only crystallization and a corresponding significant difference between compositions of the most magnesian olivine and orthopyroxene in Tongan high-Ca boninites. It seems likely that


introduction of such a hot material directly above the slab should trigger intensive degassing of the latter.

There are two major possible scenarios of mixing between primary high-Ca boninite melts produced during melting of component D, and the E2 component. The first is that this mixing occurs en route to the surface, during which different batches of primary high-Ca boninite melts mix in different proportions with earlier- or simultaneously-formed alB melts (component E2), to produce the observed chemical and mineralogical spectra of high-Ca Tongan boninites. This mechanism assumes that boninites and alB melts were formed from different parts of the plume. Boninite source mantle was depleted prior to penetration in the wedge by melting to produce Samoan magmas. This depleted hot peridotite could be exposed and drawn into the wedge as discussed above. OlBvlike melts erupted to the west of boninites were formed as a result of decompressional melting of unmodified parts of the plume which rose up as a result of plume penetration into the wedge. The spatial distribution of various boninites supports this suggestion , as the most LREE-enriched boninites are located closer to the Olfs-like lavas. Primitive alB melts usually contain 0.5-1 wt. % H20 (Danyushevsky et al., 1993) which can explain the observed decrease in boninite H20 content with increasing proportion of E2 component.

An alternative possibility is that the actual mixing took part in the mantle during boninite primary melt genesis. In this case, the E2 component represents residual melts which have not separated from their sources during alB magma genesis.

We believe that the first model is more consistent with the data summarized in this paper , although we acknowl edge that there is no tool available at present which allows a quantitative distinction between the two to be made.

A comparison of compositions of spinels included in the most magnesian olivine phenocrysts from all high-Ca boninite suites (Fig. 3b) demonstrates that:

-Extensive mixing with Olfs-like melts is an exclusive feature of Tongan boninites.

-A trend existence from the Troodos spinels through the Hunter Ridge spinels to the Ti02-poor end-member of the Tongan spinel population.

This second feature may reflect variable degrees of previous source depletion, rather than different degrees of melting during high-Ca boninite petrogenesis, since Sobolev et al. (1993) have determined the primary melt composition for the Troodos high-Ca boninites, which is very close in terms of MgO , H20 contents and crystallization temperature to that established for Tongan high-Ca boninites (Table 2) .

Although we believe that the overall petrogenesis of all three high-Ca boninite suites was similar, different degrees of depletion of their respective sources and different extents of mixing with incompatible element-enriched silicate melts imply that some, probably significant, variations exist in the petrogenetic scenario for each suite.


In conclus ion, we emphasize here that apparently no unique scheme exists for the petrogenesis of boni nites in general. It is obvious, for example, that the proposed model for the Tongan high-Ca boninites cannot be invoked to explain the petrogenesis of the almost 3000 km long belt of contemporaneous boninites in the Bonin-Mariana forearc (Pearce et al., 1992). Similar Mg# of the most magnesian olivines and low-Ca pyroxenes in low-Ca boninites from New Caledonia (Cameron, 1989) and Cape Vogel (Walker and Cameron, 1983), whose volumetric characteristic can be close to those of the Tongan high-Ca boninites, imply that not only the mantle sources of these low-Ca boninites were more depleted in the basaltic component, but the pressure of melting was significantly lower. This, in turn, indicates higher H20 contents of their primary melts, which was recently confirmed by ion probe analyses of melt inclusions in magnesian olivines from these rocks (Danyushevsky, Crawford and Sobolev, unpublished). The common features of all boninites: a depleted mantle source metasomatized by an H20-bearing component are not sufficient conditions to constrain their genesis (Crawford et al., 1989), as there are apparently many different scenarios for the evolution of these mantle sources and their interactions with enriched components; and corresponding detailed research should be performed on each particular suite.

Acknowledgements-L.V.D . and TJ.F acknowledges support from the Australian Research Council. This research was also supported by Professor D. H. Green and Dr A. J . Crawford through their ARC Large grants, and by International Science Foundation (grant MNNOOO to A .V.S .) We thank Dr J. Woodhead for performing lead isotopic analyses . A .V.S. expresses his gratitude to M. Chaussidon (CRPG, Nancy) for his assistance in ion-probe analyses. We thank Prof. Green and Dr Crawford for their comments on this manuscript. We also thank D . Ohnenstetter, W. M. White and R. N. Taylor for their constructive review of this paper.

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North Tongan high-Ca boninite petrogenesis: The role of samoan plume and subduction zone-transform...

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