kimberlites, orangeites, and related rocks || petrogenesis of orangeites and kimberlites

Whether you are really right or not doesn't matter; it s the belief that coullts.

Robertson Davies

Omnia mutantur, nos et mutamur in illis. (Anon)

PETROGENESIS OF ORANGEITES AND KIMBERLITES

Many years of study have not resulted in general agreement concerning the petrogenesis of continental alkaline rocks. Unfortunately, petrologists, acting from a stance of promoting their own particular bias, be it geophysical, experimental, geochemical, or tectonic, have devised contradictory models for a given magma type. Once introduced, such models are commonly promoted by their originators without impartial reflection on all aspects of the problem or consideration of competing hypotheses.

This unsatisfactory situation arises from the absence of modern analogues of many varieties of alkaline magmatism and the realization that the sources of alkaline magmas may be developed by a wide range of physicochemical processes operating at inaccessible depths in the mantle. Thus, there are few constraints upon the imagination of petrologists seeking to explain the genesis of alkaline rocks and the subject might be regarded as being closer to "petromancy" than an exact science.

Kimberlites have remained challenging objects for petrogenetic speculation since their discovery over 100 years ago. Hypotheses for their genesis have ranged from the bizarre, such as meteorite-generated electrical discharges (Khazanovich-Vulf 1991) and mobilized sedimentary breccias (Mikheyenko 1977), to the commonplace, such as partial melts of asthenospheric (Canil and Scarfe 1990) or lithospheric mantle (Foley 1988, Bailey 1993). Consequently, a comprehensive petrogenetic hypothesis for kimberlites has not yet been devised.

Orangeites have not previously been considered as a distinct magma type; consequently their genesis has been discussed as though they are merely varieties of kimberlite (sensu lato).

This chapter reviews hypotheses and evidence pertaining to the genesis of kimberlites and orangeites and presents some petrogenetic speculations for both magma types, based on the conclusions of the review. It is unrealistic to expect these speculations to represent the final word on the topic, and, undoubtedly, the petrogenetic schemes for both magma types will be revised in the light of further studies.

303 R. H. Mitchell, Kimberlites, Orangeites, and Related Rocks© Plenum Press, New York 1995

304 CHAPI'ER4

4.1. GEOCHEMICAL MODELS OF ORANGEITE PETROGENESIS INVOLVING LIMITED PARTIAL MELTING OF LHERZOLITIC SOURCES

During the 1970s geochemical studies of a variety of mantle-derived alkaline rocks, coupled with the recognition of metasomatized mantle-derived xenoliths, demonstrated that partial melting of a simple four-phase garnet lherzolite source rock was inadequate as a means of explaining their incompatible element contents. Consequently, older hypotheses invoking high degrees (15-30%) of partial melting of mantle lherzolite followed by extensive fractional crystallization (O'Hara and Yoder 1967, MacGregor 1970, Anderson 1984) were overthrown in favor of models favoring small «10%) amounts of partial melting of fi ve- or six -phase lherzolite sources, e.g., phlogopite garnet lherzolite (Dawson 1972). The presence of minerals such as phlogopite or amphibole was required in the sources to support the geochemical claims that incompatible elements must be sequestered in hydrous phases. The various hypotheses differ with respect to the mineralogy of the source rocks, the extent of partial melting, and the origins of the incompatible element-bearing phases. Fractional crystallization is not considered to play a significant role in the evolution of these magmas. Thus, it is widely believed that observed, whole rock, incompatible element abundances and distribution patterns can be related to the degree of melting and type of source rocks by geochemical modeling, using the forward or inverse approaches.

Most of the partial-melting models are variants of hypotheses developed to explain the geochemistry of oceanic basaltic rocks, and their relevance to exotic continental alkaline magmas has not been seriously questioned by their proponents. Typically, the hypotheses are rooted in geochemical models based on trace element abundances, especially the rare earth elements, and mineralogical and major element constraints are rarely considered.

4.1.1. Earlier Hypotheses

Earlier hypotheses, suggesting that "kimberlitic" (sensu lato) magmas are produced by small amounts of partial melting of lherzolitic sources in the mantle, have been reviewed by Mitchell (1986). Most of these hypotheses were presented prior to the mineralogical and isotopic studies, summarized in this monograph, which have demonstrated that "kimberlites" (sensu lato) are not a single group of cogenetic rocks. Consequently, common to all of these hypotheses is the assumption that kimberlites and orangeites are derived from the same source rocks by differing degrees of partial melting. Clearly, in the light of current studies this assumption is no longer valid.

Hypotheses presented by Mitchell and Brunfelt (1975), Paul et al. (1975), and Cullers et al. (1982) were devised to explain the potassium contents and high LalYb ratios of "kimberlites" (sensu lato). These hypotheses demonstrated that the requisite LalYb ratios could be produced by extremely small (<l %) degrees of partial melting of phlogopite garnet lherzolite sources with chondri tic-like REE abundances. The generally higher LalYb ratios of "micaceous kimberlites," i.e., orangeites, were considered to require lesser amounts of partial melting (0.3-0.4%) than mica-poor kimberlites (0.7-0.9%).

PETROGENESIS OF ORANGEITES AND KIMBERLITES 305

Apart from the absence of mineralogical and major element constraints, all of the above hypotheses were unable to model the observed abundances of trace elements in the rocks, and differentiation or other processes had to be appealed to, to remedy this deficiency. Subsequently, enrichment of the source in REE above chondri tic abundances by cryptic metasomatism (Dawson 1984, Mitchell and Carswell 1976) was considered as a possible means of producing melts at the lower limits of the spectrum of "kimberlite" REE abundances (Mitchell 1986). However, this model still required very small (0.5-0.2%) amounts of partial melting of the source.

The very low volumes of postulated melts have typically been regarded as a major disadvantage of the above models. Consequently, there have been doubts that such melts could ever segregate and escape from their mantle source regions. However, recent theoretical and experimental studies (McKenzie 1989, Hunter and McKenzie 1989) of the migration of small volume melts in the mantle indicate that these concerns may be groundless.

A significant conclusion of the earlier studies was that, in addition to phlogopite, a phosphate and/or titanate must be present in the source rocks of kimberlitic magmas to account for their P, Sr, REE, Th, U, Zr, Nb, and Ta contents (Fesq et al. 1975, Kable et al. 1975). This observation led Mitchell (1986) to propose that "kimberlitic" (sensu lato) magmas could be derived by 1-8% partial melting of a patently metasomatized mantle source (Dawson 1984) consisting of an apatite- and K-Ti-richterite-bearing garnet lherzolite. In this model the high LalYb ratios of the partial melts were believed to reflect those of their source, rather than resulting from element fractionation during melting of a low LalYb source. The relatively high volumes of the melts produced, compared to those required by melting of cryptically metasomatized garnet lherzolite, alleviated the supposed melt extraction problems.

In summary, although most of the earlier partial melting models are now primarily of historical interest, they were important in that they indicated kimberlites (and other alkaline rocks) must be produced by the melting of metasomatized mantle.

4.1.2. Melting of Enriched Mantle and Peridotite Entrainment

The first model concerned specifically with the petrogenesis of orangeites (referred to as "group 2 ultrapotassic kimberlites") was initially presented by Fraser (1987) and, ultimately, by Fraser and Hawkesworth (1992). The model is based on geochemical and isotopic studies of the Finsch intrusion undertaken by Fraser (1987) and Fraser et al. (1985). These data have been reviewed in Sections 3.7 and 3.8.

Fraser and Hawkesworth (1992) consider that the geochemical characteristics of the Finsch rocks result from the mixing of a melt component with 50-60% of entrained garnet lherzolite. The high Nd and Sr contents, low Sm/Nd ratios, and high TalYb ratios relative to continental crust and local country rocks are interpreted as indicating that contamination of the magmas by crustal material is not significant.

The melt component is considered to have very high incompatible element contents and be formed by very small « 1 %) degrees of melting of a source enriched in these elements relative to depleted or primitive mantle. Fraser and Hawkesworth (1992) demonstrated that the melt component dominates the incompatible trace element content

306 CBAPTER4

of the hybrid magma. Therefore, the Nd, Sr, and Pb isotopic ratios are considered to reflect those of the source. The isotopic data (see 3.8) are interpreted to suggest that the melt was derived from ancient trace-element-enriched portions of the upper mantle, most probably situated within the subcontinental mantle lithosphere. Fraser and Hawkesworth (1992), in agreement with Wyllie (1980) and Mitchell (1986), regard this source as a phlogopite magnesite peridotite.

The disaggregation and incorporation of old depleted lithospheric mantle into the melt, during ascent, generates the high compatible-element contents of the hybrid magma and the olivine macrocrysts. In contrast to Tainton and McKenzie (1994), Fraser (1987) noted that xenoliths which might represent source material are not present in the magma.

Fraser et al. (1985) suggested that enrichment of the source in incompatible elements was due to introduction of H20-rich fluids. However, Fraser (1987) stated that the origins of the trace element enrichment are unresolved and compatible with derivation from either recycled pelagic sediments or intra-mantle processes involving the migration and crystallization of small volume silicate melts. Fraser and Hawkesworth (1992) do not discuss the nature of the enrichment process.

Fraser (1987) considered that the Finsch intrusion and other orangeites were emplaced in response to the opening of the South Atlantic (see4.5.2.1). Thus, the propagation of the oceanic extensional tectonic regime into the subcontinental mantle resulted in small degrees of decompressional melting of the ancient enriched lithospheric sources. Fraser (1987) proposed a two-stage model of melt migration. An initial slow ascent of melt along grain boundaries into overlying depleted lherzolitic mantle, followed by rapid ascent of melt plus entrained lherzolite along major cracks produced during uplift of the lithosphere.

The major contributions to orangeite petrology by Fraser (1987) and Fraser and Hawkesworth (1992) were the realization that orangeites are hybrid rocks, whole rock compositions do not represent those of the magma, and parental melts are most probably derived from ancient lithospheric sources. However, the partial-melting model suggested provided no advances as it merely reiterated the earlier models, discussed above, in requiring very small degrees of melting of a cryptically metasomatized five-phase lherzolitic source.

Fraser and Hawkesworth (1992) also concluded that their model of hybridization by peridotite entrainment is directly applicable to archetypal kimberlites. Following Smith (1983), Fraser and Hawkesworth (1992) suggestthat the melt component is derived from asthenospheric sources.

4.1.3. Three-Stage Processes-Depletion, Enrichment, and Melting

Tainton (1992) and Tainton and McKenzie (1994) have presented a general model to explain the geochemistry of "kimberlites and lamproites." Note, in the latter paper unevolved orangeites are termed "group II kimberlites," the Sover North evolved orangeite is termed "lamproite," and the only bona fide lamproite considered is the atypical, altered, and contaminated (Jaques et al. 1989b) Argyle lamproite. In their model, Tainton and McKenzie (1994) make no real distinction between archetypal kimberlites and orangeites.


The approach used is a variant of the inversion model of McKenzie and O'Nions (1991), devised to explain the trace element characteristics of basaltic rocks formed by partial melting of a primitive garnet lherzolite source. The model primarily uses REE distribution patterns, in conjunction with trace element crystal-liquid distribution coefficients, to estimate the degree of partial melting, depth of melting, and the modal mineralogy of the residue. The model was modified by Tainton and McKenzie (1994) by the addition of phlogopite, apatite, and chrome spinel to the source.

In agreement with Fraser (1987), Tainton (1992) considers orangeites and kimberlites as hybrid rocks and that initial partial melts have been contaminated by entrainment of mantle-derived peridotite (see 3.7). Tainton and McKenzie (1994) assume that the observed incompatible element contents of orangeites and kimberlites may be corrected for hybridization, and the corrected composition represents that of the initial partial melts. The amount of contamination is estimated by plotting the Ni contents of bulk rocks against their macrocrystal olivine content and by assuming that the contaminant contained 2500 ppm Ni (Tainton 1992). The calculated Ni concentration of the melt for zero contamination is then used to devise a factor which is applied to the concentrations of incompatible elements to correct for the dilution by "olivine."

Using these corrected data, Tainton and McKenzie (1994) show that the trace element geochemistry of orangeites and kimberlites cannot be modeled by simple partial melting (0.3-0.4%) of a primitive mantle or MORB-type garnet lherzolite source. Consequently, they propose a three-stage model in which an initial stage of extensive partial melting, resulting in the depletion and fractionation of incompatible elements, is followed by "metasomatic enrichment" involving addition of an incompatible element-rich melt. Orangeites and kimberlites are then produced by the partial melting of this derivativeenriched source.

As an example of their model, Tainton and McKenzie (1994) calculate that the trace element geochemistry of the Bellsbank orangeites can be reproduced by 24% melting of garnet-bearing primitive mantle, followed by adding to this depleted material 7% of a metasomatic melt generated by 0.5-0.3% melting of a MORB-type source, and subsequent removal of 0.4-0.3% melt from this second-stage enriched source. The mineralogy of the latter is calculated to consist of 72% olivine, 24% orthopyroxene, 1.3% clinopyroxene, and 1.8% garnet. Broadly similar results are obtained for other orangeites, kimberlites, and Argyle lamproite (see Table 1 of Tainton and MacKenzie 1994, Tainton 1992). Each magma-forming event differs only in the degree of initial depletion and amount of secondary enrichment. The initial depletion event, involving liquids in equilibrium with garnet, is required to explain the heavy REE abundances of the rocks. Small amounts of apatite in the source are considered not likely to have an important effect on REE contents of melts, although it is recognized that phlogopite plays a significant role in controlling alkali element abundances. In all cases examined, with the exception of Sover North and Finsch, Tainton and McKenzie (1994) calculate that the observed content of potassium is less than calculated. Hence, phlogopite is required in the source rocks and may remain in the residua.

Tainton and McKenzie (1994) propose that the source of the evolved Sover North orangeites underwent 20% melting at 90-km depth, followed by the addition of 10% enriched melt, whereas the source of the unevolved Sover orangeites experienced 20%

308 CHAPTER 4

melting at 83-km depth and only 6% secondary enrichment. Both magmas were subsequently generated by 0.3-0.4% melting of these enriched sources. It follows from these conclusions that this scheme does not admit the possibility that Sover North rocks might be formed from a differentiate of an unevolved orangeite magma.

The initial extensive partial melting which generated the depleted mantle is believed by Tainton and McKenzie (1994) to represent extraction ofkomatiite during the Archean. Isotopic constraints for orangeites (and lamproites) suggest that secondary enrichment of the depleted material also occurred in Archean or Proterozoic times, and the enriched source was subsequently isolated from convecting mantle for a long time. The origins of the secondary enriching melts from MORB-type sources are ultimately placed in the asthenosphere, and Tainton and McKenzie (1994, p.813) state that "although the total amount of melt added during enrichment is well-constrained, there are no thermal or geochemical restrictions on the number of events involved."

A surprising conclusion of Tainton and McKenzie (1994) is that incompatible element-enriched alkaline rocks could be derived from depleted harzburgitic sources. Tainton and McKenzie (1994) further claim that the common, granular-depleted garnet phlogopite peridotite xenoliths present in kimberlites are examples of the source rocks of kimberlite and orangeite magmas. This conclusion is at variance with the commonly held view that these xenoliths are merely mantle-derived xenoliths unrelated to their host magmas. Arguments presented by Tainton and McKenzie (1994) for their conclusion rely upon the proposed small volumes of the melts involved. Thus, in their model, because melting is limited to grain boundaries, source rock entrainment must be a necessary consequence of melt separation from these residua. Tainton and McKenzie (1994) consider that such garnet lherwlite xenoliths have modes and major and trace element compositions similar to those calculated for their postulated sources.

An important consequence of this model is that the difference in isotopic compositions of kimberlites and orangeites requires source enrichment to have occurred at different times but not by different processes. Thus, Tainton and McKenzie (1994) believe that both magmas are generated by the same processes from the same source rocks and that archetypal kimberlites could also be lithospheric melts. This conclusion is not in agreement with hypotheses advanced by other petrologists regarding the depth of kimberlite generation (see Smith 1983, Canil and Scarfe 1990, Ringwood et al. 1992, Edgar and Charbonneau 1993, Haggerty 1994, this work).

Tainton and McKenzie's (1994) hypothesis follows an approach devised for common basaltic rocks and is therefore unlikely to be appropriate for alkaline rocks of extreme composition, such as orangeites and lamproites. It may be relevant to the generation of archetypal kimberlites, but not in the tectonic setting visualized by the authors (see 4.6.1). Mineralogical and petrological evidence suggests the hypothesis is based upon incorrect premises. The major problem being the lack of serious consideration given to sources other than phlogopite garnet lherzolite. This a priori assumption regarding the nature of the source necessarily predetermines the outcome of the mathematical modeling. Hence, the latter, which is correct and internally consistent, gives the hypothesis an unwarranted verisimilitude as a mathematically appropriate solution can always be found.

A further aspect of the model not discussed in detail by Tainton and McKenzie (1994) is the relevance of the crystal-liquid distribution coefficients, derived from basaltic


magmas, to the alkaline magmas considered. The latter differ significantly in composition from basaltic magmas, so distribution coefficients will differ correspondingly. In fact, Tainton and McKenzie (1994, pp. 808, 811) note that the distribution coefficients appropriate to these magmas are very poorly determined.

The hypothesis proposes extraction of a variety of distinct magma types from the same source. Tainton and McKenzie (1994) do not provide any realistic rationale for their choice of a lherzolitic source and do not consider any of the recent experimental and geochemical evidence pertaining to the petrogenesis of lamproites and kimberlites (see 4.5,4.6), which indicate very different sources. For example, Foley (l992a) has noted that models for the origin of ultrapotassic melts by partial melting of phlogopite-bearing lherzolites are inconsistent with the array ofliquidus experimental results on ultrapotassic compositions (see 4.2.2). This discrepancy between partial melting models of the type advanced by Tainton and McKenzie (1994) or Fraser and Hawskesworth (1992) lies in the invalid assumption that incompatible elements are homogeneously distributed in the mantle source rocks.

Unfortunately, the type of geochemical modeling employed by Tainton and McKenzie (1994) is nondiscriminatory and may be applied with equal success to many other alkaline magmas if no other constraints are imposed. For example, their model could be used for such light REE-enriched rocks as kamafugites, minettes, carbonatites, or melilitites with similar predictable conclusions, and the genesis of these magmas could thus be incorporated into their general model. This undesirable conclusion arises from the failure of the geochemical models to consider other factors, such as major element compositions of the magmas, or mineralogical and experimental evidence pertaining to their genesis (Foley 1992b).

Finally, the hypothesis is not supported by the conclusions of this work that kimberlites and orangeites are distinct magma types derived from mineralogically different sources; orangeite bulk compositions do not, even for incompatible elements, reflect those of primary melts as a consequence of high- and low-pressure crystal-liquid fractionation.

A merit of the Tainton and McKenzie (1994) hypothesis is its focus on the Archean events leading to the large-scale depletion of continental lithospheric mantle in incompatible elements and the generation of a substrate for the subsequent addition of metasomatic (sensu lato) material.

Tainton and McKenzie (1994) do not discuss in any detail the nature of the partial melting processes leading to the formation of orangeite, kimberlite, and lamproite melts. However, they do indicate that all magmas are formed in the continental lithosphere above the mechanical boundary layer and within the garnet stability field. Partial melting is believed to take place over a depth range of 125-175 km; hence, melts may be produced in either the diamond or graphite stability fields.

Tainton (1992) suggests that the metasomatically-enriched sources of orangeite are held in the lithospheric mantle at temperatures close to their solidii. The small increase in temperature required to initiate melting is provided by advection of heat from the fringes of the large asthenospheric mantle plume responsible for the opening of the South Atlantic. Subsequent segregation and ascent of the melt follows the small-volume melt-movement model of McKenzie (1989).

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4.2. EXPERIMENTAL EVIDENCE PERTAINING TO ORANGEITE PETROGENESIS

CHAPTER 4

There have been no experimental investigations of the phase relationships of orangeite compositions at low pressure and only one at high pressure. Chapters 2 and 3 emphasize the mineralogical and geochemical similarities of orangeites to lamproites; thus, some of the experimental studies directed toward understanding lamproite petrogenesis have a direct bearing on the origin of orangeites, because of their similar ultrapotassic peralkaline character and proposed derivation from similar sources.

4.2.1. Liquidus Experiments on Orangeite Compositions

To date, the only melting experiments on orangeite compositions have been undertaken by Yamashita et al. (1992) and Arima et al. (1993a), using as starting material macrocrystal and aphanitic orangeites from Makganyene. The former is, not surprisingly, greatly enriched in MgO (30.40 wt %) relative to the latter (18.44 wt %). Makganyene orangeites are evolved, so the samples used in the experiments are high in silica (39.27 and 40.90 wt %), relative to unevolved orangeites (see Tables 3.1 and 3.2).

For the aphanitic orangeite, the liquidus temperature is about 1470°C at 6 GPa and 1520°C at 8 GPa. Suprasolidus phase assemblages vary with increasing pressure as follows: from phlogopite plus liquid through phlogopite plus clinopyroxene and liquid to phlogopite, clinopyroxene, orthopyroxene, and liquid to clinopyroxene plus garnet and liquid at 1400°C. Phlogopite breakdown occurs between 1300-1400°C and 6-7 GPa by the reaction of phlogopite with clinopyroxene to form garnet and liquid. No stable K-bearing phase is observed in runs above 7 GPa, suggesting that the liquid at these pressures is extremely rich in K and volatiles.

In the experiments on the macrocrystal orangeite, olivine is stable near-liquidus phase up to 8 GPa. It coexists with phlogopite and clinopyroxene below 4 GPa and with clinopyroxene, orthopyroxene, and garnet above 6 GPa at 1400°C. The persistence of olivine as a liquidus phase to high pressures is not surprising, given the enrichment of this sample in MgO due to the presence of macrocrystal olivine.

Arima et al. (l993a) consider that the aphanitic orangeite represents a liquid composition and propose, on the basis of their experiments, that orangeites can be generated by the partial melting of phlogopite-bearing lithosphere (sic) at pressures below 6.5 GPa. As the liquids are enriched in K, those generated at higher pressures might migrate toward the lithosphere-asthenosphere boundary and cause metasomatism, including phlogopite formation. The near-liquidus assemblage of clinopyroxene and garnet is interpreted to suggest that orangeites can be equilibrated, at pressures above 6.5 GPa, with an "eclogitic source."

The value of these experiments is questionable, given the very low probability that any bulk rock composition can represent the composition of the parental orangeite magma (see 3.3); moreover the compositions investigated might not even be primary. However, the phase relationships for the aphanitic sample do indicate that olivine is unlikely to be a stable suprasolidus phase. This observation is in agreement with experimental studies of other uItrapotassic compositions (see 4.2.3), suggesting that these magmas, at their sources, cannot be in equilibrium with olivine-bearing mantle (Foley 1992a). Without


any doubt, the macrocrystal orangeite investigated is a hybrid rock containing mantlederived olivine xenocrysts; thus, the liquidus relationships have no direct bearing on the genesis of orangeite.

4.2.2. Liquidus Experiments on Lamproite Compositions

These studies are important because of the similarities in mineralogy and composition of lamproites to orangeites. However, it should be realized that lamproites are rich in F, H20, and K20 relative to orangeites, and these differences in volatile compositions might have significant effects on liquidus phase relations (see below). As it is beyond the scope of this work to summarize the many studies of synthetic and naturallamproites, the reader is referred to Foley (1990, 1992a) and Mitchell and Bergman (1991) for comprehensive reviews of the topic. The study by Luth (1967) of the system KAISi04-Mg2Si04-Si02-H20 is particularly relevant to the postemplacement crystallization of C02-poor evolved orangeites (see 4.5.4). This system is also important with regard to crystallization of ultrapotassic magmas at high pressures and has been investigated at 20-28 kbpressure, with H20, C02, or F as the volatile phase, by Sekine and Wyllie (1982), Foley et al. (1986), and Gupta and Green (1988).

The most important conclusions of the studies of synthetic systems and natural phlogopite and leucite lamproite compositions, relevant to orangeite crystallization at low and high pressure, are:

• Primary olivine is a low- or high-pressure, high-temperature phenocryst. Reaction of primary or xenocrystal olivine with the liquid to form phlogopite may occur during post-emplacement crystallization.

• Phlogopite is stable at high or low pressures and may crystallize as a phenocryst during ascent of the magma. Under some conditions early-formed phlogopite may react with liquid to form pyroxene and leucite.

• Orthopyroxene may be a high-pressure phenocryst. • Leucite, sanidine, and amphibole are low-pressure phases. • The near-liquidus phase assemblage of lamproite melts suggests they may be

derived from phlogopite harzhurgite or phlogopite wherlite sources, under H20-or F-rich conditions.

However, other interpretations of these data are possible (Foley 1992b, see 4.5.2.2). Foley (1992a), summarizing experimental data for lamproites and kamafugites, has

stressed that models for the origins of these magmas by partial melting of phlogopite lherzolite are inconsistent with the results of liquidus experiments on ultrapotassic compositions. This conclusion is based on the absence of olivine, and in many instances garnet, as a near-liquidus phase. Consequently, non-peridotitic assemblages, rich in mica and pyroxene, which in some instances may be completely free of olivine (i.e., mica clinopyroxenite, mica websterite) are considered to be the sources of potassic rocks.

LampJ:oites are currently believed to originate from a depleted source which was strongly re-enriched in incompatible elements at a later stage, thus producing mica harzburgites, and subsequently melted under H20-rich reducing conditions (Foley 1988, 1989, Mitchell and Bergman 1991). Note that this three-stage model is similar in some

312 CHAPTER 4

respects to Tainton and McKenzie's (1994) model for orangeite genesis (4.1.3), the major difference being in the composition of the second-stage source.

Foley (l992a) has reviewed the effects of volatile composition on liquidus phase relationships and noted that, for a leucite lamproite at 2 GPa and 1150°C, decreasing H20/C02 ratios lead, to the addition of orthopyroxene, to mica plus clinopyroxene assemblages. For the same bulk composition in a C02-free system, with a CIWlhO ratio of 0.22, mica is the only liquidus phase. These data suggest that oxidation and addition of C02 result in clinopyroxene formation. In the case of kamafugites, low values of H20/C02 lead to the formation of clinopyroxene, whereas high values result in orthopyroxene plus clinopyroxene and reduced olivine stability. These data demonstrate that volatile compositions play an important role in determining the liquidus-phase assemblages of ultrapotassic magmas. Hence, it is suggested that one possible reason for the existence of diopside orangeites and diopside-free orangeites may be simply one of crystallization under differing H20/C02 ratios.

Olivine lamproites are commonly believed to represent primary magmas (Jaques et al. 1986, Foley 1993), although Mitchell and Bergman (1991) have argued on mineralogical and geochemical grounds that they are hybrid derivative magmas. These rocks have some compositional and mineralogical similarities to evolved orangeites (see 3.3.5). Regardless of their origin, liquidus experimental studies by Foley (1993) are important in demonstrating that olivine, orthopyroxene, and phlogopite occur together at the liquidus at pressures in excess of 5 GPa for magmas of this composition. These data are consistent with derivation by partial melting of phlogopite harzburgite under C02-free reducing conditions. To explain the apparent contradiction to Foley's (1992a) conclusion that the source rocks of ultrapotassic magmas should not contain olivine, Foley (1993) notes that the phase relationships are also consistent with melting of a veined source in which neither vein nor wall rock consists ofphlogopite harzburgite (see 4.5.2.2). Thus, Foley (1993) prefers an interpretation in which the magma results from preferential melting of phlogopite-amphibole-clinopyroxene veins within a garnet lherzolite host. The latter, to a lesser degree, is also involved in the melting process. The liquidus-phase assemblage found in the experiments is interpreted to be that of liquids in equilibrium with a residuum of melted vein and wall rock. Such a vein-plus-wall rock melting process (Foley 1993, 1992b) is directly applicable to the genesis of orangeites and is discussed further in Section 4.5.2.2. Alternatively, the persistence of olivine to high pressures in the liquidus experiments may be a consequence of investigating a hybrid magma composition. Contamination of a MgO-poor magma at its source with xenocrystal olivine will undoubtedly move the bulk composition of the hybrid into an olivine stability field.

4.2.3. Melting of Mica Pyroxenites

Lloyd et al. (1985) determined the compositions of liquids produced by the partial melting of a phlogopite pyroxenite (Phh5Cpx52Ml7Sph4Ap2) at 2 and 3 GPa. Liquids produced by 20-30% partial melting range widely in composition, but are K20 rich (4-6 wt %), Ti02rich (3-8wt %),andpoorinAh03 (6-9wt %) and Si02 (35-39 wt %).Lloyd et al. (1985) consider that the compositions of the melts support the hypothesis that the

PETROGENESIS OFORANGElTES AND KIMBERLITES 313

potassic rocks of southwest Uganda might be formed by partial melting of phlogopiteclinopyroxene sources.

Importantly, the study supports the hypothesis of Foley (1992a) that partial melts of olivine-free ultramafic rocks are ultrapotassic and compositionally different from those produced by partial melting oflherzolitic sources. Obviously, by varying the composition of the starting material and melting conditions, partial melts equivalent to many ultrapotassic magmas may be generated. Note the large degree ofmeIting in Lloyd et al.'s (1985) study eliminates the extraction problems perceived to be associated with small degrees of melting of lherzolitic sources.

Foley (1992a), on the basis of Lloyd et al. 's (1985) experiments, draws the important conclusion that the mg numbers of primary magmas derived from pyroxenitic sources may be less (60) than the values of 65-75, which are generally taken to indicate a primary mantle-derived melt. Similarly, high Ni and Cr (200-500 ppm) contents do not necessarily

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10 20 30 40 50 60

- wt. % APATITE ----.

Figure 4.1. (A) Phase relations in the system hydroxy-phlogopite-hydroxy-apatite at 2 GPa. (B) Phase relations in the system fluor-phlogopite-fluor-apatite at 2 GPa. Fo = forsterite, Ap = apatite, PH = phlogopite, L = liquid (after Vukadinovic and Edgar 1993),

314 CHAPTER 4

imply a primitive character, as these may result primarily from contamination with peridotitic material (see 3.7). Foley (1992a) emphasizes that criteria used to identify primary basaltic magmas derived from Iherzolitic sources, may be misleading for exotic alkaline magmas.

4.2.4. Phase Relations in the System: Phlogopite-Potassium Richterite-Apatite

Vukadinovic and Edgar (1993) have investigated the pseudobinary systems: hydroxy-phlogopite-hydroxy-apatite and hydroxyfluor-phlogopite-hydroxyfluorphlogopite at 2 GPa pressure. In the F-free system, apatite, phlogopite, and forsterite coexist with a K-P-rich liquid at a pseudobinary minimum at 1225°C and PhlssAplS. In theF-bearing system this minimum occurs at 1260° at Ph166Ap34, and liquids are relatively poorer in K and richer in P. For phlogopite-rich compositions, forsterite and phlogopite coexist with liquid over a wide range of temperatures above those of these minima (Figure 4.1) due to the incongruent melting behavior of phlogopite.

Figure 4.2 illustrates pseudobinary phase relationships in the systems: apatite-fluorpotassium richterite and fluor-phlogopite-fluor-potassiumrichterite (Edgar and Pizzolato 1995). Both systems contain pseudobinary minima with K-rich liquids being present above 1150°C and I OOO°C, respectively. Orthopyroxene appears as a supraliquidus phase due to the breakdown of richterite. The F-free analogues ofthe richterite-bearing systems were not investigated by Edgar and Pizzolato (1995), but the data of Vukadinovic and Edgar (1993) suggest that minimum melting temperatures will be lower in these systems.

These data show that the presence of potassium richterite lowers the melting temperature more effectively than either phlogopite or apatite. It is to be expected that

t 1500

- 1400 U o -r- 1300

®

LIQUID

+

LIQUID

LIQUID +

ORTHO-PYROXENE

+ 1200 APATITE K-RICHTERITE

1100 +-------.l..-------i APATITE + K- RICHTERITE

50 60 70 80 90

- wt. % K - RICHTERITE ~

® L + PH + OPX 7 + OL?

1300

1200

1100

1000 PH + K-RICH +OPX + L

900 PH + K-RICH

20 40 60 80

- Wt. % PHLOGOPITE"

Figure 4.2. (A) Phase relationships in the system fluor-apatite-fluor-potassium richterite at 2 GPa. (B) Phase relationships in the system fluor-phlogopite-fluor-potassium richterite at 2 GPa. K-RICH = potassium richterite, OL = olivine, OPX = orthopyroxene, PHL = phlogopite, L = liquid (after Edgar and Pizzolato 1995).


minimum melting temperatures in the F-free pseudoternary system phlogopite-potassium richterite-apatite will be lower than lOoo°C at 2 GPa.

These studies are important with regard to the melting of mica-rich veins in the upper mantle if the hypotheses of Foley (1 992a,b ) are correct. The effects of increasing pressure or the addition of other components, i.e., diopside, K-Ba titanates, are as yet unknown. However, these data indicate that partial melting of mixtures of phlogopite, amphibole, and apatite will produce a wide range of K-rich liquids. Clearly, experimental investigation of these systems at high pressures is critical to advancing our understanding of the formation of potassic melts in the mantle.

The primary mineral assemblage of some orangeites consists principally of phlogopite, apatite, and carbonate. Their mineralogy can be in part explained if these rocks represent the products of partial melting of phlogopite (>85 wt % )-apatite veins in a dunitic or lherzolitic mantle (see 4.5.2).

4.3. PETROGENESIS OF ARCHETYPALKIMBERLITES-RECENT HYPOTHESES

Orangeite magmatism is geographically and temporally associated with archetypal kimberlite magmatism in southern Africa. To explain this relationship, some discussion of current hypotheses of kimberlite petrogenesis is required. Earlier hypotheses have been summarized by Mitchell (1986). Recent hypotheses have been concerned more with the site of partial melting rather than the details of the process.

4.3.1. Carbonated Lherzolite Sources

Recent discussion of the origin of kimberlite has centered around partial melting of carbonated lherzolite, since the discovery by Wyllie and Huang (1 975a,b ) and Eggler (1974) that dolomite and magnesite can be stable in the mantle at solidus temperatures. The numerous experimental studies of natural and synthetic peridotite-C02-H20 have been reviewed by Mitchell (1986), Wyllie (1980, 1987, I 989a,b ), Eggler (1987), and Wyllie et al. (1990). For reasons noted below, most of these studies have been devoted to determining the topology of the peridotite-C02-H20 solidus at 1-4 GPa.

Figure 4.3 illustrates phase relationships in the system peridotite-C02. The effect of C02 on the solidus curve is to produce a significant reduction in the solidus temperature at pressures at which dolomite becomes stable. This occurs where the subsolidus univariant carbonation reaction (reaction 6 of Wyllie and Huang 1976)

2Mg2Si04 + CaMgSh06 + 2C02 = 4MgSi03 + MgCa(C03h fosterite diopside enstatite dolomite

intersects the solidus. At this intersection the invariant point 16 marks the change from the solidus for peridotite-C02 to the solidus for dolomite peridotite. At higher pressures a second subsolidus univariant reaction (/0) results in the appearance of magnesite in place of dolomite (Figure 4.3). Melt compositions at the solidus have been inferred to be C02-rich undersaturated silicate melts or carbonatites at pressures greater than 16, but relatively silica rich (basanite, alkali basalt) at lower pressures. Thus, 16 is a petrogeneti-

316

--~ 30 ::J f/) f/) LLI 20 a:: a..

10

OL+OPX + CPX +GNT + MAG

@

800 1000 1200 1400

- TEMPERATURE (Oe) .....

CHAPTER 4

A Magnesite garnet lherzolite B Dolomite garnet lherzolite C Dolomite spinel lherzolite D Garnet lherzolite + C02 E Spinel lherzolite + C02 F Plagioclase lherzolite

FIELD OF LHERZOLITE + VAPOR PHASE

VAPOR ABSENT SUBSOLIDUS ASSEMBLAGE

Figure 4.3. Phase relationships in the system peridotite-C02 (0.1 %) based on experimental data and deductions from the system CaO-MgO-Si02-C02 (Wyl1ie and Huang 1976). Note that the carbonation reaction terminating at 16 on the peridotite solidus divides the diagram into vapor-absent and vapor-bearing assemblages. as all C02 is consumed in the carbonation reaction. Solidus for anhydrous peridotite (PER) is shown as a dashed line. OL = olivine. OPX = orthopyroxene. CPX = clinopyroxene. GNT = garnet, MAG = magnesite. DOL = dolomite. SP = spinel. PL = plagioclase. L = liquid. V = vapor.

cally important point as it marks the depth below which C02 plays a key role in generating primary carbonatitic to low-silica C02-rich alkaline magmas such as melilitite and kimberlite. Consequently, much effort has been devoted to locating the position of 16. The initial experimental studies undertaken in separate laboratories were not in agreement on the exact P-T of 16, in part due to the different experimental techniques utilized. Recent studies are more in agreement, although it is realized that the location also depends on other factors such as oxygen fugacity (Wyllie et al. 1990).

It is important to recognize that the location of 16 also varies with C02ilhO ratio. The solidus depression is at a maximum in the H20-free system and the carbonationrelated minimum is, of course, absent in C02-free systems. Further, complexities in the topology of the solidus of peridotite-H20 and peridotite-C02-H20 are introduced by the stabilization of amphibole (see Wyllie 1987, Falloon and Green 1989, Wyllie et al. 1990).

Fortunately, for the purposes of discussion of the genesis of kimberlite, the exact position of 16 and topology of the peridotite-C02-H20 solidus is not critical. What is of significance is the presence of the minimum and the associated "solidus ledge" (Figure


4.3) which is created at lower pressures. The solidus ledge is believed by Wyllie (1980) to play a key role in arresting the ascent of melts derived from the asthenosphere. Crystallization of melts at this ledge results in the release of vapors, which Wyllie (1980) sees as the principal agent of metasomatism in the upper lithospheric mantle.

Hypotheses of kimberlite genesis are based upon the existence of the solidus ledge (Wyllie 1980, Skinner 1989, Tainton 1992) and the observation that partial melts of carbonated lherzolite are C02-rich (Wyllie 1980, Brey et al. 1983, Canil and Scarfe 1990). Most hypotheses based upon peridotite-C02 equilibria are models of the genesis of "haplokimberlites," i.e., K-, Ti-, Fe-free melts approximating to postulated primitive magma compositions in terms of their Si02, CaO, MgO, and volatile contents. Phlogopite and other incompatible element-rich phases are not present in the experiments. It is assumed that their presence will not seriously change the melting relationships, and incompatible elements held in these phases will enter the melt along with carbonate. Determination of the actual composition of the small-volume near-solidus melts produced in these experiments is extremely difficult, as liquids quench to cryptocrystalline intergrowths of minerals rather than glass. Compositions, estimated by a variety of indirect methods, are all C02 rich and undersaturated and believed to range from carbonatite to kimberlite.

4.3.1.1. Volatile Fluxing-Diapiric Model

Wyllie (1980) presented a model for the genesis and ascent of kimberlite magmas based on volatile-induced partial melting of lherzolite. Revisions to the model have been presented by Wyllie (1989a,b) and Wyllie et al. (1990).

The Wyllie (1980) version of the model assumes that C02, H20, and the components of phlogopite are present as the constituents of trace magmas within a persistent melt zone extending from 120 to 260 km depth. This zone of melting results from the intersection of the shield geotherm, over this depth range, with the solidus of magnesite peridotite. In the presence of volatiles above 120 km the mantIe may be composed of hornblende or phlogopite dolomite peridotite and vapor, and below 260 km, of magnesite peridotite and H20-rich vapor. In the absence of volatiles the mantle would be composed of garnet lherzolite.

The starting point in Wyllie's (1980) model is the introduction of C02 and H20 to volatile-free mantle at depths below 260 km. These volatiles, of unknown origin, are considered to escape from deep within the asthenosphere. Partial melting of the asthenospheric mantle is induced at point A on the carbonated peridotite solidus (Figure 4.4). Melting results in a density inversion, followed by the uprise of a partially melted diapir along the adiabiatic path A-B (Wy lIie 1980) or along the geotherm A-C (Wy Ilie 1989a,b). Within each diapir the percentage of melting increases during ascent.

At point B (Figure 4.4) the rising diapir intersects the thermal maximum on the carbonated lherzolite solidus. If equilibrium is maintained, the magma will evolve vapor, and crystallize completely as phlogopite dolomite peridotite plus vapor. The hot solid diapir may continue to rise and will, upon recrossing the solidus at lower pressures, melt to a C02-poor liquid, possibly of alkali basaltic-nephelinitic composition.

Alternatively, if the lithosphere is under tension, the evolution of vapor at B (Figure 4.4) could result in crack formation and propagation into the overlying lithosphere, as

318 CHAPl'ER4

Explosive eruption Gas CO2-H2O

A 50

Magma crystallize

E 100 B i.' ~ - 150 G ~ -a. Q) 200 0

250 Rising

300

500 1000 1500

B 50

-E 100 ~ - 150 Graphite ~ Diamond - E a. Q) 200

0

250 Rising

300

500 1000 1500

Temperature (OC)

Figure4.4. Diapiric models for the origin and eruption of kimberlite magmas (after Wyllie 1980). (A) Partially melted diapirs rise through the mantle along the adiabatic path A-B and crystallize at the thermal maximum on the carbonated peridotite solidus. Crystallization is accompanied by evolution of C02 vapor which initiates crack propagation through the overlying lithosphere. followed by explosive eruption of diamond-free kimberlite magma from point B at about 90-km depth. (B) The system as it appears once a conduit as been established to the surface. Magmas separating from partially melted diapirs do not crystallize at the thermal maximum and are derived from progressively deeper levels in the mantle. Eventually magmas segregate from within the diamond stability field at point E and give rise to diamond-bearing kimberlites. See text for further explanation.

PETROGENESIS OF ORANGElTES AND KIMBERLITES 319

suggested by Anderson (1979). Kimberlites would erupt explosively as vesiculating partially-crystallized melts, from depths of about 90 km. The initial batches of magma erupted from 90 km will not contain diamond as they have equilibrated at depths above the stability field of diamond. Wyllie (1980) believes that once a conduit has been established by rising diapirs, further escape of volatiles will trigger diapiric ascent from successively deeper levels and, ultimately, from within the diamond stability field (Figure 4.4). Thus, Wyllie (1980) suggests that successive individual intrusions within a single kimberlite body should contain xenoliths which record increasing depths of origin.

In this model, the maximum depth of kimberlite generation must be the deepest point at which the geotherm intersects the carbonated peridotite solidus. However, batches of magma can ascend from any higher level above this depth once a conduit has been established. These magmas will ascend rapidly and will not be affected by the solidus shelf. Kimberlite magmatism in this model is analogous to a stream of bubbles rising from an ever-deepening source (Figure 4.4).

Other batches of magma equilibrate along geotherms as they ascend and neverreach the solidus ledge. These magmas will crystallize at the solidus at point F (Figure 4.4), and evolution of vapors will result in metasomatism of the adjacent mantle (Figure 4.4). In this scenario, evolution of vapor as a metasomatic agent is a consequence of kimberlite magmatism rather than a precursor. Wyllie (1989a,b), in the revised versions of his model, suggests such melts may accumulate at, and invade, the base of the lithosphere (Figure 4.5). Some of these kimberlites may escape via lithospheric cracks, whereas others remain trapped in the vicinity of the lithosphere-asthenosphere boundary (Figure 4.5). Partial melting of previously metasomatized lithosphere during subsequent thermal perturbations may generate kimberlite-like magmas, i.e., orangeites.

Though Wyllie's model has been widely accepted, it has several serious drawbacks discussed by Mitchell (1986) and Eggler (1989). The principal problems are with respect to the absence of a zone of persistent melt beneath continental shields, doubts whether rising magmas would intersect the solidus shelf and initiate crack propagation and conduit formation, and lack of geological evidence to support direct eruption of several batches of magma from deep in the mantle along a single conduit.

With regard to the latter point, WylIie's model completely ignores the known hypabyssal infrastructure of kimberlite fields which indicate that at deep crustal levels kimberlite magmas behave rheologically like many other basic magmas (Eggler 1989). Diatremes are developed above dike systems only at high levels and do not extend into the mantle. Skinner (1989) addresses these crustal emplacement problems in stage 4 of his model (see 4.4.4).

Wyllie's model is infinitely flexible with regard to the extent of the persistent melt zone. The magnitude of this may vary considerably, depending upon the slope of the carbonated mantle solidus and that of the local geotherm.

Wyllie (1980) assumes that magmatism is triggered by the ascent of volatiles into a mantle at equilibrium; if vapor is not present there is no melting. He does not discuss the alternative that a preexisting carbonated mantle diapir rising by convection from depth might cross the solidus and undergo decompressional melting.

Wyllie (1989a,b) introduced asthenospheric plumes into his model and noted that kimberlites may be generated ahead of the plume during the initial stages of melting

320

o

-E 100 .¥ -.s:::. -~ 200 o

300

CHAYfER4

CRUST

MOHO

",-

.b====i> __ . - ~e\\ ,,'-.. \....",. -- - -- ~t\eo

" r r €,t\\tO\

--- -~\,l((t------" 'I ' ,,(, Solidus - - - ......

H~fl .. , .. "/' 1.1 r:::::t I I Magma I.l:.:LI Vapor

Figure 4.5. Cross section of a craton showing lithospheric thinning resulting from upwelling of an asthenospheric plume. Note that partial melts are generated below the lithosphere-asthenosphere boundary. During the initial stages of upwelling small degrees of partial melting may give rise to kimberlitic melts which might escape through the lithosphere or remain trapped at its base. As the trapped magmas, or failed kimberlites, crystallize, they have the potential to metasomatize the adjacent lithosphere. Subsequently, further upweIling and rifting leads to the formation of nephelinitic magmas at shallow depths. See text for further explanation (after WyIlie 1989a,b).

(Figure 4.5). A consequence of the plume variant ofthe model is that initial emplacement of kimberlites is followed by lithospheric thinning, rifting, and generation of alkaline magmas above l00-km depth.

4.3.1.2. Partial Melting of Magnesite Peridotite

Brey et al. (1983), on the basis of extrapolation of liquidus experimental studies of olivine melilitite in the presence of C02 at 3-3.5 GPa, suggested that kimberlites might be generated by very small degrees of partial melting of a magnesite peridotite. Subsequent experimental studies of near-solidus phase relationships for synthetic peridotiteC02 systems to 12 GPa by Canit and Scarfe (1990) provided verification of the hypothesis.

The studies demonstrate that at all pressures from 4 to 11 GPa, the stable mantle assemblage for the system peridotite-{:02 is magnesite peridotite or magnesite garnet peridotite (Figure 4.6).

Canit and Scarfe (1990) noted that the composition, calculated by mass balance, of near-solidus partial melts (23-39%) in equilibrium with magnesite at 5-7 GPa, have lower Ca/Mg ratios than melts in equilibrium with dolomite at 3 GPa (Figure 4.6). They interpret the former to be kimberlitic and the latter melilititic in agreement with the conclusions

PETROGENESIS OFORANGEITES AND KIMBERLITES 321

CMS - CO2 CMAS -CO2

t 12 12

10 10

C a.. OL +OPX +CPX (!) OL OL+CPX - 8 + +OPX 8 +GNT

L&.J CPX +MAG+L + MAG

a:: + , OL

::::> MAG ,

+ en 6 I 6 en I. CPX OL+OPX

L&.J I + +CPX +GNT a:: I L +L a.. OL+CPX

4 +OPX 4 +L

2

1000 1200 1400 1600 1000 1200 1400 1600

TEMPERATURE (OC) ~

Figure 4.6. (Left) Phase relationships in the system CaO-MgO-Si02-C02. (Right) Phase relationships in the system CaO-MgO-AI203-Si02-C02. OL = olivine, CPX = clinopyroxene, OPX = orthopyroxene, GNT = gamet, MAG = magnesite, DOL = dolomite, L = liquid (after Canil and Scarfe 1990).

of Brey et al. (1983). Melts at 9 GPa are too rich in Mg to correspond with any proposed primary kimberlite magma (Figure 4.7). Canil and Scarfe (1990) suggest that kimberlites are generated by partial melting of magnesite-bearing peridotite at pressures of 5-7 GPa (150-250 km depth) in the upper mantle. Kimberlite cannot be generated at pressures below 3 GPa.

Melting and magma segregation, followed by ascent of kimberlite magma, commences when rising asthenospheric diapirs of magnesite peridotite become trapped at the lithosphere-asthenosphere boundary. In their model, Canil and Scarfe (1990) do not discuss Wyllie's (1980, 1989a,b) models of kimberlite genesis and emplacement.

Canil and Scarfe (1990) also suggest that magmas termed "protokimberlites" could originate at pressures above 7 GPa. Such magmas may, for tectonic reasons, remain trapped in the lower mantle. Alternatively, fractionation of olivine at high pressure may drive their bulk compositions toward those of the 5-7 GPa melts prior to eruption as kimberlites.

Clinopyroxenes crystallizing in equilibrium with melts become more subcalcic with increasing pressures. Their compositions are similar to those of subcalcic pyroxenes belonging to the megacryst suite. Thus, Canil and Scarfe (1990) suggest that the latter may represent high-pressure phenocrysts in kimberlite magmas.

Canil and Scarfe's (1990) work provides important experimental evidence that haplokimberlitic magmas may be generated by partial melting of carbonated lherwlitic sources at high pressures, consequent upon decompressional melting.

322

PARTIAL MELTS

o 5 GPo () 7 GPo • 9 GPo

CHAPTER 4

M

Cc '-------------~~----------------------~ C Ln s

Figure 4.7. Compositions of "uncontaminated" natural kimberlites. lamproites. melilitites. and partial melts formed in the system CaO-MgO-Si02-C02 projected into the ternary system CaO-MgO-Si02 from A1203. The arrows for the partial melts indicate their changing compositions with respect to increasing degrees of melting (after Canil and Scarfe 1990).

Canil and Scarfe (1990) also demonstrate that lamproites are unlikely to represent partial melts of carbonated lherzolite as their bulk compositions are richer in silica than those of haplokimberlitic liquids encountered in their study (Figure 4.7).

4.3.1.3. Partial Melting o/Carbonated Phlogopite Lherzolite Based on studies of the melting of carbonated phlogopite lherzolite, Thibault et al.

(1992) have proposed a model of cyclic metasomatism occurring beneath continental rifts. This model has similarities with one presented by Haggerty (1989a,b) in proposing the existence of distinct C02- or H20-rich metasomatically enriched horizons in the lithosphere (see 4.4.2). A major difference between the models is that Thibault et al. (1992) propose a dynamic model of metasomatism, whereas that of Haggerty is essentially static.

Melting of a carbonated phlogopite lherzolite at 3 GPa and 1060-1150°C results in the formation of alkaline dolomitic melts. The composition (wt %) of the 4% partial melt at 1l00°C is 2.57 Si02, 4.54 FeO, 15.12 MgO, 21.60 CaO, 4.93 Na20, 7.01 K20,2.66 H20, 40.31 C02. Gamet, orthopyroxene, and olivine were found in equilibrium with the melts at all temperatures, whereas phlogopite was consumed between 1125 and 1150°C.

Melting of a phlogopite lherzolite at 3 GPa from 1175 to 1250°C results in the formation of hydrous potassic calcic melts. At 1225°C, the composition (wt %) of a 7% partial melt is 47.96 Si02, 10.97 Ah03, 5.66 FeO, 13.21 MgO, 11.59 CaO, 1.04 Na20, 5.40 K20, and 2.05 H20.

Thibault et al. (1992) further showed that the alkaline dolomitic melts can metasomatize harzhurgite to olivine-rich phlogopite wherlite and calcite-bearing phlogopite


dunite. The interaction of harzburgite with the hydrous silicate melt results in enrichment in clinopyroxene and phlogopite.

On the basis of these data Thibault et al. (1992) propose a cyclic model of metasomatism active beneath a continental rift. The cycle commences with the formation of a horizon of carbonated phlogopite lherzolite at the base of the lithosphere by interaction of barren harzburgite with dense alkaline fluids derived from a hot mantle plume. Lithospheric thinning associated with plume uprise results in the uplift of the lithosphereasthenosphere boundary and melting of the carbonated phlogopite lherzolite. The lowtemperature melting components are then remobilized as small-volume volatile-rich melts which, as they migrate upward, resolidify and react with the overlying colder harzburgite. Through such cycles of melting, migration, and solidification the metasomatic front of carbonated phlogopite lherzolites eventually reaches a depth of 100 km (3 GPa).

Subsequent increase in the temperature of dolomite phlogopite lherzolite to 1150°C results in the formation of an alkaline dolomitic melt which infiltrates overlying lithosphere leaving a residual C02-free garnet phlogopite lherzolite which returns to subsolidus conditions. As the alkaline dolomitic melt reaches 65 km (2 GPa, 950°C) it solidifies and reacts with harzburgite to produce dolomite phlogopite harzburgite. The metasomatic front is now decoupled and forms two distinct horizons: a dolomite phlogopite harzburgite at 65 km and a C02-free garnet phlogopite lherzolite at 100 km.

With progressive rifting, the temperature at 65 km depth reaches lOOO°C and the dolomitic phlogopite harzburgite is transformed to olivine-rich phlogopite lherzolite and wehrlite with the release of C02-rich low-density fluids. At l00-km depth, temperatures exceed the solidus of gamet phlogopite lherzolite. Partial melts from this source infiltrate the metasome at 65 km resulting in an enrichment in phlogopite and clinopyroxene. The final result of shallow decoupled metasomatism is a wide variety of metasomatized ultramafic rocks comparable to the xenolith suites found in alkaline basaltic rocks.

Importantly, the studies indicate that carbonate-bearing and potassic magmas can be generated in the lithosphere and that carbonated phlogopite lherzolite has a very low solidus temperature at 3 GPa. Thibault et al. (1992) do not discuss the possibility of these potassic melts escaping from the lithosphere. Interestingly, the low solidus temperature suggests that melting of such sources at the fringes of large thermal plumes, in nonrifting environments, might give rise to melts which would subsequently solidify in the lithosphere as veins of complex mineralogy. The study demonstrates the complexity of metasomatic processes in the lithosphere and is thus relevant to the formation of metasomatic and veined non-Iherzolitic sources for ultrapotassic rocks.

4.3.1.4. Carbonates in the Mantle?

The partial melting models of carbonated lherzolites described above are currently the favored mechanism for generating asthenospheric kimberlites and other C02-rich alkaline magmas. Despite the popularity of such sources it should be realized that they are hypothetical concepts based entirely on experimental studies. Samples of dolomiteor magnesite lherzolite have never been reported from the xenolith suites of kimberlites. The report of the presence of a few calcite and dolomite inclusions in pyrope xenocrysts

324 CHAPTER 4

from minette diatremes in Utah (McGetchin et al. 1973, McGetchin and Besancon 1973) is hardly overwhelming evidence in favor of widespread carbonation of the upper mantIe. There is no doubt that C02 exists in the mantle and that the experimentally verified carbonation reactions must occur. Consequently, the absence of carbonate-bearing xenoliths is ascribed to either the decomposition of all carbonate during transport or failure to recognize, petrographically, small volumes of primary carbonate in altered xenoliths. Conveniently, carbonate decomposition is always considered to result in complete disaggregation of xenoliths (Wyllie et al. 1983); thus, any evidence pertaining to its former presence is always destroyed.

Canil (1990) has attempted to resolve the problem by subjecting a synthetic carbonated mantIe to various rates of isothermal decompression. A mixture of dolomite, enstatite, and diopside, held initially at 2.5 GPa and l000°C, was shown to decompose rapidly into forsterite, diopside, enstatite, and vapor at decompression rates of 1-3 GPa/hr (45-90 kmlhr). Decarbonated, decompressed experimental charges were found to be disaggregated and granular. Fluid inclusions present in newly created olivine demonstrate that fluid (C02-vapor) was present during decarbonation. Canil (1990) concludes that carbonate coexisting with silicates in mantIe-derived xenoliths could not survive entrainment, even in the fastest ascending magmas, due to rapid decarbonation upon decompression. Interestingly, decarbonation reactions may be one means of producing some macrocrystal olivines.

Another approach to the problem is to seek textural evidence for decarbonation reactions among the mineral constituents of xenoliths, but xenoliths have not been extensively studied with this object in mind. Textural relationships between the presumed products of decarbonation of dolomite, i.e., olivine and diopside, might be examined to some avail in fresh xenoliths.

Berg (1986) has suggested that rare brucite-calcite intergrowths in mantle xenoliths form when dolomite breaks down to calcite, periclase, and vapor. Periclase, being unstable in a hydrous environment, is subsequently converted to brucite. Canil's (1990) experiments do not support this reaction as being of importance in dolomite breakdown relative to carbonate-silicate equilibria. A further problem with Berg's (1986) hypothesis is the difficulty in distinguishing, on a petrographic basis, between periclase-derived brucite and that formed by serpentinization (Berg 1989).

The apparent absence of mantle carbonates in xenolith suites can also be explained if the lithospheric mantle is regarded as a barren residua purged of its volatiles by previous melting episodes (Hess 1989). Hence, the lithospheric xenolith suite sampled by kimberlites cannot be expected to contain carbonates. Any C02 entering this environment results in partial melting rather than carbonate formation, due to the depression of the mantle solidus upon carbonation (Wyllie 1980).

This conclusion does not negate the possible existence of solid carbonates at greater depth in the asthenosphere. Pristine samples of this material are unlikely ever to be present in xenolith suites because, even during slow diapiric upwelling, decompressional partial melting will inevitably occur and remove all traces of carbonate. Small modal amounts of carbonate will be consumed in the initial stages of melting. Separation of melt along grain boundaries, followed by segregation from the restite, may result in melts that migrate relatively rapidly toward the lithosphere, leaving a restite containing no trace of


carbonate. Restites are unlikely ever to be entrained by their daughter melts, although long-term mantle convection may transport restite to higher levels, where it may be sampled by subsequent batches of magma. However, such material, because of continued reequilibration, will contain no traces of its former character or depth of origin.

On the basis of the existing experimental data, melting of carbonated lherzolite appears to be suitable for the production of C02-rich archetypal kimberlites but not relatively C02-poor magmas such as orangeites or lamproites. However, detailed experimental studies on H20- and alkali-bearing systems which might confirm or deny this conclusion have not yet been undertaken. Ifindeed orangeites (and lamproites) are formed in the lithosphere, then their origins might be clarified by further experimental studies of non-Iherzolitic sources (see 4.2) with mixed volatiles.

4.3.2. Liquidus Experimental Studies at High Pressures

Inferences concerning the nature of the sources of magma, based on liquidus experimental studies of kimberlite compositions, are severely hampered by our lack of knowledge of the composition of primitive kimberlite magmas (Mitchell 1986). Further, Mitchell (1986) has concluded that kimberlite whole rock compositions do not represent those of the magmas from which they formed, as a consequence of volatile loss and crystal fractionation during differentiation. Regarding volatile loss, Brey et al. (1991) have noted that the solubility of C02 in kimberlitic melts may change dramatically at 4-5 GPa, and large amounts of C02 may be released from melts in this pressure range during ascent into the crust. Canil (1993) has observed that melting experiments on compositions with such reduced C02 contents may produce points of multiphase saturation which are not representative of the depth of origin of the magma or residual mineralogy of the source. Macrocrystal kimberlites are particularly unsuitable material for experimental studies as they are clearly contaminated rocks.

The few liquidus experimental studies that have been undertaken have utilized either natural kimberlites (Edgar et al. 1988, Edgar and Charbonneau 1993) Or cobalt-substituted synthetic analogues (Eggler and Wendlandt 1978, Ringwood et al. 1992, Kesson et al. 1994).

4.3.2.1. Liquidus Studies of Natural Kimberlite

Recognizing the above problems of sample selection, Edgar et al. (1988) and Edgar and Charbonneau (1993) have suggested that liquidus studies of an aphanitic kimberlite from the Wesselton Mine (South Africa) are meaningful as this material has a composition close to those of uncontaminated kimberlite (Mitchell 1986). Edgar and Charbonneau (1993) suggest that the absence ofxenocrystal olivine, high mg number (83.9), Ni (810 ppm), and Cr (2410 ppm) contents indicates that the parental magma of this kimberlite has not undergone appreciable fractionation. In contrast, Mitchell (1986) and Shee (1984), on mineralogical and petrographic grounds, regard the Wesselton aphanitic kimberlite as a differentiate of a less-evol ved kimberlite magma. Edgar and Charbonneau (1993) admit that some olivine and pyroxene fractionation may have occurred from the parental magma, but nevertheless state that, for the purposes of their experimental studies, "the actual origin of the aphanitic kimberlite is not as important as the similarity in its

326

4 A

\ OL " \ + \ \

5

-C 6 a. (!) 7 -lLJ 8 a::: ~ 9 en en lLJ 10 a::: a.

II

OL +

GNT + L

L

L

1300 1400 1500 1600 1700

10 -... C

..0 ~ _20

lLJ a::: ~

~ 30 lLJ a::: a.

t 40

8

o

l LIQUID A OL+L

CHAPTER 4

\ I I

\ B OL+SP+L I C OL + SP tcPX +L \ o OL+SP+CPX+CT \ E OL+SP+MO+L +LI

I F OL+SP+MO+CT+LI

I

1000 1200 1400

TEMPERATURE (OC) • Figure 4.8. Phase relationships of Wesselton aphanitic kimberlite at high (A) and low (B) pressures. OL = olivine. GNT = garnet, L = liquid. MO = monticellite. CPX = clinopyroxene. CT = calcite. SP = spinel (after Edgar and Charbonneau 1993. Edgar et al. 1988).

composition to that of other supposedly primitive kimberlites" (Edgar and Charbonneau 1993. p. 133). It is concluded here that, lacking samples of indisputably primitive composition, the Wessel ton aphanitic kimberlite remains the best available sample of a relatively unfractionated kimberlite.

Figure 4.8 illustrates the near-liquidus-phase-relationships of the Wesselton aphanitic kimberlite at 5-10 GPa, as determined by Edgar and Charbonneau (1993). The experiments were undertaken with only the volatiles present in the kimberlite (Xeo = 0.24), at an oxygen fugacity estimated as slightly above the EMOG-EMOD buffer (Eggler and Baker 1982), using presoaked Pt capsules to avoid Fe loss by alloying.

The data show that orthopyroxene and primary clinopyroxene are absent at nearliquidus phases, although quench clinopyroxene occurs in all fields. Extrapolation of the field boundaries suggests that the minimum pressure at which garnet and olivine might coexist is 12 GPa. Liquidus olivines have mg = 92.4-94.4 and are thus similar to olivine phenocrysts in the natural rock. Garnets are pyrope grossular almandines similar to those common in eclogites.

Edgar and Charbonneau (1993) interpret the data to suggest that the liquid could never have been in equilibrium with a garnet lherzolite source. At first sight, liquidusphase relations suggest garnet dunite as a potential source, although it is clearly too refractory. Consequently, Edgar and Charbonneau (1993) note that the ubiquitous presence of quench clinopyroxene up to near-liquidus temperatures is evidence that the source


material may contain clinopyroxene. The presence of olivine on the liquidus up to high pressure is believed to be a consequence of the expansion of the olivine plus liquid field, at the expense of clinopyroxene, due to C02 loss. Hence, at a higher C02 content a multi saturation point involving clinopyroxene, olivine, and garnet would be expected in experiments at lower pressures (10 GPa).

Experiments undertaken on the same starting composition at 1-4 GPa (Figure 4.8) also suggest that orthopyroxene is absent from the source (Edgar et al. 1988). The absence of Mg-carbonates at low pressure or as quench phases at high pressure is interpreted by Edgar and Charbonneau (1993) to indicate that kimberlites are not derived from magnesitebearing sources. In this context, Foley (1990) notes that calcite is a possible mineral in the source, as the reactions which limit the formation of magnesite all involve orthopyroxene.

Clinopyroxene is an abundant near-liquidus phase in low-pressure experiments at 1.5-3 GPa. To explain the absence of primary low-pressure clinopyroxene phenocrysts in kimberlites, Edgar et al. (1988) suggest the magma temperatures remained high (1250-1300°C) during rapid ascent from about 150-km depth. Under these constraints, olivine and spinel will be the only liquidus phases during ascent until they are joined by monticellite at pressures of less than 1 GPa.

In summary, Edgar and Charbonneau (1993) suggest that kimberlites are generated by the partial melting of a gametite (olivine--clinopyroxene-garnet) source at 10-13 GPa (300-330 km), which may represent deeply subducted material of basaltic composition.

4.3.2.2. Liquidus Studies of Synthetic Kimberlite

Liquidus studies of synthetic kimberlites have been undertaken by Eggler and Wendlandt (1978, 1979), Ringwood et al. (1992), and Kesson et al. (1994). All experiments use CoO as a proxy for FeO to eliminate alloying with the Pt capsules used to contain the samples in the experiments. Co-substitution results in liquidus temperatures about 50°C higher than in the analogous Fe-bearing systems. The results of, and conclusions derived from, these experiments must be regarded as preliminary as detailed investigations of the phase relationships were not undertaken.

4.3.2.2.a. Average Lesotho Kimberlite. Eggler and Wendlandt (1978, 1979) studied a bulk composition equivalent to that of the average anhydrous Lesotho kimberlite (Gurney and Ebrahim 1973) at 3 and 5.5 GPa with various C02IH20 ratios. In considering the significance of their results, it should be borne in mind that Mitchell (1986 p.279) has noted that this average composition includes macrocrystal, altered, and contaminated kimberlites and is biased toward relatively high silica and alumina contents. Thus, the composition is unlikely to represent primitive kimberlite magma.

Figure 4.9 shows this composition has olivine, orthopyroxene, and clinopyroxene near the liquidus at both pressures studied. Extrapolation of the gamet-in curve indicates that all four phases occur near the liquidus at 6 GPa. Dolomite and magnesite were found at temperatures 300°C below the liquidus. Eggler and Wendlandt (1979) intetpret the data to indicate that kimberlite could be a partial melt of a phlogopite magnesite lherzolite at about 6 Pa. During melting, phlogopite and magnesite are consumed during the initial stages and do not remain as residual phases.

328

60

~

0 .c ~ 50 -La.J a:: => en 40 en La.J a:: a..

30

LHZ + Cm +Ph

+V

1000 1200 1400

TEMPERATURE

L

1600

(OC)

CHAPrER4

Figure 4.9. Phase relationships of a cobaltian synthetic kimberlite composition containing 5 wt % C02 and 5 wt % H20 extrapolated from phase equilibria at 3.0 and 5.5 GPa. The heavy line is the univariant zone of invariant vapor composition (ZIVC) solidus for haploperidotite-C02. LHZ = lherzolite, Cm = magnesite, Ph = phlogopite, Dol = dolomite, L = liquid, V = vapor (after Eggler and Wendlandt 1979).

Eggler and Wendlandt (1979) suggest that kimberlites are generated by melting of carbonated peridotite at 5-6 GPa (160-200 km) as a rising asthenospheric diapir reaches a level at which the solidus intersects the local geotherm.

4.3.2.2.h. Ultra-high-Pressure Experiments. Ringwood et al. (1992) and Kesson et al. (1994) investigated a Co-analogue of a composition similar to that of the average group lAkimberlite of Smith et at. (1985b). This average composition is based upon the analysis of a carefully selected group of samples and provides a reasonable estimate of the bulk composition of an uncontaminated, but not necessarily primitive, kimberlite.

Experiments were undertaken at the very high pressures of 10 and 16 GPa because the discovery of majorite garnets (Moore and Gurney 1985) and xenoliths derived from depths greater than 300 km (Haggerty and Sautter 1990) suggested that kimberlites might be generated within the transition zone of the mantle. Majorite garnets are formed when the components of pyroxenes are taken into solid solution by preexisting garnets at high pressures. Pyroxenes are not stable below 500-km depth in the transition zone. Majorite garnets are solid solutions between normal garnets (Mg,Fe,CahAhSbOJ2 and an alumina-free end member (Mg,Fe,Cah[(Mg,Fe,Ca)(Si)]SbOJ2 containing octahedrally coordinated Si (Ringwood 1967).

Experiments at 16 GPa indicate that the liquidus is slightly below 1700°C, and near-liquidus phases at 1650°C are ~-Mg2Si04 and majorite garnet. Below 1600°C, clinoenstatite joins the assemblage. The solidus is located at about 1525°C with the subsolidus assemblage consisting of ~-Mg2Si04, majorite gamet, clinoenstatite, and dolomitic carbonate.


Initial experiments by Ringwood et al. (1992) at 10 GPa placed the liquidus below 1650°C. The near-liquidus assemblage consisted of olivine, clinoenstatite, and perovskite. These are joined by gamet (21% majorite) at 1500°C and by primary magnesite and clinopyroxene at 1400°C. The solidus is located between 1200 and 1300°C.

Kesson et al. (1994) conducted a second series of experiments in recognition of the discrepancy between their initial results and those of Edgar and Charbonneau (1993) at 10 GPa. The principal differences were the appearance of garnet well below the liquidus and the presence of primary perovskite.

The Wesselton aphanitic kimberlite studied by Edgar and Chabonneau (1993) is depleted in Si02 and enriched CaO relative to Ringwood et al.'s (1992) starting composition. Kesson et al. (1994) suggest that these compositional features result in the expansion of the garnet stability field toward the liquidus in Edgar and Charbonneau's experiments. However, the garnets in the latter are not pyrope rich and are completely unlike those crystallized in either of the experiments by Ringwood et al. (1992) and Kesson et al. (1994). In the latter experiments, olivine was shown to be the liquidus phase at 10 GPa and 1650°C, followed closely by pyrope-rich garnet. Perovskite was considered to be a near-solidus primary phase rather than a near-liquidus phase.

Ringwood et al. (1992) conclude that the experimental phase relationships preclude kimberlite being formed by the partial melting of a garnet lherzolite at depths ofless than 300 km. The basis for this conclusion is the absence of garnet and orthopyroxene and the presence of perovskite as near-liquidus phases. Accordingly, their preferred model is based upon their 16-GPa experiments and involves small degrees of partial melting of "refertilized" former harzburgites in the transition zone (>400 km). This source is composed of majorite garnet and ~- or y-(Mg,Fe)2Si04. Based upon these few experimental data, Ringwood et al. (1992) erected a complex hypothesis for the origin ofthe source involving interactions between subducted oceanic crust and upwelling asthenosphere at the 650-km discontinuity. This model is discussed further in Section 4.4.1.

4.3.2.3. Summary-A Cautionary Note

The differences between the three sets ofliquidus experimental studies are undoubtedly mainly related to compositional differences between the starting materials. These studies provide an excellent illustration of the problems inherent in choosing different compositions as "primitive kimberlite," as the experimental data cannot be reconciled and lead to very different conclusions regarding the composition of the source and depth of origin of kimberlites.

Importantly, none of the liquidus or partial melting experimental studies summarized above indicate that kimberlites and orangeites are formed by melting of similar source rocks.

4.4. GEODYNAMIC MODELS OF KIMBERLITE AND ORANGEITE GENESIS

Geodynamic models of kimberlite and orangeite genesis attempt to place their origins in the context of our current understanding of plate tectonics and mantle dynamics. Given our inadequate knowledge of many deep mantle processes and the long-term evolution

330 CHAPTER 4

of the Earth, many of these models are, of necessity, highly speculative; however, they are important in that they serve to direct us to new avenues of investigation which might provide useful evidence for rejecting or verifying these hypotheses.

Some recent geodynamic models of kimberlite genesis (Ringwood et al. 1992, Haggerty 1994) have placed their sources at ever greater depths in the mantle, even as deep as the core-mantle boundary. The stimulus for these hypotheses was the discovery that some diamonds in the Monastery kimberlites contain syngenetic inclusions of majorite garnet (Moore and Gurney 1985, 1989, Moore et al. 1991). On the basis of experimental studies of garnet-pyroxene solid solutions at high pressures (Irifune et al. 1989), it was suggested by Moore et al. (1991) that the inclusions formed at depths of at least 480 km, i.e., within the transition zone of the mantle.

Subsequently, Wilding et al. (1994) described inclusions within Brazilian alluvial diamonds, supposedly of kimberlitic provenance, consisting of majoritic garnets whose compositions indicated derivation from depths of 180--400 km. Other inclusions present are CaSi03, MgO-FeO solid solutions (periclase-wustite), Si02 (? stishovite), diopside, CazAbSi07, moissanite, and olivine (F08s). At depths greater than 650 km, experimental evidence indicates that mantle olivine and pyroxene are replaced by perovskite-structured "pyroxene" and ferripericlase. Therefore, Wilding et al. (1994) suggest that some members of this inclusion suite (CaSi03, MgO-FeO, Si02) may represent minerals originating from below these depths.

A deep origin for kimberlites was also considered by Haggerty and Sautter (1990) and Sautter et al. (1991), who reported the occurrence of orientated incl usions of pyroxene in garnet in xenoliths from the Jagersfontein kimberlite. These were interpreted to represent majorite garnets which crystallized at 300--400 km depth and were then transported to shallower depths where exsolution of pyroxene occurred.

Initially, the majorite garnet-bearing diamonds and xenoliths were believed to have crystallized in the transition zone, and been subsequently carried to shallower depths by convecting mantle, where they were eventually sampled by kimberlites (Moore and Gurney 1985, 1989). In this scenario, kimberlites are not formed in the transition zone. Subsequently, Moore et al. (1991) considered the possibility that the kimberlite itself was generated at depths exceeding 450 km and transported the majorite-bearing diamonds directly to the surface. Ringwood et al. (1992) commenced their high-pressure liquidus experimental studies of kimberlite (4.3.2.2) on the basis of this suggestion. Haggerty (1994) has suggested that diamonds bearing strongly-reduced mineral inclusions (SiC, silicate perovskite, wustite-periclase), might be formed in the lower mantle, and further implied that their transporting hosts might commence their ascent from these depths (see 4.4.2).

In considering hypotheses based upon the above observations, it should be kept in mind that occurrences of majoritic garnets in diamonds and xenoliths are the exception rather than the rule. Similarly, the metasomatized peridotite xenolith suite found in some southern African kimberlites (Erlank et al. 1987) may not be characteristic of all cratons. Consequently, general petrogenetic schemes erected upon the recognition of these few atypical occurrences should be regarded with due caution as to their universal, and even local, veracity.


4.4.1. Transition Zone Melting

Ringwood et al. (1992) have suggested that kimberlites are formed by partial melting of a majorite plus y-Mg2Si04 assemblage in the transition zone (400-650 km). The origins of this source are sought in a complex model, which is a variant of Ringwood's (1989) hypothesis of subduction of oceanic crust to the 650-km discontinuity, as a means of explaining the composition and evolution of the mantle.

In Ringwood et al.'s (1992) model, oceanic crust (basalt plus sediments) and lithospheric harzburgite are subducted to the 650-km discontinuity. The subducted slab, or megalith, remains at the discontinuity because of the density contrast between it and the underlying lower mantle. Here, oceanic crust recrystallizes to a volatile-bearing mixture of gametite (>85% majorite garnet + <15% stishovite) and recrystallized former sediments. At these pressures and temperatures harzburgite recrystallizes to spinel-structured y-Mg2Si04 (>80%) plus minor «20%) majorite garnet, ilmenite, and stishovite. Former harzburgite may be mixed with the garnetite layer (Ringwood 1989) or accumulate above it (Ringwood et al. 1992). Partial melting in the gametite layer results in the addition of incompatible element-rich melts to overlying harzburgitic material. These melts have higher LalYb, U/Pb, and Th/u ratios than their parental MORB-derived gametite. Ringwood et al. (1992) suggestthat the persistence of the "refertilized" lithology for times of up to 1 Ga, without further melting, may allow sufficient time for isotopic heterogeneities to develop.

Subsequently, the residual garnetite and the refertilized former harzburgite were heated by convection currents rising from the lower mantle. The gravitationally stable, and highly viscous, garnetite layer forms a barrier which deflects convection currents, causing the layer to be elevated by as much as 50 km, forming a dome-like structure several thousand kilometers in diameter (Figure 4.10). Previous melt extraction results in the gametite layer having a higher solidus temperature than overlying refertilized former harzburgite. Consequently, heat from the convection current causes remelting of only the latter, forming kimberlitic (and lamproitic) magmas. These ascend to the upper mantle, collecting xenoliths, primarily from subcontinental lithosphere, and rarely from the transition zone. The hypothesis explains the "highly fractionated" geochemical signature of kimberlite as resulting from a combination of two separate partial melting events, both occurring in the presence of garnet.

Ringwood et al. (1992) derive archetypal kimberlites, orangeites, and olivine lamproites by the same process at the same depths. Compositional and isotopic differences between archetypal kimberlites and orangeites are considered to originate from the derivation of the first partial melts in the gametite layer from regions relatively poor, or rich in, subducted pelagic sediments, respectively. Significant mixing of second-stage melts does not occur and magmas derived from isotopically distinct portions of the refertilized layer segregate and rise as separate batches of magma.

Apart from the implausibility of deriving three distinct magma types from a single source by the same process, Ringwood et al. 's (1992) hypothesis has many problems of a general geophysical-geodynamic nature related to the thermal structure and rheological character of the mantle. Discussion of these is entirely beyond the scope of this work, and comments on the hypothesis are confined to aspects specific to kimberlite genesis.

332 CHAPTER 4

Figure 4.10. Model for the generation of kimberlites by partial melting of refertilized former harzburgites at the 650-km discontinuity. Melting and uplift of the garnetite layer occur in response to temperature increases resulting from the upwelling of a lower mantle convection current. See text for further explanation (after Ringwood et al. 1992).

Ringwood et al. (1992) consider it impossible to derive kimberlites and orangeites from different sources because of constraints imposed by isotopic and major element compositions. The apparent paradox between this observation and arguments in favor of separate lithospheric and asthenospheric sources for orangeite and kimberlites, respectively, noted by Ringwood et al. (1992, p. 532), stem from considering both magmas to be genetically related. The problem disappears if the magmas are not genetically related and derived from different sources. Requiring orangeites to be derived from an asthenospheric source is not in accord with the occurrence of orangeites only in the Kaapvaal craton. Thus, the temporal and geological relationships of kimberlite, orangeite, and lamproite required by the model are not supported by observation.

For the above reasons, coupled with the mineralogical and geochemical observations given in this work and Mitchell and Bergman (1991), it is considered that Ringwood et al. 's (1992) model may have applicability to archetypal kimberlites, but is not relevant to either orangeite or lamproite genesis.

Unfortunately, even with regard to archetypal kimberlites, the model lacks specific details regarding the extent of partial melting in both stages, and the mineralogy of the refertilized source. In addition, surely discussion of the fate of subducted carbonate during


each melting episode should be a prerequisite to this model of kimberlite genesis. Interestingly, a deep asthenospheric source for kimberlites suggests that they need not be confined to continental regions.

It is not possible to reject or accept the novel hypothesis that kimberlite generation is in some manner connected with partial melting of a majorite- ~- or y-spinel source until further experiments are undertaken on the potential source material. The complexity of Ringwood et al. 's (1992) model results from the bias held-that subduction of oceanic crust is required to provide a source of incompatible elements. The simpler alternative model of decompressional melting of primitive or metasomatized mantle at these depths is not explored.

4.4.2. Metasome Melting and Mantle Plumes

Haggerty (1989a,b) has suggested that a necessary prerequisite for the genesis of kimberlitic and other varieties of continental alkaline magmatism is the formation of metasomatic mineral assemblages in depleted harzburgitic lithosphere. These assemblages are envisioned to form layers, termed "metasomes," in the continental lithospheric mantle over a depth range of 60-100 km. Xenoliths of metasomatized peridotite containing phlogopite, K-richterite, lindsleyite-mathiasite, hawthorneite-yimengite, and other titanates (Erlank et al. 1987, Haggerty 1983) found in some kimberlites are considered to be fragments of these metasomes. Metasome formation takes place when asthenospheric melts solidify in the lithosphere as they cool below the solidus of carbonated lherzolite. Their crystallization releases fluids containing dissolved incompatible elements which migrate into the surrounding lithosphere. The formation of metasomes is a variant of Wyllie's (1980) model of melt-lithosphere interaction (4.3.1.1). It differs in that Haggerty (1989a,b) requires metasomatism to be a precursor, and not a consequence, of kimberlite magmatism.

The metasomatic zone is divided into two overlapping units, a lower (100-75 km) zone representing the phlogopite-K -richterite peridotite (PKP) assemblage and an upper carbonate-rich horizon <65 km) containing zircon, ilmenite, and calcite. Both zones contain phlogopite, K-richterite, and alkali titanates. Evidence for the depths and character of the zonation is weak, as the distributions of purported assemblages are not constrained by accurate equilibration pressures and temperatures, paragenetic, or mineral stability data.

Haggerty (1989,a,b) proposes that a rapidly ascending asthenospheric "protomelt" will induce preferential partial melting of the metasome, as the latter is a region with a lower solidus temperature than the surrounding mantle. The protomelt is considered to be komatiitic for no obvious reasons other than Haggerty's (1989,a,b) requirement that this magma have higher MgO contents and temperatures than those magmas (unspecified but presumably basaltic) which generate the metasomes.

Incorporation of metasome and included volatiles into the hot komatiitic picrite then leads to "flash melting" (an undefined process) and the explosive eruption of alkaline melts. Kimberlites and lamproites are purported to be generated in the PKP zone and carbonatites in the calcite-rich zone.

334 CHAPTER 4

Haggerty (1994) has suggested that events at the core-mantle boundary are ultimately responsible for the generation of the protomelts which subsequently interact with metasomatized lithosphere to generate kimberlites. The hypothesis is based on the presence, in some kimberlites, of xenoliths and diamonds derived from, or below, the transition zone and the assumptions that there is a supposed correlation of kimberlite magmatic activity with times when the Earth's geomagnetic field remained stationary for long periods (40 Ma) of time; kimberlite magmatism is considered to be global rather than local; and there is a direct relationship between kimberlite magmatism and deep mantle plumes.

According to this model, instabilities at the D" layer, caused by crystaJIization and/or convection events in the core, result in the formation of mantle plumes which travel rapidly (25-50 Ma) from the D" layer to the lithosphere. Komatiitic melts derived from this plume are arrested in the lithosphere and cause partial melting of the PKP metasomes to form kimberlites.

Detailed criticism of many aspects of Haggerty's (1989a,b 1994) model, i.e., plume ascent, origin of the D" layer, relationship of magmatism to the cyclicity of the geomagnetic field, etc., is beyond the scope of this work and best left to the geophysical community. Other parts of the model are difficult to critique as ideas are expressed as ex cathedra pronouncements not supported by fact or reasoned argument. Further, parts of the model are simply incorrect, such as the statement that carbonatites and lamproites are closely-related (Foley 1992a,b, Mitchell and Bergman 1991).

Haggerty's concept that the genesis of kimberlite hinges upon the supposed assimilation of metasomes in a komatiitic-picritic magma is implausible. The principal argument against this process is that komatiitic magma is confined to Archean times characterized by unusually high heat flows (Hess 1989). Haggerty does not explain either how this magma is generated in post-Archean times or its relationships to the voluminous continental flood basalts which are also ascribed to the melting of mantle plumes. The details of the assimilative process are glossed over, and the concept of "flash melting" is never put on a firm physicochemical basis.

Importantly, Haggerty does not explain why flood basalts in Siberia occur in the midst of an extended period of kimberlite magmatism, kimberlites in southern Africa are preceded by basalts, and kimberlites in Colorado, Tanzania, China, Canada (Saskatchewan, Somerset Island, Lac De Gras), Zaire, and Arkhangelsk are not associated at all with flood basalts. Explanation of some of these observations is critical, given that the solidus of metasomatized mantle is undoubtedly well below the temperature of basaltic magmas passing through the lithosphere. The experimental data of Thibault et al. (1992) support this observation as they show that, at 3 GPa, the solidus of carbonated phlogopite lherzolite is as low as 1060°C. Thus, melting of metasomes by precursor episodes of basaltic magmatism is highly probable (see 4.5.2.1). Based upon these observations, Thibault et al. (1992) have formulated a model, similar to Haggerty's, in which decoupled metasomatic zones rich in either C02 or H20 are generated in the lithospheric mantle by successive cycles of melting, migration, and solidification of incompatible-element-rich melt/fluids (see 4.3.1.3). Interestingly, neither Haggerty nor Thibault et al. (1992) refer to each other's work. Finally, remobilization or assimilation of enriched, i.e., metasomatized, lithosphere during emplacement of magma has recently emerged as a fashionable


concept to explain the subtleties of flood basalt geochemistry (Cox 1992, Ellam and Cox 1991, Hawkesworth et al. 1986).

Coincidence of epochs of stable magnetic polarity with episodes of kimberlite magmatism is not surprising, given the length of the epochs relative to the time scale of any type of frequently occurring magmatism; consequently, this aspect of the model has little petrogenetic significance. (Haggerty does not discuss the errors attached to the time scales of geomagnetic epochs.)

Finally, a major omission in Haggerty's (1 989a,b, 1994) papers is the failure to discuss opposing hypotheses suggesting that plumes have absolutely nothing to do with continental alkaline magma generation-i.e., lithospheric fracture propagation (Turcotte and Oxburgh (1978) or lithospheric volatile fluxing (Bailey 1977, 1983, 1992).

The principal merit of Haggerty's model is that it draws our attention to the complexity of the metasomatic events in the upper mantle. The proposed metasomatic "stratigraphy" is innovative but, unfortunately, difficult to prove on a quantitative basis. Although the model is entirely inappropriate for generating kimberlites, it does support the concept that the melting of enriched regions of the lithosphere might playa role in the genesis oflamproites (and orangeites).

4.4.3. Hot-Spot Melting

Plate tectonic reconstructions of southern Africa, combined with geochemical data, have been used by le Roex (1986) to suggest a relationship between kimberlites and orangeite magmatism to hot-spot activity. In this interpretation, le Roex (1986) notes that the geochemical characteristics of orangeites are similar to those of oceanic island basaltic (OIB) rocks occurring within the Dupal isotopic anomaly (Zindler and Hart 1986). This anomaly encircles the Earth between the equator and 600 S latitude and is characterized by enrichment in incompatible trace elements, high 87Sr/86Sr, 206pbP04Pb ratios and low 143Nd/l44Nd ratios. The source of the Dupal anomaly has been interpreted as ancient subducted-recycled oceanic lithosphere including pelagic sediments, delaminated subcontinental lithosphere, or upwelling primordial mantle (Allegre and Turcotte 1985, Zindler and Hart 1986). Upwelling of this anomalous asthenospheric material results in hot-spot volcanism with an apparent lithospheric signature, whereas upwelling outside the Dupal anomaly region leads to volcanism with the normal depleted asthenospheric isotopic signature. Extending these observations to orangeites implies that they are asthenospheric melts derived from Dupal-type sources (Ie Roex 1986). Orangeite trace element and isotopic compositions suggest derivation from a source more enriched than those of Dupal-OIB basalts. As such they could also be considered as mixtures between MORB-type sources and an even more enriched source.

Le Roex (1986) has noted orangeites lie at the termination of compositional trends defined by Dupal-type OIBs, whereas archetypal kimberlites terminate trends defined by normal OIBs (Figure 4.11). Note that these trends are not well defined for all isotopic ratios or elemental ratios. Regardless, Ie Roex (1986) suggests that both archetypal kimberlites and orangeites are derived from asthenospheric sources.

Le Roex (1986) has attempted to correlate kimberlite and orangeite asthenospheric magmatism with hot-spot tracks in the South Atlantic (Figure 4.12) and believes that the

336 CHAPTER 4

I I I I I I I I I 0·5122 0·5126 0'5130

143 Nd / 144 Nd

20 -

1-ORANGEITE

I I I I I I I I I

0'704 0'706 0'708 0'710

87 Sr / 86 Sr

Figure 4.11. Correlations between Zr/Nb ratios and Sr or Nd isotopic compositions of South Atlantic basalts, kimberlites, and orangeites (after Ie Roex 1986).

distribution of orangeites is consistent with the Shona hot-spot paleotrack. In contrast, kimberlites show no simple correlation with either the Bouvet or Discovery paleotracks.

Similar conclusions have been drawn by Skinner (1989), who has noted that the almost linear progression of orangeite ages in southern Africa suggest that hot-spot activity may have been important with respect to their formation. In contrast to Ie Roex (1986), Skinner (1989) regards hot-spot activity to transfer only heat and volatiles to the lithosphere, where kimberlites and orangeites are generated by melting of lithospheric sources. Thus, both authors use the same data to arrive at completely different conclusions regarding sources of these magmas. In agreement with le Roex, Skinner (1989) finds no correlation between kimberlite magmatism and hot-spot tracks.

Mitchell (1986) and Bailey (1983, 1993) have discussed hot-spot models of continental alkaline volcanism and found them to be unconvincing. The principal arguments against the model are that it cannot satisfactorily explain the worldwide distribution of


o !

AFRICA

337

KIMBERLITES

Figure 4.12. Location and age ofkimberlites and orangeites (from Skinner 1989) relative to postulated hot-spot palaeotracks during the Upper Jurassic and Lower Cretaceous in southern Africa. Numbers on tracks show ages in Ma (after Ie Roex 1986).

kimberlites in particular and continental alkaline magmas in general; most kimberlites within a province have been emplaced at the same time over a wide region and not sequentially over curvilinear trends; hot spots fail to explain the repeated eruption of similar magmas over a long period of time in a relatively restricted area. Hot spots are usually conceived of as narrowly focussed heat sources (Morgan 1971) in marked contrast to the current interpretation of mantle plumes as large (1000-2000 km at their heads) upwellings of mantle.

Section 1.7 suggests that the southern African orangeite province may be better regarded as reflecting two distinct periods of magmatism rather than a single hot-spot track. The width of the province, distribution of intrusions of different age within it (see 1.7), and especially the concentration of activity at the southeastern end preclude any simple, single, small hot-spot model for its formation. Application of the concept to this province would lead to an improbable proliferation of multiple hot-spot tracks. The

apparent absence of orangeites between Swartruggens and Barkly West is also difficult

to reconcile with a single hot-spot track.

338 CHAPTER 4

Bailey (1993) rejects all plume-hot-spot-related models of continental alkaline igneous activity and suggests that the apparent age progression of orangeites may reflect progressive structural readjustments into the craton following continental breakup and the development of the Lebombo monocline. In this model, orangeites (and kimberlites) are generated in response to passive reactivation of the continental plate by external forces (see 4.5.2.1).

With regard to Ie Roex's (1986) geochemical arguments, there are simpler ways of explaining the isotopic data than by appealing to unknown asthenospheric sources. For example, apart from the conventional interpretation of distinct asthenospheric and lithospheric sources, kimberlites and orangeites could be derived from young- and ancientenriched lithospheric sources, respectively. In either instance trace element correlation diagrams involving asthenospheric Dupal-type oms and lithospheric orangeites have no petrological/geological foundation and are irrelevant, as they simply compare apples and orange(i te)s.

4.4.4. Partial Melting of Heterogeneous Lithosphere

Skinner (1986, 1989) proposed the first hypothesis attempting to provide a general model for the origin and emplacement of archetypal kimberlites (referred to as group I kimberlites) and orangeites (group II kimberlites). The model relies upon small degrees of partial melting of a phlogopite garnet lherzolite as source for the magmas. The hypothesis is based upon Wyllie's (1980) model of kimberlite genesis (see 4.3.1.1) together with an unconventional interpretation of Sr and Nd isotopic data (3.8.1).

Skinner (1989) considers that the broad similarities in the incompatible-element abundance patterns of kimberlite and orangeites (see 3.6.1), recognized by Smith et al. (1985b), indicate that both magmas are derived from "broadly similar sources" (Skinner 1989, p. 539). On this basis alone, it is assumed that both magmas are derived from compositionally different domains in the continental lithosphere, regardless of the conventional interpretation of the isotopic composition of kimberlite as indicative of asthenospheric sources.

Skinner (1989) further assumes that magmas are generated by small degrees of melting of a phlogopite-bearing source as a result of the introduction of volatile elements (C02, H20) and heat from the asthenosphere into the lithosphere (Wyllie 1980). The lithosphere is considered to be heterogeneous, as the source rocks of orangeites and kimberlites are considered to be geochemically and isotopically distinct. The sources of orangeites were enriched in K, Pb, Rb, Ba, LREE, Si02, and H20, and depleted in Ti02, Nb, and C02, relative to those of kimberlites. Enrichment of orangeite sources occurred at least 500 Ma-1A Ga ago and were isolated from those of kimberlites. Skinner (1989) does not discuss the time of formation of the kimberlite sources, except to state that they represent depleted mantle modified by older periods of metasomatism.

The crux of Skinner's (1989, p. 540) hypothesis lies in his statement that

sectors of the lithosphere that were relati vely more highly depleted with basaltic elements (i.e., essentially harzburgiticl may have acted as preferential sites for later incompatible element enrichment. Melting caused by the sti1llater introduction of H20 and C02, would be accelerated in the highly LREE enriched zones but would be retarded within those zones less enriched in LREE. The fact that Group II kimberlites predate Group I kimberlite occurrences, where both


occur in the same area, indicates that if the genesis of both magma types is initiated by the same or similar processes, then Group II magmas are generated earlier, and hence emplaced earlier than Group I kimberlites.

339

Hence, Skinner (1989) believes that orangeites and kimberlites are produced by the same process, but the rate of melting of the source rocks differs as a consequence of their differing composition. The existence of a previously metasomatized mantle is a prerequisite for Skinner's model.

Figure 4.13 summarizes Skinner's (1989) four-stage hypothesis for the genesis, ascent, and emplacement of kimberlite and orangeite. In stage 1 the introduction of asthenosphere-derived volatiles results in melt development at point A on the shield geotherm. In the LREE-, Rb-, and Ba-enriched regions of the lithosphere, slow percolation of melts derived from this source, i.e., orangeites, coalesce into small magma pockets at B. Formation of kimberlite magma takes place over the region A-C and involves a larger vertical section of lithosphere, less enriched in REE, i.e., kimberlite magma generation is slower than that of orangeites for the same volatile introduction event and the melts have isotopic signatures characteristic of depleted mantle. The evidence for magma development over an extended section of lithosphere is the apparent restriction of metasomatized lherzolitic xenoliths, and the Ti-rich macrocryst suite, to kimberlites. Skinner (1989), following Erlank et al. (1987) and Haggerty (1986), considers that the former are all generated at relatively high levels in the lithosphere by diverse metasomatic processes and the latter are xenocrysts derived from magmas unrelated to kimberlites.

50

~ 100

~

Q. 150 Q)

o 200

250

A

Buffered Peridotite

500 1000 1500

Temperature (OC)

ORANGEITE KIMBERLITE

.aJ Crust : : ~

Figure 4.13. Model of kimberlite (K) and orangeite (0) magmatism according to Skinner (1989). See text for explanation.

340 CHAPTER 4

From the paucity of xenoliths found in orangeites it follows that they must be derived from smaller volumes of mantle relative to kimberlites.

Stage 2 of the model covers the slow ascent of the magmas up to a level at which the rate of movement rapidly increases. During this stage the magmas are believed to incorporate the bulk of their xenolith suites. Orangeites following path A-D on Figure 4.13 do not sample the upper portions of the lithosphere, whereas kimberlites following path A-E sample a wide variety of xenoliths. Skinner (1989) believes that some crystallization of primary olivine, ilmenite, spinel, and phlogopite may occur during this stage.

During stage 3 it is envisaged that discrete batches of magma accelerate rapidly and move upward by means of crack propagation. The rapid ascent is initiated by an increase in the volatile content of the melt resulting from crystallization of later generations of olivine and other primary minerals. For orangeites and kimberlites the rapid ascent paths on Figure 4.13 are from D-F and E-F respectively. Relatively few xenoliths are incorporated during this stage of the ascent. Both magmas are believed to decelerate rapidly as they approach point F as a consequence of degassing into open fractures.

Stage 4 from point F-G (Figure 4.13) involves the slow intrusion of the magmas as sills and dikes, together with the formation of diatremes. Late-stage ground mass minerals and segregationary textures are formed during an extended period of post -emplacement crystallization.

Skinner (1989) does not discuss in any detail the origins of the asthenospheric volatiles which initiate the partial melting process. Allusion is made to the linear age progression of 145-110 Ma orangeites in southern Africa as being suggestive of transfer of volatiles and heat from a stationary hot-spot into a moving lithospheric plate. However, the lack of an apparent correlation between age and distribution of 95-80 Ma kimberlites is not consistent with this hypothesis. The absence of a correlation may be related to the extended time proposed by Skinner (1989) for development of kimberlites versus orangeites. Note that Skinner's model was developed primarily to explain the age and geochemical relationships between southern African J urassic-Cretaceous kimberlites and orangeites. Skinner (1989) does not discuss why the emplacement of the older kimberlite provinces of southern Africa, and elsewhere in the world, are not preceded by orangeite magmatism.

Skinner's (1989) hypothesis is appealing in that it attempts to draw together diverse aspects of both kimberlite and orangeite magmatism which must be considered in any general model of continental alkaline rock petrogenesis. Unfortunately, the model is ultimately unsatisfactory as it is based on numerous unfounded assumptions. The model stands and falls upon the initial premises of stages 1 and 2. The claim that kimberlites are lithospheric melts is simply an assumption and not supported by critical examination of the geochemical data. Similarly, the explanation of how depleted isotopic signatures might develop in the lithospheric mantle is not convincing, as no arguments are presented against the alternative hypotheses of Smith (1983).

Skinner (1989) does not present any evidence to explain why the partial melting processes leading to the generation of kimberlite and orangeite must be initiated by the contemporaneous introduction of asthenospheric volatiles. Further, there is no discussion as to why the model does not appear to be valid for other kimberlite provinces, a prerequisite for any general petrogenetic model.


Unfortunately, there is no experimental evidence to support the concept that the rate of melting is related to the composition of the mantle. Skinner (1989) does not address the question of the modal mineralogy of his heterogeneous mantle sources, but does imply that both are formed by similar metasomatic processes; hence, although their contents of incompatible elements might differ, they may be mineralogically similar. Both would therefore have very similar solidus temperatures in contradiction to Skinner's model.

Other problematic aspects ofthe model hinge upon the validity of Wyllie's (1980) model of kimberlite genesis and especially with regard to Skinner's uncritical application of this model to orangeites.

In conclusion, Skinner's (1989) model, at first sight, appears to provide a catholic interpretation of kimberlite and orangeite petrogenesis; however, upon examination it is found wanting, as it is based upon novel unsupported interpretations of geochemical and isotopic data combined with numerous poorly-defined petrogenetic processes. It is open to speculation as to why Tainton (1992), Tainton and McKenzie (1994), and Fraser and Hawkesworth (1992) have refrained from any discussion of Skinner's (1989) model.

4.4.5. Redox Melting

Green et al. (1987), Taylor and Green (1989), and Foley (1988, 1989) have proposed a mechanism for magma production in the upper mantle, termed "redox melting." In this scheme, melting of upper mantle peridotites is triggered by the oxidation of reduced C-O-H fluids (CIt! + H2) to carbon (graphite or diamond), C02, and H20. Production of water results in a reduction of solidus temperatures and initiates melting. In this hypothesis, melting is caused by volatile introduction rather than temperature increase. The reduced volatiles are considered to originate in the lower mantle as a consequence of continued degassing of the primordial earth. Operation of redox melting depends upon the existence of a reduced asthenosphere if 02 < iron-wustite buffer + 1-2 log units) and an oxidized lithosphere if 02 > magnetite-wustite buffer or quartz-fayalite-magnetite buffer). Oxidation of CIt! and H2 will occur at the interface between the reduced and oxidized domains (Figure 4.14), and carbon may be precipitated at or below the lithosphere-asthenosphere boundary as microdiamonds. C02 produced during oxidation may carbonate lithospheric peridotites or remain dissolved in the partial melt.

The initial stage of redox melting involves oxidation of CIt!-bearing fluids to H20-rich fluids. Migration of these fluids and reaction with preexisting minerals will lead to formation of hydrous phases at all levels in the lithosphere. With the introduction of more fluid, increased activities of water will eventually cause partial melting. As the water produced at the oxidation zone front migrates upward, melting will occur in the overlying oxidized zone (Figure 4.14). The nature of the melts produced will depend upon the lithosphere composition and the rate of oxidation of the fluid.

Foley (1988) suggests that the oxygen fugacity of the lithosphere is such that carbonates are present as subsolidus phases. As a H20- and CIt!-rich fluid moves slowly into the lithosphere and oxidizes to H20-rich compositions, melting should occur in the presence of H20- and C02-rich fluids. Melts will be restricted to silica-poor compositions which may be kimberlitic or melilititic at high and low pressures, respectively. If the fluid moves rapidly and equilibrium is not maintained, it will not oxidize completely and

342

OXIDISED LITHOSPHERE

~

KIMBERLITIC MELT

rn MELT

CHAPTER 4

W~\d ZONE OF PARTIAL MELTING

Figure 4.14. Contrasting origins for kimberlites and lamproites according to the redox melting model of Foley (1988). (A) Kimberlites are formed during the initial stage of melting where oxidation fronts are relatively well defmed and the oxidized zone is fertile. Partial melting is induced by the oxidation of hydrogen and methane to water. (B) Lamproites are formed in subsequent stages of melting from mantle which is depleted in major elements but enriched in incompatible elements. The reduced nature of this region results in H20-rich, C02-free melting at 1.5-2.5 GPa. C02-bearing olivine lamproites (and perhaps orangeites) may originate at 4.5-5.5 GPa, where oxidized blocks persist in regions which are primarily reduced. See text for further explanation (after Foley 1988).

methane-bearing fluids will encounter carbonated peridotites. Carbonate should be consumed by decomposition to graphite or diamond and more water will be produced. High activities of water may cause melting to occur while carbonates persist. In this disequilibrium process, restricted amounts of low-silica melt (kimberlite) may be produced before continued introduction of methane causes all of the carbonate to be reacted out. Subsequent to exhaustion of carbonate, more voluminous Si02- and H20-rich melts will be produced. These might be basanitic to nephelinitic in character.

Foley (1988) notes that lamproitic melts may form as a consequence of renewed volatile introduction when Cli4-rich fluids infiltrate a mantle region which has been previously depleted in Ca, Na, and AI. Melts produced from these regions under reducing conditions will have a depleted major element signature and resemble lamproites provided that metasomatic enrichment in K, Ti, Zr, and REE, etc., has occurred prior to the second period of melting. Foley (1988) considers that the degree of geochemical depletion and intrinsic oxidation state in the region of melting may explain the difference between kimberlitic and lamproitic melts. Olivine lamproites may originate at similar depths to


kimberlites but from a depleted and reduced mantle formed during an earlier period of melting. Leucite lamproites may form at lower pressures from similar sources in H20-and F-rich regimes (Figure 4.14).

Foley's (1988) redox melting mechanism is an extension of the volatile-induced melting hypotheses of Wyllie (1980, 1989a,b) and Bailey (1983, 1993). It provides a physicochemical explanation for melting processes which link diverse continental alkaline magmas to a common source. The model may be applicable to lamproites and, by analogy, to orangeites, if the genesis of these lithospheric magmas is initiated by volatile uprise. The model is less attractive for the generation of kimberlites, as their parental magmas are probably formed by decompressional melting of diverse carbonated-ultrabasic assemblages in the asthenosphere (4.3.1, 4.3.2).

In common with all volatile-induced melting hypotheses, the model relies on the introduction of volatiles from unknown sources and might be considered untenable on this basis alone.

4.5. PETROGENESIS OF THE ORANGEITE CLAN

The petrogenesis of the orangeite clan comprises four principal stages: development of the mantle source; melting of the source; contamination and evolution of the magma during ascent through the mantle into the crust; post-emplacement crystallization. Any hypotheses regarding the genesis of the clan must account for the following observations:

1. Orangeites are derived from sources with time-integrated Rb/Sr and SmlNd ratios which are higher and lower, respectively, than those of bulk earth. These sources have been isolated from convecting mantle for 0.5-2 Ga.

2. Individual orangeites differ in their isotopic composition within and between intrusions.

3. Multistage and/or multi component models are required to explain the Sr, Nd, and Pb isotopic compositions.

4. The bulk rock major and trace element compositions of orangeites do not represent those of their parent magmas.

5. Orangeite magmas have been extensively contaminated by xenocrysts derived from disaggregated mantle-derived ultramafic rocks.

6. The composition of the parental magma of the clan is unknown, although orangeites appear to have formed from ultrapotassic peralkaline magmas.

7. Orangeites are strongly enriched in REE, Sr, Ba, Rb, and other incompatible elements.

8. Orangeites exhibit significant negative K, Rb, and Sr anomalies in primitive mantle normalized extended incompatible-element diagrams. The negative Ta, Nb, and Ti anomalies, which are characteristic of lamproites, are not present.

9. Individual fields differ in their petrographic character. Phlogopite, spinel, apatite, and carbonate are the only principal primary liquidus phases in some orangeite magmas, whereas, in others, these minerals are accompanied by diopside. Some of the latter crystallize late-stage sanidine and potassium richterite.

344 CHAPTER 4

10. The crystallization of primary olivine probably does not playa significant role in orangeite evolution, although small amounts of primary microphenocrystal olivine may crystallize prior to emplacement. In contrast to archetypal kimberlites second-generation subhedral primary groundmass olivines are rare. Many orangeites are characterized by extensive pre-emplacement crystal accumulation of phlogopite.

11. Orangeites have closer mineralogical and geochemical affinities to lamproites than to archetypal kimberlites. Thus, experimental studies oflamproite genesis might have a bearing on orangeite petrogenesis.

12. Partial melting of H20- and carbonate-bearing, non-Iherzolitic sources might be involved in the formation of orangeites.

13. To this date orangeites have been found only in southern Africa where they are confined to particular regions of the Kaapvaal craton, the majority occurring in the southwestern area.

14. Although the Kaapvaal craton has been the site of several periods of kimberlite magmatism, orangeite magmatism is restricted to Upper Jurassic and Lower Cretaceous times.

15. Emplacement of orangeites was preceded by extensive continental flood basaltic magmatism which apparently did not interact with the sources of orangeites.

There are no simple unambiguous explanations for the above observations. Given that orangeites have only recently been subjected to detailed study, it is unreasonable to expect that, on the basis of our current knowledge, we should be in a position to devise a distinctive comprehensive petrogenetic hypothesis which will stand the test of time. Consequently, the speculations concerning their genesis presented below must be regarded as preliminary and subject to change in the light of future observations and experimental studies.

4.5.1. Development of the Source

The restriction of orangeites to the Kaapvaal craton of South Africa, coupled with the isotopic data indicating derivation from ancient enriched material, suggests that their sources are located in the nonconvecting lithosphere. There is no convincing evidence for an asthenospheric source, because, if there were, orangeitic magmatism should be widespread throughout the world. Consequently, the initial stages in the formation of the orangeite source are believed to be related to the development of the Kaapvaal craton. Detailed discussion of the latter is beyond the scope of this work, but some general points concerning the structure and evolution of the craton are required to appreciate where and how orangeite magmas are generated. Source formation is considered to have taken place in two steps: development of a deep continental root of depleted harzburgite, and production of incompatible element-rich domains in this root by metasomatism (sensu lato). The depth to which the continental root extends is extremely important in placing constraints on the maximum depths of origin of lithosphere-derived melts; therefore, evidence pertaining to the location of the lithosphere-asthenosphere boundary is discussed here at some length.


4.5.1.1. Continental Roots

Currently, continental lithospheric plates are considered to consist of the crust and uppermost part of the mantle. The crust consists of a stable ancient nucleus, termed a "craton," which is surrounded by mobile belts and may be covered with younger platform sedimentary rocks. Cratons consist of late Archean or early Proterozoic rocks which behave as rigid units and react to external tectonic forces by epeirogenic faulting and uplift rather than by folding.

Most geophysicists and petrologists agree that the lithospheric mantle consists predominantly of low-density peridotite depleted of its low-melting "basaltic" fraction. Therefore the lithospheric mantle is less dense and substantially more refractory than the underlying asthenospheric mantle. Lithospheric mantle is rigid, and heat transfer occurs mainly by conduction. The chemical boundary layer, represented by depleted subcontinental mantle, stabilizes the continental root against thermal disruption by the convecting asthenospheric mantle. Cratons are attached to the lithospheric mantle root and move with it during plate motion. Depleted lithospheric mantle is subjected to metasomatism and intrusion by asthenospheric and lithospheric melts (Jordan 1978, Jeanloz 1987, James 1989).

Although there is agreement concerning the general character of the continental lithosphere, there is none regarding the depth of the lithosphere-asthenosphere boundary. Models based upon seismological, mechanical, rheological, thermal, or compositional grounds typically are not in agreement (James 1989). However, the received opinion holds that continental roots extend no deeper than 200--250 km (Anderson 1987, Boyd and Gurney 1986). In contrast, Jordan (1978) has presented persuasive arguments holding that continental roots, termed the "tectosphere," extend to depths of 300 km and perhaps even deeper. Resolution of this question by the geophysical community is important as the depth of the boundary constrains the maximum depth of generation of lithosphericderived magmas and has important ramifications regarding the fate of underplated subducted oceanic material. Boyd and Gurney (1986) have suggested that the low geothermal gradients and deep roots, characteristic of cratons today, were initially established more than 3 Ga ago. Although the deep root may have been initiated in the early Archean (see below), current models of komatiite (Hess 1989) and lithosphere genesis (Herzberg 1993) require high heat flows at that time, in contradiction to Boyd and Gurney's (1986) hypothesis.

Cratonic root formation has been linked with the initial stages of continental crust separation and the extraction oflarge volumes ofkomatiitic (Nixon et at. 1987, Boyd and Gurney 1986, Canil 1991) or basaltic magmas (Jordan 1978). The origins of the lithospheric peridotites comprising the Kaapvaal craton root have been reviewed by Herzberg (1993), who concludes that they originated as harzburgitic rocks formed as either residues or cumulates from a source material enriched in Si02, relative to normal mantle peridotite or its pyrolite analogue. These harzburgites could have crystallized as cumulates from an ultrabasic magma richer in Si02 than most komatiites. Formation of these magmas would have required extensive (>50%) melting of the primordial mantle. Thus, Herzberg (1993) sees cratonic root formation as being intimately linked with early Archean differentiation processes and confined to these times of high heat flow. In

346 CHAPTER 4

agreement with Jordan (1978) and Pollack (1986), Herzberg (1993) suggests that the FeO-poor Kaapvaal peridotites, having a reduced density compared to prirriltive mantle, would lead to the establishment of a "buoyant raft of lithosphere" isolated from the rest of the mantle and acting as a foundation for the Archean continental crust.

Continental roots are thus considered to originate in the early Archean and consist of depleted harzburgite. Formation of garnet harzburgite and garnet peridotite from this material subsequently occurs as a result of ex solution of garnet and clinopyroxene from orthopyroxene followed by recrystallization (Canil 1991, Cox et al. 1987).

Depleted harzburgite cratonic roots subsequently form a substrate which may be impregnated by metasomatic fluids or small-volume incompatible-element-enriched melts derived from the underlying asthenosphere. These may modify the composition of the lithosphere without introducing melting. Over the course of aeons this influx of incompatible element-rich matter produces geochemically and isotopically heterogeneous domains in the non-convecting lithosphere (4.5.1.3).

4.5.1.2. Depth oj Origin ojOrangeite Magmas

The depth and topography of the lithosphere-asthenosphere boundary of the Kaapvaal craton is not well-established. The current structural model of the craton is based primarily on the depth of origin of the peridotite xenolith suite found in kimberlites.

Boyd and Nixon (1975) initially claimed that equilibration parameters for these xenoliths may be used to devise an upper mantle stratigraphy. Studies of xenoliths derived from northern Lesothan kimberlites suggested the existence of two suites of xenoliths: a relatively cool suite of granular depleted garnet peridotites and a relatively hot suite of Fe-rich or "fertile" sheared or deformed peridotites. Boyd and Nixon (1975) and Finnerty and Boyd (1987) consider that the xenolith suite defines a paleogeotherm with an inflection marking the depth of the lithosphere-asthenosphere boundary at 150-160 km. Granular and sheared xenoliths are considered to represent depleted lithospheric and asthenospheric mantle, respectively. MacGregor and Manton (1986), Shervais et al. (1988), and Taylor and Neal (1989) have suggested that the paleogeotherm inflection is due to underplating of the craton with eclogitic material derived from subducted oceanic crust.

Other studies of peridotite xenoliths have shown no general correlation between degree of deformation, equilibration temperature, and depth of origin. Carswell and Gibb (1987, and references therein), in particular, have suggested that inflected paleogeotherms are an artifact of a particular choice of geothermobarometers and have no geological significance. The "fertile" xenoliths are also considered to be formed in thermal-metasomatic aureoles associated with kimberlite intrusions rather than being samples of undepleted asthenosphere (Mitchell 1984b, Harte 1983).

Equilibration temperatures and pressures of peridotite xenoliths found in kimberlites, regardless of the geothermobarometer chosen, suggest that none have originated at depths in excess of 220-250 km. Whether the xenoliths originating below about 160-km depth are asthenospheric or lithospheric remains a moot point.

In contrast to kimberlite, peridotite xenoliths are relatively rare in orangeites, having been found only in the Finsch diatreme (Shee et al. 1982). The presence of olivine and chrome garnet macrocrysts attests to the contarrilnation of orangeite magmas by ul-


tramafic material, but it is unclear whether the disaggregated protolith is peridotite, harzburgite, dunite, or wehrlite. Garnets derived from a wehrlite paragenesis have been reported from Newlands by Schulze (1989). Orangeites also contain xenocrystal subcalcic chrome pyropes considered to be derived from disaggregated garnet harzburgites (Pokhilenko et at. 1977), yet examples of their parental xenoliths are exceedingly rare in the Kaapvaal craton (Nixon et al. 1987).

The xenolithic suite of orangeites appears to be dominated by eclogites (MacGregor and Carter 1970, Hatton 1978, Taylor and Neal 1989, Gurney 1989). It is known that xenolith suites from geographically closely-related kimberlites are significantly different as a result of random sampling of the upper mantle during their ascent (Nixon 1987, Carswell and Gibb 1987, Mitchell et al. 1980). However, it is not known whether the paucity of peridotite xenoliths in orangeites, relative to kimberlites, is the result of such random sampling of the mantle or a reflection of some fundamental difference in the rheology and ascent rates of orangeite magmas relative to kimberlites. On the basis of present evidence it appears that orangeites do not typically sample the upper lithospheric mantle postulated to be the source of the granular depleted lherzolite suite (see also 4.5.3).

Estimation of the depth of origin of eclogites is more difficult than for lherzolites because geothermobarometry is based upon garnet-clinopyroxene Fe-Mg exchange equilibria which are not pressure and temperature independent (Ellis and Green 1979). Hence, equilibration temperatures may be calculated only by assuming a particular pressure; nevertheless, eclogites hosted by orangeites have been shown to have equilibrated over a wide range of temperatures (800-1050DC) at pressures of 3.0--5.0 GPa (100--160 km) by Smyth and Caporuscio (1984), MacGregor and Manton (1986), and Taylor and Neal (1989).

The range in equilibration parameters is a reflection of the diverse origins postulated for mantle-derived eclogites. Currently, the favored viewpoint is that eclogites may represent either subducted underplated oceanic crust or "basaltic" magmas crystallizing at high pressures. Eclogites belonging to both parageneses are represented in the Bellsbank and Roberts Victor eclogite suites. Diamond-bearing eclogites clearly originate from within the diamond stability field, and Hatton and Gurney (1979) have calculated equilibration over a range of 1020-1140DC at 4.2-4.5 GPa (137-147 km) for an example from Roberts Victor. A coesite grospydite from Roberts Victor is determined by Wohletz and Smyth (1984) to have originated at 160-km depth (4.9 GPa) at 1060DC.

Xenoliths, supposedly originating in the transition zone (Moore et al. 1991, Sautter et al. 1991), have not been reported from orangeites. MARID-suite and metasomatized peridotites (Erlank et at. 1987) are also absent.

It has long been known that diamondiferous kimberlites and orangeites are found only in regions underlain by Archean cratonic basement rocks (Clifford 1966). It is now believed that most diamonds are xenocrysts in their transporting magmas and are derived from lithospheric mantle roots where they have been preserved since Archean or Proterozoic times (Richardson et al. 1984, Meyer 1987).

Consequently, a further constraint on the depth of origin of orangeite is provided by inclusions in diamonds. Meyer (1987) and Gurney (1989) reviewed this topic and noted that diamonds contain two suites of silicate inclusions, termed the "peridotitic" (P-type) and "eclogitic" (E-type) suites because of their similarity to minerals comprising peri-

348 CHAPTER 4

dotite and eclogite xenoliths, respectively. However, the inclusions are compositionally distinct from the minerals of these xenoliths. There is no correlation between xenolith suite and type of inclusions in diamonds within a given kimberlite or orangeite. Both suites of inclusions may occur in the total diamond population, although one or the other may be dominant at a particular locality. E-type diamonds may occur in intrusions that are dominated by peridotite xenoliths, and vice versa. Thus, at Roberts Victor, diamonds with P-type inclusions are the most common, yet the xenolith suite is dominated by eclogites. At Finsch, E-type diamonds become more abundant relative to P-type with increasing diamond size.

Currently it is believed that most P-type diamonds are not derived by disaggregation of coarse granular garnet lherzolite protoliths, although E-type diamonds may be derived from material similar to eclogite suite xenoliths. The source of P-type diamonds is now believed to be garnet harzburgites derived from the highly depleted cratonic roots of the continent (Pokhilenko et al. 1977, Nixon et al. 1987).

Equilibration temperatures and pressures calculated for both suites of inclusions have been summarized by Meyer (1987). The most reliable data for P-type inclusions range from about 900-1300°C at 4.5--6.5 GPa (147-212 km). At an assumed pressure of 5 Gpa most E-type inclusions fall in the stability field of diamond. It is important to note that the data imply that both suites of diamonds appear to originate at the base of the lithosphere (Meyer 1987, Gurney 1989).

In summary, xenoliths and diamond inclusion studies suggest that orangeites sample material derived from near the base of the lithosphere within the diamond stability field at depths of 150-200 km. Clearly, the magmas must be derived either from these depths or deeper. The absence of asthenospheric xenoliths precludes the latter. Consequently, it is suggested that the sources of orangeites are located in the cratonic root of the Kaapvaal craton at, or slightly above (50 km), the lithosphere-asthenosphere boundary. This region is considered to consist of ancient depleted garnet harzburgite and garnet dunite, together with eclogites representing ancient recrystallized subducted underplated oceanic crust and ponded crystallized asthenospheric melts.

4.5.1.3. Compositional Heterogeneities-Veined Harzburgites

The development of compositional and isotopic heterogeneities in the lithospheric mantle is currently considered to result from the introduction of incompatible-e1ementrich hydrous fluids or small-volume silicate melts (Hawkes worth et al. 1985, Menzies et al. 1987, Dawson 1984, Haggerty 1989a,b). Commonly, compositional changes resulting from both processes are referred to as metasomatism. However, Hawkesworth et al. (1985) correctly note that infiltration of silicate melts is not in itself metasomatism, although such melts may cause localized metasomatism in adjacent rocks.

The extensive mineralogical and geochemical evidence supporting the existence of metasomatism and melt infiltration processes in the lithospheric mantle has been summarized by Erlank et al. (1987), and Menzies et al. (1987). There is general agreement that the end result of these processes is the production of a wide variety of veins with or without metasomatic aureoles. The veins are believed to form a stockwork within the depleted harzburgitic or peridotitic mantle (Menzies et al. 1987, O'Brien et al. 1991, Foley 1992b).


The mineralogy of the veins is believed to consist primarily of widely varying modal amounts of amphibole (kaersutite or potassium richterite), Ti-phlogopite, apatite, diopside, K-Ba titanates, ilmenite, and rutile. Menzies et al. (1987) suggest that veins generated from hydrous metasomatic fluids consist of glimmerites and MARID-suite material with aureoles of phlogopite- and/or phlogopite-K-richterite peridotite. In contrast, small volume melts are believed to be essentially of basanitic compositions which crystallize to kaersuite, apatite, and mica. Peridotitic wall rocks may be subjected to Fe-Ti-rich metasomatism with the formation of pargasite/kaersutite and high Ti-mica. Carbonates may (Thibault et al. 1992) or may not (Wyllie 1980) be present.

Experimental studies have indicated that most ofthe minerals postulated to be present in the veins are stable at pressures up to and exceeding those encountered at the base of the lithosphere, i.e., 5-6 GPa. Thus, K-richterite is stable up to 14 GPa (Tr0nnes et al. 1988, Sudo and Tatsumi 1990, Foley 1991), phlogopite to 9 GPa and phlogopite plus diopside to 7 GPa at 1200°C (Luth et al. 1993), and K-Ba titanates (hollandites, crichtonites) to 5-6 GPa (Foley et al. 1994). While these experimental studies demonstrate the stability of the incompatible-element -rich minerals of interest at the base of the lithosphere, there are unfortunately few studies of the melting of vein-type assemblages at these pressures (see 4.2.3,4.2.4).

It should be stressed that there is very little quantitative evidence regarding the mineralogy and modal composition of the veins, in particular those postulated at the base of the lithosphere. Vein mineralogy certainly will vary as a function of depth and the source of the infiltrating melts and fluids. Experimental (Thibault et al. 1992) and theoretical (Haggerty 1989a,b) studies suggest that mineral zoning of metasomes will be commonplace. One undesirable, but unavoidable, aspect of this lacuna in our knowledge is that petrologists commonly use the compositions of supposed partial melts to indulge in speculation concerning the source mineralogy. This is unsatisfactory as hypotheses not constrained by experimental data are infinitely flexible with regard to the hypothetical source mineralogy.

All of the veins are considered to have high Rb/Sr and U/Pb ratios and low SmlNd, relative to bulk earth. Undoubtedly, metasomatism and infiltration is not confined to Archean times and must persist to the present day. Consequently, the Sm-Nd model ages of the enrichments range from 0.9 to 2 Ga (Menzies et al. 1987). The ages of the MARID suite and phlogopite K-richterite peridotite have been shown to be 84-87 Ma (Erlank et at. 1987, Kinny and Dawson 1992) and contemporaneous with emplacement of archetypal kimberlites in the Kaapvaal craton.

To explain the isotopic signatures of ultrapotassic lithospheric melts, some of the enriched veins are required to persist in isolation from convecting mantle for periods on the order of 1-2 Ga. This constraint is required to allow mantle domains to develop distinctive enriched isotopic signatures. Note that although Pb isotopic studies (3.8.2) suggest the initial stages of vein formation may have involved two or more episodes of infiltration and metasomatism, they have subsequently remained as closed systems apparently unaffected by Phanerozoic magmatic events.

The antiquity and suitability of the veins considered as a source of ultrapotassic melts have been demonstrated by O'Brien et al. (1991), who determined that a glimmerite vein in a mantle-derived harzburgite from the Wyoming craton was enriched in Ba, Rb, and

350 CHAPTER 4

REE, with an ENd of -33 at 52 Ma and a Nd model age of 2.57 Ga, relative to depleted mantle.

While there is general agreement among petrologists regarding the existence of a stockwork of veins in the lithospheric mantle, there is no agreement as to the ultimate source ofthe fluids and melts from which they are derived. Hypotheses include: devolatilization of underplated subducted material (Sekine and Wyllie 1982, O'Brien et ai, 1991); volatile release from the asthenosphere (Wyllie 1980, Foley 1988); crystallization of asthenospheric basaltic (Hawkes worth et al. 1985, Menzies et al. 1987, Thibault et al. 1992); kimberlitic (Menzies et al. 1987, Kinney and Dawson 1992); or carbonatitic (Haggerty 1989a,b) magmas trapped in the lithosphere or at the lithosphere-asthenosphere boundary.

It is to be expected that metasomatism and melt infiltration will not necessarily have the same character or take place to the same extent everywhere within the lithosphere. Locations may be controlled by pre-exisitng fractures or zones of weakness. The topography of the Kaapvaal lithosphere-asthenosphere boundary is not known. Most representations (Mitchell 1991b, Haggerty 1989a,b, Boyd and Gurney 1986) depict the boundary and craton root as a smooth convex-down keel; however, given the known complexity of geological processes, it is entirely possible that this representation is incorrect. Immense corrugations may exist at the lithosphere-asthenosphere boundary resulting from asthenospheric upwelling, pileup and/or sinking of underplated subducted slabs, or the presence of ancient welded mobile zones oflower-density material, e.g., the Limpopo belt. Such features may channel fluids and melts to particular portions of the lithosphere-asthenosphere boundary.

With regard to orangeites, it is evident that their primary expression is in the southwestern part of the craton. This may result from either preferential modification of the lithosphere underlying the region or selective melting of a portion of an extensively modified craton. With regard to the former it is instructive to note that Tainton (1992) and Smith (1983) consider that, on the basis ofNd isotopic compositions, the last enrichment and closure of the sources of orangeites occurred about 0.7-1 Ga ago and was possibly related to the later stages of the accretion of the N amaqua Province to the Kaapvaal craton (Tainton 1992).

Hence, devolatization of underthrust Namaqua material, according to the models of Sekine and Wyllie (1982), could have restricted the sources of orangeites to the southern margins of the Kaapvaal craton root. Such a model could explain the paucity of orangeites in the northern parts of the craton and their absence in the Zimbabwe craton. However, other explanations of the distribution are possible (4.5.2.1).

In summary, the depleted continental lithosphere is undoubtedly extensively veined by ancient and modern incompatible-element-enriched material. Partial or complete (4.5.2.2) melting of such veins provides suitable sources for incompatible-elementenriched lithospheric magmas. The sources of these veins are, as yet, undetermined and may be related to either asthenospheric or subducted material.

It is concluded that the stages in the production of the sources of orangeite magmas are:


1. Fonnation of the depleted harzburgitic roots of the Kaapvaal craton in the early Archean (Herzberg 1993, Jordan 1978). This buoyant depleted root acts as substrate for subsequent metasomatism and melt infiltration.

2. Emplacement of incompatible element-enriched veins during the interval ca. 3 Ga to 1 Ga. The exact modal mineralogy of the veins is unknown, but they undoubtedly contain potassium richterite, Ti-phlogopite, and apatite. The source of the vein material is unknown, but may be derived from asthenospheric melts contributing to further growth of the cratonic roots and/or subducted oceanic material. Oceanic crust may have been subducted under the Kaapvaal craton during collision of the Kaapvaal and Zimbabwe cratons (Helmstaedt and Schulze 1989) and subsequent to the accretion of the N amaqua Province. Thus, eclogites of two ages may be underplated and incorporated with the cratonic root. Therefore, eclogite xenoliths found in orangeites could conceivably have two provenances (however see 4.8), although only Archean isotopic signatures have so far been found for eclogites from Roberts Victor (Kramers 1979) and Bellsbank (Neal et al. 1990).

3. Stabilization of the craton subsequent to the Namaqua orogeny and geochemical isolation of orangeite sources for 0.5-1 Ga. There were apparently no thennal effects on, or compositional changes to, the sources during the Pan-African orogeny and the assembly of Gondwanaland or its subsequent dismemberment in Jurassic times.

4.5.2. Melting of the Source

Aspects of the composition and mineralogical character of orangeite magma sources may be inferred from experimental studies of orangeites and lamproites (see 4.2), in conjunction with compositional (Chapter 3) and mineralogical (Chapter 2) constraints derived from the study of rocks of the orangeite clan. However, it must be emphasized that the actual composition of any orangeite magma is not known, and thus we have no definitive knowledge of the composition of the initial partial melt or primitive magma. Consequently, hypotheses concerning the source mineralogy or extent of partial melting of that source must be regarded as speculative and preliminary.

Bearing the above in mind, the source should contain C02 and H20 and minerals which sequester alkali, alkaline earth, rare earth, and second-period transitional elements. Derivative melts must be rich in potassium and able to crystallize abundant micas, significant amounts of apatite and carbonate, and accessory titanates and spinels. Genetic models must explain the presence of diopside in some orangeites and the formation of evolved orangeites.

The simplest source having the above attributes is one composed of Ti-phlogopite and a phosphate. The latter could be apatite, a dense polymorph of apatite (Murayama et al. 1986), or a complex phosphate such the K-Ba phosphate (16-20 wt % K20, 19-21 wt % BaO, 6 wt % CaO, 9 wt % MgO), which has recently been recognized by Mitchell and Edgar (unpublished data) in 5-7 GPa near-solidus experimental studies oflamproites.

Whether the presence of other minerals is necessary, is unknown. One probable candidate which may playa particular role in the fonnation of evolved orangeites is

352 CHAPTER 4

potassium richterite, as it may contribute Ti and silica-rich melt to the magma. Foley (1991) and Trl1lnnes et al.(l988) have shown this amphibole to be stable throughout the lithosphere. Luth et al. (1993) have noted that potassium richterite is stable to greater depths than phlogopite, and, during subduction may incorporate all of the H, K, and large-ion lithophile elements released by phlogopite breakdown at about 200-km depth. Subsequent richterite breakdown, perhaps near the top of the transition zone, may cause partial melting and release of fluids and volatiles into the craton and cause reprecipitation of potassium richterite and/or phlogopite. The common occurrence of potassium richterite, together with phlogopite, in metasomatized peridotite and MARID-suite xenoliths supports the contention that this amphibole is a typical component of the vein assemblage.

Experimental studies of the systems phlogopite-apatite, K-richterite-apatite, and K-richterite-phlogopite (see 4.2.4), clearly demonstrate that veins composed of Ti-K richterite, Ti-phlogopite, and apatite will have low solidus temperatures. Melting of these assemblages will produce K~rich melts with the general compositional characteristics ascribed to orangeite magmas. Unfortunately, experimental studies of this ternary system at 5-7 GPa have not been undertaken.

The liquidus studies of lamproites and the melting of phlogopite clinopyroxenite suggest that olivine is not present in the sources of orangeites. Whether diopside occurs in the veins is unknown, as the breakdown of richterite will supply all of the components required for the crystallization of diopside from the melt at lower pressures.

The presence of carbonate in orangeites requires a source of C02. It is unknown whether this exists as solid carbonate in the veins or is subsequently introduced during the events which initiated partial melting of the veins. Both sources are, of course, possible.

It is concluded here that the incompatible element -rich veins proposed as the sources of orangeite magmas consist of varying modal amounts of Ti-K richterite, Ti-phlogopite, and apatite. Diopside and carbonate mayor may not be present. Ilmenite, rutile, hollandites, crichtonites, and wadeite are not required in the source as contributors ofTi, Ba. and K to the melt as they are unlikely to take part in near-solidus melting reactions. The absence of negative Ti-Nb-Ta anomalies in extended incompatible-element distribution diagrams for orangeites is evidence that these minerals are not present as modally significant residual phases and. hence. were unlikely to have been major components of the initial assemblage.

Production of melts from veins in a harzburgitic substrate may be considered with respect to the causes and mechanism of melting.

4.5.2.1. Causes of Melting

Partial melting of the source may be initiated by any of the following processes (Mitchell and Bergman 1991):

1. Simple thermally induced melting as ambient temperatures exceed those of the solidus. No new components are introduced. and the veins melt directly at the fringes of large mantle plumes.


2. Decompressional melting resulting from uplift, with no new components being introduced. Uplift may be associated with mantle plumes or permissive plate tectonics.

3. Introduction of volatiles, with melting resulting from reduction of the solidus temperatures. The volatile influx may result from degassing of the Earth and ascent of reduced fluids (Foley 1988, 1989, Taylor and Green 1989) or the thermally induced decomposition of subducted hydrous phases at great depth in the mantle (Ringwood 1989, Wyllie 1980).

4. Introduction of volatiles and other components. In this case the volatile flux contains alkali and other incompatible elements. These may be C02- and H20-rich limited partial melts derived from asthenospheric sources, i.e., they are similar to the melts which caused the original enrichment of harzburgitic mantle. This case may be considered as a further enrichment event.

Any combination of these processes may initiate partial melting. Processes 2 and 3 may be interrelated, as Bailey (1983) maintains that the introduction of volatiles will lead to the formation of low-density minerals and ultimately to uplift of the lithosphere. Process 4 has very important consequences given the possibility of modification of the isotopic composition of the source (see below).

Processes I and 2 suggest melting may occur at the fringes ofthe large mantle plumes which are currently believed to cause continental breakup and generation of continental flood basalt provinces. The characteristics of such plumes and how they interact with the lithosphere are reviewed by White and McKenzie (1989), Saunders et al. (1992), and Anderson et ai. (1992).

Large plumes are believed to result from the upwelling of hot asthenospheric material from the lower mantle. As the plume impinges upon the lithosphere it causes uplift, thinning, and, ultimately, rifting. Igneous rocks are generated by decompressional melting as the mantle rises passively beneath the stretched and thinned lithosphere. Plumes at the lithosphere-asthenosphere boundary are believed to spread out as very large mushroomshaped bodies, typically 2000 km in diameter (White and McKenzie 1989). In this region temperatures are raised 100-200°C above normal. The combined effects of uplift and enhanced temperature are believed to be the principal cause of the formation of the vast quantities of flood basalt associated with continental rifting.

With regard to continental lithospheric magmatism, interest in plumes is related to the elevated temperatures at, and above, the lithosphere-asthenosphere boundary resulting from the mushrooming of the plume head. Thus, partial melting may be induced in the lithospheric mantle above the head, or at the fringes, of a plume if ambient temperatures are raised above those of material with low solidus temperatures. Melting of lithosphere above plume heads is currently a popular means of generating the enriched geochemical signatures of continental flood basalts (Ellam and Cox 1991, Mahoney et at. 1989, Hawskesworth et at. 1986).

Melting at the fringes of plumes might be important for the generation of small volume magmas such as orangeites, as only small increases in temperature, coupled with slight uplift, may be sufficient to achieve suprasolidus conditions. Thus, with regard to orangeites, melting of the source, in response to the uprise of large mantle plumes, is

354 CHAPTER 4

plausible, although it is shown in Section 4.4.3 that orangeites are not related to small hot spots. Two plumes are required to explain the distribution and range in age of orangeites. The thermal effects associated with one of these plumes may also explain the absence of orangeites in many parts of southeastern Africa.

The genesis of the Late Jurassic orangeites (145-165 Ma) may be related to the breakup of Gondwana which was initiated in the mid-Jurassic (170 Ma), consequent on impingement of a large mantle plume on the eastern margin of the Kaapvaal-Zimbabwe craton (Cox 1992, White and Mckenzie 1989). This plume is believed to be the source of the vast quantities of Karroo basalts erupted in southeastern Africa. Recent geochemical studies have concluded that many of the basalts are contaminated with lithospheric mantle (Cox 1992, Ellam and Cox 1991). The contaminant has been identified by Ellam and Cox (1991) as having isotopic and trace element characteristics similar to those oflamproites, and is believed to be derived from metasomatized lithospheric mantle sources. Clearly, this component could also be identified with the sources of orangeites.

Consequently, it is suggested that the absence of eitherorangeites or lamproites from the northern Kaapvaal and Zimbabwe cratons is due to the destruction of their potential sources ahead of the rising Karroo plume. The small amounts of partial melt produced could easily be hybridized in voluminous basalt magmas without significant effects upon their major element composition, while simultaneously changing their incompatible element and isotopic abundances.

The late Jurassic orangeite magmas may be formed in the mid-Jurassic at the fringes of the plume and remain as melts trapped in the mantle or late in the breakup process of Gondwanaland as a consequence of slow conduction of heat to the outer fringes of the plume. The former seems geologically unreasonable as it requires melts to exist for 25 Ma in the mantle and some further process to initiate intrusion. However, melt persistence may be possible if solidus temperatures are low and the region is held at higher temperatures during the existence of the plume.

Alternatively, the late Jurassic orangeites may have no direct relationship with the Karroo plume and represent melts derived from remnants of original enriched mantle which survived destruction because of their location at the fringes of the plume. In this scenario, melting occurred subsequent to the decay of the Karroo plume and was initiated by other processes. The latter must have been localized, otherwise orangeites of this age would have been formed elsewhere in the craton.

It has been previously noted (Fraser 1987, Skinner 1989) that the main focus of orangeite magmatism at the southwestern margin of the Kaapvaal craton is contemporaneous with the opening of the South Atlantic and the eruption of the Parana and Etendeka flood basalts. Figure 4.15 shows that orangeites are emplaced near the outer fringes of the abnormally hot mantle associated with the Walvis (White and McKenzie 1989) or Parana plume (Cox 1992). This association and the coincidence in ages cannot be fortuitous and implies some connection between the onset of melting of the orangeite sources and uprise of the plume. Partial melting of orangeite sources could be in response to a combination of thermal effects and uplift at the fringes of the Wavis plume or be entirely decompressional and in response to plume-related uplift outside the region of enhanced temperatures. This scenario of plume-related magmatism does not involve the addition of asthenospheric material to orangeite sources at the time of partial melting.


o !

I South

2000 km !

355

• ORANGEITES

Figure 4.15. Reconstruction of the South Atlantic at 120 Ma, shortly after the onset of seafloor spreading. showing the distribution of Lower Cretaceous orangeites (after Skinner 1989) relative to the region of abnormally hot mantle (dotted) above the Walvis or Parana plume (after White and McKenzie 1989). Solid black areas show areas of extrusive basalts.

This constraint is apparently not violated for the Postmasburg, Barkly West, Winburg, Kronstad, and Boshof orangeites, but may not be satisfactory in accounting for anomalous Preiska area orangeites (see below).

In direct opposition to the above, Anderson et al. (1992) have concluded from seismic tomographic studies that there is no evidence for the involvement of plume heads or lithospheric modification in the genesis of the Karroo and Parana-Etendeka basalts. Instead, these flood basalt provinces are considered to be related to plate tectonic

356 CHAPTER 4

reorganization over broad deep regions of hot, probably partially-molten mantle. Pieces of thick continental lithosphere are believed to separate along pre-existing lines of weakness without substantial rifting or thinning. Buoyant upwelling of the underlying hot regions leads to further decompressional melting and production of voluminous basaltic magmas. Clearly, if Anderson et al. (1992) are correct, plumes cannot play any role in the genesis of orangeites.

Determining whether orangeites contain an asthenospheric component (process 4) is extremely difficult, and hypotheses in favor of, or against, commonly hinge upon particular subjective interpretations of isotopic compositions (see 3.8). Thus, the isotopic composition of orangeites may be interpreted to represent:

1. Those of the sources with no mixing of individual batches of lithospheric melt and no asthenospheric contributions.

2. Mixing of individual batches of lithospheric melt derived from veins with slightly different time-integrated isotopic compositions and no asthenospheric component.

3. Mixtures of asthenospheric and lithospheric melts. In this case the isotopic composition of the melts does not represent that of the source. Potential end members might include depleted asthenospheric MORB-type sources and a highly enriched lithospheric end member with lower 143Ndll44Nd ratios and higher 87Sr/86 ratios than observed in orangeites.

Unfortunately, all of the above can account for the intra- and inter-intrusion isotopic variations found in orangeites, although there is no compelling geological or mineralogical evidence for the existence of an asthenospheric component. Contemporaneous asthenospheric magmatism is apparently absent or restricted to younger Upper Cretaceous kimberlites and melilitites. The presence of mineralogically anomalous orangeites in the Prieska area (1.8.8) in itself does not indicate the involvement of an asthenospheric component in their genesis, as the ilmenite, gamet, and pyroxene macrocrysts could well be derived from an earlier period of kimberlitic or other type of magmatism. Note in particular that the Sr and Nd isotopic compositions of these orangeites are similar to those of other orangeites lacking macrocrysts (Figures 3.30, 3.31). Hence, even if contamination/mixing has occurred, it cannot have played a significant role in determining the isotopic compositions of the Prieska orangeites.

The isotopically and mineralogically anomalous rocks of domain V of the Preiska region (1.8.8) may represent archetypal kimberlite magmas contaminated by lithospheric material whose genesis is totally unrelated to the geographically associated younger orangeites and kimberlites. They certainly do not represent rocks derived from magmas which are "transitional" between orangeites and kimberlite magmas, as suggested by Skinner et at. (1994).

Addition of volatiles to lithospheric sources is a possible means of inducing partial melting (Wyllie 1980, Bailey 1983, Foley 1988). However, this process commonly begs the question as to the ultimate source of the volatiles, and appeals must be made to further processes, which are impossible to verify, occurring in the transition zone, the lower mantle, or at the lithosphere-asthenosphere boundary above subducted oceanic crust. It is possible, but not highly probable, that the asthenospheric events giving rise to


kimberlites may be preceded by the uprise of volatiles which add H20 and C02 to orangeite sources and initiate partial melting. This scenario requires that asthenospheric uprise and kimberlite segregation lag 20-25 Ma behind the ascent of the volatiles. The hypothesis fails to explain the general absence of a temporal association between ultrapotassic and kimberlite magmatism.

Bailey (1983, 1984, 1992, 1993) has consistently argued that continental lithospheric magmatism occurs in response to plate tectonic events such as collisions and changes in plate motion patterns. Thus, episodic intraplate magmatism is a consequence of external forces acting laterally across the lithosphere rather than being initiated by mantle plumes. Melt generation occurs in response to stress release in the lithosphere and/or volatile fluxing.

The volatile fluxing hypothesis (Bailey 1984, 1992, 1993) requires the addition to the lithosphere of volatiles derived from the deep mantle, a condition seemingly at variance with Bailey's stance that asthenospheric activity plays no role in intraplate magmatism. Regardless, Bailey's (1984, 1993) model assumes that for long periods the lithosphere acts as a lid which permits only the slow leakage of volatiles from the asthenosphere. Fractures developed or reactivated by tectonic processes external to the craton may penetrate to the base of the lithosphere and provide channels for the escape of the volatiles. These channels act as a focus for drawing volatiles and heat to one region of the lithosphere-asthenosphere boundary. The process results in the build up, by metasomatism, of the levels of volatile-bearing phases in regions adjacent to the channels. Ultimately, melting of these regions will occur due to decompression consequent upon uplift (Bailey 1984). The type of magma erupted depends on the local geothermal gradient (Bailey 1993). Thus, kimberlites, lamproites, and carbonatites lacking associated silicate melts would characterize cooler older lithosphere, and provinces with larger volumes of alkaline silicate melts would be expected to occur in regions with higher geothermal gradients, i.e., old mobile belts and rift zones. An important aspect of Bailey's (1993) version of this model is that magmatism must be contemporaneous across the whole of a tectonic plate. Accordingly, Bailey considers that Cretaceous kimberlitic and other magmatism in Africa can be correlated with collision of the African plate with Europe. Orangeites are believed by Bailey (1993) to be contemporaneous with other varieties of platewide magmatic activity, and the apparent.age trend reported by Skinner (1989) to be due to progressive structural readjustment of the craton as it drifts northward subsequent to the breakup of Gondwanaland.

Bailey's hypothesis is important in that it attempts to place all late Phanerozic magmatism in a single model; however, it fails as it requires that kimberlites, orangeites, lamproites, melilitites, and nephelinites all be generated from the same source in the lithosphere. Unfortunately, Bailey consistently ignores all the accumulated mineralogical, geochemical, and isotopic evidence pointing to the derivation of these diverse magmas from different sources and depths. Moreover, the model is specific to late Phanerozoic African magmatism, and even the proponent has not investigated whether the model has any general worldwide applicability.

With regard to orangeites, the model is unsatisfactory as, if Bailey (1993) is correct, it would be expected that the low geothermal gradients found in other cratonic regions of the African plate should result in the formation, in them, of contemporaneous

358 CHAYfER4

orangeites. However, orangeites are absent from these cratons, and other examples of African ultrapotassic magmatism, such as the Quaternary-Recent kamafugites of the Western Rift and the 220 Ma Kapamba lamproites of the Luangwa graben (Scott Smith et al. 1989), are not contemporaneous.

In conclusion, it is apparent that the sources of orangeites are located in the lithosphere, but whether melting is in response to passive or active tectonism is unknown. The contemporaneity of the magmatism with continental breakup and the eruption of flood basalt provinces is suggestive of some relationship with large asthenospheric plumes. The latter may cause uplift and heating of orangeite source regions, but do not apparently contribute asthenospheric material to the melts. Of course, there is no a priori reason why all orangeite magmas should be generated by a single process, and attempts to seek a definitive relationship to a particular tectonic mechanism may be inevitably doomed to failure.

4.5.2.2. Melting of Veined Lithosphere Foley (1992b) has reviewed the physical aspects of melt generation from a veined

lithospheric mantle and developed an important model of the melting process, with particular relevance to the genesis of ultrapotassic rocks. The model is directly applicable to the genesis of orangeites if they are formed by the partial melting of richterite-phlogopiteapatite veins in a harzburgitic substrate.

Foley (1992b) introduced his vein-plus-wall rock melting mechanism to explain the paradox of the failure ofliquidus experiments on ultrapotassic rocks to locate a peridotitelike multiphase saturation point (Foley 1992a), and the conclusion derived from geochemical (Mitchell and Bergman 1991) and other experimental studies (Edgar 1987, Foley 1990) that these magmas might be derived from an incompatible element-rich mica harzburgite.

Foley (1992b) notes that veins within a peridotitic or harzburgite mantle will have a lower solidus temperature than their host, due to the high concentration of volatiles and incompatible elements. Despite this difference, during progressive melting the veins and host do not behave as distinct systems because an initial strongly alkaline melt, originating exclusively from the vein assemblage, becomes progressively diluted by a peridotitic component derived from the wall rocks. Hence, partial melts will be hybrids with compositions dependent on the relative proportions of melts derived from the vein and wall rock.

When partial melting begins, it will be centered on the vein assemblage. Some accessory phases may be quickly eliminated, but the major minerals will remain stable due to solid solution melting reactions. These reactions occur wherever a mineral is an intermediate member of a solid solution series between two end members with very different melting points. In this context, Foley (1991) has shown that the breakdown temperatures of hydroxy- and fluor-amphibole differ by about 200°C (Figure 4.16). Similar differences in stability may be expected for apatite and mica as a function of their F/OH ratios. Hence, as minerals belonging to complex solid solution series melt, the compositions of the derivative liquids and residual minerals will continuously change. Further, the melting interval of a particular mineral may extend to temperatures above the solidus of the surrounding wall rocks. Consequently, this phase will remain stable while

PETROGENESIS OFORANGEITES AND KIMBERLITES

... C

.Q

60

~ 50

IJ..I a:: :::> 40 (f) (f)

IJ..I a:: a. 30

K - richterite F /OH

1000 1200 1400

TEMPERATURE (OC)

359

Figure 4.16. Relative stabilities of hydroxy- and fluor-potassiumrichterite at high pressures (after Foley 1991).

the vein-derived melt is hybridized by melt derived from the low-melting components of the wall rocks. Figure 4 .17 graphically illustrates Foley's (1992b) vein-wall rock melting process at various stages of melting. An important conclusion of this model is that, at advanced stages of melting, only fluor-phlogopite persists from the original vein assemblage, although olivine and orthopyroxene remain in the wall rock. Hence, melting experiments on a rock derived from a melt produced at this stage of the melting process would find only mica, olivine, and orthopyroxene as near-liquidus phases. Hence, the experimentalist might incorrectly conclude that the melt was derived by the partial melting of a homogeneous mica harzburgite. Foley (1992b) stresses that in this example neither the vein nor the wall rock consist of mica harzburgite.

Foley's (1992b) model also postulates that melt infiltrates the wall rock and begins to dissolve minerals such as olivine and orthopyroxene. Assimilation of these minerals will add a refractory wall rock component to the melt, which is compositionally very different from that derived from solid solution melting of pyroxenes and garnets. The extent of assimilation is very difficult to predict, and while Foley (1992b) does not discuss the thermodynamics of assimilation, he does note that the incorporation of refractory components and dissolution in vein-melt may commence at temperatures below the peridotite solidus. It is suggested here that small volumes of melt will be unable to assimilate much olivine or orthopyroxene without causing complete crystallization of the vein-melt. However, minerals separated from the wall rock by infiltration and porous flow of vein-melt may be mechanically incorporated into the melt during segregation and be transported.

In summary, the compositions of alkaline melts produced by vein-wall rock melting are determined by the relative proportions of:

• Melts derived from the vein assemblage (V-component) • Melts derived from the low-melting components of the wall rocks (W-component)

VEINED GARNET LHERZOLITE: UNMELTED

INTERMEDIATE STAGE OF MELTING: WALL - ROCK MELTING BEGINS

LOW - DEGREE MELTING : VEIN ONLY

ADVANCED MELTING : ONLY MICA REMAINS FROM THE VEIN ASSEMBLAGE

.... . OLIVINE

~ GARNET

_MELT

~ ORTHOPYROXENE

~ AMPHIBOLE

~ CARBONATE

j::/?/::.'I APATITE ..... ',.

PHLOGOPITE ~~;~~ CLINOPYROXENE , "

D RUTILE, ALKALI- TITANATE


• A component derived from assimilation of refractory wall rock minerals (WXcomponent)

The partial melting process results in a series of hybridized melts. Initial melts will be dominated by the vein component and have high V/(W+WX) ratios, whereas later melts will have low V/(W+WX) ratios with W>>>WX.

Foley (1992b) correctly emphasizes that the concepts of low degree and high degree of partial melting are a matter of scale. Thus, a strongly alkaline melt may be produced primarily from the vein material [very high V /(W +WX) ratios]. The scale of melting with respect to the vein is obviously very high (>50%), but low when compared to a system consisting of a few veins (1-5%) in a substrate. This observation reinforces the conclusion, advanced in this work, that concepts developed for the partial melting of homogeneous sources, and appropriate for basaltic magmatism, are inapplicable to the genesis of exotic alkaline rocks.

Foley's (l992b) model is directly applicable to orangeites. Partial melting of veins of potassium richterite, phlogopite, and apatite (± clinopyroxene) in a harzburgitic or peridotitic substrate will melt according to the vein-plus-wall rock melting mechanism. It is suggested here that all orangeite melts have high V/(W+WX) ratios; i.e., they are high-degree partial melts with respect to the vein systems. Differing modal assemblages, mineral compositions, and V/(W+WX) ratios may result in melts which differ slightly in bulk composition. Incorporation of refractory wall rock minerals might account for the presence of macrocrystal olivine in many orangeites. Differences in their macrocryst contents could result from the extent to which vein-melts infiltrate wall rocks. However, other processes, such as flow differentiation and/or gravity fractionation, can also explain these variations (see 4.5.3).

While vein-plus-wall rock melting is an attractive process for generating orangeites, it does have some limitations, especially with respect to the generation of diopside orangeites and sanidine richterite orangeites. This is because the model is not constrained and is infinitely flexible with regard to the mode of the veins. Thus, diopside orangeites may be generated either by preferential melting of richterite or veins containing diopside. Mineralogical evidence (Chapter 2) suggests that sanidine richterite orangeites are unlikely to be derived from primitive melts. However, melting of veins rich in richterite and diopside might conceivably give rise to relatively silica-rich magmas which may differentiate at low pressures to evolved sanidine richterite orangeites.

One further major problem concerns the role of orthopyroxene. Macrocrysts of orthopyroxene would be expected to be common in orangeites if olivine macrocrysts are derived from harzburgitic protoliths. Their complete absence requires that orthopyroxenes be completely dissolved, the protolith be dunite, or wall rock orthopyroxenes not be incorporated into the melt. Dissolution of orthopyroxene would generate silica-rich liquids, but the extent of assimilation would require large volumes of magma and/or superheat and is considered unlikely. A dunite protolith provides a satisfactory solution to the problem, but requires that the diamond-bearing rocks sampled by orangeites during their ascent to the crust, also be garnet dunites (see 4.8).

Further options are that either V/(W+WX) ratios are extremely high (>10,000) (i.e., essentially only the vein melts and macrocrysts are not derived from the residue of partial

362 CHAPTER 4

melting) or V IW is high and WX=O (Le., garnet and clinopyroxene in the wall rock melt with no addition of a refractory component). However, the absence of orthopyroxene still requires an explanation if macrocrystal olivine is derived from harzburgitic or peridotitic contaminants during ascent of the magma (see 4.5.3).

In conclusion, orangeite primary magmas of unknown composition are probably derived by high degrees of, or even complete, melting of richterite-phlogopite-apatite veins in a dunitic substrate. Note that the modal amounts of these minerals in the veins are also unknown.

The vein-wall rock melting mechanism of Foley (1992b) also provides an explanation of the significant negative K, Rb, and Sr anomalies found in extended incompatibleelement distribution diagrams for orangeites (3.6.1) in that, because of solid solution melting reactions, residual apatite and richterite may remain in the mantle sources.

4.5.3. Melt Segregation, Contamination, and Ascent

This stage of orangeite evolution is one of the least understood as there is very little experimental evidence pertaining to the physical processes involved in magma segregation and transit through the lithosphere into the crust. Consequently, much of the discussion which follows is highly speculative.

The formation and segregation of partial melts in the mantle has been reviewed by McKenzie (1985, 1989), Hunter and McKenzie (1989), McKenzie and Bickle (1988), Sleep (1988), Watson and Brennan (1987), and Foley (1992b).1t is concluded that melts may exist in either porous or channeled flow regimes. In the former, melts are interconnected along grain edges and form a three-dimensional network on an intergranular scale. In the channeled flow regime, melt exists as a discrete phase, filling cracks and veins.

Sleep (1988) and Foley (1992b) note that the transition between the two flow regimes is not abrupt and depends upon the scale at which a process is considered. Channeled flow dominates during the solidification of mantle veins. Porous flow may predominate during the initial stages of remelting, whereas channeled flow may prevail during later stages of melting and ascent.

It is believed that the lithosphere is dominantly a regime of channeled flow and the asthenosphere one of porous flow (McKenzie 1989, Foley 1992b). One consequence of this assumption is that the underside of the channeled flow regime may be intensively veined by asthenosphere-derived melts. This is because small volumes of melt, emanating from the asthenospheric porous flow regime, will be unable to progress very far into the lithosphere before solidifying due to lack of heat (Spera 1984).

One interesting conclusion of McKenzie's (1989) and Hunter and McKenzie's (1989) analysis of melting is that melts formed by extremely small degrees of partial melting «2%) may be extracted from the mantle. This observation is considered here to be relevant to the extraction of kimberlitic melts from asthenospheric sources, but not to orangeites from lithospheric sources (see 4.6.1).

High degrees (>50%) of melting of mantle veins in a dunitic substrate will give rise to liquids which ascend by channeled flow through the brittle lithosphere. At their source, porous flow may lead to infiltration of melt into the wall rocks, followed by disaggrega-


tion and incorporation of xenocrystal olivine into the melt. Because of the nature of the melting process, xenoliths of restite are not extracted from the source regions of the melt.

Melts derived from isotopically and modally distinct individual veins may coalesce into larger batches of magma capable of transporting xenoliths of eclogite, diamond-bearing eclogite, and diamond-bearing garnet dunite (or harzburgitic) rocks. Disaggregation of the latter will add, in addition to subcalcic chrome pyrope and diamond, a further suite of olivine xenocrysts to the melt. Thus, the macrocrystal olivine suite may consist of olivine derived from at least two sources. The reasons for the characteristic, apparently near-total, disaggregation of diamond-bearing ultramafic xenoliths in both orangeites and kimberlites remains unknown. The presence in orangeites of diamonds containing eclogitesuite inclusions clearly demonstrates that some eclogites are also disaggregated; however, the mechanism of disintegration and the fate of the silicate components in the orangeite magma are unknown.

The paucity of granular garnet peridotites and lower crustal xenoliths in orangeites suggests that there is no significant interaction between the magma and wall rocks of the conduit in the upper lithosphere and lower crust. This may reflect differences in the rheology and ascent rates of the magma in the upper lithosphere relative to those prevailing at the source. Unfortunately, we have no knowledge of these parameters. It is unlikely that upper lithosphere material was sampled and subsequently fractionated from the magma, as such a process should also remove diamond and xenocrystal pyrope. On the basis of the foregoing it is provisionally concluded that all types of xenolithic material found in orangeites are derived from locations close to the lithosphere-asthenosphere boundary and in the immediate vicinity of the sources of the magma.

Orangeite magmas will lose heat during ascent and begin to crystallize. We do not know any of the actual liquidus phases, but may surmise that the first and most important is phlogopite. Evidence for this conclusion is found in the presence of macrocrystal mica and overall high modal contents of orangeites. Olivine is unlikely to be an important liquidus phase at high-to-moderate pressures because bulk compositions of the liquids probably fall outside the primary phase volumes of olivine in ultrapotassic systems. At lower pressures, the reaction of olivine with liquid to form phlogopite suggests that olivine is more likely to be assimilated than to crystallize (see below).

Finally, nothing is known of the xenolithlxenocryst content, extent of crystallization, viscosity, volatile content, or oxygen fugacity of the magma during this stage of its evolution. Without knowing these physical parameters it is impossible to predict ascent rates and how the magma might interact with its surroundings.

The most important geochemical aspect of this stage of evolution is contamination of the primary magma by the addition of the olivine xenocrysts which compose the macrocryst suite. It is envisioned that, during ascent, the orangeite magma is a slurry consisting of xenocrysts and phlogopite macrocrysts-phenocrysts suspended in a volatile-rich ultrapotassic fluid. The xenocrystlphenocryst ratio will undoubtedly vary considerably during ascent due to flow differentiation, elutriation, and increasing amounts of phlogopite crystallization in the upper lithosphere and lower crust.

Because the magma is a slurry, it is unlikely that ascent rates will be extraordinarily high. Skinner (1989) suggests ascent times of several hours to a few days. Emplacement of orangeites as dike swarms indicates that these magmas are rheologically very similar

364 CHAPTER 4

to other ultrabasic magmas and lamprophyric magmas. Magma is thus expected to advance through the mantle by crack propagation (Anderson 1979) and to seek out zones of weakness and/or preexisting fracture systems in the lower and upper crust. Magma may flow considerable distances laterally at crustal levels in dike systems. Orangeites (and kimberlites) certainly do not ascend directly from their mantle sources in cylindrical conduits as fluidized high-velocity intrusions, as suggested by Wyllie (1980) and McGetchin (1968).

4.5.4. Low-Pressure and Post-emplacement Crystallization

At the current level of erosion the majority of orangeites occur as hypabyssal dike swarms and "blows," the latter being the lower parts of the root zones. Diatremes and the upper levels of root zones are preserved only in the Finsch-Postmasburg area. The style of intrusion is identical to that observed for kimberlites, and diatremes are believed to form above precursor dikes and root zones (see 1.9.1). Thus, it is highly probable that diatremes and craters were originally present above the dike swarms.

In areas which have undergone similar amounts of postemplacement erosion, e.g., the Barkly West-Kimberley-Boshof region, there are major differences in the facies of geographically associated kimberlites and orangeites exposed at the current level of erosion. Kimberlites are predominantly diatreme and upper root-zone facies, whereas orangeites are hypabyssal dikes and blows. This observation implies that, for these two magma types, there are significant differences in the depths at which diatremes originate. If diatremes are phreatomagmatic, this may indicate significant changes in the hydrology of the region between the Lower and Upper Cretaceous. Alternatively, the depth of diatreme initiation may be related to the level at which devolatization of the magma occurs. Mitchell and Bergman (1991) have noted that C02 is much less soluble than H20 in silicate melts, therefore, exsolution of a C02-rich fluid phase is expected to occur at relatively greater depths than for a H20-dominated melt. Thus, it is suggested that dikes of orangeite magma penetrate to much higher levels in the crust than kimberlite dikes before volatile exsolution occurs, as a consequence of the higher H20 content of the former.

Petrographic studies of orangeites indicate that the majority are emplaced as slurries of crystal-rich magmas. Typically, they consist of olivine macrocrysts and closely-packed aggregates of phlogopite set in groundmass which represents the minor former fluid component. The complex zoning and inter-crystal compositional variation found in macrocrystal and phenocrystal mica popUlations demonstrates that most of these micas have not crystallized in situ. The simplest interpretation of these observations is that the micas represent the products of crystallization of several batches of orangeite magma of broadly similar composition. Incorporation of crystals from one batch of magma into another will result in the development of epitaxial mantles, which represent the composition of the current liquidus phlogopite in the hybrid magma. Concentration of crystals from different batches of magma at slightly different stages of crystallization, together with batch mixing and hybridization ofthe magmas results, in the observed heterogeneous mica popUlation.


Support for this process is found in the observation that many orangeite dikes are composite (see 1.8), each phase being composed of modally different assemblages of similar macrocrysts and phenocrysts. Such composite dikes are particularly common at Star and Swartruggens. It is probable that dike systems remained as open conduits in which mixing of many batches of partially-crystallized magma has led to a multiplicity of hybrid magmas of broadly similar composition. Flow differentiation may result in further modal variations as mica and olivine are concentrated in different portions of the dikes. It is unlikely that flow rates will be constant, and periods of stagnation may permit the flotation of mica and the settling of olivine. The resulting mass of mica may be subsequently swept into another part of the dike system by a subsequent pulse of incoming magma, where it crystallizes as an olivine-free orangeite. From the above discussion it is concluded that the petrographic differences between orangeites and macrocrystal orangeites result entirely from flow differentiation processes.

Although some orangeites (e.g., Roberts Victor, Star) contain minor amounts of apparently primary olivine, it is considered in this work that olivine is not a significant liquidus phase during the later stages of ascent and post-emplacement crystallization of orangeites (2.3.1, 4.5.3). This is in marked contrast to kimberlites, in which primary olivine is a major component of the groundmass.

While Skinner (1989) has described mantled and zoned olivines from orangeites in the Boshof area, these do not appear to be characteristic of all orangeite fields. The mantling and zoning demonstrate that these olivines are not in equilibrium with their host magmas. This olivine population probably formed prior to emplacement, the mantles representing overgrowths upon xenocrystal substrates. Precipitation of olivine may be a consequence of the bulk composition of these particular hybrid melts being driven into the primary phase volume of olivine for this system. This may result from the assimilation of xenocrystal olivine. Alternatively, the low-pressure incongruent dissolution of orthopyroxene may cause olivine precipitation. However, given the absence of orthopyroxene xenocrysts, this latter process is considered unlikely.

The restriction of diopside-bearing orangeites to particular fields suggests that the compositional characteristics of the melt leading to diopside precipitation are imposed at the source of the magma. The presence in some orangeites of resorbed diopside crystals indicates instability in the transporting melt, and such diopsides must have formed prior to the emplacement of their current hosts.

Subsequent to intrusion of the slurry, the melt fraction crystallized phlogopite as microphenocrysts, phlogopite-tetraferriphlogopite solid solutions as mantles on all preexisting micas, and tetraferriphlogopite as a ground mass phase. Apatite, spinel, titanates, perovskite, and other accessory minerals crystallized contemporaneously, and the residual dregs of liquid formed carbonates and serpentine in the mesostasis. Intra- and inter-intrusion mineralogical variations reflect minor differences in the bulk composition of the melt and intensive parameters during cooling.

Evolved orangeites have no simple relationship to orangeites, although the absence of diamond, paucity in xenocrystal olivine, assemblage of primary minerals, and their evolutionary trends of composition suggest they are differentiates of less-evolved orangeite magmas. Unfortunately, where and how differentiation occurred cannot be determined as there is no geological evidence for the existence of magma chambers. An

366 CHAPfER4

equally unsatisfactory alternative is to consider their sources as being mineralogically different from those of orangeites, and silica enrichment is not the result of differentiation. In this scenario, evolved orangeites are melts whose silica-rich, C02-poor character is imposed at the source, perhaps by partial melting of richterite--diopside--apatite veins rather than phlogopite-apatite--carbonate veins. The answers to these problems will only be resolved by further study.

As a final point note that olivine xenocrysts are unstable in evolved orangeite magmas, as demonstrated by the presence of "dog's tooth" resorption morphologies and reaction rims of phlogopite around xenocrysts. Such features are not observed in orangeites, although olivine macrocrysts are commonly rounded and embayed in a manner suggestive of resorption. This might have occurred during transportation in the upper mantle. The differences in olivine morphology may be related to cooling rates, implying that crystallization of orangeite magma is too rapid to permit extensive reaction of olivine with the melt to occur. The formation of phlogopite reaction rims around olivine is predicted by Luth's (1967) study of the system kalsiite--forsterite-silica-H20 and is also observed in some lamproites (Carmichael 1967a, Mitchell and Bergman 1991).

4.5.5. Summary

Orangeites might be derived from phlogopite--potassium richterite--apatite--carbonate, incompatible element-rich veins within a dunitic substrate located at or above the lithosphere-asthenosphere boundary. The veins persist, in isolation from convecting lithosphere from about 1 Ga until the late Jurassic and early Cretaceous. At this time partial melting is initiated by undetermined processes. Ultrapotassic melts are produced by a vein-plus-wall rock melting mechanism and are extensively contaminated by mantle-derived olivine xenocrysts. The contaminated magma ascends to the crust as a slurry. Crystallization of phenocrystal phlogopite occurs in the lower crust. Magma mixing, flow differentiation, and low-pressure crystallization of abundant microphenocrystal phlogopite gives rise to a spectrum of hybrid crystal-rich melts which are emplaced as a system of dikes. Diatremes form above these dikes, at high levels in the crust, by undetermined processes. Evolved sanidine richterite orangeites may be either differentiates of orangeites or derived from mineralogically different sources.

Thus, outstanding problems remain with regard to the origin of the veins .and nature of the processes which initiate the partial melting events leading to the generation of orangeite magmas.

4.6. PETROGENESIS OF THE KIMBERLITE CLAN

Mitchell (1986) reviewed earlier hypotheses for the genesis of kimberlites and concluded that none was entirely satisfactory in explaining all features of the geochemistry, mineralogy, and petrology of the clan. Further, it is considered that many of the recent petrogenetic hypotheses, summarized in Sections 4.3 and 4.4 of this work, have serious flaws. Consequently, it is concluded that, although some advances have been made, many years of study have not resolved outstanding problems, such as the nature of the source, depth of melting, genesis of the megacryst suite, and origins of the

PETROGENESIS OF O~ANGEITES AND KIMBERLITES 367

characteristic hybridization. In common with orangeites, a particular impediment to understanding kimberlite genesis and evolution is our complete ignorance of the composition of primitive and derivative kimberlite magmas. Although aspects of the above problems are considered below, any attempts to devise a comprehensive model of kimberlite genesis are premature.

The low-pressure evolution of kimberlite magmas (i.e., post-emplacement crystallization and diatreme formation) is not considered in detail in this work as few major advances have been made in this area during the past decade (see 4.6.4). Of these, the recognition of phlogopite-kinoshitalite solid solutions as a common groundmass mineral (2.1.9.2) is perhaps the most significant. For reviews of aspects of this topic the reader should consult Mitchell (1986) and Scott Smith (1992). Recent specific descriptions of the mineralogy and petrology of a wide variety of kimberlites may be found in the proceedings of the Fourth and Fifth International Kimberlite Conferences published as Ross et al. (1989) and Meyer and Leonardos (1994), respectively.

4.6.1. Nature of the Source and Depth of Melting

For many years a major unresolved problem of kimberlite petrogenesis has been the question of whether the parental magma is derived from lithospheric or asthenospheric sources. Recent experimental studies (4.3) and geodynamic models (4.4) of kimberlite genesis indicate that kimberlites might be produced by the partial melting of carbonated garnet lherzolite or carbonated garnetite sources in the lower parts of the asthenospheric mantle or the transition zone. Although there is no agreement as to the nature or depth of melting of these sources, it is agreed that kimberlites are unlikely to be lithospheric magmas. This latter view, held by Skinner (1989), Foley (1988), and Wyllie (1980, 1989a,b) and extensively promoted by Bailey (1980, 1983, 1992, 1993), is based primarily upon the geographic restriction ofkimberlites to continental lithospheric plates.

Particularly strong evidence against a lithospheric source is the observation that isotopic signatures of archetypal kimberlites (Smith 1983) are similar to those of a wide variety of oceanic magmas (Zindler and Hart 1986), undoubtedly derived from asthenospheric sources. Kimberlites exhibiting an asthenospheric isotopic signature can only be derived from lithospheric sources by petrologically implausible processes. For example, these would require crystallization of an asthenospheric melt in the lithosphere, followed by the partial melting of that material within a very short time to ensure that enriched isotopic signatures are not developed. This may indeed be possible, but the initial part of this process is identical to that suggested for the origin of veined lithospheric mantle, and these sources cannot simultaneously give rise to kimberlites, orangeites, and other ultrapotassic magmas. Of course, virtually untestable, complex models involving zoning of metasomes (Haggerty 1989a,b, Thibault et at. 1992) may be considered as a means of overcoming this paradox.

A simpler answer is to place the source of all kimberlites in the asthenosphere and consider lithospheric metasomatism as a consequence of the interaction of kimberlites (and other magmas) with the lithosphere rather than being a prerequisite for their genesis. Metasomatism of the lithosphere by kimberlites is supported by the contemporaneity of

368 CHAPTER 4

the MARID and phlogopite-richterite peridotite xenolith suites and Upper Cretaceous South African kimberlites (Kinny and Dawson 1992, Erlank et al. 1987).

An asthenospheric source is further supported by the world-wide similarity in the mineralogy and geochemistry of kimberlites, which indicates that their parental magmas are produced by the same process occurring repeatedly in space and time. This observation applies to other asthenospheric magmas (MORB, tholeiitic basalt), but is in marked contrast to the unique character of the sources of orangeites and some other ultrapotassic magmas.

Accordingly, it is suggested that kimberlite magmas are generated by partial melting of asthenospheric material; however, determination of the source mineralogy and resolution of the questions as to where and how the magmas are produced is more problematic. One constraint is that the source must contain REE, Sr, K, Rb, Nb, and Ba, but never be isolated long enough to develop unusual isotopic signatures. This implies that the bulk composition of the sources has long-term integrated low Rb/Sr and high SmlNd ratios. Alkali, alkaline earth, and second-period transition elements are thus concentrated in kimberlitic partial melts derived from protoliths whose bulk compositions are relatively poor in these elements.

Kimberlites are not especially enriched in potassium «2 wt% K20) compared to lamproites, orangeites, or minettes (2-12 wt% K20; Mitchell 1986, Mitchell and Bergman 1991), but the presence in them of abundant carbonate, perovskite, and Ti-rich spinels, together with phlogopite macrocrysts and groundmass phlogopite-kinoshitalite requires that their sources contain C02, Ti, K, and Ba. Titanian phlogopite and magnesite are usually considered to host these elements, and the source is commonly believed to be a magnesite phlogopite garnet lherzolite (Mitchell 1986, Brey et al. 1983).

Whether this conclusion is correct or not remains problematical. The question of the presence of carbonate in the mantle has been discussed in Section 4.3.1.4, and it is concluded that magnesite may be present in carbonated peridotites. although the ultimate origin of the C02 is ambiguous. Phlogopite is commonly believed to be of "metasomatic" (sensu lata) origin. Whether it forms by such processes in the asthenosphere is unknown, but there is no a priori reason why metasomatism should be purely a lithospheric process. Consequently. portions of the asthenosphere may be enriched in alkali and alkali earth elements by melt migration and dehydration reactions.

At deep levels in the mantle, phlogopite is unlikely to be stable and K and Ti may be hosted by richterite (Luth et al. 1993). An alternative is to sequester K in clinopyroxene, as Edgar and Vukadinovic (1993) and Harlow and Veblen (1991) have shown that this mineral may contain significant amounts of K20 (> 1 wt %) at pressures above 6 OPa. This may provide a suitable source of potassium for relatively K-poor, C02-rich magmas such as kimberlites if magnesite and K-bearing clinopyroxene are consumed together during the initial stages of melting. At these depths Ba may be present in solid solution with carbonates, or as complex K-Ba titanates (Foley et al. 1994), rather than silicates.

The above comments indicate that the mineralogy of the mantle which might act as a source for kimberlites is likely to vary with depth. In the upper asthenosphere, the source might be magnesite phlogopite garnet lherzolite, whereas in the lower asthenosphere the same bulk composition might exist as a Ba-Sr-Ca-Mg-carbonate-bearing, K-Ti-richterite peridotite, or a carbonate-, K-Ba-titanate-bearing peridotite containing potassic diopside.


In the transition zone, potassium will probably be held in majoritic garnets and the source assemblage would be K-bearing majorite, y-Mg2Si04, and ilmenite or hollandite-structured titanates. In an oxidized mantle, magnesite would continue to host C02 at all depths as it is stable to 40 GPa and 2000°C (Katsura and Ito 1990). Limited partial melting of any of these sources could generate melts having the compositional characteristics postulated for primitive kimberlite magmas, i.e., C02-, K-, Ti-rich alkaline ultrabasic melts.

Assuming, for the purposes of this discussion, that melting of the source takes place at or above the 650-km discontinuity, i.e., the "uppermantle" of Anderson et al. (1992), then the mantle source may be juvenile primitive material upwelling from the lower mantle, metasomatized enriched "uppermantle," or a mixture of both. Formation of the enriched "uppermantle" could involve recycling of subducted material (Ringwood 1989, Ringwood et al. 1992, Anderson et al. 1992) or the circulation and crystallization of melts derived from the lower mantle and/or the "uppermantle." Unfortunately, as yet it is impossible to choose between the multitude of possible scenarios for the choice of a source, as none are constrained by experimental data.

Asthenospheric mantle is commonly considered to undergo decompressional melting as it rises toward the lithosphere as large plumes (Campbell and Griffiths 1990, Sleep 1990, Richards et al. 1989, White and McKenzie 1989) or diapirs (Green and Guegen 1974). Interestingly, recent tomographic seismic studies of the mantle appear to contradict the premises of these models. Thus, as an alternative, Anderson et at. (1992) have proposed the existence of deep (200-400 km) irregular regions of hot, perhaps partially molten mantle, termed "hot cells" or "hot regions." Discussion of the relative merits of these contradictory models is beyond the scope of this work. However, an interesting possibility arising from Anderson et at. 's (1992) hypothesis is that persistent partially molten hot cells could be the site in which kimberlites and the megacryst suite are "brewed."

Melting in the asthenosphere is probably confined to the porous flow regime, although channeled flow in the vicinity of the lithosphere-asthenosphere boundary cannot be discounted. The models of McKenzie (1989) and Hunter and McKenzie (1989) indicate that it should be possible to extract very small amounts of carbonate-rich melt from sources in the porous flow regime, a conclusion clearly applicable to asthenospheric kimberlite genesis.

In summary, partial models, such as advocated by Canil and Scarfe (1990), Edgar and Charbonneau (1993), Ringwood et at. (1992), appear to be viable for generating kimberlites in the asthenospheric mantle or transition zone. The major differences between the models lie primarily in the depth at which melting is believed to occur. The homogeneous source model of Tainton and McKenzie (1994) may also be useful if considered in an asthenospheric setting. Unfortunately, it is not yet possible to decide which of these models is preferable, and experimental studies of the problem by the inverse and forward approaches are highly desirable. Determination of the depth of melting is critical, and any new genetic hypothesis must be based on an increased understanding of the thermal and compositional character of the "uppermantle."

370 CHAPTER 4

4.6.2. The Megacryst Problem

Mitchell (1986, 1987) and Schulze (1987) have reviewed the characteristics of the Cr-poor megacrystJmacrocryst suite (1.3) and summarized hypotheses regarding their presumed cognate or xenocrystal relationship to kimberlites. Regardless of origin, the typical presence of this suite of minerals in kimberlites cannot be accidental and requires explanation. One significant conclusion emerging from studies by Mitchell (1986, 1987), Hunter and Taylor (1984), Davies etal. (1991), and Hops etal. (1992) is that themegacryst suite is itself a mixed assemblage, as significant intra- and inter-intrusion compositional and textural differences exist between megacrystal diopside, garnet, and ilmenite. Megacrysts within a given kimberlite appear to represent cumulates derived from the crystallization of several batches of their parental magma at high pressures. Disaggregation of these cumulates, and the random mixing, of the megacryst minerals within the magma which ultimately transports them, produces the hybrid heterogeneous assemblage of megacrysts found in kimberlites. Thus, megacrysts have no simple relationship to their current hosts. This observation is very important if megacrysts have a genetic relationship to kimberlites, as cumulates may crystallize from one batch of kimberlite magma and be transported by a second. Strictly speaking, the megacrysts should be considered as xenocrysts in the latter; unfortunately this designation obscures their genetic relationship to kimberlites. This observation suggests that much of the discussion concerning the cognate or xenocrystal origins of megacrysts is semantic and depends on whether they are regarded as being remotely (Hops et al. 1992, Jones 1987) or closely related to kimberlite magmas (Mitchell 1986).

Currently, megacrysts are believed to crystallize from a fractionating magma at depths of 150-200 km (Nixon and Boyd 1973, Eggler et al. 1979, Gurney et al. 1979b, Harte and Gurney 1981, Mitchell 1986, Schulze 1987, Hops et al. 1992), i.e., at or just below the lithosphere-asthenosphere boundary. The crux of the megacryst problem lies in the identification of this magma. There is a dichotomy of opinion between those who consider it to be kimberlite or protokimberlite (Hunter and Taylor 1984, Mitchell 1987, 1986, Canil and Scarfe 1990) and those suggesting it is basaltic (Harte 1983, Jones 1987, Davies et al. 1991, Hops et al. 1992).

Recent studies have not resolved the megacryst problem (Davies et al. 1991, Hops et al. 1992) as interpretations of the isotopic and geochemical data remain subjective. For example, isotopic disequilibrium between megacrysts and host kimberlites does not necessarily imply they are unrelated, given that the host is a hybrid-contaminated rock and certainly does not represent the original magma from which the megacrysts might have crystallized.

A xenocrystal origin for megacrysts has been suggested on the basis of calculation of the trace element content of their parental magma from that of the megacrysts (Kramers et al. 1981, Jones 1987). However, such hypotheses rest upon the dubious assumption that crystal-liquid distribution coefficients derived for basaltic systems at low pressures are relevant at high pressures to kimberlitic magmas.

Determination of their age is perhaps the only way in which the megacryst problem might be resolved. Hops et al. (1992) and Jones (1987) have determined that the Sm-Nd ages for garnet and diopside megacrysts from Jagersfontein and Premier are similar to


those of the host kimberlite. However, these ages might be unreliable given the possibilities of interaction of the megacrysts with the light REE-rich host and the significant errors associated with two-point isochrons. The former problem may be minimized by determination of megacryst ages by Lu-Hf geochronology using coexisting garnet and ilmenite. The latter problem will persist, given the apparent paucity of three-phase megacrysts.

Origin of the megacryst-forming magma from asthenospheric sources is now considered to be supported by the similarity of the Sr, Nd, and Pb isotopic compositions of the megacrysts to those of oceanic island basalts (Kramers et al. 1981, Jones 1987, Smith et al. 1987, Davies et al. 1991, Hops et ai. 1992).

Accordingly, Hops et al. (1992) and Jones (1987) propose that both kimberlites and megacrysts are derived from the same plume-related magmas. In these models a rising plume undergoes decompressional melting to produce alkali basaltic melts which are trapped as pools of magma at the base of the lithosphere. Crystallization of these pools, which exist in various stages of differentiation, gives rise to the Cr-poor megacryst suite. Kimberlite magmas are believed by Hops et ai. (1992) and Jones (1987) to be too magnesian and enriched in incompatible elements to be formed by extended fractionation of these magmas. Consequently, kimberlitic melts are considered to form when the residual magma infiltrates and equilibrates with the wall rock peridotites and assimilates other melts (unspecified) generated in the thermal boundary layer and at the base of the lithosphere. These processes generate hybrid kimberlitic melts which, on eruption, disrupt the megacryst cumulates and entrain them in the ascending magmas. Thus, a genetic relationship, albeit distant, between kimberlite and megacrysts is implied, although the megacrysts are correctly considered to be xenocrysts in their transporting kimberlites.

These hypotheses are attractive in explaining the heterogeneity of the macrocryst suites and the common derivation of these and kimberlites in an asthenospheric setting. Crystallization of megacrysts as high-pressure minerals from an "alkali basaltic magma" is in accord with the well-known occurrence of similar megacrysts in basanites (Green and Sobolov 1975, Parfenoff 1982, Le Blanc et al. 1982).

The most serious problem of the Hops et al. (1992) and Jones (1987) models is the assumption that kimberlites are generated by unspecified wall rock equilibration and assimilation processes. This hypothesis denies the existence of any uncontaminated primitive kimberlite magma. It is extremely unlikely that the suggested equilibration-assimilation process will be everywhere identical, and thus it cannot explain the worldwide petrological similarity of kimberlites. Hops et al. (1992) require wall rock equilibration to explain the high MgO contents ofkimberlites. However, the MgO content of primitive kimberlite magma is unknown. The observed high MgO contents of kimberlite rocks results not from an intrinsic character of the parent melt but from contamination of this magma with mantle-derived xenocrystal olivine (see 4.6.3). The relationship of kimberlites and the megacryst suite to asthenospheric basaltic magmas is considered further in Section 4.6.4.

4.6.3. Contamination of Kimberlites in the Mantle

Although Section 2.3 concludes that kimberlites and macrocrystal orangeites contain similar suites of macrocrystal olivine, a significant difference is that kimberlites appear

372 CHAPTER 4

to contain a greater proportion of macrocrysts which can be identified, on textural evidence, as phenocrysts. This observation is not surprising as kimberlites, unlike orangeites, crystallize olivine as an abundant primary mineral at low pressures. In addition, experimental studies (4.3.1.2, 4.3.1.3) indicate that olivine is undoubtedly a primary liquidus phase during passage of the magma through the upper lithosphere and may even be a primary phase in protokimberlite magmas in the asthenosphere.

It is considered in this work (1.3, 2.3, 3.7) that kimberlites are hybrid magmas whose high Ni and MgO contents result primarily from the addition of copious quantities of xenocrystal olivine to an incompatible element-rich mantle-derived magma. The problem of the identification of a source for these xenocrysts is identical to that described above with respect to orangeites (4.5.3).

Textural and compositional studies (2.3) indicate that many of the xenocrysts are derived from disaggregated ultramafic rocks. The sources of the xenoliths are placed in the upper mantle, and disaggregation is believed to occur during transport in ascending kimberlite. The mechanics of the processes which result in disaggregation are unknown. Some kimberlites contain xenocrystal diamond and subca1cic chrome pyrope; therefore, in these, some of the macrocrysts must be derived from diamond-bearing garnet dunite. There is no difference between the xenocrystal olivine suites present in kimberlites that have passed through diamond-bearing deep cratonic roots, and those emplaced in ancient mobile zones, i.e., "on-craton" and "off-craton" kimberlites. The latter lack subcalcic chrome pyrope and diamond, and the kimberlites could not have encountered diamondbearing horizons during their ascent (see 4.8). However, passage through chrome pyropebearing dunites is still required to explain the presence of olivine and apparent absence of orthopyroxene macrocrysts. Appealing to derivation of the olivine from lherzolite followed by the complete assimilation of orthopyroxene is unsatisfactory, as noted in Section 4.5.3. Thus, it is considered that the common granular and sheared peridotite xenoliths found in kimberlites are not the major source of the xenocryst suite. Clearly, kimberlites sample such peridotites during their ascent, and the preferential occurrence of these xenoliths in kimberlites, relative to orangeites, must be due to a greater extent of magma-conduit wall interaction. This may be related to the higher volatile contents, differing viscosities and ascent rates ofkimberlites relative to orangeites (and lamproites). Unfortunately, virtually nothing is known of the values of these physicochemical parameters, and further speculation on this topic is not justified.

In summary, it is suggested that asthenospheric kimberlites are contaminated, primarily near the base of the lithosphere, by garnet dunites (or extremely-orthopyroxenepoor harzburgites), some of which contain diamond. For unknown reasons deep lithospheric dunites appear to be particularly prone to disaggregation, relative to more competent granular garnet harzburgites and lherzolites encountered in the upper lithosphere. Thus, xenocrystal olivine is believed to be added primarily to kimberlite (and orangeite) magmas in the initial stages of their passage through the lithosphere, at or above the lithosphere-asthenosphere boundary. Olivines may be derived from many different dunitic protoliths, but the limited compositional variation exhibited by mantle-derived olivine precludes identification of their relative proportions.


4.6.4. Post-emplacement Crystallization

During magma ascent olivine and minor phlogopite crystallize as primary liquidus phases, and the hybrid assemblage of xenocrysts and phenocrysts is emplaced in the upper crust as a series of dikes (Clement 1982, Mitchell 1986). During post-emplacement crystallization all early-formed olivines attempt to equilibrate with their current host magma by reacting to develop mantles of olivine, whose composition corresponds to that of the final groundmass olivines (Mitchell 1986). Phenocrystal micas may be concentrated by flow differentiation during emplacement and it is this process which results in the formation of the archetypal kimberlites of "lamprophyric" appearance which are superficially similar to some orangeites.

Post-emplacement crystallization in the hypabyssal environment results in the crystallization of rutile, spinel, and perovskite followed by monticellite, apatite, and phlogopite-kinoshitalite as primary ground mass minerals (this work, Mitchell 1986). The mesostasis, in which these crystals are set, develops by the crystallization of calcite followed by primary serpophitic serpentine (Mitchell and Putnis 1988). These minerals may form a uniform groundmass or discrete segregations, depending on the local cooling rate (Mitchell 1986). Reaction of late-stage mesostasis-forming fluids with the preexisting mineral assemblage may lead to prograde carbonatization and serpentinization (Jago and Mitchell 1985).

Diatremes and volcanic craters are developed above the hypabyssal infrastructure of precursor dikes. Reviews of the diverse hypotheses proposed to explain the origin of diatremes are given by Clement (1982), Clement and Reid (1989), and Mitchell (1986). These hypotheses are not discussed further in this work as no major advances in this field have been made since publication of these reviews.

4.6.5. Summary

Kimberlites and their associated megacrysts are derived from asthenospheric sources whose composition is probably close to that of a REE-, Ti-, K-, Ba-bearing carbonated ultrabasic rock. The mineral assemblage represented by this bulk composition will vary with depth. Until the depth at which melting occurs is known the actual mineralogy of the kimberlite source cannot be determined. Regardless of the modal mineralogy of the source, small degrees of partial melting will lead to the formation of incompatible element-rich, C02-bearing, undersaturated, ultrabasic liquids. These segregate from their sources in the porous flow regime and pass into the lithosphere. The megacryst suite may be incorporated in the ascending magma either in the asthenosphere or at the lithosphereasthenosphere boundary (see below). During passage through the lithospheric mantle, the megacryst-bearing melts are extensively contaminated by xenocrystal olivines derived from the disaggregation of garnet dunites. Olivine also crystallizes as a primary phase at low pressures in the upper lithosphere. The final hybrid assemblage of phenocrysts, megacrysts, and xenocrysts is emplaced as a crystal-rich magma in a series of dikes in the upper crust.

The above summary of the processes involved in magma formation and evolution cannot be regarded as the definitive model of kimberlite genesis, as there remain many outstanding problems with regard to the nature of the source, depth of melting, and origin

374 CHAPTER 4

of the megacryst suite. Some speculations on the latter and on the relationship of kimberlites to other asthenospheric magmas are considered below.

Currently, the majority of megacrysts are considered to form at the lithosphereasthenosphere boundary. The discovery of megacrysts derived from greater depths (Sautter et at. 1991) and experimental data of Canil and Scarfe (1990) suggest that this conclusion is not entirely correct. Most current hypotheses also imply that megacrysts formed in a single event; this assumption is unlikely to be valid within any uniformitarianist models of magmatism. In the interest of generating further discussion of the megacryst problem, the following, highly speculative, scenario is presented as an alternative to current models of megacryst genesis.

Assuming that plumes or hot cells exist in the asthenosphere, it is suggested that the mantle within these may be at temperatures close to the solidus. Consequently, decompressional partial melting may be easily induced by small amounts of upwelling. This may be initiated by processes occurring in the transition zone or even triggered by the upwelling of lower mantle material. Partial melting will produce a variety of magmas, depending upon the depth and degree of melting. These melts may segregate from their sources, ascend toward the lithosphere, and commence crystallization.

It is further assumed that melts are continually produced over geological time, and all may crystallize minerals of the megacryst suite. However, because near-solidus initial melting would preferentially consume carbonate, Ti-bearing phases, clinopyroxene, and garnet, the megacryst suite may be preferentially produced from initial limited partial melts of kimberlitic composition. This is because bulk compositions of the melts would not lie in the primary phase volume of olivine for this system. This constraint does not preclude some precipitation of the megacryst suite from more advanced partial melts produced from the same source. However, crystallization of megacrysts from these may be limited due to the appearance of olivine as the dominant primary liquidus phase as the bulk composition of the partial melt changes. Thus, in such systems, the initial precipitation of megacrysts and/or increasing amounts of melting may drive liquid compositions very quickly into the primary phase volume of olivine. It might be expected that the potential for megacryst formation would decrease as partial melts changed composition in the sequence kimberlite, melilitite, basanite, as the degree of partial melting increases. Note, all of these magmas are known from geological evidence or experimental studies to have the potential to crystallize megacrysts (Green and Sobolov 1975, Nixon and Boyd 1979, Nixon et at. 1980, LeBlanc et at. 1982, Parfenoff 1982, Mitchell 1986).

Melts may crystallize completely in the asthenosphere and/or become trapped at the lithosphere-asthenosphere boundary, as it is very unlikely that all partial melts produced in the mantle will be erupted. There may be a preferential concentration of crystallizing melts at the underside of the lithosphere, where magmas are channeled into topographic irregularities at the base of the cratonic root. Some melts might enter the lithosphere and contribute to the veined metasomatic domains which have been proposed to exist at the base of the lithosphere. The end result of these processes may be pockets of megacrysts, scattered over a range of depths in the upper asthenosphere, which are derived from a variety of magmas of different ages. Fragmentation of these cumulates by subsequent batches of magma followed by entrainment and mixing will give rise to the observed hybrid assemblages of megacrysts. Note that magmas causing fragmentation of cumulates


may themselves be simultaneously crystallizing megacrysts. This model incorporates elements of hypotheses advanced by Harte and Gurney (1981), Mitchell (1986), Davies et al. (1991), and Hops et al. (1992), but differs in that the megacryst assemblage in kimberlites is considered to originate from several sources of different ages.

The above scenario suggests that kimberlite magmatism is but one facet of the broader magmatic events related to the upwelling and melting of asthenospheric mantle. In this context it is suggested that, depending upon the depth and extent of partial melting, the same source may give rise to a spectrum of partial melts. One possibility is that a mantle hot cell or rising plume may initially give rise to kimberlites which are followed by melilitites and other alkaline magmas as the extent of partial melting increases. The latter magmas may be generated only in the regimes of extensive lithospheric thinning and rifting resulting from the continued upwelling of the plume or hot cell. An alternative is that mantle hot cells may be thermally and compositionally asymmetric (laterally and vertically). Melting in one part of the cell may give rise to kimberlites, and in another to melilitites. If plumes or hot cells are not particularly buoyant, or if ascent is terminated, lithospheric thinning may not occur. Partial melts migrating ahead of such plumes or hot cells may intrude the continental lithosphere without associated rifting and nephelinitecarbonatite volcanism. Note that different partial melts will not necessarily ascend at the same rate or segregate and be emplaced at the same time. Such aborted plumes, or hot cells, might be the cause of the geographic association of kimberlite and melilititic magmatism found in southern Africa (Namibian kimberlites-Bushmanland melilitites), the Anabar Shield (Ukukitskoye-Olenyok kimberlites-northeastern Anabar melilitites), and the Hudson Bay Lowlands of Canada (Attawapiskat kimberlites-Lowlands melilitites). None of these igneous provinces are associated with rifts. This hypothesis suggests that melilitites are not lithospheric melts, and the absence of diamond in the majority of them is not simply due to genesis at depths above the stability field of diamond (see 4.8).

Kimberlites are not found in oceanic settings as upwelling of asthenospheric mantle is too rapid and partial melting too extensive to enable the initial limited partial melts to escape before being overwhelmed by basaltic magmas. Consequently, the components of kimberlite are simply diluted by and hybridized with the more voluminous basaltic magmas.

4.7. RELATIONSIDPS OFORANGEITES TO KIMBERLITES, LAMPROITES, AND OTHER ULTRAPOTASSIC MAGMAS

4.7.1. Kimberlites

The central thesis of this monograph is that orangeites and archetypal kimberlites represent distinct magma types derived from lithospheric and asthenospheric sources, respectively. Consequently, they have no genetic relationship. The only possible relationships are indirect and within the broader context of the overall evolution of the lithospheric and asthenospheric mantle. Hence, asthenospheric processes, which eventually result in kimberlite or basaltoid magma generation, may passively affect the lithospheric sources of orangeites and induce partial melting. This may result from uplift above upwelling plumes or hot cells and/or release of juvenile volatiles. Of course,

376 CHAPTER 4

asthenospheric magmas might ultimately be the parents of the ancient veined lithospheric sources of orangeites in the roots of cratons.

4.7.2. Lamproites

Orangeites have closer mineralogical, geochemical, and isotopic similarities to lamproites than archetypal kimberlites, and, in terms of their C02 and H20 contents, might be thought of as being intermediate between C02-rich kimberlites and H20-rich lamproites. In contrast to lamproites, which only rarely contain carbonates, orangeites characteristically contain primary carbonates. In addition, orangeites are poor in fluorine, relative to lamproites.

Although orangeites and macrocrystal orangeites contain some accessory minerals suggestive of a lamproitic affinity, only evolved diopside and sanidine richterite orangeites consist of a mineral assemblage similar to that of lamproites. There are significant differences between the clans with respect to the compositional variation and relative abundances of typomorphic minerals, such as phlogopite, potassium richterite, potassium feldspar, and hollandite. These differences are discussed in detail in Chapter 2.

Mitchell and Bergman (1991) have concluded that the emplacement of lamproites in a wide variety of geological environments precludes development of a universal model of their temporal and tectonic settings. Lamproites commonly occur along the margins of cratons or in accreted ancient mobile belts in regions of thick crust and lithosphere. They are not associated with modem active subduction zones, although potassic magmas formed in this environment share some of the geochemical characteristics of lamproites, i.e., Sr and Nd isotopic signatures suggestive of derivation from enriched sources, and negative Ti, Nb, and Ta anomalies. Mitchell and Bergman (1991) suggest that lamproite tectonic settings and compositional traits indicate they might represent partial melts of ancient subduction zones.

Orangeites are restricted to the Kaapvaal craton and considered to be derived from a lithospheric source located near the deepest parts of the cratonic root. It cannot be unambiguously determined whether subducted oceanic lithosphere played any role in the development of this source. In this context, extended incompatible-element distribution patterns for orangeites do not exhibit the characteristic negative Ti-Nb--Ta anomalies found for lamproitic and other subduction-related magmas.

Currently, lamproite sources are believed to be veined or metasomatized lithosphere (Fraser et al. 1985, Mitchell and Bergman 1991, Foley I 992a,b ). Lamproite primary magmas are considered by Mitchell and Bergman (1991) to be peralkaline ultrapotassic melts of intermediate silica content, which is the reason why they crystallize abundant potassium feldspar, leucite, and richterite and are mica-poor, relative to orangeites. The magma compositions must reflect the presence of greater amounts of relatively silica-rich minerals in their sources, i.e., richterite and diopside. Thus, geochemical and isotopic studies suggest that the sources of both magma types are broadly similar in mineralogical character but not mode. Those of orangeite must be richer in phlogopite and apatite and contain substantial quantities of C02 and be poor in F. Only the parental magmas of evolved orangeites might be derived from veins whose modal mineralogy overlaps that of lamproites. Unfortunately, the exact mineralogy of the source cannot be specified. In


summary, the compositions of lamproites suggest that Ti-potassium richterite, diopside, and K-Ba titanates are dominant in their sources in contrast to the phlogopite-apatite sources suggested in this work for orangeites. Partial melting of lamproite sources probably takes place by the vein-plus-wall rock mechanism of Foley (l992b).

In conclusion, because of the mineralogical and geochemical differences noted above, orangeites are not regarded as members of the lamproite clan. However, their similarities with lamproites and some other potassic rocks cannot be disregarded, and it is suggested that all of these magmas represent different expressions of a more general process of continental potassic magmatism (see 4.7.3).

4.7.3. Other Ultrapotassic Magmas

Lamproites are absent from the Kaapvaal craton, and the nearest occurrence of lamproitic rocks is at Kapamba in the Luangwa valley of Zambia (Scott Smith et al. 1989). There is also a notable absence of other types of markedly potassic magmatism in the Kaapvaal-Zimbabwe craton, in marked contrast to the situation in other cratons. It is suggested here that the absence of such rocks is directly related to the long-term metasomatic history of this craton.

In the Wyoming (U.S.A.) craton, lamproites (Leucite Hills, Smoky Butte; Carmichael 1967a, Mitchell et al. 1987, Mitchell and Bergman 1991) occur along with a wide variety of potassic rocks belonging to the shonkinite-fergusite-missourite-minettesuite (Highwood, Crazy, and Bearpaw Mountains; Larson 1940, Hearn 1989, O'Brien et al. 1991, MacDonald et al. 1992). In the Aldan (Russia) craton, many rocks with mineralogical affinities to lamproites are found in the Murun ultrapotassic complex (Shadenkov et al. 1989, Orlova 1988, Vladykin 1985), in addition to a host of geochemically defined lamproites (Bogatikov et al. 1985, 1991) and many plutonic complexes consisting of extremely undersaturated potassic rocks belonging to the kalsilite-Ieucitebiotite-<>rthoclase pyroxenite suite (Kostyuk et at. 1990). In other cratons potassic magmatism is represented primarily by lamproites-e.g., West Kimberley (Australia) craton (Jaques et al. 1986), Bhandara-Singhbum (India) craton (Mitchell and Bergman 1991), while the minette, shonkinite, and/or kalsilite pyroxenite suites are absent.

Currently, some form oflithospheric mantle metasomatism (sensu lato) is considered to play a role in the development of the sources of all of the above potassic rocks (McCulloch et al. 1983, Vollmer et al. 1984, Mitchell et al. 1987, 1994, Dudas et al. 1987, O'Brien et al. 1991, Mitchell and Bergman 1991), although opinions differ as to the relative contribution of asthenospheric or recycled subducted components (Nelson et al. 1986) to these sources during partial melting. The differing expressions of potassic magmatism in each craton suggest that, although they were all affected by a similar metasomatic (sensu lato) process, the duration, style, and extent of this process was different in each craton.

These differences are clearly expressed in the isotopic signatures of the potassic rocks. Thus, the ancient REE and Ba enrichment of the Wyoming and Aldan cratons appears not to have been accompanied by Rb enrichment, i.e., the lithosphere had long-term low Rb/Sr and Sm/Nd ratios (O'Brien et al. 1991, Mitchell and Bergman 1991, Mitchell et al. 1994). In contrast, similar REE enrichment of the sources of the Spanish

378 CHAPTER 4

and West Kimberley lamproites were associated with Rb enrichment and their sources reflect long-term high Rb/Sr and low SrnlNd ratios.

Mantle metasomatism (sensu lato) undoubtedly prepares cratons for the generation of potassic magmas. On a time scale of 1-2 Ga, different modal mineralogies will develop in veins and metasomatic horizons in different cratons. For example, one craton may be enriched in mica or potassium richterite relative to another, the differing styles perhaps reflecting the relative proportions of small-volume melt infiltration to fluid-related metasomatism. Some cratons may experience long-term persistent metasomatism (sensu lato), whereas others may undergo relatively little metasomatism. Persistent metasomatism may destroy the diamond-bearing roots of a craton if the melts and fluids involved are sufficiently oxidizing. Although metasomatism (sensu lato) as a physicochemical process will always operate in the same manner, its geological effects are likely to be unpredictable and random. Hence, the accumulated results of metasomatism (sensu lato) will never be the same everywhere.

It is concluded from the above observations that the metasomatic (sensu lato) history and evolution of a particular craton must be unique. Partial melting of these diverse, yet similar, lithospheric metasomatic assemblages will result in each craton exhibiting a distinctive style of potassic magmatism.

Magmas will be produced in different cratons, which in some instances may be broadly similar in bulk composition, but significantly different in trace element content and isotopic signatures, e.g., the lamproites of the Leucite Hills and West Kimberley provinces. Within a given craton individual lamproite fields may have similar isotopic signatures, but significantly different mineralogical character, e.g., Francis, Leucite Hills, Smoky Butte.

Different styles of metasomatism explain why extremely undersaturated potassic rocks are found in the Aldan shield, but not in the Wyoming craton. Both of these cratons appear to have experienced persistent extensive metasomatism, judging by the plethora of potassic alkaline rocks present in them compared with other cratons. Note thatthe 1-35 Ma lamproites of the Wyoming craton do not contain diamond; Devonian WyomingColorado kimberlites have very low diamond grades; and lamproite-like rocks of the Aldan Shield apparently lack diamond. These observations suggest that oxidation of any diamond originally present in the craton roots occurs during extensive and extended metasomatism.

As one might expect, there exist suites of potassic rocks which have some mineralogical or geochemical affinities to potassic rocks known from other cratons, but are significantly different in other respects. The Mata de Corda (Brazil) and Baker Lake (Canada) suites of ultrapotassic volcanism fall into this category. The Mata de Corda Formation has lamproitic and kamafugitic mineralogical and geochemical characteristics and cannot be unambiguously assigned to either ofthese clans (Sgarbi and Valenca 1991, Ulbrich and Leonardos 1991). Volcanic rocks of the Baker Lake Formation (Peterson 1994) are similarly ambiguous in their petrological character in having mineral assemblages typical of minettes, coupled with geochemical affinities to lamproites.

The existence of these apparently "transitional" rocks does not imply that kamafugite or minette magmas differentiate to lamproites, or vice versa. The occurrences are interpreted as indicating that the sources of these rocks may be unique, relative to the


sources of potassic rocks in other cratons. This conclusion is supported by recent Sr-Nd isotopic studies of kimberlites and other alkaline rocks from Brazil (Bizzi et at. 1994), which demonstrate that the Sao Francisco craton and associated accreted mobile belts have undergone a significantly different evolutionary history to the Kaapvaal craton.

These observations raise an interesting question: does each suite of apparently unique potassic rocks require a separate clan name? Clearly, if this approach is taken, it could lead to a proliferation of new magma types. Many petrologists are resistant to the introduction of new names (including the writer) and are, in the case of alkaline rocks, actively attempting to rationalize the archaic nomenclature by eliminating otiose uninformative names (Scott Smith et at. 1984, Scott Smith 1992, Mitchell and Bergman 1991, Mitchell 1994c, Woolley et at. 1995). However, the introduction of new names is entirely legitimate if a clan or suite of rocks is demonstrably genetically different to rocks belonging to other clans. The recognition ofkomatiites and boninites illustrates the value of this approach. Modal or contaminated/hybrid variants of existing clans do not warrant new names. In this context, the Mata de Corda Formation has not yet been sufficiently characterized with respect to its origin and evolution. It may be that it merely represents a kamafugitic magma that has been contaminated with lithospheric components rather than a distinctive magma type which warrants a new name.

Do orangeites require a new name, or are they merely extreme variants of the lamproite clan? The conclusion of this work, based on mineralogical and geochemical evidence, presented here and in Chapters 2 and 3, demonstrates that they are not lamproites and a new name is justified. Evolved orangeites considered in isolation from the less-evolved antecedents might indeed be described as lamproites, e.g., Tainton and Browning (1991). Note, however, that evolved amphibole-bearing minettes (Hall 1982, Nemec 1988) also have similar mineralogical and geochemical characteristics to both evolved orangeites and lamproites and could be assigned to either clan. Alternatively, evolved orangeites could be termed "minettes"! These observations do not imply that all of these rocks are related or the names synonymous, but they do illustrate the problems that arise in attempting to classify individual rocks in isolation from consanguineous antecedents and descendents (see 1.4). These apparently similar rocks reflect a mineralogical convergence resulting from differentiation of genetically diverse saturated-tooversaturated peralkaline magmas. The crystallization of coexisting groundmass richterite-arfvedsonite and potassium feldspar as the major modal phases in all cases is in response to the similar bulk major element composition of the residua. Mineralogicalgenetic classifications may be used to identify the parental clans of such petrographically similar rocks (1.4.4).

In the Kaapvaal-Zimbabwe craton, the restriction of potassic magmatism to orangeites suggests that this craton has not been subjected to the persistent and extensive metasomatism that has afflicted the Wyoming, Aldan, and Sao Francisco cratons. This may be a consequence of unusually deep, extremely durable cratonic roots acting as a near-impermeable barrier to thermal and chemical assaults from asthenospheric magmas. Similar conclusions might be drawn for the Anabar craton. It cannot be coincidental that the world's major sources of diamonds originate from the roots of these cratons and not from those of strongly metasomatized cratons characterized by extensive potassic magmatism.

380 CHAPTER 4

4.7.4. Summary

In summary, orangeites, or rocks which might be considered as mineralogical variants of the clan, do not occur outside the Kaapvaal craton. They are genetically and mineralogically distinct from lamproites and thus represent a distinct magma type. They are believed to represent the sole expression of potassic magmatism in the KaapvaalZi mbabwe craton. As a consequence of their different metasomatic histories other cratons vary with respect to the style of lamproite clan or plutonic-to-volcanic potassic and ultrapotassic magamatism.

As a final point, it may be that eventually we will be able to consider all of the diverse forms of potassic lithospheric magmatism as belonging to a family of magmas whose origins must be sought in metasomatic modifications of the lithospheric mantle. Following a suggestion by Barbara Scott Smith (pers. comm. 1994) these might be collectively referred to as the Three M-family, i.e., metasomatized mantle magmas!

4.8. PRIMARY DIAMOND DEPOSITS

As a finale to a work concerned with diamond-bearing rocks some comments are required regarding aspects of primary diamond deposits. "Primary" is used in the sense that these diamond deposits have not been formed by surficial weathering process. Strictly, the deposits should be called "pseudoprimary," as, if all diamonds are xenocrysts, the actual primary source lies at inaccessible depths in the mantle.

Detailed discussion of the genesis of diamond is beyond the scope of this work. Reviews of this topic may be found in Meyer (1985, 1987), Harris (1987), Gurney (1989), and Kirkley et al. (1991). Current models of diamond formation differ primarily with respect to the sources of carbon. One group suggests that carbon is juvenile, and deposition of it as diamond occurs as methane or other hydrocarbons are oxidized during their ascent through the upper mantle (Taylor and Green 1989) or at the lithosphereasthenosphere boundary (Haggerty 1986). These hypotheses are favored for the generation of diamonds containing the peridotitic suite of inclusions. Other possibilities for the origin of this suite of diamonds include crystallization from kimberlitic liquids (Harte et al. 1980, Arima et al. 1993b) or from ultrabasic melts during the formation of cratonic roots (this work).

A second group of hypotheses suggests that carbon is introduced into the mantle by subduction processes (Schulze 1986, Kesson and Ringwood 1989). The carbon is not juvenile and may be of biogenic origin (Milledge et al. 1983, Nisbet et al. 1994). A subduction origin is favored for the genesis of isotopically light diamonds containing the eclogitic suite of inclusions.

Although it is now evident that several diamond-forming processes exist, many important questions remain unanswered: the origin of mega- and microdiamonds; whether or not diamonds form in the lower asthenosphere transition zone; and whether diamond-forming processes are active today. The growth of diamond crystals and the mechanism of trapping silicate inclusions within them are not understood, as it is uncertain whether diamonds grow in the solid state as porphyroblasts or from liquids.


Regardless of origin, most diamonds are now believed to be xenocrysts in their transporting magmas. Current hypotheses postulate that the roots of cratons contain diamond-bearing horizons consisting of garnet dunite, garnet harzburgite, and eclogite. Diamonds may also exist just below the lithosphere-asthenosphere boundary, where they are believed by Haggerty (1986) to be formed by methane oxidation. Other diamonds may occur in subducted oceanic material underplating the flanks of a craton. The vertical and lateral disposition of all of these diamond-bearing horizons is completely unknown.

Disruption and disaggregation of diamond-bearing zones by the passage of magmas ascending from greater depths results in the incorporation of diamonds as xenocrysts in the magma. The subsequent fate of entrained xenocrysts is dependent upon the oxygen fugacity and rate of ascent of the magma toward, and through, the crust. Slow transport in highly oxidized hot magmas, e.g., lamproites, may result in the complete resorption of all diamond originally present. From the foregoing it is apparent that only magmas derived from depths below the diamond-bearing zones will contain diamonds.

Given the multiplicity of sources and the myriad of possible ascent paths, it is not surprising that individual primary diamond deposits differ greatly in character. A given deposit may contain diamonds derived from several protoliths, as the diamond suite present depends upon which diamond-bearing horizons are intersected.

This study has suggested that orangeites originate from, or just above, the lithosphere-asthenosphere boundary. Thus, there is a high probability of these magmas interacting with diamond-bearing zones. The diamond suite present in orangeites indicates disruption of both eclogitic and harzburgitic diamond-bearing horizons. Orangeites and macro crystal orangeites typically contain diamonds, whereas evolved orangeites appear to be diamond free. If evolved orangeites are low-pressure differentiates of orangeites, this may be a consequence of diamond resorption or gravitational fractionation of diamond from the magma. Of course, the parental magmas of evolved orangeites simply may not have intersected diamond-bearing horizons, although this option seems improbable as all occurrences of these rocks appear to be diamond free.

It is well known that only those kimberlites which are erupted through the diamondbearing roots of cratons contain diamonds, whereas those emplaced through accreted mobile belts lack diamonds (Clifford 1966). Importantly, there are no mineralogical differences between diamond-bearing and diamond-free kimberlites (Mitchell 1986). Similarly, the Cr-poor megacryst suite is identical in both varieties (Mitchell 1987 , Hops et at. 1992).

Boyd and Gurney (1986) have suggested that all kimberlites are derived from the lithosphere-asthenosphere boundary, which they consider to intersect the stability field of diamond only where the base of the craton underlies Archean shields (Figure 4.18). In mobile belts the lithosphere-asthenosphere boundary is located at lesser depths and within the stability field of graphite. Consequently, diamond-bearing kimberlites are confined to cratonic settings and kimberlites erupted through mobile belts are diamond free. Boyd and Gurney (1986) do not address the question of why most "on-craton" kimberlites are diamond poor or diamond free.

Boyd and Gurney's (1986) hypothesis is based on a particular interpretation of the equilibration temperatures and pressures of xenolith suites found in "on-craton" and "off-craton" kimberlites. They believe that the inflections found in the xenolith-derived

382

NW

CD +

(0)

CHAPTER 4

•• • SE NW •• • SE

• •• km

ASTHENOSPHERE ASTHENOSPHERE

DIAMOND - BEARING ASTHENOSPHERE m DIAMOND-BEARING GARNET HARZBURGITE

BARREN KIMBERLITES • DIAMOND-BEARING KIMBERLITES

(b) +

Figure 4.18. Contrasting models illustrating why diamond-bearing kimberlites are restricted to within the bounds of the Kaapvaal craton and barren kimberlites are confined to adjacent mobile belts. (A) In the Haggerty (1986) and Mitchell (1986. 1987) model. kimberlites are derived from similar depths within the asthenosphere as a result of partial melting of upwelling material. Asthenospheric diamonds are formed by methane dissociation at the lithosphere-asthenosphere boundary (LAB) in the vicinity of the deepest parts of the craton root (Haggerty 1986). Lithospheric diamonds occur only within the harzburgitic root of the craton. Only kimberlites which pass through this region can entrain xenocrystal diamond. (B) In the Boyd and Gurney (1986) model, kimberlites are derived from different depths at the lithosphere- asthenosphere boundary, the location of which is defined by the equilibration parameters of garnet lherzolite xenoliths found in kimberlites. In the mobile belts the boundary is considered to lie within the graphite stability field. Diamonds are believed to be stable only within the deepest parts of the craton root. In this model all diamonds are of lithospheric origin.

paleogeothenns define the lithosphere-asthenosphere boundary. The maximum pressure calculated for anyone suite is believed to record the maximum depth of origin of the host kimberlite.

Contrasting interpretations of the equilibration data (Carswell and Gibb 1987) suggest that the Boyd and Gurney (1986) model is not plausible. Further, it is suggested in this work and by others (Canil and Scarfe 1990, Ringwood et al. 1992, Haggerty 1986, 1994) that kimberlites are of asthenospheric origin. If this hypothesis is correct, the Boyd and Gurney (1986) model has no relevance in explaining the absence of diamonds in "off-craton" kimberlites.

A simpler model, which makes no assumptions regarding the depth of origin of xenolith suites, requires that kimberlites are all derived from similar depths in the asthenosphere (Figure 4. 18). In this model only kimberlites which pass through the craton roots traverse diamond-bearing horizons (Haggerty 1986, Mitchell 1986, 1987, 1991b). A refined version of the model presented by Mitchell (1991 b) is presented as Fig. 4.19 to illustrate how different kimberlites might develop contrasting suites of diamonds.

It is considered that the asthenospheric source model best explains the fonnation of diamond-bearing and diamond-free kimberlites. Intra- and inter-kimberlite diamond grades differ because of the heterogeneous distribution of potential sources, sorting of

PETROGENESIS OFORANGEITES AND KIMBERLITES

RIFTING CRATON

N o

- 50

- 150

- 250

- 350 km

t t K K

383

MOBILE BELT

K

DIAMOND - BEARING FACIES

[""' .... '] GARNET :;f::l HARZBURGITE

lEI LITHOSPHERIC L.:...J ECLOGITE

~ SUBDUCTED ~ ECLOGITE

ASTHENOSPHERE IGLtl GA~[;.EJOS~~~~~OLlTE

Figure 4.19. Hypothetical cross section of an Archean craton, ancient accreted mobile belt, and younger incipient rift, showing the location of the lithosphere-asthenosphere boundary (LAB, relative to the stability fields of diamond and graphite. The diagram illustrates why different kimberlites (K) differ with respect to sources of xenocrystal diamond. KJ may contain lithospheric and asthenospheric garnet lherzolite diamonds together with garnet harzburgiteldunite-derived diamonds. K2 contains diamonds from the aforementioned sources plus diamonds derived from lithospheric and subducted eclogites, i.e., five distinct sources. K3 contains only lithospheric and asthenospheric garnet lherzolite-derived diamonds. K4 does not pass through any diamond-bearing zones and is barren. Orangeites (0) are shown originating at the LAB and contain diamonds derived from garnet harzburgite/dunites and subducted eclogites. Lamproite (L) contains diamonds derived only from subducted eclogite and lithospheric garnet lherzolite sources. Melilititic magmas are considered to be derived by greater degrees of melting from the same asthenospheric sources as kimberlites. Depending upon the depth of segregation of the magma they may (Mj) or may not (M2) contain diamonds. High degrees of partial melting at shallow depths in the rift zone produces nephelinite-suite magmas (N) which are always diamond free (modified from MitchellI99Ib).

diamonds during entrainment, flow and mixing of different batches of kimberlite, and varying degrees of resorption of diamond in the ascending magma (Mitchell 1991 b).

Lamproite diamond deposits (Argyle, Ellendale, Prairie Creek, Majhgawan) do not contain subcaIcic chrome pyrope xenocrysts. Thus, derivation of their diamond suites from highly depleted garnet dunitic-harzburgitic sources is unlikely. Studies of mantlederived lherzolites, xenocrysts, and inclusions in diamonds from the Australian lamproites (Jaques 1989) suggest that their lithospheric mantle sources are less refractory than the roots of the Kaapvaal craton. The inclusion suite indicates a predominantly eclogitic source, which may represent subducted oceanic lithosphere which was eventually cratonized. This diamond-bearing material now constitutes the deeper parts of the mobile belt surrounding the Kimberley craton. A similar tectonic setting is evident for the Mahjgawan lamproite, whereas that of the Prairie Creek lamproite is less well-defined (Mitchell and Bergman 1991). Diamonds in the latter appear to be derived from both peridotitic and eclogitic sources (McCandless et ai. 1994).

The presence of diamonds derived from cratonized mobile belts in Australia, India, and the United States is in marked contrast to the situation in the Kaapvaal craton. Mobile

384 CHAPTER 4

belts accreted to the latter appear to be diamond free, as kimberlites erupted through them are barren of diamonds. This observation has two implications: eclogite suite diamonds in Kaapvaal kimberlites and orangeites might be derived entirely from the Archean eclogites incorporated into the roots of the craton during the fusion of the Kaapvaal and Zimbabwe cratons (Helmstaedt and Schulze 1989); it confirms that the sources of orangeites are not in subducted mobile belts and, therefore, are not tectonically or geochemically equivalent to those of lamproites.

It follows from the above discussion that any asthenosphere-derived magma which passes through diamond-bearing horizons in the craton root has the potential to carry diamonds. This conclusion may explain the rare occurrences of diamonds reported from alkali basaltoid and lamprophyric rocks (Kaminskii 1984, Janse 1994) and the existence of diamond-bearing melilitites in the Anabar craton (this work). The latter occur on the northeastern flank of the exposed craton and constitute the Mesozoic West Ukukit, Luchakanskoye, Dukenskoye, Kuranakhskoye, and Ari-Mastakhskoye "kimberlite" fields (BrakhfogeI1984, Kovalskii et al. 1969). Examination of many samples from these fields by the author indicates that they contain diatreme and hypabyssal facies monticellite melilitites and alnoites rather than kimberlites (this work). Most of the intrusions are not diamondiferous and those that are have very low grades (V, Kornilova pers. comm.). The province is petrologically very similar to the Hudson Bay Lowlands melilitite province (Janse et al. 1989, Reed and Sinclair 1991).

The northeast Anabar melilitites are unique among melilitite provinces in containing diamonds, a fact that must be related to the unusual stability and very high diamond content of the Anabar craton. Melilitites erupted through the mobile belts surrounding the Kaapvaal-Zimbabwe craton do not contain diamonds for the same reasons that kimberlites in that tectonic setting are diamond free (see above). Other melilitite provinces (Balcones, Bathurst Island, Rhine Graben) occur in regions of lithospheric thinning and rifting and do not contain diamonds. This may be simply due to formation by partial melting at shallow depths in regions distant from any diamond-bearing horizons (Figure 4.19). Melilitites have been insufficiently studied to determine whether the BushmanlandSutherland, Hudson Bay Lowlands, and Anabar provinces are different in character and/or derived from different depths and sources than the rift-related suites.

kimberlites, orangeites, and related rocks || petrogenesis of orangeites and kimberlites

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