There is abundant evidence for significant H2O in evolved melts from the platinum-rich UG2 chromitite and the Merensky Reef of the Bushveld Complex (South Africa), but there is no consensus about the source of H2O. We report triple-oxygen and hydrogen isotope ratios of interstitial, late-magmatic phlogopite from three localities of the UG2 layer. The phlogopite yielded δD values of −43 to −23, which is >30 higher than previously known from Bushveld rocks and far above the mantle values of ~−75. The phlogopite triple-oxygen isotope ratios are the first to be reported for Bushveld rocks, with values of Δ′17O0.5305 (17O excess relative to the reference line 0.5305) from –0.069 to –0.044 (δ18O 5.2–6.2). The oxygen data support existing models of as much as 30%–40% contamination of mantle-derived magmas in the lower to middle crust. However, the high δD values require a second step of contamination, which we attribute to brines from the marine sediments in the Transvaal Basin at the emplacement level.Understanding the petrogenesis of mafic layered intrusions and their mineralization remains a huge challenge, and among many open questions is the role of H2O in these essentially anhydrous igneous bodies (Charlier et al., 2015). The Rustenburg Layered Suite (RLS) of the Bushveld Complex (South Africa) contains the world's three largest deposits of platinum-group elements: the Upper Group 2 (UG2) chromitite and the Merensky Reef in the upper Critical Zone (CZ) in the eastern and western limbs of the complex, and the Platreef in the northern limb (Fig. 1). Crystallized melt inclusions containing phlogopite, hornblende, and Cl-rich apatite have been reported in UG2 chromite and Merensky Reef olivine, confirming a role of hydrated melts in the formation of these layers (Li et al., 2005; Schannor et al., 2018; Smith et al., 2021). These late-stage melts crystallized interstitial phlogopite, which is closely intergrown with chromite and sulfide minerals (Smith et al., 2021). It has been suggested that interaction of hydrated, Cl-rich melts with near-solidus cumulates of the UG2 and Merensky Reef caused remelting of the latter and remobilization of fluid-mobile elements (e.g., Mathez and Mey, 2005; Boudreau, 2008). Veksler and Hou (2020) confirmed experimentally that adding 4 wt% H2O to a B1 composition (parental melt composition for the CZ; Barnes et al., 2010) shifts the liquidus assemblage from silicate minerals to Cr-spinel. They inferred that localized H2O additions from the floor of the Bushveld magma chamber could have remelted pyroxenite cumulates and precipitated chromite, thus forming chromitite layers like UG2.While the evidence for late hydrous melts in the CZ is well documented, the source of H2O is not. One option is concentration of primary, mantle-derived H2O in residual melts by fractional crystallization, while the other is derivation from external, crustal sources. The latter might be suggested by long-established evidence for crustal contamination of RLS magmas from high Sr- and Os-isotope ratios (Hart and Kinloch, 1989; McCandless and Ruiz, 1991; Kruger, 1994) and from δ18O values higher than the mantle composition (Schiffries and Rye, 1989; Harris et al., 2005). Benson et al. (2020) reviewed evidence for large-scale fluid generation in the thermal aureole beneath the Bushveld intrusion and used geochemical and thermal-mechanical models to show how such fluids might have affected the RLS magmas. Surprisingly few studies have addressed the source of H2O by coupled hydrogen and oxygen isotope analysis of OH-bearing minerals, possibly because conventional methods required prohibitively large amounts of pure mineral separates given the low abundance of phlogopite. Two exceptions are studies by Mathez et al. (1994) and Willmore et al. (2002), who reported δD values of biotite from the Merensky Reef and other pyroxenite units from the CZ. There are a number of published bulk-rock values of δD, but these are harder to interpret because the mineral hosts of H2O are unknown. We report the first co-registered H and triple-O (16O, 17O, and 18O) isotopic ratios in late-magmatic phlogopite from the UG2 horizon at three localities spaced throughout the complete RLS (Fig. 1). Our analytical techniques require a sample mass of just a few milligrams, low enough to obtain high-purity phlogopite separates from the UG2 chromitite and also from silicate wall rocks.Samples were collected from industry drill cores through UG2 in the Nkwe, Khuseleka (chromitite and wall rocks), and Karee (chromitite only) mines. Details of UG2 at these mines are given by Veksler et al. (2018) and Junge et al. (2014). The spatial distribution and abundance of phlogopite in the samples were determined by micro–X-ray fluorescence element mapping on the same polished thin sections used for electron microprobe analyses. Phlogopite was separated from ~10 g of crushed, sieved rock material, and the stable isotopic ratios of hydrogen and oxygen were measured by thermal conversion elemental analyzer (TC/EA) and laser fluorination, respectively. Details of methods and analytical conditions are given in the Supplemental Material1.Phlogopite forms fresh, reddish-brown, interstitial grains (0.1–1 modal%, ~50–500 μm in size) in the UG2 chromitite and its silicate wall rocks. It is evenly disseminated in the samples, with no clustering or connectivity between grains (Fig. S1 in the Supplemental Material). Phlogopite replaces orthopyroxene in UG2 chromitite and is intergrown with chromite, plagioclase, sulfides, and platinum-group minerals (Figs. 2A and 2B). Locally, phlogopite is associated with myrmekitic K-feldspar–quartz intergrowths (Figs. 2C and 2D). The phlogopite within chromitite has Mg# [100 × Mg/(Mg + Fe2+)] of 92–97, molar K/(K + Na) of 0.81–0.97, and TiO2, F, and Cl contents of 3.0–5.1, 0.1–0.6, and 0.1–0.4 wt%, respectively (Table S1, Figs. S2 and S3). Phlogopite in silicate wall rocks has a lower Mg# (71–88), higher K/(K + Na) ratio (0.97–1), and nearly the same range of TiO2, F, and Cl contents.The total range of δD for phlogopite within chromitite is −38 to −23 (Table S2; Fig. 3), and the average values from each locality are nearly identical (Nkwe: −30 ± 4, n = 6; Khuseleka: −33 ± 1, n = 6; Karee: −29 ± 3, n = 8). Phlogopite from silicate rocks is similar in terms of range (−43 to −26) and average values (Nkwe: −33 ± 7, n = 4; Khuseleka: −32 ± 5, n = 3). All separates yielded stoichiometric H2O contents (3–4 wt%). This, combined with the lack of correlation between δD and H2O (Fig. S4) and the stoichiometric contents of 9 wt% K2O + Na2O rules out significant chloritization. We estimated the magmatic δD range as −37 to −16 using the compositionally dependent ΔDphlogopite-water (δDphlogopite – δDwater) equation at 800 °C (Table S2; Suzuoki and Epstein, 1976), based on the chromite-plagioclase O-isotope closure temperature from the Nkwe UG2 (Schannor et al., 2018), 750 °C for F-Cl-OH exchange between phlogopite and apatite (Willmore et al., 2000), and 850 °C for Ti thermometry of zircon associated with phlogopite from the UG2 footwall (Zeh et al., 2015).The δ18O values of chromitite-hosted phlogopite are practically the same in all localities (5.7 ± 0.3, n = 19; Table S2; Fig. 3), and this is also true for phlogopite in silicate rocks (5.7 ± 0.2, n = 4). The corresponding magmatic values are 6.6 ± 0.3 based on Δ18Ophlogopite-basalt of −0.9 at 800 °C (Zhao and Zheng, 2003). Not all samples yielded enough phlogopite for combined D/H and triple-O measurements, so the latter are fewer in number (Fig. 3). The Δ′17O values, expressing 17O excess relative to the reference line 0.5305, are calculated as: Δ′17O = δ′17O – 0.5305 × δ′18O, where δ′17,18O = 1000 × ln(δ17,18O/1000 + 1). They range from −0.069 to −0.044 and are similar in all localities (Nkwe: −0.053 ± 0.009, n = 6; Khuseleka: −0.051 ± 0.007, n = 2; Karee: −0.054 ± 0.003, n = 2). These Δ′17O values can be considered equivalent to magmatic values given the negligible difference in Δ′17O between magma and minerals at high temperatures (Bindeman, 2021). There are no correlations of either δ18O or δD with the chemical composition of phlogopite (Figs. S5 and S6).The triple-O isotopic ratios reported here are the first published for the Bushveld Complex. Our average δ18O values for UG2 phlogopite (5.7 ± 0.3) are similar to biotite data from CZ pyroxenites (5.6–5.7; Willmore et al., 2002), and both data sets suggest a magma composition of 6.6 ± 0.3, which is consistent with the RLS range (6.8–7.1; Schiffries and Rye, 1989; Harris et al., 2005). In contrast, the δD values of phlogopite from UG2 at all three localities studied (−43 to −23) are distinctly higher than previously reported for the RLS (Fig. 4A), although there are no published data from the UG2 layer. The range of δD values for phlogopite or biotite from the Merensky Reef, other CZ rocks, and the Platreef (−88 to −49; Mathez et al., 1994; Harris and Chaumba, 2001; Willmore et al., 2002; Pronost et al., 2008) is similar to the bulk-rock range for the entire RLS (−99 to −53; Mathez et al., 1994; Harris et al., 2005). The authors cited above suggested a magmatic source of H2O because the δD values are similar to those of mantle rocks (−75 ± 12; Loewen et al., 2019). A late-magmatic origin of phlogopite in UG2 is indicated by textures (e.g., interstitial grains with no connection outside the layer; intergrowths with chromite, sulfide minerals, and feldspar–quartz microgranophyre) and chemical composition (e.g., high TiO2, stoichiometric K2O and H2O contents). Therefore, we consider the phlogopite δD values to reflect those of the magma. Our estimate of δD for the UG2 magma based on phlogopite (−37 to −16) is at least 30 higher than the mantle value.In principle, high magmatic δD can be source related or related to processes of magma evolution including mixing and assimilation. Source-related D/H enrichment in the mantle has been attributed to subduction-related metasomatism (Loewen et al., 2019), and indeed, a link between Bushveld parental magmas and subduction-related boninites has been suggested (e.g., Willmore et al., 2002). However, if the high δD values in UG2 were source related, they should be found throughout the RLS, which is not the case. Processes that enrich D/H during magma evolution include degassing, crystallization of abundant hydrous minerals, and contamination with external, high-δD material. Degassing cannot be ruled out for the shallow-level Bushveld intrusion, but there is no reason why this process should affect UG2 more than other layers of the RLS, so it is discounted. Enrichment of D/H in residual magma due to crystallization of hydrous minerals is unlikely because all but the latest, interstitial phases are anhydrous. The same argument rules out post-cumulus equilibration of interstitial melts with the surrounding, anhydrous cumulus minerals. Therefore, contamination of the UG2 magma seems the only likely explanation.Harris et al. (2005) compiled new and existing isotopic data showing a narrow range of magmatic δ18O values throughout the RLS (7.1 ± 0.5) that are higher than the mantle composition (5.7 ± 0.2; Ito et al., 1987; Bindeman, 2008) and indicate significant crustal contamination. They argued that the uniformity in δ18O requires that contamination took place before emplacement and at a deep level, proposing 30%–40% contamination by Archean basement (with δ18O = 9.2–9.6) in the lower to middle crust. In terms of δ18O, the UG2 magma appears to be no different than that of other layers of the RLS, and can be explained by the contamination model of Harris et al. (2005). The new Δ′17O values from this study (−0.069 to −0.044) are also consistent with this model (Fig. 4B), although we note that the triple-O isotope composition of potential contaminants is poorly constrained (see the Supplemental Material for more discussion). However, the distinctly high δD values of UG2 imply that the magma was affected by a separate (and presumably later) process that was capable of increasing δD without causing a significant shift in δ18O. Mass-balance considerations suggest an aqueous fluid with high δD as the most likely contaminant, and the fact that RLS magmas intruded unmetamorphosed marine sedimentary rocks of the Transvaal Supergroup makes it likely that this contamination took place at the emplacement level. Benson et al. (2020) modeled the consequences of contact metamorphism and dehydration of the Transvaal sediments and showed how fluids could have been introduced into the overlying RLS by a combination of footwall diapirism and localized overpressure-driven channels.We suggest that seawater-derived basinal brines in the Transvaal Supergroup sediments would have appropriately high D/H ratios to raise the magmatic δD substantially without affecting oxygen isotope ratios. Furthermore, saline brines provide an alternative to subduction fluids for explaining the high Cl contents found in interstitial phlogopite and apatite from the CZ (Willmore et al., 2000, 2002). Support for this suggestion comes from a study by Gleason et al. (2011), who proposed that Mississippi Valley–type mineralization around the Bushveld Complex was related to basinal brine circulation caused by the intrusion. Estimates of the δD values of Proterozoic seawater are 0 ± 20 (Pope et al., 2012; Zakharov and Bindeman, 2019). Schiffries and Rye (1990) reported −34 δD for fluid inclusions from the Bushveld aureole, which might represent such brines. Figure 4B and Figure S7 explore the two-step contamination scenario we propose (see the Supplemental Material). The first-step scenario follows Harris et al. (2005) as discussed above. For the second step, we assumed a range of −30 to 0 for δD in Transvaal Basin brines based on the arguments above. We used 0.3 wt% H2O for the initial magma based on chemical analyses of B1-type sills (Barnes et al., 2010) and −75 for δD from average RLS values. The curves shown for δDbrine = −30 to 0 match the UG2 δD at 50%–90% mixing (brine to total dissolved H2O), with negligible shifts of δ18O (−0.1) or Δ′17O (+0.001). Melt inclusions in Merensky Reef chromite contain 2.6 wt% H2O (Li et al., 2005). If we assume a similar concentration for H2O in the late-stage UG2 melts and 0.3 wt% for initial B1 magma, then ~90% of the H2O in these hydrous melts is externally derived, which is in agreement with the mixing model.In summary, a two-step contamination process is proposed to explain the distinctive O-H isotope composition of the UG2 magma (Fig. 4B). The first step took place well below the emplacement level and involved the bulk of RLS magmas, raising their δ18O values to ~7 without much effect on δD (median RLS values: −80 to −70; Fig. 4A). The second step took place at the emplacement level, where thermal metamorphism and loading at the floor of the intrusion caused large-scale release of basinal brines from sediments of the Transvaal Supergroup. Injection of overpressured brine into the magma, possibly as proposed by Benson et al. (2020), led to an increase in δD. While this model can explain the O-H isotopic composition of late-magmatic phlogopite from UG2, there are still open questions. From the scant H-isotope data available, it appears that the Merensky Reef, although it also hosts phlogopite and hydrated melt inclusions, does not share the high-δD signature of UG2. Without more δD data from other layers of the Bushveld Complex, it is premature to speculate if and why external brines might have played a particular role in the development of UG2. We conclude by suggesting that phlogopite δD has proven value in distinguishing fluid sources in the Bushveld magmas and should be more widely applied.H. Zhou thanks the German Academic Exchange Service (DAAD) for supporting his stay in Potsdam and the German Research Centre for Geoscience (GFZ) for offsetting analytical costs. I. Veksler was supported by German Research Foundation (DFG) grant VE 619/6-1, and I. Bindeman by U.S. National Science Foundation grant EAR-1833420. We are grateful to Lutz Hecht (Museum für Naturkunde Berlin) for discussions on the origin of Bushveld chromitites and his help securing samples from Nkwe. Chris Harris and Alan Boudreau are thanked for constructive reviews.

中文翻译：

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盆地盐水对布什维尔德复合体（南非）岩浆的污染：来自 UG2 铬铁矿的金云母中的稳定同位素

有大量证据表明，富含铂的 UG2 铬铁矿和布什维尔德杂岩体（南非）的梅伦斯基礁演化出的熔体中存在大量 H2O，但关于 H2O 的来源尚未达成共识。我们报告了来自 UG2 层三个位置的间隙、晚期岩浆金云母的三重氧和氢同位素比率。金云母产生的 δD 值为 -43 至 -23，比先前从布什维尔德岩石中已知的值高 >30，并且远高于~-75 的地幔值。Bushveld 岩石的金云母三氧同位素比值是首次报道的，Δ'17O0.5305（相对于参考线 0.5305 的 17O 过量）从 –0.069 到 –0.044（δ18O 5.2–6.2）。氧气数据支持现有模型，即中下地壳中地幔源岩浆的污染高达 30%–40%。然而，高 δD 值需要第二步污染，我们将其归因于就位水平的德兰士瓦盆地海洋沉积物中的盐水。了解基性层状侵入体的岩石成因及其矿化仍然是一个巨大的挑战，其中许多悬而未决的问题是H2O 在这些基本上无水的火成岩体中的作用（Charlier 等，2015）。Bushveld Complex（南非）的 Rustenburg Layered Suite (RLS) 包含世界上三个最大的铂族元素矿床：上族 2 (UG2) 铬铁矿和位于东部上临界区 (CZ) 的梅伦斯基礁和复合体的西翼，以及北翼的 Platreef（图 1）。含有金云母、角闪石、已经在 UG2 铬铁矿和梅伦斯基礁橄榄石中报道了富含氯的磷灰石和富氯磷灰石，证实了水合熔体在这些层的形成中的作用（Li 等人，2005 年；Schannor 等人，2018 年；Smith 等人，2021 年） . 这些晚期熔体结晶出间隙金云母，与铬铁矿和硫化物矿物紧密共生（Smith 等，2021）。有人提出，水合的、富含氯的熔体与 UG2 和梅伦斯基礁的近固相线累积物的相互作用导致后者的重熔和流体流动元素的再迁移（例如，Mathez 和 Mey，2005 年；Boudreau，2008 年）。Veksler 和 Hou (2020) 通过实验证实，将 4 wt% 的 H2O 添加到 B1 组合物（CZ 的母体熔体组合物；Barnes 等人，2010 年）会将液相线组合从硅酸盐矿物转移到 Cr 尖晶石。他们推断，来自 Bushveld 岩浆房底部的局部 H2O 添加可能使辉石岩重熔并沉淀铬铁矿，从而形成 UG2 等铬铁矿层。 . 一种选择是通过分步结晶在残余熔体中浓缩原生地幔衍生的 H2O，而另一种选择来自外部地壳来源。后者可能是由长期建立的 RLS 岩浆因高 Sr 和 Os 同位素比（Hart 和 Kinloch，1989 年；McCandless 和 Ruiz，1991 年；Kruger，1994 年）和高于地幔的 δ18O 值而受到地壳污染的证据所暗示的组成（Schiffries 和 Rye，1989 年；Harris 等人，2005 年）。本森等人。(2020) 审查了 Bushveld 侵入体下方热光环中大规模流体生成的证据，并使用地球化学和热力学模型来显示此类流体可能如何影响 RLS 岩浆。令人惊讶的是，很少有研究通过对含 OH 矿物的氢氧同位素耦合分析来解决 H2O 的来源，这可能是因为鉴于金云母的丰度低，传统方法需要大量的纯矿物分离物。Mathez 等人的研究有两个例外。(1994) 和 Willmore 等人。(2002)，他报告了来自梅伦斯基礁的黑云母和来自 CZ 的其他辉石岩单元的 δD 值。有许多已发表的大块岩石 δD 值，但由于 H2O 的矿物主体未知，因此这些值难以解释。我们报告了第一个共同注册的 H 和三 O (16O, 17O 和 18O) 同位素比率来自 UG2 层位的晚岩浆金云母在整个 RLS 中间隔的三个位置（图 1）。我们的分析技术只需要几毫克的样品质量，低到足以从 UG2 铬铁矿和硅酸盐围岩中分离出高纯度金云母。围岩）和 Karee（仅铬铁矿）矿山。Veksler 等人提供了这些矿山 UG2 的详细信息。(2018) 和 Junge 等人。(2014)。样品中金云母的空间分布和丰度是通过在用于电子探针分析的相同抛光薄片上进行微 X 射线荧光元素映射来确定的。从约 10 克粉碎、筛分的岩石材料中分离出金云母，氢和氧的稳定同位素比分别通过热转化元素分析仪（TC/EA）和激光氟化法测定。补充材料 1 中给出了方法和分析条件的详细信息。金云母在 UG2 铬铁矿及其硅酸盐围岩中形成新鲜的红褐色间隙颗粒（0.1-1 modal%，尺寸约 50-500 μm）。它均匀地分布在样品中，晶粒之间没有聚集或连接（补充材料中的图 S1）。金云母取代了 UG2 铬铁矿中的斜方辉石，并与铬铁矿、斜长石、硫化物和铂族矿物共生（图 2A 和 2B）。在当地，金云母与 myrmekitic 钾长石-石英共生有关（图 2C 和 2D）。铬铁矿中的金云母的 Mg# [100 × Mg/(Mg + Fe2+)] 为 92–97，摩尔 K/(K + Na) 为 0.81–0.97，TiO2、F 和 Cl 含量分别为 3.0–5.1、0.1–0.6 和 0.1–0.4 wt%（表 S1，图 S2 和 S3）。硅酸盐围岩中的金云母具有较低的 Mg# (71–88)、较高的 K/(K + Na) 比 (0.97–1) 和几乎相同的 TiO2、F 和 Cl 含量范围。 δD 的总范围铬铁矿中的金云母为 -38 至 -23（表 S2；图 3），每个位置的平均值几乎相同（Nkwe：-30 ± 4，n = 6；Khuseleka：-33 ± 1，n = 6；卡瑞：-29 ± 3，n = 8)。硅酸盐岩中的金云母在范围（-43 至 -26）和平均值（Nkwe：-33 ± 7，n = 4；Khuseleka：-32 ± 5，n = 3）方面相似。所有分离都产生了化学计量的 H2O 含量（3-4 重量％）。这与 δD 和 H2O 之间缺乏相关性相结合（图 3）。S4) 和 9 wt% K2O + Na2O 的化学计量含量排除了显着的氯化。基于铬铁矿-斜长石 O-同位素闭合，我们在 800 °C 下使用成分依赖的 ΔDphlogopite-水（δDphlogopite – δDwater）方程估计岩浆 δD 范围为 -37 至 -16（表 S2；Suzuoki 和 Epstein，1976）来自 Nkwe UG2 的温度（Schannor 等人，2018 年），750 °C 用于金云母和磷灰石之间的 F-Cl-OH 交换（Willmore 等人，2000），以及 850 °C 用于与金云母相关的锆石的 Ti 温度测定UG2 下盘（Zeh 等人，2015 年）。铬铁矿承载的金云母的 δ18O 值在所有地区几乎相同（5.7 ± 0.3，n = 19；表 S2；图 3），这对于硅酸盐岩中的金云母 (5.7 ± 0.2, n = 4)。对应的岩浆值为6.6±0。3 基于 800 °C 下 -0.9 的 Δ18 金云母玄武岩 (Zhao 和 Zheng, 2003)。并非所有样品都能产生足够的金云母用于组合 D/H 和三重 O 测量，因此后者的数量较少（图 3）。Δ'17O 值表示相对于参考线 0.5305 的 17O 过量，计算如下： Δ'17O = δ'17O – 0.5305 × δ'18O，其中 δ'17,18O = 1000 × ln(δ17,18O/1000 + 1）。它们的范围从 -0.069 到 -0.044，并且在所有地区都相似（Nkwe：-0.053 ± 0.009，n = 6；Khuseleka：-0.051 ± 0.007，n = 2；Karee：-0.054 ± 0.003，n = 2）。考虑到高温下岩浆和矿物之间 Δ'17O 的差异可以忽略不计，这些 Δ'17O 值可以被认为与岩浆值相同（Bindeman，2021）。δ18O 或 δD 与金云母的化学成分没有相关性（图 S5 和 S6）。此处报告的三氧同位素比率是首次针对 Bushveld 复合体发表的。我们 UG2 金云母的平均 δ18O 值 (5.7 ± 0.3) 与来自 CZ 辉石岩的黑云母数据 (5.6–5.7; Willmore et al., 2002) 相似，两个数据集都表明岩浆成分为 6.6 ± 0.3，这与RLS 范围（6.8-7.1；Schiffries 和 Rye，1989；Harris 等，2005）。相比之下，来自 UG2 的所有三个研究地点（-43 至 -23）的金云母的 δD 值明显高于先前报告的 RLS（图 4A），尽管没有来自 UG2 层的已发表数据。来自 Merensky Reef、其他 CZ 岩石和 Platreef 的金云母或黑云母的 δD 值范围（-88 到 -49；Mathez 等人，1994 年；Harris 和 Chaumba，2001 年；Willmore 等人，2002 年；Pronost 等人等，2008）类似于整个 RLS 的大块岩石范围（-99 到 -53；Mathez 等，1994；Harris 等，2005）。上面引用的作者提出了 H2O 的岩浆来源，因为 δD 值与地幔岩石的 δD 值相似（-75 ± 12；Loewen 等，2019）。UG2 中金云母的晚期岩浆成因由质地（例如，在层外没有连接的间隙颗粒；与铬铁矿、硫化物矿物和长石-石英微晶石共生）和化学成分（例如，高 TiO2、化学计量 K2O 和H2O 含量）。因此，我们考虑金云母 δD 值来反映岩浆的值。我们基于金云母（-37 至-16）对 UG2 岩浆的 δD 估计至少比地幔值高 30。原则上，高岩浆 δD 可能与源有关或与包括混合和同化在内的岩浆演化过程有关。地幔中与源相关的 D/H 富集归因于与俯冲相关的交代作用（Loewen 等，2019），事实上，已经提出了 Bushveld 母岩浆与俯冲相关的 boninites 之间的联系（例如，Willmore 等., 2002)。然而，如果 UG2 中的高 δD 值与源有关，则应该在整个 RLS 中找到它们，但事实并非如此。在岩浆演化过程中富集 D/H 的过程包括脱气、大量含水矿物的结晶以及外部高 δD 物质的污染。不能排除浅层Bushveld侵入的脱气，但没有理由认为该过程比RLS的其他层对UG2的影响更大，因此打折扣。由于含水矿物的结晶，残余岩浆中的 D/H 富集不太可能，因为除了最新的间隙相之外，所有的相都是无水的。同样的论点排除了间隙熔体与周围无水积云矿物的积云后平衡。因此，UG2 岩浆的污染似乎是唯一可能的解释。哈里斯等人。(2005) 汇编了新的和现有的同位素数据，显示整个 RLS (7.1 ± 0.5) 的岩浆 δ18O 值范围较窄，高于地幔成分 (5.7 ± 0.2；Ito 等人，1987；宾德曼，2008)，并表明地壳污染严重。他们认为 δ18O 的均匀性要求污染发生在就位之前和深层，提出 30%–40% 的污染来自下地壳中的太古代基底（δ18O = 9.2–9.6）。就δ18O而言，UG2岩浆似乎与RLS的其他层没有区别，可以用Harris等人的污染模型来解释。(2005)。本研究的新 Δ'17O 值（-0.069 至 -0.044）也与该模型一致（图 4B），尽管我们注意到潜在污染物的三氧同位素组成受到的限制很差（参见补充材料）更多讨论）。然而，UG2 明显高的 δD 值意味着岩浆受到了一个单独的（可能是后来的）过程的影响，该过程能够增加 δD 而不会引起 δ18O 的显着变化。质量平衡考虑表明具有高 δD 的水性流体是最可能的污染物，RLS 岩浆侵入德兰士瓦超群未变质的海相沉积岩这一事实使得这种污染很可能发生在就位水平。本森等人。(2020) 模拟了德兰士瓦沉积物接触变质和脱水的后果，并展示了如何通过下盘底辟作用和局部超压驱动通道的组合将流体引入上覆 RLS。德兰士瓦超群沉积物将具有适当高的 D/H 比，以在不影响氧同位素比的情况下显着提高岩浆 δD。此外，盐水为俯冲流体提供了一种替代方法，用于解释在 CZ 的间隙金云母和磷灰石中发现的高 Cl 含量（Willmore 等人，2000 年，2002 年）。对这一建议的支持来自 Gleason 等人的一项研究。(2011)，他提出布什维尔德杂岩体周围的密西西比河谷型矿化与侵入引起的盆地盐水循环有关。元古代海水的 δD 值估计为 0 ± 20（Pope 等，2012；Zakharov 和 Bindeman，2019）。Schiffries 和 Rye (1990) 报告了来自 Bushveld 光环的流体包裹体的 -34 δD，这可能代表此类卤水。图 4B 和图 S7 探讨了我们提出的两步污染方案（参见补充材料）。第一步场景遵循 Harris 等人。(2005) 如上所述。对于第二步，我们根据上述参数假设德兰士瓦盆地盐水中 δD 的范围为 -30 到 0。我们使用了 0。基于 B1 型基岩的化学分析（Barnes et al., 2010），初始岩浆的 H2O 为 3 wt%，平均 RLS 值的 δD 为 -75。δDbrine = -30 到 0 的曲线与 UG2 δD 在 50%–90% 混合（盐水与总溶解的 H2O）相匹配，δ18O (-0.1) 或 Δ'17O (+0.001) 的变化可以忽略不计。Merensky Reef 铬铁矿中的熔体包裹体含有 2.6 wt% 的 H2O（Li 等，2005）。如果我们假设后期 UG2 熔体中 H2O 的浓度相似，而初始 B1 岩浆中的 H2O 浓度为 0.3%，那么这些含水熔体中约 90% 的 H2O 是外部导出的，这与混合模型一致。 ，提出了一个两步污染过程来解释 UG2 岩浆的独特 OH 同位素组成（图 4B）。第一步发生在侵位层以下，涉及大部分 RLS 岩浆，将它们的 δ18O 值提高到 ~7 对 δD 没有太大影响（中值 RLS 值：-80 到 -70；图 4A）。第二步发生在就位水平，在侵入体底部的热变质作用和载荷导致从德兰士瓦超群沉积物中大规模释放盆地卤水。将超压盐水注入岩浆中，可能是 Benson 等人提出的。(2020)，导致 δD 增加。虽然该模型可以解释 UG2 晚期岩浆金云母的 OH 同位素组成，但仍有一些悬而未决的问题。从可用的少量 H 同位素数据来看，Merensky 礁虽然也含有金云母和水合熔体包裹体，但似乎并不具有 UG2 的高 δD 特征。如果没有来自 Bushveld Complex 其他层的更多 δD 数据，现在推测外部盐水是否以及为什么可能在 UG2 的开发中发挥了特殊作用还为时过早。我们得出的结论是，金云母 δD 在区分 Bushveld 岩浆中的流体来源方面已被证明具有价值，应该得到更广泛的应用。周感谢德国学术交流中心 (DAAD) 支持他在波茨坦逗留，并感谢德国地球科学研究中心 (GFZ) 抵消分析成本。I. Veksler 得到德国研究基金会 (DFG) 授予 VE 619/6-1 的支持，I. Bindeman 得到美国国家科学基金会授予 EAR-1833420 的支持。我们感谢 Lutz Hecht (Museum für Naturkunde Berlin) 讨论 Bushveld 铬铁矿的起源以及他帮助从 Nkwe 获取样品。感谢 Chris Harris 和 Alan Boudreau 的建设性评论。