In-situ SIMS U–Pb dating of phanerozoic apatite with low U and high common Pb

Gondwana Research 21 (2012) 745–756

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Gondwana Research 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 / g r

In-situ SIMS U–Pb dating of phanerozoic apatite with low U and high common Pb Qiu-Li Li a,⁎, Xian-Hua Li a, Fu-Yuan Wu a, Qing-Zhu Yin b, Hai-Min Ye a, c, Yu Liu a, Guo-Qiang Tang a, Chuan-Lin Zhang c a b c

State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Department of Geology, University of California Davis, One Shields Avenue, Davis, CA 95616, USA Nanjing Institute of Geology and Mineral Resources, Chinese Geological Survey, Nanjing 210016, China

a r t i c l e

i n f o

Article history: Received 28 February 2011 Received in revised form 15 June 2011 Accepted 6 July 2011 Available online 20 July 2011 Keywords: SIMS Apatite U–Pb In-situ Common Pb

a b s t r a c t Apatite is a common accessory mineral in various igneous, metamorphic and sedimentary rocks. It has strong potential to provide important information for geochronology and petrogenesis. However, precise U–Th–Pb dating of apatite, especially young samples, is difficult due to low abundances of U and less radiogenic Pb, posing significant challenges for accurate common Pb correction. An additional issue is the lack of suitable apatite to serve as a standard for in-situ analyses. Here we present both analytical and data reduction protocols for SIMS insitu U–Pb analyses on apatite with low U and high common lead. With NW-1 apatite as the U–Pb age standard, which was separated from ~1160 Ma carbonatite in the Prairie Lake complex in Canada, apatites from the Kovdor carbonatite in the Kola peninsula which contains low U (mostly ~2.5 ppm) and high common Pb (20–80%) yielded weighted average 207Pb-corrected U–Pb ages of 375 ± 13 Ma (KV-8 and KV-18) and 377 ± 11 Ma (KV-A), respectively. Apatites from the Quruqtagh ultramafic intrusion in NW China containing low U (~2.3 ppm) and high common Pb (N 70%) yield 207Pb-corrected U–Pb age of 805 ± 21 Ma. The Cenozoic Durango apatite was dated at 31± 2 Ma. All these new apatite U–Pb ages are indistinguishable from independently known age constraints to within 2–4%. Our results demonstrate that a U–Pb age can be accurately and precisely measured for apatite with low U (b 3 ppm) and high common Pb (N 50%) by SIMS with a suitable standard and a careful choice of common Pb composition. © 2011 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Apatite [Ca5(PO4)3(F,Cl, OH)] is a ubiquitous accessory mineral occurring in almost all major rock types. The crystal structure of apatite can accommodate a number of substitutions, such as Sr, Mn, Mg, REE, Hf, Pb, U and Th. Radioactive decay of U and Th in apatite makes it an ideal mineral not only for thermochronology studies by U/Th–He and fission track dating (e.g. Farley and Stockli, 2002; Gleadow et al., 2002), but also for dating the emplacement age of rapidly cooled plutonic rocks using the U–Pb system (e.g., Cherniak et al., 1991; Chamberlain and Bowring, 2000; Amelin and Zaitsev, 2002; Cherniak, 2005), especially when other minerals suitable for dating are not readily available. While apatite from felsic rocks (especially pegmatite and some hydrothermal rocks) can contain hundreds and up to thousands ppm of U, apatite from mantle-derived rocks and meteorites usually has a very low U content (b10 ppm). Although the isotope dilution

⁎ Corresponding author. Tel.: + 86 10 82998443; fax: + 86 10 62010846. E-mail addresses: [email protected], [email protected] (Q.-L. Li).

thermal ionization mass spectrometry (IDTIMS) method has been used for apatite dating in the past (e.g., Chamberlain and Bowring, 2000; Amelin and Zaitsev, 2002), the capability of in situ dating is desirable in many cases when the apatite must not be destroyed or its textural context is important (e.g. Nemchin, et al., 2009). Sano et al. (1999) first reported apatite U–Th–Pb analytical protocols by SHRIMP and applied the technique to a number of meteorite investigations (e.g. Terada and Sano, 2002; Terada et al., 2003, 2005, 2007, 2008; Terada and Bischoft, 2009). Recently, Chew et al. (2011) established analytical procedures by LA–ICPMS to obtain U–Pb and Th–Pb ages with uncertainties as low as 1–2% for Palaeozoic–Neoproterozoic samples. The majority of in-situ U–Pb age-determinations on apatites so far have primarily focused on samples with old ages and/or high U contents (i.e. favorable for dating purposes), whereas young samples or those with low U were frequently considered unsuitable and thus rarely investigated. In this study, we present both analytical and data reduction protocols for in-situ U–Pb analyses on apatite using a Cameca IMS 1280 SIMS. For apatite with b3 ppm U and N50% common Pb, we demonstrate that U– Pb ages can be measured accurately and reproducibly with adequate precision for application to geological problems.

1342-937X/$ – see front matter © 2011 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2011.07.008

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2. Samples 2.1. NW-1 apatite NW-1 apatite was extracted from a carbonatite collected from the Prairie Lake alkaline–carbonatite complex in Ontario, Canada, where the PRAP apatite standard was collected (Sano et al., 1999). This complex intruded the Archean biotite–quartz–feldspar paragneiss country rock, and consists of ijolite, biotite pyroxenite, carbonatite, phlogopite carbonatite and alkaline syenite (Woolley, 1987). There have been many geochronological investigations conducted on this complex over the years, including a Pb–Pb age of 1155 ± 36 Ma for calcite from the carbonatite (Kwon et al., 1989), a SHRIMP Pb–Pb age of 1156 ± 45 Ma for the apatite from the ijolite (Sano et al., 1999), a U–Pb TIMS result of 1163.5 ± 3.5 Ma for baddeleyite and zircon from the phoscorite (Rukhlov and Bell, 2010) and a 207Pb/206Pb age of 1159 ± 5 Ma of calzirtite from the same sample (Wu et al., 2010). Therefore, the best estimate for the U–Th–Pb age of NW-1 apatite is suggested to be 1160 ± 5 Ma. 2.2. Kovdor apatites The Kovdor apatites were sampled from phoscorites (magnetite– forsterite–apatite–calcite rocks) and carbonatites of the Kovdor ultramafic–alkaline–carbonatite Massif in the southwestern Kola Peninsula, Russia (Amelin and Zaitsev, 2002). Previous geochronological investigations of the Kovdor Massif includes apatite U–Th–Pb ages of 376–380 Ma, zircon U–Pb age of 377.52 ± 0.94 Ma, baddeleyite U–Pb ages between 378.54 ± 0.23 Ma and 382 ± 3 Ma by TIMS (Amelin and Zaitsev, 2002), apatite U–Pb age of 386 ± 8.6 Ma and Th–Pb age of 380 ± 3.1 Ma by LA–ICPMS (Chew et al., 2011), baddeleyite U–Pb age of 381.4 ± 4.5 by SIMS (Li et al., 2010a) and a zirconolite Pb–Pb age of 380 ± 5 Ma (Wu et al., 2010). There are three apatite samples (KV-8, KV-18 and KV-A) from the Kovdor massif used in this study. KV-8 and KV-18 apatites used in this study are from the same sample analyzed by Amelin and Zaitsev (2002). According to their TIMS analyses, the KV-8 apatite has a U–Pb age of 382 ± 1.5 Ma and Th–Pb age of 376.3 ± 1.1 Ma, with U content of 2.8 ppm, Th content of 62.4 ppm and Th/U ratio of 22.45. The KV-18 apatite yielded a U–Pb age of 378.6 ± 1.7 Ma and a Th–Pb age of 376.5 ± 1.2 Ma, with U content of ~ 3.4 ppm, Th content of ~ 150 ppm and Th/U ratios of 41–45 (Amelin and Zaitsev, 2002). 2.3. Durango apatite Durango apatite occurs as coarse crystals within the open pit iron mine at Cerro de Mercado, Durango, Mexico. It is widely used as a standard for fission track and (U–Th)/He dating. K–Ar dating of the host rock and 40Ar– 39Ar dating of host sanidine has yielded ages between 31.0 and 31.6 Ma, consistent with the 40Ar–39Ar age of 31.44± 0.18 Ma and (U–Th)/He age of 31.02 ± 2.02 Ma for apatite (McDowell et al., 2005). Recently, Chew et al.(2011) obtained 238U–206Pb age of 30.6 ± 2.3 Ma and 232Th– 208Pb age of 32.5 ± 1.2 Ma on Durango apatite by LA– ICPMS. This apatite has been used to serve as a standard for major element determination by electron microprobe, and was recently proposed as a potential standard for trace element in situ analyses (Frei et al., 2005; Trotter and Eggins, 2006; Morishita et al., 2008). In this study, the Durango apatite was used as the U concentration standard, with a U content of 9 ppm (determined by the solution-ICPMS method, Trotter and Eggins, 2006), to calibrate U concentrations of unknowns. 2.4. Qieganbulake (QK) apatite These apatites were separated from the Qieganbulake ultramafic– mafic–carbonatite complex in Quruqtagh of the northeastern Tarim Block, NW China. This complex hosts the second-largest vermiculite

deposit in the world, as well as a medium-size phosphate ore deposit. It is considered that the apatite and baddeleyite co-crystallized from the complex, therefore baddeleyite was used to constrain the timing of mineralization, with a TIMS U–Pb age of 810 ± 6 Ma (Zhang et al., 2007) and a SIMS Pb–Pb age of 814.7 ± 3.6 Ma (Li et al., 2010a). The apatite samples studied here (QK-1 and QK-2) occur as big crystals (~2 cm in diameter) in pyroxenite from the same complex. Direct dating of the mineralization on apatite has never been performed due to its low U content and high common Pb. 3. Instrumentation and analytical procedure U–Th–Pb analyses were performed using the Chinese Academy of Sciences Cameca IMS 1280 (CASIMS) at the Institute of Geology and Geophysics in Beijing. The O2− primary ion beam was accelerated at −13 kV, with an intensity of ~10–12 nA. The aperture illumination mode (Kohler illumination) was used with a 200 μm primary beam mass filter (PBMF) aperture to produce even sputtering over the entire analyzed area. The ellipsoidal spot is about 20 × 30 μm in size. Positive secondary ions were extracted with a 10 kV potential. The samples analyzed in this study were cast in epoxy mounts. The mounts were coated with about 30 nm of high-purity gold to reach b20 Ω resistance. Sample charging effects were minimized by optimizing the energy offset to maximum transmission in a 60 eV energy window at the start of each analysis, using the 40Ca231P16O3+ reference peak at mass 159. The 40Ca231P16O3+ peak is also used as a reference peak for centering the secondary ion beam, energy and mass adjustments. The instrument was operated in high transmission (Field aperture = 5000, Max area = 40). Rectangular lenses were activated in the secondary ion optics to increase the transmission at high mass resolution (De Chambost et al., 1996). A mass resolution of ~9000 (defined at 50% peak height) was used to separate 40Ca231P16O3 peaks from isobaric interferences. This mass resolution also is sufficient to separate U, Th and Pb isotopes from isobaric interferences, such as oxides of rare earth elements (e.g., Williams, 1998; Sano et al., 1999). A single electron multiplier was used in ion-counting mode to measure secondary-ion beam intensities by a peak jumping sequence, including isotopes of Pb +, Th+, U +, ThO +, UO+, UO2+ and 40Ca231P16O3+ to produce one set of data. Each measurement consists of 10 cycles, and the total analytical time is ca. 16 min. The mass fractionations of Pb isotopes and Pb hydrides (requiring a mass resolution N30,000) were considered insignificant because a number of studies have shown that these two effects are negligible and there appears to be a mutual cancelation effect (e.g. Ireland et al., 1990; Whitehouse et al., 1997; Williams, 1998; Ireland and Williams, 2003; Li et al., 2009). 4. Calibration protocols 4.1. Pb/U calibration Significant inter-elemental fractionation between Pb and U exists in SIMS measurements. The apparent 206Pb+/ 238U+ values determined using a negative oxygen primary beam are typically many times greater than the true values of the crystalline target. The ratio is clearly matrix dependent, for example, 2–4 times for zircon (e.g. Stern and Amelin, 2003), 30–60 times for perovskite (Li et al., 2010b), 8–170 times for rutile (Li et al., 2011), 0.5–0.8 times and 7–10 times for baddeleyite with and without oxygen flooding, respectively (Li et al., 2010a). Hence, a matrix-matched standard is vital for SIMS U–Pb analyses. In this study, we use the 1160 Ma NW-1 apatite as standard, which has the lowest proportion of common Pb and highest U and radiogenic Pb contents among the dated apatite samples. The measured 206Pb +/ 238U+ values of apatite are usually 3 to 5 times higher than the true values. The linear, quadratic, or power functional relationship between 206Pb+/ 238U+ and [UOx] +/U+ (sometimes Pb +/UO + and UO2+/UO+) has been used by different investigators on different minerals (e.g. Williams, 1998). As

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indicated by SIMS rutile U–Pb analyses (Li et al., 2011), the power law [ 206Pb +/238U+ = A × ([UOx] +/238U +) E] should be the best choice and has been widely used for U–Pb dating on many minerals with different exponents by ion microprobe analyses. The 206Pb/238U of an unknown can be calibrated as: P b
U u = 206 P b
238 U st 206

238

206 P b
+ P b
+ þ 238 þ U u = U u! 206 238 þ E P b
+ UO A × 238 þx 238 þ U st U 206

238

ð1Þ

u

where, u denotes the unknowns; st denotes the standard; * denotes radiogenic Pb and x = 1 or 2. The 206Pb*/ 238U ratio of the NW-1 standard is calculated by the radiogenic function of e λt–1, where λ is the decay constant of 238U, and t (= 1160 Ma) is the age of apatite corroborated by many other minerals/methods (Sano et al., 1999; Rukhlov and Bell, 2010; Wu et al., 2010) and our Pb–Pb date (see below). As for apatite U–Pb analyses by SHRIMP, Sano et al. (1999) adopted a power law relationship between Pb/U and UO/U. For our analyses by a Cameca IMS 1280, we found a better correlation between Pb/U and UO2/U instead of UO/U, using NW-1 apatite as a standard (Fig. 1A). The exponential E, given by the slope in the linear relationship between ln(Pb/U) and ln(UO2/U) was 1.26, similar to zircon (Li et al., 2009). The external reproducibility obtained from the NW-1 apatite during the analytical session was propagated together with the statistical uncertainty on the unknown to give an overall error for the 206Pb/ 238U ratio of individual analyses. The correlation between 208Pb*/Th and ThO/Th was also constructed (Fig. 1B) and used as a power law relationship similar to that for allanite (Gregory et al., 2007) and perovskite (Li et al., 2010b). However, the NW-1 apatite is low in Th/U and high in common 208Pb, resulting in a large standard deviation (N6%) for the calibration line. Therefore, we do not consider this apatite as a suitable Th–Pb standard. 4.2. Th/U calibration The Th/U ratio is an important parameter for U- and Th-bearing minerals. For instance, Th/U ratio in zircon can be used as an index to discriminate among different genetic types of zircon (e.g. Rubatto et al., 1999). Assuming that both the Pb–U and Pb–Th isotopic systems in a mineral are concordant, the radiogenic 208Pb/ 206Pb for the analyzed domain is related to the local 232Th/ 238U by: 208




Pb = 206 P b


  T h eλ232 t −1   238 U eλ238 t −1

Fig. 1. Calibration diagram: (A) between secondary 206Pb*+/238U+ and 238U16O2+/238U+ ratios in the NW-1 apatite, while inset shows all the samples; (B) between secondary 208 Pb*+/232Th+ and 232Th16O+/232Th+ ratios in the NW-1 apatite. Error assigned to the symbol is 1σ.

232

ð2Þ

where λ232 is the 232Th decay constant (4.9475 × 10− 11 yr− 1) and λ238 is the 238U decay constant (1.55125 × 10− 10 yr− 1) (Steiger and Jäger, 1977). A plot of radiogenic 208Pb/206Pb versus 232Th/238U defines a straight line with a slope of (eλ232t − 1)/(eλ238t − 1), which is constant for a given age (t). Commonly, Th/U is determined from ThO +/UO + by a fixed calibration, for example, Th/U = 1.11 × ThO +/UO + for zircon (Compston et al., 1984). A similar calibration, Th/U = 1.02 × ThO +/UO +, was used for perovskite (Ireland et al., 1990), which was revised to Th/U = 1.14 × ThO+/UO+ by Li et al. (2010b). Fig. 2 shows the correlation in the NW-1 data between ThO+/UO+ and Th/U calculated in terms of 208Pb*+/206Pb*+ and employing the corresponding recommended U–Pb ages by previous studies (Sano et al., 1999; Rukhlov and Bell, 2010; Wu et al., 2010). These data define a linear correlation (R2 = 0.994) with a slope of 0.87 between ThO+/UO+ and Th/U, which is quite different to that of zircon (1.11, Compston et al., 1984; Williams, 1998), and perovskite (1.02 or 1.14, Ireland et al., 1990; Li et al., 2010b), indicating the presence of matrix effects. In this study, the Th/U ratios of unknowns were thus calculated using a linear relationship given by Th/U = 0.87 × ThO +/UO +.

Fig. 2. Linear regression of the adjusted Th/U vs. ThO+/UO+ for the NW-1 apatite. The socalled “adjusted” Th/U data were calculated based on Eq. (2), using the measured 208 Pb/206Pb, 204Pb-based common Pb correction method and the assigned age of 1160 Ma.

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The radiogenic 206Pb*/ 238U ratio can then be calculated from the measured 206Pb/ 238U using the relationship:

4.3. Calibration of U, Th and Pb concentrations The U, Th and Pb concentrations are additional pieces of useful information for characterization of minerals commonly measured during ion microprobe analysis for U–Th–Pb geochronology. The Pb and Th concentrations can be calculated from the 206Pb/ 238U, Th/U ratio and the Pb isotopic composition, if the U concentration is known. Thus, precise determination of the U concentration is fundamental. In general, the U concentration is calculated based on the ratio of UOx +(x = 0, 1, 2) and the intensity of a matrix ion, such as 90Zr2O + for zircon and CaTi2O4+ for perovskite (Ireland et al., 1990; Li et al., 2010b). In this study, we also observed that the UO2+/Ca2PO3+ ratio for Durango apatite with relatively a uniform U content is homogenous within 10% uncertainty during a single analytical session. Thus, the U concentration for each analysis can be calibrated based on the UO2+/Ca2PO3+ ratio using a sensitivity factor derived from the Durango data in terms of its U content (9 ppm, Trotter and Eggins, 2006).

Pb
= ð1−f206 Þ 238 U

206

! Pb λ t = e 238 −1 238 U measured

206

ð4Þ

from which, the common Pb corrected 206Pb*/ 238U ages are calculated. The only exception to this practice was for Durango apatites whose data are too clustered with low f206 (i.e. far away from the common 207 Pb/ 206Pb intercept on the y-axis of Tera–Wasserburg reverse Concordia plot) to construct a regression line, so the terrestrial Pb isotope composition (Stacey and Kramers, 1975) was used instead to estimate the common Pb composition. 5. Results

4.4. Common Pb correction Apatite contains variably high proportion of common lead, ranging from a few to tens of percent levels of total Pb. Therefore, common-Pb correction in the apatite analyses is crucial to U–Pb and Th–Pb age calculations. For ion microprobe analyses, common Pb in the measured minerals includes the initial Pb incorporated into the mineral during its crystallization and that derived from the anti-charging coating on the target surface and contaminants during sample preparation (Williams, 1998). The contribution from surface Pb can be reduced to insignificant levels by pre-sputtering of the primary ion beam over each target area prior to analysis. During U–Pb analyses of zircon, total 204Pb ion intensities are usually lower than 0.1 cps (e.g. Li et al., 2009), indicating that the contribution from surface Pb is very minor. Apart from Durango apatite, which contains very low concentrations of common Pb (Fig. 3), the intensities of 204Pb peak for other apatites range from 0.2 to 2 cps, suggesting that the initial Pb in apatites dominates the total common Pb. Highly variable proportions of common Pb can be used to deduce the precise common Pb composition by projecting a regression line through the uncorrected data on a Tera–Wasserburg plot. Then, a 207Pb-based common lead correction was performed, using the following relationship (Williams, 1998; Chew et al., 2011): ! Pb − 206 P b measured ! 207 Pb − 206 P b common

207 206

f206 =

206

P bcommon = P btotalðmeasuredÞ

! P b
206 P b
!: 207
Pb 206 P b
207

ð3Þ

Fig. 3. Diagram showing 204Pb intensities in counts per second (x-axis) and the proportion of common Pb in total Pb (ƒ206) (y-axis) for the apatites investigated in this study.

All ages were calculated using the decay constants recommended by Steiger and Jäger (1977) and the calculation routines of Isoplot/Ex (Ludwig, 2001). The resulting regression and weighted average ages are quoted at the 95% confidence interval, except where noted otherwise. 5.1. NW-1 apatite A total of 33 analyses were performed on multiple grains of NW-1 apatite. As shown in Table 1, this apatite shows variable uranium concentrations from 18 to 162 ppm, but rather homogenous thorium contents of 38± 6 ppm (1 SD) and common lead contents of 0.75 ± 0.2 ppm (1 SD). Among our many dating attempts on many apatite samples, this apatite is unique with its high U content and low Th/U ratios of 0.25–2. The 204Pb-based ƒ206 and ƒ208 are 0.5%–4.4% and 9%– 25%, respectively (Table 1 and Fig. 3). Fig. 4 shows a tight correlation between the measured 204Pb/206Pb and 207Pb/206Pb ratios for the NW-1 apatite. A least-square regression of these data based on the York method (York, 1966) gives an intercept of 0.0769± 0.0029, corresponding to a 207Pb*/ 206Pb* age of 1118 ± 76 Ma (2σ). Using the terrestrial Pb (Stacey and Kramers, 1975) as an estimate of the common-lead composition, a weighted mean 207Pb/ 206Pb age of 1160 ± 36 Ma (n= 33, MSWD = 1.2, Fig. 4 inset) can be obtained, which is consistent with the reported ages of the host rock complex (Kwon et al., 1989; Sano et al., 1999; Rukhlov and Bell, 2010; Wu et al., 2010). 5.2. Kovdor apatite The U–Th–Pb results, including 32 analyses on KV-8 apatite and 15 analyses each on KV-18 and KV-A apatite, are presented in Table 2 and plotted on the Tera–Wasserburg U–Pb Concordia diagram (Fig. 5). The KV-8 apatites have rather uniform uranium contents of 1.9 ± 0.2 ppm, Th contents of 45 ± 4 ppm, Th/U ratios of 24.3 ± 1.5 and common Pb of 0.40± 0.05 ppm (1 SD). Because of the low concentrations of 204Pb, the ƒ206 values are calculated using the measured 207Pb/206Pb ratios (Eq. (3)) yielding values of 53 ± 5% (1 SD, Fig. 3), and ƒ208 values of 25 ± 3% (1 SD). The U–Pb data points cluster tightly on the Tera– Wasserburg Concordia, resulting in imprecise lower and upper intercepts (Fig. 5A). Compared with the KV-8 apatite, KV-18 apatites yield a slightly higher U contents of 2.4 ± 0.5 ppm, but much higher common Pb contents of 1.2 ± 0.2 ppm (except KV-18@4 = 2.9 ppm). It has Th contents ranging from 52 to 197 ppm with Th/U ratios ranging from 30 to 65 (Table 2). The 207Pb-based ƒ206 values vary from 63 to 82%, and ƒ208 values from 24 to 44% (Table 2, Fig. 3). Combining these two apatite samples, a total of 47 U–Pb data were plotted on the Tera–Wasserburg diagram and yielded a lower intercept U–Pb age of 376 ± 32 Ma, and an upper intercept of 207Pb/ 206Pb = 0.843 ± 0.044 (2 σ) (Fig. 5A). With this ratio as the best estimate

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Table 1 U–Th–Pbc concentrations and Pb isotopic compositions of NW-1 standard apatite. Sample spot # NW-1@1 NW-1@2 NW-1@3 NW-1@4 NW-1@5 NW-1@6 NW-1@7 NW-1@8 NW-1@9 NW-1@10 NW-1@11 NW-1@12 NW-1@13 NW-1@14 NW-1@15 NW-1@16 NW-1@17 NW-1@18 NW-1@19 NW-1@20 NW-1@21 NW-1@22 NW-1@25 NW-1@26 NW-1@27 NW-1@28 NW-1@29 NW-1@30 NW-1@31 NW-1@32 NW-1@33

[U]a ppm

[Th] ppm

Th/U

37 30 46 55 109 58 45 49 46 18 34 58 150 50 64 117 162 46 43 60 46 44 127 50 122 94 118 152 113 62 59

30 33 56 53 35 31 26 31 37 36 41 37 37 38 38 37 50 39 40 39 42 41 31 33 43 35 42 40 42 34 39

0.83 1.11 1.23 0.97 0.32 0.54 0.58 0.64 0.79 2.01 1.22 0.63 0.25 0.77 0.6 0.32 0.31 0.86 0.92 0.65 0.9 0.93 0.25 0.67 0.36 0.38 0.35 0.27 0.37 0.55 0.67

[Pb]c ppm

ƒ206b %

ƒ208b %

204

Pb

206

Pb

0.54 0.93 0.76 0.62 0.97 0.61 0.41 0.8 0.59 0.38 0.59 0.91 0.65 0.55 0.93 0.59 0.51 0.88 0.72 0.67 0.78 0.88 1.22 0.72 1.28 0.83 0.84 0.83 0.73 0.85 0.8

2.1 4.4 2.3 1.6 1.3 1.5 1.3 2.3 1.8 2.9 2.4 2.2 0.6 1.6 2.1 0.7 0.5 2.7 2.4 1.6 2.4 2.8 1.4 2.0 1.5 1.3 1.0 0.8 0.9 1.9 1.9

14 24 12 11 24 16 13 21 14 9.4 13 19 15 12 21 15 10 19 16 16 16 18 29 18 25 20 17 18 16 21 18

0.00122 0.00255 0.00137 0.00094 0.00074 0.00088 0.00076 0.00135 0.00106 0.00172 0.00143 0.00129 0.00037 0.00091 0.00122 0.00042 0.00027 0.00157 0.00139 0.00094 0.00141 0.00164 0.00081 0.00120 0.00088 0.00074 0.00060 0.00046 0.00054 0.00114 0.00113

±σ %

207

Pb

206

Pb

20 18 19 22 17 20 23 21 28 24 19 16 20 19 17 25 21 19 19 22 19 17 17 18 16 17 18 23 23 23 19

0.1005 0.1041 0.1033 0.0971 0.0859 0.0952 0.0911 0.0926 0.0987 0.1121 0.1055 0.0918 0.0854 0.0993 0.0901 0.0836 0.0828 0.1010 0.0978 0.0927 0.1055 0.0940 0.0864 0.0943 0.0883 0.0872 0.0879 0.0859 0.0861 0.0908 0.0932

±σ %

208

Pb

206

Pb

2.5 2.7 2.2 2.0 1.5 2.0 2.3 2.2 2.3 3.1 2.5 1.9 1.7 2.1 2.2 1.6 1.4 2.4 2.4 2.1 2.3 2.4 1.6 2.2 2.6 1.6 1.5 1.4 1.5 2.0 2.3

0.312 0.387 0.416 0.317 0.113 0.198 0.209 0.224 0.272 0.668 0.409 0.222 0.089 0.272 0.212 0.108 0.104 0.302 0.319 0.222 0.329 0.329 0.101 0.240 0.126 0.135 0.127 0.094 0.127 0.201 0.225

±σ %

207 Pb/206Pbc Age (Ma)

±σ (Ma)

2.0 3.0 1.4 2.0 2.0 1.7 2.0 1.8 1.8 1.8 2.1 1.7 1.6 1.6 1.6 1.7 2.3 1.8 1.9 1.7 1.7 1.9 1.9 1.7 1.7 1.9 1.5 1.6 1.5 1.9 2.0

1269 835 1286 1284 1073 1260 1203 1016 1282 1375 1318 1017 1200 1344 998 1135 1170 1157 1141 1176 1325 935 1061 1123 1083 1109 1181 1180 1157 1051 1117

119 177 106 92 71 89 101 113 108 152 122 98 73 93 109 69 60 122 120 98 108 130 73 109 120 75 66 59 67 100 111

Error assigned to the ratio is one sigma estimated by counting statistics. a U concentrations are determined using the corresponding variations of UO2+/40Ca231P16O3+ ratios of Durango standard apatite with average 9 ppm U. b ƒ206 and ƒ208 denote the fraction of common 206Pb and 208Pb of total measured 206Pb and 208Pb (Williams, 1998; Chew et al., 2011) where (206Pb/204Pb)common and (208Pb/206Pb)common is derived with the Stacey and Kramers (1975) model. c Common Pb-correction were made using Pb compositions calculated using Stacey and Kramers (1975) model at corresponding t.

of common Pb composition, the 207Pb-corrected 206Pb/ 238U results yield a weighted average age of 375 ± 13 Ma (inset of Fig. 5A). Based on 15 measurements on KV-A apatites, this sample shows highly variable U contents ranging from 1.7 to 16.9 ppm, Th contents of 295–1715 ppm and Th/U ratios of 80–233. Interestingly, this sample contains a rather homogenous common Pb content of 1.0 ±

0.1 ppm (1 SD, Table 2), but ƒ206 values vary from 20% to 73% and are highly dependent on the U contents (Table 2, Fig. 3). Because of the high Th contents, the ƒ208 values are as low as 2–9%. The highly variable proportions of common Pb produce a wide spread in the distribution of the data points on the Tera–Wasserburg Concordia, yielding a good linear regression line with an upper intercept of 207Pb/ 206 Pb = 0.841 ± 0.047 and a lower intercept age of 378 ± 16 Ma (2 σ) (Fig. 5B). With the upper intercept employed as the common Pb composition, 207Pb-corrected 206Pb/ 238U ages result in a weighted average of 377 ± 11 Ma (Fig. 5B inset). All these three dating results of Kovdor apatites are in good agreement with the reported TIMS and LA–ICPMS U–Pb and Th–Pb ages (Amelin and Zaitsev, 2002; Chew et al., 2011). 5.3. Durango apatite

Fig. 4. 207Pb/206Pb–204Pb/206 Pb isochron for NW-1 standard apatite. The inset shows 207 Pb–206Pb ages based on 207Pb*/206Pb* ratios corrected for common Pb composition estimated using Stacey and Kramers (1975) model.

Eighteen measurements were conducted on different fragments from one Durango apatite crystal. The U content varies between 8.4 and 12.3 ppm with a 1 SD uncertainty at ±1 ppm level (Table 2) (here, U contents are calculated by self-calibration with average U = 9 ppm, Trotter and Eggins, 2006). Th/U ratios are 19.8 ± 0.8 (1 SD). This apatite contains a very low common Pb content of 0.052 ± 0.013 ppm (Table 2). However, the values of ƒ206 are relatively high, ranging from 17% to 41%, because of its young age (Table 2, Fig. 3). Values of ƒ208 are comparatively lower, between 7 and 18%, because of much higher Th contents. As shown in Fig. 6, the measured data points cluster on the Tera–Wasserburg plot, with a lower intercept age of 32 ±2 Ma with an assigned common 207Pb/ 206Pb composition of 0.84 based on the two stage terrestrial Pb evolution model of Stacey and Kramers (1975) anchored on the y-axis (Fig. 6). Accordingly, the

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Table 2 U–Th–Pb analytical results of apatite samples. Sample spot # KV-8@1 KV-8@2 KV-8@3 KV-8@4 KV-8@5 KV-8@6 KV-8@7 KV-8@8 KV-8@9 KV-8@10 KV-8@11 KV-8@12 KV-8@13 KV-8@14 KV-8@15 KV-8@16 KV-8@17 KV-8@18 KV-8@19 KV-8@20 KV-8@21 KV-8@22 KV-8@23 KV-8@24 KV-8@25 KV-8@26 KV-8@27 KV-8@28 KV-8@29 KV-8@30 KV-8@31 KV-8@32 KV-18@1 KV-18@2 KV-18@3 KV-18@4 KV-18@5 KV-18@6 KV-18@7 KV-18@8 KV-18@9 KV-18@10 KV-18@11 KV-18@12 KV-18@13 KV-18@14 KV-18@15 KV-A@1 KV-A@2 KV-A@3 KV-A@4 KV-A@5 KV-A@6 KV-A@7 KV-A@8 KV-A@9 KV-A@10 KV-A@11 KV-A@12 KV-A@13 KV-A@14 KV-A@15 Drango@1 Drango@2 Drango@3 Drango@4 Drango@5 Drango@6 Drango@7 Drango@8 Drango@9 Drango@10 Drango@11 Drango@12

[U]a ppm

[Th] ppm

Th/U

1.7 1.6 1.7 1.7 2.3 2.0 1.8 1.8 1.8 1.7 1.9 2.0 2.0 1.7 1.9 1.7 1.7 1.7 2.0 1.7 2.1 2.0 1.8 1.8 2.2 2.1 2.1 1.6 1.9 1.9 2.0 1.8 3.3 1.7 3.0 3.1 2.0 2.0 2.5 3.1 2.0 2.3 2.2 1.8 2.6 2.9 1.8 16.1 15.4 1.7 5.1 6.1 4.0 4.3 7.7 3.8 4.0 15.7 16.9 1.3 4.4 3.9 8.8 9.4 9.0 9.0 8.7 8.5 8.2 8.5 9.3 12.3 11.5 8.8

46 46 40 42 59 47 46 45 47 44 48 48 50 44 46 41 42 45 46 44 47 48 44 42 51 46 50 35 45 41 44 44 132 87 119 197 73 75 120 110 74 103 104 56 96 90 52 1660 1634 384 726 841 517 344 1116 585 698 1715 1639 295 683 341 178 184 181 186 176 176 156 168 172 221 211 179

27.7 28.4 24 24.7 25.1 23.7 25.2 24.8 25.8 25.7 25 23.4 24.9 25 23.8 24.6 23.9 25.7 23.5 25.4 21.9 23.5 24.1 23.5 23.5 21.7 23.4 22.1 23.6 22 21.8 24.6 40 52.3 40.1 64.6 36.3 36.8 48.7 35.9 36.5 44.4 46.7 30.5 36.4 30.6 29.5 103 106 233 143 138 128 80 145 155 175 109 97 224 154 88 20.2 19.6 20.1 20.6 20.3 20.7 19 19.7 18.6 18 18.4 20.3

Pbc ppm

ƒ206b %

ƒ208b %

0.4 0.5 0.5 0.5 0.4 0.4 0.4 0.4 0.5 0.5 0.5 0.5 0.4 0.4 0.5 0.4 0.5 0.5 0.6 0.4 0.5 0.4 0.4 0.4 0.5 0.4 0.5 0.5 0.5 0.5 0.5 0.5 1.6 1.4 1.3 3 1 1.2 1.1 1.1 0.9 1 1.1 1.3 1.2 1.3 1.1 1.1 1.1 1 0.8 1 1 1 1 1.1 1.2 1 0.9 0.9 1 1 0.06 0.05 0.05 0.06 0.04 0.05 0.05 0.03 0.04 0.05 0.05 0.06

54 64 62 55 47 47 53 47 57 60 60 48 52 57 51 51 55 58 58 51 51 49 53 59 53 47 51 60 60 59 53 54 69 79 72 84 68 74 72 65 69 65 70 79 66 73 76 26 26 73 44 45 58 51 38 61 61 22 20 70 50 56 22 23 30 29 19 28 31 17 18 18 19 30

23 28 29 27 21 22 22 23 29 29 26 22 25 25 27 23 26 28 28 24 24 25 26 29 26 23 23 31 27 30 26 27 29 33 31 38 32 35 25 26 31 25 25 45 29 35 39 2 2 8 4 4 6 9 3 6 6 2 2 9 5 9 9 9 8 11 9 10 10 7 7 7 7 10

238 206

U

Pb

7.53 7.43 7.28 6.96 8.50 8.25 8.16 7.74 6.91 7.33 7.45 7.43 8.81 8.45 6.91 7.44 7.03 6.89 7.07 7.72 8.45 8.31 8.14 8.43 8.72 8.15 8.01 6.53 8.01 7.28 8.23 7.02 5.04 3.27 5.54 2.93 4.70 4.41 5.39 5.96 5.17 5.50 4.83 3.97 5.16 5.62 4.30 12.7 12.1 4.19 9.47 9.69 7.76 7.70 10.3 7.53 7.26 12.7 13.0 3.69 7.99 7.20 122 151 177 151 156 162 180 166 149 164 156 154

±σ %

207

Pb

206

Pb

5.6 5.6 6.7 5.5 7.3 5.3 5.3 5.6 5.5 5.5 5.5 6.1 5.2 5.1 5.1 5.5 5.4 5.4 6.8 5.4 4.8 4.9 5.9 5.1 4.9 4.8 5.2 5.2 5.9 6.1 5.1 5.0 3.3 4.8 3.8 11.0 4.3 4.1 4.0 5.0 4.2 4.1 4.3 4.3 3.9 3.6 6.1 3.1 2.5 5.4 4.1 3.6 4.1 5.2 4.0 4.3 4.7 3.3 3.2 5.8 6.2 4.1 10 11 12 11 11 11 12 11 11 10 10 11

0.48 0.55 0.54 0.49 0.42 0.42 0.47 0.42 0.5 0.52 0.52 0.43 0.46 0.5 0.46 0.46 0.49 0.51 0.510 0.45 0.46 0.44 0.47 0.52 0.47 0.420 0.450 0.52 0.52 0.52 0.47 0.480 0.6 0.67 0.62 0.71 0.59 0.64 0.62 0.560 0.59 0.57 0.6 0.67 0.57 0.62 0.65 0.26 0.26 0.63 0.4 0.4 0.51 0.45 0.35 0.53 0.53 0.23 0.21 0.61 0.45 0.49 0.23 0.23 0.29 0.28 0.2 0.27 0.3 0.18 0.2 0.190 0.2 0.29

±σ%

207

Pb-corrcAge (Ma)

8.0 7.6 7.8 7.8 7.8 8.0 7.8 7.8 7.8 7.6 7.9 7.8 7.7 7.5 7.3 11 7.9 7.6 8.8 7.6 7.2 9.1 7.1 7.0 8.1 7.3 7.1 7.2 8.4 7.1 7.6 7.3 4.1 4.6 4.6 3.4 5.4 5.9 4.9 4.9 5.2 5.5 4.9 4.8 5.0 5.0 7.4 5.5 3.7 5.1 4.9 4.6 4.8 4.6 4.3 4.9 4.6 3.8 4.1 5.3 4.7 4.7 17 20 26 19 19 18 19 21 19 17 18 18

387 312 328 405 389 404 362 432 398 347 342 443 346 320 443 411 403 381 373 402 363 388 368 310 340 411 389 390 317 358 358 410 386 409 316 355 430 372 335 376 386 399 398 345 415 312 362 366 383 412 372 361 344 403 381 326 343 385 386 507 392 385 40.9 32.9 25.3 30 33.1 28.4 24.3 32.2 35.1 32 33.1 29

±σ (Ma) 46 50 53 50 42 39 41 42 51 48 49 45 37 40 45 58 49 51 57 42 36 43 40 39 39 36 38 51 49 47 40 46 46 87 48 90 61 74 51 44 54 51 56 73 51 50 94 15 12 71 24 21 31 32 20 34 34 14 13 81 33 33 4.8 4.4 4.6 4.4 4.1 4.0 3.9 4.0 4.3 3.6 3.8 4.2

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Table 2 (continued) Sample spot # Drango@13 Drango@14 Drango@15 Drango@16 Drango@17 Drango@18 QK-1@1 QK-1@2 QK-1@3 QK-1@4 QK-1@5 QK-1@6 QK-1@7 QK-1@8 QK-1@9 QK-1@10 QK-1@11 QK-1@12 QK-1@13 QK-1@14 QK-1@15 QK-1@16 QK-1@17 QK-1@18 QK-1@19 QK-1@20 QK-1@21 QK-1@22 QK-1@23 QK-1@24 QK-1@25 QK-1@26 QK-1@27 QK-1@28 QK-1@29 QK-1@30 QK-1@31 QK-1@32 QK-1@33 QK-1@34 QK-1@35 QK-1@36 QK-2@1 QK-2@2 QK-2@3 QK-2@4 QK-2@5 QK-2@6 QK-2@7 QK-2@8 QK-2@9 QK-2@10 QK-2@11 QK-2@12 QK-2@13 QK-2@14 QK-2@15 QK-2@16 QK-2@17 QK-2@18 QK-2@19 QK-2@20 QK-2@21 QK-2@22 QK-2@23 QK-2@24 QK-2@25 QK-2@26 QK-2@27 QK-2@28 QK-2@29 QK-2@30 QK-2@31 QK-2@32 QK-2@33

[U]a ppm

[Th] ppm

Th/U

8.9 8.7 8.8 8.5 8.7 8.4 2.2 2.1 2.2 2.2 2.8 2.7 2.1 2.1 1.8 1.9 1.9 2.3 3.0 2.2 1.7 2.0 2.0 2.3 1.9 2.4 2.4 2.4 2.5 2.2 2.3 2.4 2.4 2.2 2.2 2.1 2.0 2.9 2.6 3.5 2.6 2.1 2.4 2.4 2.3 2.2 2.6 2.5 2.1 2.2 2.3 2.4 1.5 2.7 2.3 2.0 1.9 1.9 2.0 2.6 2.1 2.6 2.0 2.5 2.0 2.5 2.6 2.2 2.3 2.2 2.4 2.4 2.3 2.3 1.7

183 179 177 177 172 169 22 22 23 23 29 28 23 22 17 19 20 23 32 22 18 21 21 24 21 24 25 25 25 24 24 25 25 22 22 21 21 30 22 32 27 19 25 24 24 23 22 27 22 23 25 24 10 23 23 22 22 20 22 27 23 24 19 20 20 20 21 23 24 24 25 25 24 24 15

20.6 20.5 20.1 20.8 19.9 20.1 10 10.4 10.3 10.6 10.6 10.5 10.9 10.7 9.4 10.1 10.5 10 10.4 10 10.5 10.7 10.1 10.4 11.1 10 10.3 10.2 10.1 10.5 10.6 10.3 10.5 10 10.1 10 10.5 10.5 8.4 9.4 10 9.1 10.4 10.2 10.6 10.5 8.6 10.7 10.1 10.4 10.6 10.2 6.8 8.6 10.2 10.8 11.3 10.7 10.9 10.3 11.1 9.2 9.3 7.9 9.9 7.9 8.1 10.6 10.2 10.6 10.6 10.6 10.6 10.4 8.6

Pbc ppm

ƒ206b %

ƒ208b %

0.05 0.04 0.04 0.06 0.08 0.08 2.5 2.4 3.0 3.3 2.1 3.0 3.0 2.8 2.1 2.6 2.9 2.4 2.6 3.1 2.9 2.8 3.4 3.1 2.5 3.5 3.1 3.1 3.0 3.6 3.6 3.1 3.1 2.9 2.7 2.2 3.1 2.6 2.1 2.4 2.5 2.4 2.9 3.0 2.8 2.9 2.7 2.7 3.2 3.3 3.5 3.4 2.1 2.8 3.1 2.9 2.5 3.0 3.1 2.5 2.9 3.0 2.5 2.7 3.1 2.7 2.8 3.1 2.9 3.0 3.0 2.7 2.5 2.5 2.5

22 21 24 32 41 38 71 73 78 78 65 75 78 77 74 74 78 69 68 81 79 80 78 78 79 80 77 80 76 81 78 75 78 78 73 69 83 68 64 63 69 73 76 75 72 78 68 70 79 81 79 81 79 72 77 76 76 81 81 70 81 72 77 73 83 71 71 79 75 79 73 74 74 71 80

10 8 7 8 15 18 63 61 67 68 52 61 67 62 66 63 67 60 54 65 68 67 67 66 68 67 67 65 64 69 70 63 67 64 62 61 71 55 59 54 57 62 65 63 60 66 63 62 69 65 65 70 76 62 65 68 63 65 71 61 63 62 67 67 70 65 65 66 62 67 65 61 62 59 72

238 206

U

Pb

133 153 185 157 163 135 2.20 2.21 1.94 1.80 2.90 2.31 1.86 2.02 2.13 1.90 1.80 2.35 2.80 2.00 1.63 1.94 1.60 1.95 2.11 1.89 2.07 2.16 2.15 1.75 1.74 2.07 2.13 2.01 2.08 2.33 1.78 2.61 2.75 3.10 2.53 2.22 2.20 2.08 1.99 1.98 2.26 2.27 1.81 1.86 1.83 1.93 1.97 2.40 1.96 1.81 2.06 1.78 1.81 2.53 2.02 2.12 2.17 2.38 1.92 2.29 2.34 1.97 2.09 2.00 2.03 2.27 2.38 2.28 1.93

±σ %

207

Pb

206

Pb

10 11 12 15 16 14 3.4 3.4 3.4 3.5 3.4 3.4 3.5 3.5 3.4 3.5 3.3 3.4 3.4 3.5 4.4 3.4 3.4 3.4 3.5 3.4 3.4 3.4 3.7 3.3 3.7 3.4 3.5 3.3 3.7 3.4 3.5 3.4 3.3 3.3 3.3 3.4 3.6 3.4 3.4 3.6 3.5 3.4 3.5 3.4 3.5 3.4 3.4 3.4 3.4 3.6 3.3 3.6 3.5 3.4 3.4 3.4 3.5 3.4 5.2 3.4 3.3 4.4 3.3 3.5 3.5 3.5 3.7 3.4 3.5

0.23 0.22 0.24 0.3 0.38 0.350 0.662 0.678 0.716 0.715 0.604 0.690 0.713 0.706 0.684 0.680 0.713 0.643 0.636 0.741 0.725 0.736 0.718 0.715 0.721 0.730 0.706 0.731 0.696 0.743 0.716 0.694 0.713 0.714 0.678 0.643 0.754 0.635 0.603 0.589 0.639 0.671 0.700 0.693 0.667 0.716 0.634 0.650 0.723 0.744 0.721 0.740 0.723 0.667 0.708 0.699 0.697 0.744 0.740 0.649 0.741 0.664 0.707 0.678 0.757 0.654 0.657 0.728 0.692 0.724 0.676 0.686 0.686 0.661 0.733

±σ%

207

Pb-corrcAge (Ma)

21 19 21 29 22 17 3.4 3.4 3.1 3.1 3.4 3.0 3.1 3.3 3.7 4.2 4.1 3.5 3.8 3.2 3.6 3.4 2.9 2.8 4.4 2.7 2.8 2.8 2.9 2.7 2.7 2.9 4.5 3.0 3.8 3.3 3.3 3.7 4.3 3.2 3.1 3.9 3.3 2.9 4.6 3.1 2.9 3.9 3.7 2.9 2.9 4.0 3.8 3.2 3.2 4.1 5.0 3.8 3.3 3.5 4.3 4.1 3.5 3.4 3.1 3.3 3.3 3.5 5.0 3.2 3.4 4.0 3.9 3.6 5.2

37.2 32.9 26.5 27.8 23 29.5 871 823 799 867 800 753 841 803 832 939 870 874 754 693 907 732 956 799 722 773 784 673 785 782 889 819 739 777 869 877 725 808 846 786 823 839 759 821 944 784 934 884 829 734 830 720 764 788 819 917 817 764 763 797 685 896 746 763 663 862 835 747 822 752 897 775 740 846 745

±σ (Ma) 4.7 4.2 3.8 6.1 5.5 5.4 85 86 96 103 61 77 100 95 95 116 124 79 70 98 129 105 111 90 115 92 84 83 82 100 99 84 114 90 99 77 114 74 75 54 68 93 87 84 116 94 73 89 117 99 98 119 110 75 96 122 125 124 110 74 120 100 93 80 105 79 78 106 123 96 93 95 90 85 145

(continued on next page)

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Table 2 (continued) Sample spot # QK-2@34 QK-2@35 QK-2@36 QK-2@37

[U]a ppm

[Th] ppm

Th/U

2.4 2.3 2.4 1.1

26 20 22 9

10.7 8.7 9.1 8.1

Pbc ppm

ƒ206b %

ƒ208b %

3.4 2.3 2.3 2.0

80 73 69 86

67 62 61 77

238 206

U

Pb

±σ %

1.96 2.60 2.49 1.66

3.4 3.5 3.4 4.0

207

Pb

206

Pb

0.737 0.673 0.644 0.782

±σ% 3.3 3.7 3.6 5.2

207

Pb-corrcAge (Ma)

722 715 822 664

±σ (Ma) 102 78 77 180

Error assigned to the ratio is one sigma based on counting statistics and calibration. a U concentrations are determined by corresponding variations of UO2+/40Ca231P16O3+ ratios of Durango standard apatite with average 9 ppm U. b ƒ206 values are calculated by 207Pb-correction method (Williams, 1998; Chew et al., 2011); ƒ208 values are estimated by ƒ206, (208Pb/206Pb)measured and (208Pb/206Pb)common. c Common Pb compositions are determined by upper intercept of the Tera–Wasserburg plot, except Durango, by Stacey and Kramers (1975) model.

207 Pb-corrected ages give a weighted average 206Pb/ 238U age of 31 ± 2 Ma (Fig. 6 inset). Our U–Pb dating results of the Durango apatite are in good agreement within errors with the 40Ar– 39Ar age of 31.44 ± 0.18 Ma (McDowell et al., 2005) and the 238U– 206Pb age of 30.6 ± 2.3 Ma and. 232Th– 208Pb age of 32.5 ± 1.2 Ma (Chew et al., 2011).

5.4. QGBLK apatite There were 36 measurements made on QK-1 apatite and 37 measurements on QK-2 apatite. As shown in Table 2 and Fig. 7, these two samples are quite homogeneous in their U (2.3 ± 0.3 ppm, 1 SD) and Th (22.7 ± 3.7 ppm, 1 SD) contents, Th/U ratio (10.0 ± 0.9, 1 SD)

and common Pb contents (2.8 ± 0.4 ppm, 1 SD). Among all the apatite samples in this study, the QGBLK apatites contain the highest common Pb concentration and proportion, with ƒ206 = 63–86% and ƒ208 = 52– 77%. The combined 73 measurements define a regression line on Tera–Wasserburg plot, with an upper intercept 207Pb/ 206Pb of 0.94 ± 0.04 and a lower intercept age of 804 ± 92 Ma (2σ) (Fig. 7). Using the upper intercept value (0.94 ± 0.04) as the common Pb composition, the 207Pb-corrected U–Pb ages (Eqs. (3) and (4)) yield a weighted average of 805 ± 21 Ma, which is consistent within errors with the TIMS U–Pb age of 810 ± 6 Ma (Zhang et al., 2007) and the SIMS Pb–Pb age of 814.7 ± 3.6 Ma for baddeleyite from the same intrusion (Li et al., 2010a). 6. Discussion The SIMS U–Pb apatite age results in this study are summarized in Table 3. It clearly shows that our SIMS U–Pb dating results are in good agreement within error with other independent age constraints from the same sample. Our work demonstrates that the in situ SIMS method is able to determine the U–Pb ages of apatites with low U and high proportions of common Pb using appropriate analytical protocols, apatite standards, calibration of Pb/U fractionation and common Pb correction. 6.1. Calibration protocols For U–Th–Pb dating by SIMS, the primary task, but one of most contentious issues, is to deduce the actual Pb/U and Pb/Th from large inter-elemental fractionations (Ireland and Williams, 2003). The basic assumption used in the ion microprobe procedure is that the measured Pb* +/U + (* denotes radiogenic) ratios scale with their

Fig. 5. U–Pb dating results for the Kovdor apatite. (A) KV-8 and KV-18 apatites; (B) KVA apatite. Data-point error bars and weighted age uncertainties are at the 2σ level.

Fig. 6. U–Pb dating results for the Durango apatite. Data-point error bars and weighted age uncertainties are at the 2σ level.

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low in Th/U and high in ƒ208 values (~15%, Table 1), resulting in a large standard deviation (N6%) of the calibration line (Fig. 1B). Thus, this apatite is inadequate as a Th–Pb dating standard. Recently, Chew et al. (2011) suggested Kovdor carbonatite apatite with high U–Th–Pb concentrations could act as suitable U–Th–Pb dating standard. However, our SIMS measurements on three apatite samples from the Kovdor carbonatite yield significantly lower U–Th–Pb concentrations than those reported by Chew et al. (2011). Further work is necessary to develop the Kovdor apatite (or other apatites) as well-characterized U–Th–Pb standards before Th–Pb age dating could be regularly employed, but the potential of in-situ Th–Pb dating of apatite is clearly there. 6.2. Common Pb corrections

Fig. 7. U–Pb dating results for the Qieganbulake apatite. Red and squared box, QK-1; Blue and circle, QK-2. Data-point error bars and weighted age uncertainties are at the 2σ level.

actual Pb*/U ratios in both the unknown and standard at the same UO2(or UO)/U (Eq. (1), see also Ireland et al., 1990). It is obvious that the Pb/U calibration could be complicated if the standard and unknown do not have same range of UO2(UO)/U, which was the situation for the SIMS U–Pb analyses on rutile (Li et al., 2011). For apatite, the Pb/U fractionation can be easily calibrated due to relatively narrow range of measured UO2(UO)/U values in all apatite samples (inset of Fig. 1, see also Sano et al., 1999; Nemchin et al., 2009). The NW-1 apatite is rather homogeneous in terms of Pb–Pb age, and has high U and low proportions of common Pb. With the NW1 apatite as a standard, reasonably good dating results have been obtained (Table 3), indicating that NW-1 could potentially be a good candidate to serve as the apatite standard for in-situ U–Pb dating. It awaits further precise age determination by ID-TIMS. It is noteworthy that apatite often contains high Th/U ratios (Table 2), making Th–Pb dating feasible because the common Pb contribution to the 208Pb total is less pronounced (Amelin and Zaitsev, 2002; Chew et al., 2011). The SIMS Th–Pb dating method has been successfully established on accessory minerals with high Th/U, such as allanite (Gregory et al., 2007) and perovskite (Li et al., 2010b). As shown in Fig. 1B, a linear relationship between ln( 208Pb* +/Th +) and ln(ThO +/Th +) analogous to that for Pb/U (Fig. 1B) was also observed in NW-1 apatite, indicating the possibility of Th–Pb dating by similar calibration protocols to the U–Pb dating. However, the NW-1 apatite is

For U–Pb dating on apatite with low U contents and consequentially high proportions of common Pb, the main difficulty is how to effectively correct for common Pb. The same challenge exists not only for lower precision in-situ SIMS and LA–ICPMS analyses, but also for high-precision, low-blank ID-TIMS analysis. The problem of estimating the initial Pb isotopic composition is circumvented if multiple analyses of a suite of cogenetic apatite grains show a significant spread in their common Pb/radiogenic Pb ratios that define a well constrained linear array of discordia or isochrons (Figs. 4–5, see also Cox and Wilton, 2006; Simonetti et al., 2006; Batumike et al., 2008). However, when such conditions are not satisfied, the common Pb correction must be undertaken based on an appropriate choice of common Pb isotopic composition. As shown in Table 2 and Fig. 3, the studied apatites contain low common Pb concentrations (b3 ppm), which results in large uncertainties in 204Pb measurements. Therefore, common Pb-correction methods using 204Pb would not be effective nor appropriate because most measurements give very low 206Pb/ 204Pb ratios (b50) with large uncertainties (N10%), particularly for young samples such as Durango apatite. Furthermore, the generally high Th/U ratios in apatite make the 208Pb-correction method inappropriate, though this method works well for some low Th/U apatites (Chew et al., 2011). Therefore, our preferred approach is to plot total 206Pb/ 238U and 207Pb/ 206Pb data on a Tera–Wasserburg inverse Concordia diagram and using the 207 Pb-correction method. As shown in Fig. 8, it is ideal when the sample has a significant spread in common Pb (ƒ206) values (Fig. 8A), which can thus produce a lower intercept age of crystallization and an upper intercept yielding the common Pb composition, as in the KV-A apatite (Fig. 5B). For samples such as KV-18 and QK apatites with very high proportions of common Pb (Fig. 3) and a limited spread of data points in Tera–Wasserburg space (Fig. 5A, Fig. 7), is difficult to deduce a precise lower intercept age. Fortunately, these uncorrected data could be regressed to project a relatively precise upper intercept on a Tera–Wasserburg Concordia, providing the common Pb component

Table 3 Summary of the ages for the apatites dated in this study. Sample

KV-8 KV-18 KV-A Durango QK NW-1

U–Th–Pbc concentrations (ppm) U

Th

Pbc

1.9 ± 0.18 2.4 ± 0.5 7.4 ± 5.6 9±1 2.3 ± 0.3 73 ± 40

45 ± 4 99 ± 36 878 ± 534 180 ± 15 22.7 ± 3.7 38 ± 6

0.4 ± 0.05 1.2 ± 0.2 1 ± 0.1 0.05 ± 0.013 2.8 ± 0.4 0.75 ± 0.2

Th/U

24.3 ± 1.5 40 ± 9 139 ± 45 19.8 ± 0.8 10 ± 0.9 0.68 ± 0.38

ƒ206 (%)

ƒ208 (%)

n

56 ± 5 72 ± 5 47 ± 17 25 ± 7 75 ± 5 1.8 ± 0.8

26 ± 3 32 ± 6 5±3 9±3 65 ± 5 17 ± 5

32 15 15 18 73 33

Age (Ma) TW Concordia

TW anchored

207

Pb-corr

Common 207 Pb/206Pb

333 ± 120 375 ± 18 375 ± 15 0.84 ± 0.04 358 ± 100 368 ± 34 372 ± 29 0.84 ± 0.04 378 ± 16 377 ± 11 377 ± 11 0.84 ± 0.04 – 32 ± 2 31 ± 2 0.84 ± 0.05 804 ± 92 805 ± 43 805 ± 21 0.94 ± 0.04 The SK model corrected Pb–Pb age = 1160 ± 36 Ma

Accepted age (Ma) 378 Ma (1) 378 Ma (1) 378 Ma (1) 31 Ma (2) 810 Ma (3,4) 1163 Ma (5,6)

Uncertainties for U–Th–Pb concentrations, Th/U ratio, ƒ206 and ƒ208 are listed at the 1σ level. The age results and common Pb composition are quoted at the 2σ level. TW Concordia denotes a Tera–Wasserburg lower intercept age without initial Pb constrains. TW anchored denotes a Tera–Wessurberg lower intercept age with assigned common Pb composition. 207 Pb-corr denotes the weighted average of 207Pb-corrected 207Pb*/206Pb* ages using the assigned common Pb composition. The common Pb compositions for Kovdor and QGBLK samples were determined by projecting a regression line through the uncorrected data on a Tera–Wasserburg plot. The SK model was used to determine the common Pb composition for Durango and NW-1 apatites. The accepted age is quoted from (1) Amelin and Zaitsev (2002); (2) McDowell et al. (2005); (3) Zhang et al. (2007); (4) Li et al. (2010a); (5) Rukhlov and Bell (2010); (6) Wu et al. (2010).

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Fig. 8. Diagrams showing the evaluation of effect on the age result by variation in common Pb proportion and isotopic composition. (A) Significant spread in ƒ206 values defines a well constrained linear array on a Concordia diagram; (B–E) show clustered data with ƒ206 = 80%, 50%, 20%, 5%, resulted in U–Pb ages with uncertainties of 2 RSD = 32%, 8.5%, 2.2%, 0.4% with the fixed common 207Pb/206Pb = 0.86 ± 0.06, respectively; (F) Influence of variation in common Pb composition in calculated ages as a function of common Pb proportion. Supposed sample age is 378 Ma (using Kovdor apatite as an example) and common Pb composition varies ranging from 0.8 to 0.92, as the Stacey and Kramers (1975) model derived 207 Pb/206Pb = 0.86 ± 0.06.

on the 207Pb/ 206Pb axis (Fig. 5A, and Fig. 7, see also Cox and Wilton, 2006; Simonetti et al., 2006; Batumike et al., 2008). Consequently, the 207 Pb-correction method could be undertaken to yield precise and accurate weighted average 206Pb/ 238U ages. For samples such as KV-8 or Durango apatites that show data points on the Tera–Wasserburg Concordia too clustered to construct a regression line, either the 207Pb-corrected method or the Tera– Wasserburg Concordia intercept age anchored through common Pb are necessary, both of which need an appropriate choice of common Pb composition. The approaches to determine the initial common Pb compositions involve either analyzing a co-existing phase with low U and high common Pb concentrations (e.g. Chamberlain and Bowring, 2000; Amelin and Zaitsev, 2002) or using the terrestrial lead composition at the time of crystallization (e.g. the widely used twostage evolution model of Stacey and Kramers, 1975, SK model hereafter). For the first approach, it is not always feasible to find a suitable mineral for this purpose. For example, co-existing U- and Th-

poor minerals in sedimentary and complex metamorphic rocks may not always be in isotopic equilibrium with apatite when they form. Therefore, the terrestrial lead composition (Stacey and Kramers, 1975) has been commonly employed and proven to be suitable in most cases (e.g. Heaman, 2009; Li et al., 2010b; Chew et al., 2011), though sometimes the calculated Pb isotopic composition is significantly different from the direct measurements of co-existing low U and high common Pb phases (e.g. Corfu and Dahlgren, 2008; Nemchin et al., 2009). In such cases, it is necessary to critically evaluate the impact on the U–Pb age results of a reasonable spread of common Pb compositions. 6.3. Influence of variation in common

207

Pb/ 206Pb

The 207Pb-correction method is equivalent to extrapolating each analysis along the common Pb mixing line on the Tera–Wasserburg Concordia diagram until 207Pb/ 206Pb reaches the nominal radiogenic

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value, then determining the corresponding radiogenic 206Pb/ 238U (Williams, 1998). In other words, it is equivalent to calculating a Tera– Wasserburg U–Pb age with an assigned common Pb composition anchored on the y-axis. Fig. 8 (B–E) show a series of assessments on a supposed 378 Ma-old sample (with Kovdor apatites as a prime example) with variable common Pb proportions when applying initial common Pb compositions ranging from 0.8 to 0.92. This range is assigned as the SK model derived 207Pb/ 206Pb = 0.86 ± 0.06, which is considered to cover the possible range of initial common Pb compositions at 378 Ma. Uncertainties in the 206Pb/ 238U and 207Pb/ 206 Pb measurements were not considered in order to check only the effect of variations in the common Pb composition. As shown in Fig. 8B, when a sample has 80% common Pb, the 207Pb-corrected U–Pb age would be between 258 and 480 Ma, with a 2 RSD = 36%, reflective of significant effect of the variation in common Pb composition. However, this effect decreases significantly with decreasing common Pb proportions, with ƒ206 = 50%, 20%, 5% corresponding to relative age uncertainties of 9.0% (Fig. 8C), 2.2% (Fig. 8D), 0.5% (Fig. 8E), respectively. If an age uncertainty of b3% (2 RSD) is acceptable for a particular geological problem, we could confidently use the SK model for samples with ƒ206 b 25% (corresponding to 207Pb/ 206Pb b 0.25 for phanerozoic samples). Taking Durango apatite with an average ƒ206 = 26% as an example, though the common Pb composition ( 207Pb/ 206Pb) could vary from 0.78 to 0.9 (0.84 ± 0.06), the average U–Pb age would range from 29.7 ± 2 Ma to 31.4 ± 2 Ma. However, if we take QK apatites with ƒ206 N 75% in this test, using the SK model of 207 Pb/ 206Pb = 0.89 at 810 Ma resulted in a U–Pb age of 688 ± 16 Ma. This age is clearly biased compared with the co-existing baddeleyite Pb–Pb age of ~810 Ma (Zhang et al., 2007; Li et al., 2010a), indicating the SK model is not appropriate here. In fact, there may be different views about a fixed value for the common Pb composition. Nevertheless, we should accept that there must be a discrepancy between the true common Pb composition and the value we used, either from the co-existing minerals or from the SK model. Taking the assumed sample with a true age of 378 Ma and corresponding common Pb composition of 207Pb/ 206Pb = 0.86 in a test, Fig. 8F outlines the influence of the supposed discrepancy between the used and “true” common Pb composition as a function of the proportion of common Pb (ƒ206). For example, a 0.02 difference in initial common 207Pb/ 206Pb ratio would result in a ~ 22% deviation from the true U–Pb age for ƒ206 = 90%, but only 5.8% for ƒ206 = 70%, 2.5% for ƒ206 = 50%, and b1% for ƒ206 b 30%. From another perspective, when an age with b5% (2 RSD) deviation from the true values could be accepted, one needn't worry much about common Pb composition when ƒ206 b 40%, but must have enough confidence in the common Pb composition used when ƒ206 N 40%, such as in the case of b0.04 discrepancy at ƒ206 ~50%, b0.02 at ƒ206 ~ 70%, only b0.005 at ƒ206 ~90% (Fig. 8F). Therefore, careful determination of the common Pb composition is the key to obtain accurate U–Pb ages when ƒ206 is high, especially for values of ƒ206 N 40%. 7. Summary We present analytical procedures and data reduction protocols for SIMS in-situ U–Th–Pb analyses on apatites with low U and high proportion of common lead. The NW-1 apatite from the ~1160 Ma carbonatite in the Prairie Lake complex in Canada has been used as a U–Pb standard. Our results demonstrate that U–Pb apatite ages can be precisely and accurately determined for apatite with low U (b3 ppm) and high proportion of common Pb (N50%) by SIMS with careful choice of the common Pb composition. Acknowledgments We thank Yuri Amelin and Bill Griffin for providing the Kovdor apatites, Ken Farley for providing the Durango apatite standard. The

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