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JOURNAL OF PETROLOGY VOLUME 40 NUMBER 5 PAGES 705–727 1999 Layered Mantle Lithosphere in the Lac de Gras Area, Slave Craton: Composition, Structure and Origin W. L. GRIFFIN1,2∗, B. J. DOYLE3, C. G. RYAN1,2, N. J. PEARSON1, SUZANNE Y. O’REILLY1, R. DAVIES1, K. KIVI4, E. VAN ACHTERBERGH1 AND L. M. NATAPOV1 1 ARC NATIONAL KEY CENTRE FOR GEOCHEMICAL EVOLUTION AND METALLOGENY OF CONTINENTS, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY, SYDNEY, NSW 2109, AUSTRALIA 2 CSIRO EXPLORATION AND MINING, PO BOX 126, NORTH RYDE, NSW 2113, AUSTRALIA 3 KENNECOTT CANADA EXPLORATION INC., 200 GRANVILLE STREET, VANCOUVER, B.C. V6C 1S4, CANADA 4 KENNECOTT CANADA EXPLORATION INC., 1300 WALSH STREET, THUNDER BAY, ONT. P7E 4X4, CANADA RECEIVED APRIL 17, 1998; REVISED TYPESCRIPT ACCEPTED OCTOBER 23, 1998 Heavy-mineral concentrates (garnets, chromites) and xenoliths from 21 Cretaceous–Tertiary kimberlite intrusions have been used to map the lithospheric mantle beneath the Lac de Gras area in the central part of the Slave Province. Analyses of Nickel Temperature ( TNi) and Zinc Temperature ( TZn) have been used to place garnet and chromite xenocrysts, respectively, in depth context. Paleogeotherms derived from both xenoliths and concentrates lie near a 35 mW/ m2 conductive model at T Ζ 900°C, and near a 38 mW/m2 model at higher T, implying a marked change in conductivity and/ or a thermal transient. Plots of garnet composition vs TNi also show a sharp discontinuity in mantle composition at 900°C. Garnets from <145 km depth are ultradepleted in Y, Zr, Ti and Ga, whereas those from greater depths (to [200 km) are similar to garnets from Archean mantle world-wide. Relative abundances of garnet types indicate that the shallow layer consists of ~60% (clinopyroxene-free) harzburgite and 40% lherzolite, whereas the deeper layer contains 15–20% harzburgite and 80–85% lherzolite. T estimates on eclogite xenoliths show that all were derived from the deeper layer. Xenolith data and garnet compositions indicate that the shallow layer is more magnesian (Fo92–94) than the deeper layer (Fo91–92), and both layers are more olivine rich than South African or Siberian Archean peridotite xenoliths. The composition and sharply defined structure of the Lac de Gras lithosphere are unique within our current knowledge of Archean mantle sections. The shallow layer of this lithosphere section is similar to peridotites ∗Corresponding author. Present address: ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents, Department of Earth and Planetary Sciences, Macquarie University, Sydney, NSW 2109, Australia. from some highly depleted ophiolites from convergent-margin settings, and may have formed in a similar situation during the accretion of the Hackett and Contwoyto terranes (magmatic arc and accretionary prism, respectively) to the ancient continental Anton terrane at 2·6–2·7 Ga. The deeper layer is interpreted as a plume head, which rose from the lower mantle and underplated the existing lithosphere at 2·6 Ga; evidence includes a high proportion of the superdeep inclusion assemblage (ferropericlase–perovskite) in the diamond population. This event could have provided heat for generation of the widespread 2·6 Ga post-tectonic granites. Proterozoic subduction from east and west may have modified the cratonic root, mainly by introduction of eclogites near its base. lithosphere; mantle; Slave Craton; kimberlites; Cr-pyrope garnet; trace elements; diamonds KEY WORDS: INTRODUCTION Geophysical data suggest that many Archean cratons are underlain by deep lithospheric keels, commonly extending to depths in excess of 200 km. Xenolith suites Oxford University Press 1999 JOURNAL OF PETROLOGY VOLUME 40 from kimberlites show that the lithospheric mantle beneath at least some of these cratons contains rock types not found beneath areas of younger crust, implying major differences between Archean and post-Archean lithospheric processes [see review by Griffin et al. (1998c) and extensive references therein]. An understanding of the nature and origin of these ‘continental roots’ is crucial to models for Earth’s early evolution, but our present knowledge of their geology is based on a very limited sample. Approximately 70% of the described xenoliths of cratonic mantle material are derived from the Kaapvaal craton of southern Africa, and most of the remainder come from a single kimberlite (Udachnaya) on the Siberian craton. These two areas show several important similarities [summarized by Boyd et al. (1997)] and together they dominate our current understanding of the Archean lithospheric mantle. However, the limitations of these samples are illustrated by comparison with more extensive data on the chemistry of mantle-derived xenocryst minerals such as garnets and chromites. For example, such data show that the composition, structure and thermal state of the Kaapvaal lithosphere changed markedly ~90 my ago (Griffin et al., 1995; Brown et al., 1998); whereas nearly all of the xenolith material described in the literature is derived from post-90 Ma kimberlites, and provides little record of the earlier state of the cratonic root. In Siberia, data from mineral concentrates show that the composition and structure of the lithospheric mantle change markedly over relatively short distances, apparently reflecting crustal terrane boundaries (Griffin et al., 1995); the picture of the craton root given by the Udachnaya xenolith population thus is not representative of the lithospheric mantle even 100 km away. So what does a ‘typical’ Archean mantle lithosphere really look like? The recent discovery of diamondiferous kimberlites in the Slave Province of northern Canada (Pell, 1997) provides an opportunity to examine material from a third Archean craton root, and to test the generality of models for such roots. In this paper we report on an integrated study of mantle-derived samples from the Lac de Gras area, near the middle of the Slave Province. The samples include xenoliths, diamonds and diamond-inclusion minerals, but the major focus is on major- and trace-element data on chrome–pyrope garnets and chromites (>1100 garnets, >600 chromites) from 21 intrusions in 19 kimberlite pipes. These data indicate that the mantle lithosphere beneath the central Slave craton is markedly different in structure and overall composition from the other well-studied examples of Archean lithospheric mantle. These differences broaden our perspective on Archean processes, and on the evolution of continental roots in general. In a companion paper (Griffin et al., 1998e) we examine the lateral extent of the unusual NUMBER 5 MAY 1999 mantle structure found beneath the Lac de Gras area, and reported here. GEOLOGICAL SETTING The Slave Structural Province (Padgham & Fyson, 1992) is a small Archean nucleus within the larger North American Craton (Fig. 1). It is bounded on the east by the Thelon Orogen (~2·2 Ga) and on the west by the Wopmay Orogen (1·9–2·1 Ga), a series of magmatic arcs and accreted terranes. Its northern boundary is defined by the overlapping Proterozoic and younger supracrustals of the Bear Province and the Arctic Platform. Proterozoic sediments also overlap the craton along its northeastern side along the Bathurst Fault, and extend part-way across the craton in a narrow NE-trending belt (the Kilohigok Basin) ~150 km north of the Lac de Gras area. The southern boundary of the Province is the Great Slave Lake Shear Zone, an ancient (1·8–2·0 Ga) continental transform that has bisected and offset the craton and juxtaposed the Archean rocks of the Slave Province against Proterozoic rocks of the Churchill Province. The oldest rocks in the Slave Province are small remnants of felsic granites and gneisses, including the 3·6–4·0 Ga Acasta gneisses, in the western part of the craton. Most of the outcrop in the central and eastern parts is made up of several supracrustal series, recognized as the Yellowknife Supergroup (~2·7 Ga), which is intruded by an extensive series of pre- to post-deformational (2·69–2·60 Ga) felsic plutons. In contrast to most cratonic areas, the supracrustal rocks of the Yellowknife Supergroup are dominated by sedimentary rocks. Except in the westernmost part of the area, no basement has been identified for this succession. The earliest sediments, found only in the western part, are quartz arenites, suggesting derivation from, and sedimentation on, a stable shelf (Padgham & Fyson, 1992) in association with banded iron formations and overlain by felsic volcanics. The most widespread unit, occurring in all parts of the craton, is a series of greywacke–mudstone turbidites with abundant volcanoclastic debris, derived from and interbedded with felsic, basaltic and andesitic volcanic rocks. More than 45% of the volcanics are felsic to intermediate, whereas most of the basalts have high SiO2 contents (>50%). Komatiites are present in the southern part of the craton, but rare (Bleeker et al., 1999; B. Kjarsgaard, personal communication, 1998) and no alkalic lavas are known (Padgham & Fyson, 1992). A younger group of quartz-rich fluvial sediments post-dates most of the granitic activity. The intrusion of voluminous plutonic rocks marked the final stabilization of the craton. Davis et al. (1994) divided them into Group 1 (2689–2650 Ma; trondhjemites and diorites), Group 2 (2610–2600 Ma; syn- to 706 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA Fig. 1. The Slave Province, showing terrane boundaries of Griffin et al. (1998e), and the location of the Lac de Gras area (shaded box; Fig. 2). Lined pattern, felsic volcanics; black, mafic volcanics; stippled, Proterozoic sediments. Stars show kimberlite occurrences: RL, Ranch Lake; T, Torrie. late-deformational monzodiorite–granodiorite and trondhjemite) and Group 3 (2599–2580 Ma; post-deformation biotite ± muscovite granites). Groups 1 and 2 have generally calc-alkaline chemistry, and trace-element patterns suggestive of modern subduction-related igneous suites, whereas Group 3 resembles many Phanerozoic post-orogenic K–U–Th-rich granite suites. Pb-isotope data on volcanogenic massive sulphide (VMS) deposits within the Yellowknife Supergroup define a major boundary running roughly along the 112°W meridian (Thorpe et al., 1992). West of this line, the Pb in these deposits and in galena from later Au deposits is significantly more radiogenic, indicating the derivation of at least part of the Pb from older felsic crust. East of the line, the Pb isotope compositions indicate a more juvenile source. The Lac de Gras kimberlites discussed here lie to the east of the Pb isotope boundary. A similar boundary, ~100 km E of the Pb line, has been noted in the Nd isotope compositions of the Group 3 granites (Davis & Hegner, 1992) in the central part of the province. The Pb line also marks the eastern limit of exposures of pre-Yellowknife gneisses, and of occurrences of quartz arenites below the turbidite sequences; it therefore appears to correspond to a significant boundary between crustal volumes of different age (Padgham & Fyson, 1992). Davis et al. (1994) suggested that the portions of the craton east of this boundary were accreted to the margins of an older continent, at ~2·7 Ga. A synthesis of the geology of the craton was presented by Griffin et al. (1998e) in terms of this model (Fig. 1), which recognizes a complex of accretionary wedge (Contwoyto terrane) and arc-related material (Hackett River terrane) abutting the ancient continental nucleus (Anton terrane) along a deformed continental-margin setting (Sleepy Dragon terrane). The Lac de Gras area discussed here lies entirely within the Contwoyto terrane. Several swarms of Early–Mid-Proterozoic (2·0–2·3 Ga; LeCheminant et al., 1995) basaltic dikes of different orientations are known in the Lac de Gras area. The 2·03 Ga Lac de Gras swarm strikes roughly N10°E, and diverges from north to south, suggesting a source beneath the Kilohigok Basin. The most important event is 707 JOURNAL OF PETROLOGY VOLUME 40 represented by the NNW-trending Mackenzie dike swarm (1·27 Ga; LeCheminant & Heaman, 1989). These dikes extend over 2300 km from a focus, interpreted as a plume head (Fahrig, 1987), located west of Victoria Island; products of this event include the Muskox intrusion and associated flood basalts, indicating a massive igneous event. In the northern part of the Province, dikes were injected essentially vertically, whereas further to the south, at distances of >500 km from the focus, injection appears to have been dominantly horizontal (Ernst & Baragar, 1992); the Lac de Gras area lies in the zone of horizontal injection. The kimberlites around Lac de Gras may belong to several generations. Rb–Sr dating of two of them has yielded Eocene dates (47–52 Ma; Davis & Kjarsgaard, 1997); this is consistent with the occurrence of Paleocene to Cretaceous fossils (flora and fauna) in the pipes (Nassichuk & McIntyre, 1995). The occurrence of pipes with both normal and reversed magnetic polarity indicates that other generations may be present, and this is confirmed by an unpublished U–Pb age of 73–75 Ma on perovskite (Davis & Kjarsgaard, 1997) and an unpublished Rb–Sr age of 82 Ma (Pell, 1997). Paleozoic U–Pb ages on zircons have been reported from pipes to the SW, across the Pb isotope boundary mentioned above [Cross pipes, 450 Ma (Ashton Mining, personal communication, 1998); Drybones pipe, 440 Ma (Davis & Kjarsgaard, 1997)]. To the north of the present area, Kopylova et al. (1998) reported a Jurassic (172 Ma) emplacement age for the Jericho pipe. This study includes material from 19 kimberlite pipes (Fig. 2); two intrusions were sampled in pipe DO-27 and in pipe A154. For convenience in presentation, we have grouped them into three sectors (Fig. 2). All of these pipes are diamondiferous, but those in the Central sector are generally higher grade than those to the east and west. These sampling localities give a picture of the upper mantle over an area of 60 km × 20 km, in the center of the Slave diamond province. They are supplemented below by data from the Ranch Lake and Torrie pipes to the northwest (Griffin et al., 1998e). ANALYTICAL METHODS Major-element data on garnets and chromites have been obtained using the CAMEBAX SX50 electron microprobe at the School of Earth Sciences, Macquarie University, using standard techniques. Similar data from this microprobe have been independently verified by crossanalysis in Australia (CSIRO) and Norway (Mineralogical–Geological Museum). The trace-element analyses used in this study have been obtained with both the HIAF proton microprobe (PMP) at CSIRO Exploration and Mining, North Ryde, NUMBER 5 MAY 1999 Fig. 2. Lac de Gras area, showing kimberlites used in this study, and line of section shown in Figs 8 and 9. This line continues across the Ranch Lake and Torrie pipes (Fig. 1). and a laser-ablation inductively coupled plasma mass spectrometer (ICPMS) microprobe (LAM) at Macquarie University. The PMP methods have been described in detail by Ryan et al. (1990a, 1990b). The proton microprobe is based on a tandem electrostatic accelerator, which provides a beam of 3 MeV protons, focused onto the sample by an electrostatic lens. The characteristic Xrays generated by the proton bombardment are collected by an Si(Li) energy-dispersive detector and displayed as spectra. Quantitative concentration data are extracted from these spectra as described by Ryan et al. (1990b). Normalization to EMP values for Fe is used to correct for differences in sample conductivity; otherwise the method is independent of standards. Typical analytical precision and accuracy are better than ±10% for most elements discussed here. In this work, the typical size of the beam spot on the sample was 30–50 lm, and beam currents were 10–15 nA. Samples were counted to a uniform accumulated live charge of 3 lC, corresponding to analysis times of 5–8 min. The LAM instrumentation and methods have been described in detail by Norman et al. (1996). The instrument uses an Nd–YAG IR laser, frequency-quadrupled to produce a beam of UV (266 nm) light that is focused through a petrographic microscope onto a polished grain or thin section in the sample cell. Ar gas flowing through the cell carries the ablated sample to the inductively coupled plasma mass spectrometer. The external standard for the garnet analyses reported here was the NIST 610 glass, and the internal standard was Ca. For the chromite analyses, NIST 610 was used as the external standard for most elements, but Zn was standardized against a well-analyzed chromite megacryst (LCR-1); Mg or Al was used as the internal standard. Typical ablation pits were 40 lm in diameter and up to 50 lm deep. Detection limits for most elements are from 1 ppm to 100 ppb, and precision at these levels is typically 708 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA Fig. 3. CaO–Cr2O3 plot for garnets from mantle xenoliths, showing fields used to assign a rock type to individual garnet xenocrysts (after Griffin et al., 1999). Sobolev, 1997 (see translation Sobolev et al., 1974). better than 10%, as reported by Norman et al. (1996, 1998). The PMP was used for most trace-element analyses performed during 1994 and 1995; the LAM has been used for all subsequent analyses. Cross-analysis of individual samples and standards shows excellent agreement between techniques for the elements that can be analyzed by both (Norman et al., 1996). The mineral assemblage from which a given garnet macrocryst was derived has been estimated using the relationship between CaO and Cr2O3 contents (Fig. 3). Various divisions between harzburgitic and lherzolitic garnets have been suggested; in this paper we adopt the one proposed by Gurney (1984), although Sobolev (1974) has pointed out that some garnets classified as harzburgitic by this method could be derived from lherzolites in which the cpx coexisting with garnet is unusually Na rich, and hence Ca poor. We further subdivide the field of harzburgitic garnets in Fig. 3 into calcic harzburgite and low-Ca harzburgite. This line encloses 50% of diamond-inclusion garnets from South Africa (Gurney, 1984; Griffin et al., 1992), and thus highlights an important class of extremely depleted garnets, genetically associated with some diamonds. Another rather arbitrary line separates Ca-rich garnets of the type usually found in wehrlite (olivine + clinopyroxene + garnet) assemblages. Low-Cr garnets are defined as those with <1·5% Cr2O3; many of these, at the low-Cr end of the lherzolite trend, may be derived from magnesian pyroxenites rather than peridotites. XENOLITHS Xenoliths used in this study come mainly from the A154 kimberlite pipe, and have been recovered from the coarse concentrate after crushing of the kimberlite. Most are fragments 1–3 cm in diameter; larger xenoliths are seen in drill core but are generally friable, and do not survive the processing. This sampling bias limits the conclusions that can be made on the proportions of rock types, and on the modal and chemical composition of individual samples. A detailed description of the petrography and mineral chemistry of the xenoliths has been given by Pearson et al. (1998), and only a summary is given here. Several lithological groups have been recognized. (1) Lherzolites (ol + opx + cpx + gnt ± chr) show a broad spectrum of microstructures (granoblastic, porphyroclastic, mylonitic) and grain size (<1 mm to >1 cm). Fo contents in olivine range from 91·5% in sheared lherzolites to 92·8% in fine-grained, cpx-poor samples. The modal abundance of Cr-diopside is low and in several samples it occurs only in intimate association with chromite. Wehrlites (ol + cpx + gnt ± chr) are relatively rare and probably represent a modal variant of the lherzolite suite, but olivine is typically more Fe rich (Fo 90·5–91·2). (2) Harzburgites (ol + opx + gnt ± chr) as defined here lack modal cpx, but most samples have garnets that lie within the lherzolite trend in Fig. 3, suggesting equilibrium with clinopyroxene that may not be seen because of the small sample sizes. Subcalcic garnets are abundant in the concentrate (see below), and Boyd & Canil (1997) have analyzed subcalcic garnets in harzburgite xenoliths from the Grizzly Pipe on the north side of Lac de Gras. Most harzburgites have fine-grained granoblastic microstructures. (3) Fine-grained granoblastic dunites (ol ± gnt ± chr) may be modal variants of the lherzolite or harzburgite suite. Large (up to 2 cm) single olivine grains commonly are strained and contain inclusions of the other phases, and thus resemble the megacrystalline dunites of Siberian kimberlites (Boyd et al., 1997). (4) Two types of websterites (opx + cpx + gnt ± chr) are recognized on the basis of mineral compositions. Websterites with Cr-pyrope garnet (Cr2O3 to 7%) and Cr-diopside are considered to be modal variants of the lherzolite suite. Bimineralic garnet clinopyroxenites with similar mineral compositions represent a further modal variant, similar to the griquaites of Nixon (1987). The other group of websterites, with more Fe-rich mineral compositions, lower Cr contents and higher XJd (up to 20 mol %) in the cpx, may be related to the eclogite suite. (5) Eclogites (cpx + gnt ± ky) also can be divided into two main types. One is characterized by <7% CaO in garnet and cpx with XJd = 15–35. The other has garnet with 9–13% CaO and Na2O > 0·1%, and cpx with XJd [ 50. Some examples of this type also contain kyanite, and as a group they are similar in mineral composition to other kyanite eclogite xenoliths from southern Africa 709 JOURNAL OF PETROLOGY VOLUME 40 NUMBER 5 MAY 1999 Temperature estimates for the eclogite xenoliths range from ~890 to 1250°C; the temperature distribution is bimodal, and the modes correspond to the compositional groups noted above. The eclogites with low-Ca garnet lie in the range 890–1050°C, whereas the eclogites with high-Ca garnet give T [ 1100°C. The step in the geotherm (Fig. 4) is unusual, and there are basically two ways to produce it (Pearson et al., 1998). The step may be a transient feature, produced by heating at the time of kimberlite magmatism. Alternatively, it may be a steady-state feature produced by a marked change in thermal conductivity over a short vertical distance. These possibilities will be evaluated below on the basis of data from the concentrate minerals. Fig. 4. P–T plots for xenoliths from pipe A154S (after Pearson et al., 1998), using two different geothermobarometer combinations (Finnerty & Boyd, 1987). Points for xenoliths from the Grizzly Pipe (north of Lac de Gras) recalculated from analytical data of F. R. Boyd (personal communication, 1997). (Pearson et al., 1995) and eastern Australia (Pearson et al., 1991). Both types show varying degrees of breakdown of omphacite to diopside + plag along grain boundaries, and compositionally distinct overgrowths on garnet; these features are interpreted as the result of decompression and/or heating in the presence of fluids. THERMOBAROMETRY OF MANTLEDERIVED XENOLITHS Detailed discussion of the P–T estimates for the xenolith suite has been given by Pearson et al. (1998). For several independent combinations of thermobarometers, the overall features of the P–T distribution are robust (Fig. 4). P–T estimates for the low-T xenoliths (<900°C) fall near Pollack & Chapman’s (1977) conductive model geotherm corresponding to a surface heat flow of 35 mW/m2, whereas xenoliths with higher equilibration temperatures lie between this geotherm and a 40 mW/m2 conductive model. The change in gradient appears as a distinct step, at a pressure close to 50 kbar. All samples with P < 45 kbar give T < 900°C, whereas all xenoliths with P > 45 kbar (with one exception) give T > 1000°C, regardless of the method used. Sheared high-T lherzolite xenoliths, although similar in microstructure and mineral chemistry to those from southern Africa, do not define a ‘kink’ away from the conductive models like that found in many African xenolith suites (Finnerty & Boyd, 1987). The high-T group also includes garnet websterites and undeformed peridotites. The deepest xenoliths give pressure estimates of ~65 kbar, corresponding to a minimum lithosphere thickness of 200 km. DIAMONDS AND DIAMOND INCLUSIONS Diamonds from pipe DO-27 (Eastern sector, Fig. 2) have been described by Davies et al. (1998). They show a wide range in morphology, from planar octahedra with minor resorption, to heavily resorbed dodecahedra with plastic deformation, as well as a large population of cubic to cubo-octahedral stones, some of which have hopper faces. A similar range of types was described from the Point Lake pipe (north of Lac de Gras) by Taylor et al. (1995). Diamonds can be divided into Types I and II on the basis of N contents; Type II have N contents below detection (typically <5 ppm), whereas Type I have measurable N contents. In the diamonds from DO-27, there is a wide range in the degree of nitrogen aggregation, which does not correlate with nitrogen contents. Many cubic stones have high N contents but low degrees of aggregation, whereas many stones with high degrees of nitrogen aggregation show extensive plastic deformation. These features are similar to those observed in Point Lake diamonds by Taylor et al. (1995). Like most diamond suites, those from the Slave Province are dominated by Type I stones. However, although Type II diamonds are ‘comparatively rare’ at Point Lake, they make up ~30% of the population studied by Davies et al. (1998). Most C-isotope compositions, including all Type II diamonds, lie in the common mantle range of d13C = –3·5 to –5·5, but more than a quarter of the stones studied have lighter carbon, extending to d13C = –21. The syngenetic mineral inclusions can be divided into three parageneses (Davies et al., 1998). The peridotitic paragenesis is represented by olivine (Fo 92·8–94·0), Crpyrope garnet (both lherzolitic and harzburgitic associations, with Cr2O3 up to 15%) and pentlandite (>25% Ni). The recognition of the paragenesis of sulfide inclusions has been discussed by Bulanova et al. (1996). The eclogite paragenesis is represented by Ca-rich garnet (9–16% CaO), omphacite and pyrrhotite (<1·9%). The 710 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA superdeep paragenesis is represented by ferropericlase, MgSiO3 perovskite and native Ni. The ferropericlase inclusions have mg-number = 0·80–0·87, 0·3–0·8% Cr2O3 and 1·2–1·5% NiO. They are thus similar to ferropericlase inclusions reported from the Orroroo, Koffiefontein and Sloan kimberlites (Scott-Smith et al., 1984; Otter & Gurney, 1989). Chinn et al. (1998) also have described ferropericlase inclusions in diamonds from pipes on the north side of Lac de Gras. This superdeep paragenesis indicates derivation of some diamonds from depths of >670 km (Scott-Smith et al., 1984; Kesson & Fitz Gerald, 1991; Harte et al., 1998); all but one of these stones are of Type II. Comparison of mineral-inclusion parageneses with data on morphology, C-isotope composition, N concentration and N aggregation state suggests that ~50% of the diamonds are eclogitic, and ~25% are of the superdeep paragenesis (Davies et al., 1998). GARNET AND CHROMITE XENOCRYSTS Geothermobarometry The key to the use of heavy-mineral concentrates in lithosphere mapping has been the development of geothermometers that can be applied to single grains of chrome–pyrope garnet and chromite. In this work we use the temperature calibrations of Ryan et al. (1996) for the Ni content of garnet, and the Zn content of chromite, to obtain equilibration temperatures for individual mineral grains. For a discussion of alternative calibrations, which would not affect the depth relations discussed below, the reader is referred to Canil (1994) and Griffin & Ryan (1996). Geotherm parameters also can be derived directly from the concentrates, using the techniques described by Griffin & Ryan (1995) and Ryan et al. (1996). This approach uses algorithms based on a combination of experimental and empirical datasets, to calculate the pressure (PCr) at which each garnet in a concentrate would have been in equilibrium with chrome spinel (Ryan et al., 1996). If no spinel was present, the calculated PCr will be underestimated. Hence, when the garnets from a concentrate are plotted in a PCr–TNi diagram (Fig. 5), the ‘Garnet Geotherm’ is defined by the envelope of maximum PCr at each TNi, assuming those garnets coexisted with chrome spinel. Garnets that did not equilibrate with chromite scatter to lower PCr (Pmin), and must be projected to the Garnet Geotherm or the xenolith geotherm to derive their depth of origin. Comparison of the TZn spectrum of chromites from the same concentrate provides additional constraints on the interpretation of the geotherm. The method gives good agreement with geotherms derived by P–T estimates on xenoliths (Griffin et al., 1996; Ryan et al., 1996). Fig. 5. PCr–TNi plot for garnet xenocrysts from pipes in the Central sector of Lac de Gras, showing concordance between Garnet Geotherm and xenolith-derived geotherm. Histogram shows temperature (TZn) distribution of chromite xenocrysts from the same pipes. A PCr–TNi plot and histogram of TZn for the Lac de Gras area (Fig. 5) shows that below 900°C, the data indicate a geotherm lying close to the 35 mW/m2 model conductive geotherm of Pollack & Chapman (1977). Lherzolitic garnets generally give lower PCr at each TNi than harzburgitic ones; this reflects the lower degree of major-element depletion, and hence the lower probability of the garnet + chromite assemblage, in the lherzolitic rocks. The relatively small spread of PCr at each TNi in this T range suggests that most of the analyzed garnets were Cr saturated or nearly so. Above 1000°C, the geotherm is less well defined, but the data to 1100°C are consistent with a geotherm lying below the 40 mW/m2 conductive model. The abundance of chromites with TZn in the interval 900–1100°C suggests that many of the garnets in this T interval probably coexisted with chromite, and that the ‘step’ in the geotherm is real, rather than an artefact of sampling. However, the large spread in PCr at each TNi above 900°C indicates that many of the garnets with TNi [ 900°C are derived from chromitefree rocks, and the absence of garnets lying along the xenolith geotherm at T [ 1100°C is consistent with the scarcity of high-T chromites. In the Central sector (Fig. 2), a distinct group of high-T (TNi > 1250°C) garnets has major- and trace-element chemistry (see below) similar 711 JOURNAL OF PETROLOGY VOLUME 40 Fig. 6. TZn–Cr2O3 plot for chromites from the Lac de Gras area; lines show expected maximum Cr2O3 contents of chromites coexisting with garnet [derived from data of Brey et al. (1991)] along the model conductive geotherms of Pollack & Chapman (1977), labeled with corresponding surface heat flow. The Lac de Gras data indicate a low geotherm (30–35 mW/m2) at T Ζ 950°C, and a higher geotherm at higher T, consistent with the garnet and xenolith data. to those of garnets in the sheared high-T xenoliths described above. The Garnet Geotherm therefore reproduces many of the effects seen in the xenolith data, and supports the conclusion that the geotherm in this area consists of two segments, with a higher geotherm at depth and a ‘step’ near 900°C. A plot of TZn vs Cr2O3 (Fig. 6) gives a rough estimate of the geotherm, because the Cr content of chromite in equilibrium with garnet is strongly pressure dependent but relatively insensitive to temperature (Brey et al., 1991). The upper envelope of Cr2O3 at each TZn therefore is controlled by the geotherm. In the Lac de Gras area, this upper envelope suggests a geotherm between the 30 and 35 mW/m2 conductive models at T Ζ 950°C, and a higher geotherm at higher temperatures, which is consistent with the xenolith and garnet data. Trace-element composition of garnets The compositional data for each garnet grain have been assigned a depth by projection of the TNi values to the Garnet Geotherm defined above. This procedure allows evaluation of the vertical variations in individual compositional parameters at each sampling point (Fig. 7), and of lateral variations across the area (Fig. 8). The trace-element data on the garnets (Fig. 7) show a sharp change in mantle chemistry at a temperature of 900°C, NUMBER 5 MAY 1999 corresponding to a depth of ~140 km. Garnets with TNi < 900°C have unusually low contents of large ion lithophile elements (LILE) and high field strength elements (HFSE). The mean contents of Y, Ga, Ti and Zr in these low-T garnets lie in the lower quartile of the values reported for garnets from Archean mantle worldwide (Griffin et al., 1999); on a world-wide scale, they are ultradepleted. Garnets with TNi > 900°C show significantly higher mean and maximum contents of all four elements, and few values as low as even the average value in the lower-T group; there is little overlap between the two sets of data (Fig. 7, Table 1). The garnets with TNi > 900°C are in fact similar in average trace-element contents to the median values of Archean garnets worldwide (Griffin et al., 1999; Table 1; Fig. 7). These traceelement data effectively divide the mantle lithosphere beneath the Lac de Gras area into a shallow ultradepleted layer and a deeper depleted layer with levels of depletion similar to Archean mantle world-wide, and this division extends across the whole Lac de Gras area (Fig. 8). It continues NW to the Ranch Lake pipe, but is interrupted in the intervening area around the Torrie pipe (Fig. 1). The garnets of the shallow layer have extremely low Y/Ga, but elevated Zr/Y, despite the extreme depletion in Zr. The garnets of the deeper layer have Y/Ga ratios more typical of Archean garnets world-wide, and higher Zr/Y, and these variables outline the deeper layer in Fig. 8. Within the deeper layer, the mean values of Y, Zr, Y/Ga and Zr/Y decrease with depth, whereas mean Ti contents increase with depth, at least in the Central sector (Fig. 7). The highest Ti contents, accompanied by high Zr, Y and Ga, are found in the group of high-T garnets noted above; these are compositionally similar to the garnets of the sheared high-T peridotite xenoliths, both in this area (Pearson et al., 1998) and world-wide (Griffin et al., 1999). By analogy with well-studied examples from South Africa, they are interpreted as reflecting metasomatism by asthenosphere-derived melts (Smith & Boyd, 1987; Griffin et al., 1989; Smith et al., 1993). These high-T, trace-element enriched garnets therefore are interpreted as reflecting interaction between lithosphere and asthenosphere (Smith et al., 1993; Griffin & Ryan, 1995), and the boundary between them and the more depleted garnets at ~1250°C (Fig. 7) is taken as the lithosphere– asthenosphere boundary (LAB). Sc contents of garnets in the two layers are similar, although the mean values in the shallow layer are slightly higher, reflecting the greater degree of depletion. However, the mean rare earth element (REE) patterns of garnets from the two layers are very different. In the shallow layer, extreme depletion of the heavy REE (HREE; reflected in high Sc/Y ratios) is accompanied in many cases by enrichment in the light REE (LREE; as shown by high Nd/Y ratios) giving sinuous REE 712 713 Y Zr Ga Sc Y/Ga (av.) Zr/Y (av.) Sc/Y Nd/Y Median values TiO2 Cr2O3 FeO MnO MgO CaO 11·6 33 8·3 122 1·3 3·6 10 0·5 0·234 6·0 7·2 0·36 20·0 5·2 West lherzolite 0·81 Deep layer: Proportion: 2·5 5·0 4·6 152 0·39 1·4 218 1·6 0·075 7·1 7·7 0·45 18·5 6·4 West lherzolite 0·53 Y Zr Ga Sc Y/Ga (av.) Zr/Y (av.) Sc/Y Nd/Y Median values TiO2 Cr2O3 FeO MnO MgO CaO Proportion: Shallow layer: 2·0 5·0 3·7 158 0·44 1·8 168 1·6 0·075 6·3 7·7 0·44 19·8 4·9 6·9 35 6·4 138 1·3 5·6 19 0·5 0·160 7·8 6·9 0·36 20·3 5·1 9·4 34 7·4 130 1·3 4·5 14 0·5 0·199 6·8 7·1 0·36 20·1 5·2 West West harzburgite 0·19 1·5 5·0 2·7 165 0·49 2·2 112 1·5 0·075 5·3 7·6 0·43 21·2 3·2 West West harzburgite 0·47 9·6 30 8·9 127 1·1 4·5 13 0·4 0·215 6·0 7·0 0·35 19·8 5·5 Central lherzolite 0·83 1·8 6·0 5·1 130 0·56 4·1 100 2·5 0·025 6·7 7·5 0·46 18·7 5·9 Central lherzolite 0·42 1·3 3·9 4·6 137 0·47 3·3 149 3·3 0·025 6·9 7·5 0·48 19·7 4·7 3·4 26 5·8 141 0·9 6·1 27 0·9 0·064 8·1 6·7 0·39 20·3 4·7 6·0 28 7·1 135 1·0 5·4 21 0·7 0·127 7·2 6·8 0·37 20·1 5·0 Central Central harzburgite 0·17 0·9 2·3 4·3 142 0·40 2·8 185 3·8 0·025 7·1 7·5 0·49 20·4 3·9 Central Central harzburgite 0·58 11·1 36 8·7 117 1·3 3·4 9 0·2 0·265 6·1 7·1 0·35 19·9 5·3 East lherzolite 0·84 2·4 8·2 5·5 125 0·89 2·9 26 1·1 0·048 5·6 8·0 0·49 18·7 5·7 East lherzolite 0·33 1·4 6·4 4·2 134 0·53 5·2 117 4·7 0·033 6·4 7·5 0·44 20·2 4·2 7·7 38 8·0 128 1·1 5·3 14 0·2 0·213 7·1 6·9 0·36 20·4 4·8 8·8 37 8·2 124 1·1 4·7 12 0·2 0·230 6·8 7·0 0·36 20·2 5·0 East East harzburgite 0·16 0·9 5·5 3·5 138 0·36 6·4 162 6·5 0·025 6·8 7·3 0·42 20·9 3·5 East East harzburgite 0·67 Table 1: Characteristics of garnet concentrates from Lac de Gras, by sector (Fig. 2) 10·6 33 8·7 122 1·2 3·9 11 0·3 0·239 6·0 7·1 0·35 19·9 5·4 All lherz. 2·2 6·7 5·2 132 0·66 3·1 94 1·8 0·044 6·3 7·7 0·47 18·7 5·9 All lherz. 5·8 33 6·8 135 1·0 5·7 20 0·5 0·143 7·6 6·8 0·37 20·3 4·8 All harz. 1·0 4·1 3·7 145 0·40 4·1 161 4·4 0·035 6·6 7·4 0·45 20·8 3·6 All harz. 7·8 33 7·6 130 1·1 4·9 16 0·4 0·183 7·0 6·9 0·36 20·2 5·0 All deep 1·5 5·1 4·3 140 0·49 3·8 140 3·5 0·038 6·6 7·5 0·46 19·9 4·6 All shallow 1·2 3·2 11 31 8·0 0·170 5·3 7·2 0·33 20·4 5·1 Archons (median) 17 54 13·1 120 1·20 3·5 8 0·2 0·625 5·2 7·5 0·29 20·4 5·3 High-T (Central) GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA JOURNAL OF PETROLOGY VOLUME 40 NUMBER 5 MAY 1999 Fig. 7. Plots of Y, Zr, Ti and Cr2O3 vs TNi in garnet xenocrysts from pipes of the Central sector, Lac de Gras, showing the intra-lithospheric boundary at 900°C, and the lithosphere–asthenosphere boundary (LAB; upper limit of Y-depleted garnets) near 1250°C. Symbols as in Fig. 5. patterns like those described from South African concentrate and xenolith garnets by Hoal et al. (1995) and Griffin et al. (1998d). Such patterns also are found in many of the garnets, especially those of harzburgitic paragenesis, from the deeper layer, but they are much less common. Most garnets in the deeper layer have REE patterns with high flat HREE and low LREE, as shown by lower mean values of both Sc/Y and Nd/Y (Table 1). Major-element composition of garnets The distribution of the Cr content of garnets with depth in the Central sector is shown in Fig. 7. The relatively high maximum Cr contents at low T are consistent with a very low geotherm (Griffin & Ryan, 1995). The minimum Cr2O3 contents of garnets from the shallow layer are ~4%, in contrast to values as low as 2% in the deeper layer; this emphasizes the highly depleted nature of the shallow layer. The vertical and lateral distribution of garnet-bearing ultramafic rock types, as derived from garnet compositions, is shown in Figs 7 and 9. The construction of Fig. 9 explicitly assumes that the proportions of different garnet-bearing rock types at depth are reflected in the relative abundance of garnet types (Fig. 3) in the concentrate. This proposition is difficult to test, but has been supported by the detailed studies of Schulze (1989, 1995) on South African xenolith collections and concentrates. In the Lac de Gras area, the high degree of depletion in the garnets, gaps in the TNi spectrum of some garnet concentrates (see below) and high chromite/garnet ratios in some concentrates suggest that many rocks in the shallow layer of the lithospheric mantle are garnet free. 714 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA In this case the relative proportion of strongly depleted rocks (probably harzburgites) estimated from the garnets is likely to be a minimum value. On average, the garnet data indicate that the shallow layer consists of about 60% harzburgite and 40% lherzolite (as defined by the presence or absence of cpx), whereas the deeper layer consists of 83% lherzolite and 17% harzburgite (Table 1). In the deeper layer, the relative abundance of harzburgite is highest near the top of the layer and decreases with depth. In the shallow layer, the relative abundance of harzburgite increases from west to east (Fig. 9), whereas in the deeper layer there appears to be no significant lateral variation in the proportion of the two rock types. Low-Ca harzburgites make up 20–40% of the harzburgitic component in the shallow layer, but <20% of the smaller harzburgite component of the deeper layer; this difference is reflected in the difference in Ca/Cr ratio of the median harzburgitic garnets in the two layers (Table 1). Harzburgitic garnets are absent in the Torrie pipe, where the twolayered structure is interrupted, but appear again in the Ranch Lake pipe; wehrlitic garnets are relatively more abundant in the deeper layer at Ranch Lake than beneath Lac de Gras. Gaul et al. (in preparation) have presented an inversion of the garnet–olivine thermometer of O’Neill & Wood (1979) that can be used to calculate the Fo content of garnet in equilibrium with a given garnet xenocryst of known TNi. When applied to the garnet concentrates from the Lac de Gras area, this technique gives a mean olivine Fo content of 92·7% in the shallow layer and 91·3% in the deeper layer, in good agreement with the more limited xenolith data. Chromite distribution and chemistry In general, the temperature distribution of the chromites parallels that of the garnets, with most chromite grains derived from the upper part of the deeper layer, and a smaller proportion from the shallow layer. In detail, some pipes show gaps in the distribution of garnet TNi, which correspond to peaks in chromite TZn (Fig. 10). This pattern suggests the presence of chromite-rich, garnetpoor horizons. As noted above, the maximum Cr content of the chromites increases with increasing TZn, and then levels off or falls above 950°C, which is consistent with a rise of the geotherm with depth (Fig. 6). Most of the chromites analyzed here have low Ga, Ti and Ni contents, consistent with derivation from harzburgitic or depleted lherzolitic source rocks (Griffin et al., 1994; Griffin & Ryan, 1995). A smaller population, with higher Ga, Ni and Ti contents, and lower Cr contents, is concentrated in the TZn interval 900–1100°C, in the upper part of the deeper layer. These resemble the so-called P2 chromites common in Group 2 kimberlites from the Kaapvaal craton, which earlier have been interpreted as a magmatic population because of their chemical similarity to some groundmass chromites (Griffin et al., 1994). However, more recent data show that similar chromites occur in metasomatized peridotite xenoliths, especially those that contain phlogopite and give equilibration temperatures in the range 900–1100°C (Schulze, 1996; Yao & W. L. Griffin, unpublished data, 1998). The garnets of such xenoliths typically have high Zr and Zr/Y, like many of those in the data presented here (Fig. 7). Both sets of data therefore are consistent with a higher degree of metasomatism near the top of the deeper layer. In the Central sector of the Lac de Gras area (Fig. 2), a third population of lower-Cr chromites with high Ga (>50 ppm), high Ni contents ([1700 ppm) and TZn >1200°C may be magmatic. Scattered examples of such chromites occur in the western sector as well. DISCUSSION A layered lithospheric mantle The garnet data presented above show that the lithospheric mantle beneath the Lac de Gras area consists of two distinct layers, differing markedly in lithology and chemical composition. The shallow layer is ultradepleted, with a high ratio of harzburgite to lherzolite, a high mean Cr content in the garnets, and a distinct scarcity of garnets with Cr2O3 < 1·5%. The garnets in this layer are extremely depleted in ‘incompatible’ elements such as Zr, Ti, and Ga, and even in elements commonly regarded as compatible in garnets, such as Y and the HREE. The median levels of Y, Zr, and Ti in these garnets are much lower than those of Archean garnets world-wide (Table 1), and as a group these garnets are among the most depleted 10% of Cr-pyrope garnets world-wide (Griffin et al., 1999). The mean Y/Ga ratio of the lherzolitic garnets from this layer falls outside the field defined by the mean values for lherzolitic garnets from all other cratons studied by us (Fig. 11; Griffin et al., 1998a, 1999). Data from the available xenoliths and the Fo contents calculated from the garnet xenocrysts both indicate that olivine of the shallow layer is relatively magnesian (Fo 92–94, mean 92·7). The deeper layer is dominated by lherzolitic rocks, and the garnets from this layer are typical in most respects of Archean garnets world-wide (Table 1); this includes the mean Y/Ga and Zr/Y of the lherzolitic garnets (Fig. 11). Xenolith data and Fo contents calculated from the garnet xenocrysts indicate that the deeper layer is less magnesian (Fo 90·5–92·3, mean 91·5) than the shallow layer. 715 JOURNAL OF PETROLOGY VOLUME 40 Fig. 8 Fig. 9 716 NUMBER 5 MAY 1999 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA Fig. 10. Histogram of T estimates for garnet and chromite xenocrysts from pipe EGO-3 in the Western Sector of the Lac de Gras area, showing a ‘gap’ in the garnet spectrum filled in by the chromite spectrum. This pattern is interpreted as reflecting a local layer of highly depleted peridotite, low in garnet but high in chromite, at the boundary between the shallower and deeper layers of the lithosphere. The T interval of ~100 °C corresponds to ~10 km thickness. The boundary between these two layers, as marked by abrupt jumps in the mean Y, Zr and Ti contents of Cr-pyrope garnet xenocrysts, lies at ~150 km depth and can be defined within ±5 km in most localities. It is essentially flat, but across the area (a distance of 60 km), it appears to rise from ~155–160 km in the western sector to 140–145 km in the eastern sector (Fig. 9). In some pipes the boundary is marked by a ‘gap’ in the TNi spectrum of the garnets, which is filled in by a corresponding abundance of chromites (Fig. 10), suggesting the presence of highly depleted peridotite carrying chromite but not garnet. The width of the gap indicates that these layers may be at least 10 km thick in some cases. Similar but less well-defined chromite-rich, garnetpoor layers also occur above the boundary, and rarely below it, across the region. The shallow layer is characterized by a geotherm lying close to the 35 mW/m2 conductive model geotherm of Pollack & Chapman (1977; Figs 4 and 5). At the boundary Fig. 11. Plot of mean Zr/Y vs Y/Ga in xenocryst garnets of the lherzolitic paragenesis world-wide (see Fig. 8) classified on the tectonothermal age of the crust penetrated by the host volcanic rock (Archons, >2·5 Ga; Protons, 2·5–1·0 Ga; Tectons, <1 Ga; after Griffin et al., 1998a). Mean compositions and standard deviations are shown for lherzolitic garnets from the shallow and deeper layers of the Lac de Gras lithosphere. between the layers, the geotherm rises to near an ~38 mW/m2 conductive model geotherm. This stepped geotherm is not typical of most Archean mantle sections. To explain the step with constant heat flow requires that the conductivity of the shallow layer be greater than that in the deeper layer by ~11% (Pearson et al., 1998). Assuming the shallow layer to be extremely olivine rich, and the deeper layer to have the average composition of Archean mantle (olivine 63%, opx 25%, cpx 2%, gnt 4%; Griffin et al., 1998c), and using the mineral conductivity data of Clauser & Huenges (1995), the difference in conductivity between the layers is still only ~2%. The addition of ~10% phlogopite and/or chromite to the deeper layer would be required to lower its conductivity to the required level. We have no direct method for estimating the abundance of these minerals in the deeper layer, but note that the high Zr/Y of many garnets in the upper part of the deeper layer is typical of garnets Fig. 8. (opposite) ‘Chemical tomography’ image showing the vertical and lateral distribution of critical elements and element ratios in garnet xenocrysts along the traverse line shown in Fig. 2. Locations of some pipes along the traverse are shown at lower right. The ultradepleted shallow layer of the lithosphere beneath the Lac de Gras area is defined by the abundances of Y, Zr and Ti; the less depleted deeper layer is best defined by the band of high Y/Ga at 150–200 km depth; the lithosphere–asthenosphere boundary is defined by the drop in Y/Ga and increase in Ti below 200 km depth. The shallow ultradepleted layer is interrupted in the area around the Torrie pipe. Data at depths <100 km and >250 km contain projection artefacts and should be ignored. Fig. 9. (opposite) ‘Chemical tomography’ image showing the vertical and lateral distribution of rock types along the traverse of Figs 2 and 8, derived from the distribution of garnet xenocrysts; rock types defined as in Fig. 3. The concentration of low-Ca harzburgites in the shallow layer of the lithosphere, and the sharp boundary between the layers (see lherzolite distribution) are clearly visible. Lower right panel shows color scale of relative proportions from low (dark) to high (white) for each rock type. 717 JOURNAL OF PETROLOGY VOLUME 40 coexisting with phlogopite in African and Siberian xenoliths (Griffin & Ryan, 1996; Griffin et al., 1998d). In depleted ultramafic compositions such as those of the shallow layer, the spinel peridotite–garnet peridotite transition will occur at depths of ~100 km (O’Neill & Wood, 1979; Kopylova et al., 1998), corresponding to temperatures of ~700°C on the geotherm derived here (Figs 4 and 5). Our samples contain many garnets with TNi [ 700°C, but few with TNi < 700°C (Fig. 7), whereas spinels give TZn down to <500°C (Fig. 6). This suggests that few garnet peridotites occur at depths <100 km, and that the geotherm derived from xenoliths and garnet xenocrysts is consistent with the predicted depth of the spinel–garnet peridotite transition. These data also imply that spinel peridotites make up most of the shallow mantle between 100 km and the crust–mantle boundary, which probably lies at 35–40 km depth (Clowes, 1997). Estimation of lithosphere thickness by the use of mineral data has been discussed by Griffin & Ryan (1995), Ryan et al. (1996) and Griffin et al. (1998c). In this work, we have adopted a geochemical approach, defining the base of the lithosphere as the temperature above which depleted garnets no longer appear. For this purpose, ‘depleted’ garnets are defined as those with <10 ppm Y, as this value is close to the median Y content of >6000 garnets from Archean cratons (Griffin et al., 1999). Higher-T garnets tend to have undepleted trace-element chemistry similar to the garnets of high-T sheared xenoliths, which are interpreted to have been infiltrated by asthenosphere-derived melts (Smith & Boyd, 1987; Griffin et al., 1989b; Smith et al., 1993). In many cratonic areas, the lithosphere thickness determined in this manner ranges from 180 to 250 km; the lithosphere– ‘asthenosphere’ boundary so defined lies within the thicker ‘tectosphere’ of Jordan (1988) and probably represents the upper limit of pronounced magma–wall rock interaction (O’Reilly & Griffin, 1996). In the Lac de Gras area, an upper limit to Y-depleted garnets can be defined beneath only the Central sector, where a higher-T undepleted garnet population is present. This boundary lies at ~1250°C (Fig. 6), corresponding to a depth of ~220 km. In the eastern and western sectors there are few garnets with TNi > 1200°C, and the data define only a minimum lithosphere thickness of ~200 km. The techniques used here do not allow a detailed estimate of the relative abundances of ultramafic and eclogitic rocks within the section, and because the Ni thermometer is not applicable to garnets from olivinefree rocks, we cannot estimate the depth of origin of eclogitic garnets as is done here for peridotitic garnets. In general, the proportion of eclogitic garnets to peridotitic garnets in the concentrates is <10%, suggesting that eclogite makes up p10% of the total section sampled by the kimberlites studied here. As noted above, the T estimates for the eclogite xenoliths show a bimodal NUMBER 5 MAY 1999 distribution, with one peak centered on ~950°C, and the other near 1200°C (Pearson et al., 1998). These temperatures suggest that eclogites are largely, and perhaps wholly, confined to the deeper layer of the lithosphere. Eclogites with high-Ca garnets ([9% CaO) give higher temperatures, and thus are most abundant near the lithosphere–asthenosphere boundary. Those containing garnets with <7% CaO are concentrated higher in the section, near the boundary between the shallow and deeper layers. Comparison with other cratons The average composition of 14 peridotite xenoliths from pipe A154S, calculated from their modal and mineral analyses, is shown in Table 2, and compared with averages of specific xenolith types from southern African and Siberian kimberlites. Our sample contains few xenoliths from the upper layer, but these data, combined with those of Boyd & Canil (1997, and unpublished data, 1997), indicate that the rocks of this layer are very olivine rich, and have moderately magnesian olivine [Fo92·8; compare Boyd & Canil’s (1997) value of Fo92·9, and the Fo92·7 calculated from garnet xenocrysts]. The xenoliths from the deeper layer, although less refractory, also have high mean olivine contents (mean Fo91·5; compare Fo91·3 calculated from garnet xenocrysts). The two groups are considered together here, because there are so few data. Both groups have lower Si and higher Mg than most other Archean xenoliths, reflecting a higher olivine/opx ratio. They are similar in this respect to some garnet harzburgites from the Udachnaya pipe. However, it should be noted that the Udachnaya xenoliths have experienced significant secondary introduction of Ca and Fe, as shown by their abnormally high Ca/Al and Fe/ Al (Boyd et al., 1997), and the primary CaO content of the Udachnaya harzburgites, calculated from modal analyses, is closer to 0·2% (Griffin et al., 1998c). A high opx/olivine ratio at high percent Fo has been regarded as one of the distinctive features of the Archean lithospheric mantle, in contrast to younger mantle, and considerable controversy surrounds the explanation for this feature (Boyd, 1989; see review by Griffin et al., 1998c). If the average composition for the Slave xenoliths studied here is meaningful, then the mantle beneath the Lac de Gras area is different from the Kaapvaal and Siberian cratons, in having a lower opx/olivine ratio, more characteristic of highly depleted types of circumcratonic mantle than of ‘typical’ Archean mantle. However, this result should be treated with caution. The available xenoliths are small (1–2 cm), and are restricted to a population that has been able to survive industrialgrade crushers; observations on drill core show that many xenoliths in the A154 kimberlite are extremely friable. 718 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA Table 2: Average compositions of xenolith suites Range Olivine Av. Kaapvaal Av. Daldyn average A154S xenoliths SD xenoliths xenoliths (n = 14) (n = 96) (n = 12)∗ 56·6–94·1 79 14 65 71 OPX 1·6–24·9 13 10 24·2 20·3 CPX 0·5–3 2·3 0·8 1·6 3·6 Garnet 1–25·4 6·6 6·1 6·8 Phlog 0–2·5 0·3 0·8 0·2 Spinel 0–0·6 5·3 — 0·2 0·2 0·2 0·1 43·5 1·9 46·5 43·7 SiO2 40·9–47·6 TiO2 0·001–0·146 0·04 0·04 0·06 0·05 Al2O3 0·28–5·42 1·28 1·31 1·40 0·92 Cr2O3 0·27–1·3 0·59 0·33 0·34 0·37 FeO 6·62–8·27 7·3 0·6 6·6 7·5 MnO 0·09–0·15 0·12 0·02 0·10 MgO 39–50 CaO 0·28–1·35 0·66 0·40 0·88 Na2O 0·01–0·13 0·04 0·04 0·10 0·003–0·23 0·03 0·06 0·02 K 2O 46·4 3·2 0·32 0·05 43·8 0·29 0·12 46·0 0·94 0·07 — NiO 0·23–0·39 mg-no. 90·6–92·8 91·9 0·6 Mg/Si 1·3–1·79 1·6 0·2 1·41 1·57 Ca/Al 0·33–1·12 0·47 0·88 0·57 0·93 Cr/(Cr + Al) 0·14–0·37 0·24 0·07 0·14 0·21 Fe/Al 2·3–17·9 4·1 5·4 3·4 5·8 T (°C) (ONW) 780–1163 92·3 0·30 92·4† ∗Boyd et al. (1997), low-T garnet peridotites. Oxides from all low-T xenoliths (n = 21). †mg-number of olivine (Fe and Ca affected by metasomatism). The population studied here therefore may be strongly biased toward olivine-rich compositions, and the modal analyses have large uncertainties as a result of the size constraint. However, as noted above, the mean Fo contents of olivine, calculated from the garnet concentrates, are in good agreement with the xenolith data. An alternative approach to estimating the composition of the lithospheric mantle uses the compositional relationships between Cr-pyrope garnets and their host rocks. Griffin et al. (1998c) showed that the Cr contents of peridotitic garnets are well correlated with the Al2O3 contents of their host xenoliths, and that other wholerock oxide contents correlate in turn with Al2O3. These relationships have been used to define a series of equations that allow calculation of a meaningful average rock composition from the median Cr content of a garnet concentrate. Table 3 shows calculated mean compositions derived from lherzolitic and harzburgitic garnets (as defined from Fig. 3) from the Kaapvaal and Siberian cratons, which can be compared with the mean compositions of lherzolitic and harzburgitic xenoliths given in Table 2. The agreement between the calculated compositions and the mean xenolith compositions is excellent, except in the case of the Fe and Ca contents of the Siberian xenoliths; as noted above, these have been affected by secondary addition of these elements (Boyd et al., 1997). Application of this approach to the concentrate garnets from Lac de Gras meets two potential problems. (1) The algorithms used for this calculation are derived from African and Siberian xenoliths, most of which have high opx/olivine ratios. If the Slave xenoliths are significantly different in this respect, the results may be biased toward unrealistically high Si/Mg, causing an artificial similarity to the Siberian and African lithospheres. (2) The method depends on the close relationship between Cr in garnet 719 JOURNAL OF PETROLOGY VOLUME 40 NUMBER 5 MAY 1999 Table 3: Calculated subcontinental lithospheric mantle (SCLM) compositions for Archean cratons Slave Upper Slave Upper Slave Lower Slave Lower Upper Lower lherz harz lherz harz mean mean Proportion: 0·4 0·6 0·83 Garnet Cr2O3 6·3 6·6 6 7·6 6·5 6·3 Rock Al2O3 0·68 0·61 0·76 0·43 0·64 0·70 SiO2 45·56 45·53 45·59 45·44 45·54 45·57 45·73 TiO2 0·03 0·02 0·03 0·02 0·03 0·03 0·04 Al2O3 0·68 0·61 0·76 0·43 0·64 0·70 1·05 Cr2O3 0·23 0·22 0·24 0·19 0·22 0·23 0·29 FeO 6·24 6·21 6·28 6·12 6·22 6·25 6·43 MnO 0·11 0·11 0·11 0·11 0·11 0·11 0·11 MgO 46·39 46·58 46·19 47·08 46·51 46·34 45·36 CaO 0·41 0·37 0·45 0·26 0·38 0·42 0·63 Na2O 0·05 0·04 0·05 0·03 0·04 0·05 0·07 NiO 0·31 0·31 0·31 0·32 0·31 0·31 0·29 mg-no. Av. Archean 0·17 92·99 93·05 92·92 93·21 93·03 92·97 92·64 Mg/Si 1·52 1·53 1·51 1·55 1·53 1·52 1·48 Ca/Al 0·55 0·55 0·55 0·55 0·55 0·55 0·55 Cr/Al 0·49 0·52 0·47 0·64 0·51 0·48 0·40 Cr/(Cr + Al) 0·18 0·19 0·18 0·23 0·19 0·18 0·16 Fe/Al 6·60 7·28 5·99 10·14 6·99 6·43 4·41 Mean comp. Daldyn Malo Bot. Kaapvaal <90 Kaapvaal >90 Liaoning Shandong Venezuela SiO2 45·68 45·81 45·98 45·87 45·49 45·65 45·55 TiO2 0·04 0·05 0·06 0·05 0·02 0·04 0·03 Al2O3 0·94 1·24 1·60 1·39 0·53 0·88 0·68 Cr2O3 0·27 0·32 0·38 0·34 0·21 0·26 0·23 FeO 6·37 6·52 6·70 6·60 6·17 6·34 6·24 MnO 0·11 0·11 0·12 0·11 0·11 0·11 0·11 MgO 45·67 44·84 43·83 44·40 46·81 45·83 46·41 CaO 0·56 0·74 0·96 0·84 0·32 0·53 0·41 Na2O 0·07 0·09 0·11 0·10 0·04 0·06 0·05 NiO 0·30 0·29 0·27 0·28 0·31 0·30 0·31 mg-no. 92·75 92·47 92·11 92·31 93·13 92·80 92·99 Mg/Si 1·49 1·46 1·42 1·45 1·54 1·50 1·52 Ca/Al 0·55 0·55 0·55 0·55 0·55 0·55 0·55 Cr/Al 0·42 0·38 0·34 0·36 0·56 0·43 0·49 Cr/(Cr + Al) 0·16 0·15 0·14 0·14 0·21 0·17 0·19 Fe/Al 4·88 3·80 3·02 3·41 8·35 5·18 6·64 and Al in the whole rock. This relationship breaks down if chromite becomes a significant phase, and as noted above, the narrow range of PCr at each TNi in the shallower layer (Fig. 5) indicates that most garnets in this layer have coexisted with chromite. If this is true then the method will underestimate the degree of depletion in the shallower layer. The median rock compositions calculated from lherzolitic and harzburgitic garnets in each of the layers are given in Table 3, and combined medians are given, 720 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA calculated using the proportions of each garnet type in the concentrates. The deeper layer, for which the estimated composition is most likely to be correct, is somewhat more depleted in terms of Ca, Al and Fe than the average Archean lithosphere of Griffin et al. (1998c). This is consistent with the lower mean Y and Ga of the garnets (Table 1). Among other sections that we have analyzed, the deeper layer is most similar to the Archean lithosphere beneath the Guaniamo Basin of Venezuela or Shandong Province, China (Griffin et al., 1998c). The composition of the shallower layer, calculated in this manner, is only slightly more depleted than that of the deeper layer; this is not consistent with the extreme depletion of the shallower garnet population in Ti, Y, Ga and Zr. The relatively high calculated values of these incompatible elements in the shallow layer directly reflect the relatively low Cr content of the depleted garnets of the shallow layer caused by their coexistence with spinel, as noted above. Both layers show calculated Mg/Si ratios similar to those of other Archean sections, which is inconsistent with the xenolith data. These discrepancies illustrate the limitations of the method, described above. The distribution of rock types with depth beneath several Archean cratons, constructed from the temperature distribution of garnets of different types, is shown in Fig. 12. Some of these, such as the Daldyn field of Siberia (Griffin et al., 1996) and Liaoning Province of China (Griffin et al., 1998f ), show a pronounced stratification defined by a concentration of harzburgitic rocks in relatively restricted depth ranges. The Lac de Gras mantle differs from these in several respects: the strong concentration of harzburgite occurs at shallower depth, the degree of depletion in the garnets of both harzburgites and associated lherzolites is much greater, and the boundary between the shallower and deeper layers is far more pronounced. However, the general increase in the abundance of lherzolites with depth, which is pronounced beneath the Lac de Gras area, is a feature shared with the Daldyn and Kaapvaal sections. The Y/Ga–Zr/Y relationships in mantle-derived lherzolitic garnets show a progressive change from Archean to Proterozoic to Phanerozoic lithospheric sections; this trend reflects a secular evolution to progressively less depleted mean mantle compositions (Fig. 11; Griffin et al., 1998a, 1998c, 1999). In this plot, the lherzolitic garnets from the deeper layer fall in the same region as those from several other Archean (and some Early Proterozoic) areas. The lherzolitic garnets from the shallower layer, in contrast, plot outside the fields defined by other data sets, emphasizing their unusually depleted nature. Within the limitations of the available data, we conclude that the deeper layer of the Lac de Gras lithosphere is broadly similar to the mantle beneath many other Archean cratons, but is somewhat more depleted than many such sections. The shallower layer, in contrast, is significantly more depleted. The sharply defined stratigraphy beneath the Lac de Gras area appears to be unique, within our present knowledge of the Archean mantle. Geophysics Geophysical data also suggest that the lithospheric mantle beneath the Slave Province is not identical to that under other cratons. Data from a teleseismic array near Yellowknife indicate that there is little difference in seismic velocity (Vp) under the Slave Province and the adjoining Proterozoic orogens to the west and south (Bostock & Cassidy, 1997). The mean Vp is higher than in the standard IASP91 Earth model of Kennett & Engdahl (1991), but lower than that beneath other Archean cratons. Hoffman (1990) suggested that a high-velocity (Vs) root occurs beneath the Slave Province, but Grand (1994) presented data that suggest a low Vs beneath the craton, relative to surrounding areas. Although the lateral resolution of the tomography is only on the same order as the dimensions of the craton, it appears that the lithospheric mantle root present in Eocene time, and sampled by the kimberlites studied here, does not show up well in seismic data. The seismic velocity of lithospheric mantle is directly related to both temperature and composition, and particularly the mean mg-number. The xenolith and garnet data cited above indicate that olivine in the shallow layer has a mean Fo content of ~92·8%, similar to the mean value of 92·5–93% common in other cratonic roots; the mean Fo content of olivine xenoliths from the deeper layer, however, is 91·5%. The xenolith data also suggest that the lithospheric mantle beneath the Slave Province is more olivine rich than other cratonic roots. The less magnesian nature of the deeper layer results in lower seismic velocities relative to other Archean roots, whereas an increase in the relative abundance of olivine will tend to increase velocities. More detailed modal data on both layers will be required to evaluate the relative contributions of these two factors to the seismic velocities beneath the craton. It also is possible that the area underlain by the ultradepleted shallow layer, and the relatively thick lithosphere beneath it, is too small to register on the seismic tomography, which has a lateral resolution of [200 km. Thompson et al. (1995) used heat flow data to infer that the lithosphere beneath the Lac de Gras area is thicker than that beneath the Yellowknife area, and Griffin et al. (1998e) found that kimberlites near the southern and SW margins of the craton have sampled a lithosphere that not only is thinner (~180 km) than beneath the Lac de Gras area, but may consist entirely of material like the deeper layer described here. 721 JOURNAL OF PETROLOGY VOLUME 40 NUMBER 5 MAY 1999 Fig. 12. Lithologic sections for representative Archean mantle lithosphere sections, constructed from the TNi distribution of xenocryst garnets and the derived geotherm; rock types defined as in Fig. 3. Unlike most other areas world-wide, the lower crust beneath the Anton terrane in the western part of the Slave province has a low conductivity, which makes it possible to use magnetotelluric data to image both the crust–mantle boundary and the base of the (electrical) lithosphere (Clowes, 1997). A large increase in conductivity at 250–300 km beneath this terrane is interpreted as the lithosphere–asthenosphere boundary, and the preliminary data from the SNORCLE transect suggest that this boundary rises to the west, and possibly to the east (Clowes, 1997). Origin of the layered structure The strongly layered lithospheric mantle beneath the Lac de Gras area is unique within our large database of sections through Archean mantle (Griffin et al., 1998c; Fig. 12). The sharp boundary between the two layers strongly suggests a two-stage process for the formation of this lithospheric root. Here we speculate on a possible mechanism for constructing such a root. The deeper layer of the lithosphere is generally similar in terms of garnet composition and rock-type makeup to other Archean lithospheres, but somewhat more depleted than most, especially the Kaapvaal craton. If the xenolith data are representative, it may have a higher mean olivine/opx ratio and less magnesian mean olivine composition than most other Archean cratonic lithosphere, but the limited data require caution about this conclusion until larger suites of large xenoliths become available. The garnet and chromite data indicate that the content of harzburgite is higher near the upper boundary, but in general the Y and Zr contents of the garnets in the upper part of this layer also are higher, suggesting a metasomatic enrichment near the intra-lithospheric boundary. Two important pieces of information, derived from the diamonds and from the eclogitic xenoliths, bear on the origin of this deeper layer. (1) The Lac de Gras diamond population contains a high proportion of stones with inclusions of the ‘superdeep’ ferropericlase ± Mgperovskite assemblage, which is believed to be stable only at depths of [650 km, below the transition zone between upper and lower mantle (Scott-Smith et al., 1984; Kesson & Fitz Gerald, 1991; Harte et al., 1998). The few available temperature estimates for diamonds from the Lac de Gras area are consistent with derivation from the deeper layer of the lithosphere (Davies et al., 1998). If we assume that this also applies to the ultra-deep diamonds (for which no independent temperatures are available), then the presence of these diamonds is strong evidence that the material of the deeper layer has been transported as a plume or diapir from the lower mantle (Haggerty, 1994). (2) T estimates for all eclogites studied so far (Pearson et al., 1998) place them in the deeper layer. One eclogite population with extremely calcic garnets, jadeitic pyroxenes and kyanite has bulk compositions similar to plagioclase-rich crustal rocks. Diamonds with inclusions of this eclogitic paragenesis have isotopically light carbon, suggesting a crustal origin (Davies et al., 1998). The diamond-inclusion data are critical in any model: they imply that the deeper layer of the Lac de Gras mantle contains a high proportion of material from the lower mantle. We therefore suggest that the deeper layer 722 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA of the Lac de Gras lithosphere represents material that rose diapirically from the lower mantle, and may consist of a mixture of previously subducted slab and entrained deep mantle material, perhaps modified by partial melting during ascent. This material was accreted to the base of the pre-existing depleted shallow layer, apparently along an essentially horizontal boundary and with little or no mixing. The eclogites (± kyanite) with Ca-rich garnets may be part of this package, representing originally plagiocaserich lithologies of the oceanic crust. However, with one exception, the diamond-inclusion garnets with corresponding high Ca contents are not majoritic (Davies et al., 1998), as would be expected if they crystallized at such great depths (Moore et al., 1991). This model therefore would require the eclogitic diamonds to have crystallized at depths <200 km in the descending slab, and to have been preserved during the subsequent rise of the diapir. Alternatively, they might represent younger oceanic crust, subducted and tectonically underplated on the base of the lithosphere during Proterozoic subduction from either the east (Thelon Orogen) or west (Wopmay orogen). Lithoprobe seismic data (Clowes, 1997) show mantle reflectors dipping eastward beneath the craton from the Great Bear magmatic arc, interpreted as related to 1·9 Ga magmatism. This subduction origin would be consistent with the position of the Ca-rich eclogites in the deepest part of the Eocene lithosphere, and with the wide range in d13C of the eclogitic diamonds, which may be a signature of post-Archean diamonds (McCandless & Gurney, 1997). If the Ca-rich eclogites do reflect Proterozoic subduction processes, it implies that the deeper layer of the lithosphere was in place by ~2 Ga. Is it possible to constrain the origin of the shallow layer? In modern lithospheric environments, the extreme depletion seen in the shallow layer is common only among abyssal peridotites and the tectonite horizons of some ophiolites from convergent margins (Menzies, 1991; Griffin et al., 1998c). This comparison may be relevant, because the turbidites and felsic to basaltic volcanic rocks that make up much of the Contwoyto Terrane have been interpreted as an accretionary wedge along a convergent margin (Padgham & Fyson, 1992; Davis et al., 1994; Griffin et al., 1998e), and the widespread pre- and syntectonic granitoids (2·6–2·7 Ga) are calc-alkaline in nature and generally resemble modern subduction-related felsic igneous rocks. This is an environment where oceanic peridotites might be expected at shallow levels. Modern ophiolitic or oceanic peridotites have several features that distinguish them from Archean peridotites of similar levels of depletion (Griffin et al., 1998c): inter alia, they tend to have higher FeO contents (median values 7–8% vs 6·5–7·5%), and higher Cr/(Cr + Al) (median values 0·2–0·8 vs 0·14–0·24). The limited xenolith data from the shallow layer do show higher FeO and Cr/(Cr + Al) than the mean values for Kaapvaal and Daldyn xenoliths (Table 1). The shallow layer of the lithospheric mantle therefore could represent the type of depleted mantle seen in such convergent-margin settings today. The extensive post-tectonic granitoids of the Slave Province, however, have isotopic and chemical characteristics more similar to K–U–Th-rich Phanerozoic post-orogenic granites, occur over a wider area, including the western part of the craton underlain by ancient continental crust, and are accompanied by widespread low-P, high-T metamorphism. Their genesis requires a major heat source that is not restricted to the former continental margin, and Davis et al. (1994) invoked lithosphere delamination as a mechanism for heating the crust. However, the unique two-layer structure of the mantle lithosphere suggests an alternative possibility, related to plume tectonics. In our model (Fig. 13), the shallow layer represents strongly depleted lithosphere formed at the active convergent margin, and accreted under the newly formed continental crust, represented by the Yellowknife Group and the 2·6–2·7 Ga granitoids. This mantle may have become hydrated as a result of processes above the subduction zone (Menzies, 1991; O’Reilly & Griffin, 1988). Late-tectonic to syn-tectonic heating was supplied by the arrival of a plume head, represented by the deeper layer of the present mantle lithosphere, at ~2·6 Ga. The rise of this plume was stopped by the presence of the refractory and buoyant lithosphere already in place; the heating may have caused remobilization of any lowmelting fraction within this layer, leading to further depletion and leaving an extremely refractory residue, which could transmit heat to the overlying crust and produce the late granitoids by remelting of the earlierformed crust. This thick refractory lithosphere might also act as a density filter, trapping basaltic melts from the rising plume and forcing them to underplate, perhaps producing the large population of low-Ca eclogites, and associated metasomatism, near the boundary between the two layers. If this model is relevant in the Lac de Gras area, it may be applicable to other cratons, especially those in which diamonds of the superdeep paragenesis are found. The stratification of rock types seen in some cratonic lithospheric sections (Fig. 12) may be related to such multi-stage generation of cratonic roots. In most cases studied so far, really sharp horizontal boundaries within the mantle, like the one beneath the Lac de Gras area, have not been observed. This might reflect greater degrees of mixing at such an interface, as a result of smaller contrasts in composition and rheology than were present beneath Lac de Gras, where the shallow lithosphere was already ultradepleted. More detailed study of the distribution of depletion signatures in mantle-derived xenocrysts from selected mantle sections can be used to 723 JOURNAL OF PETROLOGY VOLUME 40 test these models, and to evaluate the extent to which accumulation of plume-related material has contributed to the construction of subcontinental lithospheric mantle. A broader consequence of this model is its implication that diamonds of the ‘superdeep’ paragenesis may have been brought to the lower part of the lithosphere at various times throughout Earth history, and resided there NUMBER 5 MAY 1999 until entrained by much later kimberlite eruptions. This supplements the model proposed by Haggerty (1994), in which the superdeep diamonds are transported from the deep mantle in ‘superplumes’ and erupted without significant lithospheric residence, in kimberlites that are part of the ‘superplume’ process. Our model also suggests that rising plumes might have emplaced superdeep diamonds in the deeper parts of the subcontinental lithosphere episodically throughout Earth history, as part of the processes that generate the subcontinental lithospheric mantle. This process might help to explain some ‘anomalous’ occurrences of diamonds in non-cratonic settings (Griffin et al., 1998b). CONCLUSIONS (1) The mantle lithosphere beneath the Lac de Gras area has a structure and composition that are unique within our limited knowledge of Archean mantle sections. This observation suggests that the current model of Archean lithosphere, based largely on xenoliths from the kimberlites of the Kaapvaal craton and the Udachnaya pipe in Siberia, is not adequate. (2) The compositions of garnet and chromite xenocrysts, and limited xenolith data, indicate that the portion of the lithospheric mantle shallower than ~145 km beneath this area is extremely depleted in LILE and HFSE compared with other Archean mantle, has a higher harzburgite/lherzolite ratio, and may have a higher olivine/opx ratio. (3) The deeper layer of the lithosphere (~145–200 km) is generally similar to, for example, the Kaapvaal lithosphere as sampled by Group I kimberlites, but the harzburgite/lherzolite ratio decreases with depth. The limited xenolith data suggest that this layer also may have a higher olivine/opx ratio than mean Kaapvaal lithosphere. Eclogites are present mainly, and perhaps exclusively, in the deeper layer. Extremely depleted (garnet-free) peridotites may form layers as much as 10 km thick near the boundary between the two layers. Fig. 13. Schematic representation of suggested evolution of the Lac de Gras lithospheric mantle. The ultradepleted shallow layer of the lithosphere is generated as oceanic and sub-arc mantle before and during accretion of the Hackett and Contwoyto terranes (magmatic arc and accretionary wedge, respectively) to the ancient continent (Anton terrane). The deeper layer is added by a plume head ascending from the lower mantle, carrying the superdeep diamond population; associated heating produces the widespread 2·6–2·7 Ga postorogenic granitoid magmatism. Proterozoic subduction from east (Thelon orogen) and/or west (Wopmay orogen) introduces eclogite near the base of the lithosphere. 724 GRIFFIN et al. LITHOSPHERIC MANTLE OF LAC DE GRAS AREA (4) The ultradepleted shallower layer is interpreted as ancient oceanic or sub-arc mantle formed during accretion of the Hackett and Contwoyto arc–accretionary wedge terranes to the Anton continent. (5) The deeper layer contains abundant diamonds with mineral assemblages indicating derivation from the lower mantle. This layer is interpreted as a fossil plume head that rose from the lower mantle and underplated the existing ultradepleted lithosphere, probably near 2·6 Ga, providing the heat source for large-scale post-tectonic granite magmatism. (6) Eclogites may have been part of the mantle plume; the more calcic varieties alternatively may have been emplaced near the base of the cratonic root during Proterozoic subduction events (Thelon and Wopmay orogens). ACKNOWLEDGEMENTS We are grateful to Tin Tin Win for assistance with the proton microprobe analyses, and to Ashwini Sharma and Carol Lawson for assistance with the LAM-ICPMS analyses and EMP work, respectively. Sally-Ann Hodgekiss prepared the graphics with flair, style and great patience. Juanita Bellinger took care of many aspects of the sample selection and preparation, aided by several Kennecott and Aber geologists. We thank Joe Boyd and Maya Kopylova for access to unpublished data. The manuscript was improved by reviews from Don Francis, Else-Ragnhild Neumann and Martin Menzies. This work was supported by Kennecott Canada through Macquarie University collaborative grants, and by ARC funding to W.L.G. and S.Y.O’R. This is Contribution 139 from the ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents. REFERENCES Bleeker, W., Ketchum, J. W. F., Jackson, V. A. & Villeneuve, M. (1999). The central Slave Basement Complex. Part I: Its structural topology and autochthonous core. Canadian Journal of Earth Sciences 36 (in press). Bostock, M. G. (1997). Anisotropic upper-mantle stratigraphy and architecture of the Slave craton. Nature 390, 392–395. Bostock, M. G. & Cassidy, J. F. (1997). Upper mantle stratigraphy beneath the southern Slave craton. Canadian Journal of Earth Sciences 34, 577–587. Boyd, F. R. (1989). Composition and distinction between oceanic and cratonic lithosphere. Earth and Planetary Science Letters 96, 15–26. Boyd, F. 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