Journal of the Geological Society
Formation of Archaean continental lithosphere and its diamonds: the root of the
problem
D.G. PEARSON and N. WITTIG
Journal of the Geological Society 2008; v. 165; p. 895-914
doi:10.1144/0016-76492008-003
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© 2008 Geological Society of London
Journal of the Geological Society, London, Vol. 165, 2008, pp. 895–914. Printed in Great Britain.
Review
Formation of Archaean continental lithosphere and its diamonds:
the root of the problem
D. G . P E A R S O N & N. W I T T I G
Northern Centre for Isotopic and Elemental Tracing, Department of Earth Sciences, Durham University, South Road,
Durham DH1 4QE, UK (e-mail: [email protected])
Abstract: Cratonic lithospheric mantle plays an integral role in defining the physical behaviour of ancient
continents and their mineral potential. Bulk compositional data show that modern-day melting residues from a
variety of tectonic settings can be as depleted in Al and Ca as cratonic peridotites. Cratonic peridotites are
strongly affected by secondary introduction of pyroxenes and garnet such that the extent and depth of melting
cannot be reliably determined. Olivine compositions are probably the most reliable tracer of the original
melting process and indicate that typical cratonic peridotites have experienced 40% or more melt extraction.
Homogeneous levels of depletion indicated by olivine compositions, combined with mildly incompatible trace
element evidence, indicate that melting took place at shallow depths, dominantly in the spinel stability field.
Consideration of melt production models shows that shallow (,3 GPa) anhydrous melting is not capable of
producing residues dominated by large degrees of melt extraction. Instead, a critical role for water is indicated,
implicating the formation of cratonic peridotites within Archaean subduction zones. This melting occurred in
the Neoarchaean in some cratonic blocks, initially forming dunitic residues that are still evident in the xenolith
inventory of some cratons. Release and migration up-section of siliceous melt produced during orthopyroxene
breakdown metasomatizes the proto-lithospheric via re-enrichment in orthopyroxene crystallizing from this
hydrous Si-rich melt, forming the variably orthopyroxene-rich refractory harzburgites typical of most cratonic
roots. Melting in Archaean subduction zones is followed by subduction stacking to form the cratonic root.
Gravitational forces may then be responsible for the loss of imbricated mafic crust during periods of transient
thermal and physical disturbances prior to final long-term tectonic stability. Most diamonds form in the base
of these cratonic roots during pulses of thermal or tectonic activity, initially during root construction and
subsequently associated with large-scale regional lithospheric events that may be correlated to pulses in global
mantle dynamic evolution.
age of formation in relation to their associated crust, and we
examine the timing of diamond formation in relation to the
evolution of the lithosphere.
The stabilization of the first large robust continental land masses
seems inexorably linked to the protective presence of thick,
strong lithospheric roots. In addition to playing a critical role in
defining the longevity of continents, through their unwillingness
to be rifted apart, these roots beneath the continents play host to
the Earth’s reserves of diamonds (Stachel & Harris 2008). The
density, thermal structure and rheology of continental lithosphere
also exercises a profound influence on continental dynamics,
relative sea-level change and climate change. Hence, understanding the formation and evolution of continental roots is
central to many issues is Earth Sciences.
In this review we will focus primarily on the effects of melt
depletion in cratonic lithospheric mantle, together with its
consequences and implications for the formation and evolution
of cratonic roots. For several decades the depth and environment
of melting of cratonic lithosphere was a subject of fierce debate,
with a multitude of models put forward ranging from deep plume
melting (Boyd 1989; Griffin et al. 1999) to shallow melting and
subduction (e.g. Schulze 1986; Canil & Wei 1992). More
recently the consensus is moving towards the original melt
extraction taking place at shallow mantle depths prior to subduction stacking (Canil 2004; Lee 2006; Simon et al. 2008). Below
we outline recent and new constraints on the melting environments that produced the first continental roots, we address their
Cratons: records of the early history of the Earth
The history of the early evolution of the Earth is contained
within and beneath regions of continental crust in excess of
2.5 Ga old, known as cratons (Fig. 1). The oldest rocks within
these cratons are 4 Ga old (Bowring et al. 1990), although
zircons have been recovered that are older than 4.4 Ga, covering
over 96% of Earth’s history (Fig. 2). Although no cratonic
lithosphere has yet been found that approaches the age of the
oldest zircons, the recent discovery of diamonds included in
Hadean zircons (Menneken et al. 2007) has raised the possibility that deep continental material may have existed in parts of
the Earth as early as 4.2 Ga ago. The oldest terrestrial record of
melting within the mantle comes from very unradiogenic Os
isotope ratios recorded in Re-poor, Os-rich osmiridium grains
(Fig. 2). Some detrital grains within the Witwatersrand basin
yield Os model ages of 4.1 Ga (Malitch & Merkle 2004) and
are hence within 90% of the age of the Earth. We will return to
the significance of these ancient events below.
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896
D. G . P E A R S O N & N. W I T T I G
Fig. 1. Global distribution of the cratons (regions of crust .2.5 Ga old) modified after Bleeker (2003). Regional outcrops of Archaean crustal rocks are
indicated in grey (those beneath Greenland are extrapolated under the ice cap) and other definable fragments of composite cratons in brown (from Bleeker
2003). The approximate outline of units that are relatively well defined as whole, composite cratons is shown by black dotted lines. Red dashed lines show
the estimated extent of cratonic regions amalgamated from Archaean blocks during the Proterozoic. Blue dotted lines extended across oceanic areas show
links between cratonic fragments that are thought to have once formed single cratonic blocks. NAC, North Atlantic Craton.
The nature of continental roots
Xenoliths: a biased view?
Several recent reviews have documented in detail the chemical
and physical character of the rocks that make up continental
roots (Menzies 1990; Pearson et al. 2003; Carlson et al. 2005;
Lee 2006) and the reader is referred to those works for details.
Here we will focus on a limited number of features relevant to
the issues being addressed. Mineral concentrate data from
kimberlite pipes indicate that the dominant rock type forming the
lithosphere beneath cratonic crust is peridotite, with eclogite and
pyroxenitic assemblages making up less than 2% of the volume
(Schulze 1989). Although this estimate is consistent with
seismological evidence it is important to realize the nature of the
sample upon which our view is based. Almost all our samples of
cratonic lithospheric mantle originate as xenoliths carried to the
Earth’s surface by volcanic rocks, predominantly kimberlites
(e.g. Pearson et al. 2003). The elemental chemistry and mineralogy of these samples are often strongly influenced by activity
that is directly or indirectly related to the host volcanic rock and
its precursory activity. Hence the likelihood is that the xenolith
sample, in most instances, samples the wall rock of magma
conduits that have been extensively modified by the host melt
prior to, during and after eruption. Hence the sample provided by
xenoliths, or their disaggregated counterparts (mineral concentrates) is not necessarily representative of the lithosphere as a
whole. Increasing evidence indicates that a significant proportion
of the clinopyroxene and some garnet included in cratonic
peridotites are of secondary origin, related to metasomatic
events, either introducing the phases, in some cases relatively
recently (Shimizu et al. 1997; van Achterbergh et al. 2001;
Grégoire et al. 2002; Simon et al. 2003), or significantly
modifying their chemistry (e.g. Burgess & Harte 2004). Hence
using the chemistry of these phases in concentrate form to
characterize the properties of the cratonic lithospheric mantle
and the processes that have affected it in entirety (e.g. Griffin et
al. 1999) may provide a highly biased view. Nevertheless, the
xenolith sample is in most cases the only sample of cratonic
lithospheric mantle available for a given craton. The samples can
still provide essential information regarding craton origin but, as
we will document below, it is important to try to discern which
chemical and mineralogical feature might result from the original
depletion event (or events) and which are products of subsequent
re-enrichment. Lastly, the compiled averages for compositional
data presented in this paper focus on those cratons where there
are enough samples to make meaningful statistical analyses. This
means that peridotites from some cratons (e.g. North China) are
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
Fig. 2. Age ranges of the oldest terrestrial minerals and rocks compared
with the average and oldest ages reported for Archaean lithospheric
mantle. Diamond ages are from Menneken et al. (2007), oldest mantle
melting event is from osmiridium data of Malitch & Merkle (2004) and
Pearson et al. (2007). Open symbols for cratonic lithospheric mantle and
oldest crust (Slave) show the mode in the statistical distributions, which
is significantly less than the oldest ages.
897
Fig. 3. Averages and standard deviations (error bars) for olivine Mg
numbers in low-temperature peridotites from various cratons and other,
more recent samples from other tectonic settings. Shaded ovals on the
North Atlantic craton (NAC) and Slave craton lines give the statistical
modes for the most depleted peridotites (i.e. harzburgites and dunites).
Shaded vertical oblong shows the apparent range for olivines from
depleted cratonic peridotite. Literature data compiled by D. G. Pearson
are available on request. Forearc data are for peridotites from Leg 125
(Ishii et al. 1992); OIB data are from Simon et al. (2008); abyssal
average is taken from Niu et al. (2004); Iwanaidake peridotite data are
presented separately for harzburgites and dunites (Kubo 2002). Diamond
inclusion data are from Stachel & Harris (2008).
not evaluated and discussed because of the as yet relatively small
but growing database for these areas (e.g. Gao et al. 2002;
Menzies et al. 2007).
Melt depletion and chemical buoyancy
Cratons have, by definition, remained convectively stable for
periods of time in excess of 2.5 Ga. This fact, together with the
lack of observable gravity excess over these areas of anomalously
low heat flow (Pollack & Chapman 1977), led Jordan (1975) to
construct the isopycnic hypothesis. The isopycnic hypothesis
states that the negative buoyancy imparted by the low temperature of the cratonic lithospheric mantle relative to convecting
mantle is offset by the positive buoyancy generated by depletion
of the cratonic lithospheric mantle in melt-forming elements.
Peridotite that has experienced extensive melt depletion has a
significantly lower intrinsic density than more fertile, undepleted
mantle (Boyd & McAllister 1976; Lee 2003). One of the
characteristic features of cratonic mantle is the dominant signature of melt depletion, evident in a number of parameters such as
their very high average Mg number of olivine (Fig. 3), elevated
bulk MgO contents and low bulk Al and Ca contents (Fig. 4).
The important elements from a density perspective are Fe and
Al, the latter because of its requirement to make the dense
mineral garnet. The significant depletion of these elements with
respect to fertile mantle (Fig. 5) is responsible for cratonic
lithospheric mantle densities being significantly lower than that
of the ambient convecting mantle. The standard temperature and
pressure (STP) density difference between fertile peridotite with
a bulk Mg number of 88 and typical cratonic lithospheric mantle
with a bulk Mg number of 92 is c. 1.5% (3390 kg m3 v.
3340 kg m3 ), rising to .2.5% for very depleted residues with
Mg numbers approaching 94.
According to the isopycnic hypothesis, cratons are effectively
in isostatic equilibrium as a result of the thermal and chemical
buoyancy factors cancelling each other out at any given depth
within the lithosphere. Although the evidence for this has been
frequently discussed since the inception of isopycnic theory, the
consensus is that conditions are reasonably close to those
required for neutral buoyancy (e.g. Lee 2003), with scatter and
uncertainty resulting from the likelihood that the xenoliths are
derived from magma conduit wall rocks and are not fully
representative of the density of the bulk of the cratonic lithospheric mantle.
It has been generally assumed that the longevity of cratonic
lithosphere is a direct result of its anomalous degree of melt
depletion compared with other peridotites of younger age (e.g.
Griffin et al. 1999). Whereas this is true if cratonic peridotites
are compared with abyssal or orogenic peridotites, Boyd (1989,
1996), Simon et al. (2008) and the extensive compilation in
Figure 4 show clearly that peridotites from modern-day forearcs,
ophiolites and even some ocean island basalts (OIBs) are more
depleted in CaO and Al2 O3 than most cratonic peridotites.
Hence, for these elements, there is no clear temporal evolution in
the degree of depletion of melting residues. If differences in
olivine Mg number are interpreted as resulting from differences
in the levels of melt depletion in peridotites (see below) then a
clear difference does exist between cratonic samples and most
peridotites from other tectonic settings (Boyd & Mertzman 1987;
Pearson et al. 2003; Bernstein et al. 2006a; Lee 2006; Fig. 3).
898
D. G . P E A R S O N & N. W I T T I G
Fig. 4. Average and standard deviation of bulk-rock Al2 O3 and CaO for
low-temperature cratonic peridotites compared with present-day
peridotites from other tectonic settings (forearcs, OIBs, abyssal), orogenic
peridotites, ophiolites and ‘off-craton’ peridotite xenoliths predominantly
from continental rifts. NAC, North Atlantic craton (East and West
Greenland); Churchill (broadly equivalent to Rae in Fig. 1), peridotites
from Somerset Island, Churchill Province, Northern Canada. Cratonic
datasets are from Boyd & Mertzman (1987), Boyd et al. (1997),
Bernstein et al. (1998), Lee & Rudnick (1999), Schmidberger & Francis
(1999), Kopylova & Russell (2000), Bizzarro & Stevenson (2003), Irvine
et al. (2003), Pearson et al. (2004), Wittig et al. (2008b), and Boyd
(unpubl. data). Orogenic peridotites are from a dataset compiled by
Bodinier & Goddard (2003). Off-craton data are from a compilation by
Canil (2004). Forearc data are from Ishii et al. (1992), abyssal data are
from Niu et al. (2004); OIB data are taken from the Canaries dataset of
Neumann et al. (1995) and are meant to be representative of the more
depleted OIB peridotite endmember rather than OIB peridotites as a
whole (for review, see Simon et al. 2008). Ophiolite data are from Jaques
& Chappell (1980) and Marchesi et al. (2006).
None the less, few studies other than that of Boyd (1996) have
recognized that the modern-day Earth is capable of producing
residues, in ophiolites or forearcs, that are as depleted in major
elements as any cratonic peridotite, or more depleted (Fig. 3).
These ultra-depleted residues will have densities as low as or
lower than cratonic peridotites and this may explain why their
fate is to be obducted not subducted; this implication has
interesting implications for models of craton root formation in
the Archaean Earth and also for forming new stable continental
keels today (Jackson et al. 2008).
The composition of archetypal cratonic mantle
lithosphere: what are the primary phases?
To extract genetic information from cratonic lithospheric mantle
peridotites a clear understanding of the effects of mantle
metasomatic processes on the bulk composition and mineralogy
of the rocks is necessary. Only recently have petrological and
geochemical studies been able to resolve most of the likely
metasomatic effects, in particular the impact on mineralogy. In
this regard, the addition of minerals such as clinopyroxene,
orthopyroxene and garnet needs to be addressed. Possible addi-
Fig. 5. Mg/Si v. Al/Si systematics for low-temperature cratonic
peridotites compared with present-day peridotites from other tectonic
settings (forearcs, OIBs (Canary Island only), abyssal), orogenic
peridotites, ophiolites and ‘off-craton’ peridotite xenoliths predominantly
from continental rifts. Data sources are as in Figure 4. Dark grey arrows
for orthopyroxene (Opx) clinopyroxene (Cpx) and garnet (Gt) addition
start from average North Atlantic craton (NAC) and are for typical low-T
Kaapvaal peridotite mineral compositions (Boyd & Mertzman 1987).
Light green large arrow labelled ‘high-T Opx’ is the vector for addition
of high-temperature aluminous orthopyroxene estimated by Saltzer et al.
(2001). Dashed black line is the ‘geochemical fractionation’ line of
Jagoutz et al. (1979). The displacement to low Mg/Si in Kaapvaal
peridotites and Phanerozoic ophiolites should be noted.
tion of any of these phases will potentially severely disturb both
major and moderately incompatible trace element systematics in
a suite of peridotites. This creates conflicting signatures as to the
depth and environment of melt extraction. Investigating these
issues is most reliably done by following the approach of Boyd
& Mertzman (1987), who emphasized the importance of using
large (preferably over 500 g) samples for the preparation of
analytical powders, to minimize scatter from sample inhomogeneities. Alternatively, where point counting is used it is
essential to follow procedures for obtaining statistically significant data to estimate reliable modal abundances (MacKenzie &
Canil 1999).
Early studies of modal mineralogy showed that a significant
proportion of Kaapvaal peridotites contained more orthopyroxene
than could be accounted for by partial melt extraction alone
(Boyd 1989; Canil 1992). This was confirmed by further experimental petrology (Walter 1998; Herzberg 1999). Kesson & Ringwood (1989) first advocated a secondary origin for the excess
orthopyroxene in Kaapvaal peridotites, via the addition of SiO2 rich fluids in an Archaean subduction zone. Recent major
element, trace element and isotopic evidence (Kelemen et al.
1998; Lee 2006; Simon et al. 2008) indicates a secondary origin
for orthopyroxene in Kaapvaal peridotites. The addition of
orthopyroxene to Kaapvaal peridotites imparts their well-known
Si-rich characteristic, which is also a feature, albeit to a lesser
extent, of some Siberian and Tanzanian peridotites (Fig. 5). The
addition of orthopyroxene to a more olivine-rich protolith is now
widely regarded as the origin for the Si-rich nature of the
Kaapvaal peridotites (Kelemen et al. 1992, 1998; Simon et al.
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
2008). A combination of the measured orthopyroxene, clinopyroxene and garnet compositions can explain the systematic shift of
cratonic peridotites away from the magmatic fractionation line in
Mg/Si–Al/Si space (Fig. 5). These are low-T phases whose
compositions reflect conditions considerably below the peridotite
solidus. An excellent fit to the Mg/Si–Al/Si cratonic trend is
produced by addition of pre-exsolution high-temperature aluminous pyroxene of the type suggested by Saltzer et al. (2001; Fig.
5). The addition of orthopyroxene to depleted harzburgite can
also explain the anomalously low FeO of Kaapvaal peridotites
(Walter 2003; Lee 2006; Simon et al. 2008), as addition of
substantial orthopyroxene dilutes bulk-rock FeO contents (Fig. 6).
Ophiolites that have experienced extensive serpentinization
have anomalously low Mg/Si, comparable with Kaapvaal peridotites (Fig. 5). This is caused by Mg loss during alteration. The
distinctively higher FeO contents of these highly altered oceanic
rocks compared with Kaapvaal peridotites (Fig. 7) make it
unlikely that the silica enrichment or low Ca contents of some
Kaapvaal peridotites (Schulze 1986) could originate by serpentinization of their protoliths.
Kaapvaal low-temperature peridotites are so orthopyroxene
rich that their average olivine contents fall far outside average
values for other cratons (Fig. 8). In contrast, cratonic peridotites
from the North Atlantic craton (Greenland) are poor in orthopyroxene, with modal compositions that are dominated by olivine,
yet with very similar olivine Mg numbers (Fig. 8; Bernstein et
al. 2006a). In particular, peridotite xenoliths from East Greenland (Bernstein et al. 1998) and Ubekendt Ejland, West Greenland (Bernstein et al. 2006b) are very olivine rich and closely
approximate both the olivine Mg number and olivine content of
highly depleted dunites from the Iwanaidaike peridotite massif,
Japan (Figs 8 and 9). When other samples from West Greenland
are considered (Bizzarro & Stevenson 2003; Wittig et al. 2008a)
the average olivine content is still very high (.89 wt%) and the
Fig. 6. Bulk-rock FeO v. MgO for cratonic peridotites. The five curves
forming a cluster are the experimentally derived batch melting trends of
Herzberg (2004) for melting between 1 and 7 GPa. Also plotted are the
fields for olivine compositions, Kaapvaal orthopyroxene, garnet,
clinopyroxene plus amphibole compositions, and high-temperature
experimental orthopyroxenes. Data sources are as for Figures 3 and 4.
Only data from low-temperature peridotites are plotted in all cratonic
examples.
899
Fig. 7. Bulk-rock FeO v. Mg/Si for Kaapvaal peridotites (data sources as
in Fig. 4) and serpentinized ophiolitic peridotites from Cuba (Marchesi et
al. 2006). The similar range in Mg/Si but elevated FeO of the ophiolitic
peridotites should be noted. Only data from low-temperature peridotites
are plotted for Kaapvaal samples.
Fig. 8. Average olivine Mg number v. weight per cent olivine for
peridotites from different cratons and compared with peridotites from
various present-day tectonic settings (forearcs, OIBs (Canary Islands
only), abyssal), Iwanaidake massif (ophiolites containing dunites and
harzburgites), Hokkaido, Japan (Kubo 2002). Field for orthopyroxenepoor ‘dunites’ from East Greenland (Bernstein et al. 1998) and Ubekendt
Ejland, West Greenland (Bernstein et al. 2006b) is marked as a dashed
box. Other West Greenland peridotites are more orthopyroxene rich
(Bizzarro et al. 2002; Wittig et al. 2008b) and influence the less olivinerich North Atlantic craton average. Other data sources are as for Figures
3 and 4. Only data from low-temperature peridotites are plotted in all
cratonic examples.
North Atlantic craton remains the most orthopyroxene-poor
endmember of the cratonic suites, plotting close to the ‘oceanic
trend’ of Boyd (1989; Fig. 8). The Tanzania (Lee et al. 1999)
and Slave (Kopylova & Russell 2000) cratons are progressively
900
D. G . P E A R S O N & N. W I T T I G
more olivine poor whereas the Kaapvaal craton is the extreme
orthopyroxene-rich endmember with only 64 wt% olivine on
average (Fig. 8).
Interestingly, the harzburgites of the Iwanaidaike ophiolite
massif are the closest average compositions to North Atlantic
craton peridotites in terms of modal olivine and olivine Mg
number, with the depleted dunites from Iwanaidaike being very
similar in composition to the Greenland dunites from the North
Atlantic craton (Kubo 2002; Figs 8 and 9). This similarity
between some cratonic peridotites and those of the Iwanaidaike
massif, a product of Mesozoic partial melting, in a suprasubduction zone environment (Kubo 2002), has significant
implications for the origins of cratonic peridotites in general and
this will be discussed in a later section. The close compositional
similarity of average cratonic peridotites (excluding Kaapvaal
samples) to recent melt residues suggests that these relatively
orthopyroxene-poor peridotites, highly depleted in Al and Ca,
with high olivine Mg numbers, are likely to be the more
archetypal components of cratonic lithospheric roots. In contrast,
the Kaapvaal peridotites represent an extreme endmember. None
the less, the scatter in olivine contents even in the North Atlantic
craton peridotites (Fig. 9) may be a product of some orthopyroxene addition.
Other processes that act to further obscure the chemical
signature of the original depletion event (or events) carried by
cratonic peridotites are the addition of modal clinopyroxene and
garnet. There is also now clear evidence for a secondary origin
for much of the clinopyroxene within cratonic peridotites.
Addition of clinopyroxene as a metasomatic phase to cratonic
peridotites has been suggested on the basis of the varied,
generally enriched trace element patterns of the pyroxenes (e.g.
Shimizu et al. 1997). Pearson et al. (2002) showed that the
abundance of clinopyroxene in Kaapvaal cratonic residues is
highly variable and significantly exceeds that expected for
residual rocks with such high Mg numbers, even when the effects
of sub-solidus exsolution are accounted for. This indicates that
most depleted cratonic peridotites contain excess clinopyroxene.
This observation is extended to all cratonic peridotites in Figure
10, and is in contrast to the very low (2% and less) clinopyroxene abundances evident in similarly depleted harzburgites and
dunites from the Iwanaidake peridotite massif. A clear indication
that the excess clinopyroxene may have been introduced relatively recently, and cannot have all exsolved from a higher
temperature orthopyroxene (e.g. Cox et al. 1987; Saltzer et al.
2001), comes from the lack of isotopic equilibrium between the
clinopyroxene and coexisting garnet (Simon et al. 2003) together
with Sr and Nd isotope compositions that are close to those of
present-day Bulk Earth in many instances (Pearson & Nowell
2002). Most ancient isotopic signatures in these cratonic peridotites reflect enrichment rather than depletion (e.g. Menzies &
Murthy 1980; Pearson & Nowell 2002).
Although Cr-pyrope garnet is a mineral thought by many to
characterize cratonic peridotites, its presence and composition
depend on both the depth of the lithosphere at the time of
xenolith sampling and the degree of depletion and subsequent
metasomatism. Smith et al. (2008) have recently described a
suite of very depleted chromite-bearing dunites and harzburgites
from the Murowe mine, Zimbabwe, that appear to originate from
depths within the diamond stability field. In these peridotites
garnet is a scarce phase, presumably because of the high-pressure
stability of high-Cr chromite and the lack of subsequent metasomatic addition that may have lowered the bulk-rock Cr number
to the point where garnet was again stable. Experimental
petrology (Ringwood 1977; Canil & Wei 1992; Stachel et al.
1998) has also shown that high-Cr, low-Ca garnets with high
knorringite component cannot be produced as residual phases on
the liquidus of cratonic peridotite and hence must be metamorphic products of the reaction between orthopyroxene and
chromite.
The origin of more typical Ca-saturated Cr-pyropes is less
Fig. 9. Olivine Mg number v. modal olivine for peridotites from the
North Atlantic craton (Greenland; see Fig. 8 for data sources) compared
with the harzburgite–dunite sequence from the Iwanaidake peridotite
(Kubo 2002). Squares show average values for abyssal (A; Niu et al.
2004) and Kaapvaal (K) low-temperature peridotites (from Fig. 8). Only
data from low-temperature peridotites are plotted in all cratonic
examples.
Fig. 10. Olivine Mg number v. modal clinopyroxene content for cratonic
peridotites (data sources as for Figs 3 and 4) compared with the
harzburgite–dunite sequence from the Iwanaidake peridotite (Kubo
2002). The similar range in olivine Mg numbers for the two suites but
the greatly contrasting variation in modal clinopyroxene should be noted.
Only data from low-temperature peridotites are plotted in all cases.
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
clear, although in many samples such garnets have a complex
trace element signature and very few have the highly radiogenic
Nd and Hf or very unradiogenic Sr isotope systematics that
would be consistent with an origin as ancient residual phases
(e.g. Pearson & Nowell 2002). Although it may be argued that
more recent exsolution of clinopyroxene and garnet from a highT aluminous orthopyroxene may reduce the likelihood of finding
ancient residual isotopic compositions in these phases, the nature
of Sm–Nd and Lu–Hf partitioning in orthopyroxene means that
any exsolved garnet and clinopyroxene would inherit a depleted
isotopic composition from their parent phase and following
exsolution would rapidly evolve to compositions significantly
more depleted than the Bulk Earth mean values observed
(Pearson & Nowell 2002). In some peridotites textural and
chemical evidence have combined to invoke a metasomatic
origin for garnet (Bell et al. 2005).
The last phase to add to the list of minerals that, in some
circumstances, may not be of primary origin in cratonic peridotites is olivine itself. Olivines of low Mg number (86–88) have
been identified in deformed dunitic peridotites from the Kimberley area of the Kaapvaal craton (Boyd et al. 1983) and also in
some Greenland peridotites (Wittig et al. 2008a). These rocks
appear to have experienced significant olivine addition and are
hence strongly modified from their original solidus compositions.
Whereas the Kimberley dunites are rare examples of this
phenomenon, olivine addition may also account for the more
olivine-rich nature of Kaapvaal high-T peridotites (Kelemen et
al. 1992). Overall it is clear that, despite their simple mineralogy,
cratonic peridotites are rocks of extreme complexity, in which
every major mineral phase, to greater or lesser extents, may have
been modified or introduced by melt–fluid interaction. This
multi-stage history obscures much of their early history and it is
thus no surprise that their origin has been so difficult to
constrain.
901
from metasomatic activity for considerable time periods (Bernstein et al. 2006a).
Modern high-quality phase equilibria studies allow the change
in mineral chemistry of residual phases to be correlated to the
extent of melting experienced (Fig. 11; e.g. Baker & Stolper
1994; Walter 1998). This relationship was exploited by Bernstein
et al. (2006a) to estimate the amount of melt extraction
experienced by typical cratonic peridotites to be c. 40%, based
on their very consistent, high olivine Mg numbers, which average
92.8 from their database. Our own compilation of the literature
data (Fig. 3) reveals some differences from the averages recorded
by Bernstein et al. (2006a), either because of increased available
data (in the case of the North Atlantic craton and the Kaapvaal
craton) or perhaps because we have chosen to include a wider
variety of rock types, such as for the Slave craton, where the
statistical distribution of olivine Mg numbers is clearly bimodal
(Fig. 12). For Slave cratonic peridotites the lower olivine Mg
numbers belong to lherzolites whereas olivines with the highest
Mg numbers belong to the ‘ultra-depleted layer’ thought to
consist predominantly of harzburgite. This contrasts with the
mixed lherzolite–harzburgite suite from the Kaapvaal craton in
which many lherzolites have olivine Mg numbers that are just as
high as those of typical harzburgites. None the less, the low Mg
number tail in the distribution of Kaapvaal olivines is dominated
by lherzolites (Fig. 12).
Despite some interesting inter-craton variations in the statistical distribution of olivine Mg numbers, we concur with Bernstein
et al. (2006a) that although modal mineralogy varies widely,
there is remarkable consistency in the average olivine Mg
numbers for cratonic peridotites (Figs 3 and 12), the modes of
which vary from 92.5 to 92.8 for every craton. Even for cratons
Extent of melting
As outlined above, cratonic peridotites contain extensive evidence for the addition of orthopyroxene, clinopyroxene and most
probably some garnet. Because modal mineralogy varies strongly
depending on equilibration depth, it is not possible to use modal
abundance data alone to assess the likely extent of melting.
Although bulk chemistry approaches such as the significant
depletion of CaO and Al2 O3 clearly indicate extensive melt
extraction, the abundances of these elements are also sensitive to
the addition of clinopyroxene and garnet in particular. However,
at any melting depth, mantle peridotites with very high olivine
contents (.90%) probably represent residues from very extensive
melt extraction (40% or more), as long as such rocks contain
highly magnesian olivines. In fact, the chemistry of olivine may
be the most robust indicator of the extent of melting in a
peridotite. Despite the evidence of later olivine introduction into
some cratonic peridotites noted above, in the majority of cratonic
peridotites it seems likely that olivine compositions might be
relatively faithful recorders of the melting history of the rock.
The Mg number of olivine in particular is relatively insensitive
to the introduction of high Mg number phases such as orthopyroxene (e.g. Kelemen et al. 1992, 1998). Another indication of
the robustness of olivine compositions to later compositional
changes is the observation that the average Mg number of
olivines included in diamonds (Fig. 3) is very similar (92.8) to
that of the averages for peridotites from numerous other cratons
(92.5) even though the olivine in diamond has been armoured
Fig. 11. Mg number of olivine v. per cent melting (F) in residual
peridotites from experiments at varying pressures and starting
compositions. Error bars give estimated experimental uncertainties on the
melt fraction. Source of experiments: Baker & Stolper 1994; Walter
1998. Arrows mark the changes in melting reactions denoted by
clinopyroxene-out and orthopyroxene-out. It should be noted that this
plot is equivalent to that of figure 3 of Bernstein et al. (2006a) except for
the exclusion of some experimental data and the inclusion of
uncertainties on F.
902
D. G . P E A R S O N & N. W I T T I G
Fig. 13. Summary of possible effects of metasomatism on olivine Mg
number v. modal olivine relations for cratonic peridotites (s). Arrows for
different stages relate to: (1) melt extraction, probably accompanied by
melt infiltration (Niu et al. 2004); (2) melt extraction where
orthopyroxene melts incongruently leaving an olivine-rich residue;
(3) orthopyroxene addition at constant olivine Mg number (horizontal
trend; major arrow), or orthopyroxene clinopyroxene and garnet
addition at varying olivine Mg number (diagonal trends; minor arrows);
(4, 5) melt addition creating higher or lower olivine Mg numbers;
(6) melt addition creating increased olivine Mg number plus increased
modal orthopyroxene clinopyroxene and garnet.
Fig. 12. Probability density plots of olivine Mg numbers for peridotites
from the Kaapvaal craton, the Slave craton and the North Atlantic craton
(West and East Greenland). Only data from low-temperature peridotites
are plotted in all cases. A constant uncertainty of 0.2 is used for all
olivine Mg numbers. Number of analyses plotted is given as n in curve
labels. Shaded oblongs at the top of the diagram show the averages of
OIBs (Canary Islands only), abyssal peridotites and the Papuan ophiolite
belt. Data for the Iwanaidake ophiolite includes dunites and harzburgites.
Curve labelled ‘all cratons’ is a compilation of all the cratons listed in
Figure 3. Forearc data are from Leg 125 peridotites (Ishii et al. 1992).
Other data sources are as in Figure 3.
where only small databases are available for cratonic peridotite
xenoliths, such as the North China craton (Gao et al. 2002), the
average Mg number of olivine is 92.6 (Bernstein et al. 2006a).
Both the very olivine-rich East Greenland peridotites and the
olivine-poor Kaapvaal peridotites have very similar olivine Mg
numbers and so orthopyroxene addition does not appear to have
been the most significant factor dictating olivine chemistry (Figs
8 and 9). The orthopyroxene addition vector is therefore
subhorizontal on plots of olivine Mg number v. modal olivine
(Fig. 13) whereas other processes (e.g. olivine, clinopyroxene or
melt addition) may increase or decrease olivine Mg numbers and
hence create scatter. Bernstein et al. (2006a) suggested that an
olivine Mg number of 92.6 marks the orthopyroxene ‘out’ stage
in the melting of peridotite (Stage 2 of Fig. 13). Although the
constancy of the olivine compositional modes for cratonic
peridotites may support this view, there are a small number of
samples with both higher and lower olivine Mg numbers. These
variations may be caused by melt–rock interaction, with the
effect on olivine and modal compositions depending upon the
melt–rock fraction (Bodinier & Godard 2003; Fig. 13). Overall,
the main trend of constant olivine Mg number at widely varying
modal olivine composition is consistent with the idea that
cratonic peridotites melted to the stage of orthopyroxene ‘out’
(Stage 2; Fig. 13) and then experienced later, varying degrees of
orthopyroxene addition (Bernstein et al. 2006a). This multi-stage
model implies that most cratonic peridotite residues (at least
those with olivine Mg numbers close to 92.6) were dunites at
some stage. This model is supported by the statistical distribution
of olivine compositions in the Iwanaidake peridotite massif.
Olivines from this mixed dunite–harzburgite massif have a
bimodal distribution (Fig. 12) in which the main mode at Mg
number 92.5 belongs to dunites (Fig. 3) and is identical to
cratonic peridotites.
Bernstein et al. (2006a) suggested that the clear relationship
between olivine Mg number and extent of melting observed in
experiments indicates that cratonic peridotites are the residues of
40% partial melt extraction. Our plotting of the experimental
database suggests that there is some uncertainty in this estimate,
depending on the depth and hence pressure at which the melting
took place (Fig. 11). Some difficulty arises because of the
different starting compositions used by experimentalists. Using
experiments performed at different pressures and with the same
starting composition (e.g. Walter 1998) it is evident that higher
pressure melt extraction more rapidly depletes Fe in residues
leading to higher olivine Mg numbers at a given degree of
melting (compare regressions for 3 and 7 GPa, Fig. 11). This
means that for an average olivine Mg number of 92.6 the degree
of melt extraction is estimated to be between 37 and 47%,
depending on the pressure of melt extraction. Despite this
uncertainty, the indicated extent of melt extraction is clearly
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
exceptional and rarely achieved in the modern-day Earth except
in forearc environments where previously depleted residues are
remelted by water fluxing and decompression (Parman & Grove
2004; Fig. 3; see ‘Environment of melting’). Even for the
significant tail of Kaapvaal peridotites with olivine Mg numbers
of around 92, melting extents are between 27 and 36% (i.e. still
very high) if these compositions are primary and not lowered by
metasomatism.
The power of using olivine compositions to assess levels of
melt depletion in peridotites is apparent when compared with
bulk analytical data. Average CaO and Al2 O3 contents of
cratonic peridotites are certainly low, being significantly less than
those for most abyssal peridotites. However, some xenoliths
erupted by OIBs such as at Gran Canaria (Neumann et al. 1995)
also have very low CaO and Al2 O3 contents that are more
depleted than in most cratons (Boyd 1996; Simon et al. 2008;
Fig. 4). In contrast, the olivine compositions of these OIB
peridotites are significantly less magnesian (91.4 0.6) than
those of average cratonic peridotites (Fig. 12), supporting the
view that the bulk Ca and Al contents of cratonic peridotites
have been enhanced by later metasomatism. We therefore believe
that olivine compositions, used with caution, offer the most
reliable estimate of the degree of melt depletion in peridotites.
Having established that the statistical distribution of olivine
compositions for cratonic peridotites is generally tightly unimodal (with the exception of the Slave craton), it is important to
establish the melting environment and the depth of melting that
may lead to such a distribution, because this will have an
important bearing on the tectonic environment in which cratonic
roots are formed. Both theoretical and experimental approaches
have been used to quantify the melting behaviour of peridotite
(e.g. McKenzie & Bickle 1988; Iwamori et al. 1995; Herzberg
2004). The results from the two approaches are broadly in
agreement. Polybaric melting beginning at high pressure (e.g.
between 5 and 7 GPa) at a high potential temperature of 1650 8C
produces a large fraction of melting over a relatively small
Fig. 14. Fraction of melting v. depth or pressure for the polybaric
melting of peridotite. Shaded regions bounded by dashed blue (fractional)
and continuous green (batch) lines are for theoretical melting paths
calculated by Iwamori et al. (1995). Continuous black lines are melting
trajectories plotted from the experimental data summarized by Herzberg
(2004). Potential temperatures of melting trajectories are indicated (in
8C). Shaded vertical oblong denotes the region of the garnet–spinel
transition.
903
pressure drop (Fig. 14). In contrast, melting at lower potential
temperatures, which initiates melting at lower pressures (4 GPa
or less), is more linear in nature. Hence, of the two scenarios
(deep, high-pressure melting or shallow, low-pressure melting),
high-pressure melting may be more likely to yield residues where
the majority have experienced large fractions of melt depletion
Fig. 15. Whole-rock MgO (wt%) v. Al2 O3 (wt%) for cratonic peridotites.
Continuous black lines are the melting trajectories for various pressures
(2, 3, 5, 7, 10 GPa) taken from Herzberg (2004). Melt extraction equiline
of 30% is shown (larger extents of melting are to the right of the line).
Field for ‘opx’ depicts the range in orthopyroxene compositions observed
in low-T Kaapvaal peridotites (Boyd & Mertzman 1987; Simon et al.
2008). Field for ‘high T opx’ is for liquidus unexsolved orthopyroxenes
from data sources summarized by Saltzer et al. (2001) and Simon et al.
(2008). Whole-rock analytical data sources are as for Figure 4.
Fig. 16. Whole-rock Lu (ppm) v. Yb (ppm) elemental systematics in
cratonic peridotite suites. Curves are shown for polybaric fractional
melting beginning at 2, 3 and 7 GPa and isobaric melting at 7 GPa
following melting models described by Wittig et al. (2008a). Highpressure (7 GPa) isobaric or polybaric melting would produce residues
lying to the high Lu and Yb side of the Primitive Upper Mantle (PRIMA)
reference point. No cratonic peridotites are displaced in this direction.
Data sources are indicated in the diagram. Data sources are: Wittig et al.
(2008a), Simon et al. (2007), Schmidberger & Francis (1999) and Friend
et al. (2002).
904
D. G . P E A R S O N & N. W I T T I G
and therefore have similar olivine Mg numbers. Only melting at
potential temperatures in excess of 1600 8C in either low- or
high-pressure regimes can achieve melt fractions in excess of
40%, the likely average experienced by cratonic peridotites (Fig.
11). Even in these circumstances, if melting is occurring in a
ridge or plume environment, only a small fraction of the
residues, located in the axial part of the melting column, will
experience these very high levels of melt depletion. To further
constrain which melting scenario and hence tectonic environment
can best explain the statistical distribution of olivine compositions we will now try to better understand the likely depth of
melting experienced by cratonic peridotites.
of orthopyroxene, clinopyroxene and/or garnet (Fig. 15). Some of
the spread of the East Greenland data towards the high-pressure
melting trends in Al2 O3 –MgO space could be the result of minor
olivine addition that is a likely result of melt–rock interaction
(e.g. Kelemen et al. 1998). In contrast, a shift towards lowerpressure melting trends may be caused by extensive serpentinization of peridotites, inducing Mg loss and a relative increase in
the concentration of Al (e.g. Niu 2004). From these considerations it is clear that major element approaches to constraining the
pressure of melt depletion experienced by cratonic peridotites are
not likely to be accurate because major elements are relatively
insensitive to melting depth but very sensitive to modal metasomatism.
Depth of melting
Trace element considerations
Major element considerations
Although it has long been established that the cratonic lithospheric mantle is highly depleted in melt-forming elements
relative to fertile mantle (Nixon & Boyd 1973; Maaloe & Aoki
1977; Boyd 1989), the depth at which this melting took place
has been a far more contentious issue. This question is of
paramount importance in establishing an accurate picture of
craton formation and there have been a number of attempts to
constrain the issue over the last 20 years. The initial approaches
were largely based around bulk-rock major elements (Schulze
1986; Boyd 1989; Walter 1998, 1999, 2003) whereas more
recently the usefulness of moderately incompatible trace elements has been realized (Kelemen et al. 1998; Canil 2004; Lee
2006; Wittig et al. 2008b).
Even when plotting only bulk compositions that have been
reliably estimated from large samples, the effects of metasomatism on major element bulk compositions cause some significant
confusion in extracting depth of melting information. For
instance, when cratonic peridotites are compared with the major
element compositions of melt residues predicted from experiments, the data scatter across melting trends for both high- and
low-pressure melting (Figs 6 and 15). A case has been made for
using FeO as a melting depth proxy in some non-cratonic
peridotites that have experienced minimal metasomatism (Ionov
& Hofmann 2007) but for cratonic rocks this is not practical.
The low FeO contents of Kaapvaal peridotites in particular have
been cited as evidence of high-pressure melt extraction because
of the sensitivity of FeO to melting pressure (Pearson et al.
1995a; Walter 1998; Aulbach et al. 2007). Once a secondary
origin for much of the orthopyroxene is accepted, a convincing
case can be made for the dilution of FeO in these rocks by
orthopyroxene addition (Walter 2003; Lee 2006; Simon et al.
2008; Figs 6 and 15), and this removes one of the main
arguments for high-pressure melting. The compositions of orthopyroxene-poor Greenland peridotites, argued above to be more
representative of archetypal cratonic lithospheric mantle compositions (Bernstein et al. 2006), are displaced clearly towards the
low-pressure melting trend in terms of FeO–MgO relations (Fig.
6), in agreement with the generally low pressures of melting
indicated for other major elements such as Al2 O3 v. MgO (Fig.
15). However, there is still considerable scatter in the major
element bulk compositions of this and other cratonic peridotite
suites.
This scatter in cratonic peridotite datasets is partly due to latestage secondary Fe enrichment (e.g. Boyd et al. 1997) and may
also be influenced by the addition of other phases such as
clinopyroxene. Considerable scatter across the Al2 O3 –MgO
diagram is largely because of the effects of post-melting addition
The partitioning of mildly incompatible trace elements such as V,
Sc and Yb is sensitive to the presence or absence of garnet in
peridotite melting residues and hence can be used as a powerful
tool to constrain the depth of melting. Using this approach, Canil
(2004) has provided convincing evidence that most peridotite
residues reported in the literature, both cratonic and postArchaean, originate via low-pressure melting (generally
,3 GPa). This conclusion has also been made on the basis of Yb
partitioning in cratonic peridotites (Kelemen et al. 1998; Lee
2006; Simon et al. 2008). Heavy REE such as Yb are perhaps
more preferable than Sc and V for this use because they are less
susceptible to the possible introduction of clinoproxene, or
interaction with the host kimberlite, which is generally Yb-poor
(Wittig et al. 2008b). The approach has been extended using
coupled Yb–Lu elemental systematics that show clearly, for
cratonic and off-craton peridotites, that they must have experienced most of their melting under garnet-absent conditions, most
probably at pressure less than 3 GPa (Fig. 16; Wittig et al.
2008a). Hence, there is a clear consensus from workers using
mildly incompatible elements that most of the residues of
melting present within the lithospheric mantle as a whole, and
particularly cratonic mantle, form by melting at low pressures,
mostly less than 3 GPa.
Whereas estimates of melting pressure constrained by mildly
incompatible elements such as Yb and Lu are relatively insensitive to modal or host rock metasomatism, because high-pressure
(e.g. 7 GPa) melt residues evolve in the opposite direction to
low-pressure (e.g. 3 GPa) melt residues (Fig. 16) the extents of
melt extraction predicted using these elements are affected. None
the less, many cratonic peridotites retain evidence of large
extents of melt extraction, in excess of 25% or more (Wittig et
al. 2008a). In summary, the strongest evidence for the depth of
melting experienced by cratonic peridotites comes from a range
of mildly incompatible elements, with the consensus being that
melt extraction took place at pressures of 3 GPa or less (Stachel
et al. 1998; Canil 2004; Lee 2006; Simon et al. 2008; Wittig et
al. 2007, 2008a). This firm constraint can now be used to
evaluate the likely tectonic environment in which cratonic
peridotites were generated.
Environment of melting
Any successful genetic model for cratonic peridotites must
successfully explain their high and yet relatively uniform levels
of melt depletion, produced in a shallow melting environment.
Consideration of the melting relations of peridotite (Fig. 14)
indicates that single-stage melting at low pressures (3 GPa and
less) is unlikely to generate abundant residues that have experi-
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
enced c. 40% or more melt extraction, even at the high potential
temperatures that may have existed in the Archaean mantle. This
apparent contradiction can be resolved by examination of
residues produced in more recent, better understood melting
environments.
Evaluation of both bulk-rock chemistry (Fig. 4) and olivine
compositions (Figs 3 and 12) shows that the only modern-day
melting environments capable of producing melt residues of
comparable depletion to cratonic peridotites are in arc settings,
notably the Izu–Bonin forearc (Ishii et al. 1992), the Iwanaidake
peridotite massif (Kubo 2002) and ophiolites such as the highly
depleted Papuan and Cuban ophiolite massifs (Jaques & Chappell 1980; Marchesi et al. 2006). This is also confirmed from
examination of mildly incompatible element abundances (Wittig
et al. 2008a). These residues are all produced at low pressures,
with most of the melting probably occurring in the spinel
stability field, at ,3 GPa, in a mantle environment where
hydrous fluids are present.
The similarity between residues produced in some modern-day
subduction zones and cratonic peridotites has been previously
noted by Parman et al. (2004), who proposed that the production
of Barberton komatiites in an Archaean subduction zone led to
the generation of residual Kaapvaal peridotites. A multi-stage
melting model, with the final stage of melt extraction occurring
in a subduction environment, has also been suggested by Simon
et al. (2008) and Wittig et al. (2008a) to explain the modal, trace
element and isotopic characteristics of Kaapvaal and Greenland
(North Atlantic craton), respectively. This ‘wet melting’ model is
also lent support in terms of explaining the statistical distribution
of olivine compositions in cratonic peridotites that have pronounced modes at very depleted compositions. The low-pressure
(,3 GPa) melt extraction indicated by trace element systematics
means that initial melting is likely to have taken place at a
spreading centre, either at a mid-ocean ridge or in an arc-related
setting or in a shallow forearc environment. Single-stage polybaric melting at a ridge is unlikely to produce abundant residues
dominated by highly depleted compositions, even at Archaean
mantle potential temperatures (Fig. 14). The involvement of
hydrous fluids, lowering the solidus, means that melting can
extend to greater melt fractions at a given temperature (e.g.
Hirschmann et al. 1999; Parman & Grove 2004), leading to
greater levels of melt extraction and greater volumes of highly
residual peridotite. In fact, theoretical modelling of the effects of
melt depletion in peridotites using MELTS (Hirschmann et al.
1999) indicates that after clinopyroxene exhaustion, isobaric melt
productivity decreases dramatically, by a factor of four, before
rising again close to the orthopyroxene-out reaction. This drop in
melt productivity results from a decrease in the temperature
derivative of the bulk solid composition occurring as the melting
reaction changes (Hirschmann et al. 1999). Hence, in anhydrous
melting environments the clinopyroxene-out reaction is a natural
impediment to the production of residues of very large degrees
of melt extraction. The addition of water acts to lower the
temperature required to reach a given extent of melt depletion.
For instance, addition of 0.2% H2 O under isothermal conditions
at 1350 8C and 1 GPa produces an increase of 15% melting
compared with anhydrous conditions and the effect is greater at
higher temperatures (Hirschmann et al. 1999). Given that the
effect of melting past clinopyroxene exhaustion under hydrous
conditions produces the same drop in melt productivity, it seems
that the likely higher potential temperature of the Archaean
mantle will assist, along with water, in consistently producing
residues from very large extents (c. 40%) of melting.
As melt productivity rises again near the orthopyroxene-out
905
boundary (under hydrous or anhydrous conditions) and then
declines rapidly again thereafter, the orthopyroxene-out reaction
would appear to produce a natural barrier, even to melting under
hydrous conditions, and this may explain why few olivine Mg
numbers exceed the value indicated by experiments as corresponding to this reaction (Fig. 11). Hydrous melting is thus
capable of producing the strongly peaked statistical distributions
in olivine compositions seen in cratonic peridotites. This is
confirmed by examination of olivine compositions from hydrous
melting environments such as the Izu–Bonin forearc and Iwanaidake massif peridotites and many ophiolites (Figs 3 and 12). A
model involving two stages of melt extraction, with the final
stage occurring in an arc, will also encourage melting in garnetfree conditions. The stability of spinel in mantle peridotite is
enhanced at depleted, high Cr/Al bulk compositions (e.g. Ringwood 1977). Hydrous melting of depleted mantle in a subduction
zone environment means that melting is much more likely to be
taking place with spinel on the solidus at greater depths than for
normal fertile mantle. This may explain the apparent complete
absence of a strong garnet signature in any of the cratonic or
modern subduction zone peridotites (Fig. 16).
The combined weight of evidence from bulk compositions,
olivine compositions and trace element systematics points clearly
towards a shallow final stage of melting in a subduction zone
environment. Whether melting occurred in two stages, first at a
ridge and then in a subduction zone (e.g. Simon et al. 2008), or
whether it took place predominantly in a subduction zone, with
the simultaneous production of komatiites (Parman et al. 2004)
cannot easily be resolved. However, hydrous melting in presentday subduction zones often involves mantle that has experienced
previous melt depletion events (Parman et al. 2004), some of
which took place billions of years previously (Parkinson et al.
1998; Pearson et al. 2007). As noted above, melting of previously depleted peridotite also favours melting taking place in
the spinel stability field. The consequence of invoking a melting
model that involves the last major period of melt extraction
taking place at relatively shallow depths (,100 km) in a subduction zone environment is that this then requires tectonic stacking
to thicken to lithospheric root to depths of 150 km or more (e.g.
Helmstaedt & Schulze 1989). A result of subduction accretion of
residues beneath existing crustal blocks is that this process
decouples residue compositions from overlying melt compositions and also decouples the residue age from that of the
overlying crust. Both these consequences are the opposite of the
relations expected from the plume model of craton formation.
Timing of melt depletion and diamond formation in
cratonic peridotites
Lithosphere formation events
Early Re–Os isotope studies of cratonic peridotites found that
they retained signatures of Archaean melt depletion and hence
appeared to be broadly as old as the cratonic crust lying above
them (Walker et al. 1989; Carlson & Irving 1994; Pearson et al.
1995a, b). Since then, Archaean melt depletion ages have been
found in every suite of cratonic peridotites analysed and a much
larger database has been generated (for reviews, see Pearson et
al. 2003; Carlson et al. 2005). One of the newest localities from
which cratonic peridotites have become available is the North
Atlantic craton, where Archaean melt depletion ages are evident
in xenolith suites from both East and West Greenland (Hanghøj
et al. 2001; Bernstein et al. 2006b; Wittig et al. 2007; Fig. 17).
The early Re–Os studies of cratonic peridotites suggested that
906
D. G . P E A R S O N & N. W I T T I G
the main mass of cratonic lithospheric mantle was Mesoarchaean
in age, similar to much of the early crust building phase (e.g.
Pearson et al. 1995a). However, the large database now available,
especially for the Kaapvaal craton (Carlson et al. 2005), allows a
more detailed examination of the age information that can be
retrieved from these rocks and reveals that there is no direct age
relationship between crust and lithospheric mantle. In addition to
whole-rock Re–Os analyses of peridotites, another useful
development has been the use of sulphide Re–Os dating (Pearson
et al. 1998; Pearson, N. J., et al. 2002) and the two methods offer
much complementary information (Fig. 17). Although wholerock Re–Os analyses of peridotites are certainly integrations of
numerous potential events that may have affected the evolution
of a cratonic peridotite, and sulphide analyses have been
suggested to be superior in terms of retrieving age information
(Griffin et al. 2004), it is important to recognize the problems
inherent in both approaches. There is abundant evidence for the
recent introduction of significant Re (Walker et al. 1989; Pearson
et al. 1995a; Carlson et al. 2005) and some Os (Chesley et al.
1999) into whole-rock peridotites and this is why model ages that
are calculated assuming a Re content of zero (the timing of Re
depletion: TRD age) are often preferred (Walker et al. 1989).
These factors mean that the TRD age for whole-rock peridotites
is likely to be an underestimate of the actual melting age.
The sulphide dating approach also has significant pitfalls,
unless the sulphide is trapped in a diamond at the time of its
formation. The low solidi of sulphides means that they are one
of the first phases to enter the melt during peridotite melting.
This means that at the high melt fractions extracted from cratonic
peridotites, common sulphides belonging to the monosulphide
solid solution (mss) series will no longer be residual phases
(Pearson et al. 2004; Lorand & Grégoire 2006). Hence, any such
sulphides, or Ni-rich pentlandites, are metasomatic in nature
(Luguet et al. 2003; Lorand & Grégoire 2006) and so by
definition will carry a mixed signal themselves. Furthermore,
peridotites with easily located high-T sulphides are relatively
scarce among cratonic peridotites, especially when weathered.
For instance, an intensive search for high-T base metal sulphides
in a suite of peridotites from the Argyle diamond mine by
Luguet et al. (2006) was unsuccessful and whole-rock analysis
was the only possible approach. The low melting temperature of
Cu–Ni–Fe sulphides means that they are easily disturbed and
this produces a very large spread in resulting model ages. For
instance, in a study of 168 sulphides from the Kaapvaal craton,
Griffin et al. (2004) found that 33 sulphides gave ages older than
the age of the Earth and 28 sulphides yielded future ages, using
the Re–Os TMA model ages (model age using Re/Os ratio of
sample) approach, with ages ranging from 79 to +130 Ga. Part
Fig. 17. Probability density plots of Re–Os model ages for whole-rock
peridotites and sulphides for the Kaapvaal, Slave and North Atlantic
cratons. Whole-rock data sources are as follows. Kaapvaal craton:
Pearson et al. (1995a), Carlson et al. (1999), Menzies et al. (1999),
Irvine et al. (2001), Carlson & Moore (2004) and Simon et al. (2008);
Slave craton: Irvine et al. (2003), Aulbach et al. (2004) and Westerlund
et al. (2006); North Atlantic craton: Hanghøj et al. (2001), Bernstein et
al. (2006), Webb (2008) and Wittig et al. (2008b). All TRD ages are
corrected to the eruption age of the kimberlite. Model age equations for
TRD and TMA have been presented by Walker et al. (1989) and Pearson
(1999) and are calculated to reference values for chondritic mantle of
187
Os=188 Os ¼ 0:1283 and 187 Re=188 Os ¼ 0:422. Number of samples
plotted for each population is given by n.
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
of this spread may be due to open-system behaviour and part is
due to significant problems in the application of Re corrections
to small Os beams during laser ablation analyses (Nowell et al.
2008). This large fraction of aberrant ages led Griffin et al.
(2004) to screen sulphides such that only ages from sulphides
with Re/Os ,0.08 were deemed reliable. Although this narrows
the spread in sulphide ages, assessing their accuracy is difficult.
There are two factors that have caused confusion about the
relative proportions of ancient lithospheric mantle ages provided
by sulphides. One is the fact that, for the Kaapvaal craton,
despite the large number of sulphides analysed, these data come
from only 24 rocks (Griffin et al. 2004) compared with the 228
samples analysed for whole rocks. The second factor is the
misleading practice by some workers (e.g. Griffin et al. 2004) of
plotting on probability density diagrams age data where single
ages have very large differences in uncertainty. A large variation
in the apparent uncertainty of a Re–Os model age is common in
sulphide datasets but is due largely to the increased precision of
high-Os, low-Re/Os sulphides. Such data points create a false
impression of the relative proportion of lithosphere of a given
age if they are plotted alongside data of considerably worse
precision (e.g. Griffin et al. 2004). In view of the fact that the
dominant uncertainty on any Re–Os model age is our lack of
precise knowledge of the Os isotope evolution of the Earth’s
mantle, it is preferable to plot all such ages with a uniform
conservative uncertainty (Fig. 17). A final factor that has
distorted views on the relative merits of sulphide v. whole-rock
Re–Os ages is the comparison of sulphide ages and whole-rock
ages in rocks with abundant sulphides. Peridotites that have
obviously experienced such metasomatism would be avoided by
workers applying whole-rock analytical approaches.
Another issue relating to the interpretation of Re–Os ages in
cratonic peridotites is the growing evidence for considerable Os
isotopic variation in the convecting mantle (see summary by
Pearson et al. 2007). Re depletion ages of .1 Ga are now
routinely found in samples recently extracted from the convecting mantle such as abyssal peridotites (e.g. Harvey et al. 2006)
and Pt-group alloy grains (e.g. Pearson et al. 2007), and some
samples show evidence of melting events that occurred . 2Ga
ago. This creates severe difficulties in estimating the significance
of small numbers of samples that have Re depletion model ages
of 1–2 Ga occurring with other samples of younger age, within a
particular lithospheric peridotite suite (e.g. Hanssler & Shimizu
1998). In such sample suites, where all the Os isotopic diversity
is within that shown by abyssal peridotites, or where only one or
two samples are less radiogenic than the ‘convecting mantle’
range, it is difficult to have confidence that any specific lithospheric formation event is being dated without independent
constraints. However, isotopic diversity in the convecting mantle
will naturally be less in the early stages of mantle evolution.
Recent samples of convecting mantle with evidence of .2.5 Ga
melting ages are extremely rare (1% of all data: Pearson et al.
2007). These two points together with the predominance of
Archaean ages from all cratonic peridotite suites (Fig. 17) make
it very likely that the abundant Archaean Re–Os model ages for
these samples reflect melt extraction during lithospheric formation.
In summary, it is likely that whole-rock TRD ages provide
minimum estimates for the age of melting in cratonic peridotites
whereas sulphide and whole-rock TMA ages are likely to
overestimate depletion ages. Despite the fact that it is difficult to
interpret either whole-rock or sulphide model ages with high
levels of certainty, some striking observations can be made for
cratons where both types of data have been obtained in some
907
abundance (i.e. the Kaapvaal and Slave cratons). The statistical
distributions of whole-rock TMA and sulphide TMA model ages
for the Kaapvaal craton show remarkable similarities (Fig. 17),
with prominent but rather broad modes at 2.9–3.0 Ga. The
distribution of whole-rock TMA ages shows a greater proportion
of ages in the Eoarchaean to Hadean range but this probably
reflects Re addition or problems with Re interference corrections
(see Nowell et al. 2008) causing overestimation of TMA ages. In
contrast, Kaapvaal peridotite TRD ages show a tighter, more
pronounced mode at 2.7 Ga, with relatively few ages that are
Mesoarchaean or older (Fig. 18). The clustering of whole-rock
TRD model ages at around 2.7 Ga is unlikely to result from
variable metasomatic addition of sulphides as suggested by some
workers (e.g. Griffin et al. 2004), as such processes would lead
to a more diffuse spread in ages.
Although there is clearly a significant spread in ages, the fairly
good agreement between the positions of the prominent modes
for all the model age approaches could be taken to indicate that
the major melt extraction event for the majority of Kaapvaal
peridotites occurred between 2.7 and 3.0 Ga. This timing is
younger than the age for the Kaapvaal lithospheric mantle
commonly inferred from the 3.3 Ga diamond inclusion Sm/Nd
model ages obtained by Richardson et al. (1984). It is coincident
with collisional suturing of the Witwatersrand and Kimberley
blocks of the Kaapvaal craton at c. 2.9 Ga, probably along the
present-day Colesberg magnetic lineament (Schmitz et al. 2004).
The Neoarchaean has been proposed by Pearson et al. (2002) to
be the time of the major formation and stabilization of the
lithospheric mantle beneath the Kaapvaal craton, and the recently
defined palaeo-subduction zone between the two main blocks of
the Kaapvaal craton provides a mechanism to make Kaapvaal
peridotites in an arc environment at this time (Schmitz et al.
2004; Simon et al. 2008). This association is strengthened by the
petrological arguments documented above.
Further support for the timing of major subduction at 2.9 Ga is
provided by the 2.89 0.06 Ga Re–Os isochron for eclogitic
sulphides included in diamonds (Richardson et al. 2001) and
2.7–3.0 Ga Re–Os model ages for whole-rock eclogites from
this region (Shirey et al. 2001). Evidence for the involvement of
older lithospheric material, perhaps transposed from the older
Witwatersrand block, comes from pre 3 Ga Re–Os model ages
for eclogites and peridotites from the Newlands kimberlite
(Menzies et al. 1999, 2003). This may explain the 3.3 Ga Sm/Nd
model ages for peridotitic diamonds from the Kimberley block.
The Witwatersrand block of the Kaapvaal craton contains
crustal material formed between 3.7 and 3.3 Ga that seems to
have formed a coherent block by 3.2 Ga (Schmitz et al. 2004).
The older mini-peak in the mode for sulphide TMA ages at
3.1 Ga (Fig. 17) has been interpreted to represent older lithospheric mantle generation in the Witwatersrand block (Griffin et
al. 2004). Although whole-rock TMA ages are consistent with
this view of a significant fraction of Mesoarchaean lithosphere
(Fig. 17), the TRD whole-rock ages are much more tightly
clustered in the Neoarchaean. Further work is required to
investigate this issue, although it does not seem unreasonable
that some substantial fraction of lithospheric mantle from the
Witwatersrand block formed in the Mesoarchaean, as has been
suggested for lithosphere formation beneath the Siberia craton
(Pearson et al. 1995b; Pearson, N. J., et al. 2002) and the eastern
part of the North Atlantic craton (Hanghøj et al. 2001).
The mode in age distributions for both whole-rock TRD and
sulphide TMA ages are tight and in good agreement at c. 2.7–
2.8 Ga for Slave peridotites (Fig. 18). Davis et al. (2003) and
Heaman & Pearson (2008) have noted that this peak in peridotite
908
D. G . P E A R S O N & N. W I T T I G
Re depletion ages corresponds to the timing of final stabilization
of the Slave craton, as marked by intense 2.7 Ga granite
plutonism. The suggestion of the presence of older diamondiferous lithosphere beneath the Slave craton at 3.5 Ga from
diamond inclusion dating (Westerlund et al. 2006) and the
observation of intense granitic plutonism at 2.7 Ga, implying
very high thermal gradients at the base of the crust, are difficult
to reconcile (Davis et al. 2003). To explain these observations,
Percival & Pysklywe (2007) have suggested that the entire
cratonic lithospheric mantle becomes inverted during eclogite
delamination at c. 2.7 Ga. It is also possible that the 3.5 Ga
sulphide isochron of Westerlund et al. (2006) is a mixing age
produced by resetting of older lithospheric material during
diamond crystallization at 2.7 Ga, as indicated by the ubiquitous
2.7 Ga TMA model ages of these sulphides. These problems and
observations have led Davis et al. (2003) and Heaman & Pearson
(2008) to favour a Neoarchaean formation and stabilization age
for much of the Slave cratonic lithosphere.
A Neoarchaean stabilization age is also indicated by wholerock TRD ages from the North Atlantic craton (Fig. 17; Wittig
et al. 2008b). It seems a striking coincidence that peridotites
from the three cratons with the most abundant Re–Os isotope
data all have their modes in depletion ages between 2.7 and
3.0 Ga, suggesting that the Neoarchaean may have been a time
when the majority of the now extant cratonic lithospheric
mantle was produced and stabilized, mostly probably in Archaean subduction zones. There are certainly more ancient Mesoand Eoarchaean model ages evident for peridotites from the
above three cratons and from Siberia (Pearson et al. 1995b).
These may represent earlier lithospheric mantle formation
episodes; for instance, for parts of the Witwatersrand block of
the Kaapvaal craton or the northern and eastern portions of the
North Atlantic craton exposed on Greenland. Some of these
more ancient ages also may represent earlier melting events
either at ridges, or in other environments that did not result in
the accretion of these residues to the lithosphere. There are
analogues to this situation in the modern Earth, where increasing evidence is being found for ancient depleted components
existing in convecting mantle (Parkinson et al. 1998; Harvey et
al. 2006; Pearson et al. 2007).
Lithosphere modification events and diamond formation
As the great majority of diamonds are formed in the lithospheric
mantle (Stachel & Harris 2008) the formation and subsequent
evolution of the lithosphere are clearly critical factors in the
formation of diamonds. One of the best defined diamond
formation ages for the Kaapvaal craton is the 2.89 0.06 Ga
eclogitic sulphide isochron for inclusions within Kimberley
diamonds (Richardson et al. 2001). This event is most obviously
correlated with the time of proposed suturing between the
Witwatersrand and Kimberley blocks of the Kaapvaal craton
(Schmitz et al. 2004) and the Archaean diamond crystallization
event is synchronous with the likely formation of much of the
lithospheric mantle in the Kimberley block, in a subduction zone
environment. There are a number of additional post-Archaean
peaks in the statistical distribution of sulphide and whole-rock
Re–Os models ages for cratonic peridotites (Fig. 17). The most
prominent is the peak between 0.9 and 1.2 Ga observed in
sulphide TMA and whole-rock TRD ages for the Kaapvaal
craton (Fig. 17). This peak correlates with diamond-forming
events in the Kaapvaal lithosphere at 0.99 Ga (Orapa; Richardson
et al. 1990) and 1.1 Ga (Pearson et al. 1998) and corresponds to
the timing of the Namaqua high-temperature thermal event
(Tankard et al. 1982). As formation of the eclogitic hosts to these
Proterozoic diamonds took place in the Archaean (Menzies et al.
1999; Shirey et al. 2001) they are likely to be of metasomatic
origin, not associated with any discrete subduction event. Their
association with high-temperature metamorphism in the deep
crust indicates rising thermal gradients as a result of either
lithospheric rifting or plume impingement. The thermal imprint
in the lower crust of these Proterozoic thermal events is
preserved by zircon ages in 0.964–1.11 Ga granulite-facies
xenoliths (Schmitz & Bowring 2000). The broad coincidence of
all these ages, allowing for the larger inaccuracies of the mantle
xenolith Re–Os model ages, is strongly suggestive of a prolonged and significant lithospheric-scale thermal event that seems
to have involved the crystallization of diamonds in some parts of
the cratonic root. Given the relatively lowl temperatures of
diamond formation indicated by trapped inclusions (1150–
1240 8C, Stachel & Harris 2008), this growth is likely to take
place on the up-temperature side of such a thermal event, with
carbon-bearing fluids being driven off ahead of a rising plume.
The c. 2 Ga diamond-forming event at Premier (Richardson et
al. 1993; Fig. 18) correlates with the major lithospheric thermal
disturbance manifest as the Bushveld intrusion and its thermal
imprint. There are a number of Kaapvaal peridotites of this age
(Fig. 17) and new data acquired by Morel et al. (2006) indicate
significant new peridotitic lithosphere added at 2 Ga beneath the
central Witwatersrand block. This new data provide further
support for suggestions by Richardson et al. (1993) and Shirey et
al. (2002) for a link between pulses of diamond growth and
major lithospheric thermo-tectonic events. There are exceptions
to this general rule in that there seems to be no indication of a
mantle or crustal event to accompany the c. 1.5 Ga ages found
for diamonds from Finsch and Jwaneng (Richardson et al. 1990,
2004). None the less, the correlation between major lithospheric
thermo-tectonic events manifest in both crust and mantle and the
formation of diamonds is strongly indicative of a link and
indicates a pulsed nature of diamond formation associated with
such events. It seems more than coincidental that significant
secondary age peaks in the lithosphere can be identified in the
lithospheric mantle of the North Atlantic craton at similar ages
to the diamond–lithosphere ages seen in the Kaapvaal craton
(Fig. 17) and we will await diamond dating to see how these ages
may relate to the age of diamonds from this region. In a broader
context, a global correlation between mantle melting events and
crustal ages has recently been recognized through Os isotope
studies of platinum group metal alloy grains (Pearson et al.
2007), with c. 1.1 Ga and 2.2 Ga pulses of enhanced activity
evident. The correlations identified here predict that, ultimately,
there may be very strong temporal links between crust formation,
continental lithospheric mantle generation, the depletion of the
convecting mantle in general and diamond formation that are
intrinsically related to the episodic nature of mantle convection
and melting.
A model for the formation of cratonic roots
In the discussion above we have shown that the bulk compositions and olivine chemistry of cratonic peridotites are similar to
those of Phanerozoic to recent peridotites that have been melted
in shallow environments, in the presence of hydrous fluids (i.e. in
arc settings). The generation of cratonic peridotites in subduction
zone environments has been gaining support from different
angles for the past two decades (Ringwood 1977; Kesson &
Ringwood 1989; Rudnick et al. 1994; Kelemen et al. 1998; Canil
2004; Parman et al. 2004; Bell et al. 2005; Lee 2006; Simon et
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
al. 2008) and this seems to be the best overall scenario to explain
the shallow depth of melting, the very high and uniform levels of
depletion, and the variable later introduction of orthopyroxene.
Here we present a general petrogenetic (Fig. 19) and tectonic
model (Fig. 20) for the formation of cratonic peridotites that
integrates and builds on these earlier studies.
We will use a plot of Mg number of olivine v. modal olivine
through the sequence of events because it is useful to illustrate
the different processes involved and we make analogy to the
well-studied residual rocks of the Iwanaidake peridotite massif to
graphically illustrate the possible reactions between melt and
residue (Fig. 19). We start from the premise that most refractory
cratonic peridotites require two stages of melting to generate
their highly depleted, relatively uniform compositions. The first
stage of melting may have occurred at an Archaean mid-ocean
ridge where melting took place at 3 GPa or below to produce a
fertile harzburgite; Stage 1 of Figure 19. This stage of melting
produced the curved trend of increasing olivine Mg number v.
increasing modal olivine, referred to as ‘the Oceanic Trend’ by
Boyd (1989). If this melting took place in the geometrically
complex environment of an ocean ridge then it is likely that
some olivine addition also contributed to the trend defined by
Stage 1 as a result of the crystallization of olivine-saturated melt
at higher levels in the lithosphere (e.g. Niu 1994).
Stage 1 may have been the immediate precursor to the
enhanced melting represented by Stage 2 if all the melt extraction occurred during a single stage, at the wet solidus, in a
subduction zone environment. Stage 2 is extended, water-fluxed
melting to on average 40% (Bernstein et al. 2006a; Fig. 11) or
more melt removal, close to, or past, the point of orthopyroxene
exhaustion to produce orthopyroxene-poor dunitic residues. This
is the composition of archetypal cratonic peridotite suggested by
Kesson & Ringwood (1989), Bernstein et al. (2006) and Wittig
et al. (2008a). Although it is possible to produce such refractory
compositions without the aid of water because of the high
potential temperatures of the Archaean mantle (Fig. 11), such
melting is unlikely to produce the uniformly highly depleted
residues observed (Figs 9 and 12) if it is constrained to pressures
,3 GPa, and hence hydrous melting is most appropriate. The
best examples of these highly refractory Stage 2 residues are the
olivine-rich endmembers of cratonic peridotites from the East
and West Greenland parts of the North Atlantic craton (Bernstein
et al. 1998, 2006; Wittig et al. 2008a), which contain no garnet
but very high-Cr spinel. In some instances these chromite dunites
may transform at the base of cool cratonic lithosphere into highCr garnet dunites (Ringwood 1977; Schulze 1986; Kesson &
Ringwood 1989; Stachel et al. 1998).
Stage 3 is the addition of orthopyroxene to refractory dunitic
residues (Fig. 19). Kelemen et al. (1992), Bell et al. (2005) and
Simon et al. (2008) proposed that orthopyroxene formed from
high-T Si-rich fluids in the subduction zone environment,
whereas Rudnick et al. (1994) favoured siliceous melts from
basaltic slab melting. However, a more local source of Si-rich
melts is available that is capable of generating orthopyroxenerich peridotites. The melting reaction of enstatite is peritectic in
nature, forming olivine and a Si-rich melt. The extended melting
that creates refractory dunitic residues will form a melt capable
of reacting with olivine in parts of the nearby already depleted
mantle that has risen (because of its lower density) and cooled
slightly, to produce more Mg-rich orthopyroxene that will be
close to equilibrium with its host and hence unlikely to alter the
Mg number of olivine in the resulting rock. The close to solidus
temperatures of this reaction environment will be conducive to
the formation of high-temperature aluminous orthopyroxene,
909
which seems the best fit for explaining the bulk compositional
trends of many cratonic peridotites (Figs 5 and 6). A further
attraction of this model, where the orthopyroxene-forming melt
is derived relatively locally, is that melt/rock ratios could be
high, allowing extensive percolation and deposition of significant
volumes of orthopyroxene. Derivation of orthopyroxene solely
from much more distal sources, as exogenous water-rich fluids
(Bell et al. 2005; Simon et al. 2008) or silica-rich melts
(Rudnick et al. 1994; Kelemen et al. 1998) has the difficulty of
needing to extensively pervade large columns of lithospheric
mantle.
The mechanism of orthopyroxene enrichment we propose is
essentially ‘autometasomatic’ in that the Si-rich melt is locally
derived and close to equilibrium with its cooler host. This
mechanism can create the subhorizontal trend on the olivine Mg
number v. modal olivine plot that is characteristic of cratonic
peridotite suites (Boyd 1989; Menzies 1990; Boyd 1996; Figs 2,
8, 9 and 13). A single cycle of these melting and metasomatic
events will produce a mixture of dunite and orthopyroxene-rich
refractory harzburgites. Multiple cycles of this process, where
cooler early formed residues are reacting with siliceous fluids
being generated in new melting environments below, will result
in a lithospheric section dominated by refractory harzburgites
with widely varying orthopyroxene contents at relatively constant
olivine Mg number (Figs 13 and 19). Melting of newly arrived
but previously depleted mantle, as buoyant residues ascend to
form proto-lithosphere, maintains melting in the spinel facies.
Addition of olivine to fertile harzburgite that has not been fluxed
by water, or olivine formation in refractory harzburgite that has
not progressed to the orthopyroxene ‘out’ stage, will create
additional spread away from the Stage 3 vector (Fig. 19) but all
these processes are envisioned to take place in the Archaean
melting environment. Subsequent addition of minor garnet and
clinopyroxene, predominantly associated with magmatic activity
in the Proterozoic to Phanerozoic, may create more diverse
spread in modal olivine and olivine compositions (Fig. 19).
The tectonic manifestation of the above model is depicted in
Figure 20. We envisage that Stages 1–4 of ridge melting and
subduction followed by physical maturity occurred in the
Neoarchaean, between 2.9 and 2.7 Ga in the case of the
Kimberley block of the Kaapvaal craton, and possibly the Slave
and the SW part of the North Atlantic craton. It can be seen that
one obvious consequence of a model that invokes tectonic
stacking of oceanic lithosphere is the likely incorporation of
oceanic crust of abundant eclogitic slabs in the newly stabilized
‘continental’ lithosphere. Even the crudest estimate of the
abundance of this eclogitic component is of the order of 10% by
volume. Estimates of the abundance of eclogite now residing in
the cratonic lithosphere are considerably less than this, at around
1% by volume (Schulze 1989). This difference may be because
large dense slabs of eclogite are not gravitationally stable in lowdensity cratonic peridotite. Percival & Pysklywe (2007) showed
via dynamical modelling that a large horizontal slab of eclogite
located at the base of the cratonic crust could sink rapidly
through the underlying peridotite lithosphere causing significant
thermal disturbance. Although this model leads to some appealing results in terms of transient thermal spikes at the base of the
crust to produce Neoarchaean granites above older but thermally
disturbed peridotite, it is a greatly simplified scenario. A
subduction stacking model for the lithosphere is likely to have
multiple slabs of eclogite inclined at some angle from the
horizontal. This situation is also not one of long-term gravitational stability and it is likely that the lithosphere will tend to
‘cleanse itself’ of much of the dense eclogite, leaving behind a
910
D. G . P E A R S O N & N. W I T T I G
Fig. 18. Age correlations between crust,
lithospheric mantle and diamond formation
events for the Kaapvaal craton. Duration or
age uncertainty of events is indicated by
width of box. For diamond ages, line
denotes age and box width denotes
uncertainty. Colour coding for crust and
mantle events is to delineate those evident
in the Witwatersrand block (orange), the
Kimberley block (green) and the marginal
cratonic Namaqua metamorphic event
(pink). Diamond ages (grey) show ages for
silicate and sulphide inclusion dating.
Crustal geochronology is from Tankard et
al. (1982), Schmitz & Bowring (2000) and
Schmitz et al. (2004). Mantle ages are from
Figure 17 and references listed in the
caption. Diamond ages are from Richardson
et al. (1984, 1990, 1993, 2004), Pearson et
al. (1998) and Pearson & Harris (2006).
Fig. 19. Petrogenetic model for the formation of cratonic peridotites via multi-stage melting and metasomatism. The model depicts the formation of ultradepleted dunitic residues starting at an oceanic ridge (Stage 1) via hydrous melting in a subduction zone environment (Stage 2). The resulting olivine-rich
refractory lithologies from Stage 2 are probably the starting compositions for most cratonic lithosphere (Kesson & Ringwood 1989; Bernstein et al.
2006a; Figs 8 and 9), which then evolve to varying degrees of orthopyroxene enrichment as a result of metasomatism by silica-rich melts (Stage 3) driven
off from later forming dunites during hydrous melting. Melting is within the spinel facies, aided by the melting of previously depleted mantle from the
ridge. Multiple cycles of such water-fluxed melting then create ultra-depleted harzburgites with very variable orthopyroxene (and hence olivine) contents
at relatively constant and high olivine Mg numbers. Later, infiltration of small-degree Si-poor melts causes the generation of lherzolites via clinopyroxene
and some garnet crystallization. The legend gives the identity of key minerals in the upper illustrations during the evolution from fertile harzburgite to
dunite to refractory orthopyroxene-rich harzburgite to refractory lherzolite. Time sequence is from left to right. Graphs show how complexity evolves with
each stage in terms of olivine Mg number v. modal olivine; the arrows with numbers are explained in Figure 13.
A R C H A E A N C O N T I N E N TA L L I T H O S P H E R E A N D D I A M O N D S
911
Fig. 20. Tectonic schematic illustration showing the generation and evolution of the early stages of the formation of a craton focusing on the processes
taking place in the lithospheric mantle root. Initial melting to create relatively fertile harzburgite is shown taking place at an Archaean oceanic ridge in
Stage 1. The location of the two-stage melting and metasomatic model illustrated in Figure 19 is shown in the mantle wedge above an Archaean
subduction zone (Stage 2), and shows how newly created ultra-depleted lithospheric mantle might grow and cool. Stage 3 indicates a proto-cratonic root
thickened by tectonic stacking that contains a significant volume of eclogitic crust (black). This eclogitic crust is dense and hence gravitationally unstable,
sinking through the peridotite to leave a more mature craton with only minor, localized eclogite lenses (Stage 4). The transition from Stage 3 to Stage 4
probably involves significant physical and thermal instability in the mantle and lower crustal sections of the evolving craton. For the Kaapvaal craton we
envisage that Stages 1–4 took place in the Neoarchaean between 2.9 and 2.7 Ga for the Kimberley block.
more thermally and physically mature craton with a root of lowdensity peridotite dominated lithospheric mantle with minor
eclogitic pods nested within it (Stage 4 of Fig. 20). It is likely
that during the loss of the eclogite component there is significant
physical and thermal instability in the mantle and lower crustal
sections of the evolving craton. Following Percival & Pysklywe
(2007) we suggest that this activity may create one or more
thermal spikes sufficient to cause crustal melting and Neoarchaean crustal magmatism. The physical and thermal instability of
a craton during this period may be able to explain the apparently
very active early history of the Kaapvaal craton in which a major
rifting event that produced the Ventersdorp lavas was occurring
at 2.7 Ga, shortly after the supposed stabilization of the craton
(de Wit et al. 1992) following east–west suturing (Schmitz et al.
2004).
Ostensibly, the model presented in Figures 19 and 20 suffers
from a similar problem to most other models of cratonic
peridotites; that is, lack of evidence of suitable volumes of large
fraction mafic melts associated with the ultra-depleted residues
that form the lithosphere. However, although this is a seemingly
critical problem in any plume-driven hypothesis for making
cratonic mantle because there is no evidence of anomalous
crustal thicknesses beneath well-studied cratons such as the
Kaapvaal (James et al. 2001; Nguuri et al. 2001), the subduction
stacking model has the consequence of physically transporting
and decoupling residues from melts, especially if any resulting
cogenetic melts that do become incorporated into the lithospheric mantle become later ‘refined out’ by gravitational
processes (Fig. 20). This model also allows for significant
decoupling of the ages of lithospheric mantle and overlying
crust.
We thank T. Stachel, J.-L. Bodinier, Y. Niu, E.-R. Neumann, N. Simon
and D. Francis for supplying additional bulk-rock and mineral chemistry
data used in this review. S. Parman provided a critical sounding board for
the ideas presented here, and discussions with C. Garrido helped to
sharpen some of the arguments.
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Received 10 January 2008; revised typescript accepted 19 March 2008.
Scientific editing by Rob Strachan