A fresh isotopic look at Greenland kimberlites: Cratonic mantle lithosphere imprint on deep source signal

Earth and Planetary Science Letters 305 (2011) 235–248

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l

A fresh isotopic look at Greenland kimberlites: Cratonic mantle lithosphere imprint on deep source signal Sebastian Tappe a,⁎, D. Graham Pearson a, Geoff Nowell b, Troels Nielsen c, Phil Milstead b, Karlis Muehlenbachs a a b c

Department of Earth and Atmospheric Sciences, University of Alberta, 1-26 Earth Sciences Building, Edmonton, Alberta, Canada T6G 2E3 Department of Earth Sciences, Durham University, South Road, Durham, DH1 3LE, United Kingdom Geological Survey of Denmark and Greenland, Øster Voldgade 10, DK-1350, Copenhagen, Denmark

a r t i c l e

i n f o

Article history: Received 1 January 2011 Received in revised form 28 February 2011 Accepted 8 March 2011 Available online 29 March 2011 Editor: R.W. Carlson Keywords: kimberlite petrogenesis low-Cr megacrysts mantle reservoirs Sr–Nd–Hf isotope geochemistry U–Pb perovskite geochronology North Atlantic craton

a b s t r a c t The North Atlantic craton in West Greenland and northern Labrador has been subjected to deep volatile-rich melting events between ca. 610 and 550 Ma that produced compositionally diverse diamond-bearing kimberlite and aillikite magmas. Whereas kimberlite dyke intrusions appear to be restricted to the Maniitsoq area in the craton interior between 568 and 553 Ma, aillikite/carbonatite intrusives preferentially occur at Paleoproterozoic mobile belts such as in the Sarfartoq area (605–550 Ma) of West Greenland. Although there is an overlap between the major and trace element compositions of the exceptionally fresh Maniitsoq kimberlites and Sarfartoq aillikites, the latter typically show higher TiO2, Al2O3, and K2O, as well as higher Zr, Hf, Cs and Rb contents. Furthermore, the Sarfartoq aillikites are displaced toward lower εHf by ~ 3 epsilon units at similar εNd compared with the isotopically depleted Maniitsoq kimberlites and their garnet and ilmenite megacrysts. The generally lower εHf of aillikites corresponds to lower CO2/K2O and points to the involvement of a K-rich melt component in aillikite genesis, most likely derived from a cratonic metasome. In contrast, the Maniitsoq kimberlite compositions, in particular the high CO2 as well as low Al2O3 and K2O contents, resemble published carbonate-rich melt compositions that were produced experimentally from carbonated peridotite in excess of 6 GPa, i.e., under sublithospheric conditions. By utilizing published highpressure carbonated peridotite/melt trace element partition coefficients, we demonstrate that many of the hallmark geochemical features of kimberlites, such as relative Zr–Hf depletions, can be produced by lowdegree partial melting of carbonated fertile peridotite within the asthenosphere. For the Greenland-Labrador Diamond Province, we propose that a common asthenosphere-derived carbonated silicate melt component must have been present throughout the North Atlantic craton base at 610-to-550 Ma. This widespread carbonate-rich melt component variably interacted with old phlogopitebearing cratonic metasomes, giving rise to diverse suites of aillikites, i.e., hybrid carbonated potassic-silicate magmas, that locally separated out carbonate fractions to form intrusive carbonatites at crustal levels. The kimberlites, however, appear to be mixtures of this asthenosphere-derived carbonate-rich melt component and entrainment of materials from the refractory cratonic mantle lithosphere, with little or no involvement of readily fusible phlogopite-rich metasomes. The model developed herein for West Greenland highlights the importance of cratonic mantle lithosphere in exerting a major control on worldwide kimberlitic magma compositions. Moreover, the ability to examine kimberlite magma compositional variability in time and space clearly shows that decoupled Nd–Hf isotope systematics cannot be taken unconditionally as a reliable fingerprint of ultra-deep mantle processes, because this type of signal can also be imparted to kimberlitic melts by interaction with cratonic metasomes. © 2011 Elsevier B.V. All rights reserved.

1. Introduction The origin of kimberlite magma is a long-standing matter of debate. Among the most puzzling aspects are the nature of primary kimberlite melt compositions and their relationships to other mantle-derived

⁎ Corresponding author. Tel.: + 1 780 4928668; fax: +1 780 4922030. E-mail address: [email protected] (S. Tappe). 0012-821X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2011.03.005

volatile-rich magmas such as carbonatites and aillikites. Whereas knowledge of primary kimberlite melt compositions may provide the basis for high-pressure phase equilibrium studies on the mineral assemblage(s) of kimberlite source regions (Girnis et al., 1995; Kesson et al., 1994; Kjarsgaard et al., 2009; Mitchell, 2004), recognition of the range of primary isotopic signals in kimberlite and related magmas is vital for improved understanding of source-forming processes within a dynamic mantle through time (Carlson et al., 2006; Donnelly et al., 2011; Lehmann et al., 2010; Nowell et al., 2004; Tappe et al., 2007). Over

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the past few decades, a number of mantle reservoirs have been invoked in the origin of kimberlite magmas with options varying from ancient subducted oceanic crust trapped in the mantle transition zone (Nowell et al., 2004; Ringwood et al., 1992), the asthenosphere (Le Roex et al., 2003; Paton et al., 2009; Smith, 1983; Wyllie, 1980), and/or metasomatized cratonic mantle lithosphere (Griffin et al., 2000; Tainton and McKenzie, 1994). However, relatively little attention has been given to potential interactions between melts derived from these different mantle reservoirs and their relative roles in forming kimberlite and related magmas. In West Greenland and Labrador, a wide compositional spectrum of 610-to-550 Ma old diamondiferous kimberlites and aillikites, as well as carbonatites, intrudes the Precambrian crust of the North Atlantic craton (Larsen and Rex, 1992; Nelson, 1989; Nielsen et al., 2009; Tappe et al., 2006, 2008). Their co-occurrence has been explained by variable degrees of partial melting of a common carbonated peridotite upper mantle source region, i.e., by a primary melting relationship (Dalton and Presnall, 1998; Foley et al., 2009; Mitchell, 2005; Woolley and Kjarsgaard, 2008). Recently, Gaffney et al. (2007) suggested preconditioning of such a peridotite source by volatile-rich melts derived from deeply subducted oceanic crust, including carbonate-bearing sediments, similar to mechanisms invoked in the formation of oceanic island basalt (OIB) source regions (cf., Tachibana et al., 2006). However, more recent petrological investigations indicate a more complex petrogenesis of the kimberlite–aillikite–carbonatite association in the GreenlandLabrador Diamond Province (GLDP) (Nielsen et al., 2009; Tappe et al., 2008, 2009). For example, several lines of evidence, including bulk-rock and mineral compositions, suggest that the carbonatites do not represent primary melts, but instead show characteristics more typical of differentiated magmas and their cumulates, such as low bulk-rock Mg/Ca ratios and fairly Fe-rich olivine compositions (bFo88; Tappe et al., 2009). Furthermore, carbon isotope studies demonstrate that the carbonate fractions of these volatile-rich magmas have a primordial mantle origin and, thus, question the involvement of subducted carbonate sediments in the genesis of kimberlite and related magmas in the GLDP (Tappe et al., 2006, 2008). These and other inconsistencies warrant a ‘fresh look’ at Greenland kimberlite, aillikite, and carbonatite magmatism with the general goal of better understanding the origin of these magma types and their genetic relationships on a worldwide basis. In this study we report new precise U–Pb and Rb–Sr geochronology, geochemistry, and Sr–Nd–Hf–C–O isotope data for a large suite of kimberlite, aillikite, and carbonatite sheets from southern West Greenland (Sarfartoq and Maniitsoq fields). We complement these data with the first Nd–Hf isotopic analyses of low-Cr garnet and ilmenite megacrysts associated with Greenland kimberlites, plus Sr– Nd isotopic data on groundmass perovskite to better constrain potential contributions from different mantle reservoirs. Our proposed model suggests that kimberlites, aillikites and carbonatites in West Greenland and Labrador can be related to a common carbonaterich precursor magma derived from a sublithospheric mantle source region, but that their distinctive compositional characteristics developed within the cratonic lithosphere. 2. Background 2.1. North Atlantic craton of West Greenland Kimberlite, aillikite, and carbonatite samples analyzed during the course of this study come from two areas on the North Atlantic craton (NAC) of southern West Greenland: (1) the Maniitsoq area is located well inside the NAC and (2) the Sarfartoq area straddles the boundary between the northern NAC and the Paleoproterozoic Nagssugtoqidian mobile belt (Fig. 1). Recent studies on mantle peridotite xenoliths (Wittig et al., 2008, 2010) from the kimberlite and aillikite dykes investigated here, indicate that the northern NAC mantle lithosphere beneath both the Maniitsoq and Sarfartoq areas stabilized during the

Late Archean (Rhenium-depletion model ages between 3.2 and 2.7 Ga) in a forearc-like setting. These peridotite model ages are in good agreement with previously recognized collision events and major subduction-related TTG magmatism in the crust of southern West Greenland at approximately 3.0–2.72 Ga (Windley and Garde, 2009, and references therein). These Late Archean subduction and collision events between a number of Early Archean terranes led to the amalgamation of the landmass that is commonly referred to as the NAC, which also encompasses the Early Archean crustal domains of northern Labrador and Scotland (Bridgwater et al., 1973). The NAC was subsequently incorporated into the Laurentia supercraton via accretion with neighboring continental terranes such as the Superior craton between ca. 2.0 and 1.7 Ga (Hoffman, 1988). The Paleoproterozoic collision zones such as the Nagssugtoqidian orogen in West Greenland and the Torngat orogen in northern Labrador form high-grade mobile belts, along which the Archean crust had been reworked and juvenile components added. The peridotite xenoliths Re-depletion model ages from the Sarfartoq area clearly suggest a northward continuation of Archean NAC mantle underneath the Paleoproterozoic Nagssugtoqidian mobile belt (Wittig et al., 2010; Fig. 1). However, disturbed model ages of ca. 2.0 Ga suggest metasomatic overprinting and abundant phlogopite-bearing mantle peridotite xenoliths indicate pronounced K-metasomatism beneath Sarfartoq (Larsen and Garrit, 2005). Peridotite thermobarometry shows that there is no discernable difference between the Maniitsoq and Sarfartoq cratonic mantle sections in terms of lithosphere thickness (maximum thickness of ~220 km) and geothermal gradient (model heatflow of ~38 mW/m2) at ca. 600 Ma (Sand et al., 2009), ruling out impingement of a Late Neoproterozoic mantle plume (cf., Tachibana et al., 2006). The P–T estimates also suggest that some 75 km of the NAC lithospheric keel resided within the diamond stability field, consistent with the recovery of diamonds from Late Neoproterozoic kimberlites and aillikites in southern West Greenland (Grütter and Tuer, 2009; Hutchison and Frei, 2009; Steenfelt et al., 2009). 2.2. Kimberlites, aillikites, and carbonatites of the Greenland-Labrador Diamond Province Late Neoproterozoic kimberlite, aillikite, and carbonatite magmatic activity occurred in both West Greenland and northern Labrador between ca. 610 and 550 Ma (Fig. 2) (Secher et al., 2009; Tappe et al., 2006, 2008). A large number of the kimberlite and aillikite dyke and sheet intrusions tested positively for diamonds throughout the NAC (Digonnet et al., 2000; Hutchison and Frei, 2009; Tappe et al., 2008). This suggests a common deep mantle origin for these compositionally diverse volatile-rich ultramafic magmas and possibly an intimate genetic link between them within the ‘Greenland-Labrador Diamond Province’. For West Greenland, detailed petrographic and mineralogical investigations by Nielsen et al. (2009) revealed that bonafide kimberlite dykes are restricted to the Maniitsoq area, whereas aillikite dykes predominantly occur in the Sarfartoq area (Fig. 1). Taking these new findings into account, we re-classify the ‘kimberlites’ previously reported from the Sarfartoq area in isotopic studies (Gaffney et al., 2007; Nelson, 1989) as aillikites. Aillikites typically contain clinopyroxene in a carbonate-bearing groundmass and their groundmass spinel and phlogopite compositions show strong Fe- and Ti-enrichment trends, which are not seen in kimberlites (Tappe et al., 2005). Our geochemistry and isotope study on the Maniitsoq kimberlites and Sarfartoq aillikites and carbonatites is based on the mineralogical study by Nielsen et al. (2009), and the samples analyzed represent a subset of the suite investigated by Nielsen and co-workers. Unless otherwise stated, all subsequent use of the term kimberlite refers to Group-I kimberlite. 3. New age constraints A summary of the timing of kimberlite, aillikite, and carbonatite magma emplacement in southern West Greenland is given in Secher

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Fig. 1. Location of the Late Neoproterozoic Sarfartoq aillikite and Maniitsoq kimberlite fields in southern West Greenland at the margin and well inside the North Atlantic craton, respectively. The similar old central-complex Sarfartoq carbonatite intrusion is shown as open circle (not to scale). Samples discussed in this paper (black squares — this study and black circles — Gaffney et al. (2007)) are superimposed over all known kimberlite, aillikite and carbonatite occurrences from the region (based on Steenfelt et al., 2009). The open diamond depicts the occurrence of low-Cr garnet and ilmenite megacrysts along the Majuagaa kimberlite dyke discussed herein. The age dates refer to the U–Pb perovskite and Rb–Sr phlogopite analyses that are summarized in Table 1. Geographic coordinates for all samples studied here are listed in Table 2.

et al. (2009), and in Fig. 2 we have compiled all available high-quality ages from the GLDP. Our new U–Pb perovskite and Rb–Sr phlogopite data are reported in Table 1 and displayed in Supplementary file A (A.1–A.3). The analytical techniques are described in Tappe et al. (2009) and uncertainties are quoted at the 2-sigma level.

age determination of the carbonatite phase of the Sarfartoq intrusion agrees with the previously reported Rb–Sr age of 564.8 ± 4.9 Ma (Secher et al., 2009) and supports earlier notions that carbonatite and aillikite magmatism at Sarfartoq was coeval (Hutchison and Frei, 2009; Larsen and Rex, 1992; Secher et al., 2009).

3.1. Sarfartoq field (ca. 605–550 Ma) 3.2. Maniitsoq field (ca. 568–553 Ma) More than 20 high-precision U–Pb perovskite ages exist for the ~ 200 known aillikite dykes and sheets of the Sarfartoq field and indicate magma emplacement for a protracted time period between ca. 605 and 550 Ma (Fig. 2). A new U–Pb perovskite age was determined for aillikite dyke 444281 located ~ 20 km NNW of the Sarfartoq carbonatite intrusion. A multi-grain perovskite fraction yielded a 206Pb/238U age of 577.2 ± 4.8 Ma. This age is identical within uncertainty to the median emplacement age of 577.8 Ma for the Sarfartoq field, which is based on high-precision age determinations for 22 aillikite and carbonatite dykes (Fig. 2). Rb–Sr isochrons were calculated for aillikite sheet 266511 (578.7 ± 3.8 Ma) and carbonatite sheet 483828 (572.8 ± 3.8 Ma), both forming part of the centralcomplex Sarfartoq carbonatite intrusion sensu stricto (Fig. 1). The new

The Maniitsoq kimberlite field is located ~100 km south of Sarfartoq and no Late Neoproterozoic carbonatite intrusion is known from this area (Fig. 1). Previous age determinations on three kimberlite dykes suggested magma emplacement between ca. 566 and 555 Ma (Secher et al., 2009). We have further improved age constraints, in particular for the Majuagaa kimberlite dyke. Three multi-grain perovskite fractions from Majuagaa kimberlite sample 491720 yielded statistically identical 206 Pb/238U ages (Table 1; Supplementary file A.1), with a weighted average 206Pb/238U age of 558.5 ± 1.2 Ma. This concordant age improves upon the less robust 206Pb/238U age of 565.9 ± 2.0 Ma for a different segment of the Majuagaa kimberlite dyke (#491718) cited by Secher et al. (2009). A six-point Rb–Sr isochron for kimberlite dyke 483862

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(~20 km NW of the Majuagaa locality) yielded an emplacement age of 556.5 ± 3.6 Ma, identical to the new Majuagaa kimberlite age. Although much less information is available on kimberlite magma emplacement in the Maniitsoq field, the currently known dates suggest that kimberlite magmatism in the Maniitsoq area commenced some 30-to-40 Myr after initiation of aillikite and carbonatite magmatism in other parts of the NAC such as in the Sarfartoq area of West Greenland and the Torngat Mountains of northern Labrador (Fig. 2). Our new highprecision data refute the suggestion by Nielsen and Sand (2008) that Maniitsoq kimberlite magmatic activity (ca. 568–553 Ma) was coeval with the intrusion of the Sarfartoq carbonatite (572.8 ± 3.8 Ma). 4. Results Bulk-rock chemical compositions are given in Supplementary file B.1 and the isotope data are listed in Table 2, as well as online (B.2). Descriptions of the analytical techniques are available in Supplementary file C. 4.1. Major and trace element compositions The Maniitsoq kimberlites have high MgO (26.4–38.7 wt.%), Ni (591– 1962 ppm), Cr (1069–2453 ppm), and CO2 (6.4–20.3 wt.%) contents. They are SiO2 (16.7–30.3 wt.%), TiO2 (1.0–2.6 wt.%), Al2O3 (0.57– 1.73 wt.%), K2O (0.05–0.87 wt.%), and Na2O (0.01–0.18 wt.%) depleted. The Majuagaa dyke shows significant TiO2 enrichment (Fig. 3A; 3.1– 4.0 wt.%) due to abundant ilmenite megacryst fragments. Sarfartoq aillikites have high MgO (20.4–35.4 wt.%), Ni (375–1750 ppm), and Cr (750–1609 ppm), broadly overlapping the Maniitsoq kimberlite compositions. Their SiO2 (23.0–30.3 wt.%) and Na2O (b0.15 wt.%) contents also overlap the Maniitsoq kimberlite compositions. Sarfartoq aillikites typically show elevated TiO2 (1.5–2.9 wt.%), Al2O3 (1.17–3.96 wt.%), and K2O (0.54–2.87 wt.%), but lower CO2 (4.9–13.0 wt.%) contents if compared to the Maniitsoq kimberlites (Fig. 3). The two Sarfartoq carbonatite samples included in this study have elevated MgO (16.3– 17.5 wt.%) and low SiO2 (11.9–12.2 wt.%).

The primitive mantle normalized incompatible element abundance patterns in Fig. 4 demonstrate the strong incompatible element enrichment of the Maniitsoq kimberlites and Sarfartoq aillikites. Ba, Th, U, Nb, Ta, and LREE concentrations typically exceed 100 times primitive mantle. Pronounced relative depletions in the normalized patterns are apparent at K, Pb, P, Zr–Hf, and the HREE, for which concentrations drop to below 3 times primitive mantle. Although the Maniitsoq kimberlite and Sarfartoq aillikite trace element patterns are broadly similar, aillikites tend to have higher Cs–Rb contents (Fig. 4A– B). Furthermore, Maniitsoq kimberlites show a pronounced trough at Zr–Hf, which gradually ‘shallows’ toward the Sarfartoq aillikite patterns, in agreement with on average higher Ti contents. The two Sarfartoq carbonatite samples show strongly fractionated and rather ‘erratic’ trace element patterns, which only partly resemble the patterns of the associated aillikite dykes (Fig. 4B). A comparison of the Maniitsoq kimberlite and Sarfartoq aillikite trace element patterns with the coeval aillikites of Labrador shows good agreement with respect to their overall enriched nature. However, the Labrador aillikites show even stronger incompatible element enrichment that is particularly pronounced for LILE (Cs–Rb–K), LREE, and Zr– Hf–Ti. The Zr–Hf trough of the Maniitsoq kimberlites vanishes through Sarfartoq aillikites to the aillikites of Labrador (Fig. 4C). Although much more enriched in highly incompatible elements than OIBs, the Maniitsoq kimberlites and Sarfartoq aillikites show Ce/Pb (28 ± 9 and 45± 15; Fig. 3C) and Nb/U (65 ± 10 and 52 ± 11) ratios that are more similar to intraplate oceanic magmas (Ce/Pb = 25± 5; Nb/U = 47± 10; Hofmann et al., 1986) than to subduction-influenced mafic alkaline magmas such as orogenic lamproites (Ce/Pb and Nb/Ub 10; Prelevic et al., 2008). 4.2. Sr–Nd–Hf isotope compositions Unless a radiometric age exists for a given sample, the initial Sr, Nd, and Hf isotope compositions were calculated using the median emplacement ages (see Fig. 2 and Table 2 for relevant parameters). We have also recalculated the initial Nd and Hf isotope ratios for the

Fig. 2. Summary of the new U–Pb perovskite and Rb–Sr phlogopite age determinations for the kimberlite, aillikite and carbonatite dykes/sheets from West Greenland (S — Sarfartoq and M — Maniitsoq). The new age results are shown in conjunction with published high-precision 206Pb/238U perovskite ages for Late Neoproterozoic kimberlites (K), aillikites (A) and carbonatites (C) from West Greenland (Secher et al., 2009) and northern Labrador in Canada (Tappe et al., 2006, 2008). Note that the vast majority of age determinations for the West Greenland ultramafic dykes represent the Sarfartoq field (median age of 577.8 Ma, n = 22), whereas only 3 age determinations exist for the Maniitsoq kimberlite field (median age of 558.3 Ma, n = 3). The kimberlite magmatic activity in the Maniitsoq area appears to be restricted to the ‘young’ end of the known age spectrum of the Greenland-Labrador Diamond Province (see inset map). Error bars correspond to 2-sigma uncertainties. See Table 1 and Supplementary files A.1–A.3 for details of the new geochronology data.

Table 1 ID-TIMS U–Pb perovskite and Rb–Sr phlogopite results for kimberlite, aillikite and carbonatite dykes/sheets from the Sarfartoq and Maniitsoq fields, North Atlantic craton, West Greenland. Description*

U (ppm)

Th (ppm)

Pb (ppm)

TCPb (pg)

Apparent ages (Ma) **

206

Pb/

238

U

207

**

235

Pb/

U

**

207

Pb/

206

Pb

206

Pb/238U

207

Pb/235U

207

Pb/206Pb

Discordance (%)

− 75.7

10

55

2146

140

34

0.09366 ± 80

0.6863 ± 338

0.05314 ± 256

577.2 ± 4.8

530.6 ± 20.2

335 ± 107

9

221

4621

144

59

0.09176 ± 36

0.7743 ± 212

0.06120 ± 158

565.9 ± 2.0

582.2 ± 12.2

646 ± 55

13.0

31 76 32

199 158 134

4320 3037 2167

128 93 71

139 265 188

0.09066 ± 30 0.09043 ± 34 0.09018 ± 52

0.7323 ± 84 0.7224 ± 90 0.7292 ± 152

0.05858 ± 64 0.05794 ± 72 0.05864 ± 126

559.4 ± 1.8 558.1 ± 2.0 556.6 ± 3.2 WA 558.5 ± 1.2

557.9 ± 5.0 552.1 ± 5.4 556.1 ± 9.0

552 ± 24 527 ± 27 554 ± 46

− 1.5 − 6.1 − 0.5

Sample description Aillikite sheet 266511 (Sarfartoq: − 51.2473000; 66.5100800) a 578.7 ± 3.8 Ma (2σ), MSWD = 0.064, 87Sr/86Sri = 0.70299 ± 7 A. phlogopite phenocryst, light brown, b1×1 × 1 mm B. phlogopite phenocryst, light brown, b1 × 1 × 1 mm C. phlogopite phenocryst, light brown, b1 × 1 × 1 mm D. phlogopite phenocryst, light brown, b 1 × 1 × 1 mm E. phlogopite phenocryst, light brown, b 1 × 1 × 1 mm Bulk-rock Carbonatite sheet 483828 (Sarfartoq: − 51.2303886; 66.5066142) a 572.8 ± 3.8 Ma (2σ), MSWD = 0.72, 87Sr/86Sri = 0.70362 ± 4 A. phlogopite phenocryst, reddish brown, b 2 × 2 × 1 mm B. phlogopite phenocryst, reddish brown, b3 × 2 × 1 mm C. phlogopite phenocryst, reddish brown, b3 × 2 × 1 mm D. phlogopite phenocryst, light brown, b 3 × 3 × 1 mm E. phlogopite phenocryst, light brown, b 3 × 3 × 1 mm Bulk-rock Kimberlite dyke 483862 (Maniitsoq: − 52.4009161; 65.3882298) a 556.5 ± 3.6 Ma (2σ), MSWD = 0.63, 87Sr/86Sri = 0.70295 ± 3 A. phlogopite phenocryst, dark brown, b 10 × 5 × 2 mm B. phlogopite phenocryst, dark brown, b 10 × 5 × 2 mm C. phlogopite phenocryst, greenish brown, b3 × 2 × 1 mm D. phlogopite phenocryst, greenish brown, b3 × 2 × 1 mm E. phlogopite phenocryst, light brown, b 6 × 3 × 2 mm Bulk-rock

Rb (ppm)

Sr (ppm)

87

Rb/86Sr

***2σm

Sr/86Sr

***2σm

87

270.4 217.3 257.7 241.3 241.0 123.1

11.55 15.99 10.97 9.45 6.43 775.1

71.56 40.57 71.82 78.49 118.7 0.4485

± 1.07 ± 0.61 ± 1.08 ± 1.18 ± 1.8 ± 0.0067

1.284870 1.033480 1.288500 1.341200 1.672500 0.706638

± 0.000064 ± 0.000052 ± 0.000064 ± 0.000067 ± 0.000084 ± 0.000035

328.1 333.5 325.3 349.6 331.2 96.04

3.92 1.55 12.75 2.66 1.41 3945

297.2 1115 78.36 535.3 1301 0.0687

± 4.5 ± 17 ± 1.18 ± 8.0 ± 20 ± 0.0010

3.082190 9.7650 1.338680 5.002270 11.179600 0.704178

± 0.000154 ± 0.0013 ± 0.000067 ± 0.000250 ± 0.001300 ± 0.000035

794.7 790.9 740.9 560.3 770.5 3.93

29.42 29.76 14.30 20.07 19.87 2078

83.11 81.64 169.4 86.07 122.7 0.0053

± 1.25 ± 1.22 ± 2.5 ± 1.29 ± 1.8 ± 0.0001

1.354480 1.339770 2.042050 1.377310 1.657980 0.702996

±0.000068 ±0.000067 ±0.000102 ±0.000069 ±0.000083 ±0.000035

S. Tappe et al. / Earth and Planetary Science Letters 305 (2011) 235–248

Aillikite dyke 444281 (Sarfartoq: − 51.5469343; 66.6861856) 1. Black cubes/fragments; [email protected] (100) Kimberlite dyke 491718 (Majuagaa) (Maniitsoq: − 51.9998278; 65.2214971) 1. Dark brown cubes; [email protected] (29) Kimberlite dyke 491720 (Majuagaa) (Maniitsoq: − 51.9998278; 65.2214971) 1. Dark brown cubes; [email protected] (80) 2. Dark brown cubes; [email protected] (120) 3. Dark brown cubes; [email protected] (85)

Weight (μg)

*[email protected] — perovskite grains selected from magnetic fraction at 0.5 A (Frantz); numbers in parentheses are numbers of grains analyzed; geographic coordinates are in decimal degrees using WGS84 datum. **Atomic ratios corrected for fractionation (0.105%/amu Pb and 0.123%/amu U), blank (5 pg Pb; 1 pg U), isotopic tracer, and initial common Pb. ***Standard 2σ uncertainties of ±1.5% for 87Rb/86Sr and ±0.005% for 87Sr/86Sr are assigned to the reported isotope ratios. See Supplementary file C.1 for discussion. Thorium concentrations calculated based on amount of 208Pb present and 207Pb/206Pb model age; TCPb is estimated total initial common Pb based on the Stacey and Kramers (1975) terrestrial Pb evolution model. Note that the U–Pb perovskite age for kimberlite 491718 is cited in Secher et al. (2009), but no data were previously reported. a Isochron Model 1 ages were calculated with Isoplot using a decay constant of 1.402 ⁎ 10− 11 a− 1 for 87Rb (Minster et al., 1982); all uncertainties in this table are quoted at 2-sigma.

239

240

Table 2 Sr–Nd–Hf–C–O isotope composition of kimberlite, aillikite and carbonatite dykes/sheets, as well as of groundmass perovskites and low-Cr megacrysts, North Atlantic craton, West Greenland. 143

n.a. n.a. n.a. n.a. 0.703698(8) 0.703427(8) 0.703371(11) 0.703654(8) 0.703368(8) 0.703910(9) 0.706638(8) 0.702845(50) 0.705060(9) 0.704178(7)

– – – – 0.70286 0.70266 0.70238 0.70248 0.70251 0.70282 0.70299 0.70285 0.70481 0.70362

0.703897(13) 0.704361(15) 0.703732(16) 0.704070(8) 0.705401(9) 0.704474(9) 0.705216(11) 0.704747(12) 0.705139(10) 0.703641(9) 0.703986(9) 0.703883(7) 0.702996(10) 0.706608(6) 0.703426(8) 0.704422(8) 0.703897(9) 0.704049(9) 0.704003(8) 0.704108(9) 0.703815(8) 0.702794(17) 0.702803(29) n.a. n.a. n.a. n.a. n.a. n.a.

0.70260 0.70311 0.70299 0.70276 0.70386 0.70417 0.70454 0.70464 0.70455 0.70331 0.70313 0.70352 0.70295 0.70603 0.70275 0.70368 0.70331 0.70333 0.70300 0.70346 0.70331 0.70279 0.70280 – – – – – –

Type

Age (Ma)

Longitude

Latitude

87

Sarfartoq field 474514 474515 474516 474517 474560 488810 488822 488823 488825 P-Dyke 266511# 444281-1# 477425# 483827/28#

Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Aillikite Grdm perovskite Carbonatite Carbonatite

577.8 577.8 577.8 577.8 577.8 577.8 577.8 577.8 577.8 577.8 578.7 577.2 577.8 572.8

− 51.4636138 − 51.1469167 − 51.1469167 − 51.1469167 − 51.1682222 − 51.1525278 − 50.9763889 − 50.9462500 − 50.9763889 − 51.4636138 − 50.9763889 − 51.5469343 − 51.2473000 − 51.8495454

66.3530532 66.7301389 66.7301389 66.7301389 66.7439444 66.7410278 66.4871389 66.4746111 66.4871389 66.3530532 66.4871389 66.6861856 66.5100800 66.4044030

Maniitsoq field 483838 483840 483842 483843 483844 483845 483847 483848 483849 483857 483860 483861 483862 483863 483864 488582# M 491702# M 491708# M 491712# M 491722/20# M 491741# M 491720-2# M 491720-3# M 477410-ilm1 M 477410-ilm3 M 477412-ilm1 M 477412-grt1 M 477413-ilm1 M 477413-grt1

Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Kimberlite Grdm perovskite Grdm perovskite Ilm megacryst Ilm megacryst Ilm megacryst Grt megacryst Ilm megacryst Grt megacryst

558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 556.5 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3 558.3

− 52.0622301 − 52.0253501 − 51.7058803 − 51.7009071 − 51.6969592 − 51.7075937 − 51.7054775 − 51.7055186 − 51.6756639 − 51.5624339 − 51.5566915 − 51.0808631 − 52.4009161 − 52.2447033 − 52.1806395 − 51.6127990 − 51.9848940 − 51.9924354 − 51.9945019 − 51.9839207 − 51.9641854 − 51.9998278 − 51.9998278 − 51.9731229 − 51.9731229 − 51.9731229 − 51.9731229 − 51.9821409 − 51.9821409

65.0850064 65.0878718 65.2930708 65.2935580 65.2925378 65.2931242 65.2930832 65.2928204 65.2971722 65.1949045 65.1955641 65.2986155 65.3882298 65.3011030 65.0850076 65.3065100 65.2240632 65.2226266 65.2222541 65.2238429 65.2257002 65.2214971 65.2214971 65.2251400 65.2251400 65.2251400 65.2251400 65.2240273 65.2240273

Sr/86Srm

*143Nd/144Ndi

**(εNd)i

176

0.512394(5) 0.512406(7) 0.512405(5) 0.512345(7) 0.512313(7) 0.512289(6) n.a. 0.512292(6) n.a. 0.512279(8) 0.512327(5) 0.512293(15) 0.512385(7) 0.512394(7)

0.51206 0.51207 0.51206 0.51204 0.51202 0.51200 – 0.51200 – 0.51200 0.51203 0.51201 0.51203 0.51204

3.3 3.5 3.3 2.9 2.4 2.1 – 2.2 – 2.2 2.7 2.3 2.7 2.7

0.512336(12) 0.512325(8) 0.512347(17) 0.512363(11) 0.512321(8) 0.512307(8) 0.512331(8) 0.512338(11) 0.512287(6) 0.512309(4) 0.512331(9) 0.512322(7) 0.512334(8) 0.512324(7) 0.512352(5) 0.512325(8) 0.512338(5) 0.512321(5) 0.512339(5) 0.512333(3) 0.512339(5) 0.512299(9) 0.512303(7) 0.512330(12) 0.512342(10) 0.512343(15) 0.514184(29) 0.512329(6) 0.513792(34)

0.51207 0.51206 0.51208 0.51210 0.51206 0.51206 0.51202 0.51206 0.51209 0.51205 0.51206 0.51205 0.51207 0.51204 0.51208 0.51205 0.51205 0.51201 0.51203 0.51205 0.51206 0.51208 0.51209 0.51206 0.51208 0.51207 0.51213 0.51207 0.51211

3.0 2.7 3.2 3.6 2.7 2.8 2.0 2.8 3.3 2.6 2.7 2.6 2.9 2.3 3.2 2.6 2.5 1.9 2.3 2.6 2.7 3.2 3.3 2.8 3.1 3.0 4.1 2.9 3.8

Nd/144Ndm

*176Hf/177Hfi

***(εHf)i

ΔεiHf

δ13CPDB (‰)

δ18OSMOW (‰)

0.282542(3) 0.282541(3) 0.282549(4) 0.282533(5) 0.282529(5) 0.282522(5) 0.282541(13) 0.282531(6) 0.282536(11) 0.282506(4) 0.282501(6) n.a. 0.282605(10) 0.282535(9)

0.28252 0.28252 0.28253 0.28250 0.28249 0.28250 0.28251 0.28250 0.28252 0.28247 0.28248 – 0.28247 0.28248

3.6 3.7 3.8 3.0 2.5 2.7 3.3 2.8 3.5 1.8 2.3 – 1.9 1.8

−4.2 −4.3 −3.9 −4.2 −4.0 −3.4 – −3.4 – −4.3 −4.6 – −5.0 −5.1

−4.89(1) n.a. n.a. −3.86(4) −4.43(1) n.a. −4.41(2) n.a. n.a. n.a. −4.05(1) n.a. −3.20(1) −3.22(2)

8.67(3) n.a. n.a. 10.64(5) 10.21(3) n.a. 9.25(4) n.a. n.a. n.a. 16.44(2) n.a. 12.40(1) 11.63(2)

0.282561(8) 0.282553(4) 0.282543(5) 0.282539(9) 0.282528(11) 0.282555(7) n.a. 0.282514(9) 0.282564(5) 0.282578(4) 0.282585(5) 0.282553(4) 0.282602(8) 0.282595(6) 0.282597(4) 0.282618(12) 0.282591(14) 0.282556(9) 0.282565(7) 0.282561(8) 0.282570(9) n.a. n.a. 0.282568(5) 0.282565(4) 0.282551(10) 0.282842(18) 0.282558(4) 0.282790(16)

0.28254 0.28253 0.28252 0.28253 0.28250 0.28253 – 0.28249 0.28255 0.28256 0.28257 0.28253 0.28258 0.28257 0.28258 0.28260 0.28257 0.28252 0.28254 0.28254 0.28255 – – 0.28257 0.28256 0.28255 0.28256 0.28256 0.28255

3.6 3.4 3.1 3.3 2.4 3.5 – 2.0 4.2 4.5 4.9 3.4 5.1 5.0 5.0 5.8 4.7 3.2 3.9 3.9 4.2 – – 4.8 4.7 4.1 4.7 4.3 4.2

−3.6 −3.4 −4.5 −4.9 −4.5 −3.5 – −5.0 −3.5 −2.2 −1.9 −3.3 −2.0 −1.4 −2.5 −1.0 −1.9 −2.6 −2.4 −2.8 −2.7 – – −2.2 −2.8 −3.2 −4.2 −2.8 −4.2

n.a. n.a. n.a. n.a. n.a. n.a. n.a. −4.30(2) n.a. −4.43(1) −4.27(1) n.a. −5.28(2) −4.06(1) n.a. −4.05(1) −5.02(1) n.a. n.a. −4.83(1) n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

n.a. n.a. n.a. n.a. n.a. n.a. n.a. 13.42(2) n.a. 12.97(3) 12.50(7) n.a. 10.38(2) 10.57(2) n.a. 10.82(1) 10.29(4) n.a. n.a. 12.07(3) n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

Hf/177Hfm

Sr, Nd, and Hf isotope compositions were determined by PIMMS, except for samples with superscript “#”, for which Sr and Nd isotope compositions were determined by TIMS.

Measured Sr, Nd, and Hf isotope ratios are normalized to the following standard values: NBS987 = 87Sr/86Sr value of 0.71024 (Thirlwall, 1991); J&M = 143Nd/144Nd value of 0.511110, equivalent to the La Jolla 143Nd/144Nd value of 0.51185; and JMC475 = 176Hf/177Hf value of 0.28216 (Blichert-Toft et al., 1997). *Initial isotope ratios calculated for the Sarfartoq median age of 577.8 Ma and the Maniitsoq median age of 558.3 Ma, unless a high-precision radiometric age exists for an individual sample (see Figs. 1 and 2). **Initial epsilon Nd values were calculated using 147Sm decay constant of 6.54 ⁎ 10− 12 a− 1 (Lugmair and Marti, 1978); (143Nd/144Nd)CHUR = 0.512638 (Goldstein et al., 1984); (147Sm/144Nd)CHUR = 0.1967 (Jacobsen and Wasserburg, 1980). ***Initial epsilon Hf values were calculated using 176Lu decay constant of 1.865 ⁎ 10− 11 a− 1 (Scherer et al., 2001); (176Hf/177Hf)CHUR = 0.282785 and (176Lu/177Hf)CHUR = 0.0336 (Bouvier et al., 2008). ΔεiHf is defined as εiHf-(1.36 ⁎ εiNd + 3.19), such that samples with negative values fall by definition below the terrestrial Nd–Hf isotope array of Vervoort et al. (1999). Numbers in parentheses are 2-sigma-of-the-mean uncertainties for individual isotope ratio measurements; geographic coordinates are in decimal degrees using WGS84 datum. Sample numbers with superscript “M” represent the Majuagaa kimberlite dyke of the Maniitsoq field; n.a. = not analyzed.

S. Tappe et al. / Earth and Planetary Science Letters 305 (2011) 235–248

*87Sr/86Sri

Sample no.

S. Tappe et al. / Earth and Planetary Science Letters 305 (2011) 235–248

241

Fig. 3. Variation diagrams of bulk-rock TiO2 (A), Al2O3 (B), Ce/Pb (C) and εHfi (D) vs. CO2/K2O for the West Greenland kimberlite, aillikite and carbonatite samples discussed in this study. The gray and black compositional fields for experimentally derived melts in panels ‘A’ and ‘B’ are adopted from Konzett (1997) and Brey et al. (2008), respectively. Note that the elevated TiO2 contents of the Majuagaa kimberlite dyke are due to the accumulation of ilmenite macrocrysts and therefore not taken as representative of the melt composition. The indicated range of Ce/Pb = 25 ± 5 for oceanic basalts in panel ‘C’ is taken from Hofmann et al. (1986). Data for aillikites from Labrador are from Tappe et al. (2007, 2008).

Sarfartoq aillikite samples, reported as kimberlites by Gaffney et al. (2007), to an age of 577.8 Ma rather than 600 Ma. The Maniitsoq kimberlites are characterized by relatively radiogenic εNd(i) (+1.9 to +3.6) and εHf(i) (+2.0 to +5.8), with unradiogenic 87Sr/ 86 Sr(i) (0.70260–0.70464), similar to many isotopically mildly depleted South African Group-I kimberlites and modern OIBs (Fig. 5A–B). The Sarfartoq aillikites (87Sr/86Sr(i) = 0.70248–0.70299; εNd(i) = 2.1–3.5) fall within the Sr–Nd isotope compositional range of the Maniitsoq kimberlites. However, the aillikites cover the unradiogenic end of the Maniitsoq kimberlite Hf isotope compositional range (Fig. 5B) and do not exceed εHf(i) of +4.2 (εHf(i) = 1.8–3.8, this study;= 1.1–4.2, Gaffney et al., 2007). The Sarfartoq carbonatite samples (εNd(i) = +2.7; and εHf(i) = 1.8–1.9) overlap the Nd–Hf isotope compositions of the aillikites, but their 87Sr/86Sr(i) values are slightly more radiogenic (0.70362 and 0.70481), causing a horizontal array in Sr–Nd isotope space similar to the Maniitsoq kimberlites (Fig. 5A). Although such wide Sr isotope compositional spread at relatively constant εNd(i) is typically ascribed to alteration (e.g., Carlson et al., 2006), we reiterate that the Greenland samples are of exceptional freshness. It is more likely that the observed range in Sr isotope compositions is dominated by uncertainties introduced during the calculation of the initial isotopic compositions. This is particularly problematic for phlogopite-bearing rock types, for which slight powder heterogeneity can result in significant deviation

from ‘true’ Rb/Sr ratios whereas the Sm/Nd and Lu/Hf systems are insensitive to this effect (note that trace element contents and isotopic compositions were determined on different splits of the same powder). Support for relatively homogeneous 87Sr/86Sr(i) is provided by the unradiogenic values of groundmass perovskites (Fig. 5A), which yielded 0.70279 ± 2 for a Maniitsoq kimberlite (#491720) and 0.70285 ± 5 for a Sarfartoq aillikite dyke (#444281). In contrast to the variability of bulkrock initial Sr isotope ratios, εNd(i) of the kimberlitic and aillikitic perovskites (+3.2 and +2.3, respectively) fall within the range of the respective bulk-rock samples, demonstrating the robustness of bulkrock initial Nd isotope ratios. Further indication of the accuracy of the bulk-rock Nd–Hf isotope compositions is provided by low-Cr garnet and ilmenite megacrysts from the Majuagaa dyke (εNd(i) = +2.8 to +4.1 and εHf(i) = +4.1 to + 4.8), which show no systematic variation between phases and overlap the Maniitsoq kimberlite analyses (Fig. 5B). Megacryst εHf(i) values substantiate the previously noted more radiogenic nature of the Maniitsoq kimberlites compared to the Sarfartoq aillikites. Six-point Sm–Nd and Lu–Hf isochrons yield crystallization ages of 576 ± 20 Ma and 553 ± 55 Ma, respectively (Supplementary files A.3 and B.3), reinforcing the argument that low-Cr megacrysts represent the highpressure stage of a kimberlite magmatic system (e.g., Bell et al., 2004; Moore and Belousova, 2005).

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Fig. 4. Primitive mantle normalized incompatible element distribution for (A) Maniitsoq kimberlites and (B) Sarfartoq aillikites and carbonatites. The two calculated trace element patterns in panel ‘C’ represent near-solidus partial melts of carbonated peridotite at 6 and 10 GPa (for details see Section 5.2); they were obtained by utilizing the experimentally determined bulk peridotite/kimberlitic melt partition coefficients of Brey et al. (2008). The gray field in all panels represents the range of trace element patterns for Labrador aillikites (Tappe et al., 2007, 2008). Primitive mantle values are from Palme and O'Neill (2003).

4.3. Carbon and oxygen isotope compositions Carbon and oxygen isotope compositions of the Maniitsoq kimberlite carbonates range between −5.3 and −4.1‰ δ13CPDB, and +10.3 and + 13.4‰ δ18OSMOW (Table 2; Fig. 6). A similar range is observed for the Sarfartoq aillikite groundmass carbonates (−4.9 to −3.9‰ δ13CPDB; + 8.7 to +16.4‰ δ18OSMOW). The aillikite data extend toward the two

Sarfartoq carbonatite samples, which have −3.2‰ δ13CPDB and +11.6 and +12.4‰ δ18OSMOW (Fig. 6). Both Maniitsoq kimberlite and Sarfartoq aillikite carbonate fractions show mantle-like δ13C (cf., Deines, 2002) and overlap the known range of their Labrador counterparts (Tappe et al., 2006, 2008). However, δ13C values of the Sarfartoq carbonatite are marginally higher than typical mantle values. Only Sarfartoq aillikite 474514 has preserved a mantle-like δ18O value and most samples follow

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243

Fig. 5. εNdi vs. 87Sr/86Sri (A) and εHfi vs. εNdi (B) for Maniitsoq kimberlites, Sarfartoq aillikites and carbonatites. Panel ‘A’ also shows Sr–Nd isotope compositions of groundmass perovskites from a kimberlite and aillikite sample. The inset of panel ‘B’ provides a more detailed sample subdivision, i.e., the Nd–Hf isotope compositions of garnet and ilmenite megacrysts from the Majuagaa kimberlite dyke are shown separately from the bulk-rock compositions. Furthermore, the Sarfartoq ‘kimberlites’ analyzed by Gaffney et al. (2007), e.g. the ‘anomalous KIM-7’ sample, are distinguished from the Sarfartoq aillikites investigated here. The open circles represent Late Neoproterozoic aillikites from Labrador (Tappe et al., 2007, 2008). Fields for Mesoproterozoic Greenland lamproites (Nelson, 1989), Mesoproterozoic Labrador lamproites (Tappe et al., 2007), Mesozoic South African kimberlites and orangeites (Nowell et al., 2004), and modern oceanic basalts (compilation retrieved from http://georoc.mpch-mainz.gwdg.de/georoc/) are shown for comparison. Terrestrial array is after Vervoort et al. (1999).

a trend toward heavier δ18O. Wilson et al. (2007) ascribed such trends in fresh Slave craton kimberlites to magma cooling processes in the presence of CO2- and H2O-rich fluids (Fig. 6). 5. Discussion 5.1. Testing the carbonatitic–kimberlitic melting continuum hypothesis The close spatial and temporal associations of kimberlite, aillikite and carbonatite magmatism in West Greenland allow constraints to be placed on the origin of kimberlite and related volatile-rich magmas. One

of the key features of this magmatism is its CO2-rich nature throughout the NAC, which points to involvement of a CO2-rich component during melting of the mantle source region. Recent petrological studies have demonstrated that both kimberlite and aillikite magmas of the GLDP have sampled diamonds and garnet peridotite fragments en route to the surface, with the deepest analyzed xenolith having last equilibrated at ~6.3 GPa, equivalent to 215 km depth (Sand et al., 2009). Entrainment of material from the base of the North Atlantic cratonic lithosphere is indicative of an ultimate sublithospheric origin of both kimberlite and aillikite magmas, and we note that the crucial petrogenetic difference between these two magma types may not be related to the

244

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Fig. 6. Carbon and oxygen isotope composition (expressed as ‰ δ13C and δ18O relative to PDB and SMOW, respectively) of bulk-rock carbonate fractions from Maniitsoq kimberlite and Sarfartoq aillikite and carbonatite dykes/sheets, West Greenland. Opensystem CO2-degassing and concomitant Rayleigh fractionation processes, as well as low-temperature magmatic fluid-related processes have the potential to change primary stable isotope compositions as illustrated by arrows (see Deines, 1989, 2002; Wilson et al., 2007). The temperature scales at the bottom refer to two closed-system isotopic fractionation models between primary magmatic carbonates and magmatic fluids with different CO2/H2O. See Section 4.3 and Wilson et al. (2007) for more details. Estimates of the C and O isotope compositions of primary mantle carbonates, as well as compositional fields for bulk-rock carbonate fractions of global carbonatites and Late Neoproterozoic aillikite and carbonatite dykes/sheets from Labrador are shown for comparison (Tappe et al., 2006, 2008, and references therein). Symbol size is much larger than the 2σ analytical uncertainty (see Table 2).

ultimate pressure of origin. We further note that mantle xenoliths are conspicuously lacking in the GLDP carbonatites including the Sarfartoq occurrence. A widely held view on the relationship between cratonic carbonatite, aillikite, and kimberlite magmas is that they form from a common carbonated peridotite source by small, but increasing, degrees of partial melting in the above listed sequence. Such a primary melting continuum was observed in a number of melting experiments on natural and simplified peridotite under diamond-stability field conditions in the presence of CO2. The composition of melts has been shown to progress from dolomitic carbonatite with b15 wt.% SiO2 near the solidus through to carbonated silicate melts (15–35 wt.% SiO2), reminiscent of aillikite and kimberlite at higher temperatures and thus higher degrees of melting (Brey et al., 2008; Dalton and Presnall, 1998; Foley et al., 2009; Gudfinnsson and Presnall, 2005). However, a closer look at the natural carbonatite compositions from Greenland reveals that they are compositionally unlike their experimentally produced analogs. For example, there is a prominent bulk-rock compositional gap between the intrusive Sarfartoq carbonatites and associated aillikites (Fig. 7), and the carbonatites show MgO/CaO ratios b1; lower than expected for primary mantle-derived carbonatite melts (Brey et al., 2008; Dasgupta et al., 2009). The Sarfartoq carbonatites are mixed dolomite–calcite carbonatites, containing variable amounts of silicate and oxide minerals (Larsen and Rex, 1992). They typically exhibit layering and flow textures, clearly indicating that they do not represent liquid compositions; rather, they are mobilized crystal cumulates formed from one or several magma batches at crustal levels. This interpretation is in agreement with their highly fractionated and variable incompatible element abundance patterns and trace element ratios such as Ce/Pb up to 397 (Fig. 3C), which are atypical for experimentally produced, high-pressure carbonatitic melts (Brey et al., 2008; Dasgupta et al., 2009; Foley et al., 2009). In contrast, the associated Sarfartoq aillikite dykes in the vicinity of the carbonatite intrusion have mineralogical characteristics of primitive magmas (Mitchell et al., 1999; Nielsen et al., 2009). Their average major element composition falls within the experimentally determined range

Fig. 7. MgO/CaO vs. SiO2/Al2O3 for the average Maniitsoq kimberlite and Sarfartoq aillikite, as well as two Sarfartoq carbonatite samples investigated here. The average Torngat aillikite composition from Labrador is adopted from Tappe et al. (2008). Fields for experimentally produced melt compositions from synthetic and natural carbonated peridotites under high pressures are after Gudfinnsson and Presnall (2005) and Brey et al. (2008), respectively. Gray crosses are reconstructed parent kimberlite magma compositions from (1) — the Majuagaa kimberlite dyke, West Greenland; (2) — South African Group-I kimberlite; (3) — Lac de Gras high-Ti kimberlite and (4) — Lac de Gras low-Ti kimberlite, Canada; (5) — Jericho kimberlite, Canada; and (6) — Udachnaya East kimberlite, Russia. For primary data sources see compilation in Kjarsgaard et al. (2009).

of primary high-pressure melts of carbonated peridotite (Fig. 7), even if much of the olivine can be ascribed to entrainment from the lithospheric mantle (cf., Arndt et al., 2010). The carbonate fraction of the Sarfartoq aillikites has mantle-like δ13C and δ18O values (Fig. 6), whereas the carbon and oxygen isotope compositions of the carbonatite are displaced toward heavier values along a high-temperature Rayleigh fractionation trend (Deines, 1989). These observations, combined with the geographic proximity and overlapping emplacement ages (Figs. 1–2), suggest that the Sarfartoq carbonatite intrusion likely represents a differentiation product derived from a carbonated silicate parent magma akin to aillikite. Identification of the Sarfartoq aillikites as primitive magmas raises the question of whether the Maniitsoq kimberlites 100 km to the south are derived from the same source region. As emphasized previously, the Maniitsoq area is the only known occurrence of archetypal kimberlite in the GLDP and is distal to Paleoproterozoic mobile belts in contrast to all known aillikite occurrences in Greenland and Labrador. In general, the Maniitsoq kimberlites have higher MgO and CO2 contents than the Sarfartoq aillikites, and lower K, Ti, and Al. In an attempt to minimize the problem of xenocrystic olivine entrainment, we use the bulk-rock CO2/ K2O ratio as an approximation to the ‘true’ melt compositions of these superbly fresh dyke samples. Fig. 3 shows that the Maniitsoq kimberlites have high CO2/K2O ratios that partly overlap the experimentally produced near-solidus melts derived from carbonated peridotite at 6 to 10 GPa (Brey et al., 2008). The extremely low Al contents of the Maniitsoq kimberlites are also in good agreement with these experimentally produced kimberlitic melts. The Sarfartoq aillikites have much lower CO2/K2O ratios than the experimentally produced kimberlitic melts of Brey et al. (2008), and they appear to be gradational to the Maniitsoq kimberlites (Fig. 3). This apparent continuous gradation, however, is inconsistent with either a primary kimberlite–aillikite or aillikite–kimberlite melting relationship of a common carbonated peridotite source, because the most CO2-rich melts nearest to the solidus should have the lowest MgO but highest K2O and TiO2 contents (e.g., Foley et al., 2009). On the contrary, we observe a pronounced increase in TiO2 and Al2O3 with decreasing CO2/K2O from Maniitsoq kimberlites through Sarfartoq aillikites to the aillikites of Labrador (Fig. 3A–B). This trend cannot be the result of contamination with continental crustal material close to Earth's surface (or within

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the mantle source), because the generally increasing K contents are accompanied by increasing Ce/Pb ratios (Fig. 3C). Instead, this compositional trend approaches melt compositions that were experimentally produced under diamond-stability field conditions from starting materials synthesized as analogs of ultramafic MARID-type material, represented by metasomatic veins in cratonic peridotite (Konzett, 1997). In terms of the evolution of source regions, the pronounced positive correlation between bulk-rock CO2/K2O ratios and the Hf isotopic compositions (Fig. 3D) indicates that the carbonate-rich, K-poor component (as identified in Maniitsoq kimberlites) was long-term depleted, whereas the K-rich component (as identified in GLDP aillikites) was isotopically more enriched. The cratonic mantle lithosphere is known to host phlogopite-rich (and by inference K-rich), long-term enriched metasomatic assemblages with MARID material being a prime example (Grégoire et al., 2002; Pearson and Nowell, 2002). The similar overall enrichment of Maniitsoq kimberlites and Sarfartoq aillikites in incompatible elements, but the distinctly higher Cs–Rb–K and Zr–Hf–Ti contents of aillikites from Sarfartoq and in general (Fig. 4), suggest that these geochemical differences were attained by interactions between a ubiquitous carbonate-rich, K-poor sublithospheric melt component and regionally more diverse, K-rich metasomatized cratonic mantle as reported from the Sarfartoq area (Larsen and Garrit, 2005). Within such a model it does not appear to be a coincidence that the strongest K-enrichment occurs in aillikite magmas that intruded Paleoproterozoic mobile belts such as the Torngat orogen in Labrador, whereas the coeval K-poor Maniitsoq kimberlites erupted through more pristine, refractory Archean lithospheric mantle (Wittig et al., 2010). In summary, despite the fact that a primary carbonatitic–kimberlitic melting continuum exists in experiments, the compositional variation of the magmatic spectrum observed in West Greenland is not simply a result of varying degrees of partial melting of a homogeneous source. Rather, it appears that kimberlites, aillikites and carbonatites can be related to a common carbonate-rich precursor magma derived from a sublithospheric mantle source region, but that their distinctive compositional characteristics developed within the cratonic lithosphere. Within the context of such a model it seems likely that the Maniitsoq kimberlites provide the clearest isotopic signature of the sublithospheric ultimate source region of this volatile-rich magmatism beneath the NAC.

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levels. For the latter parameterization, a good fit of the patterns was achieved for the REE and HFSE including Zr–Hf; however, slight deviations are apparent for some of the LILE (not shown). On the basis of these trace element constraints, we conclude that a significant component of the carbonate-rich Maniitsoq kimberlites may have formed by low-degree partial melting of carbonated peridotite in excess of 6 GPa. The fairly radiogenic Nd and Hf isotope compositions indicate that this carbonate-rich source experienced long-term Sm/Nd and Lu/Hf depletion similar to the source region(s) of many OIBs (Fig. 5). Such a link between kimberlite and OIB source regions has been frequently proposed in the past (Carlson et al., 2006; Le Roex, 1986; Nowell et al., 2004; Ringwood et al., 1992) and is supported here by OIB-like Ce/Pb and Nb/U ratios (Fig. 3C), and high OIB-like 3He/4He of Greenland kimberlites (Tachibana et al., 2006). 5.3. Significance of Nd–Hf isotope decoupling in kimberlitic magmas An important finding from our study is that the Maniitsoq kimberlites and their megacrysts are located entirely within the Nd–Hf terrestrial array (Fig. 5B) where no ‘anomalous’ isotopic signatures are observed (cf., Gaffney et al., 2007). Departure from the Nd–Hf terrestrial array, if at all, only occurs in a few aillikite samples from Sarfartoq (e.g., KIM-7 in Figs. 5B and 8A), for which we have

5.2. A sublithospheric origin for Greenland kimberlites The close major element compositional match between the Maniitsoq kimberlites and the experimentally produced near-solidus kimberlitic melts of Brey et al. (2008) suggests melt derivation from carbonated peridotite at 6 to 10 GPa beneath West Greenland, which corresponds to the convective upper mantle, or asthenosphere. Brey and co-workers also provided a set of internally consistent ‘bulk’ peridotite/ melt partition coefficients that enable us to calculate the trace element patterns for partial melts of carbonated peridotite at 6 and 10 GPa by utilizing a batch partial melting equation (cf., Dasgupta et al., 2009). The trace element content of the source was set at primitive mantle values (Palme and O'Neill, 2003) and the degree of partial melting was fixed at 0.1% based on theoretical and experimental constraints (Gudfinnsson and Presnall, 2005; McKenzie, 1985). A first-order observation is that the calculated primitive mantle normalized trace element patterns match remarkably well the trace element distribution of Maniitsoq kimberlites, including the Zr–Hf trough (Fig. 4C). Although we note that the incompatible element concentration levels in the calculated melts are mostly lower than in the Maniitsoq samples, the deviation is by less than an order of magnitude. However, choice of a more fertile starting composition and/or further lowering of the melting degree results in a good match between model and nature. We have also performed the same type of trace element modeling using the 6.6 GPa and 8.6 GPa bulk peridotite/carbonate melt partition coefficients of Dasgupta et al. (2009), which yielded generally higher trace element concentration

Fig. 8. Mixing models in εHfi vs. εNdi isotope space for kimberlitic magma compositions from the Greenland-Labrador Diamond Province (A) and South Africa (B). The mixing hyperbolae are constructed between an isotopically depleted, radiogenic end-member (component 1; proxy for convective mantle-derived carbonate-rich melt) and a variety of ‘local’ isotopically enriched, unradiogenic end-members (components 2, 3, and 4; proxies for K-rich melts derived from cratonic mantle metasomes). The Nd and Hf contents of the invoked end-member melt compositions are given in panel ‘B’ and the underlying modeling assumptions are described in Section 5.4. Tick marks are at 5% intervals. Data for Labrador aillikites are from Tappe et al. (2007, 2008) and for South African kimberlites from Nowell et al. (2004). Terrestrial array is after Vervoort et al. (1999). The Greenland kimberlites and their low-Cr megacrysts are displayed in both panels for easier comparison.

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identified minor input from K-rich metasomatized cratonic mantle based on geochemical grounds (see Section 5.1). However, this small but significant Hf isotopic difference between the Maniitsoq kimberlites and Sarfartoq aillikites at similar Nd isotope compositions resembles on a smaller scale the pronounced decoupling of the Nd–Hf isotope systematics in South African kimberlites (Fig. 8B). Decoupling of the Sm/Nd and Lu/Hf isotope systems is uncommon in many terrestrial rocks because of their coupled fractionation in the crust–mantle system (Vervoort et al., 1999). A possible solution to this decoupling problem (as recognized for the South African data set) was that kimberlite magmas sampled a ‘hidden’ geochemical reservoir with Nd–Hf isotopic values below the terrestrial array (Nowell et al., 2004). This hidden reservoir, by definition largely untapped by MORB and OIB magmatism, had been previously postulated to resolve the apparent Nd–Hf mass imbalance between the bulk silicate Earth composition and the crust–mantle array (Blichert-Toft and Albarede, 1997). A major requirement for the putative hidden reservoir is significantly lower Lu/Hf relative to Sm/Nd and long-term isolation from mantle convection in order to facilitate the Nd–Hf isotope decoupling. Numerous authors therefore appealed to ancient subducted oceanic crust trapped in the mantle transition zone (Blichert-Toft and Albarede, 1997; Nowell et al., 2004) or at the core–mantle boundary (Bizzarro et al., 2002; Rudnick et al., 2000) as a viable candidate for this ‘missing’ reservoir. More recently, Nd–Hf isotope data from kimberlite occurrences outside of Africa have been used to make a case for the global geographic extent of the hidden reservoir (Gaffney et al., 2007; Paton et al., 2009). However, a new expanded Nd–Hf isotope data set for chondritic meteorites (Bouvier et al., 2008) is much more consistent with crust– mantle Nd–Hf isotope systematics and, thus, obviates the need for a hidden reservoir to explain the bulk silicate Earth composition. Without the need for a hidden reservoir we explore the option of whether or not interaction between convective mantle-derived magmas and cratonic metasomes can produce Nd–Hf isotope decoupling in kimberlites and related magmas (cf., Choukroun et al., 2005; Janney et al., 2002). 5.4. Understanding the cratonic mantle metasome imprint on kimberlitic magmas Recently, Prelevic et al. (2010) demonstrated for Mediterranean lamproites that Nd–Hf isotope decoupling along subhorizontal arrays in 2D isotope ratio space can be readily explained by strongly hyperbolic mixing relationships between isotopically and geochemically strongly contrasting melt components. In the Mediterranean case, the mixing is envisaged to take place between isotopically enriched melts of crustally contaminated mantle lithosphere with low Nd/Hf and a subordinate isotopically depleted convective mantlederived melt component akin to OIB with a much higher Nd/Hf. Importantly, neither of the invoked end-members were required to fall outside of the terrestrial Nd–Hf isotope array, and yet this particular mixing relationship dictates decoupling of Nd from Hf isotopes along a subhorizontal array, which principally represents the linear segment of a strongly curved hyperbola. Although there is little reason to compare the petrogenesis of orogenic lamproites with that of kimberlitic magmas from stable cratonic areas, it appears that in both petrogenetic scenarios the decoupling of Nd–Hf isotope systematics can be facilitated by mixing of melts with contrasting isotope and elemental ratios (i.e., high and low Nd/Hf). In this section we argue that a minor imprint from a K-rich cratonic melt component on carbonate-rich convective mantle-derived magma was responsible for the lower εHf of Sarfartoq aillikites compared to the Maniitsoq kimberlites; but, importantly, this interaction did not significantly affect Nd isotopes. The 87Sr/86Sr ratios are also considered immune to the interaction process outlined above due to the fact that the Sr elemental budget is largely controlled by the carbonate-rich asthenospheric melt component. In an attempt to better understand the

impact of such an interaction on the Nd–Hf isotope systematics of resultant kimberlitic magmas, we have performed simple binary mixing calculations between plausible end-member compositions relevant to the Greenland-Labrador data (Fig. 8A) and the South African kimberlite data (Fig. 8B) of Nowell et al. (2004). For Greenland-Labrador, we have demonstrated in Section 5.1 that the isotopically depleted melt component is associated with the most carbonate-rich samples (we exclude the Sarfartoq carbonatite as it is far removed from a liquid composition). Our compilation of N500 Nd isotope compositions from over 50 worldwide carbonatite occurrences shows a pronounced mode between +2.8 and +5.8 εNd, indicating that the vast majority of global carbonatites have moderately depleted isotope compositions (Bizimis et al., 2003; Nelson et al., 1988; Tappe et al., 2008). Therefore we have assigned +4 εNd and +7 εHf to the carbonate-rich end-member, which is independently supported by the fact that the Greenland and Labrador data sets converge toward this Nd– Hf isotope composition (Fig. 8A). Such a composition is also consistent with those of the low-Cr megacrysts from Maniitsoq kimberlites, which formed from a magma batch(es) that did not extensively react with the lithospheric mantle. Tappe et al. (2008) observed that the Labrador aillikites approach the unradiogenic Nd–Hf isotope compositions of Mesoproterozoic olivine lamproites from the same region (Fig. 5B), and they argued that MARID-type cratonic metasomes (typically considered a source material for anorogenic lamproite magmas; Foley, 1992) must have been involved in the genesis of the Labrador aillikites. Hence, the Nd–Hf isotope composition of a Labrador lamproite sample (−8.4 εNd and −8.5 εHf) is taken as a proxy for a melt from the K-rich metasomatized cratonic mantle beneath the Labrador segment of the NAC. A similar suite of Mesoproterozoic olivine lamproites occurs in West Greenland near Sarfartoq (Nelson, 1989). The Nd isotope compositions are even more unradiogenic (−13 εNd) than for their Labrador counterparts (Fig. 5A), supporting the idea that cratonic metasomes are compositionally highly variable within the same craton and even on a sample scale (Choukroun et al., 2005). Unfortunately, no Hf isotope data are available for the Greenland lamproites and we have inferred this parameter by extrapolating εNd to the terrestrial Nd–Hf isotope array (−14.5 εHf). This seems reasonable given the fact that the Labrador lamproite Nd–Hf isotope compositions fall close to the terrestrial array of Vervoort et al. (1999). Nd and Hf concentrations were chosen to yield a constantly high Nd/Hf ratio of 40 (100 ppm Nd; 2.5 ppm Hf) for the depleted end-member as guided by experimentally produced carbonate-rich melt compositions (Brey et al., 2008). For the enriched end-member(s), we have chosen Nd and Hf concentrations from the high- (700 ppm Nd; 20 ppm Hf) and low ends (100 ppm Nd; 25 ppm Hf) of reported lamproite compositions (Mitchell and Bergman, 1991). The following two mixing relationships were obtained: (1) A near linear Nd–Hf isotope covariation (i.e., no decoupling) was produced when the melt from the locally inferred, isotopically enriched component for the metasomatized cratonic mantle beneath northern Labrador was mixed with a melt from the isotopically more depleted component, and both end-members contributed Nd and Hf in equal proportions, i.e., Nd/Hf of 35 vs. 40 (Fig. 8A). In this model up to 20% of a K-rich melt component from a cratonic metasome is required to explain the Labrador aillikite data spread. This isotopic constraint is in good agreement with the rather significant impact that this metasomatic component left on the Labrador aillikite major and trace element compositions; i.e., they contain up to 9 wt.% TiO2 and 300 ppm Nd (Figs. 3A and 4). (2) A strongly curved hyperbolic mixing relationship was obtained when the melt from the locally inferred, isotopically enriched component for the Greenland cratonic mantle was blended with the melt from the isotopically depleted component (Fig. 8A). In this mixing scenario, it is the strongly contrasting Nd/Hf elemental ratios (4 vs. 40) of the end-members that, when

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mixed, result in departure from the terrestrial Nd–Hf isotope array. For example, the ‘depleted end’ of the mixing curve in Nd– Hf isotope space exhibits a near vertical segment due to the asymptotic behavior, which, in turn, must result in ‘hybrid’ magmas that show a wide range in Hf isotope compositions at relatively constant Nd (and by inference Sr isotope compositions). In this model, less than 5% incorporation of an isotopically enriched melt from a cratonic metasome into a convective mantle-derived, carbonate-rich magma may cause displacement of the Hf isotope composition by up to 7 epsilon units while leaving the Nd isotope compositions virtually unchanged. Such a shift corresponds to the magnitude of isotope decoupling observed between the Greenland kimberlites and aillikites. Moreover, this model reconciles well with the small, but significant major and trace element compositional differences between rock types (Figs. 3 and 4). Finally, we examined the impact that a melt contribution from a cratonic K-rich metasome would have on the Nd–Hf isotope systematics of Southern African kimberlites (Fig. 8B). For this calculation the depleted end-member is the same as for Greenland-Labrador in terms of isotopic composition and elemental ratios, consistent with the idea that such a component is rather homogenous. For the enriched end-member, however, we selected a combination of measured Nd and Hf isotope compositions from South African MARID xenoliths (−12 εNd, Grégoire et al., 2002; and −25 εHf, Choukroun et al., 2005; note that to the best of our knowledge combined Nd–Hf isotope data do not exist for MARID xenoliths). The assigned trace element concentration of this MARIDderived melt component (400 ppm Nd and 25 ppm Hf) was guided by fairly enriched lamproite compositions (Mitchell and Bergman, 1991). This combination of interacting melts shows that, as with the Greenland magmas, only minor amounts (at most 8%, or much less if isotopically more extreme MARID compositions are chosen) of melt from enriched cratonic metasomes are required to generate a disproportionately strong impact on the Hf isotope compositions of sublithospheric kimberlitic magmas. 6. Implications for the origin of kimberlites Over the past decade it has been recognized that kimberlite and related magmas tend to show decoupled Nd–Hf isotope systematics. One view is that this feature has been inherited from an ultra-deep magma source region that must contain an ancient subducted oceanic crust component. Indirect evidence for an ultra-deep origin of the decoupled Nd–Hf isotope signature in kimberlites comes from rare entrained sublithospheric diamonds and the fact that this signature is largely absent from shallower basaltic magmatism. In this study, we have also identified a small but significant decoupling of Hf from Nd isotopes in an exceptionally fresh suite of kimberlite and aillikite dykes from West Greenland. However, our geochemical data, including fairly tight correlations with the Hf isotope compositions, suggest that this signature was not attained from the sublithospheric kimberlite source region. In the Greenland example, the Nd–Hf isotope decoupling can be linked to the interaction between asthenosphere-derived carbonate-rich ‘kimberlitic’ melt and a K-rich cratonic metasome. It appears that less than 5% input from this metasome is sufficient to shift the Hf isotope composition of the kimberlitic magma by up to 7 epsilon units while leaving Nd (and Sr) isotopes largely unaffected. An important difference in the model presented here is that the Nd–Hf isotope decoupling was not inherited from any specific source component, whether deep or shallow, but resulted from the mixing relationships between compositionally highly contrasting melt components. We therefore urge caution when interpreting Nd–Hf isotope compositions of kimberlites and related rocks, because it appears that the commonly observed isotopic decoupling can have multiple origins. Hence, there is no a priori reason

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to interpret this peculiar signature as a ‘deep source signal’ and we reiterate that in many cases it may be the long-term metasomatic memory of the cratonic lithosphere that adds special chemical flavors to deep-seated magmatism observed at Earth's surface. Supplementary materials related to this article can be found online at doi:10.1016/j.epsl.2011.03.005. Acknowledgments This work is published with permission of the Geological Survey of Denmark and Greenland. Analytical work in the Radiogenic Isotope Facility at the University of Alberta is in part funded by a Major Resource Support grant and at Durham University by NERC operation grants. Chris Ottley is thanked for help with the trace element determinations in Durham. Agnete Steenfelt, Karsten Secher, Sven-Monrad Jensen, and Karina Sand are thanked for help during fieldwork and for providing additional samples. 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