fluid processes in diamond to spinel facies shallow ... - Science Direct

J. GeodynamicsVol. 20, No. 4, pp. 387-415,1995 Else&r ScienceLtd. Printed in Great Britain 0264-3707/9S $9.50+O.oO 0264-3707(95)00018-6

FLUID PROCESSES IN DIAMOND TO SPINEL FACIES SHALLOW MANTLE

MARTIN

MENZIES

and GILLES CHAZOT

Department of Geology, Royal Holloway University of London, Egham Hill, Egham, Surrey TW20 OEX, U.K. (Received for publication 29 December

1994; accepted in revised form I1 April 1995

Abstract-Topography on the lithosphere-asthenosphere boundary and resultant variations in the architecture of the lithosphere are inextricably linked to heat and mass transfer processes involving the convecting upper mantle. This is primarily accomplished by the movement of silicate and nonsilicate (i.e. carbonate) melts, derivatives of which are trapped in the shallow mantle beneath tectonically active and inactive parts of continents. In diamond facies mantle (180-220 km), melt character is inferred from solid inclusions (e.g. apatite, carbonate, mica) in megacrystic diamonds and the mineralogy and chemistry of diamondiferous peridotites and eclogites. Deep lithospheric processes have involved silicate (K-rich and hydrous) and carbonate melts. In garnet facies mantle (75-200 km) information on melt transfer processes occurs in kimberlite-borne xenoliths, e.g. phlogopite-richterite garnet peridotite, MARID (mica-amphibole-rutile-ilmenite-diopside) pyroxenites and IRPS (ilmenite-rutile-phlogopite-sulphide) pyroxenites. Polybaric fractionation and crystallization of silicate melts in propagating fractures in the lower lithosphere can explain the character of several xenolith suites. Transfer of potassic silicate melts (equiv. kimberlite/ lamproite) are thought to be closely linked to the genesis of MARIDs and hydrous garnet peridotites, and, the transfer of alkaline silicate melts (equiv. basalt) may explain the character of IRPS pyroxenites. In spinelfucies mantle ( < 75 km) kaersutite-pargasite-carbonate spine1 peridotites and amphibole-mica-apatite pyroxenites constrain the nature and origin of shallow mantle fluid processes. While transfer of mobile silicate (equiv. basalt) melts accounts for the chemistry of many spine1 peridotites and pyroxenites, highly mobile carbonate melts are believed to have played a pivotal role in the formation of apatite pyroxenites/wehrlites (converted peridotites) and carbonate-bearing peridotites (reacted wallrock).

INTRODUCTION

AND DEFINITIONS

Melt transfer processes (Frey and Green, 1974) were invoked some twenty years ago in order to account for the petrology and geochemistry of spine1 facies mantle peridotites. These silicate melts originated in the asthenosphere (1300°C) and, because of their natural buoyancy, impregnated the base of the lithosphere which had formed from melt extraction processes some time earlier. Moreover it was suggested that deep seated fractures, in this cold brittle lithosphere, had

@ Crown Copyright (1995). 387

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Mcnz~ea and G. Chazot

acted as conduits for the transfer of silicate melts (Ehrenberg, 1979; Irving, 1980; Harte, 1983) and that polybaric fractionation of small volume melts had modified the composition of the melt and migrating small volume melts had modified the adjacent wall-rock. Deep-seated equivalents of basalt, kimberlite or carbonatite (Wass et al., 1980; Menzies, 1983; Roden et al., 1990) were shown to have been primarily responsible for the effects observed in basalt-borne and kimberlite-borne xenoliths (Nixon, 1995 this issue). In this paper we will (a) define appropriate terms; (b) review data from different mantle lithologies (diamond, garnet and spine1 facies); (c) consider the metasomatic mineralogy of the mantle rocks, (d) define the type of melt (i.e. silicate or non-silicate) and its affinity to erupted equivalents (e.g. kimberlite, basalt, carbonatite), and (e) comment on the possible provenance of these melts.

Lithospherelasthenosphere In understanding the effects of fluid processes in the shallow mantle it is necessary to explain the differences between asthenosphere and lithosphere (White, 1988). The lithosphere comprises a mechanical boundary layer (MBL) (crust and mantle) that was produced by melt extraction from the mantle (Fig. 1). The continental lithosphere can reach thicknesses > 150 km beneath tectonically inactive cratons of Archaean age (e.g. Kaapvaal) or tectonically active continental regions of Phanerozoic age (e.g. Alpine front, Europe). In contrast, continental lithosphere can be thinned to e70 km in regions of extension (e.g. western USA; eastern China). According to White (1988) the mechanical boundary layer is separated from the underlying asthenosphere by a thermal boundary layer, which acts as a buffer zone and, as such, has some of the thermo-mechanical and chemical properties of both the MBL and the asthenosphere. In contrast to the lithosphere, which tends to be cold, the asthenosphere (> 1300°C) and hot spots (> 1400°C) are a continual source of high temperature melts that, because of their natural buoyancy, rise and permeate the base of the overlying lithosphere. Information regarding the character of the mechanical boundary layer has emerged from studies of erogenic massifs (spine1 facies), basalt-borne xenoliths (spine1 to garnet facies) and kimberlite-borne xenoliths (spinel, garnet and diamond facies). Partial melting of the lithosphere (Fig. 1) can be brought about by an influx of volatiles, an increase in temperature, or a loss of pressure. Mantle metasomatic processes can be associated with volatile input and an increase in temperature. Kimberlite-borne xenoliths and solid inclusions in diamonds (Fig. 2) from tectonically inactive parts of the world have P-T characteristics close to the continental/shield (i.e. cold) geotherm. Moreover these mantle samples mainly provide information on the mantle from >20 kb. Garnet peridotite xenoliths from the Kaapvaal craton and Lesotho lie along the shield geotherm with an inflexion towards higher temperatures at depth (150-180 km) which has been interpreted as the base of the lithosphere (Finnerty and Boyd, 1987 for review).

Fluid processes in diamond Temperature 1000

“C 1500

lO(

g _c E $

20(

3oc

Fig. 1. Thermal character of the lithosphere and the location of the wet and dry solidus (Wyllie, 1995 this issue). Information on the nature of the mechanical boundary layer (MBL) overlying the thermal boundary layer (TBL), is found in erogenic massifs (plagioclase to diamond facies), basalt-borne (plagioclase to garnet facies) and kimberlite-borne (spine1 to diamond facies) xenoliths. Melting of the mantle (small figures at base) may occur due to a change in solidus caused by increased volatile content (metasomatism), an increase in temperature (plume) or a decrease in pressure (extension). The path of partially melted mantle diapirs intersects the wet solidus and can thus account for metasomatism in the shallow mantle (e.g. carbonate metasomatism).

M. Men&s

and Cr. Chazot

Temperature 1000

60

(kb)

(km)

6C 80

Fig. 2. Pressure-temperature relationships in kimberlite-borne xenoliths and solid inclusions (P-type) in diamonds from the south African craton (after Finnerty and Boyd. 1987; Boyd ef al., 1985). Note that most kimberlite xenoliths are from garnet and diamond facies mantle and that the P-T estimates coincide with the shield (cold) geotherm entirely consistent with the presence of a thick, high seismic velocity lid and low heat Row (see also Haggerty. 1995 this issue: Nixon. 1995 this issue).

These garnet facies xenoliths straddle the graphite-diamond transition, a factor that is consistent with the occurrence of diamondiferous kimberlites in South Africa. Solid P-type (i.e. peridotite paragenesis) inclusions in diamonds have P-T characteristics consistent with entrapment in diamond facies mantle. In contrast, basalt-borne spine1 and garnet facies xenoliths (Fig. 3) from tectonically active parts of the world have P-T characteristics close to the oceanic/ridge (i.e. hot) geotherm with samples being representative of shallower mantle. The contrast is shown in Fig. 3 between adjacent tectonically and volcanically active areas like Kilbourne Hole close to the Rio Grande Rift, New Mexico and Lunar Crater in the Great Basin, Nevada where xenoliths are “hot”, and tectonically

Fluid processes in diamond

391

Temperature 1040

60

(km)

(kb)

120

60

180

q

THIN LITHOSPHERE THICK LITHOSPHERE

G, :

vn: 9: z: m: 33: zz::

:

80 Fig. 3. Pressure-temperature relationships in basalt-borne and kimberlite-borne xenoliths from the western U.S.A. (after Finnerty and Boyd, 1987; Boyd et al., 1985). Basalt-borne and kimberlite-borne xenoliths from beneath areas of thick lithosphere (low heat flow) tend to lie closer to the continental/shield (cold) geotherm than basalt-borne xenoliths entrained from areas of thin lithosphere (high heat flow) that lie on the oceanic/ ridge (hot) geotherm.

and volcanically inactive areas like the Thumb on the Colorado Plateau where some xenoliths are “cold”. Xenoliths from the State Line and Williams indicate the existence of thicker, diamondiferous lithosphere (15&180 km). A comparison of Figs 2 and 3 illustrates the marked differences between a tectonically inactive on-craton area (Fig. 2) where the majority of mantle xenoliths adhere to the shield geotherm and the lithosphere has attained thicknesses of 180-220 km for billions of years (Boyd et al., 1985) with a tectonically active off-craton area (Fig. 3) where the majority of mantle xenoliths from an area of thin lithosphere are proximal to the oceanic geotherm.

M. Menzies

392

and G. Chazot

ON-CRATON , (COLD)

I

I

I

I

I

90

60

70

60

50

Modalolivinewt % Fig. 4, Mg number in olivine versus modal proportion of olivinc in the shallow mantle protolith beneath cratons (on-craton), circum-cratonic (off-craton) regions and ocean basins (arrowed) (Boyd, 1989). The arrowed curve (“oceanic”) is constrained by ophiolitic and erogenic peridotites and depicts the trend interpreted as associated with the removal of tholeiitic melts from shallow mantle during formation of oceanic lithosphere (< 200 Ma). Off-craton (hot) shallow mantle defines a similar trend and may have evolved in a similar manner due to recent accretion of asthenosphere. However, on-craton (cold) lithosphere lies off this “oceanic trend” and is believed to have formed as a result of higher temperature processes that produced highly magnesian melts (i.e. komatiites) and highly magnesian residue (Boyd. 1989).

Protolith Any understanding of mantle metasomatic processes requires knowledge of the precursory mantle assemblage. The vast amount of data available in the literature leaves no doubt that the protolith is peridotitic and has a refractory character consistent with a pre-history of partial melting processes (Haggerty, 1995 this issue; Nixon, 1995 this issue). This can be adequately illustrated using modal proportions and compositional data for olivine (Boyd, 1989). In Fig. 4 the available data for off-craton and on-craton peridotites are compared with a trend for residual peridotites (i.e. ophiolitic and erogenic) produced by removal of tholeiitic melts in the production of oceanic lithosphere. Peridotites that are common as xenoliths in on-craton kimberlite pipes are characterized by a proximity to the shield (cold) geotherm and a highly magnesian composition. These contrast with the less magnesian character and proximity to the oceanic


(hot) geotherm of peridotites possible interpretation of this were produced by removal of left highly magnesian residual of the cause of these possible peridotite.

393

found in off-craton basaltic host magmas. One difference is that on-craton (Archaean) peridotites higher temperature (more magnesian) melts which peridotites. The important point is that, regardless differences, in all cases the protolith is a refractory

C-O-H repositories Experimental work has helped determine the speciation of volatile components in the upper mantle (lithosphere-asthenosphere), possible mineral repositories for these volatiles and the stability field of these minerals (Wyllie, 1995 this issue). In the upper mantle, HZ0 is mainly stored in amphibole and phlogopite, whereas CO* is contained in carbonates varying from dolomite (low pressure) to magnesite (high pressure). The presence of carbon, oxygen and hydrogen as H20, CO, and CH4 in the mantle is important as the formation of amphibole, mica and carbonate can ultimately lower the peridotite solidus (Fig. 1). The existence of hydrous melts at depth (>>2OOkm) is important as they can act as a trigger for melting of thick cratonic lithosphere (Carlson, 1995 this issue) (Fig. 1). It is apparent from recent experimental work that, at depth, any free fluid would be water-dominated, because of the stability of carbonate which acts as an effective sink for CO, (Wyllie, 1995 this issue). However, this hydrous fluid would only exist if enough potassium were present to stabilize the hydrous phase mica. Thompson (1992) reviewed the presence of water in the mantle, and the possible repositories for water in the upper mantle, and noted that dense hydrated magnesium silicates (i.e. DHMS) may act as a repository for water during recycling to depths >500 km. However, other phases are needed to account for water in the peridotitic mantle and it has been suggested that dehydration melting of richterite (K amphibole), at depth, may account for the presence of potassic, volatile-rich inclusions in diamonds. Where mantle temperatures are elevated, water may reside in melts rather than be locked up in minerals, and at around 300 km, a “water storage line” defines a level where water is dissolved in melts or in minerals. At shallow depths the free fluid would tend to be more CO2 rich due to the lack of stable carbonate and the presence of excess CO2 in silicate melts (Eggler, 1987; Wyllie, 1987; 1989). Wallace and Green (1988) showed that at 21 kb a carbonatitic melt intersected its solidus and reacted with the peridotitic matrix in decarbonation reactions with the release of CO* vapour (Fig. 1). The stability field of a sodic dolomitic carbonatitic melt coexisting with an amphibole lherzolite has been shown to lie between 21-30 kb (63-90 km) within the spine1 and garnet facies mantle (Wallace and Green, 1988). In these reactions, the carbonatitic melt reacts with enstatite or spine1 to form clinopyroxene, olivine and sometimes pargasite. These reactions modify the modal composition of the

304

M. Menzies

and CT. Chazot

rocks, from lherzolite or harzburgite towards wehrlite. The replacement of orthopyroxene by clinopyroxene is accompanied by the formation of accessory apatite. Accordingly, the chemistry of such mantle rocks would be modified, and the experiments predict an increase in Ca/Al and Na/Ca ratios (Green and Wallace, 1988). The experiments of Dalton and Wood (1993) indicated an increase in the Ca content of olivines and the Na content of clinopyroxenes due to the involvement of calcic or more sodic carbonatites. An interesting point about these experiments is that the Mg# (lOO*Mg/(Mg + Fe)) of the carbonatitic melts in equilibrium with a mantle peridotite is not very different from the Mg# of the solid silicates (Green and Wallace, 1988; Dalton and Wood, 1993). Since this contrasts with the low Mg# of a silicate melt it may be used to distinguish between carbonate-melt or silicate-melt metasomatism in the mantle. Whether carbonatites can be generated at deeper levels within the upper mantle is still unclear as much of the recent experimental data point to the presence of carbonatitic melts at ~7.5 km (Watson et al., 1990 for review). Wyllie (1989) noted that at >7.5 km silicate melts containing dissolved CO, could, upon crystallization, have produced carbonate rich silicate melt (i.e. carbonatitic fluid), or hydrous melts if sufficient potassium exists over and above that needed to stabilize mica. Carbonatites could be formed by partial melting of carbonated mantle or by cxsolution from undersaturated CO!-rich silicate melts like phonolite or nephelinite (Wyllie ef ul., 1990). Schrauder and Navon (1994) extrapolated the experimental data of Falloon and Green (1989) and suggested that carbonatitic fluids could exist in the diamond stability field (180-200 km). The data of Hauri et al. (1993) also point to mechanisms that may allow for CO? and other volatiles to survive shallow subduction, only to be recycled deep into the mantle. Clearly more experimental data are required before the full range of conditions (i.e. pressure and temperature) is defined for the genesis of carbonatites.

Wetting angles Any evaluation of the probability that a particular fluid is an appropriate metasomatic agent hinges on the olivine-olivine (i.e. peridotite) or the pyroxene-pyroxene (i.e. pyroxenite) dihedral or wetting angle (Fig. 5). Several authors (Hunter and McKenzie. 1989; Watson et al., 1990) have demonstrated that silicate or carbonate melts are the most mobile fluids due to their ability to form an interconnecting network in an olivine matrix. In contrast, CO, rich fluids are essentially immobile because of their inability to form an interconnected network (Fig. 5). With regard to hydrous fluids, Watson and Lupulescu (1993) showed that such fluids played little or no role in “larger than grain scale” chemical transport in clinopyroxene rich mantle rocks, even when NaCl, CaCl, and CO1 are added. Clearly because aqueous and CO, rich fluids form isolated pores, mantle metasomatism (sensu stricto) most likely involves silicate and carbonate melts due to their ability to form interconnecting pore networks.


395

1.2RICH FLUID

.-

t

I

20

40

I 60

80

8 Dihedral/Wetting angle Fig. 5. Olivine-olivine dihedral angles in carbonate melts, basaltic melts, water rich fluids and carbon dioxide rich fluids (after Watson et al., 1990). These data indicate that carbonate melts are mobile within a peridotite matrix due to a high level of melt connectivity along grain boundaries (top left figure) and are therefore the most effective metasomatic agents. In contrast, CO, rich fluids are immobile within a peridotite matrix due to a lack of melt connectivity along grain boundaries (top right figure) and are therefore totally ineffectual as metasomatic agents (Wyllie, 1995 this issue). Note (a) that these data are not relevant when fluids move by crack propagation through the lithosphere, and (b) that the effect of deformation on fluid mobility may be important.

Stable isotopes and fluid processes

Traditionally, oxygen isotopes have been used to identify the nature and composition of mantle fluids. Fractionation of oxygen isotopes between mantle minerals (i.e. olivine, orthopyroxene and clinopyroxene) was interpreted in terms of temperature (Kyser et al., 1981, 1982) and was used as a geothermometer. In contrast Gregory and Taylor (1986a,b) believed in an open system with exchange between peridotite and oxygen-bearing fluids. In their model, the olivine and spine1 exchanged their oxygen at a faster rate than the pyroxenes and, as a consequence, showed larger variations in oxygen isotopic composition. Kempton et al. (1989) adopted this model to explain the oxygen isotope geochemistry of spine1 lherzolites from Germany, where olivine that co-existed

396

M. Menziesand G. Chazot

with amphibole and mica, exhibited large variations in (5;“O. They suggested that the mantle under the Eifel had experienced several metasomatic episodes leading to LREE enrichment and to important changes in the primary mantle oxygen isotopic composition. However, in this open system model, it is difficult to understand how a fluid with different 6’*0 composition may circulate in the mantle without being affected/buffered by the 6180 composition of ohvine which represents in excess of 80% of the mantle (Mattey and Macpherson, 1993). Recent advances in oxygen isotope analysis by laser-fluorination techniques (Sharp, 1990; Mattey and Macpherson, 1993) have allowed for very precise measurement of c?“O in mantle minerals - especially refractory olivine. Analyses of olivine in hydrous and anhydrous spine1 and garnet lherzolites (Mattey et al., 1994) reveal that the 6”O ratio in mantle ohvine is not so variable as previously reported. This observation is confirmed in a recent study of spine1 lherzolites from Ataq, Yemen and Nunivak, Alaska (Chazot et al., 1995a,b), which reveals no 6”O variation in co-existing olivine or pyroxene whether in u hydrous (wet) or an anhydrous (dry) spine1 lherzolite. Because of this and the fact that hydrous minerals like amphibole and mica have a homogeneous isotopic composition it is unlikely that, in the majority of cases, oxygen isotope data can be used to track fluid movement in the upper mantle (Mattey and Macpherson, 1993; Mattey et al., 1994; Chazot et al., 1995a,b).

TECTONICALLY

INACTIVE

CONTINENT

Diamond-fiicies peridotites (180-220 km) Diamond megacrysts provide valuable information about the deep lithosphere (Nixon, 1995 this issue) because of their ability to trap high pressure primary phase assemblages (e.g. garnet, clinopyroxene, olivine) and/or secondary phase assemblages (e.g. carbonate, mica, apatite and quartz) (Richardson et al., 1984; Navon et al., 1988; Guthrie et al., 1991; Schrauder and Navon, 1994). Navon et al. (1988) analysed inclusions in megacrystic diamonds, from Zaire and Botswana, and noted that they were rich in H20, CO,, SiO,, K20, CaO and FeO. Infrared spectra pointed to the presence of hydrated sheet silicates (phlogopite?), carbonates and phosphates (apatite?) and some molecular CO, as overgrowths on the cores of coated diamonds (Navon et al., 1988; Schrauder and Navon, 1994). In a further study of fibrous diamonds from Botswana, Schrauder and Navon (1994) reported the presence of water, carbonates, silicates, apatite, chlorine and CO,. Although they were unable to positively identify dolomiteankerite they concluded that the bulk of the CO, resided in carbonate at these pressures. Several inclusions had element ratios compatible with the presence of Fe-, Mg- and C&bearing carbonates, mica (phlogopite-biotite), rutile and enstatite, although the latter phase may have existed due to reaction. The phases reported by Schrauder and Navon (1994) are extremely rich in potassium with an extrapolated temperature and pressure of formation around 1000°C and 4.5


397

GPa. With the use of elemental data Schrauder et al. (1994) and Schrauder and Koeberl (1994) noted that the fluid compositions of Yakutian and Indian diamonds are similar to that of African diamonds. Melts of various compositional affinities (i.e. water-carbonate-potassium rich fluids) were identified by Navon et al. (1988) who interpreted the diamond inclusion data to mean that a potassic melt similar to erupted kimberlite or lamproite may have acted as a source for the potassic and hydrous inclusions in diamonds. More recently, Schrauder and Navon (1994) reported data in support of carbonatitic and waterrich hydrous fluids in fibrous diamonds from Zaire. According to these authors the compositional range between these two extreme fluid types may have resulted from partial melting of a wet carbonated peridotite, reaction of a hydrous fluid with carbonate-peridotite, mixing of carbonatitic and hydrous fluids, and fractional crystallization of a water-rich carbonatitic fluid.

Megacrystic diamonds Considerable debate exists as to the type of fluid present during the formation of diamonds and their inclusions (e.g. Harte et aE., 1980; Javoy et al., 1984; Richardson et al., 1984; Boyd et al., 1987; Ballhaus, 1993; Griffin et al., 1993; Deines and Harris, 1994; Lowry et al., 1994). While Harte et al. (1980) believed in the existence of CO*-bearing melts during the formation of peridotitic diamonds, in the case of Zairean diamonds, stable and radiogenic isotopic data suggest a close genetic link with kimberlite (Javoy et al., 1984; Boyd et al., 1987), the implication being that the diamonds formed as phenocrysts or cognate megacrysts. Other investigators have been rather non-committal about the “fluid” involved in chemical enrichment processes in diamond facies mantle and have argued that some of the trace element characteristics in garnet solid inclusions resulted from sub-solidus processes involving an “alkali, LREE and CO2 rich melt” (Richardson et al., 1984; Shimizu and Richardson, 1987; Richardson, 1990). Richardson et al. (1984) discussed the link between diamond formation and melt extraction but they did not elaborate on the “fluid” involved in diamond formation, other than to allude to the introduction of “diamondforming components” by metasomatic processes. Lowry et aZ., (1994) reported stable isotope data for almost one hundred diamonds from Finsch Mine, RSA and found disequilibrium features in some solid inclusions of peridotitic (P-type) and eclogitic (E-type) paragenesis. They stressed that interaction between the host lithologies (eclogite, peridotite) and the diamond-forming fluid was an on-going process such that the solid inclusions trapped inside megacrystic diamonds are “snapshots” of this process. They concluded that the diamond growth is associated with IgO depletion, 13C enrichment and Ca enrichment in both P- and E-type inclusions at Finsch. Reaction between carbonate-rich fluid and host rock would best explain such trends and is supported by the presence of CO2 rich fluids in garnet pyroxenite and eclogite (Luth and Scarfe, 1994). In a recent contribution, Deines and Harris

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(1994) commented that it is more likely that equilibria relevant to volatile rich systems, rather than in melts, can best be applied to processes involved in diamond growth. Ballhaus (1993) noted that the oxygen fugacity in the deep cratonic lithosphere is most likely governed by the activities of carbonate and COz (carboncarbonate-silicate), and that elemental C may be present during metasomatism given the solubility of carbon in metasomatic C-H-O fluids. Ballhaus (1993) suggested that the chemical character of the deep cratonic lithosphere could be the result of prolonged, widespread impregnation and concomitant depletion (i.e. melt loss) of the mantle by carbonatitic melts. In addition, Griffin et al. (1993) explained the chemical characteristics of diamond indicator minerals (i.e. garnets and chromites) as a result of the influx of carbonatitic fluid or melt into a mantle residue. Indeed it was proposed by these and other workers (Zhou et al., 1994) that silicate melt activity in diamond facies mantle produced a lower grade and quality of diamond (e.g. Shandong and Guizhou, E. China), while silicate melt inactivity maintained a higher grade and quality of diamond (e.g. Liaoning, E. China). The presence of silicate melts at the base of the cratonic lithosphere accounted for the loss of diamond facies mantle. Griffin et al. (1993) noted that both in the case of South Africa and Siberia, diamonds grew in response to short lived thermal effects as would be associated with magmatic intrusion.

Diamondiferous eclogites Apart from megacrystic diamonds, diamondiferous eclogites and peridotites provide valuable information on the processes active in the deep upper mantle. MacGregor and Manton (1986) showed that eclogites from Roberts Victor had a chemical signature consistent with derivation from subducted lithospheric material. They suggested that a MARID component may have been derived from eclogite and that this component may have acted as a metasomatic fluid in the overlying lithosphere (Thompson, 1992). The presence of jadeitic pyroxene in some diamonds may link the genesis of fluid bearing variety III diamonds with Pearson et al. (1991) studied erogenic peridotiteeclogite paragenesis. pyroxenite massifs and noted that pyroxenitic host rock may have formed from a melt derived from altered and subducted oceanic lithosphere. In order to account for the Pb and C isotopic data they suggested that a hemi-pelagic sediment component may have been involved in the recycling process. In summary, evidence for the nature and origin of melts in diamond facies mantle (tectonically inactive region) is being discovered from silicate (mica, apatite, garnet) and non-silicate (carbonate) inclusions in megacrystic diamonds and chemical effects in the constituent minerals of diamondiferous peridotites and eclogites. The type of melt varies from silicate (potassic and hydrous) to non-silicate (carbonatitic) and the erupted equivalents are similar to kimberlite, lamproite or carbonatite which probably have a sub-lithospheric provenance within the convecting upper mantle or recycled lithosphere.

399

Fluid processes in diamond Table 1. Mineralogy of xenoliths and megacrysts from tectonically inactive continental character Mantle petrology Garnet facies (75-200 km) Peridotite-pyroxenite

Mineralogy

Fluid character

Diopside, K-richterite, mica, Mg-ilmenite Phlogopite, edenitic-pargasitic amphibole diopside, K-richterite, lindsleyite-mathiasite (Dawson and Smith, 1977; Haggerty er al., 1983)

H,O-CO,, REE rich fluid (Dawson, 1987) Kimberlitic melt (Dawson and Smith, 1977; Jones ef al., 1982; Kramers ef al., 1983) K-rich hydrous fluid (kimberlitel lamproite) (Menzies et al., 1987) Potassic, LIL rich and water rich fluids (Erlank et al., 1987) Ultrapotassic melt (lamproite) (Waters, 1987) Fe-Ti rich silicate.melt (Harte er al., 1987).

Ilmenite, rutile, phlogopite, sulphides (i.e. IRPS) Diamond facies (180-220 km) Solid inclusions

regions and fluid

Sub-calcic garnets

Hydrated sheet silicates (mica?) carbonates, phosphates (apatite?) (Navon et al., 1988) Garnet, olivine, clinopyroxene Fluid inclusions

-

Diamond megacrysts

-

Alkali, REE, COz rich melt (Richardson et al., 1984) Sub-solidus fluids (Shimizu and Richardson, 1987) Potassic melt (Navon et al., 1988) Component of proto-kimerlite melts (Lowry et al., 1994) Carbonatitic to hydrous fluids (Schrauder and Koeberl 1994; Schrauder et al., 1994) COZ rich fluid (Harte er al., 1980) Kimberlite (Javoy et al., 1984; Boyd er al., 1987) Carbonate-melt (Ballhaus 1993; Griffin et al., 1993; Lowry et al., 1994) C-H-O fluid (Ballhaus 1993; Deines and Harris, 1994)

Garnet facies pyroxenites (75-200 km) Due to the ubiquity of garnet facies mantle rocks in kimberlite-borne xenolith suites, a considerable amount of information is available on the possible role of mantle fluids in the cratonic lithosphere. Hydrous garnet peridotite, micaamphibole-rutile-ilmenite-diopside (MARID) and ilmenite-rutile-phlogopitesulphide (IRPS) pyroxenite xenoliths provide valuable information about the nature and provenance of mantle fluids in the garnet facies mantle. In many cases the metasomatic changes observed in peridotites are inextricably linked to the presence or absence of cross-cutting pyroxenite veins. Consequently, evidence for the nature of mantle fluids in the garnet facies mantle is frequently drawn from the MARID/IRPS pyroxenite database. In one of the earliest contributions to the origin of MARIDs, Dawson and Smith (1977) coined the acronym MARID and suggested derivation from a kimberlitic magma under oxidizing conditions. While Jones et al. (1982) and

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Kramers et al. (1983) verified this conclusion with elemental and isotopic data, other workers suggested alternative origins for the MARID rocks on the basis of geochemical data. In the most detailed study of metasomatized peridotites and pyroxenites from Kimberley, Erlank et al. (1987) concluded that water-rich fluids, highly charged with potassium and other incompatible elements, were the best source fluids for the MARID pyroxenites (Navon et al., 1988). Erlank et al. (1987) did not propose a petrogenetic affinity between these MARID fluids and kimberlite or lamproite. After a comprehensive study of MARIDs, Waters (1987) opted for their derivation from lamproitic melts that she believed were present in garnet facies mantle, as a result of melting of wet garnet peridotite or hybridization of asthenospheric melts during transfer through the lithosphere. In contrast to some of the earlier work, Waters (1987) concluded that a simple relationship to kimberlite was not valid. However, Menzies et al. (1987) supported the earlier conclusions of Dawson and Smith (1977), and used available elemental and isotopic data to demonstrate a link between MARID pyroxenites and the influx of a kimberlite into garnet-facies mantle. A kimberlitic parentage for MARIDs is supported by recent experimental work (Sweeney et al., 1993) which showed that whilst peridotites occupy the field of “basalts” in the molecular tetrahedron, experimental melts from MARIDs lie in the field of melt compositions similar to the hydrous and carbonatitic fluids found in diamond facies mantle (Schrauder and Navon, 1994). Furthermore Sweeney et al. (1993) suggested that olivine fractionation from a kimberlite magma would leave a residual melt similar to MARID in composition. This contrasts with MacGregor and Manton (1985) who found that much of the LREE, K, Rb enriched component in eclogites was interstitial and that this MARID like component may have metasomatized the garnet facies mantle. Van der Laan and Foley (1994) performed experiments on MARID compositions and found that whilst the MARID solidus lies above the wet peridotite solidus, it is hundreds of degrees below the dry solidus (Fig. 1). At 50 kb it coincides with the shield geotherm and at 15 kb it is higher than the oceanic geotherm. During partial melting MARIDs can produce highly magnesian (Mg# = 80-90) silicate melts. For a review of mantle melting see Cox (1995).

Garnet facies peridotites Studies of garnet facies peridotites (Dawson, 1987; Erlank et al., 1987; Menzies et al., 1987) indicate a close association between the fluids involved in MARID genesis and the fluids responsible for the growth of phlogopite and richterite in garnet peridotites. The elemental fractionation in such small volume melts will, with time, generate greater isotopic heterogeneity in the shallow mantle (Fig. 6). While Dawson (I987) opted for a potassic, LIL and water-rich fluid, Erlank et al. (1987) suggested a potassic hydrous fluid. Menzies et al. (1987) suggested that the metasomatism observed in the wall rock garnet peridotite was part of a continuum of processes associated with the formation of


401

+I0 ALKALI BASALTS OCEANIC & CONTINENTAL) 0

KIMBERLITES -10

-20

-30

I

0.800

0.850

*‘Sr/*%r

Fig. 6. Chemical (i.e. Sr and Nd isotopes) heterogeneity in the lithospheric mantle (cold) compared to relative homogeneity in the asthenosphere (hot) (after Menzies et al., 1987). Melt stagnation in the mechanical boundary layer of the lithosphere, over time, has allowed elemental heterogeneity to develop into isotopic heterogeneity which is apparent in “cold” xenoliths entrained from beneath cratons. These mantle xenoliths and solid inclusions in diamonds provide information on “small-scale” mantle heterogeneity, while lamproites and micaceous kimberlites (Group II) provide information on “large scale” mantle heterogeneity. The dotted area shows the relative homogeneity in the source regions of melts derived from beneath the lithosphere. Peridotites associated with the genesis of such asthenosphere-derived melts, or peridotites modified by their movement, will have a restricted Sr-Nd isotopic composition. This accounts for the narrow range in Sr-Nd isotopes observed in “hot” spine1 and garnet facies xenoliths from off-craton areas where young lithosphere may have formed by accretion of such material.

MARID veins, involving kimberlitic melts. Ballhaus (1993) proposed that the characteristic depletion of the sub-cratonic lithosphere may be due to segregation and reaction with carbonate-rich kimberlitic or carbonatitic melts, a process that would effectively strip out all the elements soluble in carbonatite (Mg, Ca and Fe) and leave a residue enriched in silica and orthopyroxene. Mantle xenoliths from the Matsoku kimberlite (Lesotho) provided evidence of somewhat different melts in the garnet facies mantle beneath a stable continental region. Harte et al. (1987) described in detail the mineralogical characteristics of the peridotite-pyroxenite suite and provided evidence that the formation of the pyroxenite sheets (i.e. IRPS: ilmenite, rutile, phlogopite, sulphide) involved intrusion of silicate melts. They compared the Matsoku spine1 lherzolite xenoliths with those from the Basin and Range province of the western U.S.A., a

M. Menzies

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tectonically active region, where the peridotites have been affected by the movement of basaltic, and not kimberlitic, melts. In summary, evidence for the nature and origin of melts in garnet facies mantle (tectonically inactive region) can be found in metasomatic minerals in garnet peridotites and pyroxenites (i.e. low titanium mica, richteritic amphibole, apatite, titanates, rutile. ilmenite, sulphides and diopside). The general consensus is that the fluids are primarily silicate melts with affinities to kimberlite, lamproite or basalt. The melts with kimberlite or basalt affinity appear to have originated in sub-lithospheric sources. while those melts with affinities to lamproite or micaceous (Group II) kimberlite may represent complex hybrid melts with contributions from the lithosphere and the asthenosphere (Carlson, 1995 this issue).

TECTONICALLY

ACTIVE

CONTINENTS

AND OCEANIC

LITHOSPHERE

Garnet facies peridotite (75-200 km) In contrast to kimberlite-borne xenolith suites from stable cratonic areas, basalt-borne xenolith suites from tectonically active regions tend to be dominated by shallow, spine1 facies peridotites and pyroxenites with a paucity of deeper, garnet facies rocks. In a study of garnet-spine] peridotite xenoliths from the Sahara, Dautria et al. (1992) reported the presence of minor phlogopite and interpreted the chemical changes in the peridotites as being caused by the movement of carbonatitic melt. No primary carbonate was reported. In contrast, silicate melts were proposed by Ehrenberg (1979) as being primarily responsible for the distinctive chemical characteristics of metasomatized garnet lherzolites from the Colorado Plateau. Ehrenberg proposed that the silicate melt intruded hot deformed Fe-Ti rich peridotites. Evidence for the character of melts in garnet facies mantle (tectonically active region) is contained within metasomatic minerals in garnet-spine] peridotites. The type of fluid is primarily a silicate or non-silicate melt similar to an alkaline basalt or carbonatite, respectively. While the silicate melts appear to be primarily derived from sub-lithospheric reservoirs, because of their similarities to ocean island basal@, the carbonate melts may be produced by intra-lithosphere melting or may be generated as an immiscible fraction from sub-lithospheric basaltic melts. Insufficient isotopic data exists on the carbonate melt derivatives to decide on these alternatives.

Spine1 facies peridotites and pyroxenites (< 75 km) Spine1 facies mantle has received more attention than other mantle facies because of the ubiquity of spine1 peridotites and pyroxenites in basalt-borne xenolith suites. Fluid processes are apparent as the growth of metasomatic minerals in peridotites and chemical enrichment of primary minerals. The very


403

Table 2. Mineralogy of xenoliths and megacrysts from tectonically active continental and ocean basins and fluid character Mantle petrology Spinel-facies ( < 75 km) Peridotite

Mineralogy

Fluid character

Kaersutite, Fe-Ti phlogopite, apatite, magnetite, rutile, diopside -

Basaltic intrusions (Ehrenberg 1979, Wilshire, 1984) and/or carbonatitic melts (Menzies, 1983) Volatile-rich silicate melt or soluterich supercritical fluid (Lloyd et al., 1987) Carbonatitic melt (Wass et al., 1980; Yaxley et al., 1991; Hauri et al., 1993; Ionov et al., 1993; Rudnick et al., 1993; Chazot et al., 1995a,b)

Carbonate, apatite, Na-Cr clinopyroxene (Al poor), amphibole, mica, Cr spinel, monazite, (alkali and P-rich glass); Mg-Ca rich olivine Kaersutite, diopside, rutile, ilmenite, carbonate (magnesite) Amphibole, clinopyroxene, mica

Pyroxenite

Amphibole, apatite, biotite, diop side, ilmenite, titanomagnetite, rutile, sulphides (Wass et al., 1980) Clinopyroxene, orthopyroxene, sphene, titanomagnetite, apatite, calcite, feldspar (Lloyd, 1987) Mg diopside, apatite, Cr-spinel, Ca-Na-K aluminosilicates (Green and Wallace, 1988) Garnet-facies Peridotite

Silicate-carbonate melt (Schiano et 1994) Ti and alkali-rich hydrous fluid (Varne, 1970, Conquere, 1971) Alkali basalts (Irving, 1980; Fabries et al., 1989; Bodinier et al., 1990) Kimberlitelcarbonatite (Wass et al.. 1980) al.,

Volatile-rich silicate melt or soluterich supercritical fluid (Lloyd, 1987) Carbonatite melts (Green and Wallace, 1988)

(75-200 km) Phlogopite (Dautria et al., 1992)

Carbonatitic melt (Dautria 1992)

et al.,

high concentration of trace elements and the occurrence of amphibole and apatite in some lherzolites led to the conclusion that the fluids circulating in the mantle should be CO2 (Menzies and Wass, 1983) and/or HZ0 rich (Andersen et al., 1984). Although the principal fluid in the shallow upper mantle is C02, due to the lack of stable carbonate, experimental data reveal that these CO, (+ H,O) fluids cannot move in the mantle (Watson et al., 1990; Watson and Lupulescu, 1993). Consequently the principal metasomatic agents are interpreted to be silicate and non-silicate melts of basaltic, andesitic and carbonatitic composition. Silicate melts (equivalent

to intraplate basaltic melts). Frey and Green (1974) and

Irving (1980) proposed that alkaline silicate melts (basalts) were responsible for the chemical effects observed in spine1 lherzolite xenoliths from Victoria, Australia. Irving (1980) suggested that many of the pyroxenites found in xenolith suites were part of mantle vein networks that had been disrupted during entrainment and that the parental magma for those veins was an alkaline basalt. This was confirmed by isotopic data (Fig. 6) and was adopted by Wilshire (1984)

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M. Menzies and G. Chazot

to explain the vein systems in erogenic peridotite massifs (Conquere, 1971). The link between metasomatic processes and volcanism was emphasized by Menzies and Murthy (1980) in the basalt-xenolith suites of Ataq (Yemen) and Nunivak (Alaska) where the Sr-Nd isotopic composition of the basalts are clearly in equilibrium with some metasomatic or metasomatically modified minerals of the entrained peridotite xenoliths. More recently, Menzies et al. (1985) proposed a model of metasomatism along ramified veins of basanitic magma with repeated passage of silicate melts producing the range of chemical and isotopic compositions observed in mantle xenoliths. Near these veins, hydrous conditions allow the deposition of hydrous minerals such as amphibole or mica. Far from the vein, the increasing C02/C0, + HZ0 ratio of the metasomatic fluid produced enrichment of the peridotite in LREE in an anhydrous environment. Cooling of such conduits will form pegmatitic veins with amphibole, apatite, mica, feldspar. These observations of percolation of a silicate melt from a vein have been modelled as a chromatographic process by Navon and Stolper (1987) and as a percolation-diffusion process by Vasseur et al. (1991) from field and chemical observations in the Lherz erogenic peridotite (Bodinier et al., 1990). These models are able to describe the type of trace-element enrichment at various distances from a vein of infiltrated magma but are difficult to apply to natural observations. From a detailed study of xenoliths from Dish Hill in California, Nielson et al. (1993a,b) proposed that variations in the volume and chemistry of the metasomatizing melts can produce the chemical variations observed in these spine1 lherzolite xenoliths. The genetic link between erupted magmas, pegmatitic veins and megacrysts is further supported in many cases by similarities in the Sr-Nd-Pb isotopic data (Menzies et al., 1985; Ben Othman et al., 1990). Ion probe studies of the constituent minerals in Australian xenoliths led Vannucci et al. (1994) to conclude that metasomatic processes had involved both alkaline (basaltic) and carbonatitic melts. Sen et al. (1993) investigated the constituent minerals in spine1 peridotites from Oahu and assigned the characteristic spoonshaped REE profiles of the clinopyroxenes to metasomatic reaction with silicate melts similar to the Honolulu volcanics. Chromatographic effects were suggested to be the cause of any changes in the REE character of xenoliths from different sites within the same island. The xenoliths from Hawaii are not so markedly depleted as those from abyssal locations, in having higher modal amounts of clinopyroxene which in itself has higher LREE content, sodium and titanium than the abyssal peridotites. These authors concluded that the oceanic mantle must be vertically zoned from a shallow LREE depleted upper plagioclase facies zone through a less LREE depleted spine1 facies zone to a lower garnet facies zone with a marked enrichment in the LREE perhaps due to melt influx. Silicate melts (equivalent

to continental crustlandesite). Vidal et al. (1989) investigated a suite of harzburgites, dunites and lherzolites from Batan Island, Philippines which contained metasomatic mica and diopside and were cross-cut by pargasite bearing veins, in turn cross-cut by mica-bearing veins. Because of


405

the similarity in isotopic chemistry between the metasomatic minerals and western Pacific island arc volcanic rocks, the metasomatic fluid was believed to have originated in the subducted slab (sediment and oceanic crust). Arc processes were also invoked by Nimz et al. (1993) as an explanation of the geochemistry of chemically enriched pyroxenites from Mexico. They went so far as to suggest that arc-related basalts, or fluids derived from these rocks in the Cenozoic, may have affected the lithosphere beneath the western U.S.A. While this contrasts with the suggestion of Proterozoic-age processes it is consistent with the trend towards “crustal” compositions (i.e. high Sr and low Nd isotopic ratios) (Menzies et al., 1985) and the considerable range in C isotopes (Menzies, 1990) comparable to the range observed in E-type diamond inclusions, whose origin is believed to involve recycling processes. Schiano et al. (1994) investigated the melt inclusions in spine1 facies mantle minerals from oceanic (and continental) xenoliths from the Society Islands, Canary Islands, Kerguelen and the Comores. The secondary inclusions could be sub-divided into silicate melt inclusions, carbonate-rich inclusions and CO2 fluid inclusions. The most common daughter mineral assemblage in these inclusions was kaersutite + diopside + rutile + ilmenite + carbonate (magnesite) somewhat like the MARID assemblage except that the amphibole is kaersutite and no mica was reported. Such minerals were precipitated from silicate-carbonate melts thus accounting for their difference to the host peridotite assemblage. Carbonate melts. The presence of pyroxenite xenoliths derived from basaltic or andesitic melts can be contrasted with apatite pyroxenites derived from carbonatites (Wass et al., 1980). Mineralogical and geochemical evidence for peridotite interaction with carbonatitic melts has been reported from several spine1 lherzolite xenolith suites. Apatite and amphibole are common phases, sometimes disseminated in the matrix, or in veins in association with other metasomatic minerals. While amphibole is always contained in melt pockets surrounded by silicate glass and newly formed olivine, clinopyroxene and spinel, many xenoliths contain melt pockets without residual hydrous minerals. This may (CominChiaramonti et al., 1986; O’Reilly, 1987) or may not (Ionov et al., 1995) be due to total melting of amphibole or phlogopite. Many xenoliths are very orthopyroxene-poor and show the destabilization of the orthopyroxene into clinopyroxene and olivine (Yaxley et al., 1991; Hauri et al., 1993; Ionov et al., 1993; Rudnick et al., 1993). A surprising feature of these xenolith suites is the lack of primary carbonates. Dolomitic carbonate, accompanied by carbonatitic melt, has only been reported as an important phase in xenoliths from Spitsbergen (Ionov et al., 1993). This may be explained by very rapid dissociation of the carbonates during syn-entrainment decompression melting of the xenoliths (Canil, 1990). The chemical composition of the secondary minerals formed by interaction between the primary mineralogy and the carbonatitic melts is in agreement with experimental data, with high Mg and Ca content in the secondary olivines (Hauri et al., 1993; Ionov et al., 1993), and low Al (Hauri

406

M. Menzies

and G. Chazot

et al., 1993; Rudnick et al., 1993) and high Na (Yaxley et al., 1991) content in the secondary clinopyroxene. The whole rock mineralogy is also sensitive to decarbonation reactions, with the replacement of orthopyroxene by clinopyroxene leading to high Ca content and high CaIAI ratio for the metasomatized peridotites (Yaxley et al., 1991; Rudnick et al., 1993; Ionov et al., 1995). Varne (1970) and Varne and Graham (1971) described xenoliths from Ataq, Yemen modified by infiltration of hydrous silicate melts, with crystallization of pargasite and enrichment of the clinopyroxenes in REE. A detailed reinvestigation of the constituent minerals in spine1 facies xenoliths from the same diatremes has revealed a more complex history. Using laser fluorination, oxygen isotope data and ion microprobe elemental data, Chazot et al. (1995a,b) found that these peridotites had been modified initially by reaction with a fluid leading to crystallization of amphibole, apatite and clinopyroxene. All these minerals were in elemental and oxygen isotopic equilibrium and show large negative anomalies for Zr, Hf and Ti on normalized trace element diagrams, and overall have a trace element character identical to that of carbonatites. The initial melt to invade the spine1 facies mantle was a carbonatite and it produced textural relationships in these natural peridotites that compare favourbly with the textures observed in experiments where carbonatitic melts are reacted with peridotite (Wallace and Green, 1988; Dalton and Wood, 1993). Subsequent decompression melting (i.e. wet melting) of these amphibole peridotites generated glasses with a silica rich composition similar to those reported by Schiano et al. (1994a, b). In spine1 facies mantle, the effects of decarbonation reactions have been noted for a long time as a decoupling between the major and trace element composition (Frey and Green, 1974). The reaction between carbonatitic melts and peridotites produced characteristic trace element patterns in good agreement with experimental data but different from the patterns produced by interaction with a silicate melt. While the behaviour of trace elements is highly dependent on the mineralogy (e.g. amphibole), the carbonated metasomatic melts have a high content of incompatible elements including Th and U and a low content of Zr, Hf, Ta and Ti (Yaxley et al., 1991; Hauri et al., 1993; Ionov et al., 1993,1995; Rudnick et al., 1993). One interesting feature is the systematic fractionation of the Zr/Hf ratio. In addition Sweeney et al. (1992) reported low partition coefficient for Nb between carbonatitic melt and garnet and amphibole, which is in good agreement with the high Nb content of extrusive carbonatites and the high Nb content of carbonatite-metasomatized peridotites (Yaxley et al., 1991; Rudnick et al., 1993). However, others report very low Nb content in carbonatebearing metasomatized lherzolites (Hauri et al., 1993; Ionov et al., 1993, 1995). Hauri et al. (1993) reported a secondary phase assemblage (clinopyroxene, spine1 and apatite) in spine1 facies xenoliths from the Pacific and elegantly demonstrated that these were in equilibrium with carbonate-rich melts. The Sr-Nd-Pb isotopic characteristics of these secondary phases may have originated from carbonate-bearing sources in mantle plumes, the ultimate cause of much of


401

the volcanism in this area. Interestingly the isotopic characteristics of the Samoan xenoliths are similar to global sediments and/or processes found in island arcs, leading to the conclusion that a recycled component may be part of this system. Hauri et al. (1993) concluded that the isotopic data were consistent with recycling of sediments, and that carbonate metasomatism pointed to CO2 transport through subduction zones and into the upper mantle recycling system. Carbonatitic melts have clearly played a vital role in metasomatic processes within the spine1 facies mantle but a considerable amount of evidence also exists for silicate melts of basaltic and andesitic composition having been transfered through the lithosphere by crack propagation. However, many samples described as modified by silicate melts have chemical characteristics compatible with carbonatitic interactions with typical Ti and Zr depletion (for example Roden et al., 1984; Vannucci et al., 1991) and should be reconsidered in the light of the new data about carbonatitic melts in the mantle. The petrogenetic link between silicate melt and carbonate melt metasomatism is not yet understood. Considering the possibility that carbonate fluids are generated by exsolution from silicate melts, carbonate metasomatism is likely to be a common phenomenon around uprising mantle diapirs and should shortly precede the entrainment of the xenoliths by the silicate magmas. Ionov et al. (1995) proposed that fluid influx prior to entrainment of the xenoliths may induce the formation of melt pockets around hydrous minerals or even in anhydrous peridotite around clinopyroxene and spinel. In the case of amphibole- or mica-bearing xenoliths, we should therefore envisage at least two stages of metasomatism with crystallization of hydrous minerals from a hydrous silicate melt followed by carbonate metasomatism. Since no real distinction exists between the chemical signatures produced by fluid processes involving various mantle fluids (e.g. silicate melts, carbonatitic melts, hydrous fluids) Zinngrebe and Foley (1994) studied metasomatized peridotites in order to evaluate the effects of low oSiOz melts. They compared reactions that led to loss of orthopyroxene with orthopyroxene free, low orSiOz metasomatism and demonstrated that dunite-wehrlite parageneses can be explained by interaction with orthopyroxene-undersaturated silicate melt alone and the need for carbonatitic melts is questioned. Clearly the time has arrived when many of the classic metasomatized suites need to be carefully evaluated using experimental and ion microprobe techniques (Vannucci et al., 1994). In summary, evidence for the nature and origin of fluids in spine1 facies mantle (tectonically active continental regions and ocean basins) is contained within metasomatic minerals in spine1 peridotites and pyroxenites (i.e. high titanium mica, kaersutite-pargasite amphibole, Fe-Ti pyroxene, apatite and carbonate) and fluid inclusions. The type of fluid is primarily a silicate or non-silicate melt similar to an alkaline basalt or carbonatite respectively. While the silicate melts appear to be primarily derived from sub-lithospheric reservoirs because of their similarities to ocean island basalts, the carbonate melts may have a more complex origin.

408

M. Menzies and G. Chazot Table 3. Polybartc mantle metasomatism: continents and ocean basins

Mantle petrology Garnet facies (75-200 km) Garnet peridotites Garnet pyroxenites

Diamond facies (180-220 km) Diamonds and diamond inclusions

(a) Tectonicallv inactive continent (craton) Kimberlite Basalt

Dawson, lYX7 Menzies er al., 1YX7 Harte er ul.. 1975 Dawson and Smith, IY77 Jones et al., 1982 Kramers et al., 1083 Sweeney ef ul., 1993

Harte cr al.. lY87

Carbonatite

Dautria et al., 1992 -

-

-

Harte er ~1.. 1980 Navon et al., 1088 Schrauder and Koeberl. lYY4 -

Ballhaus, 1993 Griffin et al., 1993 Schrauder and Navon, 1994 Lowry et al.. 1994 Schrauder et al., 1994

-

(b) Tectonically active continent or ocean basin Spine1 facies (475. Schrauder M. and Koeberl C. (1994) Trace element analyses of fluid-bearing fibrous diamonds from Jwaneng (Botswana) by neutron activation analysis. Mineral. Mag.,58A, X1 I-812. Schrauder M. and Navon 0. (1994) Hydrous and carbonatitic mantle fluids in fibrous diamonds from Jwaneng. Botswana. Geochim. Cosmochim. Acfa 58, 761-771. Schrauder M., Navon 0, Szafranek D.. Kaminsky F. V and Galimov E. M. (1994) Fluids in Yakutian and Indian diamonds. Mineral. Mug. 58A. XI.%X14. Schiano P., Clocchiatti R. and Shimizu N. (1994a) Melt mclusions trapped in mantle minerals: a clue to identifying metasomatic agents in the upper mantle heneath continental and oceanic intraplate regions.

Mineral. Msg. 58A. 807-808. Schiano P.. Clocchiatti R. and Shimizu N., Weis D. and Mattielli N. (1994b) Cogenetic silica-rich and carbonate-rich melts trapped in mantle minerals in Kerguelen ultramafic xenoliths: implications for metasomatism in the upper oceanic mantle. Earfh Plurrer. Sci. Left. 123, 167-17X. Sen G.. Frey F. A.. Shimizu N. and Leeman W. P. (lY93) Evolution of the lithosphere bencath Oahu. Hawaii: rare earth element abundances in mantle xenoliths. Eurfh Planet. Sci. Left. 119, 53-69. Sharp 2. D. (1990) A laser based microanalytical method for the in sifu determination of oxygen isotope ratios of silicates and oxides. Geochim. ~‘osmochim. A~tr 54. 13% 1.357. Shimizu N. and Richardson S. H. (IYX7) Trace element abundance patterns of garnet inclusions in peridotitesuite diamonds. Geochim. C‘o.smochim Acftr 51. 75.5-7X. Sweeney R. J., Thompson A. B. and Ulmer P. (1993)Phase relations of a natural MARID composition and implications for MARID genesis. lithospheric melting and mantle metasomatism. Confrih. Mineral. Pefrol. 115, 225-241. Thompson A. B. (lYY3) Water in the Earth’s upper mantle. Nature 358, 2YS-302. Vannucci R.. Shimizu N.. Bottazzi P., Ottolini I... Piccardo G. B. and Rampone E. (IYYl) Rare earth and trace element geochemistry of clinopyroxencs from the Zabargad peridotite-pyroxenite association. J. Pefrol. Special Lherzolites Issue. 255-269. Vannucci R.. Bottazzi P.. Ottolini L., Dal Negro A. and Piccirillo E. M. (lYY4) The trace element variations in clinopyroxenes from spine1 peridotitc xcnoliths from western Victoria (Australia). Mineral. Mug. 58A. Y32-

933. Varne R. (lY70) Hornhlende lhcrzolite and the upper mantle. (‘onfrih. Minerul. Perrol. 27, 45-51, Varne R. and Graham A. L. (1971) Rare earth abundances in hornblende and clinopyroxene of a hornblende lherzolite xenolith: implications for upper mantle fractionation processes. Earth Planet. SC;. Leff. 13. 1l-18. Vasseur G.. Vernieres J. and Bodinier J.-L. (1991) Modelling of trace element transfer between mantle melt and heterogranular peridotite matrix. J. Pefrol. Special Lherzolites Issue, 41-54. Vidal Ph., Dupuy C., Maury R. and Richard M. (19X9) Mantle metasomatism above subduction zones: trace-element and radiogenic isotope characteristics of peridotite xenoliths from Batan Island (Philippines). Geology 17, II IS-I 118. Wallace M. E. and Green D. H. (IYXX) An experimental determination of primary carbonatite magma composition. Nature 335, 34.%346. Wass S. Y., Henderson P. and Elliott C. J. (IYXO) Chemical heterogeneity and metasomatism in the upper mantle: evidence from rare earth and other elements in apatite-rich xenoliths in basaltic rocks from eastern Australia. Phil. Trans. Roy. Sot. (Land.). 297. 33sF-346. Waters F. G. (1987) A geochemical study of metasomatized peridotite and MARID nodules from the Kimberley Pipes, South Africa. Ph.D. Thesis, University of Cape Town (unpubl.). Watson E. B. and Lupulescu A. (1993) Aqueous fluid connectivity and chemical transport in clinopyroxenerich rocks. Earth Planet. Sci. Left. 117, 279-294. Watson E. B., Brenan J. M. and Baker D. R. (IYYO) Distribution of fluids in the continental mantle. In Continental Ma&e (Menzies M. A.. Ed.). pp. I1 l-125. Clarendon Press, Oxford. White R. S. (1988) The Earth’s crust and lithosphere. 1. Pefrol. Special volume: Oceanic and continental lithosphere: similarities and differences. I-IO. Wilshire A. G. (1984) Mantle metasomatism: the REE story. Geology 12. 395-398.


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Wvllie P. J. (1987) Metasomatism and fluid generation in mantle xenoliths. In Mantle Xenoliths (Nixon, P. H., Ed.), pp. 609-623. Wiley. Wyllie P. J. (1989) Origin of carbonatites: evidence from phase equilibrium studies. In Carbonatites: Genesis nnd Evolution (Bell, K., Ed.). pp. 500-545. Unwin Hyman, London. Wyllie P. J., Baker M. B. and White B. S. (1990) Experimental boundaries for the origin and evolution of carbonatites. Lithos 26, 3-19. Yaxley G. M., Crawford A. J. and Green D. H. (1991) Evidence for carbonatite metasomatism in spine1 peridotite xenoliths from western Victoria, Australia. Earth Planet.Sci. Lett. 107, 30.5-317. Zhou J., Griffin W. L., Jaques A. L., Ryan C. G. and Win T. T. (1995) Geochemistry of diamond indicator minerals from China. Proceedings of the Fifth International Kimerlite Conferece, Zinngrebe E. and Foley S. (1994) Metasomatism in two natural peridotites: the effects of low-a!%02 melts. Mineral. Mug. 58A, 1006-1007.