Geodynamic controls on diamond deposits: Implications for Australian occurrences

Tectonophysics 404 (2005) 217 – 236 www.elsevier.com/locate/tecto

Geodynamic controls on diamond deposits: Implications for Australian occurrences C.J. O’Neilla,*,1, L. Moresib, A.L. Jaquesc a The University of Sydney, School of Geosciences, Building F05, NSW, 2006, Australia School of Mathematical Sciences, Monash University, Building 28, VIC, 3800, Australia c Geoscience Australia, GPO Box 378, Canberra, ACT 2601, Australia

b

Received 10 November 2004; received in revised form 5 April 2005; accepted 21 April 2005 Available online 23 June 2005

Abstract Conventional diamond exploration guidelines predict that economic diamond occurrences will be restricted to Archaean cratons, where the lithosphere is thick and cool, and diamond is the stable form of carbon in the lower portions of the lithosphere. However, Australia’s current economic diamond deposits are not well predicted by these conventional exploration guidelines. Tomographic images show that Australia’s economic diamond deposits lie at step changes in lithospheric thickness within dominantly cratonized Proterozoic provinces with thick (z 200 km) lithosphere. The thickest portions of the seismic lithosphere in Australia occur not under the major Archaean cratons, rather the central Proterozoic regions of the continent. We use a numerical code to show that such features are stable, and that the longevity of the diamond stability field is dependent on distance to the continent–ocean boundary, local depth of the chemical boundary layer (CBL), and proximity to changes in CBL depth. We also show that abrupt changes in lithospheric thickness focus lithospheric stress gradients, affecting melt migration paths, and that continental melt production is enhanced in regions adjacent to major cratons. Diamond pipes occur where conditions conducive to diamond stability and deep-seated alkaline volcanism (kimberlite or lamproite) occur simultaneously, and the common confluence of these factors at abrupt changes in lithospheric thickness marks them as potential exploration targets. D 2005 Elsevier B.V. All rights reserved. Keywords: Diamond; Geodynamic; Mantle lithosphere; Australia

1. Introduction * Corresponding author. Tel.: +1 713 348 4567; fax: +1 713 348 5214. E-mail address: [email protected] (C.J. O’Neill). 1 Present address: Department of Earth Sciences, Rice University, 6100 Main St Houston, MS 126, Houston, TX, 77005, USA. 0040-1951/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2005.04.010

Southern Africa has long been one of the world’s most important producers of diamonds, and many existing diamond exploration strategies are the result of years of distilled experience searching for diamonds in such terrains. One of the best known explo-

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ration guidelines, known as Clifford’s rule (Clifford, 1966; Janse, 1994) loosely states that diamondiferous kimberlites are restricted to the ancient cratons, whereas pipes occurring off-craton in adjacent mobile belt terrains will be barren (Clifford, 1966; Janse and Sheahan, 1995). The reason for this is that in cratonic terrains the lithosphere is old, thick, and comparatively cold (Helmstaedt and Gurney, 1995; Morgan, 1995). This is thought to be due to the stabilization at depth of these cratonic brootsQ or bkeelsQ by virtue of their high viscosity or chemical buoyancy (Jordan, 1978, 1981, 1988; Griffin et al., 2003). In such regions the continental geotherms will pass the graphite–carbon phase transition, into the diamond stability field, within the depth of the lithosphere, meaning that diamonds will be the stable form of carbon in the deepest parts of the lithosphere. In contrast, the adjacent mobile belts are much younger, and the lithosphere in such regions is typically significantly thinner than under cratons. As a result, the continental geotherms will not pass into the diamond stability within the depth of the lithosphere in these areas. This interpretation has found support in xenolith barometry (e.g. Finnerty and Boyd, 1987; Priestley and McKenzie, 2002) and seismic tomography (James and Fouch, 2002; Shirey et al., 2002) which both indicate a much thicker lithosphere in cratonic regions than other areas in southern Africa. The lithosphere in the surrounding regions is both thinner, hotter and younger (Griffin et al., 2003). Rhenium–Osmium isotopic dating of peridotite xenoliths has yielded Re model depletion (T RD) ages that indicate that significant amounts of the thick lithosphere under the Kaapvaal craton were formed in Neoarchaean time (3.0–2.5 Ga) in contrast to that of surrounding areas which have Proterozoic T RD ages (Pearson et al., 2002). This simple model, which generally describes the distribution of economic diamond deposits, was extended by Janse (1994) who recognized Archons (N2500 Ma), Protons (Proterozoic cratons, 2500– 1600 Ma) and Tectons (1600–800 Ma) and noted that economically viable kimberlites were restricted to Archons while a few economically viable lamproites also occur on Protons. However, a number of important mines, both in South Africa and elsewhere, lie near the reworked margins of the major Archaean cratons or cratonic nucleii (e.g. Gurney et al., 1991) and recent dating of inclusions in diamonds

from a number of these pipes indicates the presence of diamondiferous lithosphere of Proterozoic as well as Archaean age in the diamond source regions (Richardson et al., 2004; Shirey et al., 2002, 2004). More importantly, the location of Australia’s three economic deposits to date is not well predicted by conventional guidelines (Jaques, 2002; Jaques and Milligan, 2004). Australia is also unusual in that the thickest portions of the seismic lithosphere do not occur under the major Archaean cratons, but rather the central Australian Proterozoic mobile belts (Kennett, 2003; Simons et al., 2002, 1999). This suggests that, in contrast to South Africa, Australia’s Proterozoic regions may be potentially prospective for diamond. The purpose of this paper is to assess the geodynamic controls on diamond deposits, and compare model predictions to the observed distribution of Australian diamondiferous pipes. We employ a fully dynamic numerical code (Moresi et al., 2002) to model the evolution of continents in a convecting mantle. We track the variability shown in continental geotherms, and the time they spend in the diamond stability field (DSF), throughout the time of each simulation. We also explore the distribution of sub-continental melt production, and factors affecting the early-stage migration of alkaline melts. The results from these numerical experiments are presented in a regional exploration framework. 1.1. Australian diamond occurrences Diamondiferous (mostly of very low grade and not economic) intrusions of kimberlite, lamproite and related rocks occur within or at the margins of most of the Precambrian provinces of Australia (Pidgeon et al., 1989; Atkinson et al., 1990; Jaques, 2002; Jaques and Milligan, 2004) but no economic diamond deposits have been identified within the major Archaean cratons to date. Australia’s very significant diamond production (first in world by weight) comes from the Argyle deposit in the Kimberley region of Western Australia (Boxer and Jaques, 1990). There is also limited production from the much smaller mine at Ellendale, in the West Kimberley region (Fig. 1). Both these deposits differ from the more conventional cratonic kimberlite pipes: they are hosted in lamproite rather than kimberlite and lie in Palaeoproterozoic mobile belts at the margins of a craton. The 1180

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Fig. 1. Image map of heat flow for the Australian continent (from Cull, 1982). Also shown are the major geological features discussed in the text, including the Pilbara, Yilgarn, Kimberley and Gawler cratons, Australia’s three economic diamond occurrences (Argyle, Merlin and Ellendale fields, shown as diamonds), other primary diamond occurrences discussed in the text (large circles), and related non-diamondiferous sub-alkaline volcanism (light circles). Data are from Jaques (2002).

Ma Argyle AK1 lamproite pipe lies with the Palaeoproterozoic Halls Creek Orogen at the southeastern margin of the Kimberley Craton which is thought to have, at least in part, a basement of late Archaean age although none is exposed (Graham et al., 1999). Inclusions in the diamonds themselves have been dated at around 1.580 Ga (Richardson, 1986), indicating a residence time in the mantle lithosphere of ~ 400 Myr. The younger (ca. 815 Ma) Bow Hill lamprophyres also lie in the Hall’s Creek Orogen (Fielding and Jaques, 1989). They are, however, barren, indicating either a different source and/or melt migration path, or that the lithosphere had experienced some thinning or a thermal event (Fielding and Jaques, 1989). The 20 Ma Ellendale lamproite pipes are part of a large West Kimberley lamproite province that extends south from Palaeoproterozoic King Leopold Orogen at the southwestern margin of the Kimberley Craton in the Palaeozoic Canning Basin (Jaques et al., 1986). Australia’s other economic diamond deposit–the ca. 350 Ma Merlin kimberlite field

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which was mined on a trial basis between 1998 and 2003–lies within the dominantly Proterozoic North Australian craton (Lee et al., 1997, 1998). All three deposits appear to lie at the margins of or on major gradients between lithospheric domains defined by seismic tomography (Jaques and Milligan, 2004). Other lower grade primary diamond occurrences (see Fig. 1 and Jaques and Milligan, 2004) include the ca. 800 Ma North Kimberley kimberlite field (Jaques et al., 1986; Wyatt et al., 1999), the 1900–1700 Ma Nabberu kimberlites (Shee et al., 1999) and the 1300 Ma Jewill kimberlites (Graham et al., 1999) at the northern margin of the Yilgarn Craton, the 1900 Ma Brockman kimberlite dyke in the Pilbara Craton (Wyatt et al., 2004), the Jurassic kimberlites at Orroroo in the Adelaide Fold Belt (Scott-Smith et al., 1984), Timber Creek in the Victoria River Basin in the Northern Territory (Berryman et al., 1999; Belousova et al., 2001) and, possibly, the 164 Ma Wandagee pipes of the Carnarvon Basin (Jaques et al., 1989). 1.2. Geodynamic controls on diamond deposits Diamonds can be formed in a number of highpressure processes, and have been documented in paleo-subduction zones (Davies et al., 1999), highpressure metamorphic shear zones (Sobolev and Shatsky, 1990), and impact sites (Lipschutz, 1964). However, most diamonds of economic interest form in the deep continental lithosphere or, in some cases, the asthenosphere (e.g. Stachel, 2001), and are sampled and brought to the surface by alkaline volcanism (Helmstaedt and Gurney, 1995). Helmstaedt and Gurney (1995) list five essential ingredients for an economic diamond deposit; we consider a modified version of this. These necessary ingredients are: diamond formation, preservation in the continental roots, production of alkaline melt at sub-lithospheric depths, sampling of the diamondiferous roots during the ascent and eruption of the melt, and preservation of the surface expression of the primary deposit. Isotopic dating of inclusions in diamond has yielded both ancient Archaean (3.2–3.3 Ga) and much younger (Proterozoic and Phanerozoic) ages with the older inclusions being typically of peridotitic composition and the younger ages typically of eclogitic paragenesis (e.g. Richardson et al., 1984, 1993,

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2004; Pearson et al., 1999). The Archaean ages of inclusions from several populations of South African diamonds, erupted in Cretaceous kimberlites, have led to the suggestion that the cratonic root beneath South Africa, and the thermal conditions within it, have been stable for several billion years (Boyd et al., 1985). Direct evidence of the long-term stability of cratonic roots has been provided by Re–Os systematics of peridotite xenoliths from kimberlites intruding the craton (Pearson et al., 2002). Griffin et al. (2003) used the composition of garnet xenocrysts to map the compositional structure and thermal state of the sub-continental lithospheric mantle beneath the Kalahari Craton and showed that the oldest components of the lithosphere are the most refractory (chemically depleted) and that the younger surrounding mobile belts are underlain by more fertile lithosphere which they suggested might reflect the reworking of Archaean lithosphere. The range in ages of inclusions in diamond and their apparent correlation with crustal events and craton structure, as determined by seismic tomography, has led to the proposition that diamond formation is episodic and intimately linked to craton development and evolution (Richardson et al., 2004; Shirey et al., 2002, 2004). The stability of continental regions is a complicated issue, and depends strongly on the tectonic history of the continent, and its structure (Lenardic et al., 2000). Jordan (1978, 1981, 1988) and more recently O’Reilly et al. (2001) argued that the buoyancy of the chemically-depleted continental roots imparts thermal and mechanical stability resulting in their long-term survival of Archaean sub-continental lithosphere. Studies of the composition of mantle peridotite xenoliths (e.g. Boyd, 1989, 1997) and more recently analysis of mantle pyrope garnet xenocrysts (Griffin et al., 1998, 2003) confirm that Archaean sub-continental lithospheric mantle is compositionally distinct from younger mantle, being more refractory (chemically depleted), of lower density and significantly buoyant compared with the underlying asthenosphere, enhancing its likelihood of surviving tectonic processes. Griffin et al. (2003) proposed that Archaean sub-continental lithospheric mantle formed as residues and/or cumulates from large-scale melting at 150–250 km depth in response to major mantle overturns or mega-plumes and that this very buoyant, refractory lithosphere

(termed dlife raftsT) formed the core or nuclei of the ancient cratons. Later events result in reworking of or accretion to the craton. However, although the buoyancy of Archaean cratonic roots makes them resistant to destruction they are not indestructible: they may be transformed and/or destroyed by mechanical disaggregation (by lithospheric thinning and/or rifting) and infiltration of upwelling fertile mantle (O’Reilly et al., 2001). Griffin et al. (1999, 2003, 2004) have presented evidence for progressive re-fertilisation of formerly refractory Archaean sub-continental lithospheric mantle, especially during episodes of intraplate magmatism, that results in a loss of the unique chemical and density differences compared with younger mantle, potentially enabling its destruction by recycling as evidenced by the recycling of the Sino-Korean craton by Phanerozoic subduction (Griffin et al., 1998). Clearly, the mechanical stability of the roots does not necessarily imply thermal stability. O’Neill and Moresi (2003) have shown that, while the conditions for diamond stability can exist in continents under Archaean mantle conditions, thermal fluctuations (Richter, 1985) within stable roots are of sufficient magnitude to potentially destroy the diamonds they contain. Many examples exist where the mantle root has apparently been destroyed, or perturbed beyond the diamond stability field, by a tectono-magmatic event (O’Reilly et al., 2001; Gaul et al., 2003; Griffin et al., 2004). In some cases the later events/processes have apparently destroyed the diamondiferous source regions but in other cases they have not. For example, the economic diamondiferous mid-Paleozoic kimberlites of the Archaean Magan and Anabar provinces of the Siberian Craton are underlain by thick (190–240 km) cold, refractory lithosphere whereas lithosphere beneath the Mesozoic kimberlites in the northern Olenek province to the northwest is some 50–60 km thinner, as a consequence of Devonian rifting and Triassic volcanism and the pipes are barren (Griffin et al., 1996; Milashev, 1974). The root to the Archaean Wyoming province has been strongly modified and partially destroyed in a later collisional orogeny (Eggler et al., 1988) that has resulted in the tectonic stacking of wedges of Archaean lithosphere within younger, more fertile lithosphere (Griffin et al., 2004) but Devonian kimberlites intruding the province contain diamonds. Similarly, the Bushveld igne-

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ous complex event has had a marked impact resulting in a significant P-wave velocity anomaly (Helmstaedt and Gurney, 1995; James and Fouch, 2002), resetting of Archaean lithospheric mantle ages to Proterozoic (~ 2.05 Ga; Carlson et al., 1999), iron metasomatism and, according to Griffin et al. (2003), probable destruction of most pre-2 Ga diamonds in the mantle roots beneath the Premier kimberlite. However, as noted by James and Fouch (2002), the abundance of diamonds in other surrounding younger kimberlite pipes indicates that the thermal impact of the event responsible for the Bushveld intrusion has not been sufficient to raise the geotherm of the cratonic root sufficiently to move it out of the diamond stability field. Nor did the plume responsible for the McKenzie dike swarm affect the diamonds within the Slave Province cratonic root (Helmstaedt and Gurney, 1995), some of which contain inclusions of Archaean (~ 3.4 Ga) age (Westerlund et al., 2003). Similarly, the widespread Cambrian flood basalt magmatism tectonothermal event on Northern Australia (Handley and Wingate, 2000) has not destroyed the diamond in the mantle source regions of the younger Timber Creek and Merlin kimberlites or alternatively the diamonds are very young. The impact of a mantle-root destructive event on diamond survival therefore is likely to be determined by proximity to the tectonothermal event and its magnitude (Helmstaedt and Gurney, 1995). This is further complicated by the potential for multiple diamond-forming events (Pearson et al., 1999) and the possibility that the ages of inclusions in diamond might not necessarily be the age of diamond formation (e.g. Spetsius et al., 2002). The evidence for post-Archaean diamond-forming events, the existence of diamond deposits in regions where evidence of unequivocal widespread Archaean lithosphere is lacking, and the stability of the lithosphere in these regions, both thermal and mechanical, is an important factor in assessing their economic importance. There is broad agreement that kimberlite magmas are small volume melts rich in H2O and CO2 formed by low degrees of partial melting of metasomatised mantle peridotite (Eggler, 1989; Tainton and McKenzie, 1994). The high pressures and comparatively low temperatures hinder melt production beneath thick lithosphere, and either high excess temperatures or the presence of volatiles to lower the peridotite solidus are required to generate significant melt in these

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regions. Exactly what triggers and controls the distribution of kimberlite volcanism is unclear and a number of differing tectonic models have been proposed. In some areas, such as western Africa and the Siberian shield, kimberlite volcanism appears to take advantage of pre-existing weaknesses when erupting through the cratons, as suggested by their proximity to major intercratonic fault zones, and a relationship between kimberlite volcanism and continental rifting has been proposed (e.g. Haggerty, 1992; Kaminsky et al., 1995; White et al., 1995). Elsewhere, observed systematic changes in kimberlite ages across cratons have been inferred to indicate the migration of mantle plumes (Heaman and Kjarsgaard, 2000) or, alternatively, traces of deep-seated subduction (Helmstaedt and Gurney, 1997; McCandless, 1999). Heaman et al. (2004) concluded that no single tectonic process could adequately explain the distribution of kimberlite events in North America, but showed that in some provinces kimberlite events could be linked to known mantle heat sources; notably, mantle plume hotspots and upwelling asthenosphere associated with continental rifting. Models of melt migration under mid-ocean ridges have emphasised the importance of stress focusing and strain localization beneath a ridge to localize melt migration patterns (Scott and Stevenson, 1989; Sleep, 1988; Sparks and Parmentier, 1991; Spiegelman and McKenzie, 1987; Stevenson, 1989). Similar controls probably apply to kimberlites and related volcanism, where deep-seated sub-lithospheric stress gradients result in large velocity gradients in the mantle around a craton, focusing early-stage melt migration. These lithospheric stresses also have a first-order influence on the structural controls which in some areas influence the distribution of kimberlite volcanism (Haggerty, 1992; Kaminsky et al., 1995; White et al., 1995). These magmas must ascend with a velocity of around ~ 30 km/h in order for the diamonds entrained within them to survive (e.g. Kelley and Wartho, 2000). Such velocities can only be attained by a fracture propagation mechanism (Eggler, 1989). Finally, a kimberlite pipe erupted into an unstable environment may be eroded, deformed by later tectonic activity, or covered by subsequent sedimentation or volcanic activity. Thus, the chances of identifying such a pipe are greater in stable continental regions, since the survival of the surface expression of a pipe is greater in these terrains.

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Many factors in the formation of economic diamond occurrences are influenced by geodynamic controls. We model three of these factors in this paper: (1) the thermal stability of the continental lithosphere, (2) the focusing of stress gradients (and, subsequently, strain, affecting melt migration paths) within and immediately beneath the deep lithosphere, and (3) the production of melt from peridotite at depth.

0

Stress (MPa) 400

0 τ0

τ = τ0 + τ1P

Depth (km)

Ductile lower crust

τ1 70 Yield point Root zone? Ductile mantle

2. Description of model

Liq

Tp

Pressure (GPa)

The models we present in this paper are based on a particle-in-cell finite element code, previously described elsewhere (Moresi et al., 2002, 2001; Moresi and Solomatov, 1998). The code combines the traditional generality of the finite element method with a Lagrangian point integration method, which allows the tracking of the material and strain history throughout the large deformations encountered in mantle convection (Moresi et al., 2002). We use a viscoplastic rheology with an extreme temperature dependence based on the Frank-Kamenetski approximation (Frank-Kamenetskii, 1969), and which varies over five orders of magnitude. A strainweakening mechanism is incorporated, which results in more stable, long-lived fault zones. Our rheological model is shown in Fig. 2A. The continents are modeled as chemically distinct simple blocks, with a high enough strength to resist deformation and convective recycling. This is achieved by large values of cohesion (s 0) and coefficient of friction (s 1), and viscosity. We do not consider the effects of a ductile lower crust in this paper as we are only interested in the deep continental stress field. We model the unique refractory, buoyant nature of Archaean sub-continental lithospheric mantle by extending the depth of the chemical boundary layer in Archaean terrains. This effectively extends the depth to viscous, ductile deformation in these regions (see Fig. 2). This has the important effect of extending the thermal boundary layer in these regions (see Fig. 3), but also influences the thermal structure of the rest of the continent. Our melt production model is similar to that outlined in de Smet et al. (1998, 1999), and the details of it are shown in Fig. 2B. We only consider a simple linear solidus, and liquidus, in this paper, based on

140 uid

us

Ts

5

0.8

10

TL

1.0

F

So

lid

0.6

us

0.4

15

0.2 0.4 0.6 0.8

θ 1000

1400 1800 Temperature (°C)

Fig. 2. (A) Typical yield strength envelopes for the simulations presented. Whereas realistic rheologies include a ductile lower crust, we only consider mantle lithosphere stresses, for stable continental areas. Hence the viscosity, cohesion (s 0) and the coefficient of friction (s 1) are chosen large enough so that the continents can resist viscoplastic deformation. Deep continental roots extend the depth of the mantle brittle–ductile transition to greater depths (shown). (B) An illustrative geotherm (solid line), and the solidus and liquidus used in this study (adapted from Takahashi and Kushiro, 1983). The supersolidus temperature h is defined in the text, for the temperature of a particle, the solidus and the liquidus at a given pressure. The inset shows its relation to the melt fraction F, based on the linear model of Jaques and Green (1980, dashed) and McKenzie and Bickle (1988, solid line).

Takahashi (1990) and Takahashi and Kushiro (1983). This linearized solidus is quite close to the third-order solidus of Herzberg and Zhang (1996) for the depth range we’re interested in. When a particle’s temperature T p exceeds that of the solidus at that pressure, then we can define the supersolidus temperature h as: h¼

T p  TS ð P Þ TL ð PÞ  TS ð PÞ

ð2:1Þ

where T S and T L define the solidus and liquidus temperature at pressure P. McKenzie and Bickle

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Fig. 3. Evolution of the temperature field of a convecting system over time (note the use of a non-linear greyscale to highlight TBL features). Surface blocks are chemical distinct continental material, cratonic lithosphere is modeled by extending the depth of the CBL. Basal Rayleigh number is 107, depth of the system is 670 km. Average surface heat flux over the length of the run is shown at the top. Times shown are T = 75, 200 and 325 Ma respectively.

(1988) showed that a simple parameterization exists relating the supersolidus temperature h, and the degree of melting f, and this is shown diagrammatically in Fig. 2B. We show the results for two models, the linear approximation of Jaques and Green (1980), and the third-order polynomial fit of McKenzie and Bickle (1988). We do not consider the ongoing depletion of the residual material after melting, nor do we consider the effect the presence of melt has on the mantle, or melt migration. We are only interested in the distribution of instantaneous melting in this paper. We include a latent heat term in the energy equation, similar to de Smet et al. (1999). Model parameters are listed in Table 1.

3. Results 3.1. Thermal structure of the continental lithosphere The average geotherms for the model in Fig. 3 are shown in Fig. 4 (top). The PT estimates are from xenoliths entrained in kimberlite pipes, the examples are from on-craton diamondiferous pipes (Kimberley, Kaapvaal craton; Finnerty and Boyd, 1987), pipes

peripheral to a craton (East Griqualand; Finnerty and Boyd, 1987), and barren off-craton Namibian pipes (Hanaus and Anis Kubub; Franz et al., 1996). The average off-craton geotherm correlates well with xenolith PT estimates from the Namibian Hanaus and Anis Kubub pipes. The average off-craton geotherm is considerably elevated in temperature compared to its oncraton counterpart. The shallower portions of both geotherms are underestimated due to the simple, uniform heat production assumed for the continental lithosphere. The on-craton geotherm passes through the graphite–diamond transition within the depth of the thermal lithosphere whereas the average off-craton geotherm does not. However, a significant problem with such a simplistic representation is that it does not convey the true variability of the continental geotherms. For example, the on and off-craton geotherms of the model shown in Fig. 3 pass in and out of the diamond stability field many times over the lifetime of the simulation. This behavior is shown in Fig. 4 (bottom). The variations in continental heat flow are a result of the thermal configuration of the mantle around the continent. These thermal perturbations by nature must affect the lower continental thermal field and, consequently, the

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Table 1 Model parameters Symbol

Description

Value

dm dc DT

Depth of convecting upper mantle Depth of chemical boundary layer Temperature change across upper mantle Density of mantle Density of chemically distinct lithosphere Density of crust Thermal expansivity Thermal conductivity Thermal diffusivity Continental crustal heat production Chemically distinct continental lithosphere heat production Mantle heat production

670 km 40–120 km 2200 8C

q mantle q c_litho q crust a k j H crust H cont_mantle H mantle g

A b C0 Bp T 0(solidus) T 1(solidus) T 0(liquidus) T 1(liquidus) L

Viscosity given by FrankKamenetski approximation, g = Ae  bT Prefactor in Frank-Kamenetski viscosity law Temperature coefficient in viscosity law Cohesion Coefficient of friction Peridotite solidus at atmospheric pressure Slope of linearized peridotite solidus Peridotite liquidus at atmospheric pressure Slope of linearized peridotite liquidus Latent heat of melting

3300 kg/m3 3100 kg/m3 2900 kg/m3 3.5  10 5 K 1 4 Wm 1 K 1 1  10 6 m 2s 1 0.1–10  10 10 W/kg 1.5  10 13 W/kg 7.46  10 12 – 7.46  10 13 W/kg 10 19 – 10 23 Pa s

1023 Pa s 5.233  10 2 K 1 20 MPa 0.2 800–1670 8C 0–1.125 8C/MPa 1200–2070 8C 0–1.125 8C/MPa 320 kJ kg 1

continental heat flow in Fig. 4 (bottom) is correlated with the time the geotherms are in the diamond stability field. Though the cratonic geotherms spend more time in the DSF than their non-cratonic counterparts, both exhibit time-dependent behavior. The amount of time the geotherms spend in the DSF is dependent on a number of factors. The proximity to the continental margin is one of the more important primary effects. For example, the creation of continental margins through rifting, and the timedependence of (later) subduction at such margins results in large temperature perturbations in the uppermost mantle in these regions. Terrains distant from these processes exhibit a much lower degree of vari-

ation. This effect is illustrated in Fig. 5A which shows a snapshot of the temperature field for one of our models, together with the percentage time the geotherms in each continental region spend in the DSF. The finite lifetime of the DSF in the proximity of the continental margins is due to the subduction of the cool, oceanic boundary layer. The geotherms in these regions are extremely erratic and depend strongly on the configuration of the subducting slabs. The thermal boundary layer in the continental interior is more stable, and consequently more long-lived, towards the center of the continent. This is reflected in the long lifetime of the DSF in these regions. Fig. 5B shows the combined effect of distance to the continental margin and depth of the stable (nonconvecting) chemical boundary layer (CBL). For a given CBL thickness, the DSF is again less stable and more short-lived near the margins, as opposed to interior regions. However, the depth of the CBL exerts an overwhelming influence, in this example, on the time the geotherms spend in the DSF. This model is an example of the dynamics behind Clifford’s rule. In this case the Archaean craton, with its convectively stable root zone, hosts conditions suitable for diamond preservation over the lifetime of the simulation, whereas the adjacent continental region, without a chemically distinct mantle lithosphere, is far less thermally stable. It is interesting, however, that the thickness of the thermal boundary layer in the non-cratonic terrain is influenced by proximity to the craton. Fig. 5C shows a model with 3 different CBL thicknesses. Proterozoic mantle lithosphere has been shown to exhibit similar chemical distinctions to Archaean mantle lithosphere, though to a lesser degree (Poudjom-Djomani et al., 2001; Griffin et al., 2003). They both can be considered distinct from contemporary mantle lithosphere, even though the thickness of the CBL is not as great in Proterozoic regions as in Archaean. Hence Fig. 5C simulates the relative stability of the chemical mantle lithosphere from these different times. This configuration mimics that of Australia, with the Eastern portion of the continent dominated by Phanerozoic fold belts characterized by a thin mantle lithosphere, the center of the continent predominantly Proterozoic fold belts and basins, and the westernmost portion consisting of Archaean cratons—comprising a thick chemical boundary layer. The thermal litho-

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225

0

Depth km

Griqualand, Kaapvaal Periphery Kimberley, Kaapvaal craton Hanaus & Anis Kubub, Namibia Diamond-graphite transition Peridotite solidus

100

Off-craton On-craton

Heat Flow (mW/m2)

200 0

500

1000 Temperature C

1500

2000

100

Average surface HF 50

Average continental mantle HF Time in DSF 0

0

100

Off-craton 200 Time Myr

300

On-craton 400

Fig. 4. Top: time averaged on and off-craton geotherms, with standard deviations, for the model shown in Fig. 1. Also shown are the peridotite solidus and the graphite–diamond transition (from Boyd et al., 1985), and xenolith PT estimates from the Kimberley and Griqualand pipes (from Finnerty and Boyd, 1987), and the Hanaus and Anis Kubub pipes (from Franz et al., 1996). Bottom: average surface heat flux, and mantle contribution to the continental heat flux, for the model shown in Fig. 1. Also shown are the times the geotherms spend in the diamond stability field (DSF) for the on and off-craton regions.

sphere, while influenced by the chemical structure, is distinct from it, and also depends on the evolution of the thermal boundary layer beneath the continent. The change in the CBL between the Proterozoic and Phanerozoic regions results in a lateral temperature gradient which, in the snapshot shown, initiates a downwelling beneath the continent. It is interesting that the thickness of the thermal boundary layer in the central Proterozoic region is similar to that of the Archaean terrains, despite the difference in their respective CBL. Importantly, the geotherms in the Proterozoic region nearest the Archaean craton spend just as long a time in the DSF as their Archaean-cratonic counterparts. The stability of the geotherms in these

regions is a combination of their interior position and proximity to the Archaean craton. Fig. 5D illustrates the relative importance of continental structure versus the physical properties of the lithosphere at any given location. The cratonic region on the left side is subjected to a long-lived thermal upwelling beneath it. Despite the thickness of the CBL, and its mechanical stability, the thermal field is considerably elevated in this region throughout the length of the simulation. As a result the length of time the geotherms spend in the diamond stability field is negligible. However, the central region is protected from the lateral motion of mantle material at depth due to the surrounding cratonic roots. This again

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Fig. 5. The percentage of time the geotherms spend in the DSF (bar graphs) for different parts of the continent, for a variety of continental configurations. Basal Rayleigh number in each example is 107, with periodic side boundary conditions. Initial thermal configuration for each simulation is based on a previously calculated temperature field from a similar convection model (note the use of a non-linear greyscale to highlight mantle features). The asymmetry in the models arises from long-lived subduction zones, or plumes (e.g. D, left hand side craton) locking onto the continent for most of the simulation.

stabilizes the thermal boundary layer in this region, and the geotherms spend a great portion of the simulation in the DSF. How sensitive are these geotherms to our assumed thermal configuration, particularly to variability in heat production? To give an example, variation in crustal heat production will shift the geotherms in Fig. 4 (top) by a certain amount, which may be estimated by: dc dqc =k:

ð3:1:1Þ

Here d c is the variability in the mean depth of crustal heat production, d qc is the variability in the total crustal heat production, and k the thermal conductivity. Assuming representative values of d c = 5 km, d qc = 20 mW/m2, and k = 3, the variability in the

mantle geotherm is around 30 8C, which is within the variance of the geotherms shown in Fig. 4 (top), and much less than the systematic variation between on and off-craton geotherms at depth. 3.2. Stress focusing within the continental lithosphere Fig. 6 shows snapshots of the temperature fields and melt production fields for simulations with various cratonic root depths. Three horizontal stress profiles are shown at the top of the temperature plots, and geotherms at marked positions are shown to the right. Horizontal melt production profiles are shown at the top of the melt production plots. The presence of heterogeneities in the system affects the propagation

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Fig. 6. Temperature field (left, note the use of a non-linear greyscale to highlight mantle features, continentals are chemically distinct blocks) and melt production (right, light material is supersolidus) for mobile-lid convecting systems, with basal Ra = 1 e7, for different depths (d) of the thickened cratonic root. Normalized horizontal stress profiles are shown to the top of the temperature field for each example, at the depths marked to the left. Normalized geotherms are shown to the right of the temperature field, at the positions shown on the continent. Normalized horizontal melt production profiles are shown to the top of the melt production images, for the depths shown.

of near-surface stress, and this can be seen in the focusing of deviatoric stress near the boundaries of the continent. However, abrupt changes in lithospheric thickness also result in large stress gradients, and this is most clearly seen for d = 120 km in Fig. 6. The details of the deep continental stress field are important in diamond exploration for two reasons. Firstly, the continental stress fields control styles and magnitude of deformation. Many kimberlite occurrences follow major fault networks, particularly trans-lithospheric structures, which are largely controlled by continental stresses. Secondly, melt migration paths, particularly for early stage diffuse melt, are largely controlled by strain localization, and hence the local stress field. This has been demonstrated as important in the localization of melt at mid-ocean ridges (Scott and Stevenson, 1989; Sleep, 1988; Sparks and Parmentier, 1991; Spiegelman and McKenzie, 1987; Stevenson, 1989), and is most likely of fundamental importance in the early migration of kimberlitic melts.

A useful parameter for defining the degree of localization of the stress field is stress ratio c, where c is defined as the average stress within a region located within 33 km of our lithospheric discontinuity, divided by the average cratonic stress. Thus values of c much greater than 1 suggest much greater stresses near the lithospheric discontinuities than those experienced on the craton. This factor obviously depends on the material contrast between the adjacent regions, and also the nature of the change in lithospheric thickness. In these simulations our yield parameters are high enough to stop the continents deforming, and the stress fields are similar to that expected for a stable continental region. Fig. 7A shows the variation in the stress ratio c for different values of the cohesion term s 0 in our yield law, and for different material viscosities. The parameters have been normalized to our starting values, 50 MPa and 1  1022 Pa s. The different values of the stress ratio largely reflect the difference in mantle heat production between the two model suites. There is no

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0.8 0.75

A

Stress ratio

0.7 0.65 0.6 0.55 0.5 Cohesion Viscosity

0.45 0.4 0.001

0.01

1 0.1 10 B0 or Viscosity ratio

100

5

B

Q=0.5 Q=5.0

Stress ratio

4 3 2 1 0 53.6

67.0

80.4 93.8 107.2 120.6 Root depth ( km)

134.0

Fig. 7. (A, top) Normalized cohesion ratio (B0, diamonds) or viscosity ratio (squares), versus the stress ratio, defined as the average stress within 33 km of a lithospheric discontinuity, divided by the average cratonic stress. (B, bottom) Root depth versus the stress ratio, for non-dimensional mantle heat production values of 0.5 and 5.0.

systematic variation in the stress ratio with either of these two descriptors of material strength. These parameters both define a deforming material’s response to applied stress. As such, for our stable continents, there is little relationship between the viscosity, or cohesion, of the lithosphere, and the partitioning of continental stress. Variations in the elastic properties may have more effect on the stress ratio (e.g. Zuber et al., 1989), but this was not explored in this paper. Fig. 7B shows the variation of the stress ratio with the thickness of the cratonic root. The model suites are shown for non-dimensional mantle heat production values of 0.5 and 5.0 (corresponding to mantle heat production values of 7.46  1013 W/kg and 7.46  1012 W/kg respectively). In both cases, a dramatic localization of stress occurs as the cratonic

root thickness increases. This effect is also seen in the snapshots of Fig. 6. The lateral contrast of the mantle lithosphere in the margin regions, with the high viscosity cratonic lithosphere, creates large stress gradients at the boundary between the two. The root also acts as a significant obstacle to mantle flow beneath the continent. Since the deep continental stresses are largely affected by the velocity gradient between the mantle flow field and the stable lithosphere, large stress gradients develop at such lateral discontinuities. Large stress gradients at craton boundaries lead to strain localization, and both these factors have a significant affect on melt migration at depth (Scott and Stevenson, 1989; Sleep, 1988; Sparks and Parmentier, 1991; Spiegelman and McKenzie, 1987; Stevenson, 1989). When a critical melt and volatile density are reached, and the dynamic pressure of the kimberlite is sufficient to fracture the surrounding rock, they will ascend at rapid speeds by fracture propagation. At this stage, the local stress field will have little direct affect on the rapid ascent of the magma. It may have an indirect effect through largescale structural features, which have been shown in some cases to exert some control over kimberlite propagation and emplacement. Prior to this fracturepropagated ascent, the lithospheric stress field will play an important role in the porous flow of extracted magma (Stevenson, 1989), and its shear localization to form channels and interconnected networks (Kelemen et al., 1997, 1995). We suggest that the large stress gradients at lithospheric discontinuities will focus sub-continental melt towards these regions, and provide favorable conditions for the pooling of melt. 3.3. Sub-lithospheric melt production The melt-production field shown in Fig. 6 is based on the peridotite solidus and liquidus shown in Fig. 2, and highlights some important points. Firstly, melt production beneath continental regions is significantly depressed with respect to oceanic regions. There are two factors at work here: the continental geotherms at these depths are significantly lower than their oceanic counterparts, and also the continental lithosphere is significantly thicker, so the depths to which hot, upwelling mantle can be advected to are not as shallow. These factors also

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0.7 0.04

0.8

Slope of solidus 0.9 1

A Q=5.0

Mp

0.03

Q=0.5

0.02

1.1

T0 varies Margin Q=0.5 Craton Margin Craton Q=2.5 Margin Craton Q=5.0 Slopevaries Margin Q=0.5 Craton

0.01 Q=2.5

0 533

800

1067 1334 T0(°C)

1600

1867

0.04

B

Margin, Q=5.0 Craton, Q=5.0 Margin, Q=0.5 Craton, Q=0.5

0.03 Mp

apply to the contrast between melt production beneath a thick craton, and a thinner continental margin, as seen in Fig. 6. Fig. 8A shows the effect of different solidus and liquidus parameters on the melt production beneath cratons and continental margins. The dry peridotite solidus at atmospheric pressure (T 0) is given by Takahashi and Kushiro (1983) as 1115 8C. For more depleted peridotite (i.e. greater olivine fraction), this value of T 0 may be N 1600 8C: we predict no melt beneath the craton for this value and minimal melt under the margins. In order to generate significant amounts of melt, we had to use values of T 0 less than ~ 1100 8C. The presence of volatiles in the mantle would have the effect of lowering the solidus temperatures for a given pressure within the mantle— albeit in a more complicated manner than we have modeled. The ZIVIC peridotite solidus (from Boyd et al., 1985, in the presence of CO2 and water), for instance, has a T 0 of b1000 8C, and we generate significant melt, at least beneath the margins, for these low values. The mantle heat production naturally has an important effect in the magnitude of melt production, in governing the mantle internal temperatures. We also varied the slope of the solidus and liquidus (always assumed parallel, top scale). As the slope of the solidus–liquidus (T 1) decreases, the melt fraction increases. We note the values of T 1 vary considerable over the upper mantle, even becoming negative for the wet peridotite solidus. Fig. 8B shows the variation in melt production with increasing root thickness, for non-dimensional mantle heat production values of 0.5 and 5.0 (corresponding to mantle heat production values of 7.46  1013 W/kg and 7.46  1012 W/kg respectively). The melt production parameter Mp is the percentage melt by volume in the sub-lithospheric mantle. In both cases, the melt production in the continental margin regions is significantly higher than in the cratonic regions. This is the case even when the two are the same thickness, and is a consequence of the proximity of the margins to the oceanic regions. The contrast in the melt production between margins and cratons increases with increasing root depth—for large imposed root depths the cratonic melt production is zero. This highlights the difficulty in generating significant amounts of melt beneath thick cratons, unless volatiles are introduced to the system.

229

0.02

0.01

0 53.6

67

80/4 93.8 107.2 Root depth (km)

120.6

0.2

Fig. 8. (A, top) Surface temperature of solidus (T0, bottom axis) or slope of solidus (or liquidus, top scale) versus the sub-continental total volumetric melt fraction. The relationship is shown for sublithospheric melting beneath the cratonic and continental margin regions (see Fig. 6), for a variety of mantle heat production values (listed). (B, bottom) Depth of the cratonic root, versus melt fraction, for both cratonic and continental margin regions.

The production of kimberlite/lamproite magmas is likely to be severely constrained by continental keel geometry. In our models, the excess temperature of a hot plume or upwelling may be sufficient to generate minor melt beneath thick cratonic lithosphere. For high volatile contents, this scenario may be enough to generate small-volume kimberlitic melts. However, significant volumes of melt only occur once a hot upwelling material is advected around the thick root, into regions of thinner lithosphere. For low volatile contents, melting may only initiate when the mantle adiabatically decompresses as it advects up around the continental root zone. The depth range for the generation of kimberlitic volcanism is poorly constrained (between the base of the lithosphere and ~ 670 km;

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Eggler, 1989; Tainton and McKenzie, 1994), and the depth of melt generation in our models (~ 120–400 km) is broadly consistent with what is known of kimberlite formation. How much water is required to generate significant melt beneath cratons? A parameterization of the peridotite solidus as a function of water content is given by Katz et al. (2003). To borrow an example of theirs, at 1250 8C, 1.5 GPa and 0 wt.% (by weight) water content, the melt fraction is 0. As the water content increases to 0.1 wt.%, the melt fraction becomes 0.05. For 0.25 wt.% water the melt fraction is 0.1. Assuming a similar magnitude of dependence for sub-cratonic conditions, this implies the addition of around 0.1–0.25 wt.% water in the mantle for production of kimberlitic magmas.

4. Discussion The interaction of continental geodynamics and mantle convection plays a large role in determining where, and for how long, diamonds will be stable in the continental lithosphere. Similarly, continental alkaline volcanism is also controlled by mantle dynamics, in providing thermal conditions conducive to melt production, delivering volatiles to the sub-continental mantle, and focusing melt migration in regions of high stress gradients and shear localization. Our results suggest that continental regions away from active margins, with deep continental roots, or near regions with deep root zones, are all conducive to the survival of diamonds over long periods of time. We also suggest that melt production will be extremely limited in regions of thick lithosphere, and would require either significant volumes of volatiles in the sub-lithospheric mantle or a major focused thermal input from the asthenosphere (plume or up-welling). Not surprisingly, continental regions of thin lithosphere will produce significantly more melt than regions with thick continental roots. How do these results apply to Australia? Fig. 9A shows the observed velocity variations beneath Australia at 200 km depth (from Simons et al., 2002, 1999; see also Kennett, 2003). A primary feature is that the thickest portions of Australia lie not under the major Archaean cratons, but beneath the central Proterozoic domains. The boundary of this velocity perturbation to

the north lies within what is known as the North Australia craton (Fig. 9A). Our modeling suggests that this feature is stable (see Fig. 9B); it lies away from active margins, and is situated between major cratons (the West Australian craton incorporating the Archaean Pilbara and Yilgarn cratons, the North Australian craton, and the South Australian craton incorporating the Gawler craton to the south). The west, north and south Australian cratonic elements are believed to have amalgamated by 1.60 Ga, with the South Australian craton then breaking away between 1.45 and 1.10 Ga but reattaching by 1.10 Ga (Myers et al., 1996; Betts et al., 2002). Such a thick, stable lithosphere is not a favorable situation for (alkaline) volcanism, and both the paucity and distribution of known Australian alkaline volcanism (Fig. 1; Jaques et al., 1995) postdating the amalgamation of the Precambrian terrains supports this interpretation. Neoproterozoic sub-alkaline mafic magmatism–mafic dyke swarms in the Gawler craton, the Musgrave province and the Curnamona craton, and basaltic volcanism in the Amadeus Basin–has been linked to development of rift and sag basins along the eastern margin and in the interior of the continent as part of a continental breakup episode (Betts et al., 2002). Geochemical tomographic mapping indicates that the Gawler Craton lithosphere was thinned by some 10–15 km between the Permian and Jurassic, probably as a consequence of Pangean rifting (Gaul et al., 2003). The widespread distribution of Phanerozoic alkaline magmatism, notably Tertiary basaltic volcanism, in the Tasman orogenic system of Eastern Australia (Johnson et al., 1989) is also consistent with our result that sub-continental melt production is enhanced in regions of thinner lithosphere. Although it is not possible, within the limits of the datasets, to conclusively establish whether alkaline volcanism follows boundaries in lithospheric thickness, some features certainly suggest this may be the case. Most of the known Precambrian alkaline intrusions in Australia lie at or near the margins of individual Precambrian cratons and terrains. For example, alkaline magmatism (kimberlite, lamproite, lamprophyre, carbonatite) appears to be focused, with few exceptions, around the margins of the Kimberley craton, and at the northern and eastern margins of the Yilgarn craton (Jaques et al., 1995; Graham et al., 2004; Jaques, 2002). The Phanerozoic kimberlites and related rocks are largely located at or near the

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231

Fig. 9. (A) Tomographic depth slice at 200 km through the model of Simons et al. (1999). The major fast velocity anomalies at this depth are beneath the central Australia and the southern portion of the Yilgarn craton. White diamonds indicate the position of the Argyle (C) and Merlin (D) diamond mines. (B) Predicted time spent in the diamond stability field, based on the results of our modelling. Effects include distance to the continental–oceanic crustal boundary, nature of the lithosphere (probable Archaean cratonic lithosphere or not), and distance to a craton. Effects are equally weighted and gridded on a 1  1 degree grid. (C and D) Cross-sections though the tomographic model of Simons et al. (2002). Position of cross-sections shown in Fig. 4 (top). Dark solid lines correspond to the position of the Argyle and Merlin diamond mines, shown in Fig. 4 (top) as white diamonds. Depth of sections is 400 km, all tomographic images use same scale bar.

margin of the Precambrian shield. For example, the Jurassic kimberlites and lamprophyres, with the apparent exception of the Timber Creek kimberlites, lie at the southern and western margins of the Precambrian shield and appear to be associated with the early stages of breakup of Gondwanaland (Jaques, 1994). Can we say which of our modeled effects, melt production or melt focusing, may be more relevant to the associated of kimberlites and lamproites with cratonic margins? A hint is given by the Ellendale lamproites. These young (~ 20 Ma) volcanics formed in a Proterozoic mobile belt. There is little evidence to suggest significant deformation across this belt for most of the Cenozoic. Thus it is unlikely focusing of melt in high strain regions played a significant role in the generation of the Ellendale kimberlites. It is probable that the distribution of the Ellendale lamproites, at least, is largely the result in enhanced

melt production at the periphery of a thickened lithosphere. Fig. 9C and D show N–S cross-sections of the tomography model of Simons et al. (2002, 1999). These cross-sections pass through the locations of the Argyle and Merlin deposits, respectively. The locations of these deposits are characterized by abrupt depth changes of the shear wave velocities. Jaques and Milligan (2004) observed that most of the Australian diamondiferous intrusions are located along such gradients in seismic velocity. The position of Argyle at the edge of the Kimberley craton suggests that it also formed at a steep gradient in lithospheric thickness. While upper mantle shear wave velocity variations are generally associated with temperature perturbations, the variation in near-surface anomalies could also be explained by a variation of the CBL in these regions. The thick seismic lithosphere begins in

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the North Australian craton (Simons et al., 2002, 1999; Kennett, 2003), which contains scattered late Archaean inliers, and extends southward into the central Australian Proterozoic basins where it attains maximum thickness under the Arunta region. Whether this is related to a compositional variation in the lithosphere or, as seems more likely, a result of Australia’s geodynamical evolution, is unclear. The association of diamond occurrences with these velocity gradients suggests a confluence of effects; stable diamonds in the thick lithosphere on one side, favorable conditions for melt production on the other, and steep stress gradients and melt focusing in the transition zone. The distribution of non-economic diamond occurrences is also consistent with this relationship. Timber Creek and Roper, in the Northern Territory, both fall along the margin of the thick central Australian seismic lithosphere (Jaques and Milligan, 2004). The association of diamond occurrences with the margin of the Kimberley block is well documented (Jaques et al., 1986; Janse and Sheahan, 1995; Jaques, 1994). As noted earlier, the Wandagee pipes (Fig. 1; Jaques et al., 1989) lie on the western margin of the Gascoyne Complex between the Pilbara and Yilgarn cratons, the Nabberu pipes are on the northern margin of the Yilgarn Craton (Jaques, 2002), and the Orrorro kimberlites lie to the east of the Gawler Craton (ScottSmith et al., 1984). Most other reported diamonds are either alluvial finds of uncertain providence or, in the case of eastern Australia, of uncertain origin, and possibly subduction related (Davies et al., 1999; Barron et al., 1996), and are inferred to have been carried to the surface by Cenozoic alkali basalts associated with hotspots (Sutherland, 1996). An important question, and one that is difficult to assess through geodynamic modelling, is how the diamonds are originally formed. Many of the more important South African diamond mines lie within a low Pvelocity anomaly that coincides with the region of the Bushveld intrusion and extends NW to the Okwa Belt and possibly the Magondi-Limpopo zone in Botswana at the northwest and western margins of the Kaapvaal and Zimbabwe cratons respectively (James and Fouch, 2002). At the surface this region registers a strong Proterozoic (~ 2 Ga) overprint but the lithosphere sampled by the Jwaneng and Premier kimberlites, although Archaean depleted peridotite, includes eclogitic dia-

mond inclusions dated at 2.9 and 1.5 Ga (Jwaneng), and also 1.9 Ga lherzolitic and 1.2 Ga eclogitic diamond inclusions (Richardson et al., 2004). These data, and isotopic dating of inclusions in diamonds from the Slave and Siberian cratons point to discrete periods of formation of diamond in the Proterozoic: these may be superimposed on or added to earlier generations of diamond stored in the lithospheric mantle keel (Richardson et al., 2004; Shirey et al., 2002, 2004). Many of the younger generations of diamond are eclogitic and are inferred to be products of Proterozoic subduction at the craton margins. Limited information is available on the age of formation of Australian diamonds. Taylor et al. (1990) used nitrogen aggregation characteristics to identify time–temperature populations and show that there was more than one Ellendale diamond-forming event. They identified at least two diamond-forming events–Proterozoic (peridotitic) and Phanerozoic (eclogitic)–in the Ellendale diamond populations. These diamonds are distinct from the Argyle eclogitic diamonds which formed under higher temperatures with inclusions dated at 1.58 Ma (Richardson, 1986). The peridotitic diamonds at Argyle are inferred to be sourced from lithospheric mantle of late Archaean age (Jaques et al., 1990; Graham et al., 1999). Given the evidence for multiple episodes of diamond formation that extend into the Proterozoic and perhaps Phanerozoic, and the existence of diamondiferous intrusions (lamproites and sub-economic kimberlites) within dominantly Proterozoic provinces (Jaques and Milligan, 2004), we suggest that exploration guidelines based strictly on bClifford’s ruleQ should be used with some caution. Diamond formation is intricately tied with the assembly and formation of the continental lithosphere, and a more complete understanding of these processes and their relation to diamond formation (and subsequent survival in the mantle source regions) is required to enable confident prediction of regions likely to host diamond deposits.

5. Conclusion While Australia’s economic diamond deposits are not well predicted by conventional exploration guidelines, their distribution is dependent on Australia’s lithospheric structure and geodynamic evolution. The

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thickest portions of the Australian seismic lithosphere occur under central Australia, not Australia’s major Archaean cratons. We have used a numerical mantle convection code (Moresi et al., 2002) to show that such structures are stable, and host conditions conducive to diamond preservation over geological timescales. Australia’s economic diamond mines occur at step changes in lithospheric thickness, probably as the conditions favoring diamond stability and alkaline volcanism simultaneously occur in these regions.

Acknowledgements The authors greatly appreciate the contribution of Frederick Simons in providing the tomography figures based on his most recent model. We also thank Barry Drummond and Peter Milligan for helpful comments on the draft manuscript, and Mike Sandiford and two anonymous reviewers for helpful comments. Jaques published with the permission of the CEO, Geoscience Australia.

References Atkinson, W.J., Smith, C.B., Danchin, R.V., Janse, A.J.A., 1990. Diamond deposits of Australia. In: Hughes, F.E. (Ed.), Geology of Mineral Deposits of Australia and New Guinea. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 69 – 76. Barron, L.M., Lishmund, S.R., Oakes, G.M., Barron, B.J., Sutherland, F.L., 1996. Subduction model for the origin of some diamonds in the Phanerozoic of eastern New South Wales. Australian Journal of Earth Sciences 43, 257 – 267. Belousova, E.A., Griffin, W.L., Shee, S.R., Jackson, S.E., O’Reilly, S.Y., 2001. Two age populations of zircons from the Timber Creek kimberlites, Northern Territory as determined by laserablation ICP-MS analysis. Australian Journal of Earth Sciences 48, 757 – 765. Berryman, A.K., Stiefenhofer, J., Shee, S.R., Wyatt, B.A., Belousova, E.A., 1999. The discovery and geology of the Timber Creek kimberlites, Northern Territory, Australia. In: Gurney, J.J., Gurney, J.L, Pascoe, M.D., Richardson, S.H. (Eds.), The J.B. Dawson Volume, Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 30 – 39. Betts, P.G., Giles, D., Lister, G.S., Frick, L.R., 2002. Evolution of the Australian lithosphere. Australian Journal of Earth Sciences 49, 661 – 695. Boxer, G.L., Jaques, A.L., 1990. The argyle (AK1) diamond deposit. In: Hughes, F.E. (Ed.), Geology and Mineral Deposits of

233

Australia and New Guinea, Monograph, vol. 14. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 697 – 706. Boyd, F.R., 1989. Composition and distinction between oceanic and cratonic lithosphere. Earth and Planetary Science Letters 96, 15 – 26. Boyd, F.R., 1997. Origin of peridotite xenoliths: major and trace element considerations. In: Ranalli, G., Ricci Lucchi, F., Ricci, C.A., Trommsdorff, T. (Eds.), High Pressure and High Temperature Research on Lithosphere and Mantle Materials. University of Siena, pp. 89 – 106. Boyd, F.R., Gurney, J.J., Richardson, S.H., 1985. Evidence for a 150–200 km thick Archaean lithosphere from diamond inclusion thermobarometry. Nature 315, 387 – 389. Carlson, R.W., Irving, A.J., Hearn, B.C., 1999. Chemical and isotopic systematics of peridotite xenoliths from the Williams Kimberlite, Montana: clues to processes of lithosphere formation, modification and destruction. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), The J.B. Dawson Volume; Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 90 – 98. Clifford, T.N., 1966. Tectono-metallogenic units and metallogenic provinces of Africa. Earth and Planetary Science Letters 1, 421 – 434. Cull, J.P., 1982. An appraisal of Australian heat flow data. BMR Journal of Australian Geology and Geophysics 7, 11 – 21. Davies, R.M., O’Reilly, S.Y., Griffin, W.L., 1999. Diamonds from Wellington, NSW: insights into the origin of eastern Australian diamonds. Mineralogical Magazine 63 (4), 447 – 471. de Smet, J.H., van den Berg, A.P., Vlaar, N.J., 1998. Stability and growth of continental shields in mantle convection models including recurrent melt production. Tectonophysics 296, 15 – 29. de Smet, J.H., van der Berg, A.P., Vlaar, N.J., 1999. The evolution of continental roots in numerical thermo-chemical mantle convection models including differentiation by partial melting. Lithos 48, 153 – 170. Eggler, D.H., 1989. Kimberlites: how do they form? In: Ross, J., Jaques, A.L., Ferguson, J., Green, H., O’Reilly, S.Y., Danchin, V., Janse, A.J.A. (Eds.), Kimberlites and Related Rocks, Special Publication-Geological Society of Australia, vol. 14. Geological Society of Australia, Sydney, NSW, Australia, pp. 489 – 504. Eggler, D.H., Meen, J.K., Welt, F., Dudas, F.O., Furlong, K.P., McCallum, M.E., Carlson, R.W., 1988. Tectonomagmatism of the Wyoming Province. Colorado School of Mines Quarterly 83, 25 – 40. Fielding, D.C.., Jaques, A.L., 1989. Geology, petrology and geochemistry of the Bow Hill lamprophyre dikes, Western Australia. In: Ross, J., Jaques, A.L., Ferguson, J., Green, H., O’Reilly, S.Y., Danchin, V., Janse, A.J.A. (Eds.), Kimberlites and related rocks, Special Publication-Geological Society of Australia, vol. 14. Geological Society of Australia, Sydney, NSW, Australia, pp. 207 – 219. Finnerty, A.A., Boyd, F.R., 1987. Thermobarometry for garnet peridotites: basis for the determination of thermal and compositional structure of the upper mantle. In: Nixon, P.H. (Ed.), Mantle Xenoliths. John Wiley and Sons Ltd., Chichester, UK, pp. 381 – 402.

234

C.J. O’Neill et al. / Tectonophysics 404 (2005) 217–236

Frank-Kamenetskii, D.A., 1969. Diffusion and Heat Transfer in Chemical Kinetics. Plenum, New York. Franz, L., Brey, G.P., Okrusch, M., 1996. Re-equilibration of ultramafic xenoliths from Namibia by metasomatic processes at the mantle boundary. Journal of Geology 104 (5), 599 – 615. Gaul, O.F., O’Reilly, S.Y., Griffin, W.L., 2003. Lithosphere structure and evolution in southeastern Australia. Special Publication-Geological Society of Australia 22. Special PaperGeological Society of America 372, 185 – 202. Graham, S., Lambert, D.D., Shee, S.R., Smith, C.B., Reeves, S., 1999. Re–Os isotopic evidence for Archean lithospheric mantle beneath the Kimberley Block, Western Australia. Geology 27, 431 – 434. Graham, S., Lambert, D., Shee, S., 2004. The petrogenesis of carbonatite, melnoite and kimberlite from the Eastern Goldfields Province, Yilgarn Craton. Lithos 76 (1–4), 519 – 533. Griffin, W.L., Kaminsky, F.V., Ryan, C.G., O’Reilly, S.Y., Win, T.T., Ilupin, I.P., 1996. Thermal state and composition of the lithospheric mantle beneath the Daldyn kimberlite field, Yakutia. Tectonophysics 262, 19 – 33. Griffin, W.L., Zhang, A., O.’Reilly, S.Y., Ryan, C.G., 1998. Phanerozoic evolution of the lithosphere beneath the Sino-Korean Craton. In: Flower, M., Chung, S.L., Lo, C.H., Lee, T.Y. (Eds.), Mantle Dynamics and Plate Interactions in East Asia, Geodynamics Series, vol. 27. American Geophysical Union, Washington, DC, pp. 107 – 126. Griffin, W.L., Ryan, C.G., Kaminsky, F.V., O’Reilly, S.Y., Natapov, L.M., Winn, T.T., Kinny, P.D., Ilupin, I.P., 1999. The Siberian lithosphere traverse: mantle terranes and the assembly of the Siberian Craton. Tectonophysics 310, 1 – 35. Griffin, W.L., O’Reilly, S.Y., Abe, N., Aulbach, S., Davies, R.M., Pearson, N.J., Doyle, B.J., Kivi, K., 2003. The origin and evolution of Archean lithospheric mantle. Precambrian Research 127, 19 – 41. Griffin, W.L., O’Reilly, S.Y., Doyle, B.J., Pearson, N.J., Coopersmith, H., Kivi, K., Malkovets, N., Pokhilenko, N., 2004. Lithosphere mapping beneath the North American plate. Lithos 77, 873 – 922. Gurney, J.J., Moore, R.O., Otter, M.B., Kirkley, M.B., Hops, J.J., McCandless, T.E., 1991. Southern African kimberlites and their xenoliths. Magmatism in Extensional \Tectonic Settings: The Phanerozoic African Plate, pp. 495 – 536. Haggerty, S.E., 1992. Diamonds in West Africa: tectonic setting and diamond productivity. Russian Geology and Geophysics 33, 35 – 49. Handley, L.M., Wingate, M.T.D., 2000. SHRIMP zircon age for an Early Cambrian dolerite dyke: an intrusive phase of the Antrim Plateau Volcanics of northern Australia. Australian Journal of Earth Sciences 47 (6), 1029 – 1040. Heaman, L.M., Kjarsgaard, B.A., 2000. Timing of eastern North American kimberlite magmatism: continental extension of the Great Meteor hotspot track? Earth and Planetary Science Letters 178, 253 – 268. Heaman, L.M., Kjarsgaard, B.A., Creaser, R.A., 2004. The timing of kimberlite magmatism in North America: implications for global kimberlite genesis and diamond exploration. Lithos 76 (1–4), 377 – 397.

Helmstaedt, H.H., Gurney, J.J., 1995. Geotectonic controls of primary diamond deposits: implications for area selection. Journal of Geochemical Exploration 53, 125 – 144. Helmstaedt, H.H., Gurney, J.J., 1997. Geodynamic controls of kimberlites—what are the roles of hotspot and plate tectonics? Geologiya I Geofizika 38 (2), 461 – 476. Herzberg, C., Zhang, J., 1996. Melting experiments on anhydrous peridotite KLB-1: compositions of magmas in the upper mantle and transition zone. Journal of Geophysical Research 101, 8271 – 8295. James, D.E., Fouch, M.J., 2002. Formation and evolution of Archaean cratons: insights from southern Africa. In: Flower, C.M.R., Ebinger, C.J., Hawkesworth, C.J. (Eds.), The early earth: physical, chemical and biological development, Special Publication-Geological Society of London, vol. 199. Geological Society of London, London, pp. 1 – 26. Janse, A.J.A., 1994. Is Clifford’s Rule still valid? Affirmative examples from around the world. In: Meyer, H.O.A., Leonardos, O.H. (Eds.), Diamonds: Characterisation, Genesis and Exploration, Proceedings of the Fifth International Kimberlite Conference, Brazil. Companhia de Pesquisa de Recursos Minerais, Brazil, pp. 213 – 235. Janse, A.J.A., Sheahan, P.A., 1995. Catalogue of world wide diamond and kimberlite occurrences: a selective and annotative approach. Journal of Geochemical Exploration 53, 73 – 111. Jaques, A.L., 1994. Diamonds in Australia. In: Solomon, M., Groves, D.I. (Eds.), The Geology and Origin of Australian Mineral Deposits. Clarendon Press, Oxford, pp. 787 – 820. Jaques, A.L., 2002. Australian diamond deposits, kimberlites, and related rocks. 1:5 million map, Geoscience Australia, www.ga. gov.au/pdf/RR0114.pdf. Jaques, A.L., Green, D.H., 1980. Anhydrous melting of peridotite at 0–15 kb pressure and genesis of tholeiitic basalts. Contributions to Mineralogy and Petrology 73, 287 – 310. Jaques, A.L., Lewis, J.D., Smith, C.B., 1986. The kimberlites and lamproites of Western Australia. Bulletin-Geological Survey of Western Australia 132, 239 – 246. Jaques, A.L., Kerr, I.D., Lucas, H., Sun, S.-S., Chappell, B., 1989. Mineralogy and petrology of picritic monchiquites from Wandagee, Carnarvon Basin, Western Australia. In: Ross, J., Jaques, A.L., Ferguson, J., Green, H., O’Reilly, S.Y., Danchin, V., Janse, A.J.A. (Eds.), Kimberlites and related rocks, Special Publication-Geological Society of Australia, vol. 14, pp. 121 – 139, Sydney, NSW. Jaques, A.L., O’Neill, H.St.C., Smith, C.B., Moon, J., Chappell, B.W., 1990. Diamondiferous peridotite xenoliths from the Argyle (AK1) lamproite pipe, Western Australia. Contributions to Mineralogy and Petrology 104, 255 – 276. Jaques, A.L., Creaser, R.A., Ferguson, J., Smith, C.B., 1995. A review of the alkaline rocks of Australia. Transactions of the Geological Society of South Africa 88, 311 – 334. Jaques, A.L., Milligan, P.R., 2004. Patterns and controls on the distribution of diamondiferous intrusions in Australia. Lithos 77 (1–4), 783 – 802. Johnson, R.W., Knutson, J., Taylor, S.R., 1989. Intraplate Volcanism in Eastern Australia and New Zealand. Cambridge University Press, Cambridge, United Kingdom.

C.J. O’Neill et al. / Tectonophysics 404 (2005) 217–236 Jordan, T.H., 1978. Composition and development of the continental tectosphere. Nature 274, 544 – 548. Jordan, T.H., 1981. Continents as a chemical boundary layer. Philosophical Transactions of the Royal Society of London 301, 359 – 373. Jordan, T.H., 1988. Structure and formation of the continental tectosphere. Journal of Petrology Special Volume, 11 – 37. Kaminsky, F.V., Feldman, A.A., Varlamov, V.A., Boyko, A.N., Olofinsky, L.N., 1995. Prognostication of primary diamond deposits. Journal of Geochemical Exploration 53, 167 – 182. Katz, R.F., Spiegelman, M., Langmuir, C.H., 2003. A new parameterization of hydrous mantle melting. Geochemistry, Geophysics, Geosystems 4 (9), 1073. doi:10.1029/2002GC000433. Kelemen, P.B., Shimizu, N., Salters, V.J.M., 1995. Extraction of mid-ocean-ridge basalt from the upwelling mantle by focused flow of melt in dunite channels. Nature 375, 747 – 753. Kelemen, P.B., Hirth, G., Shimizu, N., Spiegelman, M., Dick, H.J.B., 1997. Review of melt migration processes the adiabatically upwelling mantle beneath oceanic spreading ridges. Philosophical Transactions of the Royal Society of London. A 355 (1723), 283 – 318. Kelley, S.P., Wartho, J.A., 2000. Rapid kimberlite ascent and the significance of Ar–Ar ages in xenolith phlogopites. Science 289 (5479), 609 – 611. Kennett, B.L.N., 2003. Seismic structure in the mantle beneath Australia. In: Hills, R.R., Mueller, R.D. (Eds.), Evolution and Dynamics of the Australian Plate, Special Paper-Geological Society of America, vol. 372; Special Publication-Geological Society of Australia 22, 7 – 23. Lee, D.C., Milledge, H.J., Reddicliffe, T.H., Scott-Smith, B.H., Taylor, W.R., Ward, L.M., 1997. The Merlin kimberlites, Northern Territory, Australia. Proceedings of the Sixth International Kimberlite Conference. Allerton Press, New York, pp. 82 – 96. Lee, D.C., Reddicliffe, T.H., Scott Smith, B.H., Taylor, W.R., Ward, L.M., 1998. Merlin diamondiferous kimberlite pipes. In: Berkman, D.A., MacKenzie, D.H. (Eds.), Geology of Australian and Papuan New Guinean Mineral Deposits, Monograph, vol. 22. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 461 – 464. Lenardic, A., Moresi, L., Muehlhaus, H., 2000. The role of mobile belts for the longevity of deep cratonic lithosphere; the crumple zone model. Geophysical Research Letters 27 (8), 1235 – 1238. Lipschutz, M.E., 1964. Origin of diamonds in Ureilites. Science 143, 73 – 86. McCandless, T.E., 1999. Kimberlites; mantle expressions of deepseated subduction. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), The P.H. Nixon Volume, Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 545 – 549. McKenzie, D.P., Bickle, M.J., 1988. The volume and composition of melt generated by extension of the lithosphere. Journal of Petrology 29, 625 – 679. Milashev, V.A., 1974. The Kimberlite Provinces. Nedra, Leningrad, pp. 548 – 574. Moresi, L., Solomatov, V., 1998. Mantle convection with a brittle lithosphere: thoughts on the global tectonic styles of

235

the Earth and Venus. Geophysical Journal International 133 (3), 669 – 682. Moresi, L., Muelhaus, H., Dufour, F., 2001. Particle in cell solution for creeping viscous flows with internal interfaces. Bifurcation and Localization in Soils and Rocks. Balkema, Rotterdam, pp. 345 – 354. Moresi, L., Dufour, F., Muhlhaus, H.B., 2002. A Lagrangian integration point finite element method for large deformation modeling of viscoelastic geomaterials. Journal of Computational Physics 184, 476 – 497. Morgan, P., 1995. Diamond exploration from the bottom up: regional geophysical signatures of lithosphere conditions favourable for diamond exploration. Journal of Geochemical Exploration 53, 145 – 165. Myers, J.S., Shaw, R.D., Tyler, I.M., 1996. Tectonic evolution of Proterozoic Australia. Tectonics 15, 1431 – 1446. O’Neill, C.J., Moresi, L., 2003. How long are diamonds stable in the continental lithosphere? Earth and Planetary Science Letters 213, 43 – 52. O’Reilly, S.Y., Griffin, W.L., Poudjom Djomani, Y.H., Morgan, P., 2001. Are lithospheres forever? Tracking changes in subcontinental lithospheric mantle through time. GSA Today, 4 – 10. Pearson, D.H.., Shirey, S.B., Bulanova, G.P., Carlson, R.W., Milledge, H.J., 1999. Dating and paragenetic distinction of diamonds using the Re–Os isotope system: application to some Siberian diamonds. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), The P.H. Nixon Volume, Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 637 – 643. Pearson, D.G., Irvine, G.J., Carlson, R.W., Kopylova, M.G., Ionov, D.A., 2002. The development of lithospheric keels beneath the earliest continents: time constraints using PGE and Re–Os isotope systematics. In: Flower, C.M.R., Ebinger, C.J., Hawkesworth, C.J. (Eds.), The Early Earth: Physical, Chemical and Biological Development, Special Publication-Geological Society of London, vol. 199, pp. 65 – 90, London. Pidgeon, R.T., Smith, C.B., Fanning, C.M., 1989. Kimberlite and lamproite emplacement ages in Western Australia. In: Ross, J., Jaques, A.L., Ferguson, J., Green, H., O’Reilly, S.Y., Danchin, V., Janse, A.J.A. (Eds.), Kimberlites and Related Rocks, Special Publication-Geological Society of Australia, vol. 14. The Geological Society of Australia, Sydney, NSW, pp. 369 – 381. Poudjom-Djomani, Y.H., O’Reilly, S.Y., Griffin, W.L., Morgan, P., 2001. The density structure of the subcontinental lithosphere through time. Earth and Planetary Science Letters 184, 605 – 621. Priestley, K., McKenzie, D., 2002. The structure of upper mantle beneath southern Africa. In: Flower, C.M.R., Ebinger, C.J., Hawkesworth, C.J. (Eds.), The Early Earth: Physical, Chemical and Biological Development, Special Publication-Geological Society of London, vol. 199, pp. 45 – 64, London. Richardson, S.H., 1986. Latter-day origin of diamonds of eclogitic paragenesis. Nature 322, 623 – 626. Richardson, S.H., Gurney, J.J., Erlank, A.J., Harris, J.W., 1984. Origin of diamonds in old enriched mantle. Nature 310, 198 – 202.

236

C.J. O’Neill et al. / Tectonophysics 404 (2005) 217–236

Richardson, S.H., Harris, J.W., Gurney, J.J., 1993. Three generations of diamonds from old continental mantle. Nature 366, 256 – 258. Richardson, S.H., Shirey, S.B., Harris, J.W., 2004. Episodic diamond genesis at Jwaneng, Botswana, and implications for Kaapvaal craton evolution. Lithos 77, 143 – 154. Richter, F.M., 1985. Models for the Archean thermal regime. Earth and Planetary Science Letters 73, 350 – 360. Scott, D.R., Stevenson, D.J., 1989. A self-consistent model of melting, magma migration and buoyancy-driven circulation beneath mid-ocean ridges. Journal of Geophysical Research 94 (B3), 2973 – 2988. Scott-Smith, B.H., Danchin, R.V., Harris, J.W., Stracke, K.J., 1984. Kimberlites near Orroroo, South Australia. In: Kornprobst, J. (Ed.), Kimberlites and Related Rocks. Elsevier Science, Amsterdam, pp. 121 – 142. Shee, S.R., Vercoe, S.C., Wyatt, B.A., Hwang, PH.H., Campbell, A.N., Colgan, E.A., 1999. Discovery and geology of the Nabberu Kimberlite Province, Western Australia. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), The P.H. Nixon Volume, Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 912 – 922. Shirey, S.B., Harris, J.W., Richardson, S.H., Fouch, M.J., James, D.E., Cartigny, P., Deines, P., Viljoen, F., 2002. Diamond genesis, seismic structure, and evolution of the Kaapvaal–Zimbabwe craton. Science 297, 1683 – 1686. Shirey, S.B., Richardson, S.H., Harris, J.W., 2004. Integrated models of diamond formation and craton evolution. Lithos 77, 923 – 944. Simons, F.J., Zielhuis, A., van der Hilst, R.D., 1999. The deep structure of the Australian continent from surface wave tomography. Lithos 48, 17 – 43. Simons, F.J., van der Hilst, R.D., Montagner, J.P., Zielhuis, A., 2002. Multimode Rayleigh wave inversion for heterogeneity and azimuthal anisotropy of the Australian upper mantle. Geophysical Journal International 151 (3), 738 – 754. Sleep, N.H., 1988. Tapping of melt by veins and dikes. Journal of Geophysical Research 93, 10255 – 10272. Sobolev, A.V., Shatsky, V.S., 1990. Diamond inclusions in garnets from metamorphic rocks: a new environment of diamond formation. Nature 243, 742 – 746. Sparks, D.W., Parmentier, E.M., 1991. Melt extraction from the mantle beneath spreading centers. Earth and Planetary Science Letters 105, 368 – 377. Spetsius, Z.V., Belousova, E.A., Griffin, W.L., O’Reilly, S.Y., Pearson, N.J., 2002. Archean sulfide inclusions in Paleozoic zircon megacrysts from the Mir kimberlite, Yakutia: implications for

the dating of diamonds. Earth and Planetary Science Letters 199, 111 – 126. Spiegelman, M., McKenzie, D., 1987. Simple 2-D models for melt extraction at mid-oceanic ridges and island arcs. Earth and Planetary Science Letters 83, 137 – 152. Stachel, T., 2001. Diamonds from the asthenosphere and the transition zone. European Journal of Mineralogy 13, 883 – 892. Stevenson, D.J., 1989. Spontaneous small-scale melt segregation in partial melts undergoing deformation. Geophysical Research Letters 16, 1067 – 1070. Sutherland, F.L., 1996. Alkaline rocks and gemstones, Australia: a review and synthesis. Australian Journal of Earth Sciences 43 (3), 323 – 343. Tainton, K.M., McKenzie, D., 1994. The generation of kimberlites, lamproites, and their source rocks. Journal of Petrology 35 (3), 787 – 817. Takahashi, E., 1990. Speculations on the Archaean mantle: missing link between komatiite and depleted garnet peridotite. Journal of Geophysical Research 95, 15941 – 15945. Takahashi, E., Kushiro, I., 1983. Melting of dry peridotite at high pressure and basalt magma genesis. American Mineralogist 68, 859 – 879. Taylor, W.R., Jaques, A.L., Ridd, M., 1990. Nitrogen-defect aggregation characteristics of some Australasian diamonds: time– temperature constraints on the source regions of pipe and alluvial diamonds. American Mineralogist 75, 1290 – 1310. Westerlund, K., Shirey, S.B., Richardson, S.H., Gurney, J.J., Harris, J.W., 2003. Re–Os isotope systematics of peridotitic diamond inclusion sulphides from the Panda kimberlite, Slave Craton. Extended Abstract, Eighth International Kimberlite Conference, Victoria. White, S.H., de Boorder, H., Smith, C.B., 1995. Structural controls of kimberlite and lamproite emplacement. Journal of Geochemical Exploration 53, 245 – 264. Wyatt, B.A., Sumpton, J.D.H., Stiefenhofer, J., Shee, S.R., Smith, T.W., 1999. Kimberlites in the Forrest River Area, Kimberley Region, Western Australia. In: Gurney, J.J., Gurney, J.L., Pascoe, M.D., Richardson, S.H. (Eds.), The P.H. Nixon Volume, Proceedings of the Seventh International Kimberlite Conference, Cape Town. Red Roof Design, Cape Town, pp. 912 – 922. Wyatt, B.A., Baumgartner, M., Anckar, E., Grutter, H., 2004. Compositional classification of bkimberliticQ and bnonkimberliticQ ilmenite. Lithos 77, 819 – 840. Zuber, M.T., Bechtel, T.D., Forsyth, D.W., 1989. Effective elastic thicknesses of the lithosphere and mechanisms of isostatic compensation in Australia. Journal of Geophysical Research 94, 9353 – 9367.