Earth and Planetary Science Letters 401 (2014) 132–147

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Earth and Planetary Science Letters

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LIMA U–Pb ages link lithospheric mantle metasomatism to Karoo

magmatism beneath the Kimberley region, South Africa

Andrea Giuliani a,∗, David Phillips a, Roland Maas a, Jon D. Woodhead a, Mark A. Kendrick a,b, Alan Greig a, Richard A. Armstrong b, David Chew c, Vadim S. Kamenetsky d, Marco L. Fiorentini e

a School of Earth Sciences, The University of Melbourne, Parkville, 3010 Victoria, Australiab Research School of Earth Sciences, The Australian National University, Acton, 0200 ACT, Australiac Department of Geology, School of Natural Sciences, Trinity College Dublin, Dublin 2, Irelandd School of Earth Sciences, University of Tasmania, Hobart, 7001 Tasmania, Australiae Centre for Exploration Targeting, Australian Research Council Centre of Excellence for Core to Crust Fluid Systems, School of Earth and Environment, University of Western Australia, Crawley, 6009 WA, Australia

a r t i c l e i n f o a b s t r a c t

Article history:Received 3 February 2014Received in revised form 23 May 2014Accepted 24 May 2014Available online xxxxEditor: B. Marty

Keywords:crichtonite-series LIMA titanateU–Pb geochronologymantle metasomatismKimberleyKaroo magmatism

The Karoo igneous rocks (174–185 Ma) of southern Africa represent one of the largest continental flood basalt provinces on Earth. Available evidence indicates that Karoo magmas either originated in the asthenosphere and were extensively modified by interaction with the lithospheric mantle prior to emplacement in the upper crust; or were produced by partial melting of enriched mantle lithosphere. However, no direct evidence of interaction by Karoo melts (or their precursors) with lithospheric mantle rocks has yet been identified in the suites of mantle xenoliths sampled by post-Karoo kimberlites in southern Africa. Here we report U–Pb ages for lindsleyite–mathiasite (LIMA) titanate minerals (crichtonite series) from three metasomatised, phlogopite and clinopyroxene-rich peridotite xenoliths from the ∼84 Ma Bultfontein kimberlite (Kimberley, South Africa), located in the southern part of the Karoo magmatic province. The LIMA minerals appear to have formed during metasomatism of the lithospheric mantle by fluids enriched in HFSE (Ti, Zr, Hf, Nb), LILE (K, Ba, Ca, Sr) and LREE. LIMA U–Pb elemental and isotopic compositions were measured in situ by LA-ICP-MS methods, and potential matrix effects were evaluated by solution-mode analysis of mineral separates. LIMA minerals from the three samples yielded apparent U–Pb ages of 177 ±12 Ma, 178 ±29 Ma and 190 ±24 Ma (±2σ ). A single zircon grain extracted from the ∼190 Ma LIMA-bearing sample produced a similar U–Pb age of 184 ± 6 Ma, within uncertainty of the LIMA ages. These data provide the first robust evidence of fluid enrichment in the lithospheric mantle beneath the Kimberley region at ∼180–190 Ma, and suggest causation of mantle metasomatism by Karoo melts or their precursor(s). The results further indicate that U–Pb dating of LIMA minerals provides a new, accurate tool for dating metasomatic events in the lithospheric mantle.

© 2014 Elsevier B.V. All rights reserved.

1. Introduction

Constraining the temporal relationship between mantle pro-cesses and major crustal events is a fundamental challenge in Earth geochemistry. For example, it is likely that large-scale partial melting events in the mantle coincide with major crust-forming events and/or periods of active tectonism (e.g., Pearson et al., 2007). In the Kaapvaal craton (southern Africa; Fig. 1), Re–Os model ages of mantle xenoliths and their sulphides sampled by

* Corresponding author. Tel.: +61 3 90359873; fax: +61 3 83447761.E-mail address: andrea.giuliani@unimelb.edu.au (A. Giuliani).

http://dx.doi.org/10.1016/j.epsl.2014.05.0440012-821X/© 2014 Elsevier B.V. All rights reserved.

kimberlite magmas mirror the timing of major tectonic events recorded in crustal rocks, such as craton amalgamation in the Archean and subsequent accretion events in the Proterozoic result-ing in the formation of orogenic belts (e.g., Griffin et al., 2004;Pearson and Wittig, 2008 and references therein; Simon et al., 2007). The Re–Os isotopic system is thus thought to provide robust geochronological constraints on major partial melting events in the mantle (see reviews by Rudnick and Walker, 2009, and Shirey and Walker, 1998).

In contrast, precise dating of metasomatic (i.e. fluid enrich-ment) events recorded in mantle rocks remains challenging, thus hindering the correlation of mantle metasomatic episodes with magmatic/tectonic events in the crust. Radiogenic isotopic systems

A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147 133

Fig. 1. Schematic map of southern Africa showing the location of major Karoo lavas, dykes and sills (adapted from Jourdan et al., 2004). The estimated boundaries of the Kaapvaal craton and the position of Kimberley and of some clusters of orangeites (or Group II kimberlites; Barkly West, Kroonstad, Swartruggens) are also shown.

commonly applied to date metamorphic and magmatic events in the crust, such as Rb–Sr and Sm–Nd, are generally not consid-ered reliable for dating the formation of metasomatic minerals in mantle xenoliths. This is due to the potential for disturbance and resetting of these chronometers during xenolith residence at man-tle temperatures and/or xenolith transport in host kimberlite or alkali basalt magmas (e.g., Allsopp and Barrett, 1975; Kramers et al., 1983; Lazarov et al., 2012; McDonough and McCulloch, 1987;O’Reilly and Griffin, 2013; Pearson et al., 1995b; Schmidberger et al., 2003). Claims that the 40Ar/39Ar dating technique can re-liably date mantle metasomatic events (e.g., Hopp et al., 2008) have yet to be tested rigorously, with spurious ages likely, be-cause of incorporation of excess 40Ar – i.e. Ar not produced by in situ decay of 40K – in minerals (e.g., Phillips and Onstott, 1988;Phillips, 2012). Re–Os ages of sulphides and bulk rocks are also un-likely to provide temporal constraints for fluid enrichment event(s), because metasomatic sulphides have much lower Os concentra-tions than pre-existing refractory sulphides (Alard et al., 2000; Rudnick and Walker, 2009).

Given its well-established high closure temperature (>900 ◦C; Cherniak and Watson, 2001), U–Pb dating of mantle zircons prob-ably provides the most robust age constraints for mantle metaso-matic events (e.g., Katayama et al., 2003; Kinny and Dawson, 1992;Kinny and Meyer, 1994; Kinny et al., 1989; Konzett et al., 1998;Rudnick et al., 1999; Simonetti and Neal, 2010; Spetsius et al., 2002; Zheng et al., 2006). For example, the studies of Liati et al.(2004) and Liu et al. (2010) have revealed distinct generations of zircons in mantle xenoliths from Namibia and the North China Cra-ton, respectively, with ages corresponding to major tectonic events in each region. However, even mantle zircons appear susceptible to partial resetting in some cases (Konzett et al., 1998, 2000, 2013).

In this study, we document the U–Pb isotope systematics of minerals (LIMA, clinopyroxene, zircon) in intensely metasomatised, phlogopite-rich mantle peridotites from the Kimberley kimberlites (South Africa). New U–Pb ages for metasomatic LIMA obtained in situ by LA-ICP-MS were validated by isotope dilution analyses of LIMA separates. Clinopyroxene, co-genetic with LIMA, was also analysed in an attempt to define the common-Pb composition of LIMA. LIMA U–Pb ages vary between ∼180–190 Ma, consistent with a single zircon SHRIMP U–Pb age of ∼184 Ma, and provide the first robust evidence for metasomatism of the southern African lithospheric mantle coeval with Karoo flood basalt magmatism in the region.

1.1. Phanerozoic magmatism on the Kaapvaal craton

The Phanerozoic evolution of the Kaapvaal craton is punctuated by three main magmatic cycles, namely the eruption of the Karoo large igneous province in the mid–early Jurassic (174–185 Ma – Jourdan et al., 2005, 2007a and references therein), the intru-sion of orangeites (or Group II kimberlites) from ∼115–200 Ma and the emplacement of Group I (or archetypal) kimberlite mag-mas mainly at ∼80–115 Ma (e.g., Allsopp and Barrett, 1975;Le Roex, 1986; Phillips et al., 1998, 1999; Smith et al., 1985, 1994). The Karoo magmas were erupted over an extensive area of south-ern Africa (Fig. 1), over a period of ∼5–6 Myr (peak of volcanic ac-tivity between 178–184 Ma – Jourdan et al., 2007a; Svensen et al., 2012). Conversely, orangeites and kimberlites represent punctuated magmatic events each formed in a short (less than few Myr) period of time. Geochronology studies of mantle xenoliths, mainly from the 80–90 Ma kimberlites of the Kimberley cluster (Fig. 1), have attempted to link apparent ages for mantle metasomatic events to these major Phanerozoic magmatic events (Hamilton et al., 1998;Hawkesworth et al., 1990; Kinny and Dawson, 1992; Konzett et al., 1998, 2000; Kramers et al., 1983) and to older tectonic events such as the Eburnian (∼1.7–2.0 Ga) and Kibaran (∼1.3–1.0 Ga) orogenies (Hopp et al., 2008; Shu et al., 2013). The xenolith-based ages ob-tained from zircon U–Pb and whole-rock Rb–Sr dating range from ∼145 to 80 Ma, and have been variably interpreted to reflect meta-somatism of the lithospheric mantle by kimberlitic and orangeitic melts (Hamilton et al., 1998; Kinny and Dawson, 1992; Konzett et al., 1998, 2000, 2013). Some mantle zircons in xenoliths from the Kimberley kimberlites provided U–Pb ages up to 142 Ma (Konzett et al., 1998), distinctly older than the entraining kimberlite and the orangeites (i.e. Barkley West cluster, 118–120 Ma – Phillips et al., 1999) emplaced in the proximity of Kimberley (Fig. 1). If some de-gree of resetting during transport in the host kimberlite magma is accepted, such ages might indicate older mantle metasomatism, possibly by Karoo melts (Konzett et al., 1998, 2000). However, ro-bust geochronological links between Karoo magmatism and mantle metasomatism are lacking.

1.2. LIMA minerals: previous work

LIMA (LIndsleyite–MAthiasite) phases are titanate minerals of the crichtonite series with the general formula AM21O38. The M site is occupied predominantly by Ti, Cr, Fe2+, Fe3+, Zr and Mg, whereas the A site hosts larger cations (K, Ba, Sr, Ca, Na, REE, U, Pb – Grey et al., 1976; Haggerty, 1983, 1991). Minerals of the crichtonite series are named according to the dominant A-site cation – e.g., lindsleyite for Ba, mathiasite for K, crich-tonite for Sr, loveringite for Ca (Haggerty, 1983; Haggerty et al., 1983 and references therein). LIMA phases appear to be the main crichtonite-series minerals occurring in the mantle. LIMA min-erals have been identified in intensely metasomatised peridotite xenoliths entrained by kimberlite magmas (Erlank et al., 1987;Haggerty et al., 1983; Jones et al., 1982; Konzett et al., 2000, 2013), in mantle polymict breccias (Giuliani et al., 2013a), in kimberlitic heavy mineral concentrates (Haggerty, 1983; Griffin et al., 2014;Zhou et al., 1984), and as inclusions in diamond (Leost et al., 2003;Sobolev et al., 1997) and mantle garnet (Wang et al., 1999).

LIMA minerals are highly enriched (hundreds to thousands of ppm) in incompatible trace elements, including LILE (K, Ba, Sr, Na, Pb), HFSE (Ti, Zr, Hf, Nb), LREE, U and Th (Jones and Ekam-baram, 1985; Konzett et al., 2013; Griffin et al., 2014). In mantle rocks LIMA minerals generally replace or overgrow chromite and are typically associated with phlogopite, clinopyroxene and rutile, and less commonly with K-richterite, ilmenite and other metaso-matic phases such as carbonates, sulphides and Zr-rich minerals (Erlank et al., 1987; Haggerty et al., 1983; Jones et al., 1982;

134 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Konzett et al., 2000, 2013). LIMA minerals in the mantle are of un-doubted metasomatic origin and appear to have crystallised from silicate or silicate–carbonate melts enriched in LILE, HFSE and LREE perhaps related to the PKP (Phlogopite–K-richterite Peridotites) style of metasomatism (Erlank et al., 1987; Haggerty et al., 1983;Jones, 1989). Experimental investigations have shown LIMA min-erals are stable to depths of ∼300 km, i.e. throughout the litho-spheric mantle (Konzett et al., 2005). Given the extreme enrich-ment in incompatible elements and the wide range of P–T condi-tions for which LIMA minerals are stable, these phases could be an important mantle source component for alkaline magmas such as kimberlites and carbonatites (Foley et al., 1994; Haggerty, 1983;Jones, 1989; Konzett et al., 2005).

Due to relatively high concentrations of U (>500 ppm; Griffin et al., 2014), LIMA minerals are potential candidates for dat-ing mantle metasomatism using U–Pb geochronology. However, elevated concentrations of Pb, and therefore of “common” (i.e. non-radiogenic) Pb, incorporated during crystallisation, represent a challenge for U–Pb dating of LIMA minerals. In recent years a number of minerals with variable and sometimes low U/Pb ra-tios have been dated successfully by in situ LA-ICP-MS (both single and multi-collector) methods, including perovskite (Batumike et al., 2008; Cox and Wilton, 2006), apatite (Chew et al., 2011), ti-tanite (Simonetti et al., 2006; Storey et al., 2006), allanite (Cox et al., 2003) and loparite (Mitchell et al., 2011). These studies have shown that reliable Tera-Wasserburg concordia ages can be ob-tained when there is sufficient spread in U/Pb ratios such that a robust regression through the data can be attained. Isochron di-agrams using 204Pb as common denominator (e.g., 207Pb/204Pb vs206Pb/204Pb) are not easily applied to LA-ICP-MS analyses because measurements of 204Pb are compromised by isobaric interference from 204Hg.

In a recent attempt to date LIMA minerals by LA-ICP-MS U–Pb methods, Griffin et al. (2014) measured the isotopic compositions of 15 LIMA macrocrysts from the Jagersfontein kimberlite (South Africa). However, the measured grains provided limited spread in U/Pb ratios, resulting in a poorly constrained age. In contrast, the LIMA grains in mantle xenoliths from the current study provide a much greater range in U/Pb ratios, thus allowing the acquisition of robust ages for LIMA formation and, hence, mantle metasomatism.

2. Sample selection

Three LIMA-bearing peridotite xenoliths, XM1/341, XM1/345 and XM1/362, were collected from the Bultfontein Dumps (Kim-berley, South Africa), which contain waste material from mining of the Bultfontein kimberlite. The Bultfontein kimberlite is part of the Kimberley cluster of kimberlites, which also includes the De Beers, Dutoitspan, Wesselton, Wesselton Floors and Kimberley kimberlites, along with a number of smaller pipes (Field et al., 2008). The Kimberley cluster is located in the southwestern part of the Kaapvaal craton and in the southwestern portion of the Karoo magmatic province (Fig. 1). The Kimberley kimberlites have been classified as Group I or archetypal kimberlites on the basis of their Sr–Nd isotopic signatures (Smith, 1983). A variety of dating tech-niques (i.e. zircon U/Pb, perovskite U/Pb, phlogopite Rb/Sr, phlogo-pite 40Ar/39Ar) have provided emplacement ages between 81 and 90 Myr for the Kimberley kimberlites (Allsopp and Barrett, 1975;Batumike et al., 2008; Davis, 1977; Fitch and Miller, 1983; Smith et al., 1989). The Kimberley kimberlites are well known for hosting a range of mantle xenoliths with distinct metasomatic styles (e.g., Giuliani et al., 2012, 2013b, 2013c, 2014; Grégoire et al., 2002;Pearson et al., 2003; Rehfeldt et al., 2008).

3. Analytical methods

3.1. Major oxide and U–Th–Pb concentrations of LIMA minerals

Electron microprobe (EMP) analyses of LIMA grains were car-ried out on thin sections and epoxy mounts, using a Cameca SX50 electron microprobe at the University of Melbourne and em-ploying the same conditions described by Giuliani et al. (2013a). Uranium, thorium and lead elemental abundances in LIMA grains were measured in situ, using an Agilent 7700× quadrupole ICP–MS interfaced with an excimer 193 nm UV laser ablation probe (Woodhead et al., 2007). Laser ablation conditions were as follows: ablation time of 60 s; fluence ∼3 J/cm2, repetition rate of 5 Hz; standard delay for sample washout of 15 s; longer delay for back-ground measurements (50 s) every 5 analyses; beam size of 42 μm. The synthetic glass BCR-2G was used as the calibration material, and 49Ti was the internal standard with Ti concentrations deter-mined from EMP analyses. Natural and synthetic glasses NIST610 and BHVO-2G were analysed as unknowns to verify data integrity (Table S1, Supplementary Material).

3.2. LA-ICP-MS U–Pb ages of LIMA grains

U–Th–Pb age determinations of LIMA grains were carried out using the same analytical set-up described above, but in a separate session optimised for geochronology. In addition to 206–207–208Pb, 232Th and 238U, 202Hg and 204Pb + Hg were monitored in an at-tempt to constrain common-Pb compositions based on 204Pb. How-ever, the uncertainties in Hg-corrected 204Pb were too large to be useful for 204Pb-based common-Pb corrections. Down-hole frac-tionation of the Pb/U ratio for LIMA was initially assumed to be similar to that of zircon, and the zircon reference material 91 500 (Wiedenbeck et al., 1995) was used to both monitor and correct for down-hole fractionation and instrument drift (Paton et al., 2010). Plesovice zircon (Slama et al., 2008) was measured as a secondary reference material to check data integrity (Table S2, Supplementary Material). Raw data were reduced using the Iolite software pack-age (Paton et al., 2011). Fractionation-corrected 207Pb/206Pb and 238U/206Pb ratios for LIMA were plotted on Tera-Wasserburg dia-grams and regression lines calculated using the Isoplot software (Ludwig, 2012a).

3.3. Solution ICP–MS U–Th–Pb isotope analyses of LIMA and clinopyroxene grains

Solution-mode U–Th–Pb data were collected for two aliquots of XM1/362 LIMA and three samples of clinopyroxene, one from each studied xenolith. High-purity separates of LIMA (0.10 and 0.15 mg) and clinopyroxene (3.2 to 8.6 mg) grains were cleaned with hot 2M nitric acid (1 h), rinsed with ultra-pure water and dissolved in 0.5–1.0 ml of (3:1) HF–HNO3 and 6M HCl on a hotplate. This cleaning procedure was designed to remove any labile Pb from the grains; comparison with the results of step leaching experiments (Frei et al., 1997) suggests that the isotopic ratios intrinsic to the sample are unlikely to be affected by leaching with such dilute acid for such a short period. All samples produced clear solutions. A weighed split of each sample (∼10 wt.%) was removed and anal-ysed for trace elements using an Agilent 7700× quadrupole ICP–MS. USGS basalt BCR-2, analysed in the same batch, yielded results consistent with published values (i.e. GeoReM preferred values – http :/ /georem .mpch-mainz .gwdg .de). The remainders (∼90 wt.%) of the LIMA solutions were equilibrated with a 233U–205Pb tracer; no spike was added to the clinopyroxene solutions. Lead was ex-tracted using a double pass over small (0.1 ml) columns of AG1–X8 (100–200 mesh) anion-exchange resin, using conventional HBr–HCl chemistry. Total Pb blanks were <20 pg. Extraction of U from the

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LIMA fractions was achieved using EICHROM TRU resin (Pin and Santos Zalduegui, 1997). Isotopic analyses were carried out at the University of Melbourne on a Nu Plasma multi-collector ICP–MS coupled to a CETAC Aridus desolvator equipped with a low-uptake PFA nebulizer (Maas et al., 2005). Instrumental mass bias in Pb iso-tope runs was corrected using thallium doping for clinopyroxene (see Woodhead, 2002) and standard bracketing for LIMA aliquots. U/Pb, Th/Pb and Th/U ratios for clinopyroxene and Th/Pb ratio for LIMA samples are based on quadrupole ICP–MS trace element anal-yses of the same solutions and have a precision of ∼±2% (2sd). Isotope dilution calculations were performed using the EarthTime spreadsheet documented by Schmitz and Schoene (2007). Further details are provided in Table 3.

3.4. Zircon U–Pb dating and Hf isotopes

A single zircon grain (∼100 ×60 μm) was retrieved after crush-ing a portion of sample XM1/362. The zircon was encased in epoxy and imaged with SEM cathodoluminescence (CL) to reveal internal structures. U–Pb age determinations were carried out us-ing a SHRIMP II instrument at the Australian National University. A spot size of ∼20 μm was employed. The data were reduced in a manner similar to that described by Williams (1998 and ref-erences therein). The U/Pb ratios were calculated relative to the AS3 zircon (Paces and Miller, 1993); U and Th concentrations were determined relative to the SL13 standard. Pb*/U ratios calculated using the 204Pb, 207Pb and 208Pb correction methods were indis-tinguishable within uncertainties; the 204Pb-corrected results are listed in Table 4. Age calculations and data representation were ac-complished using the Isoplot software (Ludwig, 2012a).

In situ Lu–Hf isotopic measurements of the zircon grain were undertaken at the University of Melbourne, using a Nu Plasma multi-collector ICP–MS interfaced with an excimer 193 nm UV laser ablation probe (Woodhead et al., 2004). Laser ablation condi-tions were as follows: ablation time of 60 s; fluence ∼3 J/cm2, rep-etition rate of 5 Hz; standard delay of 30 s for sample washout and background measurements; beam size of 55 μm. Raw data were reduced using the Iolite software package (Paton et al., 2011). Iso-baric interferences of 176Yb and 176Lu on 176Hf were assessed and corrected following the procedure described by Woodhead et al.(2004) and Woodhead and Hergt (2005). Plesovice zircon standard (Slama et al., 2008) was analysed to correct for instrumental drift, whereas Temora-2 and 91 500 zircon references were measured as unknowns to assess data quality and returned values within error of the solution values provided by Woodhead and Hergt (2005) – Table S3, Supplementary Material.

4. Results

4.1. Petrography

The LIMA-bearing peridotite xenoliths are small (less than 5 ×5 × 2 cm) off-cuts of larger samples from the De Beers rock store. The xenoliths display coarse to moderately coarse granular textures (Fig. 2a). Samples XM1/341 and XM1/362 contain abundant, large (up to 1 mm across) grains of phlogopite (∼10 vol.%) and clinopy-roxene (∼5 vol.%; Fig. 2a), which are associated with, and over-print, protolith olivine and orthopyroxene grains (Fig. 2b). LIMA grains are up to 0.5 mm in size, but mostly 50–250 μm, and are closely associated with phlogopite and clinopyroxene. LIMA oc-curs interstitial to the other phases or overprinting olivine and orthopyroxene, and as inclusions in/intergrowths with phlogopite and clinopyroxene (Fig. 2c, d). Spinel and rare K-richterite only oc-cur in xenolith XM1/341.

Sample XM1/345 is a spinel harzburgite with large (cm-sized) domains dominated by phlogopite (∼90 vol.%; Fig. 2e, f) with mi-

nor spinel, LIMA and scarce clinopyroxene and ilmenite grains. LIMA grains are of similar size to the other two samples, and are commonly included in phlogopite or in contact with phlogopite and clinopyroxene. In all three xenoliths, LIMA grains are often rimmed by rutile.

4.2. Thermobarometry

The equilibration conditions of the peridotite minerals were constrained by: i) the olivine-spinel Fe–Mg exchange thermome-ter (Ballhaus et al., 1991; samples XM1/341 and /345), and ii) the ‘orthopyroxene–clinopyroxene solvus’ thermometer (Brey and Kohler, 1990; Wells, 1977). For sample XM1/341, the two inde-pendent thermometers produced different results (T = 870–930 ◦Cfor olivine-spinel and 800–840 ◦C for the two pyroxenes solvus – Table S4, Supplementary Material). Sample XM1/345 yielded olivine-spinel temperatures of ∼850 ◦C (Table S5). Similar values of 820–860 ◦C were obtained using the two pyroxene solvus-based orthopyroxene-only thermometer of Brey and Kohler (1990),whereas the orthopyroxene–clinopyroxene solvus thermometer provided consistently lower temperatures (≤750 ◦C; Table S5), probably indicating disequilibrium between clinopyroxene and or-thopyroxene. Finally, equilibrium temperatures of 800–850 ◦C were calculated for orthopyroxene-clinopyroxene in sample XM1/362 (Table S6).

Based on a heat flow of 40–41 mW/m2, which is considered appropriate for the Kaapvaal craton (e.g., Lazarov et al., 2009), the calculated temperatures correspond to pressure ranges of 3.3–3.7 (∼110–120 km; olivine-spinel equilibrium) or 3.0–3.2 GPa (∼100–105 km; orthopyroxene–clinopyroxene solvus) for sample XM1/341; 3.1–3.4 GPa for XM1/345; and 3.0–3.3 GPa for XM1/362. Therefore, it appears that the three xenoliths were sampled from similar depths.

4.3. Chemical compositions of LIMA minerals

The crichtonite series minerals documented in the three xeno-liths show compositional features typical of LIMA minerals found in kimberlitic heavy mineral concentrates and in mantle xenoliths (Erlank et al., 1987; Griffin et al., 2014; Haggerty, 1983, 1991; Haggerty et al., 1983; Jones and Ekambaram, 1985; Konzett et al., 2013). The LIMA grains contain high concentrations of Cr2O3(∼14–16 wt.%), FeO (∼8–13 wt.%) and ZrO2 (∼4–5 wt.%; Table 1). In each xenolith, LIMA grains show large variations in major ox-ide compositions, both within and between grains, consistent with previous investigations (Jones et al., 1982; Konzett et al., 2000, 2013). Distinct compositional zoning in several grains is evident from back-scattered electron SEM imaging (Fig. S1, Supplementary Material).

In samples XM1/341 and /345 two distinct populations of LIMA minerals could be distinguished based on major element and U–Th–Pb concentrations (Table 1). The majority of LIMA grains analysed in sample XM1/341 are dominated by the loverin-gite (Ca) and mathiasite (K) end-members (∼40% each); these grains are relatively enriched in CaO (1.15–1.55 wt.%) and Al2O3(0.36–0.71 wt.%) and host minor amounts of Na2O (≤0.18 wt.%). LIMA grains of this “Ca–Al-rich” group have similar concentra-tions of U (281–375 ppm) and Pb (199–377 ppm), resulting in U/Pb ratios of between 0.9–1.6, but significantly lower Th contents (≤77 ppm). The grains of the “Sr–Ba–Fe-rich” group (Table 1) are solid solutions of lindsleyite (Ba), mathiasite (K) and crichtonite (Sr). These grains display much lower CaO and Al2O3 concen-trations than the Ca–Al-rich LIMA grains, with Na2O generally below EMP detection limits, and SrO and BaO up to 1.45 and 3.40 wt.%, respectively (Fig. 3a). The two groups contain similar

136 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Fig. 2. Optical photomicrographs showing textural details of xenolith samples XM1/341, /345 and /362. a) Granular texture of sample XM1/341; b) phlogopite (Phl), clinopy-roxene (Cpx) and LIMA overprinting orthopyroxene (Opx) grains; c) large grain of LIMA partially intergrown with phlogopite, interstitial to orthopyroxene and olivine (Ol); d) inclusion of LIMA in phlogopite; e), f) domains of phlogopite with minor spinel and LIMA in sample XM1/345.

U concentrations, whereas the Sr–Ba–Fe-rich LIMA grains exhibit higher Pb concentrations (≥501 ppm; Fig. 3d) and U/Pb ratios of 0.4–0.5.

LIMA grains in xenolith sample XM1/345 are either “Ba-rich” (BaO = 2.72–3.74 wt.%, corresponding to a ∼45–50% of lindsleyite end-member; K2O < 0.29 wt.%) or “K-rich” (K2O = 0.56–1.08 wt.%, corresponding to ∼55–60% of mathiasite component; BaO <

2.16 wt.%; Fig. 3b). The other oxides show very similar concentra-tions in the Ba-rich and K-rich groups. The U–Th–Pb concentrations also largely overlap (U = 269–332 ppm; U/Pb = 0.4–0.6; Fig. 3e).

The LIMA grains in sample XM1/362 do not form distinct com-positional clusters. Nonetheless, significant compositional varia-tions are evident, particularly for the LILE oxides such as BaO (1.20–3.46 wt.%), SrO (0.83–1.57 wt.%) and K2O (0.51–0.88 wt.%; Fig. 3c). Likewise, U, Th and Pb vary widely (U = 212–368 ppm;

Pb = 245–567 ppm) and without apparent inter-elemental correla-tions (U/Pb = 0.5–1.2; Fig. 3f).

4.4. U–Pb isotope systematics of LIMA

4.4.1. In situ LA-ICP-MS age determinations for LIMA grainsThe results of in situ U–Pb isotopic analyses of LIMA grains are

reported in Table 2. No obvious correlation exists between ma-jor oxide composition and U–Pb isotopic systematics. Common-Pb contents in the LIMA grains analysed by laser ablation are rela-tively high (207Pb/206Pb ∼0.74–0.88; Table 2), but display suffi-cient variation to produce well-constrained regression lines (Fig. 4; spreading factors between 5 and 13% – cf. Jourdan et al., 2009). Sample XM1/341 yields a lower intercept age of 177 ± 12 Ma(±2sd) and an extrapolated initial 207Pb/206Pb ratio of 0.935 ±

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Table 1Major oxide (wt.%) and U–Th–Pb (ppm) compositions of LIMA grains in xenoliths XM1/341, /345 and /362.

XM1/341 LIMA XM1/345 LIMA XM1/362 LIMA

Ca–Al-rich group (n = 17) Sr–Ba–Fe-rich group (n = 6) Ba-rich group (n = 14) K-rich group (n = 12) single group (n = 16)

average variation average variation average variation average variation average variation

SiO2 0.02 <0.05 0.02 <0.02 0.02 <0.04 0.02 <0.05 0.04 <0.11TiO2 59.96 57.98-62.34 56.28 55.06-58.34 56.32 55.43-57.62 57.57 56.77-58.74 58.17 56.02-60.63Al2O3 0.52 0.36-0.71 0.18 0.16-0.21 0.16 0.13-0.18 0.15 0.11-0.17 0.35 0.32-0.38Cr2O3 16.54 14.24-17.46 15.11 14.91-15.24 15.31 14.55-15.82 14.98 13.88-15.41 15.97 14.48-17.14FeO 8.78 7.53-9.60 12.67 12.09-13.17 13.14 12.95-13.28 12.86 12.35-13.14 9.46 8.56-10.25MnO 0.04 <0.08 0.07 0.05-0.10 0.08 0.03-0.13 0.08 0.02-0.14 0.06 <0.11MgO 3.66 3.39-4.43 2.94 2.87-3.05 2.81 2.72-2.93 2.79 2.68-2.90 3.56 3.44-3.76NiO 0.05 <0.08 0.05 <0.07 0.05 <0.13 0.06 0.01-0.12 0.05 0.02-0.10SrO 0.61 0.51-0.79 1.21 0.99-1.45 0.96 0.81-1.10 0.89 0.72-1.25 1.24 0.83-1.57BaO 0.60 0.12-1.33 2.89 2.13-3.40 3.27 2.72-3.74 1.77 1.26-2.16 2.31 1.20-3.46CaO 1.42 1.15-1.55 0.37 0.34-0.40 0.27 0.24-0.30 0.40 0.28-0.67 0.64 0.51-0.93Na2O 0.12 0.08-0.18 0.01 <0.01 bdl bdl 0.04 <0.07 0.07 <0.17K2O 0.64 0.54-0.79 0.46 0.30-0.66 0.27 0.26-0.29 0.81 0.56-1.08 0.67 0.51-0.88ZrO2 4.85 4.66-5.14 4.42 4.28-4.59 3.73 3.56-3.90 3.95 3.82-4.11 4.61 3.93-5.04V2O3 0.54 0.41-0.63 0.57 0.49-0.63 0.55 0.46-0.60 0.55 0.49-0.62 0.46 0.30-0.52Nb2O5 0.52 0.42-0.72 0.59 0.48-0.69 0.24 0.16-0.30 0.27 0.19-0.31 0.42 0.34-0.52Total 98.87 96.66-100.07 97.84 95.44-98.82 97.17 96.41-98.00 97.19 95.56-98.16 97.61 96.04-99.13

U 330 281-375 323 299-353 291 269-307 311 280-332 268 212-368Th 61 42-77 100 89-110 292 267-316 290 268-328 90 45-156Pb 276 199-377 534 501-579 653 596-698 616 488-758 386 245-567U/Th 5.5 4.5-7.4 3.3 2.9-3.9 1.0 0.9-1.1 1.1 1.0-1.2 3.5 1.5-7.2U/Pb 1.2 0.9-1.6 0.6 0.6-0.7 0.4 0.4-0.5 0.5 0.4-0.6 0.7 0.5-1.2

bdl: below detection limit.

0.009 (MSWD = 1.9; n = 26/26; Fig. 4a). For sample XM1/345 the lower intercept age is 178 ± 29 Ma, with an initial 207Pb/206Pb ra-tio of 0.912 ± 0.008 (MSWD = 3.6; n = 29/29; Fig. 4b). The lower intercept age for LIMA grains from sample XM1/362 is 190 ±24 Ma(MSWD = 5.3; n = 16/19; Fig. 4c) with an initial 207Pb/206Pb inter-cept of 0.911 ± 0.011, which is indistinguishable to that for sam-ple XM1/345. The large uncertainties (12–29 Ma) associated with the lower intercept ages result from the high common-Pb con-tents of the LIMA grains (85–96% of total Pb; Table 2). As a result all samples give ages that are indistinguishable at the 95% con-fidence level (Fig. 4). MSWD values between 1.9 and 5.3 indicate very minor scatter in the dataset for each sample beyond that ex-pected from analytical uncertainties alone, probably due to minor modification of some LIMA grains after crystallisation (e.g., during xenolith entrainment and transport by the kimberlite magma). The initial 207Pb/206Pb ratios of LIMA minerals are above the ranges de-fined by southern African kimberlites (0.807–0.841) and orangeites (0.879–0.899 – Smith, 1983), but partly overlap with the large 207Pb/206Pbi interval of Karoo basalts (0.861–0.920 – Ellam, 2006;Jourdan et al., 2007b).

4.4.2. Potential matrix effects for in-situ U–Pb analyses of LIMA minerals

LA-ICP-MS analyses of minerals, such as LIMA, with different structures and compositions to zircon reference materials, require careful assessment of possible matrix effects. These potential ef-fects were evaluated by i) inspecting the time-resolved (down-hole) U/Pb fractionation trends during laser ablation analysis of individual LIMA grains; and ii) performing isotope dilution analyses of LIMA grains separated from xenolith sample XM1/362, where the absence of spinel enabled preparation of a pure LIMA separate.

i) The 206Pb/238U time-resolved signals of LIMA analyses were calculated assuming that U–Pb mass bias and down-hole frac-tionation behaviour in LIMA minerals are identical to that of the primary zircon standard. If significant differences exist, LIMA 206Pb/238U trends with time, after down-hole correction, may not be flat. Examination of the measured trends (“uncorrected 206Pb/238U” in Fig. 5) is complicated by the possibility of within-hole variations in common-Pb contents, which would be expected

to produce undulating time-resolved signals. However, the occur-rence of relatively flat time-resolved 206Pb/238U trends for some LIMA analyses (Fig. 5a) is encouraging. To examine this issue fur-ther, we employed the recently developed ‘VizualAge_UcomPbine’ data reduction package for Iolite (Chew et al., 2014). This pack-age can apply a variety of common-Pb corrections to both primary reference materials and unknowns. In this case no common-Pb correction was applied to the zircon primary reference material (91 500), which is considered free of common Pb. A reference down-hole fractionation model was calculated using 91 500 zircon and was then applied to the LIMA analyses. VizualAge_UcomP-bine then employs the 207Pb/206Pbi value calculated from the Tera Wasserburg Concordia plot to subtract common Pb from the time-resolved LIMA analyses. Following removal of common Pb, the 206Pb/238U time-resolved signals of the LIMA analyses show flat trends (“207Pb-corrected 206Pb/238U” in Fig. 5). This indicates that the undulating “uncorrected 206Pb/238U” time-resolved signals are solely due to down-hole variations in common-Pb contents.

ii) The results of U–Th–Pb solution analyses for two LIMA sep-arates of sample XM1/362 are reported in Table 3 and plotted with the in situ results in Fig. 4c. The LIMA solution analyses plot well within the 238U/206Pb–207Pb/206Pb field of the in situanalyses and on the same regression line as the in situ analyses. These results confirm that the in situ analyses were not biased significantly by element fractionation due to matrix differences be-tween LIMA and zircon, as documented previously for other min-erals such as perovskite (Batumike et al., 2008; Cox and Wilton, 2006) and apatite (Chew et al., 2011). The two solution anal-yses of XM1/362 LIMA separates yield a 238U/204Pb–206Pb/204Pb 2-point age of 205 ± 22 (±2sd) Ma (Fig. S2, Supplementary Ma-terial), within uncertainty of the lower intercept age provided by the in situ analyses (190 ± 24 Ma; Fig. 4c). This provides further validation of the in situ analytical protocol employed here.

4.4.3. Common Pb correctionTera-Wasserburg age estimates can be improved by better con-

straining the initial Pb compositions of the unknown (LIMA in this study). Common-Pb compositions can be constrained by measuring the Pb isotopic composition of cogenetic phases containing mini-

138 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Fig. 3. Major oxide (a, b, c) and U–Pb co-variation diagrams (d, e, f) for LIMA in xenolith samples XM1/341, /345 and /362. Two distinct groups of LIMA grains with different major oxide compositions are distinguishable for samples XM1/341 and XM1/345, whereas sample XM1/362 contains a single group of LIMA grains with variable composition.

mal U or very low U/Pb ratios (e.g., Chen and Simonetti, 2013;Corfu and Dahlgren, 2008). In the studied xenoliths, phlogopite, clinopyroxene and LIMA show textural features (e.g. intergrowths; Fig. 1c and Fig. S1, Supplementary Material) indicating that these phases were cogenetic. Clinopyroxene in Kimberley mantle xeno-liths show relatively low U/Pb ratios (≤0.25, but mostly <0.10 – Giuliani et al., 2013c; Grégoire et al., 2002; Simon et al., 2003). Therefore, we analysed the U–Pb isotopic compositions of clinopy-roxene separates from each sample by solution-mode methods (Ta-ble 3; Fig. 4).

The clinopyroxene from sample XM1/362 has an unusually high U content (1.7 ppm) and U/Pb ratio resembles that in bulk LIMA from this sample (238U/206Pb = 2.05, compared to LIMA values of 2.37–2.71; Fig. 4c). This means that the clinopyroxene data do not improve the estimate of the common-Pb ratio for XM1/362 LIMA. Conversely, XM1/341 and XM1/345 clinopyroxene separates con-tain very low U concentrations (<0.1 ppm) and low U/Pb ratios

(0.04) and represent promising targets to constrain common-Pb compositions. However, on Tera-Wasserburg plots, the clinopyrox-ene analyses plot away from the regression lines through the LIMA data of samples XM1/341 and XM1/345 (Fig. 4a and b). This indi-cates that clinopyroxene and LIMA were not in isotopic equilibrium at the time of kimberlite entrainment. The Sr isotopic composition of clinopyroxene in mantle xenoliths can be partially reset imme-diately prior to, or during, entrainment by the kimberlite magma (Schmidberger et al., 2003). This process is also likely to affect the Pb isotopes, because Pb diffuses at a similar rate to Sr in diopside at high temperature (Cherniak and Dimanov, 2010 and references therein).

An alternative way to correct for common Pb is to employ Pb isotopic compositions estimated from models for the evolu-tion of crustal Pb, such as that of Stacey and Kramers (1975). This approach requires knowledge of the approximate age of the sample, which is provided by the Tera-Wasserburg lower inter-

A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147 139

Table 2U–Pb isotopic compositions of LIMA grains analysed in situ by LA-ICP-MS.

Sample Label Major oxide comp.b

206Pb/238U ±2sd 207Pb/206Pb ±2sd 208Pb/206Pb %“common Pb”c 207Pb-corrected age (Ma)

±2sd

XM1/341 341-LIMA2 Ca-rich 0.2447 0.0023 0.8350 0.0052 2.0648 89.3 176.4 16.4341-LIMA3 Ca-rich 0.2418 0.0027 0.8307 0.0044 2.0416 88.8 181.7 15.5341-LIMA9 Ca-rich 0.2190 0.0024 0.8222 0.0062 2.0101 87.9 177.9 15.5341-LIMA14 0.2314 0.0030 0.8245 0.0056 2.0178 88.1 184.1 15.8341-LIMA18 Ca-rich 0.1874 0.0025 0.7998 0.0061 1.9604 85.5 182.4 13.0341-LIMA19 0.1760 0.0019 0.7958 0.0053 1.9662 85.1 176.4 11.6341-LIMA23 Ca-rich 0.1999 0.0024 0.8080 0.0060 1.9996 86.4 182.8 13.9341-LIMA6 Ca-rich 0.1884 0.0024 0.8017 0.0048 1.9778 85.7 180.8 12.1341-LIMA8 0.3007 0.0034 0.8543 0.0049 2.0820 91.3 175.1 20.1341-LIMA13 0.1950 0.0025 0.8133 0.0054 1.9920 86.9 171.0 13.1341-LIMA21 0.1912 0.0030 0.8069 0.0055 1.9893 86.3 176.4 13.0341-LIMA24 0.2178 0.0027 0.8173 0.0057 2.0044 87.4 184.6 14.9341_2 LIMA1 Sr–Ba–Fe-rich 0.2455 0.0029 0.8444 0.0059 2.0619 90.3 160.5 17.3341_2 LIMA3 Ca-rich 0.2345 0.0030 0.8373 0.0070 2.0471 89.5 165.2 17.7341_2 LIMA5 Ca-rich 0.3830 0.0060 0.8710 0.0068 2.1259 93.1 177.1 29.1341_2 LIMA8 Sr–Ba–Fe-rich 0.1994 0.0021 0.8164 0.0065 1.9984 87.3 170.4 14.3341_2 LIMA9 0.2318 0.0035 0.8305 0.0047 2.0404 88.8 174.6 15.2341_2 LIMA12 Ca-rich 0.2040 0.0031 0.8167 0.0078 2.0105 87.3 173.8 16.0341_2 LIMA13 0.1991 0.0030 0.8132 0.0056 2.0109 86.9 174.7 13.6341_2 LIMA15 0.2245 0.0033 0.8235 0.0056 2.0396 88.0 180.3 15.4341_2 LIMA18 Sr–Ba–Fe-rich 0.3176 0.0042 0.8487 0.0056 2.0820 90.7 197.4 22.0341_2 LIMA18_2 0.2235 0.0033 0.8264 0.0055 2.0619 88.3 174.9 15.3341_2 LIMA21 Sr–Ba–Fe-rich 0.4332 0.0038 0.8752 0.0051 2.1589 93.6 187.1 29.8341_2 LIMA22 Sr–Ba–Fe-rich 0.2342 0.0021 0.8292 0.0053 2.0525 88.6 178.5 15.7341_2 LIMA24 Ca-rich 0.2113 0.0018 0.8255 0.0050 2.0383 88.3 166.8 13.9341_2 LIMA24_2 0.3755 0.0033 0.8733 0.0041 2.1482 93.4 167.5 24.5min 0.1760 0.7958 1.9604 85.1 160.5max 0.4332 0.8752 2.1589 93.6 197.4weighted mean 0.2250 0.0200 0.8300 0.0092 2.0378 88.6 176.6 3.0

XM1/345 345_2 LIMA 1 Ba-rich 0.5321 0.0034 0.8705 0.0036 2.1409 95.5 162.2 32.8345_2 LIMA 2 0.4544 0.0056 0.8591 0.0035 2.1155 93.9 178.4 26.3345_2 LIMA 3 Ba-rich 0.5445 0.0039 0.8724 0.0040 2.1386 95.7 158.4 34.4345_2 LIMA 5 0.4559 0.0052 0.8660 0.0049 2.1268 95.0 154.2 30.2345_2 LIMA 6 0.4674 0.0033 0.8676 0.0041 2.1313 95.2 152.6 29.6345_2 LIMA 7 Ba-rich 0.5924 0.0039 0.8705 0.0041 2.1409 95.5 180.4 37.4345_2 LIMA 9 0.4729 0.0058 0.8679 0.0048 2.1236 95.2 150.1 30.5345_2 LIMA 10 mixture 0.6705 0.0093 0.8781 0.0056 2.1386 96.3 156.9 44.0345_2 LIMA 11 mixture 0.6484 0.0072 0.8749 0.0050 2.1336 96.0 169.7 41.7345_2 LIMA 12 mixture 0.6295 0.0055 0.8700 0.0047 2.1478 95.4 193.9 41.1345_2 LIMA 13 0.5604 0.0048 0.8683 0.0045 2.1209 95.2 179.7 36.2345_2 LIMA 14 0.3125 0.0036 0.8361 0.0045 2.0475 91.7 174.5 19.7345_2 LIMA 16 Ba-rich 0.5214 0.0056 0.8650 0.0052 2.1079 94.9 179.9 35.0345_2 LIMA 18 K-rich 0.5372 0.0050 0.8694 0.0045 2.1354 95.3 168.1 34.8345_2 LIMA 19 Ba-rich 0.4985 0.0044 0.8638 0.0045 2.1151 94.7 176.5 32.1345_2 LIMA 24 Ba-rich 0.5783 0.0043 0.8655 0.0046 2.1308 94.9 197.2 37.4345_2 LIMA 25 Ba-rich 0.6417 0.0056 0.8708 0.0031 2.1327 95.5 193.9 38.5345_2 LIMA 28 Ba-rich 0.5553 0.0036 0.8685 0.0035 2.1281 95.3 177.3 34.0345_2 LIMA 29 0.5418 0.0034 0.8679 0.0033 2.1368 95.2 175.4 32.8345_2 LIMA 32 mixture 0.6227 0.0051 0.8710 0.0053 2.1340 95.5 187.3 42.2345-LIMA1 K-rich 0.3453 0.0042 0.8356 0.0039 2.0521 91.6 193.9 21.0345-LIMA4 K-rich core;

Ba-rich rim0.7100 0.0053 0.8725 0.0034 2.1322 95.7 205.5 43.3

345-LIMA10 Ba-rich 0.5539 0.0056 0.8675 0.0055 2.1254 95.1 180.9 38.0345-LIMA17 Ba-rich 0.5389 0.0061 0.8662 0.0033 2.1227 95.0 189.5 34.1345-LIMA3 mixture 0.6755 0.0054 0.8659 0.0035 2.1249 95.0 228.0 41.0345-LIMA6 0.5559 0.0061 0.8696 0.0045 2.1254 95.4 170.0 35.3345-LIMA7 Ba-rich 0.6247 0.0075 0.8723 0.0065 2.1128 95.7 181.9 45.8345-LIMA8 0.5615 0.0055 0.8676 0.0054 2.1200 95.2 183.0 38.2345-LIMA13 0.5475 0.0055 0.8708 0.0059 2.1177 95.5 165.6 38.6min 0.3125 0.8356 2.0475 91.6 150.1max 0.7100 0.8781 2.1478 96.3 228.0weighted mean 0.5310 0.0350 0.8662 0.0035 2.1228 95.0 177.0 6.1

XM1/362 362_2 LIMA1 0.3887 0.0042 0.8458 0.0048 2.1008 92.8 187.0 32.0362_2 LIMA2 0.4257 0.0042 0.8495 0.0049 2.1160 93.2 193.1 35.3362_2 LIMA4 0.1567 0.0028 0.7431 0.0063 1.8765 81.6 194.0 12.9362_2 LIMA5 0.2611 0.0045 0.8177 0.0053 2.0387 89.8 179.8 21.5362_2 LIMA6a 0.2882 0.0036 0.7780 0.0046 1.9681 85.4362_2 LIMA7 0.4348 0.0050 0.8548 0.0051 2.1195 93.8 180.3 36.5362_2 LIMA8 0.5095 0.0063 0.8505 0.0045 2.1245 93.4 226.9 41.5362_2 LIMA9a 0.2905 0.0038 0.7727 0.0048 1.9493 84.8362_2 LIMA10 0.2801 0.0033 0.8191 0.0047 2.0479 89.9 189.9 22.4362_2 LIMA13 0.3025 0.0034 0.8315 0.0044 2.0877 91.3 177.5 24.2

(continued on next page)

140 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Table 2 (continued)

Sample Label Major oxide comp.b

206Pb/238U ±2sd 207Pb/206Pb ±2sd 208Pb/206Pb %“common Pb”c 207Pb-corrected age (Ma)

±2sd

362M LIMA1 0.3584 0.0053 0.8480 0.0072 2.0986 93.1 166.8 32.8362M LIMA2a 0.1873 0.0029 0.7387 0.0085 1.8801 81.1362M LIMA3 0.4552 0.0059 0.8503 0.0075 2.1044 93.3 203.7 42.2362M LIMA4 0.3148 0.0039 0.8391 0.0071 2.0786 92.1 167.2 28.5362M LIMA5 0.5286 0.0084 0.8638 0.0076 2.1222 94.8 184.1 49.8362M LIMA5bis 0.4256 0.0068 0.8460 0.0082 2.0903 92.9 194.4 38.8362M LIMA6bis 0.4714 0.0065 0.8621 0.0079 2.1182 94.6 170.2 45.0362M LIMA7 0.4546 0.0039 0.8456 0.0054 2.0934 92.8 219.0 38.1362M LIMA8 0.4007 0.0036 0.8317 0.0057 2.0644 91.3 233.8 33.6min 0.1567 0.7387 1.8765 81.1 166.8max 0.5286 0.8638 2.1245 94.8 233.8weighted meanc 0.3420 0.0570 0.8360 0.0140 2.0568 91.9 190.1 8.7

Note: Rejected analyses not included.a Rejected analyses, based on >3sd difference compared to values on the regression line in Tera-Wasserburg plot.b “Major oxide comp.” indicates the major-oxide compositional group to which each grain belongs, where an EMP analysis is available; for xenolith XM1/362 no distinct

compositional groups were recognised.c % common Pb for each analysis is given by the ratio of measured 207Pb/206Pb to the initial 207Pb/206Pb (i.e. the upper intercept value of the regression line through

uncorrected U–Pb values).

cept value for the LIMA samples (Fig. 4). The major shortcoming of this approach is that the isotopic evolution of Pb in the crust is a) model dependent and b) may not be the same for con-temporaneous mantle rocks. The Tera-Wasserburg upper intercept 207Pb/206Pb values for the current LIMA samples (0.911–0.935) dif-fer significantly from the Stacey–Kramers 207Pb/206Pb values at ∼150–200 Ma (0.845–0.849). Therefore, the Stacey–Kramers model is deemed unsuitable for the LIMA samples.

A widely used method to correct for common Pb employs the upper intercept 207Pb/206Pb value of the Tera-Wasserburg regression line to improve precision on individual uncorrected analyses (e.g. Batumike et al., 2008; Chen and Simonetti, 2013;Chew et al., 2011; Simonetti et al., 2006; Cox and Wilton, 2006;Mitchell et al., 2011). However, Ludwig (2012b) pointed out that such approach is mathematically circular and cannot improve the precision or accuracy of the age estimates. For completeness, we have reported the ages for LIMA samples corrected using this method in Table 2, but do not discuss these 207Pb-corrected ages further. The U–Pb ages discussed below are Tera-Wasserburg lower intercept ages; no other common-Pb correction was applied.

4.5. Zircon U–Pb age and Hf isotopes

The zircon grain extracted from sample XM1/362 shows elon-gate prismatic habit (Fig. S3, Supplementary Material). The CL im-age of this grain is featureless. Three measurements of U–Th–Pb isotopic compositions were carried out on this zircon grain using the SHRIMP (Table 4). Both U and Th are high (∼4000 ppm and ∼5100 ppm, respectively) and the analyses are essentially free of common Pb. The three U–Pb analyses are concordant (Fig. S4) and yield 206Pb/238U ages of 181.6 ± 4.4–186.6 ± 3.8 Ma.

Two measurements of the Lu–Hf isotopic compositions provide present-day 176Hf/177Hf ratios of 0.28256 ±0.00011 and 0.28259 ±0.00012, and very high 176Lu/177Hf ratios (0.0038–0.0048; Ta-ble S3, Supplementary Material). Initial εHf values calculated for a zircon age of ∼184 Ma are of −3.5 and −2.4, respectively. These values are in the restricted εHf(i) field defined by zircon megacrysts from the Bultfontein kimberlite (−3.5–−1.7; Wood-head, unpublished data) and also fall within the broad εHf(i)field of southern African kimberlites (∼−6–+8; Nowell et al., 2004) and Karoo lavas (∼−11–+4; Ellam, 2006; Jourdan et al., 2007b – Fig. S5). Conversely, the εHf(i) values of the XM1/362 zircon differ from analyses of zircon grains in strongly metasoma-tised MARID (mica–amphibole–rutile–ilmenite–diopside; Dawson and Smith, 1977) mantle xenoliths from the Kimberley kimberlites

(∼−11–−19; Choukroun et al., 2005; Giuliani and Woodhead, un-published).

5. Discussion

Isotope-dilution analyses of LIMA grains and 207Pb-corrected time-resolved signals of in situ LIMA analyses suggest that reli-able U–Pb age determinations for titanates of the crichtonite se-ries can be obtained by LA-ICP-MS using zircon as the primary calibrant. In addition, we tested the accuracy of the LIMA U–Pb ages by dating a zircon grain from one of the samples (XM1/362). However, it should be reiterated that the zircon grains in xenolith XM1/362 have no clear textural relationship with the LIMA miner-als, clinopyroxene or phlogopite.

The zircon U–Pb age of 184 ± 6 Ma (sample XM1/362; Ta-ble 4) is consistent with crystallisation during the Karoo magmatic event that affected the southern African lithosphere in the early Jurassic (∼174–185 Ma – Jourdan et al., 2005, 2007a and refer-ences therein; Fig. 6). The Hf isotopic composition of the zircon (εHf(i) = −3.5–−2.9) suggests crystallisation from a melt derived from a mantle source with similar time-integrated Lu–Hf ratios to zircon megacrysts recovered from the Bultfontein kimberlite. Based on the ‘Karoo’ age and similar εHf(i) values to Karoo vol-canic rocks (∼−11–+4; Ellam, 2006; Jourdan et al., 2007b), the XM1/362 zircon may have crystallised from a Karoo magma that crystallised at mantle depths. However, U and Th concentrations in the zircon are much higher than that of most mantle zircons anal-ysed to date (Griffin et al., 2000; Hamilton et al., 1998; Heaman et al., 1990; Katayama et al., 2003; Kinny and Dawson, 1992;Kinny and Meyer, 1994; Kinny et al., 1989; Konzett et al., 1998, 2000, 2013; Liati et al., 2004; Robles-Cruz et al., 2012; Schärer et al., 1997; Simonetti and Neal, 2010; Spetsius et al., 2002;Zheng et al., 2006). The only exceptions are: i) zircons from man-tle xenoliths studied by Rudnick et al. (1999), which contained U and Th concentrations up to 3950 and 3300 ppm, respectively, and were assumed to be inherited grains from recycled crustal rocks; and ii) zircons in mantle pyroxenites from the Hannouba alkali basalts (China), which have U and Th contents up to 23 000 and 16 500 ppm, respectively, and were attributed to carbonate meta-somatism (Liu et al., 2010).

The very high U and Th concentrations of XM1/362 zircon might suggest a crustal rather than a mantle origin. It could be argued that the zircon crystallised in the Karoo dolerite intrusion that was later intersected by the Kimberley kimberlites (Clement, 1982; White et al., 2012). This interpretation is consistent with

A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147 141

Fig. 4. 207Pb/206Pb vs 238U/206Pb Tera-Wasserburg Concordia diagrams for in-situ LA-ICP-MS analyses of LIMA grains from xenolith samples XM1/341 (a), XM1/345 (b) and XM1/362 (c). The upper intercept on the Tera-Wasserburg diagrams represents the initial (i.e. common) 207Pb/206Pb ratio, whereas the lower intercept provides the average 206Pb/238U age (“intercept age”). The ellipses represent two sigma uncertainties; red and blue ellipses are accepted and rejected analyses, respectively. The results of solution-mode analyses for clinopyroxene (Cpxsol) from each sample and two LIMA separates (LIMAsol) from sample XM1/362 are also shown (uncertainties smaller than symbol sizes). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

142 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Fig. 5. Examples of time-resolved signals for XM1/345 LIMA analyses showing uncorrected 206Pb/238U ages and 207Pb (i.e. common Pb)-corrected 206Pb/238U ages. “Uncor-rected 206Pb/238U” ages were calculated assuming all 206Pb is radiogenic. “207Pb-corrected 206Pb/238U” ages were calculated using the initial 207Pb/206Pb (i.e. common Pb) value given by the upper intercept of Tera-Wasserburg Concordia plot for XM1/345 LIMA analyses (see Fig. 4b). Except for diagram a), the “uncorrected 206Pb/238U” time-resolved signals show non-flat trends with undulating shape. After common Pb subtraction, the “207Pb-corrected 206Pb/238U” time-resolved signals appear reasonably flat. This indicates that the variable shapes of “uncorrected 206Pb/238U” time-resolved signals are solely due to variations in common Pb composition (i.e. there is no difference in U–Pb down-hole fractionation between LIMA and zircon standard).

the high U and Th concentrations (up to 22 000 and 30 000 ppm, respectively) measured for some zircons from Karoo dolerite sills (Svensen et al., 2012). Alternatively, the zircon may have crys-tallised in the lower crust after impingement by Karoo melts and was later sampled by the ascending kimberlite magma. However, it is difficult to envisage how a crustal zircon xenocryst could be fortuitously entrapped in a mantle xenolith hosting LIMA minerals that yield indistinguishable U–Pb ages. Furthermore, the XM1/362 zircon displays a homogeneous CL structure, which is at odds with the oscillatory growth zoning common to magmatic zircons (Corfu et al., 2003). Therefore we suggest a mantle metasomatic origin for the XM1/362 zircon, as per the conclusion of Liu et al. (2010) for similar zircon occurrences.

We conclude that U–Pb dating of LIMA minerals provides a ro-bust and reliable method for determining the timing of mantle metasomatism, despite high common-Pb (∼90%) contents. A disad-vantage of the LIMA U–Pb dating method is the limited occurrence of LIMA minerals in metasomatised mantle xenoliths (Erlank et al., 1987; Haggerty et al., 1983; Jones et al., 1982; Konzett et al. 2000, 2013); the same situation applies to metasomatic zircons.

The LIMA U–Pb apparent ages of 177 ± 12, 178 ± 29 and 190 ±24 Ma (Fig. 6) are significantly older than the emplacement ages of the host Kimberley kimberlites (81–90 Ma – Allsopp and Barrett, 1975; Batumike et al., 2008; Davis, 1977; Fitch and Miller, 1983;Smith et al., 1989), as well as orangeite clusters in the region (Barkly West, 118–120 Ma; Kroonstad, 130–135 Ma; Swartruggens, 142–145 Ma – Phillips et al., 1998, 1999 and references therein – Fig. 1). Reported U–Pb ages of zircon grains in metasomatised mantle xenoliths from the Kimberley and nearby Kamfersdam kim-berlites, range between 80 and 142 Ma (Hamilton et al., 1998;Kinny and Dawson, 1992; Konzett et al., 1998, 2000, 2013), i.e. they approach, but do not overlap the LIMA ages (Fig. 6). However, Konzett et al. (1998, 2000) suggested that some resetting may have affected at least some of these zircon ages, with possible forma-tion ages as old as the Karoo magmatic event (∼174–185 Ma – Jourdan et al., 2005, 2007a and references therein; Fig. 6). This interpretation is also supported by a ∼170 Ma Re–Os model age for a MARID xenolith from Kimberley (Pearson et al., 1995a) and a ∼180 Ma Pb isotope model age for LIMA macrocrysts in the nearby Jagersfontein kimberlite (Griffin et al., 2014). The forma-

A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147 143

Table 3U–Th–Pb concentrations and isotopic compositions of LIMA and clinopyroxene separates analysed by solution mode.

Sample Label U (ppm)

Th (ppm)

Pb (ppm)

Th/U 238U/204Pb ±2sd (%)

206Pb/204Pb ±2sd (%)

207Pb/204Pb ±2sd (%)

XM1/341 341 Cpx 0.05 0.13 1.26 2.61 2.41 2.00 17.30 0.10 15.50 0.10

XM1/345 345 Cpx 0.09 0.26 2.09 2.82 2.76 2.00 17.54 0.10 15.51 0.10

XM1/362 362 Cpx 1.71 1.04 2.88 0.60 37.34 2.00 18.20 0.10 15.53 0.10362 LIMA 1 227 115 321 0.51 43.62 0.53 18.42 0.08 15.54 0.12362 LIMA 2 243 101 298 0.42 50.44 0.52 18.64 0.08 15.56 0.13

Sample Label 232Th/204Pb ±2sd (%)

208Pb/204Pb ±2sd (%)

235U/204Pba ±2sd (%)

238U/206Pb ±2sd (%)

207Pb/206Pb ±2sd (%)

XM1/341 341 Cpx 6.50 2.00 37.65 0.15 0.02 2.00 0.139 2.00 0.896 0.05

XM1/345 345 Cpx 8.04 2.00 37.83 0.15 0.02 2.00 0.158 2.00 0.885 0.05

XM1/362 362 Cpx 23.32 2.00 37.93 0.15 0.27 2.00 2.052 2.00 0.854 0.05362 LIMA 1 23.10 2.00 38.01 0.15 0.32 0.53 2.368 0.53 0.844 0.04362 LIMA 2 22.10 2.00 38.12 0.15 0.37 0.52 2.707 0.52 0.835 0.06

U–Th–Pb elemental concentrations measured by quadrupole ICP–MS; Pb concentrations based on the sum of 206Pb, 207Pb and 208Pb count rates.Pb isotopic ratios and, for LIMA only, U/Pb isotopic ratios analysed by MC–ICP–MS; Th/Pb ratios analysed by quadrupole ICP–MS.MC–ICP–MS Instrumental mass bias in Pb and U runs was corrected by standard bracketing with SRM981 standard (spiked Pb in LIMA), Tl-doping (unspiked Pb in clinopy-roxene; see Woodhead, 2002) and internal normalisation to 238U/235U = 137.88 (spiked U in LIMA). Uncertainties for Pb isotope ratios in clinopyroxene (in-run 2se up to ±0.19% for 206Pb/204Pb) are larger than typical external precisions of the Tl-doping method (≤ ±0.05% for total Pb signals near 10 V) because of the relatively small Pb signals obtained (≤6.3 V). USGS basalt BCR-2 yields 206Pb/204Pb = 18.759 ± 0.007 (±2sd), 207Pb/204Pb = 15.621 ± 0.010 and 208Pb/204Pb = 38.730 ± 0.034 (n = 22), con-sistent with TIMS and MC–ICP–MS reference values. The decay constants used are 0.155125 × 10−9/yr for 238U, 0.98485 × 10−9/yr for 235U and 0.049485 × 10−9/yr for 232Th.

a Calculated by multiplying 238U/204Pb values for 1/137.88.

Fig. 6. Age distribution for magmatic and mantle metasomatic events in the Kimberley region during the Jurassic and Cretaceous. The Kimberley kimberlites were emplaced between 81 and 90 Ma (Allsopp and Barrett, 1975; Batumike et al., 2008; Davis, 1977; Fitch and Miller, 1983; Smith et al., 1989). The Barkly West orangeites (or Group II kimberlites) were emplaced in close proximity to the Kimberely cluster (Fig. 1) at between 118–120 Ma (Phillips et al., 1999). The Kroonstad (e.g., Lace, Voorspoed) and Swartruggens clusters of orangeites, north–east of Kimberley (Fig. 1), were emplaced at ∼130–135 and 142–145 Ma, respectively (Phillips et al., 1998, 1999 and references therein). Bulk rock Rb–Sr analyses of strongly metasomatised, phlogopite ± K-richterite-rich peridotite (PKP and PP) xenoliths from the Kimberley kimberlites produced a crude isochron with an age of 144 ± 23 Ma (Hawkesworth et al., 1990). SHRIMP U–Pb studies of zircon in metasomatised mantle xenoliths from the Kimberley and adjacent Kamfersdam kimberlites produced ages between 80 and 142 Ma (Hamilton et al., 1998; Kinny and Dawson, 1992; Konzett et al., 1998, 2000, 2013). The emplacement ages of Karoo magmatic rocks (lavas, dykes and sills) are from Duncan et al. (1997), Encarnacion et al. (1996), Jourdan et al. (2005, 2007a and references therein) and Svensen et al. (2012).

tion of LIMA minerals at ∼180–190 Ma beneath the Kimber-ley area requires the migration of melts enriched in HFSE, LILE and LREE through the southern African lithospheric mantle prox-imal to the time of Karoo magmatism (Fig. 6); thus indicat-ing a probable link between mantle metasomatism and mantle melting.

It is debated whether the Karoo magmas formed within the enriched lithospheric mantle (Ellam, 2006; Ellam and Cox, 1989;Hawkesworth et al., 1984; Riley et al., 2006), or originated in the asthenosphere before contamination by metasomatised man-tle lithosphere during ascent (Ellam et al., 1992; Neumann et al., 2011; Sweeney et al., 1991). Consequently, it is possible that the LIMA-forming melts could have been a precursor to the Karoo magmas (i.e. they metasomatised the lithospheric mantle source of Karoo magmas); alternatively, the LIMA-forming melts might rep-

resent Karoo melts that underwent extreme incompatible elements enrichment during ascent into the lithospheric mantle.

6. Conclusions

The new geochronological data in this study provide the first robust evidence of fluid/melt enrichment in the lithospheric man-tle beneath the Kimberley region (South Africa) at the time of the widespread Karoo magmatic event. LIMA grains from three intensely metasomatised phlogopite-rich peridotite xenoliths from the ∼80–90 Ma Kimberley kimberlites show variable major oxide and U–Pb isotopic compositions. Regressions through LIMA anal-yses on Tera-Wasserburg plots produced apparent U–Pb ages of 177 ± 12, 178 ± 29 and 190 ± 24 Ma. The in situ U–Pb dating protocol was validated by solution-mode analyses of LIMA grains

144 A. Giuliani et al. / Earth and Planetary Science Letters 401 (2014) 132–147

Tabl

e4

Sum

mar

y of

SHRI

MP

U–P

b da

ta fo

r XM

1/36

2 zi

rcon

.

Labe

lU

(ppm

)Th

(p

pm)

232

Th23

8U

206

Pb*

(ppm

)%

206

Pbc

206

Pb23

8U

Age

(M

a)a

±2sd

238

U20

6Pb

(unc

orr.)

±2sd

207

Pb20

6Pb

(unc

orr.)

±2sd

206

Pb23

8U

a±2

sd23

8U

206

Pba

±2sd

207

Pb20

6Pb

a±2

sd20

7Pb

235

Ua

±2sd

362Z

_1.1

3448

4266

1.28

84.6

0.00

181.

64.

435

.00

2.40

0.04

961

1.00

0.02

862.

4035

.00

2.40

0.04

969

1.02

0.19

572.

6036

2Z_1

.246

2959

611.

3311

7 .0

0.00

186.

63.

834

.04

2.00

0.04

978

0.84

0.02

942.

0034

.05

2.00

0.04

974

0.86

0.20

142.

2036

2Z_1

.339

0750

841.

3497

.30.

0018

4.2

4.0

34.5

02.

200.

0495

31.

060.

0290

2.20

34.4

92.

200.

0496

51.

080.

1985

2.60

wei

ghte

dm

ean

3995

5104

1.32

117 .

018

4.4

6.2

Pbc

and

Pb* i

ndic

ate t

he co

mm

on an

d ra

diog

enic p

orti

ons, re

spec

tive

ly.

aCo

mm

on P

b co

rrec

ted

usin

g m

easu

red

204

Pb.

separated from one of the xenoliths. An independent constraint on the accuracy of LIMA U–Pb ages was given by the SHRIMP U–Pb age (184 ± 6 Ma) of a zircon extracted from LIMA-bearing sample XM1/362. The LIMA ages are consistent with metasomatism of the lithospheric mantle by melts enriched LILE, HFSE and LREE, coeval with Karoo magmatism. We conclude that LIMA (and crichtonite-series minerals) U–Pb dating by in situ LA-ICP-MS using zircon as an external standard is a new and promising method to date man-tle metasomatic events.

Acknowledgements

The Authors acknowledge Graham Hutchinson for support with microprobe analyses and Bence Paul for assistance with Iolite data reduction. Matthew Felgate is thanked for insightful discussion on U–Pb dating and for drafting one of the figures, and Bill Griffin for providing a preprint of his article on LIMA. The Authors ac-knowledge De Beers Consolidated Mines for providing access to the studied samples. Reviews by Fred Jourdan and Jan Kramers, and editorial handling by Bernard Marty are greatly appreciated. A.G.’s PhD research was supported by an International Australian Postgraduate Award and the Albert Shimmins Memorial Fund. This is contribution 462 from the ARC Centre of Excellence for Core to Crust Fluid Systems (http :/ /www.ccfs .mq .edu .au).

Appendix A. Supplementary material

Supplementary material related to this article can be found on-line at http://dx.doi.org/10.1016/j.epsl.2014.05.044.

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