Journal of Petrology Advance Access originally published online on July 25, 2007
Journal of Petrology 2007 48(9):1725-1760; doi:10.1093/petrology/egm036
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Crystallization and Breakdown of Metasomatic Phases in Graphite-bearing Peridotite Xenoliths from Marsabit (Kenya)
1Institut De Géologie Et D’Hydrogéologie, Université De Neuchâtel, Rue Emile-Argand 11, CH-2009 Neuchâtel, Switzerland
2Institute of Geological Sciences, University of Bern, Baltzerstrasse 1, CH-3012, Bern, Switzerland
RECEIVED AUGUST 4, 2006; ACCEPTED JUNE 15, 2007
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ABSTRACT |
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Mantle-derived xenoliths from the Marsabit shield volcano (eastern flank of the Kenya rift) include porphyroclastic spinel peridotites characterized by variable styles of metasomatism. The petrography of the xenoliths indicates a transition from primary clinopyroxene-bearing cryptically metasomatized harzburgite (light rare earth element, U, and Th enrichment in clinopyroxene) to modally metasomatized clinopyroxene-free harzburgite and dunite. The metasomatic phases include amphibole (low-Ti Mg-katophorite), Na-rich phlogopite, apatite, graphite and metasomatic low-Al orthopyroxene. Transitional samples show that metasomatism led to replacement of clinopyroxene by amphibole. In all modally metasomatized xenoliths melt pockets (silicate glass containing silicate and oxide micro-phenocrysts, carbonates and empty vugs) occur in close textural relationship with the earlier metasomatic phases. The petrography, major and trace element data, together with constraints from thermobarometry and fO2 calculations, indicate that the cryptic and modal metasomatism are the result of a single event of interaction between peridotite and an orthopyroxene-saturated volatile-rich silicate melt. The unusual style of metasomatism (composition of amphibole, presence of graphite, formation of orthopyroxene) reflects low P –T conditions (
850–1000°C at < 1·5 GPa) in the wall-rocks during impregnation and locally low oxygen fugacities. The latter allowed the precipitation of graphite from CO2. The inferred melt was possibly derived from alkaline basic melts by melt–rock reaction during the development of the Tertiary–Quaternary Kenya rift. Glass-bearing melt pockets formed at the expense of the early phases, mainly through incongruent melting of amphibole and orthopyroxene, triggered by infiltration of a CO2-rich fluid and heating related to the magmatic activity that ultimately sampled and transported the xenoliths to the surface. KEY WORDS: graphite; peridotite xenoliths; Kenya Rift; modal metasomatism; silicate glass
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INTRODUCTION |
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During recent decades numerous studies have described different styles of metasomatism in the Earth's upper mantle. These studies, in combination with experimental work, show that the agents responsible for mantle metasomatism include mafic silicate melts, carbonate melts or C–O–H-rich fluids (e.g. reviews by Menzies & Hawkesworth, 1987
; Luth, 2003). Integration of textural, mineralogical and geochemical data from mantle samples highlights the complexity of mantle metasomatism. The data show that metasomatic features are controlled by: (1) the composition of the initial metasomatizing agent; (2) the pre-metasomatic composition and heterogeneity of the mantle rock; (3) the evolution of the physico-chemical parameters during metasomatism (e.g. porosity, fO2, P, T); (4) element fractionation processes during melt (fluid)–rock interaction; (5) mineral reactions. Several recent studies have demonstrated that compositional differences between metasomatic products in a single suite of mantle rocks do not necessarily imply different metasomatizing events (e.g. Bedini et al., 1997; Ionov et al., 2002, 2006; Bodinier et al., 2004; Rivalenti et al., 2004). Processes such as chromatographic fractionation and melt/fluid–rock reaction can generate compositionally distinct products, starting from one initial melt or fluid (e.g. Navon & Stolper, 1987; Harte et al., 1993; Godard et al., 1995; Vernières et al., 1997).
We present a study of metasomatized, graphite-bearing xenoliths from Marsabit (northern Kenya), to investigate melt–rock reaction and fractionation processes. The products of metasomatism are texturally and compositionally variable and unusual. An early Na–Si-rich assemblage comprises volatile-bearing phases (amphibole, phlogopite, apatite) and metasomatic orthopyroxene, as well as graphite. Associated with these phases is a second metasomatic assemblage of carbonate-bearing silicate glass patches (‘melt pockets’), similar to those described from many other xenolith suites (e.g. Frey & Green, 1974; Dautria et al., 1992; Ionov et al., 1993, 1994; Zinngrebe & Foley, 1995; Chazot et al., 1996a; Draper & Green, 1997; Yaxley et al., 1997; Coltorti et al., 2000; Laurora et al., 2001; Bali et al., 2002; Ban et al., 2005). Our approach is based on detailed investigation of trace element variation in minerals and integration of textural constraints. It emphasizes that the ‘exotic’ nature of the metasomatic products reflects a particular pre-existing physical and chemical environment—i.e. low ambient temperatures (and probably pressures) and relatively low oxygen fugacities—rather than an exotic initial metasomatic agent. We will show that metasomatism in this case does not necessitate multiple metasomatic events of different nature, but simply reflects the product of melt–rock reaction, possibly initiated by common, continental rift-related basanites. We further present mass-balance calculations, which show that the melt pockets were formed by in situ partial melting of the early assemblage (mainly amphibole and orthopyroxene), triggered by infiltration of a CO2-rich fluid.
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SAMPLE CONTEXT |
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The investigated mantle xenoliths were collected from scoriae of basanitic to alkali-basaltic Quaternary cinder cones of the Marsabit shield volcano (Volker, 1990
; Henjes-Kunst & Altherr, 1992). Magmatic activity in Marsabit started in the late Miocene, whereas rocks from the shield volcano yield Pleistocene to Quaternary ages (e.g. Key et al., 1987). Magmatism is thus clearly related to the development of the Kenya rift (i.e. the eastern branch of the East African Rift system, EARS; see Fig. 1). The Marsabit shield volcano is, however, located eastward of the main axis of the EARS, within the Anza Graben, the eastward continuation of the Turkana depression. The evolution of the lithosphere in this region includes multiple accretion during Pan-African orogenesis (formation of the Pan-African Mozambique belt from 720 to 550 Ma; e.g. Meert, 2003), followed by several phases of continental rifting, including formation of the Mesozoic–Paleogene Anza Graben, later cross-cut by the Tertiary–Quaternary EARS [see Fig. 1 and Morley (1999) for a review]. A more detailed description of the geological and geodynamic context has been given by Kaeser et al. (2006).
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In our previous study (Kaeser et al., 2006
) we used the Marsabit peridotite xenoliths to constrain the evolution of the lithospheric mantle beneath Marsabit. The peridotites were subdivided into four groups, including the porphyroclastic spinel harzburgites and dunites (Group III) investigated here (Table 1). For detailed descriptions of the textures and mineral chemistry of the primary assemblages of all four groups the reader is referred to Kaeser et al. (2006).
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Thermobarometry calculations revealed that all the porphyroclastic rocks experienced decompression and cooling from high pressures and temperatures (mainly preserved in formerly garnet-bearing Group I lherzolites; see Fig. 2) to low mantle P–T conditions of <900°C at <1·5 GPa (Kaeser et al., 2006
). The present study focuses on metasomatic features in the low-temperature Group III xenoliths. Textural and compositional characteristics suggest a genetic relationship with late EARS-related heating and cryptic metasomatism, leading to U–Th–Nb and light rare earth element (LREE) enrichment in clinopyroxene in the recrystallized Group II lherzolites. The latter were interpreted to reflect metasomatism by percolation of alkaline mafic melts, possibly parental to the Quaternary magmas of the Marsabit volcano (Kaeser et al., 2006), post-dating rift-related decompression and deformation. In contrast, metasomatic phases in the porphyroclastic Group I lherzolites (i.e. Ti-pargasite and rare phlogopite; Kaeser et al., 2006) are texturally and compositionally distinct, suggesting a different event. The trace element signatures of Ti-pargasite suggest that this metasomatic assemblage formed earlier, prior to rift-related decompression and deformation (Kaeser et al., 2006).
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ANALYTICAL METHODS |
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Modal compositions (Table 1) were determined from digitized images of thin sections and rock slice scans (for details, see Kaeser et al., 2006
). Minerals were analysed for major elements using a CAMECA SX 50 electron microprobe equipped with four wavelength-dispersive spectrometers (Mineralogisch-Petrographisches Institut, Universität Bern) and a CAMECA SX 51 electron microprobe with five wavelength-dispersive spectrometers (Mineralogisches Institut, Universität Heidelberg). Pyroxene, olivine and spinel were analysed using routine operating conditions (focused beam, accelerating voltage of 15 kV, beam current of 20 nA, and counting times of 20 s for most elements). Silicate glass was analysed with a larger beam size (10–15 µm) and a lower beam current (15 nA) to avoid Na loss during analysis. Natural and synthetic silicates and oxides were used as standards. Carbonates were analysed with a beam current of 10 nA and a beam size of 10 µm, using natural carbonates and sulfates as standards. The concentrations of CO2 (in carbonates) and Fe3+ (in spinel) were calculated based on stoichiometry. The raw data of all microprobe analyses were corrected using routine PAP software.
Trace element contents in minerals were analysed in-situ on polished thin sections (40–50 µm thick), using a laser ablation (LA) instrument equipped with a 193 nm ArF excimer laser (Lambda Physik, Germany) coupled to an ELAN 6100 (Perkin Elmer, Canada) quadrupole inductively coupled plasma mass spectrometer (ICPMS) at the Institut für Isotopengeologie und Mineralische Rohstoffe, ETH Zürich [see Pettke et al. (2004) and references therein for the instrumental setup, capabilities and operating conditions]. Raw data were reduced using the LAMTRACE program. Laser pit diameters were between 14 and 110 µm, depending on grain size and the absence or presence of mineral, fluid or melt inclusions, and of exsolution lamellae.
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PETROGRAPHY |
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Group III harzburgites and dunites display two texturally different types of modal metasomatic modification. The first type of metasomatism formed clusters and veinlets of volatile-bearing minerals (amphibole, phlogopite, ± graphite, ± apatite; Fig. 3a–e), associated with metasomatic orthopyroxene. The second type is characterized by the occurrence of ‘melt pockets’ (e.g. Frey & Green, 1974
; Ionov et al., 1994): patches and networks of silicate glass associated with micro-phenocrysts of clinopyroxene, olivine and chromite (Fig. 4a–f). Glass further contains globules of carbonates and vugs (Figs 4a–e and 5a, b). Melt pockets form at least partly at the expense of volatile-bearing minerals. Therefore, the volatile-bearing assemblage is referred to in the following as ‘early metasomatic’ (including metasomatic orthopyroxene referred to as Opxm). The later glass-bearing assemblage is termed ‘late metasomatic’ (Na-rich Cr diopside, Dim; olivine, Olm; chromite, Chrm). The pre-metasomatic peridotitic phases are termed ‘primary’ (olivine, ol-I; clinopyroxene, cpx-I; orthopyroxene, opx-I; spinel, spl-I).
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Based on petrographic investigations, we subdivided Group III peridotites into four types (Type III-a to -d). Textural and geochemical features show that the transition from Type III-a to Type III-c (+d) mainly corresponds to the transformation of modally non-metasomatized spinel harzburgite to pervasively overprinted amphibole harzburgite–dunite. Twenty out of 27 Group III samples investigated petrographically show evidence for modal metasomatism (i.e. they belong to Type III-b or -c).
Type III-a
Type III-a is classified as porphyroclastic, cpx-I-bearing spl harzburgite, devoid of volatile-bearing phases, but showing geochemical evidence of cryptic metasomatism (Kaeser et al., 2006, and this study). This rock type was originally garnet-bearing and was subjected to decompression, as is indicated by the presence of rare spl–opx–cpx symplectites (Table 1 and Kaeser et al., 2006).
Type III-b
Xenoliths of Type III-b are cpx-I-bearing spinel harzburgites; however, they contain metasomatic volatile-bearing phases (Table 1). One xenolith (Ke 1965/15) contains veinlets of pale green amphibole, phlogopite, graphite ± apatite (Figs 3e and 4b). Graphite occurs as macroscopic flakes (up to 500 µm long; Fig. 3e), similar to those described from cratonic, kimberlite-hosted peridotite xenoliths (Pearson et al., 1994). Cpx-I in the vicinity of veinlets is strongly resorbed and develops patches or lamellae of Ti-poor amphibole, indicating metasomatic replacement (Fig. 4b). In all Type III-b samples amphibole and clinopyroxene are partially replaced by the late glass-bearing assemblage (e.g. Fig. 4a). Cpx-I is overgrown by a secondary, green Na-rich Cr diopside (Dim) occasionally containing fluid, glass and/or carbonate inclusions (Fig. 2c). Dim further occurs as micro-phenocrysts in melt pockets, sometimes enclosing (resorbed?) orthopyroxene.
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Type III-c
Type III-c samples are harzburgites and dunites devoid of cpx-I. Samples transitional between Type III-b and -c still contain relics of cpx-I. Metasomatic products in Type III-c xenoliths generally occur as clusters of amphibole, phlogopite, graphite and rare apatite. The primary assemblage is dominated by ol-I. Green, Ti-poor amphibole and phlogopite typically form aggregates around relic, skeletal spl-I (Fig. 3c and d). Dunite xenolith Ke 1965/25 is unusual as it contains centimetre-scale clusters of large, euhedral twinned amphibole crystals (Fig. 3a) with small phlogopite inclusions. Graphite is associated with amphibole and phlogopite (sometimes found as inclusions in the latter; Fig. 3d) or orthopyroxene (Fig. 3c). Straight contacts between amphibole, phlogopite and ol-I indicate textural equilibrium. Phlogopite also occurs as euhedral, randomly dispersed flakes in orthopyroxene (Fig. 3b). This, together with compositional arguments (e.g. very low Al contents; see below), suggests that orthopyroxene is related to metasomatism (Opxm). Sample Ke 1965/3 also contains early metasomatic interstitial apatite.
Late metasomatic assemblages are ubiquitous in Type III-c xenoliths. They form melt pockets, interconnected by glass films along olivine grain boundaries. Euhedral apple-green Na-rich Cr diopside micro-phenocrysts (Dim; up to 0·3 mm long) in individual melt pockets commonly have the same crystallographic orientation (Fig. 4a), suggesting growth at the expense of an earlier phase. Back-scattered electron (BSE) imaging reveals patchy zoning (Figs 4e and 5a). Melt pockets in several samples enclose relic, resorbed amphibole grains (Fig. 4d). Phlogopite (e.g. in Ke 785*) appears to be less resorbed (Fig. 4d). Early amphibole is rimmed by a kaersutitic selvage (Fig. 4d), indicating that the early amphibole is not in chemical equilibrium with the silicate glass. When in contact with glass, Opxm is also strongly resorbed. Several features suggest incongruent breakdown by a process similar to the reaction opx + melt 
ol + cpx + modified melt (e.g. Shaw et al., 1998): thin glass films are present on orthopyroxene cleavage planes, late euhedral Olm + Dim ± Chrm nucleated within orthopyroxene and Olm overgrows orthopyroxene (Fig. 3b; development of strongly concave opx grain boundaries).
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Glass in Type III-b and -c samples contains considerable amounts of rounded carbonates (e.g. Figs 4e and 5a), texturally similar to interstitial, glass-related carbonate globules from other xenolith suites (e.g. Ionov et al., 1993
, 1996; Norman, 1998; Lee et al., 2000; Laurora et al., 2001; Bali et al., 2002; Demény et al., 2004). BSE imaging and qualitative energy-dispersive spectrometry mapping highlight that carbonate globules consist of clusters of sometimes euhedral Mg-calcite or dolomite crystals, occasionally showing oscillatory zoning (Fig. 5a and b). In sample Ke 1965/15, Mg-calcite is in contact with graphite flakes (Fig. 4c), indicating nucleation of carbonate on graphite. Glass in several samples contains rounded blebs of apatite (Fig. 4b), as well as minor amounts of Fe–Ni sulphide globules. Abundant vugs (e.g. Fig. 4a and d) in all types of silicate glass are possibly remnants of an exsolved vapour phase. Sample Ke 1959/27 (Type III-c) is unusual in several aspects. It contains kaersutite phenocrysts hosted in the glass. Further, Dim micro-phenocrysts are overgrown by Ti-augite rims (Fig. 4f), and the glass is of an untypical dark brown colour. This xenolith records interaction with a melt similar to the host basanite during ascent to the surface.
Type III-d
Type III-d is a unique sample (Ke 1970/6) of metasomatic coarse-grained spinel harzburgite (grain size 
4 mm; Fig. 3f). Olivine and orthopyroxene are almost strain free. Spinel occurs as rounded, black grains (chromite) surrounded by a cluster of apple-green Na-rich Cr diopside (Fig. 3f) containing glass inclusions. Volatile-bearing phases are absent. Several orthopyroxene grains show numerous melt inclusions forming continuous films or trails along cleavage planes. These remarkable textural features can be explained by almost complete annealing of a former Type III-c spinel harzburgite containing silicate glass + Dim micro-phenocrysts, which subsequently recrystallized to form the clinopyroxene clusters around spinel and orthopyroxene with melt inclusions.
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MINERAL COMPOSITION |
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Compositional data (major and trace elements) for non-metasomatic primary phases from all peridotite types, as well as for Ti-pargasite and phlogopite from Group I samples have been given by Kaeser et al. (2006
). A more complete dataset with respect to the metasomatic phase assemblages discussed in this paper is available at http://petrology.oxfordjournals.org.
Early metasomatic assemblages
Clinopyroxene
‘Cryptically’ metasomatized cpx-I (Type III-a and -b) are diopsides (Table 2) with high Mg-numbers and low Ti contents typical for mantle harzburgite (e.g. Pearson et al., 2003). The major element compositions of ‘unmetasomatized’ cpx-I from Type III-a and -b xenoliths are similar. However, Type III-b samples occasionally contain clinopyroxene that is texturally identical to cpx-I but contains significantly higher amounts of Na2O (up to 2·54 wt%; see Fig. 6a), pointing to cryptic metasomatism. Trace element analyses (Table 2) show enrichment of U, Th, Sr, and LREE relative to middle REE (MREE), variable depletions in heavy REE (HREE), and marked negative anomalies for high field strength elements (HFSE; Nb, Ta, Zr, Hf, Ti; Fig. 7a). Trace element enrichment is strongest in the high-Na cpx-I (Fig. 7a).
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Orthopyroxene
Opx-I from Type III-a and -b xenoliths is enstatite-rich (Mg-number of 91·2–92·5). It is LREE-depleted and slightly enriched in U and Th, indicating cryptic metasomatism (Table 3). Opxm (Types III-c and -d) has higher Mg-numbers than opx-I (up to 93·1) and is extremely Al-poor (0·2–1· 0 wt% Al2O3). LREE and MREE are slightly enriched relative to HFSE, which is unusual for mantle-derived orthopyroxene (Rampone et al., 1991
).
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Amphibole
Green amphibole (Type III-b and -c) is mostly Ti-poor (Table 4) with very high Si and Na contents and high Mg-numbers (Fig. 8a and b). Low-Ti amphibole in Type III-c is Mg-katophorite (Fig. 8a; according to Leake et al., 1997
), which is an unusual composition for mantle-derived amphibole. Type III-b samples contain less silicic edenite–Mg-katophorite (Fig. 8a). Cr contents (Fig. 8b) are high only in amphiboles related to metasomatic replacement of spl-I. Fluorine and Cl were not detected by electron microprobe. Trace element patterns show enrichment in U and Th, strongly negative Hf and Zr anomalies, and enrichment of LREE over HREE (Fig. 9a and b). Amphibole from sample Ke 1965/15 has the lowest HREE concentrations, but in contrast, exhibits a conspicuous positive Ti anomaly (Fig. 9b). Nb and Ta contents vary considerably, with the highest abundances in euhedral Mg-katophorite (Table 4; Fig. 9c). Compared with typical mantle amphiboles (e.g. Ionov et al., 1997) the Mg-katophorites show a large compositional range (e.g. Fig. 10a). They are, however, remarkably depleted in large ion lithophile elements (LILE; e.g. 1·26–83 µg/g Ba compared with 75–1400 µg/g as reported for mantle amphibole from several off-craton xenolith suites; see Ionov et al., 1997), and show low LILE/LREE ratios (Fig. 10b).
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Phlogopite
Phlogopites (Table 5) have very high Na2O contents (up to 2·88 wt%). Ti concentrations are mostly very low, and Mg-numbers are high (>93). Fluorine and Cl were not detected by electron microprobe. Generally, Na-rich phlogopites have higher U and Th concentrations and lower HFSE contents than phlogopite from the Group I (grt)–spl lherzolites (Kaeser et al., 2006
; Fig. 9a). Some trace elements show very broad ranges between samples (e.g. Nb ranging from 1· 02 µg/g in Ke 785* to 47·9 µg/g in Ke 1965/3).
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Apatite
Considerable differences were found between the two textural types of apatite (early in sample Ke 1965/3 and phosphate globules associated with glass in Ke 1959/15). Microprobe analyses only poorly satisfy apatite stoichiometry and show fairly low totals (Table 5). This is probably due to the abundant inclusions, which are inevitably included in the analysis. Early apatite is F-rich (1·89–2·53 wt% F) with variable Cl contents (0·92–1·57 wt%). It has extremely high REE and SrO contents (3·38–5·24 wt%; see Table 5). The MgO, Na2O and FeOtot contents are typical for upper mantle apatites (O’Reilly & Griffin, 2000
). In contrast, apatite globules in silicate glass are Cl-rich (2·96–3·14 wt% Cl) and have lower SrO and Na2O, but slightly higher FeOtot contents (Table 5). Both types of apatite are strongly LREE enriched [(La/Yb)N up to 783] and are the major host for U and Th (up to 33 500 times primitive mantle values), whereas HFSE concentrations are very low (Fig. 11b).
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Late (glass-bearing) assemblages
Silicate glass
Representative results of microprobe and LA-ICPMS analyses of silicate glass are given in Table 6 and selected compositional plots are displayed in Fig. 12a–f. Generally, glass is siliceous (up to 66 wt% SiO2) and peraluminous (12·45–25·07 wt% Al2O3) with characteristics (very low MgO and FeO, high alkalis) similar to those in many other xenolith suites (e.g. Ionov et al., 1994
; Zinngrebe & Foley, 1995; Chazot et al., 1996a; Neumann & Wulff-Pedersen, 1997; Draper & Green, 1997; Varela et al., 1999; Coltorti et al., 2000; Laurora et al., 2001; Shaw & Klügel, 2002; Ban et al., 2005). Mg-numbers are commonly high (up to 70; Fig. 12b). Considerable compositional heterogeneity can occur on a very small scale within individual samples. In sample 785*, for examples, glass in the vicinity of Opxm contains up to 65 wt% SiO2, whereas a nearby melt pocket (at less than 5 mm distance) contains glass with 58 wt% SiO2.
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With respect to trace elements, glass is always LREE enriched [(La/Yb)N up to 77] at variable absolute REE concentrations (Fig. 13; YbN from 1· 4 to 91· 0). Zr, Hf and Ti are depleted relative to the REE, and Nb and Ta relative to U and Th. Trace element compositions always mirror the composition of the associated earlier phases (i.e. amphibole in most cases but also clinopyroxene, orthopyroxene, phlogopite and apatite).
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Glass in Group III xenoliths is clearly different compared with Ti-rich subalkaline glass occasionally found in Group I lherzolites from Marsabit (related to melting of Ti-pargasite; see Kaeser, 2006
, and Fig. 12a–f). Also, glass is generally clearly distinct from the xenolith-hosting basanites (Fig. 12a–f). The exception is glass within the Type III-c dunite Ke 1959/27 and from the margin of xenolith Ke 1965/15 (contact with host lava), which show some important differences compared with glass from the other samples; that is, higher FeO and lower SiO2 contents, similar to the composition of the host basanite (Figs 12a–f and 13).
Clinopyroxene micro-phenocrysts
Clinopyroxene micro-phenocrysts (Dim) are Na-rich Cr diopsides (up to 3·23 wt% Na2O and 6·45 wt% Cr2O3; Fig. 6a and b) with low TiO2 contents (<0·5 wt%) and high Mg-numbers (up to 94·9). In all samples, Dim are strongly zoned with Cr and Na contents generally decreasing from cores to rims (Fig. 6b). Dim in contact with basanitic glass in the kaersutite-bearing dunite (Ke 1959/27) has resorbed rims, overgrown by Ti-augite (Fig. 4f) with significantly lower Na2O (<1 wt%) contents and Mg-numbers (
85; see Table 2).
Trace element patterns of Dim are similar in shape to the associated glass (except LILE: Cs to U; Fig. 13). Negative HFSE anomalies are mostly due to higher REE abundances; that is, the absolute concentrations of HFSE are similar to those of cpx-I from Type III-a and -b (see Fig. 7a; no real HFSE depletion). Texturally equilibrated clinopyroxene in the Type III-d spl harzburgite (Ke 1970/6; Fig. 3f) is a Na-rich Cr diopside with similar major (Fig. 6a) and trace element characteristics (Fig. 7b) similar to those of the glass-hosted micro-phenocryst cores.
Olivine, spinel, kaersutite
Compared with ol-I, glass-related Olm is more magnesian (Table 3; up to Fo94·6) and enriched in Ca. Ol-I in contact with glass develops strong compositional gradients approaching the composition of Olm at the rims. Spinel micro-phenocrysts (Chrm) are Mg-rich chromites (Cr-number up to 66·5), enriched in (calculated) Fe2O3 compared with spl-I.
Kaersutite occurs as tiny rims on amphibole in Type III-b and -c, and as micro-phenocrysts in Type III-c sample Ke 1959/27. Reaction rims are characterized by higher Mg-numbers than the original amphibole. They are too narrow to be analysed for trace elements. Kaersutite micro-phenocrysts in sample Ke 1959/27 have low Mg-number (about 82) and high K2O contents (Table 4). Major element characteristics as well as trace element patterns (convex-upward REE patterns, high Nb–Ta, low U–Th; Fig. 9d) are similar to those of hornblendites and amphibole megacrysts from alkaline magmas (e.g. Shaw & Eyzaguirre, 2000).
Carbonates
Most glass-related carbonate globules contain Mg-calcite (Table 7) with an average Mg-number of 96·5 and Ca/(Ca + Mg) between 0·87 and 0·98. In sample Ke 1965/3, MgO correlates positively with SrO (up to 5·35 wt% MgO and 0·89 wt% SrO). Other elements show very low concentrations (Table 7). One sample (Type III-c Ke 1959/15) contains zoned crystals ranging from dolomitic to calcitic in composition (Fig. 5a). Carbonates have low REE contents with flat to slightly LREE-enriched patterns (Fig. 14). Extended trace element plots exhibit a very broad range of concentrations in different samples (e.g. Pb; Fig. 14). Large positive peaks for U, Pb and Sr are typical. In addition, U and Nb are strongly enriched compared with Th and Ta, respectively. Hf, Zr and Ti show very low concentrations. Overall, major and trace element characteristics of the carbonates are similar to those of carbonates found in other peridotite xenoliths (e.g. Ionov, 1998; Lee et al., 2000; Laurora et al., 2001; Ionov & Harmer, 2002; see Fig. 14).
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FORMATION OF THE EARLY ASSEMBLAGES |
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Despite large compositional variations, numerous common textural and compositional features indicate that the early metasomatic assemblages in all Group III samples formed during a single metasomatic event. In the discussion below, we argue that during a first metasomatic stage, the primary (pre-metasomatic) phase assemblage in Group III spinel harzburgites was partially replaced by metasomatic minerals, finally resulting in clinopyroxene-free amphibole- and phlogopite-bearing dunite and harzburgite (Type III-c). The initial metasomatic agent was probably an evolved CO2–H2O-rich melt, and the transition from peridotites of Type III-a to -c can be explained by progressive melt–rock reaction.
Transition from cryptic (Type III-a) to modal (Type III-b and -c) metasomatism
Formation of new phases requires (1) crystallization from a melt or fluid phase (decreasing melt or fluid mass) and/or (2) melt–solid or fluid–solid reaction. Textural features such as spinel being replaced by amphibole and phlogopite (e.g. Fig. 3c) indicate melt(fluid)–rock reaction. The geochemical signatures observed in the different peridotite types thus probably reflect pre-existing compositional heterogeneities coupled with metasomatic modification. This is illustrated best in sample Ke 1965/1 (Type III-a), the least metasomatized xenolith in this study. It can be subdivided into three parts containing clinopyroxene with different REE signatures (Zones A, B and C in Fig. 15). The LREE are increasingly enriched from Zone A to C, indicating increasing cryptic metasomatism. HREE abundances, on the other hand, are very heterogeneous in the least metasomatized Zone A, moderately depleted in Zone B, and homogeneously low in the most metasomatized Zone C (containing graphite; Fig. 15), similar to REE patterns in clinopyroxene from the amphibole- and phlogopite-bearing Type III-b sample (Ke 1965/15; Fig. 7a).
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The decoupling of strongly (LREE, Th, U) from moderately incompatible elements (HREE) can be explained by chromatographic fractionation during percolation of small amounts of melts or fluids through an initially non-metasomatized peridotite (Navon & Stolper, 1987
; Godard et al., 1995; Bedini et al., 1997; Vernières et al., 1997; Ionov et al., 2002). Upon reaction with the wall-rock, a fluid or melt will become progressively enriched in incompatible elements as a function of increasing distance from its source. This is due to the fact that elements with higher Dmineral/melt values (e.g. HREE) are selectively removed from a percolating liquid when it exchanges with the surrounding minerals. The results are faster-moving chemical fronts of incompatible elements compared with those of more compatible elements. This is consistent with the geochemical signature of cpx-I from the least metasomatized part of sample Ke 1965/1 (Zone A; Fig. 15), which already started to re-equilibrate with respect to the incompatible elements (i.e. addition of Th, U and LREE) whereas more compatible elements still reflect pre-metasomatic heterogeneities. In this case, these are extreme HREE variations, probably reflecting the earlier garnet breakdown (Kaeser et al., 2006). On the other hand, more intensely metasomatized parts (e.g. Zone C in sample Ke 1965/1 and sample Ke 1965/15) are characterized by clinopyroxene trace element patterns where mildly incompatible elements have also been modified (e.g homogeneously depleted HREE; Figs 7a and 15).
Similar fractionation processes can explain the modal and compositional differences between Type III-b and -c. Compared with Type III-b, higher abundances of metasomatic phases in Type III-c suggests that higher amounts of fluid or melt crystallized in, or passed through, these rocks. The amphibole-rich dunite (Ke 1965/25), for example, contains low-Ti Mg-katophorite exhibiting the lowest (La/Yb)N and (La/Ce)N ratios, and high (Nb/La)N and (Nb/Th)N ratios (Fig. 10a and c), indicating a moderately fractionated metasomatic agent. Mg-katophorite from the phlogopite-rich harzburgite (Ke 785*; Type III-c), and amphiboles from the more cpx-rich harzburgites (Type III-b) show progressively more fractionated patterns (e.g. Fig. 10c) and decreasing abundances in ‘amphibole-compatible’ elements (e.g. Nb; Fig. 9a). Thus, the transition from Type III-b to -c could reflect decreasing distance between the site at which the xenolith was sampled and the source of the ‘initial’ metasomatizing agent (e.g. Zanetti et al., 1996; Ionov et al., 2002; Bodinier et al., 2004). Judging from the abundance of modal metasomatic phases, the amphibole dunite (Ke 1965/25) was probably closest to the source of the metasomatizing agent. The phlogopite-rich nature of Ke 785*, which indicates higher K and LILE (Cs, Rb, Ba) contents in the remaining agent, is in agreement with a more fractionated melt or fluid, at somewhat greater distance from the metasomatic source.
The Nb abundances in amphibole also agree with this model as well. For example, amphibole from Type III-c Ke 1965/25 has comparatively high Nb concentrations (Fig. 9c), in line with relatively high DNbamph/melt values constrained by different studies (e.g. Ionov & Hofmann, 1995; LaTourrette et al., 1995). In Ti-poor systems, such as those of this study, Nb uptake by amphibole is even enhanced as a result of increased DNb values (Tiepolo et al., 2001). Consequently, in an event during which amphibole crystallizes continuously, the remaining fluid or melt will become progressively depleted in Nb (Ionov et al., 2002), resulting in formation of Nb-depleted amphiboles further away from the source (Zanetti et al., 1996; Ionov et al., 2002); such Nb-depleted amphibole is found in sample Ke 785*, as well as in Type III-b xenoliths (e.g. Ke 1959/15; Fig. 9a). Amphibole signatures from xenolith suites modelled successfully in the context of percolative trace element fractionation (e.g. Spitsbergen; Ionov et al., 2002) show similar ranges (e.g. in Nb/Th ratios; Fig. 10a) to those documented for our samples. This further supports the hypothesis that similar percolation mechanisms acted in the mantle beneath Marsabit.
The role of pre-existing heterogeneities and exotic minerals
Although a model based on trace element fractionation during melt percolation is supported by textures and can account for many compositional features, some compositional peculiarities remain unexplained. First, clinopyroxene and amphibole in several samples (Ke 785*, 1965/1 and 1965/15) are characterized by extremely low HREE abundances, more than one order of magnitude lower than in amphiboles from the strongly metasomatized amphibole dunite Ke 1965/25 (Fig. 9a–c). The main phases that crystallize during metasomatism are low-Ti Mg-katophorite and phlogopite. However, DHREEamph/melt data (e.g. LaTourrette et al., 1995) indicate that amphibole crystallization alone is unlikely to deplete the metasomatic agent to such a high degree in HREE. The same probably accounts for the extreme Nb depletion in amphiboles from Ke 785*, relative to the other Type III-c xenolith (Ke 1965/25; see also Ionov et al., 2002).
Further, clinopyroxene, phlogopite and amphibole in the Type III-b sample Ke 1965/15, on one hand, all show strongly fractionated REE patterns (e.g. high La/Ce in amphibole; Fig. 10c) but, on the other hand, are characterized by a positive Ti anomaly, decoupled from other HFSE (e.g. Hf and Zr; Fig. 9b). Trace element fractionation during amphibole crystallization (i.e. formation of Type III-c amphibole dunites), however, should lower Ti contents in the remaining fluid or melt (e.g. Moine et al., 2001).
Based on textures, mineralogy and dominant common compositional features, the hypothesis that Type III-b and -c xenoliths were metasomatized by completely independent metasomatic agents can be ruled out. Therefore, two other possibilities remain to account for the inconsistent observations described above: (1) inference of other phases, reactants or products, that are able to fractionate specific trace elements, present in the mantle beneath Marsabit but not present in the few investigated samples; (2) pre-metasomatic compositional heterogeneities, similar to the inherited HREE variation in Type III-a clinopyroxene (see above).
Garnet would be an important sink for HREE, but metasomatism occurred in the spinel field, after decompression, with maximum pressures of <1·5 GPa (Kaeser et al., 2006), so garnet had already broken down to symplectites (Fig. 15). Percolation of melt through garnet-bearing websterites, prior to metasomatism in Ke 1965/1 and 1965/15, could be an alternative. Websterites have been interpreted to occur in close spatial relationship with porphyroclastic peridotites in the shallow mantle beneath Marsabit (Kaeser et al., 2006), and they occasionally contain rutile and ilmenite, which could further account for the strong Nb and Ta depletion in some samples (e.g. Ke 785*). The available compositional data from the websterites, however, do not indicate metasomatic modification of websteritic garnet (Olker, 2001; B. Kaeser, unpublished data) nor compositional similarities to Group III peridotites.
Careful study by SEM rules out the presence of exotic phases such as zircon or Ti-oxides, described from other xenolith suites (e.g. Bodinier et al., 1996; Rudnick et al., 1999). Nevertheless, it cannot be excluded that the metasomatizing agent percolated through mantle domains containing these phases before infiltrating the studied rocks.
The preservation of pre-metasomatic compositional heterogeneities is another possibility that must be considered. That such ‘rock signals’ partly overlap with the ‘metasomatic signal’ is indicated by the HREE variation in Type III-a clinopyroxene (see above) and also by other compositional parameters. For example, both Mg-katophorite and phlogopite growing at the expense of spinel have very high Cr2O3 contents, whereas euhedral Mg-katophorite in the amphibole dunite (Ke 1965/25) or in phlogopite inclusions in Opxm is comparably Cr2O3-poor (Fig. 8b; Tables 4 and 5). Pre-existing heterogeneities would essentially be preserved through elements that are relatively compatible in a given mineral (i.e. that are not mobilized during metasomatism). This would be in agreement with the preserved HREE variations in Type III-a clinopyroxene, high Cr2O3 in amphiboles and phlogopite, and Ti spikes in amphibole from Ke 1965/15.
Further, very high Mg-numbers in replacive amphiboles support melt (fluid)–rock interaction at low melt (fluid)/rock ratios, where Mg and Fe contents are buffered by the ambient peridotite (Kelemen et al., 1995; Tiepolo et al., 2001). The slightly lower Mg-number of euhedral amphiboles in Ke 1965/25, combined with their lower Cr2O3 contents (Fig. 8b), suggests a higher melt (fluid)/rock ratio, which is further supported by considerably lower Mg-number in associated ol-I (Fo88; Table 3).
The limited number of investigated samples does not allow us to decide which of the processes described above acted, and to what extent. However, given the heterogeneity observed in the numerous non-metasomatized xenoliths from Marsabit (Kaeser et al., 2006), pre-metasomatic features certainly affect the final product of the melt–rock reaction process to a considerable degree.
The nature of the metasomatizing agent(s)
Graphite stability and oxygen fugacity considerations
The Marsabit xenoliths are unusual in that they occur in an off-cratonic setting yet bear macroscopic graphite. The latter usually either occurs in kimberlite-hosted peridotites (on-craton) or in pyroxenites from alpine-type peridotite bodies (e.g. Pearson et al., 1994). In the phlogopite peridotites from Finero, graphite has recently been reported as nanoscopic layers in phlogopite (Ferraris et al., 2004). In xenoliths from Marsabit, graphite occurs in the same textural relationship as in kimberlite-hosted peridotite xenoliths (i.e. millimetre-sized flakes, typically associated with orthopyroxene; see Fig. 3c–e and Pearson et al., 1994). Graphite inclusions in phlogopite and Opxm could indicate crystallization during metasomatism by a ‘reduced’ C-bearing fluid or melt. However, precipitation of graphite more probably reflects the relatively reducing conditions in the uppermost mantle prior to metasomatism. We used the oxygen barometer of Ballhaus et al. (1991), based on ferric iron contents in spinel coexisting with olivine, to constrain oxygen fugacity (fO2) conditions prior to metasomatism in the Marsabit xenoliths (Fig. 16). The results show that non-metasomatized Group I and IV samples reflect similarly reducing conditions to the cryptically metasomatized Type III-a samples.
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Although assessing Fe3+/Fe2+ ratios in spinel based on microprobe analyses can be problematic (Wood & Virgo, 1989
; Ballhaus et al., 1991), all microprobe analyses were obtained using the same operational settings and, therefore, give at least the relative fO2 variation between samples. The relatively reducing conditions in the pre-metasomatic, cold and highly deformed mantle, well below carbon–carbonate–silicate equilibria (Mg2Si2O6 + 0·5 CaMg(CO3)2 = Mg2SiO4 + 0·5 CaMgSi2O6 + C + O2; Eggler & Baker, 1982; see Fig. 16) and close to the CCO oxygen buffer are in agreement with the presence of graphite in the Group III harzburgites. Moreover, the P–T conditions at which the Group III xenoliths partially re-equilibrated are close to the subsolidus decarbonation of model carbonated lherzolite (Fig. 2; see Lee & Wyllie, 2000, and references therein), in agreement with the observed phase assemblage including orthopyroxene, olivine, clinopyroxene and graphite.
Evidence for low-temperature SiO2–Na2O–CO2–H2O-rich melts or fluids
Potential metasomatic agents, inferred from the presence of graphite, are C–O–H-rich fluids, carbonate melts or CO2-bearing silicate melts. The resulting phase assemblages and geochemical signatures are the result of infiltration of fluid or melt into a cold (750–850°C), shallow (<1·5 GPa; Fig. 2), heterogeneous and locally reduced lithospheric mantle. Metasomatism under such conditions in an open system must result in complex geochemical signatures that reflect the degree of fluid–rock reaction. This makes it particularly difficult to constrain the exact nature of the ‘original’ metasomatic agent.
Type III-c xenolith Ke 1965/25 contains the highest fraction of metasomatic phases and the least fractionated amphibole trace element patterns. The ‘magmatic’ textures of the Mg-katophorites in this sample suggest direct crystallization from a melt or a fluid rather than replacement of a pre-existing phase. Therefore, hypothetical liquids in equilibrium with the amphiboles in this xenolith could yield the composition of the initial metasomatic agents. Figure 17 illustrates the trace element patterns of such melts and fluids, calculated using sets of experimentally or empirically determined mineral/melt (fluid) partition coefficients, including carbonate melt, silicate melt and hydrous fluid (for references see caption to Fig. 17).
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Common features for all calculated metasomatic agents are negative Zr–Hf–Ti anomalies and LREE enrichment (Fig. 17). However, previous studies have shown that these signatures are indicative of almost all types of metasomatism, including interaction of peridotite with carbonatitic melt (Green & Wallace, 1988
; Dautria et al., 1992; Hauri et al., 1993; Rudnick et al., 1993; Ionov et al., 1996; Yaxley et al., 1998; Coltorti et al., 1999), with hydrous fluids and/or melts, for example, originating from dehydration of altered subducting lithosphere (Maury et al., 1992; Zanetti et al., 1999; Grégoire et al., 2001; Laurora et al., 2001; Parkinson et al., 2003), and interaction of peridotite with volatile- and Si-rich melts that evolved through reactive porous flow (Bedini et al., 1997; Ionov et al., 2002, 2006; Bodinier et al., 2004; Rivalenti et al., 2004).
In the case of the Marsabit xenoliths, integration of the trace element characteristics, thermobarometric inferences and mineralogical constraints favours a Na2O–SiO2–CO2–H2O-rich silicate melt. Carbonated melts can probably be ruled out on major element and mineralogical considerations. First, amphiboles are Ca-poor, in conflict with carbonate melt metasomatism, which is generally characterized by Ca enrichment (Rudnick et al., 1993; Neumann et al., 2002). Further, orthopyroxene is a stable phase and probably even crystallizes during metasomatism in some samples (e.g. Ke 785*). At near-solidus conditions coexistence of orthopyroxene and carbonate melt is possible only at pressures 2·0 GPa (see Fig. 2; e.g. Lee & Wyllie, 2000, and references therein), which is considerably higher than the maximum pressures indicated by the composition of spinel (<1·5 GPa). At such low pressures, metasomatism by carbonated melts will convert lherzolite and harzburgite immediately into wehrlite, which is the major peridotite type in all xenolith suites where carbonate melt metasomatism is convincingly demonstrated (Rudnick et al., 1993; Yaxley et al., 1998; Neumann et al., 2002). In Marsabit, wehrlites are absent, however.
An Na2O–SiO2–CO2–H2O-rich melt or fluid, on the other hand, would explain the Na-rich nature of amphibole and phlogopite, as well as the presence of orthopyroxene. Growth of metasomatic (and Al-poor) orthopyroxene has been described from several xenolith localities where peridotites were infiltrated either by hydrous fluids or Si-rich melts (Smith et al., 1999; McInnes et al., 2001; Morishita et al., 2003). Na-rich phlogopite is commonly found as a reaction product between anhydrous mafic to ultramafic rocks and H2O-rich evolved liquids (Costa et al., 2001, and references therein) or H2O-bearing fluids (e.g. Odling, 1994). The latter study was an experiment at 2·0 GPa simulating the effects of decreasing temperatures on the phase assemblages and their composition during metasomatism using a model peridotite + C–O–H fluid. The coolest part of Odling's (1994) experiment (<1075°C) contained a strikingly similar phase assemblage (katophorite, phlogopite with 2 wt% Na2O, and Al-poor orthopyroxene) to the Marsabit samples, interpreted to have formed during sub-solidus fluid–solid interaction. A compositionally similar phase assemblage (high amounts of opx, Si-rich hornblende, Na-rich phlogopite) resulted from experiments simulating peridotite interaction with Si-rich melt (dacite) at lower P–T conditions (950°C at 0·92 GPa; Prouteau & Scaillet, 2003).
In this study, the trace element signatures of Group III-c Mg-katophorite and the calculated metasomatic agents further suggest that the metasomatizing agent was a melt rather than a fluid. Hypothetical fluids in equilibrium with euhedral Mg-katophorite from sample Ke 1965/25 have high Nb/LILE (e.g. Rb, Ba) and low LILE/REE ratios (Fig. 17). Because experimental data predict the opposite for fluids (LILE enriched over HFSE and REE; e.g. Manning, 2004, and references therein), a fluid with a signature such as that indicated in Fig. 17 is unrealistic. At the relatively low maximum pressures (<1·5 GPa), a supercritical liquid is not expected either (e.g. Kessel et al., 2005).
Finally, a relatively hydrous metasomatizing agent would be consistent with formation of Mg-katophorites with low Ti and presumably low Fe3+ contents (given the relatively reducing conditions), which, in turn, imply high (OH)– contents. In other words, partial dehydrogenation, typically observed in common (Ti-rich) mantle amphibole (e.g. Popp et al., 1995; Tiepolo et al., 2001, and references therein), cannot have played an important role in this case, as H deficiency is mainly the result of coupled substitutions involving Ti and Fe3+ (Popp et al., 1995). In conclusion, the metasomatizing agent was most probably a Na2O–SiO2–CO2–H2O-rich melt.
Comparison with experimental studies further constrains the low-temperature regime in which metasomatism occurred. In ultramafic (e.g. Niida & Green, 1999), basaltic (e.g. Rapp, 1995) and Si-enriched systems (e.g. trachybasaltic; Barclay & Carmichael, 2004) amphibole (in the latter study together with phlogopite) is confined to temperatures lower than 1000–1050°C at pressures <2·0 GPa (Fig. 2). Thus, crystallization of the early metasomatic phase assemblage can be explained by infiltration of a silicate melt that rapidly cooled upon contact with the ambient peridotite (750–850°C prior to metasomatism; Fig. 2). However, this low-T stage is based on Ca-in-opx thermometry (see Brey & Köhler, 1990; Kaeser et al., 2006; Table 1) and therefore probably underestimates the effective temperature conditions during metasomatism. This can be explained by low diffusion rates and proximity to the closure temperature for the diopside–enstatite exchange reaction at temperatures lower than 850°C (e.g. Smith, 1999, and references therein). The lower temperature limit at which silicate melt metasomatism is possible (i.e. defined by the melt's solidus and viscosity) depends strongly on volatile and SiO2 contents. Experiments by Draper & Green (1997), for example, showed that highly silicic melts (52–64 wt% SiO2) can be in equilibrium with mantle harzburgite at temperatures as low as 850°C (at 1· 0–2·0 GPa; under anhydrous and C–O–H-fluid-bearing conditions). These are, however, extreme compositions for mantle melts. The actual SiO2 content of the metasomatizing agent in this case is difficult to constrain from the metasomatic mineral assemblage, as compositions may have varied considerably during reactive percolation (e.g. including Si-releasing reactions such as cpx replacement by amphibole versus Si-consuming reactions such as Opxm crystallization). Moreover, it is unlikely that a melt persists at temperatures much below 900°C, which is the solidus of fertile lherzolite (i.e. primitive upper mantle composition) under H2O-saturated conditions at 1·5 GPa (Grove et al., 2006; see Fig. 2). Below the solidus, residual volatile-enriched fluids may percolate further; for example, forming the less metasomatized Type III-c xenoliths (e.g. Ke 785*), as well as Type III-b and -a peridotites (see also Bailey, 1987). Relatively cold conditions would also explain the very high compositional gradients developed on small scales in Type III-a samples (Fig. 15). Given the absence of chlorite in all investigated samples, the lower temperature limit of such fluid metasomatism is 800°C (e.g. Fumagalli & Poli, 2005; see Fig. 2).
Finally, the presence of apatite, even though often taken as indicator for carbonatite-type metasomatism (e.g. Yaxley et al., 1998), is not in conflict with metasomatism by silicate melts. Apatite has also been shown to crystallize from volatile and alkali-rich derivatives formed during reactive porous flow of originally mafic silicate melts (Woodland et al., 1996; Bedini et al., 1997; Bodinier et al., 2004; Ionov et al., 2006).
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FORMATION OF THE LATE, GLASS-BEARING ASSEMBLAGES |
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Textures strongly suggest that the glass-bearing assemblages formed later and at the expense of the early metasomatic assemblages (Fig. 4a and d). Features such as the reappearance of clinopyroxene (Na-rich Cr diopside micro-phenocrysts), a phase replaced by the early metasomatic assemblage, indicate that melt pocket formation reflects a significant change in P–T–X space in parts of the mantle beneath Marsabit after earlier metasomatism. Based on geochemical arguments and mass-balance considerations, we will show that (1) the late, glass-bearing assemblages are closely related to early metasomatic assemblages, mainly representing breakdown products, and (2) breakdown was induced by infiltration of a CO2-rich fluid in the mantle, probably followed and enhanced by heating and decompression melting during xenolith transport in the host magma.
Evidence for in situ partial melting
BSE imaging reveals that the phases becoming unstable are amphibole, orthopyroxene and, to minor extents, spl-I, phlogopite, cpx-I and apatite. Compositional (inter- and intra-xenolith) heterogeneity of silicate glass is texture controlled and varies as a function of the associated precursor phases (Fig. 12a–f). TiO2, for example, shows higher concentrations in glass from xenolith Ke 1965/15 containing Ti-enriched amphibole (Fig. 12c), whereas K2O contents are highest in phlogopite-rich samples (e.g. Ke 785*; Fig. 12d). Elevated P2O5 contents are found in glass adjacent to apatite (Fig. 12e), which mirrors at the same time the high Th and LREE contents of apatite (Fig. 12f). Significant Si enrichment is found in glass coating orthopyroxene (Fig. 12d), supporting incongruent melting of orthopyroxene (see below). Even though the melt pockets are connected along grain boundaries, the strong compositional variation (on inter- and intra-xenolith scale) suggests that the mobility of the melt was limited, or that the melting event was very short-lived (for timing constraints see below).
The co-variation plots in Fig. 12 further reveal considerable compositional differences between all types of xenolith glass and the host basanite. These data indicate that the glass cannot be the product of xenolith–host lava interaction during ascent to the surface. Also, back-mixing of micro-phenocrysts into the melt fails to derive the glass from a host-related melt (Fig. 12b). Glass from xenolith margins, however, is compositionally closer to the bulk host basanite (e.g. Fig. 12b), in line with a direct derivation from xenolith–host lava interaction. Its more evolved major element composition (e.g. higher SiO2, lower Mg-number; Fig. 12b) is probably the result of fractionation of basanite phenocrysts (e.g. olivine, augite and titanomagnetite; Volker, 1990). Host lava–xenolith interaction is further supported by the trace element signature of glass at the margin of the xenoliths, which is closely similar to that of the Marsabit basanites (Fig. 13; Volker, 1990). The same characteristics are observed for intra-xenolith glass in the unusual kaersutite-bearing sample Ke 1959/27.
Xenolith glass unrelated to the host lava and melt pockets similar to those reported here, in association with residual phases (e.g. amphibole or clinopyroxene), has been reported from many xenolith localities (e.g. Frey & Green, 1974; Ionov et al., 1994, 1996; Chazot et al., 1996a, 1996b; Yaxley et al., 1997; Laurora et al., 2001; Bali et al., 2002). Residual phases within the melt pockets strongly indicate in situ partial melting within the xenoliths. However, a variety of processes may trigger such in situ melting, including: (1) closed-system decompression melting of either anhydrous (Maaløe & Printzlau, 1979) or hydrous phases (Frey & Green, 1974; Francis, 1976) during xenolith ascent in the host lava; (2) melting during interaction of the xenoliths with the host basaltic melt during xenolith ascent (e.g. Zinngrebe & Foley, 1995; Shaw & Klügel, 2002; Shaw et al., 2006); (3) melting in the mantle prior to xenolith entrainment, induced by an infiltrating fluid (i.e. flux melting; e.g. Jones et al., 1983; Ionov et al., 1994; Demény et al., 2004). However, most studies agree that that partial melting occurred very shortly before and/or during xenolith exhumation by the host magma. Thus a combination of metasomatically triggered melting followed and/or assisted by decompression melting during xenolith transport may be a realistic alternative in some cases (e.g. Dawson, 2002).
Recently, Ban et al. (2005) have shown, in a very detailed mass-balance study on glass-bearing xenoliths from the West Eifel, that the weight proportions of glass and the associated micro-phenocrysts match the composition of the preserved parent amphibole. We used the same approach and performed simple mass-balance calculations (least-squares fitting), to constrain the conclusions drawn from petrography, and to test which phase contributed, and to what extent, to the formation of the melt pockets. The results, reported in Table 8, show that in most cases amphibole-melting can account successfully for the melt pocket composition, and only a minor amount (< 20%) of other phases (orthopyroxene and/or phlogopite) needs to be added to the reactant. Such a contribution of other phases agrees with textural observation (e.g. Fig. 4c showing Dim growing on resorbed Opxm), as well as with the conclusions of Ban et al. (2005), who stated that more Si-rich glass (i.e. >54 wt%) necessitates orthopyroxene contributing to the melt.
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Trace element patterns of the bulk glass + micro-phenocryst pockets, calculated using the measured trace element abundances of glass and micro-phenocrysts and the respective mass-balance parameters (Table 8), are broadly identical to the measured amphibole composition (Fig. 18a and b). Higher LILE (Cs, Rb, Ba) concentrations in the calculated bulk melt patches in sample Ke 785* can be balanced by adding
7 wt% of phlogopite (Fig. 18b), which is in line with mass balance (Table 8). The dominant control of amphibole to account for the trace element budget of the melt pockets is similar to results found in other suites of glass-bearing xenoliths (Chazot et al., 1996a; Yaxley & Kamenetsky, 1999; Laurora et al., 2001). Further, the obtained phase proportions for the melt pockets are in good agreement with the results from image analysis (Table 1). In addition, they are similar to the glass–cpx–ol proportions recorded in recent experiments on the melting behaviour of amphibole in wehrlites at temperatures between 1175 and 1250°C and pressures of 0·5–1· 0 GPa (Médard et al., 2006). These temperatures are in good agreement with the estimates for the formation of olivine and chromite micro-phenocrysts in the melt pockets (Fig. 16). The heat supply required for amphibole melting in the formerly low-T metasomatized Group III xenoliths (
T 200–300°C) may come from local heating shortly before or during xenolith entrainment in the host lava or from fluid infiltration that triggered melting (see discussion below).
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Finally, formation of the most Si-enriched melt (adjacent to Opxm in Group III-c samples) can be modelled assuming progressing interaction between infiltrating melt and orthopyroxene where the latter melts incongruently:
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| (1) |
).
Partial melting triggered by fluid flux?
Mass-balance calculations provide an explanation for melt pocket formation by in situ partial melting. Thus, one might conclude that melting is simply the result of decompression and heating during xenolith transport in the host magma. However, there are two points that argue that melting was triggered by the infiltration of a fluid phase shortly before the xenoliths were sampled by the host magma: (1) the presence of occasionally high amounts of carbonate included in the melt pockets; (2) the extremely Na-rich clinopyroxene (Dim) compositions, either included as micro-phenocrysts in the melt pocket or present as equilibrated grains around chromite in the single Type III-d xenolith (Fig. 3f). Both arguments will be discussed in the next section.
Formation of carbonates
Textures and compositions of carbonate globules can be explained by carbonate precipitation from a CO2-bearing fluid. As demonstrated by several studies, high Ca/(Ca + Mg) values (Table 7) and low REE concentrations are strong arguments that the carbonates are cumulate crystals, precipitated from a CO2-rich melt or a fluid and not quenched carbonated melts (Lee et al., 2000; Ionov & Harmer, 2002). Oscillatory zoning (Fig. 5a and b) of individual carbonate crystals within single globules corroborates this interpretation. The rounded contacts toward glass (globules) are also in line with these findings. Similar globular forms were reported from experimental studies, and do not a priori imply the presence of immiscible carbonate and silicate melts (e.g. Lee & Wyllie, 1996).
The high amount of carbonate in some melt pockets, however, (e.g. Fig. 4e) does need further explanation, as the incongruent melting reaction (1) does not account for the CO2 required for carbonate precipitation during amphibole melting. Therefore, amphibole breakdown either occurred in an open system with addition of at least CO2 (e.g. Laurora et al., 2001; Ban et al., 2005), or another pre-existing phase acted as a source of carbon. Oxidation of graphite, texturally identified as a part of the early assemblage (e.g. Fig. 3d), could account for some of the CO2 necessary for carbonate formation. A change in fO2 conditions is also indicated when the composition of glass-related spinel (Chrm) and Olm are used for oxygen barometry, which yields oxygen fugacities well on the oxidized side of the carbon–carbonate–silicate buffer (Fig. 16). However, increasing fO2 triggered by closed-system decompression-induced melting alone does not appear feasible. Also, textures, such as straight contacts between graphite and Mg-calcite (Fig. 4c), suggest that carbonate grows on graphite (i.e. indicating local fO2 conditions corresponding to carbon–carbonate equilibrium) rather than replacing it. We therefore interpret melt pocket formation to be triggered by the infiltration of a CO2-rich fluid (i.e. flux melting; see also Ionov et al., 1994). This interpretation also accounts for the presence of abundant vugs in the glass that could have contained a fluid (Ionov et al., 1994, 2006), which was presumably lost during sample preparation. However, we emphasize here that late fluid infiltration had only subtle effects on the trace element composition of Type III-b and -c xenoliths.
Formation of Na-rich Cr diopside
Infiltration of CO2-rich fluid would also explain the Na-rich composition of Dim micro-phenocryst cores (Fig. 6a and b). Ionov et al. (2005) noted recently that clinopyroxenes in melt pockets are commonly characterized by Na2O contents <1 wt%, mainly as a result of low cpx/melt partition coefficients at low pressures (e.g. Blundy et al., 1995). Thus, very high Na2O contents in Dim micro-phenocrysts (Fig. 6a) suggest replacement of amphibole by Dim and eventually Olm and Chrm during reaction with a fluid at subsolidus conditions. An origin of micro-phenocrysts by replacement rather than crystallization from a melt is indicated by the identical crystallographic orientation of individual Dim crystals in single melt pockets (Fig. 4a).
Once the solidus of the precursor phase(s) was reached (i.e. 1050°C for amphibole; Niida & Green, 1999) silicate melt was generated in situ and led to the formation of clinopyroxene rims with lower Na contents (Fig. 6b) and the development of euhedral shapes in contact with melt (e.g. Fig. 4a and e). Eventual immiscibility between remaining fluid and newly formed melt (e.g. during xenolith ascent) would then explain the presence of carbonate globules and vugs enclosed in glass. Likewise, different origins of melt and carbonate were recently proposed for similar melt pockets from Hungarian xenoliths by Demény et al. (2004), who found oxygen isotope disequilibrium between carbonate globules and the host glass. Such a two-step model (fluid-induced replacement followed by melt generation) can be explained by CO2-fluid infiltration occurring within the mantle at sub-solidus conditions, ultimately followed by xenolith entrainment in the host magma and transport to the surface. The host magma possibly supplied the additional heat necessary for melt generation. Fast decompression would further enhance melting and could also account for the lower Na contents in Dim micro-phenocryst rims as a result of decreasing cpx/melt partition coefficients at decreasing pressures (Blundy et al., 1995).
Rapid replacement of amphibole (± phlogopite) by an anhydrous phase assemblage indicates that the fluid was water-poor (e.g. low H2O/CO2 ratio; see Ionov et al., 2006). This is corroborated by the fact that Mg-katophorite (presumably OH-rich; see above) is occasionally rimmed by a residual kaersutitic selvage (Fig. 4d). Such Ti-rich mantle amphiboles are commonly OH-poor as a result of partial dehydrogenation (e.g. Tiepolo et al., 2001, and references therein).
Finally, the presence of the Type III-d xenolith, containing texturally equilibrated clinopyroxene of almost identical composition to Dim micro-phenocrysts in melt pockets (Table 2; Figs 6a and 7b), further strengthens the argument that CO2-fluid percolation started in the mantle before entrainment (i.e. enough time elapsed to allow this rock to anneal completely after the late metasomatic assemblage had formed). A similar conclusion can be drawn for the clearly host-related glass found in xenolith Ke 1959/27 (e.g. Fig. 13b). This Type III-c sample is characterized by the presence of Dim micro-phenocrysts partly overgrown by mineral rims clearly related to the host basanite magma (kaersutite, Ti-augite and olivine with Fo88·5; Fig. 4f; Tables 3 and 4, and supplementary data at http://petrology.oxfordjournals.org). Kaersutite has major and trace element characteristics similar to typical basanitic and alkali basaltic amphibole phenocrysts (Fig. 9d). The special phase assemblages in this xenolith indicates that host-related basanitic melts intruded some Type III-c xenoliths and partly overprinted the assemblage formed slightly earlier by in situ derived alkali-rich peraluminous melts, fully preserved in other samples (e.g. Ke 1965/3 or Ke 785*).
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ORIGIN OF METASOMATIZING AGENTS: TIMING AND TECTONIC SETTING |
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 TOPABSTRACTINTRODUCTIONSAMPLE CONTEXTANALYTICAL METHODSPETROGRAPHYMINERAL COMPOSITIONFORMATION OF THE EARLY...FORMATION OF THE LATE,... ORIGIN OF METASOMATIZING AGENTS:... SUPPLEMENTARY DATAREFERENCES |
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As noted above, late metasomatism was probably short-lived, started at mantle depths and probably continued during xenolith transport in the host magma. The relation between fluid infiltration and xenolith entrainment suggests that the source of the fluid is intimately related to the host magma. Volatile components such as CO2 play a major role during wall-rock disruption and xenolith sampling at depth (e.g. Spera, 1984
; Lensky et al., 2006). Thus the CO2-rich fluids may be derived directly from the rising basanitic magma that ultimately sampled the xenoliths (Fig. 19). Alternatively, the CO2-fluid may be derived from decarbonation reactions (CO2 release) at greater depths where rising carbonate-rich melts reach the solidus of carbonated lherzolite (see Fig. 2 and Lee & Wyllie, 2000, and references therein). Finally, we would like to stress that silicate melt (now glass) was formed in situ, within the xenoliths and, thus, does not represent an extraneously derived metasomatizing agent.
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In contrast to the melt pockets, the age of formation of the early metasomatic assemblages is more difficult to assess. Relative age information can be obtained from textural relationships. Metasomatism partly overprints the strongly deformed fabric of the pre-metasomatic Group III rocks. Because pervasive deformation probably occurred during the development of the Mesozoic Anza Graben (Kaeser et al., 2006
; Fig. 1), metasomatism should be younger. In the geodynamic context of Marsabit, this would then most probably be in association with the young (Miocene–Pleistocene) magmatic activity of the Kenya Rift (EARS).
The occurrence of orthopyroxene-saturated volatile-rich melts in a region characterized by voluminous alkaline magmatism is enigmatic, however. It seems most plausible that such melts evolved from alkaline mafic melts through melt–rock reaction, as in the models of Bedini et al. (1997) or Bodinier et al. (2004). In their study, Bedini et al. (1997) explained metasomatism in the enriched and deformed harzburgite xenoliths from southern Ethiopia (200 km north of Marsabit) by complex reactive porous flow processes, starting with an ‘ordinary’ ocean island basalt type mafic silicate melt. Their hypothetical melts in equilibrium with the metasomatized xenoliths have trace element compositions strikingly similar to the calculated silicate melts in equilibrium with euhedral Mg-katophorite (Fig. 17). In Marsabit, the heated (1000–1200°C) and recrystallized Group II lherzolites provide evidence for recent, EARS-related metasomatism through alkaline mafic melts, possibly parental to the basanites extruded at the surface (Kaeser et al., 2006). Similar to the process invoked for the Ethiopian xenoliths, continuous melt–rock reaction and fractionation could have led to the formation of small fractions of highly evolved melts infiltrating and metasomatizing the cold and deformed uppermost mantle represented by Group III harzburgites (Fig. 19; see also Ionov et al., 2006, for a similar model). A relationship between metasomatism in Group II and III peridotites is supported by several xenoliths showing transitional textures and trace element signatures (i.e. strong U–Th–LREE enrichment in some Group II clinopyroxenes; Kaeser et al., 2006).
Alternatively, Si- and volatile-enriched melts may be residual liquids from the crystallization of alkali basaltic or basanitic veins (i.e. hornblendite or clinopyroxenite), ultimately metasomatizing the adjacent peridotitic wall-rock (e.g. Harte et al., 1993; Wulff-Pedersen et al., 1999; Bodinier et al., 2004). This remains speculative, however, as such veins or dykes have not yet been reported from Marsabit. It can thus not be resolved whether early metasomatism reflects a large, kilometre-scale process (e.g. an upwelling melt percolation front as proposed for the mantle beneath southern Ethiopia; Bedini et al., 1997), or whether metasomatism is restricted to narrower zones adjacent to small (centimetre- to metre-scale) melt conduits (e.g. wall-rock metasomatism).
Extensive melt–rock reaction in the mantle beneath the EARS, however, is indicated by numerous xenolith suites containing abundant and variably metasomatized rocks (e.g. Tanzania: Dawson & Smith, 1988; Rudnick et al., 1993; Lee & Rudnick, 1999; Dawson, 2002, and references therein; Ethiopia: Bedini et al., 1997; Rooney et al., 2005). In the case of northern Kenya, seismic studies point to a broad zone of low seismic wave velocities in the relatively shallow mantle 600 km across and reaching depths of <40 km depth beneath Lake Turkana and 90 km beneath Marsabit (see Simiyu & Keller, 1997). This anomaly is commonly interpreted as a zone of hot rising mantle containing small amounts (3–5%) of melt (Green et al., 1991). Upwelling of such melts into the lithospheric mantle followed by reaction could thus account for the metasomatism seen in the Marsabit mantle xenoliths. Such interaction of sub-lithospheric melts or fluids (asthenospheric or plume-related) with colder lithospheric material is one of the major processes invoked to account for the heterogeneous isotopic composition of primitive lavas from numerous parts of the EARS (e.g. Macdonald et al., 2001; Furman et al., 2006, and references therein).
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SUPPLEMENTARY DATA |
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Supplementary data for this paper are available at Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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We would like to thank Edwin Gnos and Alfons Berger (Bern), and Hans-Peter Meyer (Heidelberg) for assistance with microprobe work, as well as André Villard, who provided top-quality thin sections. Discussion with Othmar Müntener, Arjan Dijkstra, Rainer Altherr, Jean-Louis Bodinier and Alan Woodland helped to clarify many aspects of this work. We thank Hilary Downes, Tanya Furman and Dmitri Ionov for constructive reviews, and Marjorie Wilson for valuable comments and editorial handling. This publication is part of a project financed by the Swiss National Science Foundation (grant No. 200021-100647/1 to A.K.). Further financial support from the SNF (grants 21-26579.89 and 200021-103479/1) for the electron microprobe facility at the Institute of Geological Sciences, University of Bern, is acknowledged.
*Corresponding author. Present address: School of GeoSciences, University of Edinburgh, West Mains Road, Edinburgh EH9 3JW, UK. Telephone: 0044 (0) 131 650 7339. Fax: 0044 (0) 131 668 3184. E-mail: v1bkaser{at}staffmail.ed.ac.uk
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