The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

3.6 The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths DG Pearson, University of Alberta, Edmonton, AB, Canada N Wittig, Carleton University, Ottawa, ON, Canada ã 2014 Elsevier Ltd. All rights reserved.

3.6.1 3.6.2 3.6.2.1 3.6.2.1.1 3.6.2.1.2 3.6.2.1.3 3.6.2.1.4 3.6.2.2

Introduction Modification of CLM Mantle Processes: Evidence for Metasomatic Modification of Modal Mineralogy Clinopyroxene Orthopyroxene Garnet Olivine Metasomatism, Cooling, and Diffusion: Some Constraints from Nontraditional Stable Isotopes (Li, Mg, and Fe) 3.6.2.3 Metasomatic Effects of Diamond-Forming Fluids 3.6.3 Primary Compositions of Cratonic Peridotites and Their Melting Environment 3.6.3.1 Composition of Cratonic Mantle: Some Comparisons 3.6.3.2 Composition: Depth Variations and the Origin of Cratonic Mantle 3.6.3.3 Trace Element Evidence 3.6.3.4 The Role of Subduction: Was the Melting Environment Hydrous? 3.6.4 Constraining the Timing of Lithosphere Formation 3.6.4.1 Information Derived from Lu–Hf Isotope Systematics 3.6.4.2 Age Constraints from Re–Os Isotopes 3.6.4.2.1 Data presentation 3.6.4.2.2 Analytical approach: Whole-rock and sulfide analyses 3.6.4.2.3 Os isotope heterogeneity of the mantle 3.6.4.3 In Situ Pb–Pb Dating of Clinopyroxenes 3.6.4.4 Case Studies: Dating Cratonic Mantle 3.6.4.4.1 The Kaapvaal craton 3.6.4.4.2 North China craton 3.6.4.4.3 North Atlantic craton 3.6.5 Models for the Formation of Cratonic Roots Acknowledgments References

3.6.1

Introduction

Mantle xenoliths provide rare glimpses into the nature of the continental lithospheric mantle (CLM). For the Earth’s cratons (stable regions of continental rocks with basement ages of at least 2.5 Ga), such as xenoliths, together with diamonds, are the only known direct samples of the basal lithosphere and so they provide a unique perspective on models for craton formation. This chapter provides an addition to the review of the geochemistry of mantle xenoliths provided by Pearson et al. (see Chapter 3.5). It focuses on new evolutionary models and the improved understanding of mantle processes that have been published in the intervening 10 years and hence should be viewed neither as a data update nor as a comprehensive summary of the literature on the geochemistry of mantle xenoliths published over that decade. The aim is to identify significant unresolved issues concerning the genesis of cratonic roots that can be addressed via geochemical investigation of mantle xenoliths. As with the previous contribution, the reader is referred to

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classic works such as Nixon (1973, 1987), Sobolev (1974), Menzies and Hawkesworth (1987), and Menzies (1990a) as benchmark publications on the nature of mantle xenoliths and the information that they provide. Newer syntheses that use summary geochemical databases for mantle xenoliths and diamonds to address the question of lithospheric mantle formation have been published by Carlson et al. (2005), Lee (2006), Stachel and Harris (2008), Pearson and Wittig (2008), Griffin et al. (2009), Lee et al. (2011), and Aulbach (2012). A review of internally consistent mantle thermobarometers has been presented by Nimis and Grutter (2009). Summary evaluations of new methods for constraining cratonic mantle geotherms are given by Michaut et al. (2007, 2009), and Mather et al. (2011). Here, the emphasis is on understanding the evolution of cratonic CLM because of the important role of cratons in crustal evolution and as a reflection of the economic importance of cratonic crust and mantle. As such, the authors will deal with select parts of the literature that are pivotal (in their

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view) to defining the boundary elements of genetic models for cratonic roots. To orientate the reader in terms of the cratonic regions that will be referred to, Figure 1 summarizes the current consensus of the location of the major cratons and shows locations that contain xenoliths or other mantle samples indicating the presence of Archean CLM underpinning them. The significance of these locations and their ages is discussed in the ensuing text. In addition, Figure 1 shows xenolith locations where Proterozoic ages have been obtained. In general, there is a broad correlation between the presence of Archean CLM beneath Archean crust and Proterozoic CLM beneath crust of Proterozoic age, indicating a basic tenet of continental geology that has been evident for almost three decades, namely that there is long-term coupling between the crust and mantle. Nonetheless, there are exceptions to these general observations that will be highlighted in the following text. In assembling data to understand the genesis of CLM, it is important to use data that reflect, as far as possible, the primary composition, that is, the bulk composition that resulted from the dominant processes that lead to the genesis and stabilization of CLM. To focus on primary compositions, it is necessary to identify the secondary processes that are likely to modify these original compositions, and hence, the first task will be to try to trace the mineralogical and compositional changes that result from these processes. The authors then try to constrain what the original bulk compositions of CLM may have been and what they tell us about the craton-forming processes in particular. The age constraints on the CLM are reviewed and some of the advances in knowledge are illustrated with specific examples while highlighting major caveats that have arisen, particularly with attempts to date Phanerozoic CLM. Finally, the genetic models, in particular for cratonic mantle and how they may be integrated into more general craton-formation models, are discussed.

3.6.2

Modification of CLM

3.6.2.1 Mantle Processes: Evidence for Metasomatic Modification of Modal Mineralogy The introduction of new minerals into mantle xenoliths, after initial melt depletion, by the process of metasomatism is a well-known process (e.g., Dawson, 1984; Erlank et al., 1987; Harte et al., 1975; Menzies, 1983). Most of these studies focused on hydrous minerals, such as mica (phlogopite) and amphibole, and on more exotic phases. More recently, the focus has moved back to assessing whether anhydrous phases, thought to be intrinsic to the ‘primary’ mineralogy of peridotite, that is, clinopyroxene (cpx), orthopyroxene (opx), and garnet, are residual or metasomatic in origin.

3.6.2.1.1 Clinopyroxene Within cratonic peridotites, there is now abundant evidence based on trace element and isotopic studies that the cpx within peridotites containing highly magnesian olivine has been introduced, in many cases within a few million years of eruption, by small degree melts that may be related to their host magmas or earlier carbonatitic melts (Boyd et al., 1997; Gre´goire et al., 2002; Malarkey et al., 2011; Pearson and Wittig, 2008; Pearson

et al., 2002b; Rudnick et al., 1993; Simon et al., 2003; van Achterbergh et al., 2001). The evidence for cpx addition comes in various forms. Some diopsides have unusual major element compositions that are indicative of carbonatite metasomatism (Rudnick et al., 1993) or major element and textural disequilibrium with other phases (e.g., Boyd et al., 1997). Trace element evidence includes the observation that very few diopsides in peridotites from any tectonic setting sampled as xenoliths by alkaline volcanic rocks have the depleted rare earth element (REE) patterns required by residues. In addition, there is a general lack of trace element equilibrium with other phases (e.g., Agranier and Lee, 2007; Gre´goire et al., 2002; Pearson et al., 2002b; Rehfeldt et al., 2008; Simon et al., 2003; van Achterbergh et al., 2001). Very few diopsides have Sr–Nd isotopic compositions compatible with having evolved for significant periods as light rare earth element (LREE)-depleted residual phases (Pearson and Nowell, 2002). There is also significant Sr isotopic (Schmidberger et al., 2003), oxygen isotopic (Rehfeldt et al., 2008), and Sr–Nd isotopic (Malarkey et al., 2011) variability between and within individual crystals in the same peridotite. This evidence is in keeping with the widespread indications of refertilization via pyroxene addition and physical intermixing of pyroxenites with peridotites in massif peridotites (e.g., Kornprobst, 1969; Le Roux et al., 2007; Pearson et al., 1993; Riches and Rogers, 2011). Such refertilization, in particular with the addition of new pyroxene from melts, probably explains the much more restricted range of Nd isotopic compositions observed in cpx from peridotite xenoliths (see following texts).

3.6.2.1.2 Orthopyroxene Boyd (1989) highlighted the opx-rich nature of Kaapvaal peridotites. A secondary origin for this opx, linked to Archean subduction-zone metasomatism, was initially suggested by Kesson and Ringwood (1989) and subsequently advocated at a more general level by Rudnick et al. (1994). Kaapvaal lowtemperature peridotites are so orthopyroxene-rich that their average opx contents (and hence Mg/Si ratios) fall far outside the average values for other cratons. In contrast, cratonic peridotites from the North Atlantic craton (NAC) (Greenland) are low in opx, with modal compositions that are dominated by olivine, yet with very similar olivine Mg numbers (Figure 2; Bernstein et al., 2006). In particular, peridotite xenoliths from E Greenland (Bernstein et al., 1998), Ubekendt Ejland, W Greenland (Bernstein et al., 2006), and the Sarfartoq region of W Greenland (Bizzarro and Stevenson, 2003; Wittig et al., 2008a,b) are very olivine-rich such that the average olivine content of NAC peridotites (>89 wt%) makes them the most olivine-rich, opx-poor end-member of the cratonic peridotite suites so far analyzed, plotting close to the ‘oceanic trend’ of Boyd (1989). The Tanzanian (Lee and Rudnick, 1999) and Slave (Kopylova and Russell, 2000) cratons are progressively more olivine-poor, whereas the Kaapvaal craton is the extreme opxrich end-member with only 64 wt% olivine on average. Recent major trace element and isotopic studies (Bell et al., 2005; Kelemen et al., 1998; Lee, 2006; Simon et al., 2007) indicate a secondary origin for the excess opx in Kaapvaal peridotites and, probably, other peridotites that follow distinct trends on Mg/Si versus Al/Si plots toward the addition of

j

Exposed Archean crust Well-defined cratons, part of composite cratons Composite craton outline; Proterozoic amalgamation Craton correlations from Pangea Kimberlite or alkaline rock cluster with mantle xenoliths yielding Archean ages

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Figure 1 Global distribution of the cratons (regions of crust >2.5 Ga). The regional outcrop of Archean crustal rocks is indicated in blue (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 cratons is shown by black dotted lines. Red dashed lines show the estimated extent of cratonic regions amalgamated from Archean blocks during the Proterozoic. Blue dotted lines extended across oceanic areas show links between the cratonic fragments that are thought to have once comprised single cratonic blocks (Bleeker, 2003). Also indicated (blue diamonds) are cratonic regions that contain kimberlite-hosted (or in a few cases alkali-basalt-hosted) xenolith suites that have produced Archean TRD Os model ages (Kaapvaal – Walker et al., 1989; Wyoming – Carlson and Irving, 1994; Siberia – Carlson et al., 1999; Carlson and Moore, 2004; Pearson et al., 1995b; Tanzania – Chesley et al., 1999; Superior – Pearson et al., 1995c; E North Atlantic – Hanghøj et al., 2001; W African – Barth et al., 2001 (eclogite TMA Os model ages); North China craton – Gao et al., 2002; Slave – Aulbach et al., 2004b; Baltic Shield (Karelian craton) – Peltonen and Bru¨gmann, 2006; W North Atlantic craton – Bernstein et al., 2006; Wittig et al., 2010b; and Kalahari (Murowa, Zimbabwe) – Smith et al., 2009). Modified after Bleeker W (2003) The late Archean record: A puzzle in ca. 35 pieces. Lithos 71: 99–134; Pearson DG and Wittig N (2008) Formation of Archean continental lithosphere and its diamonds: The root of the problem. Journal of the Geological Society 165: 895–914.

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Mg number olivine

94.0 93.5

Greenland 'dunites'

93.0 92.5 92.0

Abyssal 91.50 Forearcs OIB (Canaries) Cratons all 91.00ex Kaapvaal Cratons Kaapvaal craton NAC 90.50 Siberian craton Tanzania Churchill 90.00 Slave1.0 102 95 Iwanaidake -HARZ Iwanaidake -DUNITES

90

85

80 75 Olivine (wt%)

70

65

60

Figure 2 Average olivine Mg number versus weight percent olivine for peridotites from different cratons compared with peridotites from present-day tectonic settings (forearcs, OIB – Canary Islands only, abyssal), ophiolites and dunites and harzburgites from the Iwanaidake peridotite, Hokkaido (Kubo, 2002). Field for orthopyroxene (opx)-poor ‘dunites’ from E Greenland (Bernstein et al., 1998) and Ubekendt Ejland, W Greenland (Bernstein et al., 2006), marked as dashed box. Other W Greenland peridotites are more opx-rich (Bizzarro and Stevenson, 2003; Wittig et al., 2008a,b) and influence the less olivine-rich NAC average. Data sources given in Pearson and Wittig (2008). Only data from low-temperature peridotites are plotted in all cratonic examples.

aluminous opx (Aulbach, 2012; Pearson and Wittig, 2008). Whether the opx originates from subduction-related fluids (e.g., Kesson and Ringwood, 1989; Rudnick et al., 1994) or SiO2-rich melts produced by either harzburgite (Pearson and Wittig, 2008) or eclogite melting during craton formation (Aulbach, 2012; Rehfeldt et al., 2008) is currently debated.

3.6.2.1.3 Garnet Garnet has typically been regarded as a primary residual phase in garnet peridotites, and this interpretation is consistent with the persistence of garnet on the peridotite solidus during extensive melting at high pressures (Walter, 1998). Nonetheless, many garnets, especially in cratonic peridotites, display REE and trace element patterns that are consistent with either extreme metasomatic alteration of preexisting garnet (e.g., Hoal et al., 1994) or new garnet formation, either directly from metasomatic melts (Klein-BenDavid and Pearson, 2009; Rehfeldt et al., 2008; Shimizu et al., 1997) or through fluidassisted metamorphic transitions (Malkovets et al., 2007). In some Kaapvaal peridotites, textural and chemical evidence have been combined to make a clear case for a metasomatic origin for garnet (Bell et al., 2005). Lastly, only a restricted fraction of mantle garnets have Nd and Hf isotope compositions consistent with long-term depletion histories (see following text).

3.6.2.1.4 Olivine Finally, even olivine may have been added to cratonic lithosphere or at least had its composition, for example, Mg number, altered by metasomatism. This means that care must be taken in using olivine compositions to determine melting extents and environment. For instance, Boyd et al.

(1983) and Rehfeldt et al. (2007) document the addition of relatively Fe-rich (Fo86–89) dunites to the Kaapvaal craton subcontinental lithospheric mantle (SCLM), probably due to the high-pressure fractional crystallization of Karoo magmas in the mantle. Olivines of low Mg number (86–88) have also been identified in some Greenland peridotites (Sand et al., 2009). These rocks appear to have experienced significant olivine addition and are hence strongly modified from their original solidus compositions. Such variations are unusual, however, and the great majority of olivines from peridotite xenoliths, especially those from cratonic regions, have compositions more refractory than estimates of fertile mantle (Figure 2). An analysis of global olivine data for cratonic xenoliths (Figure 3; Bernstein et al., 2007; Pearson and Wittig, 2008) shows a very pronounced mode in Mg number at 92.6, but with a wide spread, to higher and lower values. The mode can be interpreted to reflect the most common level of melt depletion (e.g., Bernstein et al., 2007; Pearson and Wittig, 2008). The tail to lower Mg numbers in Figure 3 could be interpreted as evidence of metasomatic Fe enrichment on previously depleted peridotites (e.g., Aulbach, 2012). While such effects have clearly been documented in cratonic peridotites (e.g., Boyd et al., 1997), if the olivine compositions are representative of the lithosphere as a whole (an extrapolation), then mass balance dictates that very large volumes of melt are required to shift olivine compositions from an average Mg number of 92.6 downward. This seems unlikely. A more favorable explanation for the tail to lower Mg number olivines might be that they represent less depleted restite compositions produced during polybaric mantle melting. Support for this notion comes from the observation that the secondary mode in olivine Mg number observed in some cratonic mantle sections (e.g., the

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

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NAC (Greenland) n = 78

Relative probability

Kaapvaal craton n = 171

Slave craton n = 85

88

89

90

91

Abyssal

92

OIB

93

94

95

96

Papua ophiolite

Relative probability

Iwanaidake n = 93

Kaapvaal craton n = 171 Forearcs n = 74 All cratons n = 493

88

89

90

91 92 93 Olivine Mg number

94

95

96

Figure 3 Probability density plots of olivine Mg numbers for peridotites from the Kaapvaal craton, the Slave craton, and the North Atlantic craton (W and E Greenland). Only data from low-temperature peridotites are plotted. A constant bandwidth of 0.2 is used for all olivine Mg numbers. Number of analyses plotted given as ‘n’ in curve labels. Shaded oblongs at the top of the diagram show the averages of the OIB (Canary Islands only), the abyssal peridotites, and the Papuan ophiolite belt (data and sources from Pearson and Wittig (2008) and Simon et al. (2008)).

Slave craton; Heaman and Pearson, 2010) occurs at a petrologically significant interval, that is, the cpx-out point at between 25 and 30% melting of peridotite (e.g., Afonso and Schutt, 2012; Herzberg, 2004; see section on melt depletion in the following text). In addition to variations in olivine composition resulting from melt depletion, studies of modal abundance variation in orogenic peridotite massifs have shown clear evidence of large changes in modal olivine abundance and bulk-rock (and hence olivine) Mg number that cannot be explained by melting and must be related to melt–rock infiltration and reaction (Bodinier and Godard, 2003). The refertilization of melt residues involving the consumption of olivine and melt and the production of opx and cpx can decrease, maintain, or increase bulk Mg number. This depends on whether the melt-induced

refertilization process is reaction dominated (mass ratio of crystals formed to melt introduced close to 1) or percolation dominated (mass ratio of crystals formed to melt infiltrated much less than 1), as shown in Figure 4. In extreme examples of melt–peridotite interaction during purely percolative flow, with negligible reaction (R values, where R is the mass ratio of crystals formed to melt infiltrated, of
0.1; Figure 4), olivine mass is effectively conserved. Thus, bulk Mg number and hence olivine Mg number can be increased by up to 1% absolute. Maximum increases of Mg number likely to be caused by melt– rock reaction are typically 0.7% or less in orogenic peridotites (Bodinier and Godard, 2003). The potential to change olivine Mg number via the melt– peridotite interaction (Bodinier and Godard, 2003) led Bernstein et al. (2007) and Aulbach (2012) to suggest that

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

Olivine consuming

Olivine forming

92 Refertilization percolation dominated R = 0.1

91

80 R = 0.001

75

R = 0.2

Mg number

R = 0.5 R=

masscrystals =1 massmelt

R=1

90

tion d te rtiliza Refe domina n io t c rea

Ol rea ivine cti -for on m mo ing de ls

g

ltin

Me

73

89

Fertile mantle Melt Mg#70 Tok L–W

88 50

60

80 70 Olivine (wt%)

90

100

Figure 4 Theoretical models of depletion versus refertilization and melt–rock reaction trends for bulk-rock 100  Mg number versus modal olivine (wt%) in spinel-facies peridotites. Because bulk-rock and olivine Mg number track each other closely in olivine-rich peridotites, these trends are applicable, in a close sense, to olivine 100  Mg number versus modal (wt%) olivine plots. The fertile source has a starting 100  Mg number of 89.3. ‘Refertilization’ trends (blue/gray continuous lines) were calculated using the ‘Plate Model’ of Verniers et al. (1997) and represent crystallization of clinopyroxene (cpx) and orthopyroxene (opx) as olivine and melt are consumed. The trends emanate from a residue of 25% melt extraction at 1.5 GPa. Following Bodinier and Godard (2003), the R values represent the mass ratio of crystallized minerals to infiltrated melt. When R is
1, the interaction is ‘percolation dominated,’ whereas the more that R approaches 1, the more reaction and mineral crystallization dominate the process. The olivine-forming reaction trends (green dashed lines) involve crystallization of olivine at the expense of opx at nearly constant to decreasing melt mass. The initial peridotite composition for reaction begins at the residue of 20% melt extraction at 1.5 GPa. The trends shown are from melt compositions with variable 100 Mg numbers (shown). For further details the reader is referred to Bodinier and Godard (2003; their Figure 9). Blue squares are some of the metasomatized lherzolite–wherlite series from the Tok locality (Ionov et al., 2005b) that appear to have interacted with melt of relatively low Mg number (63), producing Fe-rich metasomatic rocks with Mg numbers as low as 84 (not plotted). Modified after Bodinier J-L and Godard M (see Chapter 3.4).

olivines trapped within diamonds, which have slightly higher Mg numbers (Figure 5), might be more reliable indicators of the level of depletion due to their preservation from the effects of Fe metasomatism. The claim that diamond-inclusion olivine compositions preserve unique compositions that reflect only melt depletion (Aulbach, 2012; Bernstein et al., 2007) does not stand up to closer statistical scrutiny. Xenolith studies of carbonatite-metasomatized peridotites from the Olmani locality, Tanzania, show highly magnesian olivines, with Fo94 olivines being present in both dunites and wherlites (Rudnick et al., 1993). A comparison (two-tailed Student’s t-test, unpaired data with unequal variance) between Mg numbers of the diamond-inclusion olivine database (Stachel and Harris, 2008) and olivines from the Olmani peridotites (Jones et al., 1983; Rudnick et al., 1993) shows that their means are indistinguishable (Figure 5), consistent with Olmani olivine Mg numbers being increased during interaction with Fe-poor magnesiocarbonatites. Hence, the elevated average Mg number, typical of diamond-inclusion olivines, cannot be uniquely ascribed to their depletion and lack of subsequent exposure to silicate melts. Moreover, the link between carbonatites and diamond genesis (e.g., Klein-BenDavid et al., 2007, 2010; Weiss et al., 2011) means that one might expect to observe similarities between inclusions within diamonds and shallower mantle affected by carbonatite metasomatism.

Clearly, the addition of all the above phases will have a significant effect on bulk-rock major element compositions and hence must be considered both in genetic models and in tracing the secular evolution of bulk compositions. For instance, while there is a clear transition in olivine Mg number between the Archean mantle of the Kaapvaal craton and the circumcratonic Proterozoic mantle beneath central Namibia (Figure 5; Boyd et al., 2004; Janney et al., 2010; Pearson et al., 2004), the average bulk CaO and Al2O3 contents are the same or lower for these otherwise contrasting mantle sections (Figure 6). This difference seems likely to be due to metasomatic effects, increasing bulk-rock Ca and Al more than the olivine Mg numbers in both the Archean and Proterozoic mantle columns (Boyd et al., 2004). This means that measures of the possible secular evolution of mantle depletion have to be chosen carefully and secondary metasomatic effects have to be well understood. Bulk-rock Ca and Al contents are not good choices for evaluating these trends (contrast Figures 5 and 6). Despite their simple mineralogy, cratonic peridotites are clearly 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 multistage history obscures much of their early history and it is thus no surprise that their origin has been so difficult to constrain.

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

96

261

Olivine Mg number n = 687

95 n = 387

n = 22

n = 92

94 n = 22

n = 92

Iwanadaike HD-2

Circumcratonic

93

92

91 n = 95

90

89

Diamond Carbonatite inclusions metasom. perids.

Cratonic spinel

Cratonic garnet

Forearc

Figure 5 Box and whisker plots of 100 Mg number for olivines from mantle samples from a variety of tectonic environments. Central line within box denotes median value. Outer limits of box define the upper quartile and lower quartile, respectively (i.e., 50% of the data lie within the box). Whiskers represent the upper and lower quartiles and outliers are plotted as open circles. Number of samples (n) given for each category. Diamond-inclusion database from Stachel and Harris (2008). Carbonatite-metasomatized peridotites are those from Tanzania, published by Rudnick et al. (1993) and Jones et al. (1983). Cratonic and forearc peridotite data sources summarized by Canil (2004) and Pearson and Wittig (2008). Circumcratonic peridotites are those kimberlite-hosted xenoliths from locations immediately surrounding the Kaapvaal craton (Figure 19) and are data from Pearson et al. (2004) and Janney et al. (2010). Iwanadaike HD-2 is the most depleted harzburgite–dunite mantle section from the Phanerozoic Iwanadaike ophiolitic peridotite (NE Japan) studied by Kubo (2002).

3.5

AL2O3 (wt%) 3.0 2.5 2.0 1.5 1.0 0.50 0.0

Cratonic all Kaapvaal Circ-cratonic

NAC

OIB perid

Ophiolites

Forearcs

Figure 6 Box and whisker plots of bulk-rock Al2O3 contents (wt%) for peridotites from a variety of tectonic settings. Kaapvaal craton and North Atlantic craton (NAC) peridotite xenoliths are compared to the database for all cratons (low-temperature peridotites only). Cratonic dataset from data sources summarized in Chapter 3.5 and Pearson and Wittig (2008) and additions from Smith et al. (2009), Ionov et al. (2010), and Maier et al. (2012). OIB dataset from Simon et al. (2008). Ophiolites are Phanerozoic ophiolites (Gruau et al., 1998; Hanghøj et al., 2010; Jaques and Chappell, 1980; Marchesi et al., 2006). Forearc data from sources given in Canil (2004).

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3.6.2.2 Metasomatism, Cooling, and Diffusion: Some Constraints from Nontraditional Stable Isotopes (Li, Mg, and Fe) All studies that have been made on the Li isotope composition of mantle peridotites have found significant variation in d7Li (Aulbach and Rudnick, 2009; Ionov and Seitz, 2008; Jeffcoate et al., 2007; Lundstrom et al., 2005; Rudnick and Ionov, 2007; Seitz et al., 2004; Teng et al., 2007). The mobility of Li due to low-T alteration in the crust means that mantle Li isotope studies clearly warrant a well-grounded understanding of element mobility and leaching behavior if the signal from secondary alteration is to be excluded. Seitz et al. (2004) show changes in d7Li of up to 4% after washing carefully handpicked cpx grains with 18.2 M-ohm water, emphasizing the relative ease of removing surficial contamination for Li. As such, all Li isotope studies should be conducted on ultrapure mineral separates that have been thoroughly washed with ultrapure water. Harsher leaching would require an understanding, for each type of sample, of the efficiency of removal of surficial versus lattice-held Li. Significant progress has been made in understanding the geochemistry of Li in the mantle through the study of lithospheric mantle xenoliths. As melting-induced isotopic fractionation is thought to be minimal (
1%; Jeffcoate et al., 2007), the dominant process controlling Li isotope variations in mantle minerals appears to be kinetic fractionation of the two isotopes during diffusion (e.g., Aulbach and Rudnick, 2009; Ionov and Seitz, 2008; Jeffcoate et al., 2007; Lundstrom et al., 2005; Parkinson et al., 2007; Rudnick and Ionov, 2007). This process induces large ranges of up to 40% in d7Li within pyroxenes (e.g., Jeffcoate et al., 2007), recording diffusion profiles that can only have persisted for a few years at mantle temperatures. Two mechanisms favored to generate diffusion-driven kinetic Li isotope fractionation in peridotite xenoliths are (1) diffusion from grain-boundary melt films to mantle minerals during entrainment in the host magma or very recent metasomatism and (2) subsolidus intermineral Li partitioning, which can produce both isotopically light and heavy Li through Li addition or loss, respectively (Aulbach and Rudnick, 2009). Recent Li-partitioning experiments between olivine and diopside over a wide range of temperatures have found that Li distribution between these two minerals is independent of temperature (Jakob et al., 2012). This observation favors opensystem behavior during either fluid infiltration or redistribution in response to the changing oxygen fugacity. Olivines may, in some cases, preserve a depletion signature in refractory samples characterized by low Li concentrations and light d7Li values. Because of the susceptibility of Li isotopes to fractionation during these processes, it is very difficult to confidently identify any primary mantle source effects, such as recycling signatures, using Li isotope compositions (Rudnick and Ionov, 2007). While Li element and isotopic zoning offer great potential for constraining xenolith ascent velocities, and will be a useful comparison to estimates provided by hydrogen zoning (e.g., Peslier and Luhr, 2006), there have been no detailed applications so far. In contrast to the spectrum of Li isotope compositions displayed by minerals in mantle peridotites, magnesium

isotope studies have shown much more restricted compositions. The small range in d26Mg data in mantle-derived rocks indicates that partial melting does not fractionate Mg isotope ratios significantly beyond the analytical uncertainty (Teng et al., 2007, 2010; Wiechert and Halliday, 2007). The larger, apparent intermineral fractionation of Mg isotopes indicated by early laser ablation Mg isotope measurements (Pearson et al., 2006) appears more likely related to unresolved analytical issues (Bourdon et al., 2010; Handler et al., 2009; Norman et al., 2006). In high-precision studies, cpx–olivine fractionations are typically small (0–0.15%). Recently, Liu et al. (2011b) found that both silicate and silicate–oxide intermineral Mg isotope fractionations correlated with equilibration temperature, implying equilibrium fractionation controlled by MgO bond strengths, with stronger bonds favoring heavier isotopes. These authors suggest the possibility of a spinel–olivine Mg isotope thermometer. Previous measurements and theoretical calculations of Mg isotope partitioning by Young et al. (2009) support the idea that the differences between spinel and olivine may be due to temperature-related equilibrium fractionations and they also proposed a spinel– olivine thermometer based on Mg isotopic partitioning. However, the experiments of Jakob et al. (2012) appear to rule out temperature-dependent fractionation as an explanation of significant intrasilicate Mg–Li isotope fractionations and other explanations must be sought. Because the diffusion of Mg in olivine and pyroxene is relatively slow compared with Li (e.g., Parkinson et al., 2007), Mg isotopes appear much less disturbed than Li isotopes in mantle peridotites. High-precision solution-based measurements of mantle olivine reveal small but significant differences between the Earth’s mantle, as represented by peridotite xenoliths, and other planetary materials, including chondrites (e.g., Handler et al., 2009; Wiechert and Halliday, 2007; Young et al., 2009). The magnitude of the differences in 25 Mg/24Mg between Earth and chondrites differs between studies. Further work is warranted to resolve these discrepancies before their significance can be understood in the context of planetary formation and differentiation. Young et al. (2009) propose that some of the 25Mg/24Mg ‘excess’ observed in terrestrial rocks can be explained by the same process that accounts for the 29Si/28Si ‘excess’ relative to chondrites. This process is rejected as being core formation for Mg isotopes by Young et al. (2009) and these authors infer that the Mg and Si isotope fractionations are caused by the same, as yet unknown, process. The Fe isotope compositions of mantle peridotite xenoliths have proven to be relatively homogeneous (Figure 7), with generally (0–0.2%) small intermineral fractionations (e.g., Beard and Johnson, 2004; Williams et al., 2005). Spinel is an exception and can be up to 1.5% heavier in d57Fe, with variations that appear to be related to oxygen fugacity (Williams et al., 2004). Both oxygen fugacity and mantle metasomatism via melt infiltration appear to be the main causes of significant silicate intramineral Fe isotope fractionations (Weyer and Ionov, 2007; Williams et al., 2005). Zhao et al. (2012) have recently recorded variations of d57Fe values of over 1.3% among silicates from peridotite xenoliths from China. These authors found that olivine was systematically lighter than other silicates (Figure 7), with some samples having olivines with

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

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Figure 7 Summary of literature of Fe isotope data (d57Fe, per mil relative to IRMM-14) for peridotite minerals (ranges indicated for olivine, orthopyroxene and clinopyroxene), summarized in Williams et al. (2009) and Zhao et al. (2012) and compared to new data from Beiyan (eastern North China craton (NCC); Zhao et al., 2012; see Figure 20 for location) and Sanyitang (western NCC; see Figure 20 for location). The Sanyitang and Beiyan locations were highlighted for their unusually large variations for single localities. Uncertainties on individual data points are 2 standard deviations.

d57Fe as low as 1.0% (vs. IRMM-14). The light Fe isotopic composition of olivines in these samples correlated positively with CaO and Rb, supporting fractionation induced via extreme metasomatic enrichment and producing olivines with Mg numbers as low as 86. There is evidence for small but significant variations in Fe isotopic composition resulting from melt depletion in peridotites (e.g., Weyer and Ionov, 2007; Williams et al., 2005). Some of the total variation was ascribed to metasomatism and resulting Fe enrichment, while the Fe isotope variation that could confidently be attributed to melt extraction was small compared with the overall range seen by Weyer and Ionov (2007). As such, their proposed isotopic fractionation factor – lnamantle–melt of 0.1–0.3% – is less than the original estimate by Williams et al. (2005). Whereas the isotope fractionation effects of partial melting in peridotite systems seem small, Williams et al. (2009) have proposed that a positive linear relationship between d18O and d57Fe in a suite of South African eclogite xenoliths is predominantly related to isotopic fractionation via disequilibrium partial melting, although they could not rule out modification by melt percolation processes. Hence, it seems that before Fe

isotopes can be used to provide definitive evidence of mantle processes, further detailed studies are required to better understand the effects of the various styles of metasomatic activity on Fe isotope variations in both eclogites and peridotites.

3.6.2.3

Metasomatic Effects of Diamond-Forming Fluids

There is abundant major and trace element evidence that documents the various agents of metasomatism that have affected continental lithosphere through time (e.g., Menzies and Hawkesworth, 1987). Most recently, the study of solid and fluid inclusions within diamonds has thrown new light, both on the variety of deeply derived metasomatic fluids that affect cratonic mantle and on how such fluids might influence the lithospheric compositions estimated from xenoliths and diamond inclusions. Studies by Stachel and coworkers have carefully documented the major and trace element variability in silicate inclusions within numerous suites of diamonds from Africa and Canada (Banas et al., 2007; Donnelly et al., 2007; Stachel and Harris, 2008; Stachel et al., 2003, 2004; Tappert et al., 2005). Such studies, along with the comparison between garnets included in diamonds and garnets from the peridotite

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

hosting the diamonds in xenoliths (Creighton et al., 2008), clearly document differences between garnets included within diamonds and the host-rock garnets. The garnets included in diamonds are characterized by being Ca-poor compared with the host-rock garnets, whereas the Cr contents were found to be similar by Creighton et al. (2008). This indicates a style of metasomatism that introduces Ca without significantly changing bulk-rock Cr or Cr/Al ratio and is analogous to the xenocryst and xenolith garnet zonation trends documented in detail by Schulze (1995) and Burgess and Harte (2004), respectively. In the case of Diavik diamonds and diamondiferous peridotites, sinusoidal REE patterns are present in both diamond inclusion and host garnets (Creighton et al., 2008; Donnelly et al., 2007). This indicates that multiple metasomatic activities affected both groups of garnets in a similar way and points to the minimal influence of later silicate-melt metasomatism in the case of the host peridotitic diamonds. Studies of other diamond inclusions and xenoliths (Klein-BenDavid and Pearson, 2009; Shimizu et al., 1997; Simon et al., 2003; Stachel and Harris, 1997; Stachel et al., 1998) show a significantly greater variation in xenoliths, from rare sigmoidal to ‘classic’ equilibrated garnet REE patterns. Such compositional differences between diamond inclusions and host rocks or other mantle xenoliths are generally interpreted as reflecting the action of much later metasomatic activity on the xenoliths, which are essentially open systems. However, in the specific case of diamondiferous xenoliths, where silicate inclusions in the host diamonds have also been studied (Creighton et al., 2008), the major compositional difference seems to be the elevation of Ca in the host-rock garnets. In this case, an alternative model is possible wherein the enriched Ca is a reflection of continued metasomatism by diamond-forming fluids that formed first the diamonds and then continued to interact with the host silicate assemblage. This may also explain the more Ca-rich compositions of the peridotite garnet diamond inclusions at Diavik and the Lac de Gras region in general compared to those at Kimberley (Stachel et al., 2003, 2004; Tappert et al., 2005). In this context, Creighton et al. (2008) have noted the Ca-rich nature of fluids trapped within some Diavik diamonds (Klein-BenDavid et al., 2004). Such fluids are also greatly enriched in Na and K (Klein-BenDavid et al., 2007; Navon et al., 1988), but these elements are not partitioned into garnet. The significant enrichment of lherzolitic diopside inclusions in diamonds in K2O (up to 1.68 wt%, median 0.1%, and s ¼ 0.27%; data from Stachel and Harris, 2008) seems likely to reflect the effects of diamond-forming fluids (Harlow, 1997). In contrast, diopsides from nondiamondiferous lherzolites from equivalent depths in the lithosphere usually have K2O contents below electron probe routine detection limits. Harlow’s (1997) experiments suggest that a Cr-diopside with 1 wt% K2O would require an equilibrium C-rich fluid with between 15 and 28 wt % K2O. Another example of the imposition of unusual compositions on mantle silicates included in diamonds is the occurrence of distinctive Cr- and Cl-rich phlogopite inclusions within octahedral diamonds reported by Sobolev et al. (2009) and attributed to interaction with a Cl-rich diamond-forming fluid. The diamond-forming fluid became enriched in Cr due to fluid–rock interaction. This hypothesis is supported by the recent experimental finding of significant Cr solubility in

KCl-bearing water at 1000–1200  C and 4–6 GPa (KleinBenDavid et al., 2010). A genetic link has been made for the association between low-Ca, high-Cr (G10) garnets and enrichments from diamond-forming fluids, further implicating such fluids as effective metasomatic agents in the lower CLM. Both Pearson et al. (1995b) and Malkovets et al. (2007) have suggested fluidassisted reactions involving precursor chromite to form these garnets. The extreme Nd and Hf isotope compositions measured in these garnets by Klein-BenDavid et al. (2009) indicate a very complex and ancient metasomatic evolution of the precursor, but the isotopic compositions do not uniquely constrain the age of garnet formation. Hence, both ancient (Richardson et al., 1984) and recent (Malkovets et al., 2007; Pearson et al., 1995b; Shimizu and Sobolev, 1995) formation models have been proposed. Fluids within fluid-rich diamonds have complex and widely varying major element (Klein-BenDavid et al., 2004, 2007), trace element (Klein-BenDavid et al., 2010; McNeill et al., 2009; Rege et al., 2005, 2008; Resano et al., 2003; Schrauder et al., 1996; Tomlinson et al., 2005, 2006, 2009; Weiss et al., 2008, 2011; Zedgenizov et al., 2007), and radiogenic isotope characteristics (Akagi and Masuda, 1988; Klein-BenDavid et al., 2010) that are likely caused by many processes. Some fluid end-member compositions have similarities with carbonatites (Klein-BenDavid et al., 2007; Weiss et al., 2008, 2011) but the diamond fluids have greatly enriched alkalis in comparison. The wide variations in Sr, Nd, and Pb isotopic compositions of diamondforming fluids highlight the importance of mixing of fluids from at least two different sources – both ancient lithospheric and convecting mantle – in their genesis (KleinBenDavid et al., 2010).

3.6.3 Primary Compositions of Cratonic Peridotites and Their Melting Environment The metasomatic addition of one or all of (1) opx, (2) cpx, and (3) garnet to cratonic peridotite makes bulk-rock compositions of elements, such as Al and Ca, very unsuitable for use in estimating primary bulk compositions of cratonic residues and hence for reconstruction of the processes that shaped their origin (Pearson and Wittig, 2008; Wittig et al., 2008a,b). For this reason, using these elements as a base for establishing genetic models is unwise. For instance, as noted by Janney et al. (2010), the mean Al2O3 values for peridotites from the Kaapvaal craton (1.28 wt%) and circumcratonic (Proterozoic) peridotites surrounding the Kaapvaal craton (1.26 wt%) overlap (Figure 6); their interquartile ranges are very similar and their means cannot be distinguished (tval ¼ 0.2989; tprob ¼ 0.7654; and degrees of freedom (dof) ¼ 147) despite clearly lower average olivine Mg numbers, whose means are highly unlikely to be drawn from the same population (tprob < 0.0001). Moreover, the mean Al2O3 value of all cratonic peridotites (1.048) is even slightly more elevated than that for peridotite xenoliths from modern ocean island basalt (OIB) (0.912; Figure 6; excluding Hawaiian and Tahitian peridotites), as noted by Simon et al. (2008), such that it is not safe to say that cratonic peridotites

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

represent the most depleted of mantle samples using this depletion index. In contrast to bulk-rock Al2O3 contents, olivine Mg numbers are distinctive between peridotite xenoliths from OIB and cratonic peridotites and between Proterozoic peridotites and Archean peridotites (Pearson and Wittig, 2008; Pearson et al., 1994; Figure 5) such that there is very little likelihood that their means could be drawn from the same population using any statistical test. Hence, despite possible alteration due to melt–rock reaction, olivine Mg number has been proposed as perhaps the most reliable tracer of peridotite evolution (Bernstein et al., 2007; Pearson and Wittig, 2008). Some authors have developed more sophisticated approaches using ‘internally consistent’ criteria for multiple parameters. These include Al2O3 contents, along with FeO–MgO relations (Herzberg, 2004; Herzberg and Rudnick, 2012), or heavy screening using compositional cutoffs for Al2O3. However, the authors believe that the inclusion of Al2O3 is dangerous because of its relative insensitivity to melting pressures, its apparent ubiquitous disturbance, and uncertainties in the primitive mantle Al2O3 value (e.g., Wittig et al., 2008a,b). While olivine has also been added to some lithospheric peridotites or the existing olivine compositions have been altered (Boyd et al., 1983; Rehfeldt et al., 2007, 2008), the resulting olivine is usually very distinctive in having olivine Mg numbers below 90. Nonetheless, in Section 3.6.2, the possibilities of changing bulk-rock and hence olivine Mg numbers during peridotite–melt infiltration and reaction, with the clear possibility that Mg numbers can both increase and decrease, are documented. An examination of such effects in cratonic peridotite bulk composition and modal olivine variation (Figure 4) against the theoretical models of Bodinier and Godard (2003) shows that melt infiltration, whether olivine consuming or producing, is likely to be dominated by the relatively high mass ratios of crystals formed (opx or olivine) to melt infiltrated, that is, reaction effects are greater than percolation effects. This is clear from the near-constant olivine Mg number while modal olivine varies over 50 wt%. The theoretical model provides a reasonable explanation for the very large range in modal olivine contents of cratonic peridotites (Bernstein et al., 2007; Boyd, 1989; Boyd et al., 1997; Pearson and Wittig, 2008) at an almost constant Mg number. Furthermore, this trend suggests that the dominant influence on bulk-rock and hence olivine Mg number in cratonic peridotites is not melt–rock reaction but partial melting. This conclusion contrasts with the variation seen in some orogenic peridotites but might be expected because the extent of melting in cratonic peridotites is more than a factor of 2 greater than in most modern melting environments (Bernstein et al., 2007; Herzberg and Rudnick, 2012; Pearson and Wittig, 2008). Olivine Mg number is used rather than bulk rock because of common grain-boundary enrichments in FeO in many peridotites. The remaining complication with olivine Mg number is its temperature sensitivity and compositional sensitivity to exchange with garnet (O’Neill and Wood, 1979) and spinel (Fabries, 1979). While these effects can be significant at extremes of temperature and must be borne in mind when comparing mean Mg numbers from rocks of grossly different equilibration conditions, the effects are predictable and systematic (O’Neill and Wood, 1979).

265

3.6.3.1 Composition of Cratonic Mantle: Some Comparisons Some workers have used olivine compositional data from xenocrysts dispersed within kimberlite to assemble large databases of cratonic mantle compositions (e.g., Gaul et al., 2001; Griffin et al., 2009). While undoubtedly accessing much data, there are several complications to this approach that lead one to prefer using xenolith olivine compositions. Among the problems with the xenocryst approach are the following: (1) the difficulty of knowing whether olivines are being derived from a few large xenoliths, thereby magnifying the inherent problem with xenolith sampling to the extreme, (2) greater uncertainties in depth estimation due to numerous assumptions made in the estimate of a depth of origin, (3) no means of cross-checking the depth estimate using other thermobarometers, (4) no procedure for checking compositional equilibrium, and (5) spinel-facies peridotites have been excluded from xenocryst compilations or have not been specifically identified as such. In contrast, using peridotite xenoliths allows equilibrium to be checked for and equilibration conditions can often be multiply constrained. In this analysis, only the ‘low-T granular’ peridotites of Boyd et al. (1997) are considered due to the complex metasomatic effects recognized in the ‘high-T deformed’ category of peridotites (e.g., Smith and Boyd, 1992). An issue that affects both the xenoliths and xenocrysts is the potential effect of metasomatism. Bernstein et al. (2007) and Aulbach (2012) propose that olivine included within diamonds is a more reliable indicator of lithospheric mantle composition and hence origin. When simple statistical tests are made, support is given to the contention of Bernstein et al. (2007) and Aulbach (2012) that olivine inclusions in diamond are compositionally distinct from olivines within cratonic peridotites in general (probability of the means being the same is very low: tprob ¼ <0.0001 and dof ¼ 1215). While a difference is clear, as shown in Section 3.6.2, it seems likely that this difference could be due to the effects of carbonatite metasomatism associated with diamond-forming fluids. Moreover, the diamond-inclusion olivine dataset cannot be claimed to be unique. A two-way Student’s t-test comparison of the olivine database of diamond inclusions from Stachel and Harris (2008) with olivines from cratonic spinel-facies peridotites (tprob ¼ 0.0955; dof ¼ 171) shows that one cannot confidently reject the null hypothesis that the means are the same. Hence, the authors take the view that olivines included in diamonds are influenced by other processes rather than just melt extraction and that their compositions are not, in fact, uniquely depleted. Supporting evidence for the extensive interaction of harzburgitic silicate inclusions in diamond with enriched melts comes from the extreme LREE enrichment observed in harzburgitic diamond inclusions (see Stachel and Harris, 2008, for review). From this combined evidence, the authors are cautious about building craton genesis models on such a restricted and unusual sample set. Furthermore, they note that where particularly pristine bulk compositions can be found in cratonic residue suites, such as the low-opx harzburgite sample F05JM9 (0.18% Al2O3 and 0.38% CaO; Gibson et al., 2008), their olivine Mg number (92.4) is identical to the mean cratonic olivine Mg number of 92.6 (Bernstein et al.,

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

2007; Pearson and Wittig, 2008), indicating that the xenolith olivine population is likely to be fairly reliable as an indicator of protolith depletion. Once diamond-inclusion olivines are excluded from consideration, the available data show some fundamental differences between olivine populations from different lithospheric depths (Figure 8). The subtleties of these variations are discussed in more detail below, but the basic differences must be highlighted before comparing cratonic lithosphere with other lithosphere compositions. The most obvious subdivision to examine in the cratonic lithosphere is that between the spinel facies (e.g., Boyd et al., 2004) and garnet-facies peridotites. When Mg number of olivine for the two different facies of peridotites are compared (Figure 5) across all cratons and within individual cratons, there is very little likelihood that the means for the two facies can be drawn from the same population (e.g., for all cratons, tprob ¼ < 0.0001 and dof ¼ 142), with average olivine Mg number being significantly higher in spinelfacies peridotites. The difference in means (0.59 Mg number units) might be related to temperature-related partitioning effects; however, the spinel-facies peridotites, in general, do not have the pronounced tail to lower Mg numbers in their olivines and, in some cases, for example the Kaapvaal craton, extend to unusually high values. The generally more elevated olivine Mg numbers in spinel-facies peridotites might be due to carbonatite metasomatic effects. Alternatively, the spinel-facies cratonic peridotites are either a significantly more depleted endmember of a suite of peridotites related by more extensive polybaric melting because they have upwelled further than the residues beneath them or they have formed from a different process than garnet-facies peridotites. Lastly, the garnet peridotites could have, on average, low olivine Mg numbers because they have experienced more melt refertilization than the spinelfacies rocks above them. How uniquely depleted are cratonic peridotites compared to melting residues produced at other times in Earth history? Although the sample set is relatively small, spinel-facies cratonic peridotites cannot be confidently distinguished from the most depleted Phanerozoic mantle samples, that is, very high olivine Mg numbers (Figure 5) in the HD-2 section of the Iwanadaike peridotite massif (Kubo, 2002; tprob ¼ 0.317; dof ¼ 68). Garnetfacies cratonic peridotites, in particular, appear significantly less depleted than this Phanerozoic suite of recent melt residues (Figure 5). The Iwanadaike peridotites are products of hydrous melting in a forearc environment (Kubo, 2002), and hence, one possibility is that the shallowest mantle beneath cratons is also formed in this environment (i.e., a subduction-zone melting environment), whereas the deeper cratonic mantle is formed in a separate environment. This will be considered further in the next section.

3.6.3.2 Composition: Depth Variations and the Origin of Cratonic Mantle Assuming that Mg number of olivine offers the most reliable but still imperfect tracer of melting processes, the variation of this parameter with depth in the lithosphere can be compared with experimental parameterizations of mantle-melting products under different conditions (Figures 8 and 9; Herzberg, 2004; Herzberg and Rudnick, 2012; Herzberg et al., 2010). In contrast

to the Herzberg approach, which examines all of the data without the context of depth, the authors reason that, in both endmember models for generating depleted peridotite-dominated cratonic keels, melting is polybaric and that, in any single tectonic setting, the residues that have experienced the greatest degree of isentropic upwelling will have experienced the most melt depletion. This should lead to characteristic variations of composition with depth in a column of residual mantle. Hence, depth versus depletion relationships are viewed as an important part of resolving the origin of cratonic peridotites. Figure 9 shows compilations of olivine Mg numbers versus equilibration depths for the cratonic and circumcratonic peridotite xenolith locations that have the most data points. Also plotted are the melting trends for olivine Mg number calculated using the polybaric fractional-melting parameterizations of Herzberg (2004). These melting trends track the progressive compositional change with depth of residues accumulating beneath the crust due to melting in tectonic environments equivalent to oceanic ridges, hot ridges (where Archean mantle potential temperatures are assumed), plumes, and ‘hot plumes.’ The shapes of the trends are perhaps most important, along with the range of total Mg numbers and the maxima generated. The relative relationships between residues are more significant than the absolute depths because, although melting at ridge environments brings residues to shallower depths than melting in plumes, tectonic thickening or subduction accretion of such residual mantle will emplace this column of residues beneath the much thicker proto-continental crust. Also, there is the potential for the telescoping of the residue trends if cratonic keels form via simple compression and thickening of preexisting lithosphere, as proposed by Jordan (1978) and McKenzie et al. (2005). Hence, the pressure of melting is unlikely to be equivalent to the pressure/depth of equilibration. For this reason, all the calculated melting trends in Figures 8 and 9 are terminated beneath 40-km-thick crust, typical of the cratonic mantle that they now underpin. While the data are scattered, several conclusions can be drawn from their comparison with the theoretical depletion trends. Firstly, melting at a ridge environment with similar thermal properties to modern mantle is not capable of generating residues with high enough Mg numbers, in agreement with the analysis of Boyd (1989), Herzberg (2004), Carlson et al. (2005), Bernstein et al. (2007), Simon et al. (2007), Aulbach et al. (2007), Pearson and Wittig (2008), and Griffin et al. (2009). Secondly, of the several hundred samples, only a few spinel-facies peridotites from the Kaapvaal craton have compositions depleted enough to have been produced from a ‘hot plume,’ such as might have existed in the Archean (Herzberg, 2004; Korenaga, 2008). Even the majority of diamond inclusions are much less depleted than the very elevated Mg numbers (94 and above) produced in over 100 km thickness of lithosphere that should be accreted beneath a craton in the ‘hot plume’ environment. Hence, the predicted very thick column of extremely depleted residues (Figure 8) in this melting environment, where olivine Mg numbers of >92 extend to at least 250 km depth, is not observed or sampled beneath any craton. The authors therefore agree with Herzberg and Rudnick (2012) that the formation of cratonic lithosphere in Archean plumes seems unlikely.

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Figure 8 Olivine Mg number versus depth of equilibration plots for peridotite xenoliths from the Kaapvaal craton (the 1200 Ma premier location is plotted separately), Slave craton (Lac de Gras region; MacKenzie and Canil, 1999; Menzies et al., 2003; Pearson et al., 1999), north Slave craton (Jericho; Kopylova et al., 1999), Siberia (Boyd et al., 1997; Ionov et al., 2010), NAC (Bernstein et al., 2006; Hanghøj et al., 2001; Sand et al., 2009), and Namibian craton (Gibeon pipes and Farm Louwrensia; Boyd et al., 2004 and references therein). Equilibration depths for spinel-facies peridotites were assumed to be 50 km. Spinel– garnet transition marked by thick wavy line. Thin wavy line marks crust–mantle boundary at 40 km based on James et al. (2001) and Nguuri et al. (2001). Straight solid line shows the isochemical temperature dependence of olivine Mg number with increasing depth (O’Neill and Wood, 1979). Thick, gray irregular lines are the depth versus olivine Mg number trends calculated from olivine xenocryst data by Gaul et al. (2001). Upper-left Kaapvaal craton panel has melt depletion trends for Mg number in olivine during polybaric fractional melting using the parameterizations of Herzberg (2004) and Herzberg and Rudnick (2012) and for thermal structures at Phanerozoic mid-ocean ridges (MOR), Archean MOR (hot MOR), Phanerozoic plumes (plumes), and Archean plumes (hot plumes).

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

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Figure 9 Average olivine Mg number (1 standard deviation bars) versus average depth of equilibration (20 km depth bins) for peridotite xenoliths from the Kaapvaal craton, Slave craton, and North Atlantic craton. Melt depletion trends for Mg number in olivine during polybaric fractional melting using the parameterizations of Herzberg (2004) and Herzberg and Rudnick (2012) and for thermal structures at Phanerozoic mid-ocean ridges (MOR), Archean MOR (hot MOR), Phanerozoic plumes (plumes) and Archean plumes (hot plumes) have been moved upward (for plume trends) and downward (for MOR trends) to underpin a crust of 40 km thickness to account for the potential presence of such residual lithospheric mantle columns beneath a typical cratonic crust, assuming no additional melting in the case of the plume trends.

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

This conclusion agrees with other analyses of basalt-residue compositions and geological factors (Rollinson, 2010). Most of the peridotite xenolith olivine data from the Kaapvaal craton, Slave craton, Siberian craton, and NAC (Figure 8) yield compositional-depth trends in which the compositional ranges and residue column thicknesses are best approximated by the ‘hot ridge’ residue parameterization. This is especially the case if the trend is moved downward to account for the presence of a melting residue column in the modern Earth, that is, beneath 40 km of cratonic continental crust. This finding agrees with that of Herzberg and Rudnick (2012), who used other bulk-rock compositional parameters, such as Al2O3, from a database filtered for consistency between melt-extraction indicators. The trends for Siberian peridotites, those from Jericho (Slave craton), and the NAC are particularly clear. They are consistent with the vertical extent and compositional range of the modeled ‘hot ridge’ residues, especially the slope of the tail to lower Mg numbers (Figure 9). The trend for the Neoarchean Somerset Island peridotites of the Rae craton (Irvine et al., 2003), while scattered, is also bound by the hot-ridge trend. This conclusion, for these cratons, is amplified if the compositional data is averaged in blocks of 20 km depth (Figure 9). The depth-averaged olivine Mg number trend for the Slave craton (Figure 9) fits the steep, lower-pressure polybaric melting trend of a hot ridge in its upper part and then continues downward at higher Mg number than expected for a residue

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column produced from single-stage melting. Instead, there may be two different lithosphere sections stacked under the central Slave province, as suggested on the basis of garnet xenocryst studies (Griffin et al., 1999; Grutter et al., 1999, 2004) and xenolith studies (Aulbach et al., 2007; Heaman and Pearson, 2010; Kopylova and Caro, 2004). In some models, the upper layer may have been produced via shallow melting, while the lower layer was generated in a plume environment, either in the Archean (Aulbach et al., 2007, 2009) or in Mesoproterozoic times (Griffin et al., 1999). An examination of Figure 10 shows that formation of the lower portion of the Slave lithosphere in an Archean plume is unlikely due to the lack of sufficiently depleted compositions and the inappropriate slope of the compositional trend. The lower portion may fit part of the ‘Phanerozoic’ plume residue trend but equally possible is that the lower layer represents another section of ‘hot ridge’ residual melting, formed in Archean times or later, that has been underthrust beneath a preexisting lithosphere. This option will be explored further when the age of cratonic lithosphere is considered. The Mg number versus depth trend for the Kaapvaal craton is the most scattered of any of the localities, with some of the data, especially beyond 150 km depth, extending toward the ‘plume’ trend and one outlier even extending toward the ‘hot plume’ melting trend. Hence, using the extremes of the data, a plume origin could be advocated (e.g., Aulbach, 2012; Aulbach et al., 2007; Boyd, 1989; Griffin et al., 2003, 2009;

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Yb (ppm) Figure 10 Whole-rock Lu (ppm) versus Yb (ppm) elemental systematics in cratonic peridotite suites. Curves shown for polybaric fractional melting beginning at 2, 3, and 7 GPa and for isobaric melting at 7 GPa following melting models described by Wittig et al. (2008a,b) based on Walter (1998) experiments and Herzberg (2004) parameterization. High-pressure (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. While garnet exhaustions moved into the ‘spinel’ field at pressures above 4 GPa, garnet remains as a stable residual phase to melting extents in excess of 60% (Walter, 1998). Data sources indicated on diagram (see Wittig et al. (2008a,b) for more detailed discussion). Modified after Pearson DG and Wittig N (2008) Formation of Archean continental lithosphere and its diamonds: The root of the problem. Journal of the Geological Society 165: 895–914.

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Pearson et al., 1995a). However, a different picture emerges if the data are averaged in 20 km depth intervals (Figure 9). In this case, a ‘kinked’ trend is apparent, with shallow, highly depleted residues in the spinel-facies field descending to lessdepleted residues that are displaced again to very depleted compositions between 130 and 150 km. Below 150 km, compositions again tail off to lower Mg numbers. The transition to more depleted compositions at 130–150 km in the Kaapvaal mantle lithosphere coincides with the depth interval for harzburgites containing low-Ca, high-Cr garnets (Boyd et al., 1993). A subducted spinel-facies mantle protolith has been argued for the low-Ca harzburgites and dunites on the basis of their garnet compositions, which can only be stabilized at extreme bulk-rock Cr/Al (Bulatov et al., 1991; Stachel et al., 1998). Hence, in contrast to the conclusions of Griffin et al. (2009) and Aulbach et al. (2009), even for the lithospheric keel beneath the Kaapvaal craton, a plume origin is far from clear and some evidence may indicate the subduction stacking of two mantle residue columns produced at a hot ridge. This model is strongly supported by peridotites from the Finsch kimberlite in the far western part of the west block of the Kaapvaal craton (Gibson et al., 2008). The location of the Finsch mantle column was distal from the effects of the subduction-zone metasomatism that appears to have caused much of the opx enrichment in the Kaapvaal craton (e.g., Bell et al., 2005; Kelemen et al., 1998; Simon et al., 2007). The preservation of some more pristine peridotite bulk compositions at Finsch allows better comparison with experimental melting models and supports depletion occurring between 4.5 and 1.5 GPa in a hot-ridge setting (Gibson et al., 2008).

3.6.3.3

Trace Element Evidence

Tainton and McKenzie (1991), Kelemen et al. (1998), and Canil (2004) have pointed out that the low concentration of moderately incompatible elements, such as Sc and Yb, in cratonic peridotite xenoliths is consistent with melting in the absence of garnet. Hence, these authors argued for an origin of cratonic lithosphere via the subduction accretion of melting residues produced at relatively shallow levels, that is, an Archean mid-ocean ridge (MOR). Aulbach (2012) has recently argued that garnet will disappear as a residual phase during melting, and hence, the concentration of Sc, Yb, and heavy rare earth elements (HREE) in cratonic residues is merely a function of melting to garnet-out at high pressure. However, this view of mantle melting is at odds with the experimental evidence (Walter, 1998) that shows clearly that if melting is induced in the garnet stability field, garnet remains a stable residual phase at 5 GPa and above, up to melting extents of 70%. Hence, even highly depleted melt residues, which experienced melt extraction in the garnet facies, should retain the signature of residual garnet. Additional studies by Simon et al. (2007) and Wittig et al. (2008a,b) and an extensive compilation of cratonic peridotite data by Pearson and Wittig (2008) show that low HREE concentrations and, in particular, fractionation of Lu from Yb (Figure 10) can only be generated by extensive melting in the spinel peridotite stability field. This feature is a ubiquitous signature in cratonic peridotites. Hence, the trace element database for cratonic peridotites is consistent with the analysis

from major elements, indicating that melting in a hot-ridge environment is most likely. In the elevated geothermal conditions of the Archean, the spinel to garnet phase transition may have been displaced to greater depths. Modeling by Klemme (2004) indicates that for a mantle that is 300  C hotter, with a Cr/Cr þ Al ratio of  0.2, the garnet-in reaction is displaced to  0.7 GPa higher pressure so that a larger residual column would be produced bearing the signature of spinel-facies melting. With increased melt extraction and more depleted bulk compositions (Cr/Crþ Al > 0.7), the garnet-in reaction is displaced to pressure >2.4 GPa at any temperature above 1100  C. Hence, a hotter Archean mantle producing more depleted residues could have significantly expanded the depth of residual peridotite with a spinel-facies melting signature to 100 km or more. While it is tempting to conclude from this analysis that the absence of a garnet signature in cratonic peridotite trace element systematics may be due to a spinel–garnet transition displaced to a greater depth, this effect should be seen in the experiments of Walter (1998). In these natural system experiments, the large melt-fraction experiments require elevated temperatures and create very depleted compositions, yet garnet remains a stable residual phase to pressures of 7 GPa and melt fractions >50%. Hence, there is a dichotomy between experiments on natural systems and thermodynamic modeling. Further work is required on the behavior of the spinel–garnet transition during large degrees of melting at elevated temperatures.

3.6.3.4 The Role of Subduction: Was the Melting Environment Hydrous? The association between anomalously depleted cratonic mantle and high-MgO Archean magmas, which have been interpreted by some workers as melts from Archean subduction zones (e.g., Isua boninites (Polat et al., 2002) or some Kaapvaal komatiites (Parman and Grove, 2004; Wilson et al., 2003)), has led to models where cratonic peridotites were either generated in whole (Parman et al., 2004) or in part (Carlson et al., 2005; Pearson and Wittig, 2008; Simon et al., 2007; Wittig et al., 2008a,b) in subduction zones. The generation of cratonic mantle in a convergent margin setting has the added attraction of also explaining the opx enrichment of some lithosphere sections (e.g., Kaapvaal, Boyd, 1989; Siberia, Boyd et al., 1997) and links in with the apparent subduction-related metasomatic enrichment of cratonic lithosphere in sulfides carrying radiogenic Os (Carlson et al., 2005; Griffin et al., 2003; Richardson et al., 2001; Westerlund et al., 2006). However, the enrichment of opx seems unique to the Kaapvaal lithosphere and to the docking of the ‘Kimberley’ and ‘Witwatersrand’ cratonic blocks (Bell et al., 2005; Schmitz et al., 2004; Simon et al., 2007). New, more extensive analyses of Siberian peridotites (Ionov et al., 2010) and the very opx-poor nature of the NAC mantle (Wittig et al., 2008a,b) indicate that cratonic mantle is not as opx-enriched as previously thought, and hence, making opx enrichment an intrinsic part of the genesis of cratonic lithosphere (e.g., Pearson and Wittig, 2008) is not necessary or desirable. While the most depleted mantle residues produced in modern (Phanerozoic) Earth occur in subduction-zone environments (Canil, 2004; Ishii et al., 1992; Kubo, 2002; Pearson et al., 2004), such residues are very scarce and seem unlikely to generate thick,

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

extensive lithospheres even in the Archean. Melting in the mantle wedge of subduction zones is limited to the zone where the wet mantle solidus is exceeded and so cannot explain the presence of an extensive, shallow residual mantle beneath the deeper, less-depleted residues, unless wedge residues are allowed to rise unimpeded to accumulate under the arc crust by buoyancy. Moreover, it is still unclear whether hydrous melting is any more capable of producing ultraresidual peridotites than via extensive melting in anhydrous conditions. Water lowers the solidus of mantle peridotite (Kushiro et al., 1968), but the incompatibility of water in peridotite systems (Hirschmann et al., 1999) means that at the high integrated melting extents operating to produce cratonic peridotites there will be little or no water remaining to influence melt productivity. Without a better understanding of the hydrous melting of depleted peridotite and of the dynamics of melting in the Archean subduction zones, it seems hard to do more than speculate on how such a tectonic environment might produce 200-km-thick residual mantle keels. A survey of all commonly measured chemical parameters in cratonic peridotites reveals that there is no distinctive signature in any known cratonic peridotite suite, which is a strong indication of genesis in a subduction zone. There is abundant data that suggest modification of cratonic mantle in convergent settings (e.g., Kelemen et al., 1998; Rudnick et al., 1994; Simon et al., 2007). Given this and the abundant and consistent evidence for residue genesis in a hot-ridge environment (Gibson et al., 2008; Herzberg, 2004; Herzberg and Rudnick, 2012; Rollinson, 2010), the most conservative tectonic interpretation for the genesis of cratonic mantle is at an Archean MOR, with subduction accretion occurring in a convergent setting, accompanied to varying degrees by subduction-related metasomatic modification. The ramifications for this conclusion in the context of models for craton genesis as a whole (crust and mantle) will be discussed after the timing of cratonic keel formation have been constrained.

3.6.4 Constraining the Timing of Lithosphere Formation 3.6.4.1 Information Derived from Lu–Hf Isotope Systematics The timing of continental lithosphere stabilization, that is, the moment of melt extraction from the convecting mantle and coupling of the refractory and buoyant residues to the continental crust, is paramount in understanding the geodynamic and geochemical evolution of Earth’s crust and mantle. The mineralogical and elemental compositions of peridotite residues directly correspond to the degree of melt extraction, which may be reconstructed from experimentally determined mineral partition coefficients, mineral equilibria, and modal melting reactions (Herzberg, 2004; Walter, 1998). The isotope systematics of Sr, Nd, and Pb in mantle silicates document the sensitivity of these elements to mantle metasomatism, which, except in rare examples (e.g., Deng and Macdougall, 1992; McCulloch, 1989), obliterates the initial trace element signature of melting (e.g., Menzies and Hawkesworth, 1987; Menzies and Murthy, 1980; Pearson and Nowell, 2002). In fact, these isotope systems may be used to acquire information

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on the timing of the most recent metasomatic events in continental mantle roots (e.g., Carlson et al., 2004; Schmidberger et al., 2007; Wittig et al., 2008b). By contrast, initial studies using the Lu–Hf isotope system in mantle rocks appeared to more faithfully record the chemical signature and timing of melt extraction (Salters and Zindler, 1995; Schmidberger et al., 2002; Wittig et al., 2006, 2007). After nearly a decade of Lu–Hf geochronology on constituent mantle minerals, the extent to which Lu–Hf isotope dating rivals the usefulness of whole-rock Re–Os isotope systematics in dating mantle melting and lithosphere formation is examined. Lithophile isotope data from peridotites is typically generated from leached, hand-picked minerals, chiefly garnet and cpx, although opx data are emerging (e.g., Chu et al., 2009). Mineral analyses are preferred because they avoid the measurement of intergranular contamination and alteration from percolating melts and metasomatic liquids. Lu and Hf concentrations in peridotites may serve as a first-order proxy for mantle-melting signatures in residual mantle and therefore as a suitable target for Lu–Hf dating depletion and hence lithosphere formation. Mantle melting produces Lu–Hf fractionations and, consequently, time-integrated Hf isotope compositions in cpx, which would be less radiogenic than those of garnet assuming the same time and degree of depletion. A large range of Lu–Hf in a suite of xenoliths would be particularly appealing because it will rapidly produce radiogenic Hf isotope signatures in the most depleted samples and facilitate precise and accurate dating of relatively recent events. This is advantageous considering the problems in obtaining Re–Os model ages in young mantle rocks generated by Os isotope heterogeneity in modern Earth. The issue of whether cpx Hf isotope analyses are an adequate representation of spinel peridotite Hf isotope evolution has been discussed recently by Stracke et al. (2011). The cpx is the major host for incompatible trace elements in spinel lherzolites (e.g., McDonough et al., 1992). Both opx and olivine can become significant hosts for moderately incompatible elements, such as Lu, particularly in cpx-poor peridotites. According to modeling performed by Stracke et al. (2011) in rocks where modal cpx abundances are greater than 3–5%, more than 70% of the total peridotite Hf and Nd are concentrated in the cpx, whereas a significant fraction of the moderately incompatible elements (e.g., up to 30% of the HREE) are hosted in opx and olivine. Partition coefficients for the HREE are significantly higher than for the middle REE and LREE in both opx and olivine, and so, these two phases are generally expected to have higher Lu/Hf and Sm/Nd ratios than cpx in these rocks. Modeling by Stracke et al. (2011) shows that, compared to the whole rocks, Lu/Hf ratios in cpx are between 25 and 70% lower (depending on cpx content), whereas Sm/Nd ratios in cpx are within about 6% of the whole-rock value for modal cpx abundances greater than 3% and within 20% for cpx abundances less than 3%. From this, it can be predicted that over long time periods, the time-integrated Hf and Nd isotopic ratios of both opx and olivine should be more radiogenic than cpx in spinel lherzolites. For Hf isotopes, 50–100 My is required to elapse before isotopic differences outside of analytical error can be measured, whereas the higher Sm/Nd ratios generate measureable differences in Nd isotopic ratios in less than 50 My.

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Because the Sm/Nd ratios in cpx track the whole rock more closely than Lu/Hf ratios such that during the evolution of residual spinel peridotite, the cpx Nd isotopic composition deviates less from the whole-rock composition than for Hf isotope ratios (e.g., Stracke et al., 2011). Despite the obvious interest in analyzing opx for Hf isotopes, the low Hf abundance of this mineral has precluded accurate measurements being obtained. New generation ICP instrumentation should allow data to be produced, but there is a clear possibility that the low Hf contents make opx much more susceptible to the effects of metasomatic enrichment and hence the expected gain in chronological information might be

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marginal. In the absence of opx data, the published data that mostly focus on cpx measurements, with subsidiary garnet data also published, will be reviewed. The accumulated Lu and Hf data of peridotite minerals show only a small number of cpx that actually exhibit the high Lu/Hf ratios (large degree of fractionation), which would be expected if melting extended to approximately 10– 15% fractional melt extraction (Figure 11). In fact, most published mantle cpx appear to exhibit Lu–Hf abundances that are consistent with the modeled primitive mantle cpx composition, that is, they do not seem to record explicit evidence of extensive melt extraction. In these samples, other trace element

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Figure 11 Compilation of Lu/Hf versus Hf and Lu (ppm) of off-cratonic mantle cpx (a and b) and cratonic garnet (c and d). Melting trends are taken from Wittig et al. (2006; spinel-facies depletion of cpx) and Wittig et al. (2008a; 3GPa garnet-facies depletion of garnet (red) and cpx (green)) and are shown as 5 and 4% increments, respectively, and compared to kimberlite garnet (Nowell et al., 2004; Scully et al., 2004). Data for spinel-facies offcraton cpx from xenoliths entrained in intraplate volcanic rocks is filtered for those samples that also yield Nd–Hf isotope compositions (Bianchini et al., 2007, 2010; Bizimis et al., 2003; Shaw et al., 2007; Teklay et al., 2010; Wittig et al., 2006, 2007, 2010a). The cpx data from the North China craton is shown with the post-Archean data from spinel-facies xenoliths (Choi et al., 2008; Chu et al., 2009). Data for garnet from cratonic xenoliths is filtered for those samples that also yield Nd–Hf isotope data (Aulbach et al., 2004a,b; Bedini et al., 2004; Carlson et al., 2004; Ionov et al., 2005a; Lazarov et al., 2009; Simon et al., 2007; Wittig et al., 2008b).

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from a theoretical point of view. In Figure 13, the authors present a simple numerical model that calculates Hf isotope composition of variably fertile reference peridotite, which has been depleted at times that correspond with enhanced production of continental and oceanic crusts (0.35, 1.2, 1.9, and 2.8 Ga; Condie, 1998 and references therein; Pearson et al., 2007). The model focuses on the whole-rock Hf isotope evolution of melting residues. In order to acquire a conservative estimate of the Hf isotope composition that may be expected from continental mantle roots depletion is calculated as fractional modal melting in the spinel facies and the initial Lu–Hf isotope composition is taken as chondritic (Blichert-Toft and Albare`de, 1997). A more detailed model that relates the elemental and isotopic compositions of cpx to whole rocks and the effects of secondary metasomatic processes involving melt–rock reaction is given by Stracke et al. (2011). A first-order observation from this thought experiment is that only very young and small depletion events – for example, 0.35 Ga and 5% – will yield Lu/Hf ratios and, consequently, Hf isotope compositions that are akin the present-day convecting mantle (eHfpresent þ11). Assuming 5% depletion at 1.2, 1.9, and 2.8 Ga, whole-rock eHfpresent of approximately þ38, þ60, and þ89, respectively, may be expected. Older and/or more extensive melting results in more radiogenic 176 Hf/177Hf than seen in any convecting mantle material and also exceeds most values typically found in mantle xenolith minerals. For example, many cratonic peridotites indicate Os depletion ages of 2.8 Ga (Carlson et al., 2005; Pearson et al., 2004) and platinum-group element (PGE) compositions that indicate melting extents that exceed 20–25% melting, that is,

markers also point to negligible depletion and/or metasomatic enrichment. Not surprisingly, Hf isotope systematics in these cpx are akin to either the composition of their host volcanic rocks or the convecting mantle, as defined by the mid-ocean ridge basalt (MORB)/OIB array (Figure 12). Importantly, those few cpx with radiogenic Hf isotopic compositions, which are resolved from the mantle array, typically have Nd isotope compositions that are within the range of the convecting mantle, that is, Hf and Nd isotopes are decoupled in a fraction of the lithospheric mantle (Figure 12; Bedini et al., 2004; Carlson et al., 2004; Choi et al., 2008, 2010; Ionov et al., 2005a; Salters and Zindler, 1995; Schmidberger et al., 2002; Stracke et al., 2011; Wittig et al., 2006; see also Chapter 3.5). In order to generate this decoupling of Nd from Hf isotopes, Nd and other REE are required either to have drastically different diffusion coefficients and hence closure temperatures relative to Hf (Bedini et al., 2004) or to achieve faster equilibration with percolating liquids relative to Hf and Lu (e.g., Stracke et al., 2011). Whichever mechanism operates, the result is that Nd isotopic compositions are reset more readily than Hf isotope compositions. This does not imply that the Lu–Hf isotope system is immune to the effects of mantle metasomatism but suggests that metasomatic liquids, such as certain carbonatitic and hydrous fluids, may carry minimal high field strength elements and Lu, while silicate melts have the capacity of overprinting Lu and Hf together with other trace elements. To understand the potential for the Lu–Hf system to date lithospheric mantle formation, it is important to examine more closely the behavior of Lu and Hf during mantle melting

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

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beyond cpx exhaustion. Such a melting event would likely produce a eHfpresent value on the order of 1.3  105, stemming from a whole-rock 176Lu/177Hf of 71. While such extreme values may not necessarily be expected from off-craton spinel-facies peridotites, moderate depletion in the Proterozoic (10% and 1.9 Ga) should produce eHfpresent values of c.300. Yet neither the lavas from convecting mantle nor the offcratonic SCLM appear to systematically record this type of Hf isotope signature (e.g., Bianchini et al., 2007, 2010, 2011; Choi et al., 2008, 2010; Ionov et al., 2005a; Zhang et al., 2011). Some circumcratonic peridotite suites contain a few exceptions. For instance, two samples (out of 40) from the Penglai and Shanwant Cenozoic basalt locations surrounding the NE China craton have eHfT values of þ553 and þ655 (Chu et al., 2009). But in general, such suites appear to have been overprinted by metasomatism. These melt-extraction models only apply to cpx depleted during spinel-facies melting or to spinel-facies whole-rock peridotite. Yet, mantle garnet has received specific interest because of the high compatibility of Lu and the resulting high initial and residual Lu/Hf ratios (Figure 11). This causes

Lu to accumulate during melt extraction while Hf is being extracted preferentially (Figure 11). Garnet-facies depletion models may be calculated (e.g., 3 GPa; Figure 11); however, Lu and Hf concentrations in garnet are pressure–temperature– composition (P–T–X)-dependent, and so, the simple model presented here should only be viewed as indicative, regarding the extent of depletion. While many cratonic peridotites do contain garnet with radiogenic Hf isotopic compositions (e.g., Carlson et al., 2004; Simon et al., 2007), only in a few samples do the garnets show high Lu/Hf and radiogenic Hf isotopes that are significantly greater than the highest values for cpx from (off-cratonic) spinel peridotites (Figures 11 and 12). Melting beginning at much greater depth would yield larger Lu/Hf fractionation than shown in Figure 11 for shallow garnet-facies melting. An explanation for the low Lu/Hf in garnet from cratonic peridotites may be that melting in the shallow garnet facies will support an initial modal abundance of 10%, which is exhausted from the residue after only 12% melt extraction (Figure 11; note that garnet persists as a residual phase to extensive melting at pressures >5 GPa; Walter, 1998). Perhaps a more likely explanation of the prevalence of

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

convecting mantle, such as signatures, or values that are within the range shown by abyssal and oceanic peridotites (Bizimis et al., 2003; Stracke et al., 2011), even in cratonic garnets, is the metasomatic origin of many peridotite garnets (e.g., Pearson and Nowell, 2002; Shimizu et al., 1997; Simon et al., 2003). This distinction must be made for each cratonic xenolith locality based on supporting major and trace element data. A recent metasomatic origin or extreme overprinting of residual garnets seems necessary to explain the Hf isotope composition of most cratonic garnets from the cluster of cratonic peridotite Hf isotope data that are close to chondritic, strongly overlapping the Lu–Hf elemental and isotopic composition of garnet megacrysts from kimberlites (Nowell et al., 2004; Scully et al., 2004). Although the bulk of the cpx and garnet Hf isotope data appears to be significantly less extreme than would be predicted from simple time-integrated depletion models and much of the data overlaps the range for convecting mantle (Figure 12), some notable exceptions occur. Mantle cpx from the northern Massif Central, France, have strongly correlated Lu–Hf elemental and isotopic systematics that indicate substantial depletion at approximately 360 Ma, identical to the closure of the Variscan ocean and its subduction-zone volcanism (Wittig et al., 2006). In Figures 11 and 12, the Massif Central cpx mark the most extreme Lu/Hf fractionation and eHf values. The most extreme radiogenic cpx from Penglai, N China (Chu et al., 2009), give c.1.3 Ga TDM Lu–Hf model ages that are consistent with the Re–Os systematics of these samples. The more fertile samples from Penglai show little Lu–Hf fractionation, pointing to a smaller degree of initial depletion, but Lu–Hf isotope ratios that are reasonably well correlated along an errorchron, consistent with an age for the suite of c.1300 Ma. Spinel-facies peridoties from Svalbard show broad but very scattered Lu–Hf isotope correlations that suggest Paleo- to Mesoproterozoic depletion, with eHf values ranging from þ22 to þ149 (Choi et al., 2010). Spinel peridotites from the Keluo potassic lavas in the Central Asian Orogenic Belt of N China appear to have experienced Paleoproterozoic melt depletion from Re–Os isotope constraints, yet their eHf values show a narrow range between 8 and þ39, suggesting overprinting of the Lu–Hf system (Zhang et al., 2011). Only a single sample in this suite (08KL-13) has Nd (TDM), Hf (TDM), and Os (TRD) model ages that concur at c.1.2 Ga. In cratonic peridotites, Hf isotope signatures in constituent minerals vary widely and no statistically robust Lu–Hf isochrons have been reported. Nonetheless, numerous samples show indications of ancient signatures. The peridotite xenoliths from the Wyoming craton analyzed by Carlson et al. (2004) are extremely disturbed, with no Archean Hf TDM model ages, despite clear indications of their Archean heritage from Re–Os isotopes (Carlson and Irving, 1994; Carlson et al., 2004). Garnet lherzolites from the Kimberley and Lesotho areas of the Kaapvaal craton appear equally variable in their Lu–Hf isotope systematics. The cpx in these xenoliths appear to have been formed close to the age of the host kimberlite (e.g., Simon et al., 2003) and to have, in general, Hf isotope compositions that span the mantle array. In contrast, garnets in the same rocks can have measured eHf values of up to 300 and some have Neo- to Mesoarchean TDM Lu–Hf model ages, consistent with the Re–Os ages, but the spread in model ages is very large (Simon et al., 2007). Jacob et al. (2005) found that minerals from cratonic eclogite xenoliths

275

from Roberts Victor, South Africa, have equally varied Hf isotope compositions (initial eHf values 37 to þ2561) but without systematic and accurate age information. Examples of much more consistent age information being produced from the Lu–Hf isotope system in cratonic peridotites do exist. Lazarov et al. (2009) focused only on single grains of low-Ca, high-Cr (G10) garnets from the Finsch Mine to provide considerably more consistent Lu–Hf isotope systematics. These authors screened samples using Ca–Cr systematics along with the shapes of REE patterns. They found that the garnets with complex ‘sigmoidal’ REE patterns gave much more consistent Lu–Hf isotope systematics than those with silicate-melt equilibrated REE patterns. The garnets with sigmoidal REE patterns gave a well-defined Lu–Hf isochron of 2.5 Ga and an elevated initial Hf isotope composition (eHf ¼ þ20; Lazarov et al., 2009). One garnet gave a model age of 3.5 Ga, interpreted by the authors as the age of primary melting of the suite. On the basis of a moderately elevated initial Hf isotope composition for the garnet isochron, these authors interpret the Neoarchean isochron as the time of the second stage of melt extraction of these peridotites during craton amalgamation, following an original melting event, to form oceanic lithosphere, in the Mesoarchean. There are several caveats to the interpretation of single-mineral Hf isotope data from polyphase peridotites. Elevated initial eHf values of þ20 or more (depending on the melting model) can be generated from the residues of 5–10% fractional melting in 50–100 My, hence the need for c.700 My of prehistory as cratonic lithosphere prior to isochron generation (as suggested by Lazarov et al., 2009) is not, at present, clear for the Finsch locality. Nonetheless, this novel study, using single garnets carefully filtered for systematic trace element compositions, highlights that it is possible to extract more consistent age information from the Lu–Hf isotope system in cratonic and post-Archean peridotites. Selection of samples that appear to have experienced similar geochemical histories is key to obtaining clearer geochronological information from silicates. New instrumentation should make it possible to analyze other phases, such as opx, to better interrogate the age information offered by garnet and cpx. Where circumstances combine to yield sample peridotite or pyroxenite suites that have experienced little metasomatism, then the Lu–Hf isotope system is capable of defining melt depletion ages for Phanerozoic rocks (Wittig et al., 2006) because of the extreme isotopic diversity produced in the melt residues of moderate to ancient age. In contrast, Re–Os model ages are unlikely to yield reasonable age information from such young rocks because of significant variation in the Os isotopic composition of Phanerozoic to modern convecting mantle (e.g., Pearson et al., 2007). However, from the available data, Os isotope studies appear far more likely to define Archean ages in cratonic peridotites than Lu–Hf isotope studies, unless highly focused Lu–Hf studies are made, perhaps analyzing single-mineral grains of both garnet and opx, from only harzburgites.

3.6.4.2

Age Constraints from Re–Os Isotopes

Following initial work by Walker et al. (1989), Reisberg et al. (1991), Carlson and Irving (1994), and Pearson et al. (1995a,b), the Re–Os isotope system has become the most regularly

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

6

4 3 2 1

3.6.4.2.1 Data presentation (a)

Yangtze block n = 16

0 5

Uniform bandwith = 200 My

Number

4

Relative probability

The geochronology community has steadily adopted the use of probability density diagrams for presenting age information. These plots, generated from kernel density estimates, give smoothed plots that do not require the use of end-point bins. They are easily generated by software such as ‘ISOPLOT.’ Despite these attractions, probability density diagrams suffer from the same drawback as histograms, namely, that the number of apparent modes is susceptible to the choice of bandwidth. In addition, they should not be plotted for small numbers of samples, especially without some clear label indicating the numbers of samples plotted. The smoothed nature of the plots creates the impression of high densities of data points, whereas sometimes very few data points are plotted and yet this is not clearly specified on the plots. This leads to misleading conclusions regarding the significance of data peaks. The drawbacks of probability density diagrams, in terms of potentially depicting misleading representations of data, have been summarized from geochronological and geochemical contexts by Galbraith (1998) and Rudge (2008). Despite their drawbacks, probability density plots are useful and equally valid ways of comparing geochronological data as long as some important guidelines are followed. Perhaps the most important two rules for presenting Os isotope model ages on probability density diagrams are to reserve the plots for large numbers of samples (>100; Galbraith, 1998) and to use a uniform bandwidth of appropriate width (Sheather and Jones, 1991). In other words, researchers should avoid plotting the uncertainties for individual model ages on such plots because this can lead to huge overemphasis on a single data points that have analytical errors that are much smaller than other data. An example of this latter behavior is common in the depiction of Re–Os ages obtained using laser ablation inductively coupled plasma mass spectrometry (ICP-MS) analysis on sulfides. The large variation in Os concentrations within mantle sulfides (e.g., Alard et al., 2002, 2005; Aulbach et al., 2009; Burton et al., 2000; Pearson et al., 1998) means that analytical precision varies by a factor of 100 or more, where a single, very precise analysis can appear as a large ‘spike’ on a probability density diagram. If an over optimistic view of the knowledge of the mantle Os isotope composition is taken, then analytical precision dominates the uncertainty in estimates of model age uncertainty. This, in turn, leads to variations in over 100 of the bandwidths on probability density diagrams that plot the model ages, creating a misleading appearance and hence potential confusion in the significance of the ‘age spikes.’ In fact, the ‘spikes’ in sulfide ages portrayed in this way come from very narrow peaks, often for single samples, where overly optimistic age precision for that sample creates a ‘spike’ superimposed on broader peaks made from some sulfides that have

Small “n”; individual 1s used as bandwiths

5

Relative probability

applied isotope system for dating lithospheric mantle. Despite its success, the system has been shown to have several complications, reviewed recently by Rudnick and Walker (2009), meaning that care must be taken in the application of the system and the presentation and interpretation of results. Before summarizing recent results and their implication for cratonic lithosphere genesis, the authors amplify and add to some of the caveats raised by Rudnick and Walker (2009).

Number

276

3 2 1 0 0

(b)

0.8

1.6 2.4 3.2 TRD Os model age Ga

4.0

Figure 14 Probability density plots of sulfide data for 15 grains from peridotite xenoliths from the ‘Yangtze’ block, eastern China (Xu et al., 2008b). Panel (a) shows the probability density plot, after Xu et al. (2008b), constructed using model age uncertainties for individual sulfides that are dominated by analytical errors. Panel (b) shows the same data plotted using a uniform 200 My bandwidth that is close to the optimum bandwidth calculated using the approach of Sheather and Jones (1991) and is accompanied by an underlying histogram.

much lower model age uncertainties, sometimes as large as 1.3 Ga (e.g., Xu et al., 2008b). An example of these problems is illustrated in Figure 14(a), where a small dataset from a single xenolith locality has been plotted using model age uncertainties that are dominated by analytical precision (Xu et al., 2008b) and attention is drawn by the authors to the apparent ‘peaks’ at 0.6 and 1.2 Ga. Replotting of this data at a more appropriate and uniform bandwidth, which more adequately reflects the uncertainty in mantle reservoirs in the Meso- to Neoproterozoic (Figure 14(b)), gives a very different looking graph with a single mode at c.0.6 Ga. This clearly illustrates the importance of avoiding the use of varying bandwidths for different samples on probability density diagrams depicting age information. An additional example of likely oversmoothing of age data on probability density diagrams is the abyssal peridotite data presented by Pearson et al. (2007). When using Re–Os model ages to evaluate the relative proportions of lithosphere forming at different times, it is important to avoid overemphasis of single samples that may yield many sulfide grains. If more than one sulfide model age per xenolith is plotted and there are numerous grains that yield a particular age, then there will be an overrepresentation of that age relative to others (e.g., Griffin et al., 2003). Moreover, if the highly depleted rocks have no sulfide, then they will not

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

register on plots of only sulfides and hence their compositions will be unrepresented on such diagrams. Sulfides are most common in metasomatized peridotites (Lorand and Gregoire, 2006), which are less likely than highly depleted samples to yield undisturbed ages. The abundance of sulfides in certain samples can seriously skew the portrayal of age populations. Nonetheless, it is clear from sulfide studies that, where they are present, they can lead to much valuable chronological information about enrichment events in the lithospheric mantle (e.g., Aulbach et al., 2009).

3.6.4.2.2 Analytical approach: Whole-rock and sulfide analyses The analysis of mantle sulfides for Os isotope ratios (e.g., Alard et al., 2002, 2005; Burton et al., 2000; Griffin et al., 2003; Luguet et al., 2008; Pearson et al., 1998, 1999, 2002a; Richardson et al., 2001) has provided much complementary information to whole-rock analyses, and where both phases are present in the rocks, then data for both types of sample should be obtained. A recent review by Rudnick and Walker (2009) makes some comparisons between the two approaches, and only a few points are highlighted here. While whole-rock Re–Os analyses of peridotites are integrations of numerous potential events that may have affected the evolution of a cratonic peridotite, there is abundant evidence for the recent introduction of significant Re (Carlson et al., 2005; Pearson et al., 1995a; Walker et al., 1989) and some Os (Chesley et al., 1999) into whole-rock peridotites. This is why model ages that are calculated assuming a Re content of zero (the TRD age) are preferred for cratonic peridotites (Carlson and Irving, 1994; Pearson et al., 1995a,b; Walker et al., 1989). Much of the Re addition is recent in these highly depleted residues (Carlson et al., 2005), and this, together with the low Os abundances in metasomatic sulfides, translates to very little disturbance for depleted, low-S samples. In contrast, the TRD age approach should not be adopted in peridotites that have experienced little to even moderate melting, where Re is unlikely to have been quantitatively extracted from the melting residue. Whole-rock Os isotope studies can be combined with PGE analyses (Pearson et al., 2004) to more fully evaluate the likelihood of Re disturbance in depleted residues. For example, Liu et al. (2010) proposed that the very fractionated PGE patterns (Os, Pd, and Re depletions relative to Ir) in peridotites from Yangyuan, N China, combined with their low S and Se contents, suggest that they experienced sulfide breakdown due to an interaction with a S-undersaturated melt/fluid. These factors mean that TMA ages (Walker et al., 1989) for wholerock peridotites are likely to be unreliable and one should use TRD ages as relatively conservative estimates of minimum melt depletion ages. An additional point to be aware of when evaluating wholerock Os isotope data is the issue of the poor comparability between data obtained via negative thermal ionization mass spectrometry (N-TIMS) and sparging into an ICPMS instrument. Liu et al. (2010) provide a comparative study of samples analyzed via both techniques from the N China craton. These authors find that Os concentrations and Os isotope compositions analyzed by N-TIMS, from two locations in their study, differed significantly from the sparging data for the same

277

samples published by Xu et al. (2008a) and Zhang et al. (2009). Further work is needed in understanding the sparging technique before the Os isotope data produced in this way can be used with the same confidence as N-TIMS data. The sulfide dating approach also has significant pitfalls that need to be carefully considered when interpreting Re–Os age information. The low solidi of the monosulfide solid solution (mss) mean that it is one of the first phases to enter the melt during peridotite melting well below the silicate solidus. As such, the high melt fractions extracted from cratonic peridotites mean that mss will no longer be present as a residual phase (Lorand and Gregoire, 2006; Pearson et al., 2004). This view is also confirmed by the experimental studies of Os and Ir in sulfide melts, which show, along with Re/Os systematics in oceanic basalts, that Os must be hosted in an alloy after partial melting (Fonseca et al., 2011). Hence, any mss or Ni-rich pentlandites found in cratonic peridotites are metasomatic in nature (Lorand and Gregoire, 2006; Luguet et al., 2003) and so by definition will carry a mixed isotopic signal. Furthermore, peridotitic sulfides are relatively scarce among cratonic peridotites, especially when weathered. An intensive search for high-T base metal sulfides in a suite of >20 peridotites from the Argyle Diamond Mine by Luguet et al. (2009) was unsuccessful and whole-rock analysis was the only possible approach. When sulfides are present in cratonic peridotites, they can offer a wealth of additional information about melting and metasomatism (e.g., Aulbach et al., 2004a). The tendency of mantle sulfides to exsolve on cooling means that it is imperative to sample the entire sulfide in order to obtain a valid isotopic measurement for use in age calculations (e.g., Richardson et al., 2001). Hence, any laserbased method is likely to subsample a sulfide and to generate additional analytical uncertainty that must be considered in evaluating the significance of any age. This may contribute to the very large differences in replicate analyses reported in some sulfide Re–Os studies that translate into model age differences in ablation spots on the same sulfide of up to 600 Ma for TRD or TMA (Xu et al., 2008b). Analysis of whole sulfides by solution yields improved analytical uncertainties and is less vulnerable to exsolution problems (Pearson et al., 1998). It should be clear from the above discussion that both whole-rock and sulfide Re–Os dating suffer from drawbacks that make it necessary to apply the approach carefully on well-characterized rocks. While whole-rock ages tend to be disturbed to younger ages by radiogenic Os addition (e.g., Chesley et al., 1999), sulfide model ages can scatter to both younger and older ages. The above summary shows that there is significant disturbance and uncertainty in both whole-rock and sulfide Re–Os dating approaches. The requirement that Os is hosted in microto nanoalloy phases in mantle residues that have experienced significant melting (Fonseca et al., 2011; Luguet et al., 2007) means that whole-rock analyses are best conducted on residual dunite–harzburgite, accompanied by PGE abundance data to evaluate late-stage Re- and P-PGE addition. Sulfides, where possible, should be analyzed as whole grains, unpolished, to avoid removing portions with differential Re/Os. When both whole rocks and sulfides are analyzed on large suites of peridotites, a detailed view of the melting and enrichment history of continental lithosphere should be possible.

278

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

3.6.4.2.3 Os isotope heterogeneity of the mantle A significant issue relating to the interpretation of Re–Os ages in peridotites (whole rock or sulfide) is the growing evidence for considerable Os isotopic variation in the convecting mantle through time, especially the preservation of ancient-depleted domains. This variability is manifested in the now substantial variation observed in recent samples of the convecting mantle. The presence of a significant fraction of samples with unradiogenic Os isotope compositions, indicating long-term melt depletion, was first highlighted by Meibom et al. (2002) and Meibom and Frei (2002) for PGE-alloy phases from ophiolites. The earlier indications of such compositions had come from mantle xenoliths in subduction zones (Parkinson et al., 1998) and from abyssal peridotites (Snow and Reisberg, 1995). Subsequent studies of abyssal peridotites (Brandon et al., 2000; Harvey et al., 2006; Liu et al., 2008) and their sulfides (Alard et al., 2005) and ophiolitic chromites (Walker et al., 2002), together with the publication of new, extensive data for PGEalloys (Brandon et al., 2006; Pearson et al., 2007), expanded the range. Pearson et al. (2007) noted that the modes in Os model ages coincide with the age peaks in continental crust production. Bizimis et al. (2007) and Ishikawa et al. (2011) have both noted that depleted domains yielding depletion ages as old as c.2.0 Ga are also present in the residual peridotites accreted by modernday plumes beneath Hawaii and the Ontong Java Plateau. This very large variability in the Os isotope composition of modern convecting mantle (Figure 15) strongly overlaps the global range for off-craton mantle peridotites making the prospect of dating lithosphere formation using the Re–Os system very difficult, unless there is a clear isochronous relationship. The latter is an unlikely outcome for peridotites in the modern Earth because this inherent variability will cause very significant variations in initial isotopic ratios. The problem becomes less severe further back in Earth history as heterogeneity reduces (e.g., Day et al., 2008; Rudnick and Walker, 2009). The inherent Os isotope variability in convecting mantle, with around 10% of ‘modern’ mantle having Os isotopic compositions that yield model ages of between 1 and 1.3 Ga (Pearson et al., 2007), means that it is very difficult to confidently assign a Proterozoic depletion age to a suite of peridotites where only a few samples yield ages between 1 and 2 Ga and the rest have Os isotope compositions similar to the convecting mantle average. For example, with the new database for convecting mantle, it is difficult to make a case for a Proterozoic age for the mantle lithosphere beneath the Kerguelen Plateau (Hassler and Shimizu, 1998). Similarly, the single xenolith yielding an Archean TRD model age from Spitsbergen, western Svalbard (Choi et al., 2010) is not strong evidence of the presence of an Archean mantle beneath this region without corroborating evidence from other isotope systems or additional samples being found with Archean ages. More confidence can be placed in sample suites showing a significant proportion of Archean ages because 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 clear difference in the mean and overall distribution of Os isotope compositions between cratonic peridotites and modern convecting mantle (Figure 15; t-value 2.084;

0.14 187Os/ 188Os

0.135 0.13 0.125 0.12 0.115 0.11 0.105

n = 1445

n = 210

n = 70

n = 501

0.1

Modern mantle Off craton Circumcratonic

Cratonic

Figure 15 Box and whisker plots (parameters defined in Figure 7) of 187 Os/188Os ratios for peridotites from a variety of tectonic settings. Number of samples (n) given for each category. Modern convecting mantle comprises Pt-group alloy grains (data from Brandon et al., 2000, 2006; Meibom et al., 2002; Pearson et al., 2007; Pearson, unpublished data; Shi et al., 2007), Phanerozoic ophiolites (Dijkstra et al., 2010; Hanghøj et al., 2010; Marchesi et al., 2006; Schulte et al., 2009; Walker et al., 2002), OIB xenoliths (Bizimis et al., 2007), and arc peridotites (Brandon et al., 1996; McInnes et al., 1999; Parkinson et al., 1998; Widom et al., 2003). Cratonic samples from data sources summarized in Pearson and Wittig (2008), Menzies et al. (1999, 2004), Gao et al. (2002), Westerlund et al. (2006), Peltonen and Bru¨gmann (2006), Smith et al. (2009), Luguet et al. (2009), Chu et al. (2009), and Wittig et al. (2010b). Off-craton samples from McBride et al. (1996), Becker et al. (2001), Handler et al. (1997, 2003, 2005), Handler and Bennett (1999), Lee et al. (2000, 2001), Meisel et al. (1996, 2001), Peslier et al. (2000a,b), and Reisberg et al. (2004, 2005). Non-Archean samples erupted through the NCC have been omitted due to their complex provenance. Circumcratonic samples are kimberlite-hosted xenoliths from immediately surrounding the Kaapvaal craton (Janney et al., 2010; Pearson et al., 2004).

dof ¼ 161; and tprob ¼ 0.0387), make it very likely that the abundant Archean Re–Os model ages for cratonic peridotites reflect melt extraction during lithospheric formation in Archean times. It is not possible, in the authors’ opinion, to resolve within 200 Ma when the depletion took place. For instance, the occasional sample that yields a Mesoarchean depletion age within a suite that yields average depletion ages in the Neoarchean is not definitive evidence of lithosphere formation at that time followed by subsequent disturbance, which was the prevalent interpretation in early datasets (e.g., Pearson et al., 1995a,b). The occasional older ages in cratonic peridotites may equally reflect preserved heterogeneity from earlier melting events, where residues remained in the convecting mantle, to be later accreted as lithosphere.

3.6.4.3

In Situ Pb–Pb Dating of Clinopyroxenes

Recently, the unique characteristics of the U–Pb system have been exploited to date low U–Pb mantle minerals via laser

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

ablation multicollector ICPMS (Hunt et al., 2012; Schmidberger et al., 2007; Tappe et al., 2011). This builds on the foundation laid by Jacob and Jagoutz (1994) who used Pb–Pb systematics in multigrain cpx separates analyzed by TIMS to date the age of subduction of Roberts Victor eclogites. Clinopyroxene has been the focus of interest and the method has yielded useful geochronological information for both peridotitic and eclogitic rocks. Improvements in the sensitivity of new instrumentation should permit extension of the technique to other samples. The resulting Pb–Pb isotope relations are often complex, and careful supplementary petrography and trace element characterization are required to ensure that ‘disturbed’ grains can be eliminated. Nonetheless, numerous data points can be obtained for the same sample or between multiple samples that, in the cases studied so far, define secondary Pb–Pb arrays whose intersection with model Pb growth curves is taken to reflect a major event in the evolution of the host rock. This event could either be the age of major melt extraction from a peridotite or eclogitic residue (Hunt et al., 2012; Jacob and Foley, 1999; Tappe et al., 2011) or, if the cpx is metasomatic, some integrated age that reflects both the metasomatic event and the age of the protolith. Once such dating studies proliferate, it will be necessary to evaluate the Pb–Pb model ages with chronological information derived from other systems or other phases to better understand what event is being dated by the Pb–Pb arrays.

Kaapvaal craton

Relative probability

Sulfides TMA n = 71

v

Whole rocks TMA n = 228 Kaapvaal craton

Relative probability

Whole rocks TRD n = 228

Sulfides TMA n = 71

3.6.4.4 Slave craton

Relative probability

Due to its abundance of large mantle xenoliths, originating from historical diamond mines, the Kaapvaal craton (Figure 1) has been the subject of the most intense and systematic Re–Os dating study of any craton in terms of the diversity of locations where dating studies have been undertaken. Early Re–Os studies of Kaapvaal peridotites suggested that the main mass of cratonic lithospheric mantle was Mesoarchean in age, similar to much of the early crust-building phase (e.g., Pearson, 1999; Pearson et al., 1995a). However, the large database now available (Carlson and Moore, 2004; Carlson et al., 1999, 2005; Janney et al., 2010; Pearson and Wittig, 2008; Simon et al., 2007) allows a more detailed examination of the age information that can be retrieved from these rocks. Only a small portion of the data yield Mesoarchean ages (Figure 16), and hence, craton evolution is undoubtedly more complex than the early models suggested.

Whole rocks TRD n = 36

Relative probability

N Atlantic Whole rocks TRD n = 50

0.5

1.0

1.5

2.0 2.5 3.0 Model age (Ga)

3.5

4.0

Case Studies: Dating Cratonic Mantle

3.6.4.4.1 The Kaapvaal craton

Sulfides TMA n = 48

0.0

279

4.5

Figure 16 Probability density plots of Re–Os model ages for whole-rock peridotites and sulfides for the Kaapvaal craton, Slave craton, and N Atlantic cratons. Whole-rock data, Kaapvaal craton – Pearson et al. (1995a), Carlson et al. (1999), Irvine et al. (2001), Carlson and Moore (2004), Menzies et al. (1999), Simon et al. (2007); Slave craton – Irvine et al. (2003), Aulbach et al. (2004b); Westerlund et al. (2006); N Atlantic craton – Hanghøj et al. (2001), Bernstein et al. (2006), Wittig et al. (2010b). All TRD ages corrected to the eruption age of the kimberlite. Model age equations for TRD and TMA presented in Walker et al. (1989) and Pearson (1999) and are calculated to reference values for chondritic mantle of 187Os/188Os ¼ 0.1283 and 187Re/188Os ¼ 0.422. Number of samples plotted for each population given by ‘n.’

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

The statistical distribution of whole rock TMA and sulfide TMA model ages for the Kaapvaal craton shows remarkable similarities (Figure 16), with prominent but rather broad modes at 2.9–3.0 Ga. The distribution of sulfide TMA ages shows a greater proportion of ages in the Eoarchean to Hadean range, but this likely reflects Re addition or problems with Re interference corrections (see Nowell et al., 2008) causing overestimation of TMA ages. In contrast, the Kaapvaal peridotite TRD ages show a tighter, more pronounced mode at 2.7 Ga, with relatively few ages that are Mesoarchean or older (Figure 16). The clustering of whole-rock TRD model ages at around 2.7 Ga is unlikely to result from variable metasomatic addition of sulfides, as suggested by some authors (e.g., Griffin et al., 2004), as such processes would lead to a more diffuse spread in ages. Furthermore, most metasomatic sulfides either would have insufficient Os to significantly affect the bulk-rock composition (Pearson et al., 2004; Reisberg et al., 2005; Rudnick and Walker, 2009) or would produce very different trends to those observed for whole rocks on a Re–Os isochron diagram (Carlson et al., 2005). 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), although the latter ages would be overestimates if the garnets formed from mantle enriched by subducted crustal material at 2.9 Ga, thereby reconciling the Os and Nd ‘ages.’ The collisional suturing of the Witwatersrand and Kimberley blocks of the Kaapvaal craton occurred at c.2.9 Ga (Schmitz et al., 2004). The coincidence of this crustal suturing age and the major mode of Re–Os model ages may suggest that the Re–Os system is dating lithosphere formation related to this suturing event. The Neoarchean has been proposed by Pearson et al. (2002b) and Pearson and Wittig (2008) to be the time of major formation and stabilization of lithospheric mantle beneath the Kaapvaal craton. The recently defined paleosubduction zone between the two main blocks (western and eastern) of the Kaapvaal craton provides a mechanism to accrete Kaapvaal peridotites, originally formed at an Archean MOR, in a collisional environment at this time (Gibson et al., 2008; Schmitz et al., 2004; Simon et al., 2007). Further support for the timing of major subduction at 2.9 Ga, which probably caused the enrichment of opx in Kaapvaal peridotites (see earlier), is provided by the 2.89  0.06 Ga Re–Os isochron for eclogitic sulfides included in diamonds (Richardson et al., 2001) and by the 2.7–3.0 Ga Re–Os model ages for whole-rock eclogites from this region (Shirey et al., 2001). Some evidence exists for the involvement of older lithospheric material, perhaps transposed from the older Witwatersrand block, in the form of pre-3 Ga Re–Os model ages for eclogites and peridotites from the Newlands kimberlite (Menzies et al., 1999, 2003). 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 sulfide TMA ages at 3.1 Ga

(Figure 16) has been interpreted to represent older lithospheric mantle generation in the Witwatersrand block (Griffin et al., 2004). While whole-rock TMA ages are consistent with this view of a significant fraction of Mesoarchean lithosphere (Figure 16), the TRD whole-rock ages are much more tightly clustered in the Neoarchean. Given the tendency for sulfide ages to spread back well past the age of the Earth due to difficulties in interfering element corrections and the likelihood that the Archean convecting mantle contained regions of previously depleted residues (but to a lesser extent than modern-day mantle), it is hard to place significance on these older ages. A careful analysis of existing and newer data by Carlson and Moore (2004) concluded that there was no definitive evidence for a difference in the age of the lithosphere between the two cratonic blocks. Lastly, the coincidence of peridotite xenolith suites with mean (or median) TRD Os model ages that are Archean and the boundaries of the craton based on crustal studies are striking (Figure 17). This intuitive relationship is strong support that the Re–Os model age approach in cratonic peridotite suites is a valid way of obtaining lithosphere formation ages when sufficient numbers of samples are analyzed (preferably at least 12–15 per location). In fact, this observation can be extended further – all cratonic regions that contain kimberlitehosted xenolith suites have produced a significant number of peridotites with Archean TRD ages (Kaapvaal – Walker et al., 1989; Wyoming – Carlson and Irving, 1994; Siberia – Pearson et al., 1995b; Tanzania – Chesley et al., 1999; Superior – Pearson et al., 1999; E North Atlantic – Hanghøj et al., 2001; W African – Barth et al., 2001; North China craton (NCC) – Gao et al., 2002; Slave – Aulbach et al., 2004b; Baltic Shield – Peltonen and Bru¨gmann 2006; W NAC – Wittig et al., 2010b; and Kalahari – Smith et al., 2009). This attests to the efficacy of the Re–Os system in identifying the presence of Archean lithospheric mantle, where it exists, beneath the cratons. The method is powerful enough to demonstrate the presence of

ZIMBABWE CRATON

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KAAPVAAL CRATON Newlands CARIEP BELT R. Victor N. Lesotho Finsch Kimberley Hoedkop Jagersfontein E. Griqualand Uintjesberg Monastery Hebron NAMAQUA-NATAL BELT Gansfontein Melton Wold

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Figure 17 Schematic map of the Kaapvaal craton and surrounding geological terranes depicting average TRD Os model ages for peridotite xenolith suites on and around the craton. Data sources in Janney et al. (2010). One exception is Roberts Victor where the age is a 2.7 Ga secondary Pb–Pb isochron for eclogites (Jacob and Jagoutz, 1994). Modified after Carlson RW, Pearson DG, James DE (2005) Physical, chemical, and geochronological characteristics of continental mantle. Reviews of Geophysics 43: 1–24 and references therein.

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

Archean mantle beneath regions of Proterozoic basement, for instance, beneath the Argyle Diamond Mine in the Paleoproterozoic Halls Creek ‘mobile belt,’ where Luguet et al. (2009) found clear evidence of the presence of Archean continental mantle. This makes the Argyle diamond deposit a more ‘conventional’ diamond deposit in that it is located above an area of Archean lithospheric mantle. One location within the Kaapvaal craton where peridotites do not have an average TRD age that is Archean is the Premier kimberlite, where Archean ages exist but most TRD ages are Mesoproterozoic (Carlson et al., 1999; Morel et al., 2008; Pearson et al., 1995a). This ‘age anomaly’ coincides with a distinct seismic anomaly beneath the central Kaapvaal craton (Carlson et al., 2005; James et al., 2001) that is coincident with the spatial extent of the Bushveld intrusion and likely represents the addition of new lithospheric residues and the modification of former Archean lithosphere in the Paleoproterozoic (2.05 Ga). Kimberlites sampling the lithospheric mantle surrounding the edges of the Kaapvaal craton have peridotite xenolith suites that yield Paleoproterozoic TRD ages. This attests to the strong long-term physical coupling of the crust and mantle over billions of years. In addition, it indicates the likelihood of subvertical lithotectonic boundaries at the craton edge, where the Archean crust and mantle transitions to a Paleoproterozoic crust and mantle over a small distance (Figure 17).

3.6.4.4.2 North China craton The last 10 years have seen a surge in Re–Os dating studies of the NCC (Figures 1 and 18). Major element and mineral chemical characteristics of the kimberlite-borne peridotite

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Figure 18 Map of the North China craton showing a shallow topography background with major tectonic divisions of the eastern block, the western block, and the central region. The central region – ‘central orogenic belt (COB)’ – is assumed here to be formed between the eastern and western blocks at 2.5 Ga (Kusky and Li, 2003). The Khondalite belt is a Paleoproterozoic metamorphic belt. Mantle xenolith localities shown as squares (Paleozoic eruption age), stars (Mesozoic eruption age), and circles (Cenozoic eruption age). For alternate interpretations and nomenclature of terrane boundaries, see summary in Liu et al. (2011a). Modified after Liu J, Rudnick RL, Walker RJ, et al. (2011) Mapping lithospheric boundaries using Os isotopes of mantle xenoliths: An example from the North China Craton. Geochimica et Cosmochimica Acta 75: 3881–3902.

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xenoliths of the NCC suggest the presence of thick Archean lithosphere that was subsequently removed in a c.120 km lithosphere-loss event in the Phanerozoic (Menzies et al., 1993). Proof of the Archean age of the NCC lithosphere came with the Re–Os study of Gao et al. (2002), subsequently enhanced by Chu et al. (2009), that showed that the kimberliteerupted deep-seated peridotites were Archean in age, whereas the shallow spinel peridotites erupted by Cenozoic alkalibasalt magmas all give Os (TRD) and Hf (TDM) model ages ranging from Mesoproterozoic to recent (Chu et al., 2009; Gao et al., 2002; Liu et al., 2011a; Wu et al., 2003, 2006; Xu et al., 2008b). Chu et al. (2009) interpret the range of Os isotope compositions from peridotites erupted in the Cenozoic (E NCC: Penglai and Shanwang), that give model ages ranging from Mesoproterozoic to recent, as an indication that this region of the craton is underlain by modern convective upper mantle-containing ancient-depleted domains rather than being accreted in the Proterozoic. This is consistent with Sr–Nd isotope constraints (Menzies et al., 2007). In contrast, Liu et al. (2011a) interpret the range of Os isotope compositions for peridotites in the more northern NCC, within the TransNorth China Orogen (Hannuoba and Yangyuan; Figure 18) as reflecting Paleoproterozoic formation rather than the refertilization of Archean peridotites (Xu et al., 2008a; Zhang et al., 2009). The NCC thus appears to be one of the most dynamic cratons in terms of the history of its lithospheric root, as discerned from Re–Os isotope constraints. The eastern block appears to have experienced almost complete removal of thick, diamond-bearing Archean lithosphere (Gao et al., 2002; Menzies et al., 1993). The steep trend of 187Os/188Os versus Al2O3 in the eastern NCC non-Archean peridotites leads Chu et al. (2009) to conclude that refertilization or peridotite–melt interaction was unlikely to have caused the replacement of the NCC lithosphere, as suggested by some models (e.g., Tang et al., 2008; Ying et al., 2006; Zhang et al., 2008). Instead, Chu et al. (2009) concur with the models of Gao et al. (2002) and Wu et al. (2006) that invoke thinning of the eastern NCC driven by foundering (delamination) of the deep crust and lithospheric mantle. Recently, Liu et al. (2010) have provided a detailed picture of crust–mantle relations across the NCC from a study of almost 100 xenoliths. This dataset and the studies cited earlier show that the Neoarchean ages obtained from the southern central section of the NCC (Hebi, Fushan, and Fansi; Figures 18 and 19) have ages that match those of the overlying crust, with a transition at the Fansi locality to lithosphere of Paleoproterozoic age that underlies a crust of at least 2.5 Ga (Figure 19). This indicates that in the central region of the NCC, the Archean mantle was replaced in the northern section by N–S collision during the assembly of the Columbia supercontinent at 1.8–1.9 Ga or from extrusion of 1.9 Ga lithospheric mantle from the Khondalite belt (Figure 18) beneath the Trans-North China Orogen during the 1.85 Ga collision of the eastern and western blocks of the NCC (Figure 18; Liu et al., 2011a). Peridotites erupted in locations outside the bounds of the NCC, such as at Panshishan, Lianshan, and Fangshan (Subei Basin; Figure 18) give maximum TRD Os model ages that are Paleoproterozoic, leading to suggestions that the lithospheric

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

mantle was formed at this time (Reisberg et al., 2005). However, in contrast to the Paleoproterozoic lithosphere surrounding the Kaapvaal craton (Janney et al., 2010), it is not possible to statistically distinguish these compositions from modern convecting mantle and accretion of recent convecting mantle is equally possible. The difficulty in adequately distinguishing Proterozoic mantle from modern convecting mantle, especially in small suites of peridotite xenoliths with only the occasional sample yielding a Mesoproterozoic age, casts some doubt on the suggestion of Senda et al. (2007) that Proterozoic mantle extending from SE and NE China underlies parts of SW Japan. Clearer evidence for the presence of Paleoproterozoic lithosphere was found by Zhang et al. (2011) at Keluo, in the

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Os isotope studies of spinel harzburgites and dunites from E Greenland (Hanghøj et al., 2001) and W Greenland (Bernstein et al., 2006) established that the shallow mantle beneath the NAC (Figure 1) is Archean in age, with Mesoarchean ages being obtained from some peridotites. An extensive study of more deeply derived peridotites from the SW Greenland part of the NAC (Figures 1 and 20) established the presence of Archean mantle into the diamond stability field (Wittig et al., 2010b). The total distribution of ages (Figure 16) shows a pronounced mode in TRD ages at c.2.7 to 2.8 Ga that corresponds broadly to the peak in crustal U–Pb zircon ages derived from the regional trondhjemite–tonalite–granodiorite (TTG) crust (Figure 21). No Paleoarchean TRD ages are recorded, yet the sampling locations for the peridotite xenoliths do not directly correspond to the exposure of Paleoarchean crust. Hence, it is not possible, from the available data, to rule out the presence of Paleoarchean lithospheric mantle. Significantly, there is a strong secondary mode in peridotite ages at c.2.0 Ga that appears to correlate with the tectonic activity associated with the amalgamation and rifting of the W Greenland section of the NAC (Kangaˆmiut dykes 2.0 Ga and Nagssuqtoqidian orogeny 1.8 Ga; Wittig et al., 2010b). The peridotites from the southernmost margin of the NAC in W Greenland are dominated by

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Figure 19 Schematic cross section across line A–A0 in Figure 21, which shows the crust–mantle age relations across the central zone of the North China craton. Reproduced from Liu J, Rudnick RL, Walker RJ, et al. (2011) Mapping lithospheric boundaries using Os isotopes of mantle xenoliths: An example from the North China Craton. Geochimica et Cosmochimica Acta 75: 3881–3902.

North Atlantic Craton Maniitsoq [UML ca. 580 Ma]

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eastern part of the Central Asian Orogenic Belt, N China, where 50% of the 20 samples analyzed yielded Paleo- to Mesoproterozoic TRD Os model ages.

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Figure 20 Schematic cross section across the North Atlantic craton and the Nagssugtoqidian mobile belt toward Ubekendt Ejland on the Disko craton showing the total field magnetic profile (in nanotesla) of the continental crust. Vertical columns through the lithosphere show the estimated depth and sampling range (dark gray) of the lithospheric mantle (Bizzarro and Stevenson, 2003; Nielsen et al., 2008; Sand et al., 2009), combined with TRD Os model ages (Wittig et al., 2010b) and calculated at the eruption age of the ultramafic lamprophyres (UMLs) for the selected samples from the UML dykes located at Pyramidefjeld, Nigerlikasik, Maniitsoq, Sarfartoq, Kangerlussuaq, and the alkali-basalt hosted xenolith locality at Ubekendt Ejland (Bernstein et al., 2006). The oldest TRD Os model ages derived from xenoliths sampled at individual dykes are given in brackets. Typical summary PGE patterns for whole-rock xenoliths also shown. Reproduced from Wittig N, Webb M, Pearson DG, et al. (2010) Formation of the North Atlantic craton: Timing and mechanisms constrained from Re–Os isotope and PGE data of peridotite xenoliths from S.W. Greenland. Chemical Geology 276: 166–187.

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

Nunatak-1390 eclogites NAC peridotites

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Age (Ga) Figure 21 Summary of geochronological data for the crust and mantle components of the W Greenland section of the NAC. Crustal age probability density diagram uses data from detrital zircon database of Nutman et al. (2004) and indicates a peak in late Archean crustal growth in the central NAC. Open rectangles indicate TTG magmatism and NAC accretion in the Tasiusarsuaq terrane (Naeraa and Schersten, 2008). Range and mean for peridotite mantle TRD Os model ages from Wittig et al. (2010b). Eclogite age is for a suite of eclogites from Nunatak 1390 in SW Greenland and is a Pb–Pb secondary isochron for cpx (Tappe et al., 2011). Reproduced from Tappe S, Smart KA, Pearson DG, Steenfelt A, Simonetti A (2011) Craton formation in Late Archean subduction zones revealed by first Greenland eclogites. Geology 39: 1103–1106.

c.2.0 Ga depletion ages and may represent strongly modified former Archean mantle. Newly discovered Archean (2.7  0.3 Ga Pb–Pb isochron on cpx) eclogite xenoliths from the SW Greenland part of the NAC add an additional dimension to the link between the crust and mantle in this region (Tappe et al., 2011). The age of the eclogites is broadly coeval with the major regional episode of TTG magmatism noted earlier (Figure 21). Major and trace element characteristics of the most pristine (nonkimberlite, altered) eclogites show a very refractory nature, in keeping with the very depleted compositions of the peridotite xenoliths from the same region (Wittig et al., 2008a,b). The eclogites have elevated d18O values in garnet and these, together with the complementary major and trace element compositional relationship with the SW Greenland TTG crust, indicate a protolith–melt relationship between the two (Tappe et al., 2011). The combined compositional and temporal relations between the crust and the eclogite and the peridotite portions of the mantle lithosphere in SW Greenland provide a strong indication for the growth of Neoarchean crust via the melting of oceanic crust in a subduction zone. The underlying oceanic mantle lithosphere, formed at a hot ridge, could then accrete to form, in this craton, a compositionally and chronologically coupled crust–mantle column.

3.6.5

Models for the Formation of Cratonic Roots

Any successful model for the formation of Archean cratons must comprehensively explain all aspects of the geochronology and geochemistry in both the crustal and mantle lithosphere sections of cratons. No model published to date has succeeded in this aim and so all models thus far are incomplete. The origins of cratonic keels and their role in craton formation,

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stability, and destabilization have been reviewed recently by Carlson et al. (2005), Arndt et al. (2009), Pearson and Wittig (2008), Foley (2008), Griffin et al. (2009), Lee et al. (2011), Aulbach (2012), and Herzberg and Rudnick (2012), with reviews of the dynamical considerations presented by Cooper et al. (2006) and Cooper and Conrad (2009). As such, it is not the authors’ intention to repeat all aspects of these discussions. Section 3.6.3 of this chapter concluded that the most likely tectonic setting for the formation of most cratonic mantle peridotites was an Archean MOR followed by accretion at a convergent margin, in agreement with numerous previous workers (e.g., Canil, 2004; Canil and Lee, 2009; Carlson et al., 2005; Gibson et al., 2008; Helmstaedt and Schulze, 1989; Herzberg and Rudnick, 2012; Kelemen et al., 1998; Lee et al., 2011; Simon et al., 2007; Stachel et al., 1998). The additional melting at a subduction zone (Parman et al., 2004; Pearson and Wittig, 2008) does not appear to be warranted despite the fact that the closest match to depleted cratonic mantle compositions in the modern Earth are forearc peridotites. The high Archean mantle potential temperatures in the Meso- to Neoarchean predicted by the Herzberg et al. (2010) thermal model allow melting at hot ridges to produce sufficiently depleted residua (Herzberg, 2004; Herzberg and Rudnick, 2012; Rollinson, 2010) and thick enough columns of residual peridotite to account for the depth-compositional relations seen in most well-studied cratonic keels (Section 3.6.3). If the Korenaga (2008) prediction of Earth’s thermal history is correct, then this peak in mantle potential temperature would coincide with most estimates of when cratonic mantle lithosphere was stabilized (see earlier and Pearson and Wittig, 2008). An attraction of the ‘hot ridge’ environment for making cratonic peridotites is that the compositions of the magmas produced in this setting, calculated from the experiment-based parameterizations, are a much better fit as residua to the most common ‘nonarc’ basalt composition found in Archean cratons (Herzberg and Rudnick, 2012; Rollinson, 2010). In contrast, the plume model (e.g., Arndt et al., 2009; Aulbach, 2012; Aulbach et al., 2009; Griffin et al., 2009; Pearson et al., 1995a), using the experimental parameterization of Herzberg (2004), cannot reproduce the most commonly observed Archean basalt compositions. The much more MgO-rich magmas demanded by the plume hypothesis must be somehow disposed of in this model. Furthermore, plume melting produces anomalous thicknesses of highly refractory residue compositions that are not observed in the xenolith or xenocryst sample set. The plume model has the requirement that both basalt and residue ages should coincide, whereas this is only the case in some cratons and not in others. The ‘hot ridge’ model has the advantage that, in some cases, crust–mantle coupling may be preserved during lateral convergence of terranes to form cratonic roots as seems to have been the case for the NAC (Tappe et al., 2011; Wittig et al., 2010b). In other cases, this accretion may have decoupled the crust and mantle, for example, where the Archean mantle subcreted beneath the Paleoproterozoic Halls Creek mobile belt at the edge of the Kimberley craton, W Australia (Figure 1; Luguet et al., 2009). Some authors (e.g., Aulbach et al., 2009; Griffin et al., 2009) invoke ‘hybrid models’ for certain lithospheric mantle sections that show compositional stratification, such as the Slave craton lithosphere. In these models, the ‘ultradepleted’

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The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

shallow Slave mantle would represent low-pressure melt residues subcreted beneath the Slave cratonic crust in the Mesoarchean, whereas the deeper lithosphere represents vertically accreted plume residue from the Mackenzie Large Igneous event. The Slave mantle lithosphere does appear to be compositionally stratified (Figures 10 and 11), but this variable composition could be generated by the subduction staking of two melt-residua columns of depleted residues from a hot ridge. For both ridge-accretion and plume end-member models of cratonic lithosphere formation, there is the problem of the fate of the large volume of basaltic to komatiitic (depending on the model) melts produced during extensive melting in either environment (Arndt et al., 2009; Boyd, 1989; Herzberg and Rudnick, 2012; Pearson and Wittig, 2008; Rollinson, 2010). The volume of Archean basalt (equivalent to at least 40 km thickness of any lithosphere section) is not present in any estimate of Archean crustal structure and composition (e.g., James et al., 2001; Nguuri et al., 2001; see Chapter 4.1). Furthermore, estimates of the abundance of eclogite now residing in the cratonic lithosphere are very low, at around 1 vol% (Jacob, 2004; Schulze, 1989, 1995). This discrepancy may be because large, dense slabs of eclogite are not gravitationally stable in low-density cratonic peridotite. Percival and Pysklywe (2007) showed via dynamical modeling that a large horizontal slab of eclogite located at the base of the cratonic crust could sink rapidly through the underlying peridotite lithosphere causing a significant thermal disturbance. While this model leads to some appealing results in terms of transient thermal spikes at the base of the crust, which could generate Neoarchean granites above older but thermally disturbed peridotite, it is a greatly simplified modeling scenario. Further modeling of this type is required using a variety of realistic rheologies in three dimensions to better understand the constraints they offer. Pearson and Wittig (2008) note that any subduction stacking model for Archean lithosphere is likely to have multiple slabs of eclogite inclined at some angle from the horizontal. This situation is not one of long-term gravitational stability and it is possible that the lithosphere will tend to ‘cleanse itself’ of much of the dense eclogite (Carlson et al., 2005; Pearson and Wittig, 2008). A potential result would be a resulting thermally and physically mature craton with a root of lowdensity peridotite-dominated lithospheric mantle containing minor eclogitic pods nested within it. Significant physical and thermal instability is likely to take place in the mantle and to lower crustal sections of the evolving craton during the loss of the eclogite. An alternative solution to the basalt/eclogite problem in cratons has been proposed by Bedard (2006). This model also seeks to address the deficit in eclogitic or mafic amphibolite (Foley et al., 2002) compositions that are required as the source of the abundant Archean TTG magmatism. The premise of the model involves delamination of eclogite but beneath a thick oceanic plateau formed by plume melting. Continued delamination of eclogite induces refertilization of underlying residual peridotite, allowing further melting to take place. The additional melt ‘catalyzes’ the production of more melt from the basaltic plateau that, in turn, produces additional eclogite, which sinks into peridotite, inducing further melting and so on. While intuitively appealing and able to create the temporal

links between crust, eclogite, and peridotite required by some cratonic lithospheric sections, the model has some weaknesses. It relies on a plume setting for the melting environment, which is questionable from the earlier arguments. The continued refertilization of residues in this process seems not conducive to the generation of large thicknesses of highly depleted dunite to harzburgite residues. In addition, Rollinson (2010) has pointed out that the Bedard (2006) model has difficulty in reconciling a suitable mafic source for the resulting TTG compositions in that it uses basalt from the Superior craton, which is strongly enriched in large-ion lithophile elements with a slight negative Nb anomaly. The basalt-TTG melting model proposed by Smithies et al. (2009) also relies on a similarly very incompatible element-enriched mafic source. In contrast, Rollinson (2010) manages to make TTG of the requisite trace element characteristics by melting an Archean tholeiitic basalt that fractionated from a picritic precursor and that was produced at a hot Archean ridge. The melting of this fractionated basalt in an Archaean subduction zone produces the TTG magmatic suite and a rutile-eclogite residue. In summary, while no current model is capable of explaining all aspects of the crust–mantle sections observed in cratons, there is a consensus that the likely mode of origin of cratonic peridotites was at the hot Archean MOR (Canil, 2004; Herzberg and Rudnick, 2012; Kelemen et al., 1998; Rollinson, 2010; Simon et al., 2007), involving relatively shallow melt generation, mostly at pressures less than 4 GPa. This conclusion is supported by the profiles of olivine composition versus depth for different cratonic lithosphere sections presented earlier. Such residual mantle, along with its attendant oceanic crust, would then need to be accreted, by convergence either with other oceanic lithosphere or with oceanic plateaux, to form the building blocks of the cratons that are seen today. Even the relatively thick oceanic lithosphere likely to be formed at a hot Archean ridge compared to the modern mantle (Herzberg et al., 2010) requires additional thickening to achieve the present-day observed thicknesses of cratonic lithosphere (McKenzie and Priestly, 2008). Pearson and Wittig (2008) note that such thickening via lateral movement is consistent with the general mismatch in the Neoarchean mode between the Os model ages of cratonic peridotites (Figure 16) and the age of TTG crust in cratons such as the Kaapvaal craton, and yet can also explain the coincidence in crust and mantle ages noted in other cratons, such as the NAC (Tappe et al., 2011; Wittig et al., 2010b). Compressional thickening of Archean lithospheric mantle (Jordan, 1978) is also consistent with the ‘block accretion’ required by the terranes of multiple ages that make up present-day cratons (e.g., the NCC – Liu et al., 2011a; Kaapvaal craton – de Wit et al., 1992; Schmitz et al., 2004; Siberian craton – Rosen, 2002), where full amalgamation can be as late as the end of the Paleoproterozoic. This ‘stitching together’ of cratonic blocks can lead to further compositional modification that blurs both the nature of the original peridotite protoliths and the timing of their formation. A problem with the accretion of former oceanic lithosphere to make the thick lithospheric building blocks for cratons is the fate of the abundant eclogite within such keels. This should be the focus of new dynamic modeling studies.

The Formation and Evolution of Cratonic Mantle Lithosphere – Evidence from Mantle Xenoliths

Acknowledgments The authors thank Rick Carlson for his unending persistence and patience in extracting this article from them and for then summoning the enthusiasm to provide a very detailed review that help focus the manuscript. They thank Jingao Liu, Kathy Mather, and Sarah Woodland for assistance in the preparation of this contribution.

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