Journal of Petrology Volume 41 Number 4 Pages 477-509 2000
© Oxford University Press 2000
Trace Element Residence and Partitioning in Mantle Xenoliths Metasomatized by Highly Alkaline, Silicate- and Carbonate-rich Melts (Kerguelen Islands, Indian Ocean)
1GEMOC ARC NATIONAL KEY CENTRE, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY, NORTH RYDE, N.S.W. 2109, AUSTRALIA
2DEPARTMENT OF GEOLOGY–UMR 6524, UNIVERSITY J. MONNET, 23 RUE P. MICHELON, 42023 ST-ETIENNE, FRANCE
Received January 21, 1999; Revised typescript accepted September 21, 1999
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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Mantle xenoliths in alkaline lavas of the Kerguelen Islands consist of: (1) protogranular, Cr-diopside-bearing harzburgite; (2) poikilitic, Mg-augite-bearing harzburgite and cpx-poor lherzolite; (3) dunite that contains clinopyroxene, spinel phlogopite, and rarely amphibole. Trace element data for rocks and minerals identify distinctive signatures for the different rock types and record upper-mantle processes. The harzburgites reflect an initial partial melting event followed by metasomatism by mafic alkaline to carbonatitic melts. The dunites were first formed by reaction of a harzburgite protolith with tholeiitic to transitional basaltic melts, and subsequently developed metasomatic assemblages of clinopyroxene + phlogopite ± amphibole by reaction with lamprophyric or carbonatitic melts. We measured two-mineral partition coefficients and calculated mineral–melt partition coefficients for 27 trace elements. In most samples, calculated budgets indicate that trace elements reside in the constituent minerals. Clinopyroxene is the major host for REE, Sr, Y, Zr and Th; spinel is important for V and Ti; orthopyroxene for Ti, Zr, HREE, Y, Sc and V; and olivine for Ni, Co and Sc.
KEY WORDS: mantle xenoliths; mantle metasomatism; partition coefficients; Kerguelen Islands; trace elements
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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Mantle xenoliths provide unique information about the chemistry and mineralogy of deep lithospheric rock types (e.g. Nixon, 1987
; O’Reilly & Griffin, 1996). Studies of upper-mantle xenoliths in alkali basalts, kimberlites, lamproites and carbonatites have improved our understanding of materials and processes involved in the geochemical evolution of the mantle (e.g. Downes & Dupuy, 1987; Ionov et al., 1993; Chalot-Prat & Boullier, 1997). The variation and magnitude of geochemical heterogeneities in the lithospheric mantle reflect the composition of mantle melts and fluids and the efficiency of heat and mass transfer. Mantle plumes are important for initiating such transfer processes. The Kerguelen plume is remarkable because of its volume, the persistence of volcanic activity for at least 115 My, and its migration across diverse geotectonic environments through time as a result of spreading of the Indian Ocean (Weis et al., 1992). In this paper we report a trace element study of bulk rocks and constituent minerals of clinopyroxene-bearing harzburgite, clinopyroxene-poor lherzolite and phlogopite + clinopyroxene-bearing dunite from the mantle beneath Kerguelen. These rocks show evidence for partial melting and mantle metasomatism related to the Kerguelen mantle plume. Our data yield insights into the distribution of trace elements in peridotites and concerning element partitioning between minerals under upper-mantle P–T conditions.
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GEOLOGICAL SETTING |
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The Kerguelen Islands are located in the oceanic domain of the Antarctic Plate (Fig. 1). They are the exposed part of the Kerguelen oceanic plateau, which is the second largest (25 x 106 km3) after the Ontong Java plateau (Coffin & Eldhom, 1993
). The Kerguelen Islands drifted from a location near the South East Indian Ridge (SEIR) to their present-day intraplate setting. Their magmatic activity has extended over 45 My (Giret, 1993). Therefore, the Kerguelen Islands combine characteristics of both the Iceland and Hawaiian hotspots (Giret et al., 1997). Magmatism has progressively changed from tholeiitic to alkaline (Gautier et al., 1990; Weis et al., 1993).
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, other xenolith outcrops. Inset: location of Kerguelen Islands. SWIR, South West Indian Ridge; SEIR, South East Indian Ridge.Ultramafic and mafic xenoliths from the Kerguelen Islands are found in dykes, lava flows and breccia pipes of the youngest and more alkaline basaltic rocks (Grégoire et al., 1994
, 1998; Mattielli et al., 1996). We collected xenoliths from 10 different localities in the archipelago (Fig. 1). The xenoliths are subrounded in shape and range from 10 to 20 cm. They are Type I Kerguelen mantle xenoliths of Grégoire et al. (1997).
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SAMPLING AND ANALYTICAL METHODS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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Samples (50–100 g) from the central parts of xenoliths were ground in an agate mill. Major and minor elements (Cr, Ni) in bulk rocks were analysed by X-ray fluorescence spectrometry (XRF) at Macquarie University [see O’Reilly & Griffin (1988)
for methods]. The concentrations of 29 minor and trace elements [rare earth elements (REE), Ba, Rb, Th, U, Nb, Ta, Pb, Sr, Zr, Ti, Y, Sc, V, Co, Cu and Zn] were analysed by inductively coupled plasma mass spectrometry (ICP-MS) with a Perkin Elmer Sciex Elan 6000 instrument at Macquarie University. The sample preparation for ICP-MS was as follows: 100 mg of sample powder was dissolved in concentrated HF–HNO3 in 17 ml Savillex Teflon screw-top beakers. Following digestion, samples were evaporated to incipient dryness, dissolved in 6 N HNO3, and again evaporated to incipient dryness. Remaining residues were dissolved in 2% HNO3. The sample solutions were transferred to 125 ml polypropylene bottles. A known weight of internal standard solution was added and the solution was diluted with 2% HNO3. Final sample/solution ratios were 
1000–1200. The technique yields results that agree well with recommended values for the Kilauea basalt KIL-1 [93-1489 of Eggins et al. (1997)], which was used as a standard, and also measured along with the samples to assess precision (<2·5% RSD). The reference standard W-1 was prepared and analysed with the samples, together with three reagent blanks. No oxide corrections were used. Detection limits (taking into account chemical blank) were 1–5 ppb for most REE, Y, Nb, Ta, Th and U; 5–30 ppb for Ce, Sm, Zr, Rb, Sr, Ba, Pb and Sc; 50–100 ppb for V and Co; 200 ppb for Cu and Zn; and 500 ppb for Ti.
Mineral compositions were determined by a Cameca Camebax SX 50 microprobe at Macquarie University using wavelength-dispersive spectrometry (WDS). The microprobe was used with 15 kV accelerating voltage, sample current of 20 nA, a beam diameter of 2–3 mm, and natural and synthetic minerals as standards. Matrix corrections were done by PAP (Pouchou & Pichoir, 1984) procedures. Count times were 20–40 s and no values are reported below detection limits (0·01–0·04 wt %).
Concentrations of 29 trace elements (REE, Ba, Rb, Th, U, Nb, Ta, Pb, Sr, Zr, Hf, Ti, Y, Sc, V, Co and Ni) and of Al and Ca in olivine were determined in >100 mm polished sections by ICP-MS with a Perkin Elmer Elan 5100 instrument (16 samples) and a Perkin Elmer Elan 6000 instrument (three samples) at Macquarie University. Both ICP instruments were coupled to a Continuum Surelite I-20, Q-switched Nd:YAG laser ablation system (LA-ICP-MS). A typical analysis consists of 120 replicates, with each replicate representing one sweep of the mass range at a dwell time of 50–100 ms per mass. For each sample, 30–35 replicates were counted on the carrier gas (argon) alone to establish the background, followed by 85–90 replicates for ablation. The NIST 610 glass standard was used to calibrate relative element sensitivities. Each analysis was normalized using either CaO (clinopyroxene) or MgO (orthopyroxene, olivine, spinel, phlogopite, amphibole) values determined by electron microprobe. Typical detection limits are in the range 10–20 ppb for REE, Ba, Rb, Th, U, Nb, Ta, Pb, Sr, Zr, Hf and Y; 100 ppb for V and Sc; 2 ppm for Ti, Ni, Co and Cr; and 5 ppm for Al and Ca. The typical precision and accuracy for a laser microprobe analysis range from 2 to 7%. A more detailed description of laser operating conditions, calibration values for the NIST 610 glass standard and error analysis have been given by Norman et al. (1996) and Norman (1998). The LAMTRACE program developed by S. E. Jackson (e.g. Longerich et al., 1996) was used for data reduction.
Modal compositions (Table 1) were calculated by mass balance based on major element bulk-rock compositions and electron-microprobe analyses of constituent minerals.
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PETROGRAPHY AND MODAL COMPOSITION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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Protogranular and poikilitic harzburgite xenoliths show local gradation to porphyroclastic microstructures in which a fine-grained mosaic of olivine and orthopyroxene neoblasts (<1 mm) surrounds larger porphyroclasts (2–10 mm). Crystals in protogranular harzburgites show curvilinear grain boundaries. The grain size of olivine and orthopyroxene typically varies from 2 to 10 mm. The poikilitic texture is similar to that described for some harzburgitic xenoliths from the French Massif Central (Coisy & Nicolas, 1978
). Poikilitic microstructure differs from the protogranular microstructure by the presence of olivine grains up to 5 cm long that contain inclusions of orthopyroxene. Clinopyroxene grains typically contain inclusions of resorbed orthopyroxene and spinel grains. Vermicular spinel grains in both types of harzburgites occur between olivine and pyroxene crystals and frequently form clusters with orthopyroxene and clinopyroxene. Phlogopite rarely occurs only in poikilitic harzburgites. Millimetre-sized interstitial phlogopite crystals contain inclusions of spinel, olivine and orthopyroxene, and are in textural equilibrium with clinopyroxene. More rarely, phlogopite forms thin veins (<0·5 mm). Amphibole was found in a single sample (OB-93-5) of phlogopite-bearing harzburgite. Most of the amphibole grains are rounded interstitial, but a few occur as inclusions in phlogopite grains. Clinopyroxene-bearing harzburgite and clinopyroxene-poor lherzolite xenoliths from the Kerguelen Islands contain melt and fluid inclusions trapped in olivine and pyroxene (Schiano et al., 1994). Poikilitic samples are especially rich in such inclusions. Clinopyroxene + phlogopite-bearing dunite samples are coarse-grained rocks made up mostly of 2–5 (rarely to 10) mm olivine grains. Clinopyroxene, orthopyroxene, spinel, phlogopite and amphibole grains range from 0·1 to 2 mm. Small, isolated, globular to anhedral grains of spinel are either interstitial or occur as inclusions in olivine, or more rarely, in clinopyroxene. Clinopyroxene is interstitial and locally concentrated in thin layers (<0·5 mm) between olivine crystals. Clinopyroxene commonly contains exsolution lamellae of spinel. Phlogopite grains are commonly euhedral and interstitial but some are found as inclusions in olivine. Sample MG-91-143 contains both amphibole and phlogopite. The amphibole is interstitial and commonly appears to have replaced interstitial clinopyroxene. Samples MM-94-54, MM-94-97 and MM-94-101 contain numerous sulphide mineral grains.
Protogranular and poikilitic harzburgite samples contain 67–84% olivine, 13–29% orthopyroxene, 1·5–5% clinopyroxene and 0·5–1·5% spinel. Phlogopite (0·5%) is present in samples OB-93-3 and OB-93-5, and amphibole (0·5%) in sample OB-93-5. Poikilitic sample JGM-92-1c is the first true lherzolite reported from the Kerguelen Islands [according to Streckeisen’s (1976) classification]. It is poor in clinopyroxene (olivine 79%, orthopyroxene 13%, clinopyroxene 7%, spinel 1%). Kerguelen dunites (olivine 94–97%) contain 1–3% of clinopyroxene and 1–2% of spinel. Phlogopite (0·5–1·5%) is found in all dunite samples, whereas amphibole has been observed in only one sample (MG-91-143).
Many samples display interstitial patches and/or veins of fine-grained material with a complex mineralogy. In harzburgite samples OB-93-3, OB-93-5, GM-92-501 and GM-92-502, this material consists of feldspar + olivine2 + rutile + ilmenite + Cr-armalcolite + Cr–Ca-armalcolite + Ti-chromite. Other samples of harzburgite and dunite display patches <50 µm wide that consist of clinopyroxene2 ± olivine2 ± amphibole ± biotite ± chromite ± ilmenite ± rutile ± feldspar ± carbonate ± glass. Electron microprobe analyses show that compositional effects of these assemblages on the original minerals of the host peridotite are restricted to the adjacent 50–100 µm of host minerals.
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MAJOR ELEMENT COMPOSITION |
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Whole-rock samples
The protogranular and poikilitic harzburgites and clinopyroxene-poor lherzolite have similar major element contents, with CaO/Al2O3 ratios that range from 0·60 to 1·30 (Table 1). All are strongly depleted in ‘basaltic’ components (CaO < 1·35 wt %; Al2O3 < 1·45 wt %; Na2O < 0·25 wt %) compared with model compositions for the undepleted upper mantle (CaO 3·23–3·60 wt %; Al2O3 4–4·46 wt %; Na2O 0·33–0·61 wt %, Jagoutz et al., 1979
; McDonough & Sun, 1995). However, protogranular harzburgite samples display mg-numbers [100Mg/(Mg + Fe2+)] of 91·4–92, higher than that for primitive upper mantle (89·8). The mg-numbers of poikilitic harzburgite samples vary from 89·9 to 91·8, and poikilitic, clinopyroxene-poor lherzolite sample JGM-92-1c has an mg-number of 88. Dunite samples exhibit low CaO (0·24–0·53 wt %), Na2O (0·15–0·21 wt %), Al2O3 (0·44–0·82 wt %) and TiO2 (0·04–0·07 wt %) contents, which suggest a refractory character, but their mg-numbers (86–88·5) are lower than those for both primitive mantle (89·8) and the harzburgite xenoliths. The CaO/Al2O3 ratios of dunite samples are 0·40–0·89.
Minerals in clinopyroxene-bearing harzburgite
The mg-numbers of olivine range from 86 to 92 in the two types of harzburgite. Olivine compositions are more uniform in protogranular harzburgites (mg-number = 90·5–92) than in poikilitic peridotites (mg-number = 86–92). The lowest mg-number value is for olivine from lherzolite sample JGM-92-1c. Orthopyroxene displays a similar distribution. Clinopyroxene mg-numbers range from 91·5 to 92·5 in protogranular harzburgite samples, and from 87·5 to 92·5 in poikilitic ones (Table 2). Orthopyroxene from poikilitic harzburgite samples is typically richer in Na2O, Al2O3 and TiO2 than those of protogranular harzburgites (Table 2).
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Clinopyroxene ranges in composition from diopside in protogranular harzburgite samples to magnesian augite in poikilitic harzburgites and lherzolite samples (Table 2). The mg-number of diopside is high and homogeneous (93–95) for protogranular harzburgite, but shows a wider range and lower values (86–91·5) for harzburgite and lherzolite samples that contain poikilitic Mg-augite. The mg-number of clinopyroxene is systematically higher than those of olivine and orthopyroxene in protogranular harzburgite samples. In contrast, the mg-number of the clinopyroxene from poikilitic, Mg-augite harzburgite and lherzolite samples is lower than or equal to that of olivine and orthopyroxene. The Mg-augite is richer in Al2O3, Na2O, TiO2 and Cr2O3 than is Cr-diopside from protogranular harzburgite samples (Table 2).
Spinel grains are Mg–Al chromites that display mg-numbers of 66·5–72 and cr-numbers [100Cr/(Cr + Al)] of 40–52 in protogranular harzburgite samples, and mg-numbers of 58–79 and cr-numbers of 20–49 in poikilitic harzburgite and lherzolite samples. The TiO2 contents of spinel are smaller in protogranular harzburgite samples (<0·07 wt %) than in poikilitic harzburgite and lherzolite samples (0·20–2·80 wt %).
The amphibole in poikilitic harzburgite sample OB-93-5 is pargasite of fairly constant TiO2, Na2O and Cr2O3 contents (Table 2)
The phlogopite in poikilitic harzburgite samples OB-93-3 and OB-93-5 displays high mg-numbers and Cr2O3 contents (Table 2).
Minerals in clinopyroxene + phlogopite-bearing dunite
The mg-numbers of olivine in dunite samples range from 86 to 90. These samples contain an Al–Cr diopside in which mg-numbers range from 87·5 to 90 (Table 2). Spinel is Mg-chromite that shows a wide range of composition and significant amounts of Fe2O3 and TiO2 (Table 2). Phlogopite (mg-number 85–90) has Cr2O3 of 1·40–2·20 wt %. Phlogopite associated with amphibole in sample MG-91-143 has less TiO2, Al2O3 and K2O, and more Na2O and SiO2 than that found in amphibole-free dunite samples (Table 2). The disseminated amphibole in sample MG-91-143 (mg-number 84·75) is pargasite (Table 2).
Equilibration temperatures
Equilibration temperatures of clinopyroxene-bearing harzburgite and clinopyroxene-poor lherzolite samples were estimated using two-pyroxene geothermometers (Brey & Köhler, 1990a), the Ca-in-orthopyroxene geothermometer (Brey & Köhler, 1990b), orthopyroxene–spinel equilibria (Sachtleben & Seck, 1981) and olivine–spinel equilibria (Fabriès, 1979; Ballhaus et al., 1991). Core compositions of large grains in the harzburgites were used. Our goal was only to establish relative temperatures that are consistently estimated by all methods.
The majority of the mineral assemblages in the protogranular harzburgite samples re-equilibrated at T = 845–1005°C (Table 3). The range of equilibration temperatures is systematically higher for poikilitic harzburgite and cpx-poor lherzolite samples (T = 1015–1135°C). The olivine–spinel and orthopyroxene–spinel geothermometers yield temperatures of 925–940°C for sample GM-92-502.
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In samples of clinopyroxene + phlogopite-bearing dunite, the absence of orthopyroxene makes temperature estimates less reliable. Olivine–spinel equilibria yield equilibration temperatures of 940–1090°C (Table 3).
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TRACE ELEMENT COMPOSITIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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Transition trace elements in clinopyroxene-bearing harzburgite and clinopyroxene-poor lherzolite
The harzburgites and the clinopyroxene-poor lherzolite show very similar bulk-rock transition element contents (Table 4). They are samples richer in Ni and Co, and poorer in Sc and V than the postulated primitive upper mantle (Table 4; McDonough & Sun, 1995
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Olivine (Table 5) and spinel (Table 6) from protogranular harzburgite samples are respectively poorer in Cr and Ni than those from poikilitic harzburgite and lherzolite samples. Ca and Al contents of olivine are greater and vary more widely in poikilitic than in protogranular harzburgite samples (Table 5). Orthopyroxene shows the most marked difference in composition between the two textural types. Sc is higher and Co and Ni are commonly lower in protogranular harzburgite samples than in poikilitic peridotite samples (Table 7). There is no systematic difference in the transition trace element content of the Cr-diopside occurring in the protogranular harzburgite samples and of the Mg-augites in the poikilitic peridotite samples, but the latter display a wider range of compositions (Table 8). Spinel and phlogopite (Tables 6 and 9) of Kerguelen harzburgites and clinopyroxene-poor lherzolite are rich in V and Co and poor in Sc. Sc is concentrated in clinopyroxene (45–73 ppm) and amphibole (47 ppm) (Tables 8 and 9).
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Transition trace elements in clinopyroxene + phlogopite-bearing dunite
Dunites have bulk-rock Ni, Sc and V contents similar to those of harzburgites; the two samples rich in sulphides (MM-94-51 and MM-94-101) are higher in Cu (Table 4). Clinopyroxene grains in dunite are richer in Sc than those of harzburgites (Table 8). Olivine (Table 5) from dunites displays the same Cr contents as those from protogranular harzburgite. Its Ca and Al contents vary in a range, which approximately covers those of the two types of harzburgites. Ni/Co ratios of dunitic olivine range from 12·3 to 19·85, and are typically below the primitive mantle value of 18 (McDonough & Sun, 1995) and lower than those of the two types of harzburgites (19·4–22). Spinel in Kerguelen dunites is also rich in V, Co and Ni, and contains little Sc (Table 6). Phlogopite (Table 9) in amphibole-bearing dunite contains less Sc and V than in amphibole-free samples. Amphibole in the phlogopite-bearing sample (Table 9) contains significant amounts of Sc, V, Co and Ni.
Incompatible trace elements (REE, Y, HFSE, LILE)
The two types of harzburgites and the lherzolite from Kerguelen have different incompatible trace element characteristics, but all are enriched in the more incompatible trace elements (Table 4; Figs 2 and 3). Both clinopyroxene grains and bulk-rock samples display U-shaped REE patterns in protogranular harzburgite and light REE (LREE)-enriched REE patterns in poikilitic harzburgite and lherzolite (Figs 2 and 3). In Kerguelen dunite samples, only the Rb and Pb values of bulk-rock analyses (and Th in sample MM-94-54) are close to those of primitive mantle. Other elements are significally depleted relative to the primitive mantle (Fig. 4). Dunite samples display trace element patterns enriched in the more incompatible trace elements (Fig. 4). Olivine and spinel from dunite typically have very low contents of incompatible trace elements, most below detection limits (Tables 5 and 6). Spinel displays a large range of Ti contents (Table 6). Some spinel grains may contain small amounts of Zr (0·50–4 ppm) and Nb (0·25–1 ppm). Ti content of olivine ranges from 9·50 to 68·35 ppm.
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, sample BOB-93-666; • sample OB-93-58; 
, sample OB-93-426; 
, sample OB-93-67b. In primitive mantle-normalized trace element plots, the sequence of elements is related to their decreasing incompatibility during partial melting of upper-mantle peridotites (Sun & McDonough, 1989
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, sample GM-92-453; 
, sample MG-91-260.
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, sample MG-91-143 (phlogopite + amphibole sample).
Incompatible trace elements in protogranular, clinopyroxene-bearing harzburgite
Incompatible trace element contents are commonly below detection limits for olivine, orthopyroxene and spinel (Tables 5–7). Bulk-rock samples and clinopyroxene grains of protogranular harzburgite display very uniform heavy REE (HREE) contents. Lucpx and Lubulk rock range from 0·5 to 1·5 and from 0·03 to 0·08 times the primitive mantle value, respectively. Sample OB-93-67b shows the lowest La and Ce contents and the highest Yb and Lu contents, but all samples are richer in LREE than in middle REE (MREE) (Fig. 2). Sample OB-93-426 displays the highest REE contents, in both clinopyroxene and the bulk rock. Clinopyroxene from protogranular harzburgite samples has pronounced depletion in Ba, Zr and Ti, and enrichment in U and Pb (Fig. 2). Two other LREE-rich clinopyroxenes (samples OB-93-426 and OB-93-58) also show a significant depletion in Nb. Samples OB-93-426, BOB-93-666, OB-93-67b and OB-93-279 show a depletion in Sr. Bulk-rock trace element patterns display high enrichment in Pb and Sr and commonly an enrichment in Ti (Fig. 2). Sample BOB-93-666 has high Rb and Ba contents. Orthopyroxene from protogranular harzburgite samples shows significant amounts of Ti and around 1 ppm of Ga (Table 7). Other incompatible trace element contents of orthopyroxene are very low but the LREE content of orthopyroxene from sample OB-93-426 is higher than that of orthopyroxene from the other protogranular harzburgite samples (Table 7). Olivine has low Ti contents and very low Zr and Pb contents (Table 5). Spinel from sample OB-93-426 contains more Ti than those from the four other protogranular harzburgite samples (Table 6). Spinel from protogranular harzburgite samples also has low Rb, Sr, Zr, Nb and Pb contents (Table 6).
Incompatible trace elements in poikilitic clinopyroxene-bearing harzburgites and clinopyroxene-poor lherzolite
Incompatible trace element (ITE) contents of poikilitic peridotites and their constituent minerals are much higher than those of protogranular samples (Tables 4–8, Figs 2 and 3). Olivine from poikilitic peridotite samples contains low amounts of Ti, Zr and Pb (Table 5). Spinel grains of poikilitic peridotite samples are richer in Ti than those of protogranular harzburgite samples. They contain low amounts of Rb, Sr, Zr and Nb (Table 6). The poikilitic, clinopyroxene-poor lherzolite contains clinopyroxene and orthopyroxene richer in Ti than those of poikilitic harzburgite samples (Tables 7 and 8).
REE compositions of both bulk rock and minerals are very uniform (Figs 3 and 5). Clinopyroxene is poor in Ba, Nb, Pb, Sr, Zr and Ti (Fig. 3).The bulk rocks are uniformly poor in Ba but not in Nb, Pb, Sr, Zr and Ti (Fig. 3). For example, the anhydrous sample GM-92-501 shows no depletion in Nb and Pb, depletion in Sr and Zr, and enrichment in Ti (Fig. 3). On the other hand, the amphibole + phlogopite-bearing sample OB-93-5 displays no depletion in Nb and Sr but it is depleted in Pb, Zr and Ti (see Fig. 3 for other samples). Orthopyroxene grains have lower REE contents than primitive upper mantle and display LREE-depleted REE patterns (Fig. 3). They are also relatively poor in Ba, Sr and Pb but rich in Zr and Ti. However, orthopyroxene from samples OB-93-3 and OB-93-22 is poor in Ti, and that from sample OB-93-5 is poor in Zr (Fig. 3). Phlogopite from samples OB-93-5 and OB-93-3 displays REE contents lower than those of the primitive mantle (Fig. 5 and Table 9). The phlogopite grains are rich in Rb, Ba, Nb and Ta, and contain significant amounts of Sr, Ti, Pb, Zr and Hf, but are poor in Th, U and Y (Table 9). Amphibole from sample OB-93-5 displays high ITE contents and LREE-enriched patterns, but is poor in U, Ta, Pb, Zr and Hf (Fig. 5).
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, MG-91-143 phlogopite + amphibole sample. Amphibole of dunite:
Incompatible trace elements in clinopyroxene + phlogopite-bearing dunites
Phlogopite-bearing, amphibole-free dunite samples have uniform REE and trace element bulk-rock contents. They display LREE-rich or upward convex REE patterns, and are rich in Ti and Pb, and poor in Zr (Fig. 4). Sample MM-94-54 is rich in Th, U and Sr. Sample GM-92-480 is rich in Sr, Rb, Ba, Nb and Ta (Fig. 4). Clinopyroxene grains in this suite display LREE-rich or upward convex REE patterns, are poor in Pb and Ti, but otherwise have incompatible trace element patterns similar to those of clinopyroxene from the poikilitic harzburgite (Fig. 4). Samples BOB-93-640.1, MM-94-54, MM-94-97 and MM-94-101 contain clinopyroxene that is poor in Nb and Zr and of variable Sr contents. Clinopyroxene from sample GM-92-480 is poor in the most incompatible elements (Rb, Ba, Th, U and Nb) and in Zr (Fig. 4). Clinopyroxene from sample GM-92-468 is the richest in moderately incompatible trace elements (from Sr to Lu, Fig. 4). It is also rich in Hf and poor in Sr. Phlogopite grains from this suite resemble counterparts from poikilitic harzburgite (Fig. 5). Both are rich in Rb, Ba, Nb, and Ta, and poor in Th, U and REE (Table 9, Fig. 5).
Clinopyroxene and amphibole from amphibole- and phlogopite-bearing dunite sample MG-91-143 display similar REE contents and LREE-rich REE patterns (Figs 4 and 5). Clinopyroxene is poorer in Rb, Ba, Nb, Ta, Ti, Zr and Hf than amphibole. Phlogopite grains (Fig. 5) contain less Ti, Nb, Ta, Zr and Hf, and more Rb, Pb and U than those found in amphibole-free dunite and poikilitic harzburgite samples.
Origin and evolution of Kerguelen mantle xenoliths
Partial melting characteristics
Both types of clinopyroxene-bearing harzburgites and the clinopyroxene-poor lherzolite display features of residues of high degrees of partial melting. Bulk-rock compositions rich in MgO, Ni and Co and poor in CaO, Al2O3, Na2O, Sc and V result from the preferential melting of clinopyroxene and spinel and retention of olivine and orthopyroxene (Mysen & Kushiro, 1977; Presnall et al., 1978; Kostopoulos, 1991). The high mg-numbers of the bulk rocks, olivine and pyroxenes, along with low and uniform HREE content, imply that a significant amount of basaltic melt was removed from protogranular harzburgites. Fifteen to 25 wt % melting of Kerguelen protogranular harzburgite was estimated by Grégoire et al. (1997). This result is consistent with the presence of clinopyroxene in even the most depleted protogranular harzburgites (Elthon, 1993). Kerguelen mantle xenoliths lack the LREE-depleted patterns of mantle peridotites from oceanic settings (e.g. abyssal peridotites, Johnson et al., 1990; Hawaiian peridotites, Sen et al., 1993; Yang et al., 1998) that result from a simple residue model.
Metasomatic characteristics
The consistent enrichment of highly incompatible trace elements in the harzburgites and clinopyroxene-poor lherzolite requires metasomatic reactions [see also Hassler & Shimizu (1998)]. In addition to incompatible trace element enrichment, poikilitic peridotite samples display lower bulk-rock and mineral mg-numbers than those of protogranular harzburgite, and contain Na-, Cr-, Al- and Ti-rich magnesian augite, ± phlogopite and amphibole. U-shaped patterns similar to those of protogranular harzburgite samples are found in metasomatized refractory mantle spinel peridotites, both as xenoliths entrained in alkali basalts (Downes & Dupuy, 1987; Siena et al., 1991; Xu et al., 1998) and in orogenic lherzolite massifs (Bodinier et al., 1991; Downes et al., 1991). The LREE-enriched or upward convex REE patterns of both bulk-rock samples and clinopyroxene from the poikilitic harzburgite closely resemble those of alkaline cumulate rocks and peridotite infiltrated by alkaline silicate melts (Bodinier et al., 1988; Fabriès et al., 1989; Xu et al., 1998). The difference between the two types of REE patterns can be explained by a chromatographic effect (Navon & Stolper, 1987; Bodinier et al., 1988).
The presence of phlogopite ± amphibole and the major and trace element compositions of minerals and bulk rocks preclude an origin for Kerguelen dunite as single residues of partial melting processes [see also Grégoire et al. (1997)]. Furthermore, minerals in most of phlogopite-bearing dunite samples display incompatible trace element characteristics similar to those of poikilitic Mg-augite-bearing harzburgite (except for the phlogopite + amphibole-bearing dunite sample).
To assess the nature of the metasomatic agents we calculated melt compositions in equilibrium with clinopyroxene in poikilitic harzburgite, clinopyroxene-poor lherzolite and dunite. We used two sets of clinopyroxene–melt partition coefficients to assess the possible effect of the range of these values on the results. The first set is average Dcpx/sil = Dclinopyroxene/mafic silicate melts, with most elements from the compilation of Chazot et al. (1996), the value for Ti from Hart & Dunn (1993) and the value for Ta from Chalot-Prat & Boullier (1997). The other set of coefficients is Dcpx/carb = Dclinopyroxene/carbonatitic melts, with most elements from Klemme et al. (1995), except for the value for Rb, from Green et al. (1992).
Melts display trace element contents 10–5000 times those of the primitive mantle (except for Ti, Ta and Zr calculated with Dcpx/carb; Fig. 6). The two calculated melt compositions for amphibole + phlogopite-bearing dunite sample MG-91-143 are similar. Both are depleted in Nb, Ta, Zr, Hf and Ti (Figs 6 and 7; see Hauri et al., 1993; Chazot et al., 1994; Baker et al., 1998).
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Model melt compositions for poikilitic harzburgite, clinopyroxene-poor lherzolite and phlogopite-, clinopyroxene-bearing dunite suggest that all these rocks could have been in equilibrium with a CO2-rich silicate melt (Figs 6 and 7). However, the high field strength element (HFSE) contents of the model melt depend on the partition coefficient dataset. Using the Dcpx/carb we obtained an equilibrium melt poor in HFSE, but not as HFSE poor as melt in equilibrium with sample MG-91-143 (Fig. 7). However, Dcpx/sil values yield model melts that are relatively undepleted in Nb, Zr and Hf, and only slightly depleted in Ti. The calculated equilibrium melt for Dcpx/sil resembles that of ultramafic and alkaline lamprophyre in its incompatible trace element composition but not that of carbonatites (Fig. 7). The equilibrium melts calculated with Dcpx/carb have geochemical signatures consistent with either carbonatite or ultramafic and alkaline lamprophyre (Fig. 7).
In summary, the amphibole + phlogopite dunite (MG-91-143) probably was metasomatized by a CO2-rich silicate melt. The poikilitic harzburgite, clinopyroxene-poor lherzolite and phlogopite-bearing dunite samples probably were metasomatized by highly alkaline mafic silicate melts. Such magmas have been already described from oceanic settings (Yagi et al., 1975; Nixon et al., 1980; Rock, 1987), and young dykes with ‘lamprophyric affinities’ are found in the Kerguelen Islands (Leyrit, 1992; Moine, 2000).
Mantle history of the Kerguelen xenoliths
The origin and evolution of the two types of Kerguelen clinopyroxene-bearing harzburgites and clinopyroxene-poor lherzolite can be summarized in two main steps, namely, early partial melting (15–25%, Grégoire et al., 1997) followed by metasomatism as a result of percolation of highly alkaline silicate melts into the previously depleted upper mantle. The metasomatism is cryptic in protogranular harzburgite and probably related to the percolation of a small volume of melt into the protogranular harzburgite, because it is manifested only in a slight enrichment in the most incompatible trace elements. The metasomatism is modal in poikilitic harzburgite and clinopyroxene-poor lherzolite, because incompatible trace enrichment is accompanied by crystallization of Mg-augite ± phlogopite ± amphibole. Some harzburgite samples display evidence of a later discrete metasomatic event evidenced by the crystallization of feldspar + olivine + rutile + ilmenite + armalcolite + chromite paragenesis.
The origin of clinopyroxene + phlogopite ± amphibole-bearing dunite is not clear. The dunite may represent a high-pressure cumulate of basaltic magma, as has been suggested for Hawaiian dunites (Sen, 1987; Clague, 1988), which was later metasomatized by highly alkaline silicate melts or carbonate-rich melts. Alternatively, anhydrous dunite may have formed by a reaction between harzburgites and basaltic melts that led to the dissolution of orthopyroxene and crystallization of olivine (Grégoire et al., 1997), and then been metasomatized by highly alkaline silicate melts or carbonate-rich melts that produced clinopyroxene + phlogopite or clinopyroxene + phlogopite + amphibole.
We favour the second hypothesis because: (1) the studied dunites display Ni, Sc and V contents similar to those of harzburgites; (2) in composite xenoliths, dunite hosts magmatic veins that show affinities with the tholeiitic–transitional and alkaline magmatic series from the Kerguelen archipelago (Grégoire, 1994; Grégoire et al., 1997); (3) not all Kerguelen dunite is similar to Type II Kerguelen xenoliths (peridotites, pyroxenites and metagabbros), which are high-pressure cumulates from the tholeiitic–transitional and alkaline magmatic series of the archipelago (Grégoire, 1994; Grégoire et al., 1997, 1998).
The reaction of mantle harzburgite with basaltic melts to form dunite, together with a large volume of basaltic intrusions at the crust–mantle boundary, may explain the very low MgO contents (typically 4–5%) of Kerguelen basalts that lack primary mantle-melt compositions (Gautier et al., 1990; Weis & Frey, 1996; Grégoire et al., 1998; Yang et al., 1998).
Trace element partition coefficients
Mineral-pair partition coefficients
Our comprehensive set of trace element data for mantle peridotites allows us to determine partition coefficients for a wide range of elements in mantle clinopyroxene, orthopyroxene, olivine, spinel, amphibole and phlogopite. To establish meaningful intermineral partition coefficients, chemical equilibrium is required. The protogranular, clinopyroxene-bearing harzburgite samples display mg-number in olivine < mg-number in orthopyroxene < mg-number in clinopyroxene, as do numerous Type I mantle peridotites that represent equilibrium phase assemblages (e.g. Frey & Prinz, 1978; Brown et al., 1980; Grégoire, 1994). The protogranular harzburgite samples are cryptically metasomatized but their mineral compositions do not vary within or between grains (Grégoire, 1994; Grégoire et al., 1997). This indicates major element equilibration. We therefore assumed trace element equilibration and calculated two-mineral partition coefficients for ol–opx, ol–cpx, opx–cpx, sp–ol, sp–opx and sp–cpx in these rocks.
Poikilitic cpx-bearing harzburgite, clinopyroxene-poor lherzolite and dunite were modally metasomatized. This is evidenced by: (1) the addition of clinopyroxene ± phlogopite ± amphibole; (2) reaction of orthopyroxene and spinel inclusions in poikilitic cpx; (3) the high Fe2O3 and TiO2 contents of spinel; (4) the higher Na2O, Al2O3, TiO2 and incompatible trace element contents of orthopyroxenes relative to protogranular orthopyroxenes; (5) the low mg-number of olivine. In dunite sample MG-91-143, amphibole has replaced interstitial clinopyroxene in a reaction relationship. Therefore, we calculated mineral-partition coefficients only for mineral pairs that crystallized from the same metasomatic agent: phlogopite–clinopyroxene in harzburgite and phlogopite-bearing, amphibole-free dunite samples, and amphibole–clinopyroxene and amphibole–phlogopite in the poikilitic harzburgite sample OB-93-5.
Partition coefficients (D) for trace elements in orthopyroxene and clinopyroxene pairs from protogranular harzburgite are <1 (Table 10) except for Co (2·53) and Ni (1·99). Values for other REE are two or three times those for La and Ce (Table 10). The greatest D values for incompatible trace elements are for Rb, Ba, Ti and Zr (Table 10). Our D values are much larger than those of Eggins et al. (1998). However, many interelement relationships are similar, e.g. DTi and DZr > DREE. This may reflect a difference in the composition of clinopyroxene. In our samples, it is a Cr-diopside that is poor in incompatible trace elements, whereas Eggins et al. studied an Mg-rich augite that contains significant amounts of incompatible trace elements [compare Tables 2 and 8 of this study with table 1 of Eggins et al. (1998)]. Spinel strongly concentrates V and HFSE (Nb, Ta, Ti and Zr), and has Dsp/ol and Dsp/opx > 1 for many trace elements (Table 10). Ni is partitioned preferentially into olivine but Co seems to be preferentially incorporated into spinel (Dsp/ol = 1·92).
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Partition coefficients between amphibole and clinopyroxene range from 0·45 to 3, except for Rb, Ba, Nb, Ta and Ti (Table 11 and Fig. 8a). Our results for Rb, Ba, Sr, Pb, U, Nb, Ti, Zr, La, Ce, Er, Yb and Lu agree with literature values, except that our D for Th is larger, and those of Hf and most REE (Nd to Ho) are smaller (Table 11 and Fig. 8a).
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Average partition coefficients between phlogopite and clinopyroxene for poikilitic harzburgite, phlogopite-bearing dunite and amphibole + phlogopite-bearing dunite show that phlogopite strongly concentrates Ba, Rb, Nb, Ta, Co, Ni and Ti, and perhaps Pb, but incorporates lesser amounts of REE and Y, Zr and Hf relative to clinopyroxene (Table 11, Fig. 8b). Values of Dphl/cpx for poikilitic harzburgite and phlogopite-bearing dunite resemble those found by Ionov et al. (1997)
. Partition coefficients between phlogopite and amphibole indicate that phlogopite is enriched in Co, Ni, Rb, Ba, Pb, Nb, Ta and Ti relative to amphibole (Table 11). Other trace elements are preferentially incorporated into amphibole, except for V, which is equally partitioned between the two (Fig. 8c). Our results for Ti, Rb, Sr and Ba agree with those from the literature except that our DZr is slightly larger and our DNb is >1 (Table 11).
Calculated mineral–melt partition coefficients
Data for trace element partitioning between olivine, orthopyroxene, spinel, amphibole, phlogopite and mafic silicate and carbonatitic melts are rare compared with those for clinopyroxene–melt partitioning (Green, 1994; Ionov et al., 1997). We have calculated mineral–melt D values for these phases by multiplying the average mineral–clinopyroxene ratio for each element by clinopyroxene–melt partition coefficients from the literature. For harzburgite and phlogopite-bearing dunite we used the Dcpx/sil (as discussed above); for the phlogopite + amphibole-bearing dunite MG-91-143 we used the Dcpx/carb melts (as discussed above). We assumed that Dmineral/cpx values at the subsolidus temperature are similar to those at near-liquidus temperature in basaltic systems. This assumption is controversial: some researchers have argued that temperature and mineral composition have little effect on D values (Ionov et al., 1997; Johnson, 1998), and others have provided experimental data indicating the opposite (e.g. Green, 1994; Blundy et al., 1998).
The calculated partition coefficients between olivine, opx, spinel and basic silicate melt are very low, except for DTiopx/melt, DZrsp/melt and the DTisp/melt (Table 12). Dopx/melt values are very small for La and Ce but much greater for other REE. The values compare well with the GERM partition coefficient compilation (http://www-ep.es.llnl.gov/germ/partitioning.html).
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Our estimates of phlogopite–mafic silicate melt partition coefficient for Rb, Ba, Th, U, Nb, Ta and Pb (Table 12) are larger than other published data (La Tourette et al., 1995; Foley et al., 1996; Ionov et al., 1997; see Appendix). Values for Sr are smaller and those for Ti (Table 12) are similar to those of Ionov et al. (1997). Our DRbphl/sil agrees with the value of 8·2 calculated by Ionov et al. (1997). The D value for Ce agrees with that calculated by Ionov et al. (1997; D 0·0006) but values for other REE are systematically larger (Table 12 and Appendix). The experimental data of La Tourrette et al. (1995) for La and Nd, and the measured data of Foley et al. (1996) for Ce, are larger than our values. The partition coefficients between phlogopite from sample MG-91-143 and carbonatitic melt indicate that Rb (and probably Ba), Nb, Ta and Ti behave as compatible elements in this chemical system and that phlogopite is a good residence site for those trace elements (Table 12). D values for Zr and Hf are less than those for basic silicate melt. The D values for REE progressively increase from LREE to HREE.
The calculated partition coefficients between amphibole and mafic silicate melt are given in Table 12. Ba and Ti are concentrated in amphibole relative to melt; other trace elements display D values ranging from 0·010 (U) to 0·44 (Rb). The D values for REE increase from La to Eu with nearly constant values from Gd to Lu. The same trend has been observed by Chazot et al. (1996). Zr and Hf have similar partition coefficients but those for Ta are less than those for Nb (Table 12). D values for Pb are larger than those for Th and U. D values for Ba and Th are close to values proposed by Chazot et al. (1996); DBa up to 1·59 and Brenan et al. (1995); DTh 0·017. D values for LREE and HREE agree with those of Chazot et al. (1996) and Witt-Eickschen & Harte (1994) but D values for MREE are smaller (Table 12). Although D values for Sr and Hf are less than literature values, those for Zr and Nb are similar to those obtained by Adam et al. (1993), Dalpe & Baker (1994), Witt-Eickschen & Harte (1994), Brenan et al. (1995), La Tourrette et al. (1995), Chazot et al. (1996) and Ionov et al. (1997).
Trace element residence sites
The trace element contents of constituent minerals (olivine + clinopyroxene + spinel ± orthopyroxene ± phlogopite ± amphibole) and their modal proportions can be used to calculate whole-rock compositions for comparison with the whole-rock analyses, to evaluate mass balance and the proportions of elements that reside in constituent minerals.
The clinopyroxene-poor lherzolite JGM-92-1c and the phlogopite-bearing-dunite MM-94-54 have incompatible trace element bulk-rock contents that can be easily explained by their constituent minerals, but all other samples display significant discrepancies between calculated and measured bulk-rock trace element compositions (Figs 9 and 10). Discrepancies are evident for Rb, Ba, Sr, Nb, Ta, Zr, Hf, Th, U, Pb and LREE; the MREE, HREE and transition trace elements are generally concordant. Samples that deviate from calculated trace element compositions contain veins and patches of metasomatic phases that were not analysed. Harzburgite samples GM-92-501, OB-93-3 and OB-93-5 show variable proportion of veins and patches that contain feldspar, olivine, ilmenite, rutile and armalcolite and chromite (Grégoire et al., 2000). However, harzburgite samples GM-92-453 and GM-92-502 display rare patches, which contain clinopyroxene ± carbonate ± chromite. Finally, phlogopite-bearing dunite samples GM-92-480 and MM-94-101 display few patches, which contain clinopyroxene ± amphibole ± biotite ± chromite ± rutile ± carbonate. We therefore conclude that metasomatic veins and patches contribute significantly to the Rb, Ba, Sr, Nb, Ta, Zr, Hf, Th, U, Pb and LREE contents of these seven peridotite samples. For example, a significant amount of Sr is probably hosted by feldspar, carbonate, phlogopite and amphibole occurring in the metasomatic veins and patches. Therefore, the major mineral phases (including the metasomatic phases) can account for the trace element budget without considering glass or fine-grained areas of metasomatic minerals. In addition, these samples do not show any significant trace element component in fluid and solid inclusions or at grain boundaries [compare with O’Reilly et al. (1991) and Eggins et al. (1998)]. These observations are consistent with those of Rosenbaum et al. (1996), who argued that fluid inclusions will dominate the incompatible element budget of typical mantle peridotite only if present in greater than sub-weight percent quantities.
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Our mass balance calculations estimate the fractional contribution by specific mineral phases to the calculated bulk-rock composition in samples where the calculated composition is close to the bulk-rock analysis (clinopyroxene-poor lherzolite JGM-92-1c, harzburgite GM-92-453 and dunite samples MM-94-54, MM-94-101 and GM-92-480). Clinopyroxene is the dominant host of REE, Sr, Y, Zr and Th; opx and olivine host HREE, especially Yb and Lu (Fig. 11). Olivine, the dominant mineral phase, is the major host of Ni (and Co) as well as for Pb and Sc. Olivine, when present in high modal abundances, contributes significantly to bulk-rock Nb, Rb, Ba and Th contents (Fig. 11), despite its very low contents of these elements. Opx in harzburgite and clinopyroxene-poor lherzolite is an important host for Sc, Ti, V, Zr, Y, Rb and HREE, again because of its high modal proportion in these rocks. In dunite samples, phlogopite is the principal residence site for Rb, Ba, Nb and Ti, but not Zr or Pb. Spinel hosts significant amounts of V and Ti. Despite its relatively large DHFSE compared with olivine, clinopyroxene and orthopyroxene, spinel does not contribute much of the bulk-rock budget of Nb, Ta, Zr and Ti, owing to its small modal abundance.
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SUMMARY AND CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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The trace element signatures of mantle xenoliths indicate that pervasive mantle metasomatism by melts with intraplate alkali mafic silicate affinities occurred on a large scale in the Kerguelen lithospheric mantle. The inferred metasomatic fluids range from highly alkaline mafic silicate magmas to CO2-rich silicate melts (‘carbonatitic’). Carbonatitic mantle metasomatism beneath the Kerguelen archipelago (Schiano et al., 1994
; Mattielli, 1996; Mattielli et al., 1999) is probably restricted in its impact. The dominant metasomatic agent is the mantle equivalent of highly alkaline silicate magmas that have erupted in the Kerguelen Archipelago and are observed as late-stage ‘lamprophyric dykes’. The petrological and geochemical characteristics of the two types of Kerguelen harzburgites and clinopyroxene-poor lherzolite can be explained by a two-stage process related to the origin and evolution of the Kerguelen archipelago. In stage I, partial melting associated with the formation of the Kerguelen oceanic lithosphere in the vicinity of the South East Indian Ridge results in harzburgite. In stage II, reaction between these harzburgitic residues and various alkaline mafic silicate to carbonatitic melts reflects a contribution from the Kerguelen mantle plume. Anhydrous dunites may also have formed by reaction between harzburgite and basaltic melts in the upper mantle during the early tholeiitic–transitional magmatic activity that formed the archipelago. Later metasomatism of the anhydrous dunites by alkaline mafic silicate to carbonatitic melts may then be related to the activity of the Kerguelen mantle plume in the within-plate setting of the islands.
The trace element budgets of the Kerguelen mantle rocks can be quantitatively accounted for by the major mineral phases in samples that do not contain significant metasomatic veins and patches. In these samples there is no significant concentration of trace elements on grain boundaries or in fluid and solid inclusions. This result is similar to that obtained by O’Reilly et al. (1991) for a more restricted element set for 12 xenoliths from western Victoria (Australia) and by Eggins et al. (1998) for two mantle peridotites from southeastern Australia.
Our results emphasize the major role of clinopyroxene as host for trace elements such as REE, Sr, Y, Zr and Th. Spinel is an important residence site of V, Pb, Sc and Ti, but its small modal abundance indicates that it has but a small effect on the trace element budget of these rocks. Olivine may be an important residence site of Pb and Sc, despite its paucity of trace elements, because it constitutes the major part of modes. If orthopyroxene is a major constituent, it may contribute significantly to rock budgets for Sc, V, Zr and Ti, as has been proposed by Rampone et al. (1991) and discussed by Xu et al. (2000). Phlogopite contributes to rock budgets for Ba, Nb, Ta and Ti.
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APPENDIX A |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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APPENDIX B |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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ACKNOWLEDGEMENTS |
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This work has been made possible by the generous assistance and technical expertise of N. J. Pearson, A. Sharma and C. Lawson (GEMOC Geochemical Analysis Unit). We thank T. H. Green, D. A. Ionov and N. J. Pearson for their helpful comments and English improvement. We very much appreciate the thoughtful and constructive reviews by Doug Smith, S. Sorensen and D. G. Pearson, which helped to improve the paper and clarify presentation. This work has been supported by Macquarie University Research Fellowship and Grant Schemes (M.G.), Australian Research Council Large and Small Grants (S.Y.O’R.) and the ARC International Fellowship Scheme (M.G.). Bertrand Moine acknowledges support from the ‘Région Rhône-Alpes, Programme EMERGENCE’. We also thank the following institutions for their support: the French Polar Research and Technology Institute (IFRTP, Brest, France), the French CNRS UMR-6524, the French Ministry of Education and Research and the University Jean Monnet (St Etienne, France). This is Publication 175 of the ARC National Key Centre for the Geochemical Evolution and Metallogeny of Continents (GEMOC).
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FOOTNOTES |
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*Corresponding author. Telephone: +61-2-9850-8362. Fax: +61-2-9850-8428. e-mail: sue.oreilly{at}mq.edu.au or soreilly{at}laurel.ocs.mq.edu.au
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REFERENCES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL METHODSPETROGRAPHY AND MODAL...MAJOR ELEMENT COMPOSITIONTRACE ELEMENT COMPOSITIONSSUMMARY AND CONCLUSIONSAPPENDIX AAPPENDIX BREFERENCES |
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