Download Layered Mantle Lithosphere in the Lac de Gras

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
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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
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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
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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
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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,
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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
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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
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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
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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
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JOURNAL OF PETROLOGY
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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
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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.
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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.
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Fig. 8
Fig. 9
716
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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.
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JOURNAL OF PETROLOGY
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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.
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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
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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.
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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
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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
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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.
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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.
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