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Article

Geochemistry of Magmatic and Xenocrystic Spinel in the No.30 Kimberlite Pipe (Liaoning Province, North China Craton): Constraints on Diamond Potential

by
Ren-Zhi Zhu
1,2,
Pei Ni
1,*,
Jun-Ying Ding
1 and
Guo-Guang Wang
1
1
State Key Laboratory for Mineral Deposit Research, Institute of Geo-Fluids, School of Earth Science and Engineering, Nanjing University, Nanjing 210023, China
2
Department of Earth and Environmental Sciences, University of Milano-Bicocca, Milan 20126, Italy
*
Author to whom correspondence should be addressed.
Minerals 2019, 9(6), 382; https://doi.org/10.3390/min9060382
Submission received: 29 April 2019 / Revised: 22 June 2019 / Accepted: 23 June 2019 / Published: 24 June 2019
(This article belongs to the Special Issue Minerals of Kimberlites: An Insight into Petrogenesis and the Diamond Potential of Deep Mantle Magmas)
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Abstract

:
There are two genetic types of spinel (magmatic spinel crystallizing directly from kimberlite magma and xenocrystic spinel derived from mantle xenoliths) in the No.30 kimberlite pipe (Liaoning Province, North China Craton). Their geochemistry is investigated to reveal processes of diamond capture and resorption during kimberlite magmatism to constrain the diamond potential. Magmatic spinels are mostly euhedral to subhedral, 20 to 60 µm in size, and have compositional zones: the cores are classified as chromite with high Cr and Mg contents, and the rims are classified as magnetite with low Cr and high ferric Fe. The compositional trends suggest that magmatic spinel and olivine phenocrysts are crystallized contemporaneously during the early stages of kimberlite crystallization. During this period, temperature (T) and oxygen fugacity (fO2) values calculated at an assumed pressure of 1 GPa are in the range of 994–1274 °C and 1.6–2.4 log fO2 units below the nickel-nickel oxide (NNO) buffer, respectively. The high values of fO2 suggest heavy diamond resorption during kimberlite magmatism. Estimated temperatures of xenocrystic spinel range from 828 to 1394 °C, and their distributions indicate that only a small proportion of xenocrystic spinels are derived from the diamond stabilization field, which suggests a low potential of diamond capture. The low diamond capture and heavy diamond resorption during kimberlite magmatism contributed to the low diamond grade of the No.30 kimberlite.

1. Introduction

Diamond exploration relies on mantle-derived “indicator minerals” because they commonly occur as inclusions within diamonds, and thus, play an important role in diamond formation. The widely utilized indicator minerals include garnet, chromite, clinopyroxene, and ilmenite [1,2,3,4,5,6,7,8,9,10,11,12,13]. The use of chromite in diamond exploration has concentrated on the recognition of grains similar in the major-element composition to chromite inclusions within diamonds [14]. Combined major and trace element compositions for chromite are used to distinguish the source rock of chromite and to discriminate mantle-derived xenocrysts from magmatic spinels [4]. In addition, the Zn-in-chromite thermometer (proposed by Ryan et al. [15] based on the strong temperature-dependence of the partitioning of Zn between chromite and olivine) further divides chromite xenocrysts from kimberlite into those derived from the diamond stability field and those from shallower, barren levels of mantle [4,5]. Single-grain temperature estimates are projected onto known geotherm to obtain a depth of origin for each grain [16].
At present, the discrimination of mantle-derived materials (xenoliths and xenocrysts) entrained by kimberlite is the main tool to predict the grade and quality of diamonds transported by the kimberlite to the Earth’s surface. However, in recent years, more and more studies focusing on various morphological forms and surface features in natural diamond recovered from kimberlites, show that diamond resorption (dissolution) widely occur during kimberlite magma ascent, which greatly influences the grade and value of diamonds [17,18]. The diamond resorption is greatly controlled by the oxygen fugacity (fO2) of the kimberlite magma [17,19,20,21]. A good understanding of the magmatic history of the kimberlite is necessary to ascertain whether diamonds could survive the rapid journey from the upper mantle to the surface of the Earth. It has been shown that spinel commonly forms throughout the kimberlite crystallization [22], and thus the morphology and composition of magmatic spinel may be useful in deciphering the kimberlite crystallization history [23,24], and the temperature and oxygen fugacity in kimberlite magma during magmatic spinel formation could be estimated [25].
In a previous study on xenocrystic spinels from the No.30 kimberlite (Liaoning Province, North China Craton), the authors suggested that the cores of the spinel (classified as magnesiochromite) were remnants of primary mantle xenocrysts, and the intermediate zones (classified as chromite) and the rims (classified as magnetite) formed during kimberlite magmatism [26]. In this study, we report the trace element concentrations of xenocrystic spinels and discuss the differences between magnesiochromites derived from the diamond stability field and those from the barren levels of the mantle [4]. The major element compositions of magmatic spinels are presented and compared with those of the rims in xenocrystic magnesiochromites. The temperature and oxygen fugacity of the No.30 kimberlite magma are estimated in order to discuss the diamond resorption. The diamond potential of the No.30 kimberlite is discussed on the basis of the trace elements signature and temperature estimates of the xenocrystic garnet xenocrysts [27] and spinel (this paper).

2. Samples and Methods

The samples in this study consist of 14 magmatic spinels and 51 xenocrystic spinels from the No.30 kimberlite pipe. The magmatic spinels were prepared as double-polished thin sections and their major compositions were analyzed in this study. Xenocrystic spinels were single crystals obtained by heavy mineral concentration and separation, and their morphology and major elements have been reported in a previous study by Zhu et al. [26]. In this study, we proceeded to analyze their trace element concentrations.
The major elements of magmatic spinel were analyzed at the State Key Laboratory for Mineral Deposit Research, Nanjing University, China, using a JAX 8100 Electron Probe Micro Analyzer (EPMA). The EPMA work was conducted by wavelength dispersive spectrometry (WDS). The analytical conditions were as follows: beam accelerating voltage 20 kV, beam current 20 nA, beam diameter 1 μm, and a counting time of 30 s. A number of natural, as well as synthetic standards, were used for calibration. For the operating conditions, the detection limit for major elements was 0.01 wt. %, and the accuracy was better than 1%. The proportion of Fe2+ and Fe3+ in the chrome spinel was calculated using stoichiometric and charge balance equations [28].
Trace elements in xenocrystic spinel were investigated at the State Key Lab of Ore Deposit Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences, using a Coherent GeoLasPro 193-nm Laser Ablation system coupled with an Agilent 7700× ICP-MS. For spot analysis, a 40 μm spot size was applied with an energy density of ~100 mJ per pulse and a repetition rate of 6 Hz. Each analysis contained 20 s for measuring the gas blank and 40 s for data acquisition. Different silicate glass reference materials, including BHVO-2G, BIR-1G, GSE-1G, and NIST610 were used as external standards and 57Fe as an internal standard to calibrate element contents. Quality control of time-dependent drift of sensitivity and mass discrimination was monitored by GSE-1G repeated after every eight analyses. The off-line data processing was performed using the ICPMSDataCal program [29]. Values were cross-checked against standards BCR-2G. Analytical precision of ≤6% was indicated for most elements (Table S1 in Supplementary Materials).

3. Results

3.1. Morphology of Spinel

The No.30 kimberlite was heavily hydrothermally altered. However, the xenocrystic and magmatic spinels were very resistant to alteration and survived as primary minerals. The morphology of xenocrystic spinel has been described by Zhu et al. [26] and will not be repeated here. The magmatic spinels are pervasive in kimberlite, usually euhedral to subhedral, 20 to 60 µm in diameter, and strongly zoned from the chromite core to the magnetite rim (Figure 1A, Table 1). The chromite cores in some magmatic spinels are further divided into Ti-Al-Mg chromite core and Ti-Al chromite intermediate zone (Figure 1B, Table 1). Some spinels show small holes (Figure 1C) that may represent an incipient “lagoon” in an atoll-texture [23], which is the characteristic feature of magmatic spinel in Group I kimberlite [22,30].

3.2. Chemical Composition of Spinel

Magmatic spinels and xenocrystic spinels show compositional zones (Figure 1 and Figure 2). From the cores through the intermediate zones to the rims in the magmatic spinels, there is a continuous decrease in Cr2O3 (from 49.8–53.4 through 16.4–46.0 to 0.2–2.7 wt. %) and MgO contents (from 8.1–13.9 through 0.3–4.6 to 0.1–0.7 wt. %); the Al2O3 concentrations increase from the cores (4.6–6.5 wt. %) to the intermediate zones (4.5–8.5 wt. %) and decrease sharply in the rims (0.01–1.5 wt. %); the TiO2 concentrations increase from the cores (3.5–4.6 wt. %) to the intermediate zones (5.2–10.3 wt. %) and decrease sharply in the rims (0.2–4.6 wt. %) (Figure 2, Table 1); and there is a continuous increase in Fe2O3 contents (from 7.2–9.7 through 10.2–13.1 to 57.8–65.9 wt. %). The cores of the xenocrystic spinels are discriminated by higher Cr2O3 and MgO and lower TiO2 and Fe2O3 compared with magmatic spinels (Figure 2), which supports the xenocrystic origin [26]. The intermediate zones and rims of xenocrystic spinels are suggested to form in the host kimberlite magma and can be linked to the magmatic spinels [26]. As shown in Figure 2, the rims of xenocrystic spinels show low Cr2O3, MgO, Al2O3, and high Fe2O3 contents, which can be linked to the rims in the magmatic spinels. The intermediate zones of xenocrystic spinels show a wider range in Al2O3 and Cr2O3 concentrations, which shows more affinity to the cores of the xenocrystic spinels rather than the cores of the magmatic spinels; however, their MgO, TiO2, and Fe2O3 concentrations are consistent with the latter. Moreover, in the spinel prisms ([23] and references therein), the cores of the magmatic spinel and intermediate zones of xenocrystic spinels are classified as chromite, the rims of the magmatic spinels and xenocrystic spinels are classified as magnetite, and the cores of the xenocrystic spinels are classified as xenocrysts (Figure 3).
The trace element concentrations in the xenocrystic spinels from the No.30 kimberlite are shown in Figure 4 and Table 2, with Zn (320–1047 ppm), Ga (6.4–102 ppm), Zr (0.1–11.3 ppm), Nb (0.1–7.1 ppm), and Ni (668–2306 ppm). Most element concentrations are too low to meet the detection limit, and thus, are not shown here.

4. Discussion

4.1. Early Crystallization Sequence

An understanding of the texture and chemical composition of kimberlite minerals is difficult because of the lack of detailed experimental work on the crystallization history of kimberlite magma [22,31]. Commonly, the first minerals to crystallize from the kimberlite magma are olivine phenocrysts and chromite [22]. Within the No.30 kimberlite pipe, chromite in the magmatic spinel was found in the cores of zoned magmatic spinels. The zoning trend in magmatic spinels from the chromite core to the magnetite rim is typical for the kimberlite and magnetite rim and is considered to be a product of late-stage crystallization [22,23]. In addition, an atoll-texture shown in zoned spinel (Figure 1B) is typical in Group I kimberlite [22].
The spinel prism-diagram showing the variation in the composition of spinel is widely used to infer crystallization processes in kimberlite [22,23,32,33]. It uses either ferric iron (oxidized prism) or titanium (reduced prism) as the vertical axis. There are two typical compositional trends in kimberlite: trend 1 from chromite to magnesio-ulvospinel-magnetite and trend 2 of increasing titanium and ferric iron as a function of increasing Fe2+/(Fe2+ + Mg) (Figure 3, [23]). Trend 1 is influenced by the high carbonate content of kimberlite melt which causes rapid crystallization of the minerals [23], and trend 2 is generally known to form by the co-crystallization of phlogopite that depletes melt in Mg and Al [22]. As shown in Figure 3, with the increasing of Fe2+/(Fe2+ + Mg), Cr/(Cr + Al) decreases and Fe3+/(Fe3+ + Al + Cr) increases from the cores to the rims in the magmatic spinels. This compositional trend in Figure 3 is similar to trend 2 but slightly deviates to the right of the outline suggested by Mitchell [22]. This indicates that besides the co-crystallization of phlogopite and spinel, another Mg-rich mineral, such as olivine, must be involved. This is consistent with the proposal by Roeder and Schulze [23] and Bussweiler et al. [34] where trend 2 was shown to be caused by co-crystallization of olivine and phlogopite in kimberlites. In the Al-Cr-Fe3+ diagram (Figure 3C), the intermediate zones in the magmatic spinel show the compositional variation to trend 1 but appears different in the other plots (Figure 3A,B). Here, we explained that because the crystallization of large phlogopite prevented new phlogopite nucleation in the vicinity, local Al contents in the kimberlite magma increase during the crystallization of minerals (Table 1), and thus data points in Figure 3C are shifted from Trend 2 to the Trend 1 [23]. In summary, we conclude that spinel, olivine, and phlogopite are crystallized contemporaneously at the beginning of kimberlite crystallization.

4.2. Temperature and Oxygen Fugacity

Co-crystallized spinel-olivine pairs can be used to estimate the temperature and oxygen fugacity (fO2) of the magma at the time of crystallization. The oxygen fugacity of kimberlite is an important parameter that influences diamond resorption [17,35,36]. Magmatic chromite compositions in this study suggest crystallization together with olivine phenocrysts (see Section 4.1). However, virtually all the olivine phenocrysts in the No.30 kimberlite were altered, and thus, their composition could not be obtained. An alternative method is to use the composition of olivine inclusion within chromite [37] which shows an average Mg# (=Mg/(Mg + Fe) × 100) of 90. In this study, equilibrium temperatures are estimated using the spinel–olivine Fe/Mg exchange thermometer of O’Neill and Wall [38], as simplified by Ballhaus et al. [25]. As shown in Table 3, the calculated temperatures range from 994 to 1274 °C assuming a pressure of 1 GPa. Pressure has a fairly small effect on temperature (1.5% decrease in temperature from 10 kbar to 1 bar). The oxygen barometer is based on the olivine-orthopyroxene-spinel equilibrium developed by Ballhaus et al. [25]. This barometer requires the presence of orthopyroxene, and for orthopyroxene-undersaturated rocks, the barometer gives maximum oxygen fugacity values [25]. Because the silica activity (aSiO2) of kimberlites is well below that required to stabilize orthopyroxene [39,40,41,42,43], a correction has to be made to the calculated fO2 values. Silica activity in the kimberlite can be limited by the groundmass mineral assemblages at a given temperature (T) and pressure (P) [22,30]. In this study, the upper limit of silica activity is constrained by the presence of monticellite, which is below the diopside-monticellite (Di-Mont) buffer. Silica activity for the Di-Mont buffer was calculated using the thermodynamic data of Holland and Powell [44], and the method of fO2 adjustment in this study was described in detail by Fedortchouk and Canil [35]. As shown in Table 3, the adjusted fO2 of the No.30 kimberlite ranges from1.6 to 2.4 log units below the nickel-nickel oxide (NNO) buffer.

4.3. Trace Element Concentration in the Diamond Window

In diamond exploration, it is important to discriminate the indicator mineral that is derived from diamond stability field and that from the diamond-barren or no-diamond field. The composition characteristics of minerals in the diamond stability field are based on mineral inclusions in the diamond [14]. However, the relative scarcity of mineral inclusions in the diamond limits a comprehensive understanding of these characteristics. Griffin and Ryan [4] describe the diamond stability field as the “diamond window”, which can be defined as the range of TNi (Ni-in-garnet thermometer) between the intersection of geotherm with the diamond-graphite equilibrium curve and the base of the lithosphere (unmodified lithosphere). According to the estimate by Griffin and Ryan’s for the Liaoning Province, there is a diamond window with temperatures between 900–1250 °C [4]. In this study, we use the major and trace element composition in xenocrystic spinels to redefine a slightly narrower temperature range (950–1200 °C) for the diamond window. In contrast to the approach by Griffin and Ryan, we use a multi-element approach to constrain the temperature range.
As shown in Figure 4, the distribution of major and trace element concentrations show an abrupt change when TZn (Zn-in-chromite temperature, [15]) moves through 950 and 1200 °C. The compositions in the xenocrystic spinel with temperature 950–1200 °C are nearly constant: Cr2O3 (60–66 wt. %), Al2O3 (0–8 wt. %), TiO2 (0–1 wt. %), Ni (500–1000 ppm), Zr (0–3 ppm), and Nb (0–2 ppm), and are consistent with that of chromite inclusions in diamonds [14]. The Ni, Zr, and Nb concentrations in spinels with temperatures between 950–1200 °C are lower than those with temperatures higher than 1200 °C (Figure 4), which may indicate that higher-T spinels have interacted with asthenosphere-derived melts [4,45,46,47]. For the spinels with temperatures below 950 °C, the spinels show the hybrid features: inner feature of spinels in the graphite field (low Cr2O3 and high Al2O3 contents) and low-T modification (increasing Zr and Nb concentrations). Due to the paucity of studies on chromite metasomatism, it is difficult to distinguish between metasomatism and original trace element signatures. Because some low-T metasomatism, especially carbonatite or COH metasomatism can be related to diamond formation [11,12,13,48,49,50], the lower T boundary of the diamond window is unclear, and further studies are needed to constrain this more precisely.

4.4. Diamond Potential in Kimberlite

Diamond formation in the mantle is generally considered to be a metasomatic process [51,52]. The metasomatic agents react with the mantle rocks in which they infiltrate, and the diamond crystallizes as a consequence of different redox reactions. The oxidation state is likely controlled by Fe0-Fe2+-Fe3+ components in oxide minerals like chromite [12]. Malkovets et al. [16] suggested that garnet and chromite play an important role in the diamond formation. Therefore, garnet and chromite are important indicators of minerals. In general, diamondiferous kimberlites sample both garnet and chromite, but low-grade pipes tend to sample only chromite [16]. As discussed in Section 4.3, the diamond window beneath the Liaoning Province is redefined. If kimberlite captures garnet and chromite from this field at the same time, it has great potential to capture diamonds without considering other factors.
The equilibrium temperatures in the garnet xenocrysts from the No.30 kimberlite pipe calculated using single grain Ni-in-garnet thermometer [53] ranges from 1100 to 1350 °C, with a peak at about 1150 °C (Figure 5; [27]). However, the calculated equilibrium temperatures of xenocrystic chromites from the No.30 kimberlite using single grain Zn-in-chromite [15] vary from 828 to 1394 °C, with a peak at about 950 °C (Figure 5, Table 2). The temperature distribution curves show that garnets and chromites are not distributed in the same way, which is similar to the low-grade Zarnitsa and Solokha kimberlites, and different from the high-grade Udachnaya and Lomonosova kimberlites [16]. It indicates that garnets have not overgrown with chromites and diamonds haven’t widely formed in mantle rock samples by the kimberlite (Figure 5) [16]. Thus even though the No.30 kimberlite sampled the garnet/chromite-bearing mantle rock from the diamond stability field, the diamond grade in the kimberlite is low because the sampled mantle rocks are diamond-barren [16] and the diamond window is narrow.
As previously suggested by Zhu et al. [26], the oxidizing environment also contributes to low-grade diamonds in the No.30 kimberlite. The estimated fO2 of the No.30 kimberlite magma is 1.6–2.4 log units relative to NNO buffer during the early magmatism (Table 3). Ideally, the fO2 in the No.30 kimberlite is compared with that of other kimberlites with different diamond grades in the North China Craton. However, there is no such study on fO2 or related studies. So in this study, we compare fO2 of the No.30 kimberlite magma with fO2 estimates for the Lac de Gras kimberlites (Slave craton, Canada) with different diamond grades [17]. As shown in Figure 6, the comparison shows that the fO2 of the No.30 kimberlite magma is most similar to the fO2 of the low-grade Ranch Lake kimberlite and significantly higher than the fO2 of high-grade Misery and Beartooth kimberlites in the Lac de Gras area. As suggested by many researchers, high fO2 promotes diamond resorption [17,20,21]. We conclude that diamond resorption in the No.30 kimberlite was strong, which contributed to the generally low diamond grade.

5. Conclusions

Magmatic spinels in the No.30 kimberlite pipe are euhedral to subhedral, 20–60 µm in diameter. These spinels are zoned from Ti-Al-Mg chromite through Ti-Al chromite to magnetite, suggesting co-crystallization of spinel and olivine phenocrysts during early kimberlite crystallization.
At an assumed pressure of 1 GPa, temperatures and oxygen fugacity of kimberlite magma are estimated in a range of 994–1274 °C and 1.6–2.4 log units below the NNO buffer, respectively, suggesting strong diamond resorption during kimberlite magmatism.
Temperature distribution curves of xenocrystic garnet and chromite show that the sampling of garnet does not overlap with that of chromite, which suggests the mantle rock sampled by the kimberlite is diamond-barren.
Low-diamond capture and heavy-diamond resorption during kimberlite magmatism lead to low-diamond grade in the No.30 kimberlite.

Supplementary Materials

The following are available online at https://www.mdpi.com/2075-163X/9/6/382/s1, Table S1: Trace elements chemistry for standard.

Author Contributions

R.-Z.Z., P.N., G.-G.W., and J.-Y.D. did the field work; R.-Z.Z. analyzed the samples and interpreted the results; R.-Z.Z. and P.N. wrote the paper; R.-Z.Z., P.N., G.-G.W. and J.-Y.D. revised the manuscript. All authors read and approved the manuscript.

Funding

This research was funded by the Ministry of Land and Resources of the PRC (Grant No. 201404025) and China Scholarship Council (No. 201706190172).

Acknowledgments

We thank N.K., S.-N.L., T.B., A.-D.Z., B.H., and L.-L.C. for their help during field work. We would like to thank W.-L.Z. for their technical guide during the analysis of the major elements of spinel. We appreciate L.L. and J.-F.G. for their help during the analysis of the trace elements of the spinel. Thanks to the State Key Laboratory for Mineral Deposit Research, Nanjing University for their assistance in the experiment. We appreciate the constructive comments from the Academic Editor and two anonymous reviewers.

Conflicts of Interest

The authors declare no conflict of interest.

References

  1. Hardman, M.F.; Pearson, D.G.; Stachel, T.; Sweeney, R.J. Statistical approaches to the discrimination of crust-and mantle-derived low-Cr garnet–Major-element-based methods and their application in diamond exploration. J. Geochem. Explor. 2018, 186, 24–35. [Google Scholar] [CrossRef]
  2. Grütter, H.S.; Gurney, J.J.; Menzies, A.H.; Winter, F. An updated classification scheme for mantle-derived garnet, for use by diamond explorers. Lithos 2004, 77, 841–857. [Google Scholar] [CrossRef]
  3. Schulze, D.J. A classification scheme for mantle-derived garnets in kimberlite: A tool for investigating the mantle and exploring for diamonds. Lithos 2003, 71, 195–213. [Google Scholar] [CrossRef]
  4. Griffin, W.; Ryan, C. Trace elements in indicator minerals: Area selection and target evaluation in diamond exploration. J. Geochem. Explor. 1995, 53, 311–337. [Google Scholar] [CrossRef]
  5. Griffin, W.; Sobolev, N.; Ryan, C.; Pokhilenko, N.; Win, T.; Yefimova, E. Trace elements in garnets and chromites: Diamond formation in the Siberian lithosphere. Lithos 1993, 29, 235–256. [Google Scholar] [CrossRef]
  6. Gurney, J.J. A correlation between garnets and diamonds in kimberlites. In Kimberlite Occurrence and Origin: A Basis for Conceptual Models in Exploration; University of Western Australia: South Perth, Australia, 1984; Volume 8, pp. 143–166. [Google Scholar]
  7. Sobolev, N.V.; Lavrent’ev, Y.G.; Pokhilenko, N.P.; Usova, L.V. Chrome-rich garnets from the kimberlites of yakutia and their parageneses. Contr. Mineral. Petrol. 1973, 40, 39–52. [Google Scholar] [CrossRef]
  8. Carmody, L.; Taylor, L.A.; Thaisen, K.G.; Tychkov, N.; Bodnar, R.J.; Sobolev, N.V.; Pokhilenko, L.N.; Pokhilenko, N.P. Ilmenite as a diamond indicator mineral in the Siberian craton: A tool to predict diamond potential. Econ. Geol. 2014, 109, 775–783. [Google Scholar] [CrossRef]
  9. Wyatt, B.A.; Baumgartner, M.; Anckar, E.; Grutter, H. Compositional classification of “kimberlitic” and “non-kimberlitic” ilmenite. Lithos 2004, 77, 819–840. [Google Scholar] [CrossRef]
  10. Schulze, D.J. Origins of chromian and aluminous spinel macrocrysts from kimberlites in southern Africa. Can. Mineral. 2001, 39, 361–376. [Google Scholar] [CrossRef]
  11. Gurney, J.J.; Helmstaedt, H.H.; Richardson, S.H.; Shirey, S.B. Diamonds through Time. Econ. Geol. 2010, 105, 689–712. [Google Scholar] [CrossRef]
  12. Shirey, S.B.; Cartigny, P.; Frost, D.J.; Keshav, S.; Nestola, F.; Nimis, P.; Pearson, D.G.; Sobolev, N.V.; Walter, M.J. Diamonds and the geology of Mantle carbon. In Carbon in Earth; Hazen, R.M., Jones, A.P., Baross, J.A., Eds.; Mineralogical Society of America: Chantilly, VA, USA, 2013; Volume 75, pp. 355–421. [Google Scholar]
  13. Gurney, J.J.; Helmstaedt, H.H.; Le Roex, A.P.; Nowicki, T.E.; Richardson, S.H.; Westerlund, K.J. Diamonds: Crustal distribution and formation processes in time and space and an integrated deposit model. Econ. Geol. 2005, 100, 143–177. [Google Scholar]
  14. Gurney, J.J.P.; Zweistra, P. The interpretation of the major element compositions of mantle minerals in diamond exploration. J. Geochem. Explor. 1995, 53, 293–309. [Google Scholar] [CrossRef]
  15. Ryan, C.G.; Griffin, W.L.; Pearson, N.J. Garnet geotherms: Pressure-temperature data from Cr-pyrope garnet xenocrysts in volcanic rocks. J. Geophys. Res. Solid Earth 1996, 101, 5611–5625. [Google Scholar] [CrossRef]
  16. Malkovets, V.G.; Griffin, W.L.; O’Reilly, S.Y.; Wood, B.J. Diamond, subcalcic garnet, and mantle metasomatism: Kimberlite sampling patterns define the link. Geology 2007, 35, 339–342. [Google Scholar] [CrossRef]
  17. Fedortchouk, Y.; Canil, D.; Carlson, J.A. Dissolution forms in Lac de Gras diamonds and their relationship to the temperature and redox state of kimberlite magma. Contr. Mineral. Petrol. 2005, 150, 54–69. [Google Scholar] [CrossRef]
  18. Gurney, J.J.; Hildebrand, P.R.; Carlson, J.A.; Fedortchouk, Y.; Dyck, D.R. The morphological characteristics of diamonds from the Ekati property, Northwest Territories, Canada. Lithos 2004, 77, 21–38. [Google Scholar] [CrossRef]
  19. Alam, M.; Sun, Q. The kinetics of chemical vapor deposited diamond-oxygen reaction. J. Mater. Res. 1993, 8, 2870–2878. [Google Scholar] [CrossRef]
  20. Sonin, V.; Pokhilenko, L.; Pokhilenko, N.; Fedorov, I. Diamond oxidation rate as related to oxygen fugacity. Geol. Ore Depos. 2000, 42, 496–502. [Google Scholar]
  21. Kozai, Y.; Arima, M. Diamond dissolution in kimberlite and lamproite melts at deep crustal conditions. In Proceedings of the 8th International Kimberlite Conference, Victoria, BC, Canada, 22–27 June 2003. Extended Abstracts. [Google Scholar]
  22. Mitchell, R.H. Kimberlites; Plenum: New York, NY, USA, 1986. [Google Scholar]
  23. Roeder, P.L.; Schulze, D.J. Crystallization of groundmass spinel in kimberlite. J. Petrol. 2008, 49, 1473–1495. [Google Scholar] [CrossRef]
  24. Van Straaten, B.I.; Kopylova, M.G.; Russell, J.; Webb, K.J.; Smith, B.H.S. Discrimination of diamond resource and non-resource domains in the Victor North pyroclastic kimberlite, Canada. J. Volcanol. Geotherm. Res. 2008, 174, 128–138. [Google Scholar] [CrossRef]
  25. Ballhaus, C.; Berry, R.F.; Green, D.H. High pressure experimental calibration of the olivine-orthopyroxene-spinel oxygen geobarometer: Implications for the oxidation state of the upper mantle. Contr. Mineral. Petrol. 1991, 107, 27–40. [Google Scholar] [CrossRef]
  26. Zhu, R.Z.; Ni, P.; Ding, J.Y.; Wang, D.Z.; Ju, Y.; Kang, N.; Wang, G.G. Petrography, chemical composition, and Raman spectra of chrome spinel: Constraints on the diamond potential of the No. 30 pipe kimberlite in Wafangdian, North China Craton. Ore Geol. Rev. 2017, 91, 896–905. [Google Scholar] [CrossRef]
  27. Zhu, R.Z.; Ni, P.; Ding, J.Y.; Wang, G.G.; Fan, M.S.; Li, S.N. Metasomatic processes in the lithosphere mantle beneath the No.30 kimberlite (Wafangdian region, North China Craton). Can. Mineral. 2019, 57, 1–19. [Google Scholar] [CrossRef]
  28. Droop, G. A general equation for estimating Fe3+ concentrations in ferromagnesian silicates and oxides from microprobe analyses, using stoichiometric criteria. Mineral. Mag. 1987, 51, 431–435. [Google Scholar] [CrossRef]
  29. Liu, Y.S.; Hu, Z.C.; Gao, S.; Günther, D.; Xu, J.; Gao, C.G.; Chen, H.H. In situ analysis of major and trace elements of anhydrous minerals by LA-ICP-MS without applying an internal standard. Chem. Geol. 2008, 257, 34–43. [Google Scholar] [CrossRef]
  30. Mitchell, R.H. Kimberlites, Orangeites, Lamproites, Melilitites, and Minettes: A Petrographic Atlas; Almaz Press: Thunder Bay, ON, Canada, 1997; p. 243. [Google Scholar]
  31. Mitchell, R.H. Kimberlites and orangeites. In Kimberlites, Orangeites, and Related Rocks; Plenum Press: New York, NY, USA, 1995. [Google Scholar]
  32. Haggerty, S.E. The chemistry and genesis of opaque minerals in kimberlites. Phys. Chem. Earth 1975, 9, 295–307. [Google Scholar] [CrossRef]
  33. Irvine, T. Chromian spinel as a petrogenetic indicator: Part 1. Theory. Can. J. Earth Sci. 1965, 2, 648–672. [Google Scholar] [CrossRef]
  34. Bussweiler, Y.; Stone, R.S.; Pearson, D.G.; Luth, R.W.; Stachel, T.; Kjarsgaard, B.A.; Menzies, A. The evolution of calcite-bearing kimberlites by melt-rock reaction: Evidence from polymineralic inclusions within clinopyroxene and garnet megacrysts from Lac de Gras kimberlites, Canada. Contr. Mineral. Petrol. 2016, 171, 65. [Google Scholar] [CrossRef]
  35. Fedortchouk, Y.; Canil, D. Intensive variables in kimberlite magmas, Lac de Gras, Canada and implications for diamond survival. J. Petrol. 2004, 45, 1725–1745. [Google Scholar] [CrossRef]
  36. Kozai, Y.; Arima, M. Experimental study on diamond dissolution in kimberlitic and lamproitic melts at 1300–1420 °C and 1 GPa with controlled oxygen partial pressure. Am. Miner. 2005, 90, 1759–1766. [Google Scholar] [CrossRef]
  37. Zhao, D.; Zhang, Y.; Essene, E.J. Electron probe microanalysis and microscopy: Principles and applications in characterization of mineral inclusions in chromite from diamond deposit. Ore Geol. Rev. 2015, 65, 733–748. [Google Scholar] [CrossRef]
  38. O’Neill, H.S.C.; Wall, V. The Olivine-Orthopyroxene-Spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth’s Upper Mantle. J. Petrol. 1987, 28, 1169–1191. [Google Scholar] [CrossRef]
  39. Sharygin, I.S.; Litasov, K.D.; Shatskiy, A.; Safonov, O.G.; Golovin, A.V.; Ohtani, E.; Pokhilenko, N.P. Experimental constraints on orthopyroxene dissolution in alkali-carbonate melts in the lithospheric mantle: Implications for kimberlite melt composition and magma ascent. Chem. Geol. 2017, 455, 44–56. [Google Scholar] [CrossRef]
  40. Kamenetsky, V.S.; Yaxley, G.M. Carbonate–silicate liquid immiscibility in the mantle propels kimberlite magma ascent. Geochim. Cosmochim. Acta 2015, 158, 48–56. [Google Scholar] [CrossRef]
  41. Kamenetsky, V.S.; Kamenetsky, M.B.; Sobolev, A.V.; Golovin, A.V.; Sharygin, V.V.; Pokhilenko, N.P.; Sobolev, N.V. Can pyroxenes be liquidus minerals in the kimberlite magma? Lithos 2009, 112, 213–222. [Google Scholar] [CrossRef]
  42. Brett, R.C.; Russell, J.; Moss, S. Origin of olivine in kimberlite: Phenocryst or impostor? Lithos 2009, 112, 201–212. [Google Scholar] [CrossRef]
  43. Barker, D.S. Calculated silica activities in carbonatite liquids. Contr. Mineral. Petrol. 2001, 141, 704–709. [Google Scholar] [CrossRef]
  44. Holland, T.; Powell, R. An internally consistent thermodynamic data set for phases of petrological interest. J. Metamorph. Geol. 1998, 16, 309–343. [Google Scholar] [CrossRef]
  45. Agashev, A.M.; Ionov, D.A.; Pokhilenko, N.P.; Golovin, A.V.; Cherepanova, Y.; Sharygin, I.S. Metasomatism in lithospheric mantle roots: Constraints from whole-rock and mineral chemical composition of deformed peridotite xenoliths from kimberlite pipe Udachnaya. Lithos 2013, 160–161, 201–215. [Google Scholar] [CrossRef]
  46. Sharygin, I.S.; Litasov, K.D.; Shatskiy, A.; Golovin, A.V.; Ohtani, E.; Pokhilenko, N.P. Melting phase relations of the Udachnaya-East Group-I kimberlite at 3.0–6.5 GPa: Experimental evidence for alkali-carbonatite composition of primary kimberlite melts and implications for mantle plumes. Gondwana Res. 2015, 28, 1391–1414. [Google Scholar] [CrossRef]
  47. Nixon, P.H. A review of mantle xenoliths and their role in diamond exploration. J. Geodyn. 1995, 20, 305–329. [Google Scholar] [CrossRef]
  48. Stachel, T.; Aulbach, S.; Brey, G.P.; Harris, J.W.; Leost, I.; Tappert, R.; Viljoen, K.S. The trace element composition of silicate inclusions in diamonds: A review. Lithos 2004, 77, 1–19. [Google Scholar] [CrossRef]
  49. Stachel, T.; Harris, J.W. The origin of cratonic diamonds—Constraints from mineral inclusions. Ore Geol. Rev. 2008, 34, 5–32. [Google Scholar] [CrossRef]
  50. Stachel, T.; Harris, J.W. Formation of diamond in the Earth’s mantle. J. Phys. Condens. Matter 2009, 21, 364206. [Google Scholar] [CrossRef]
  51. Haggerty, S.E. A diamond trilogy: Superplumes, supercontinents, and supernovae. Science 1999, 285, 851–860. [Google Scholar] [CrossRef]
  52. Stachel, T.; Brey, G.P.; Harris, J.W. Inclusions in sublithospheric diamonds: Glimpses of deep Earth. Elements 2005, 1, 73–78. [Google Scholar] [CrossRef]
  53. Canil, D. The Ni-in-garnet geothermometer: Calibration at natural abundances. Contr. Mineral. Petrol. 1999, 136, 240–246. [Google Scholar] [CrossRef]
Minerals 09 00382 g001 550
Figure 1. Backscatter electron (BSE) images of magmatic spinels (A,B) and xenocrystic spinels (C,D) from the No.30 kimberlite. Magmatic spinels show compositional zones from the Ti-Al-Mg chromite core through the Ti-Al chromite intermediate to the magnetite rim, and spinel xenocrysts show compositional zones from xenocrystic magnesiochromite core through the Ti-Al-Mg chromite intermediate to the magnetite rim.
Figure 1. Backscatter electron (BSE) images of magmatic spinels (A,B) and xenocrystic spinels (C,D) from the No.30 kimberlite. Magmatic spinels show compositional zones from the Ti-Al-Mg chromite core through the Ti-Al chromite intermediate to the magnetite rim, and spinel xenocrysts show compositional zones from xenocrystic magnesiochromite core through the Ti-Al-Mg chromite intermediate to the magnetite rim.
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Figure 2. Variations of MgO (A), Al2O3 (B), TiO2 (C), and Fe2O3 (D) versus Cr2O3 contents in the magmatic and xenocrystic spinels from the No.30 kimberlite.
Figure 2. Variations of MgO (A), Al2O3 (B), TiO2 (C), and Fe2O3 (D) versus Cr2O3 contents in the magmatic and xenocrystic spinels from the No.30 kimberlite.
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Minerals 09 00382 g003 550
Figure 3. Three spinel projections of oxidized spinel prism for magmatic and xenocrystic spinels from the No.30 kimberlite pipe: (A) Cr/(Cr + Al) versus Fe2+/(Fe2+ + Mg), (B) Fe3+/(Fe3+ + Al + Cr) versus Fe2+/(Fe2+ + Mg) and (C) Fe3+-Al-Cr. The iso-potential lines (dotted lines) indicating that the composition of the spinels that is in equilibrium with olivine (Fo 90 and 80) at a constant temperature of 1200 °C and 1100 °C is after Irvine (1965). Curves of trend 1 and trend 2 are after Mitchell [22] and Roeder and Schulze [23]. Abbreviations: Xen = xenocryst peridotite spinel; MUM = magnesium-ulvöspinel-magnetite; Chr = chromite; Xen’ = Metasomatized xenocryst peridotite spinel; Mag = magnetite.
Figure 3. Three spinel projections of oxidized spinel prism for magmatic and xenocrystic spinels from the No.30 kimberlite pipe: (A) Cr/(Cr + Al) versus Fe2+/(Fe2+ + Mg), (B) Fe3+/(Fe3+ + Al + Cr) versus Fe2+/(Fe2+ + Mg) and (C) Fe3+-Al-Cr. The iso-potential lines (dotted lines) indicating that the composition of the spinels that is in equilibrium with olivine (Fo 90 and 80) at a constant temperature of 1200 °C and 1100 °C is after Irvine (1965). Curves of trend 1 and trend 2 are after Mitchell [22] and Roeder and Schulze [23]. Abbreviations: Xen = xenocryst peridotite spinel; MUM = magnesium-ulvöspinel-magnetite; Chr = chromite; Xen’ = Metasomatized xenocryst peridotite spinel; Mag = magnetite.
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Figure 4. Variations of Cr2O3 (A), Al2O3 (B), TiO2 (C), Ni (D), Zr (E), and Nb (F) versus TZn in the cores of the xenocrystic spinels from the No.30 kimberlite. The TZn are the temperature estimates based on the Zn-in-chromite thermometer [15].
Figure 4. Variations of Cr2O3 (A), Al2O3 (B), TiO2 (C), Ni (D), Zr (E), and Nb (F) versus TZn in the cores of the xenocrystic spinels from the No.30 kimberlite. The TZn are the temperature estimates based on the Zn-in-chromite thermometer [15].
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Figure 5. Temperature (depth) distribution of xenocrystic garnets and chromites from the No.30 kimberlite. Equilibrium temperatures of garnets [27] are calculated using the Ni-in-garnet thermometer [53], and equilibrium temperatures with the cores of xenocrystic spinels are estimated using Zn-in-chromite thermometer [15].
Figure 5. Temperature (depth) distribution of xenocrystic garnets and chromites from the No.30 kimberlite. Equilibrium temperatures of garnets [27] are calculated using the Ni-in-garnet thermometer [53], and equilibrium temperatures with the cores of xenocrystic spinels are estimated using Zn-in-chromite thermometer [15].
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Figure 6. Estimated log fO2 oxygen fugacity conditions (ΔNNO; derived from Ol–Sp thermobarometry at an assumed pressure of 1 GPa) for the No.30 kimberlite (this study) compared with Misery, Beartooth, Panda, Grizzly, Leslie, and Ranch Lake kimberlites from the Lac de Gras area [17] (the numbers in parentheses represent diamond grades of kimberlites in units of carats/ton).
Figure 6. Estimated log fO2 oxygen fugacity conditions (ΔNNO; derived from Ol–Sp thermobarometry at an assumed pressure of 1 GPa) for the No.30 kimberlite (this study) compared with Misery, Beartooth, Panda, Grizzly, Leslie, and Ranch Lake kimberlites from the Lac de Gras area [17] (the numbers in parentheses represent diamond grades of kimberlites in units of carats/ton).
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Table
Table 1. Mineral chemistry (oxide wt. %) of magmatic spinels from the No.30 kimberlite.
Table 1. Mineral chemistry (oxide wt. %) of magmatic spinels from the No.30 kimberlite.
Oxide (wt. %)WFD148-SP1WFD148-SP2WFD148-SP3WFD148-SP4WFD148-SP5WFD148-SP6WFD148-SP7
Core RimCore RimCore RimCoreRimCore RimCore RimCore Rim
TiO23.84 0.353.92 0.793.99 0.693.510.344.20 0.393.75 0.244.04 0.77
Al2O36.49 0.035.76 0.225.30 0.346.450.084.58 0.436.21 0.055.38 0.15
Cr2O351.86 1.5751.67 0.7250.11 0.5350.040.2349.75 0.7050.13 0.9752.22 0.95
Fe2O37.47 65.877.38 63.388.92 63.368.2364.889.71 63.488.90 64.818.24 64.27
FeO21.21 30.7221.92 29.3423.21 29.1423.8029.3923.48 28.4622.02 29.6821.57 29.72
MgO9.85 0.179.18 0.518.18 0.567.630.288.09 0.749.07 0.149.64 0.57
MnO0.58 0.010.56 0.180.81 0.090.800.080.75 0.120.63 0.160.69 0.12
Total101.30 98.71100.38 95.14100.53 94.71100.4695.27100.56 94.32100.69 96.04101.77 96.55
Ti0.096 0.0100.100 0.0240.102 0.0210.0900.0100.108 0.0120.095 0.0070.101 0.023
Al0.255 0.0020.230 0.0110.213 0.0160.2590.0040.185 0.0210.247 0.0020.212 0.007
Cr1.366 0.0481.383 0.0231.353 0.0171.3490.0071.348 0.0221.337 0.0311.378 0.030
Fe3+0.187 1.9300.188 1.9190.229 1.9250.2111.9690.250 1.9330.226 1.9530.207 1.917
Fe2+0.591 1.0000.620 0.9870.663 0.9840.6790.9910.673 0.9630.621 0.9940.602 0.985
Mg0.489 0.0100.463 0.0300.416 0.0340.3880.0170.413 0.0450.456 0.0080.480 0.034
Mn0.016 0.0000.016 0.0060.024 0.0030.0230.0030.022 0.0040.018 0.0050.019 0.004
Total3.00 3.003.00 3.003.00 3.003.003.003.00 3.003.00 3.003.00 3.00
Cr/(Cr + Al)0.84 0.970.86 0.690.86 0.510.840.660.88 0.520.84 0.930.87 0.81
Fe2+/(Fe2+ + Mg)0.55 0.990.57 0.970.61 0.970.640.980.62 0.960.58 0.990.56 0.97
Mg/(Mg + Fe2+)0.45 0.010.43 0.030.39 0.030.360.020.38 0.040.42 0.010.44 0.03
Fe3+/(Fe3+ + Al + Cr)0.10 0.970.10 0.980.13 0.980.120.990.14 0.980.12 0.980.12 0.98
Oxide (wt. %)WFD165-SP8WFD165-SP9WFD165-SP10WFD165-SP11WFD165-SP12WFD165-SP13WFD165-SP14
CoreIntermediateRimCoreIntermediateRimCoreIntermediateRimCoreRimCoreIntermediateRimCoreIntermediateRimCoreIntermediateRim
TiO24.367.061.454.1810.340.464.556.391.474.291.974.185.220.924.517.664.214.148.854.63
Al2O35.787.530.246.067.170.014.765.190.415.340.265.194.540.135.457.871.545.998.490.15
Cr2O353.1832.160.4152.9116.360.1652.5040.440.9849.990.1853.3745.982.7252.6527.810.3452.9021.040.45
Fe2O37.2314.4564.278.8023.0666.958.5311.5263.729.3863.488.9710.1563.508.5617.6356.937.9320.1757.84
FeO17.3235.8031.7815.7539.7730.8919.0531.5231.9420.7132.2716.5629.1130.2916.6136.1934.1416.4238.3334.20
MgO12.661.090.0313.850.370.0511.553.290.069.870.0513.224.630.6813.341.280.1513.210.300.16
MnO0.401.950.060.311.310.060.542.040.021.020.010.311.490.110.291.900.010.262.040.17
Total100.92100.0598.24101.8798.3898.58101.48100.3798.59100.6098.22101.80101.1298.35101.40100.3497.32100.8699.2297.59
Ti0.1080.1900.0430.1020.2860.0140.1130.1700.0430.1090.0580.1020.1370.0270.1110.2060.1230.1020.2410.136
Al0.2240.3180.0110.2310.3110.0010.1860.2170.0190.2120.0120.2000.1870.0060.2100.3310.0710.2310.3630.007
Cr1.3820.9110.0131.3520.4760.0051.3751.1350.0301.3320.0061.3751.2720.0841.3590.7850.0101.3700.6030.014
Fe3+0.1790.3901.8910.2140.6391.9670.2130.3071.8650.2381.8670.2200.2671.8570.2100.4731.6720.1950.5511.707
Fe2+0.4761.0731.0390.4261.2251.0090.5280.9351.0390.5841.0550.4510.8520.9840.4531.0801.1140.4501.1631.121
Mg0.6210.0580.0020.6670.0200.0030.5710.1740.0030.4960.0030.6430.2410.0390.6490.0680.0090.6450.0160.009
Mn0.0110.0590.0020.0080.0410.0020.0150.0610.0010.0290.0000.0080.0440.0040.0080.0580.0000.0070.0630.006
Total3.003.003.003.003.003.003.003.003.003.003.003.003.003.003.003.003.003.003.003.00
Cr/(Cr + Al)0.860.740.540.850.600.900.880.840.620.860.320.870.870.930.870.700.130.860.620.67
Fe2+/(Fe2+ + Mg)0.430.951.000.390.981.000.480.841.000.541.000.410.780.960.410.940.990.410.990.99
Mg/(Mg + Fe2+)0.570.050.000.610.020.000.520.160.000.460.000.590.220.040.590.060.010.590.010.01
Fe3+/(Fe3+ + Al + Cr)0.100.240.990.120.451.000.120.190.970.130.990.120.150.950.120.300.950.110.360.99
Table
Table 2. Trace element concentrations and equilibrium temperature estimated using the Zn-in-chromite method [15] in the xenocrystic spinels from the No.30 kimberlite.
Table 2. Trace element concentrations and equilibrium temperature estimated using the Zn-in-chromite method [15] in the xenocrystic spinels from the No.30 kimberlite.
ppmL30-01L30-02L30-03L30-04L30-05L30-06L30-07L30-08L30-09L30-10L30-11L30-12L30-13L30-14L30-15L30-16L30-17
Zn1047 452 662 453 571 657 910 850 803 482 698 677 609 358 375 363 359
Ga85.6 22.7 20.1 40.9 17.3 12.1 52.5 25.9 75.1 27.9 37.8 46.5 24.5 27.4 26.5 27.0 39.2
Zr0.23 0.83 6.13 1.44 0.36 1.12 1.97 0.17 0.15 1.80 0.51 0.16 0.24 5.35 5.31 5.44 3.09
Nb2.74 1.22 2.43 2.06 1.57 0.30 1.87 0.07 0.62 1.59 0.12 0.38 1.00 2.39 2.26 2.18 2.07
Ni696 957 994 1147 788 843 731 668 1327 977 1490 1373 750 1432 1437 1425 1821
TZn (°C)828 1177 995 1176 1059 998 874 898 919 1143 973 986 1031 1316 1286 1307 1316
ppmL30-18L30-19L30-20L30-21L30-22L30-23L30-24L30-25L30-26L30-27L30-28L30-29L30-30L30-31L30-32L30-33L30-34
Zn551 610 751 764 723 648 595 729 611 588 478 416 479 655 806 411 320
Ga15.4 41.5 62.7 29.1 101.5 61.0 38.1 13.8 10.1 51.4 17.2 20.8 26.3 35.5 38.3 65.9 88.0
Zr4.02 0.11 0.07 0.20 1.32 0.28 1.62 2.36 9.31 1.41 0.41 2.74 2.35 3.25 2.55 4.82 10.39
Nb1.19 0.65 0.07 0.49 3.39 0.71 2.16 1.37 2.32 1.38 1.00 1.38 1.04 1.37 1.46 2.95 3.83
Ni957 1435 978 702 1494 832 927 778 858 1491 808 955 1024 1563 1361 2041 2306
TZn (°C)1077 1030 944 938 959 1004 1041 956 1029 1047 1147 1224 1146 1000 918 1231 1394
ppmL30-35L30-36L30-37L30-38L30-39L30-40L30-41L30-42L30-43L30-44L30-45L30-46L30-47L30-48L30-49L30-50L30-51
Zn660 655 815 695 478 815 520 419 903 452 751 397 484 736 809 793 524
Ga6.4 7.4 53.3 79.0 29.0 46.0 29.7 15.1 49.6 13.1 18.6 85.5 17.0 87.8 23.3 39.3 29.6
Zr1.30 11.3 5.37 0.35 2.08 0.08 0.28 4.56 0.94 0.36 0.70 5.93 0.95 0.05 10.66 2.68 2.16
Nb0.35 3.21 5.05 3.99 1.80 0.21 1.36 2.38 1.64 0.63 3.25 3.33 1.15 0.10 2.13 7.07 1.82
Ni846 922 1059 1048 981 1317 931 1308 976 838 701 1451 858 1089 1064 1075 888
TZn (°C)996 999 913 975 1147 914 1105 1220 877 1177 944 1251 1141 952 916 924 1100
Table
Table 3. Equilibrium temperatures and oxygen fugacities estimated for the No.30 kimberlite magma by Ol-Sp thermobarometry assuming a pressure of 1 GPa.
Table 3. Equilibrium temperatures and oxygen fugacities estimated for the No.30 kimberlite magma by Ol-Sp thermobarometry assuming a pressure of 1 GPa.
SampleOl-Sp T (°C) aOl-Sp Oxygen Fugacity blog aSiO2 cCorrected d
Δlog fO2FMQ log fO2Δlog fO2NNODi-MontEn-FoΔlog fO2FMQΔlog fO2NNO
WFD148-SP11249 1.5 −9.8 1.0 −1.40 −0.41 −1.5 −2.0
WFD148-SP21218 1.5 −10.1 1.1 −1.46 −0.43 −1.6 −2.0
WFD148-SP31168 1.9 −10.5 1.5 −1.54 −0.46 −1.4 −1.8
WFD148-SP41102 1.7 −11.7 1.3 −1.65 −0.50 −1.7 −2.1
WFD148-SP51181 2.0 −10.1 1.6 −1.52 −0.45 −1.1 −1.6
WFD148-SP61214 1.8 −10.0 1.4 −1.46 −0.43 −1.3 −1.7
WFD148-SP71265 1.6 −9.4 1.2 −1.38 −0.40 −1.3 −1.7
WFD165-SP81170 1.2 −6.9 0.8 −1.54 −0.46 −2.0 −2.4
WFD165-SP91274 1.5 −5.5 1.1 −1.36 −0.40 −1.4 −1.8
WFD165-SP101107 1.6 −7.3 1.2 −1.65 −0.50 −1.8 −2.2
WFD165-SP11994 1.9 −8.6 1.5 −1.84 −0.57 −1.9 −2.4
WFD165-SP121233 1.6 −5.9 1.2 −1.43 −0.42 −1.5 −1.9
WFD165-SP13 1240 1.5 −5.9 1.1 −1.42 −0.42 −1.5 −1.9
WFD165-SP141217 1.3 −6.3 1.0 −1.46 −0.43 −1.7 −2.1
a Temperatures calculated from the Fe/Mg exchange Ol-Sp thermometer of O’Neill and Wall [38] and Ballhaus et al. [25]; b Oxygen fugacities at 1 GPa calculated for the fayalite-magnetite-quartz (FMQ) buffer from the oxygen barometer by Ballhaus et al. [25]; c Silica activity of Diopside-Monticellite (Di-Mont) and Enstatite-Forsterite (En-Fo) buffers calculated using thermodynamic data of Holland and Powell [44]. d Corrected values of oxygen fugacities of kimberlite calculated for silica activity of Diopside-Monticellite buffer [35]. The equation for nickel-nickel oxide (NNO) buffer is from Ballhaus et al. [25].

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Zhu, R.-Z.; Ni, P.; Ding, J.-Y.; Wang, G.-G. Geochemistry of Magmatic and Xenocrystic Spinel in the No.30 Kimberlite Pipe (Liaoning Province, North China Craton): Constraints on Diamond Potential. Minerals 2019, 9, 382. https://doi.org/10.3390/min9060382

AMA Style

Zhu R-Z, Ni P, Ding J-Y, Wang G-G. Geochemistry of Magmatic and Xenocrystic Spinel in the No.30 Kimberlite Pipe (Liaoning Province, North China Craton): Constraints on Diamond Potential. Minerals. 2019; 9(6):382. https://doi.org/10.3390/min9060382

Chicago/Turabian Style

Zhu, Ren-Zhi, Pei Ni, Jun-Ying Ding, and Guo-Guang Wang. 2019. "Geochemistry of Magmatic and Xenocrystic Spinel in the No.30 Kimberlite Pipe (Liaoning Province, North China Craton): Constraints on Diamond Potential" Minerals 9, no. 6: 382. https://doi.org/10.3390/min9060382

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