Petrogenesis and metallogenesis of the Taihe gabbroic intrusion associated with Fe–Ti-oxide ores in the Panxi district, Emeishan Large Igneous Province, southwest China

Ore Geology Reviews 49 (2012) 109–127

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Petrogenesis and metallogenesis of the Taihe gabbroic intrusion associated with Fe–Ti-oxide ores in the Panxi district, Emeishan Large Igneous Province, southwest China Tong Hou a, Zhaochong Zhang a,⁎, John Encarnacion b, M. Santosh a, c a b c

State Key Laboratory of Geological Process and Mineral Resources, China University of Geosciences, Beijing, 100083, China Department of Earth and Atmospheric Sciences, Saint Louis University, 3642 Lindell Boulevard, St. Louis, MO 63108, USA Division of Interdisciplinary Science, Kochi University, Kochi 780-8520, Japan

a r t i c l e

i n f o

Article history: Received 6 February 2012 Received in revised form 25 June 2012 Accepted 18 September 2012 Available online 25 September 2012 Keywords: Taihe Panxi Emeishan Fe–Ti oxide Geochemistry Petrogenesis Metallogenesis

a b s t r a c t The Taihe layered gabbro intrusion in the northernmost part of the Panxi district in southwest China is part of the 260 million year old Emeishan Large Igneous Province. This intrusion hosts a giant Fe–Ti oxide deposit with 810 million tonnes of ore reserves, which makes it one of the largest deposits in the Panxi district. The intrusion covers an areal extent of ~13 km2 and has a vertical stratigraphic thickness of ~1400 m. It can be divided into a lower zone (LZ) of coarse-grained gabbro, apatite-bearing gabbro, troctolite and intercalated gabbro and clinopyroxene-bearing troctolite, followed upward by a middle zone (MZ) of gabbro and intercalated clinopyroxenite, plagioclase-bearing clinopyroxenite with major oxide layers, and an upper zone (UZ) of olivine gabbro and layered gabbro including unmineralized leucogabbro and melanogabbro, with some small oxide ore bodies in the lower part. Each of these ‘zones’ contains oxide minerals and relatively similar lithologies. Ore textures and associated mineral assemblages indicate that the ore bodies formed by crystallization of Fe–Ti–V-rich melt under high oxygen fugacity and a volatile-rich environment during the late-stage of magmatic differentiation. A general systematic variation of major oxides is seen through the intrusion as reflected by a slight overall decrease in MgO and Fe2O3 (as total iron) and an increase in SiO2, Na2O, Al2O3, and CaO upward in the layered sequences. Based on lithology and bulk-rock geochemical features, such as positive Eu anomalies, the Taihe intrusion is inferred to have been derived from a ferropicritic melt and became more evolved in chemistry upward following a tholeiitic differentiation trend with enrichment in Fe, Ti, and V. The Taihe gabbros define a small range of age-corrected εNd(t) (t= 260 Ma) from −0.6 to 0.7 and (87Sr/86Sr)t ratios ranging from 0.7040 to 0.7050. The relatively lower εNd(t) values and higher (87Sr/86Sr)t ratios compared to those from Lijiang picrite which represents the initial product of Emeishan plume head, combined with the enrichment in light rare earth element (LREE) relative to heavy rare earth element (HREE) as well as negative high field strength element (HFSE; e.g., Nb, Ta, Zr, and Hf) anomalies, suggest that subduction-related material was involved in the source region. We propose that the parental picritic magma was generated from the interaction of the ~260 million year Emeishan mantle plume with the lithospheric mantle, and the picritic magma interacted with an eclogitic component in the lithospheric mantle. In our view, the eclogitic component was derived from the earlier Neoproterozoic subduction. The junction of these subduction-modified lithospheric mantle sources and the Emeishan plume was a possible crucial factor leading to the production of large Fe–Ti oxide deposits in the Panxi area. © 2012 Elsevier B.V. All rights reserved.

1. Introduction Magmatic Fe–Ti oxide ores are commonly associated with, or hosted in, layered mafic intrusions or Proterozoic anorthosite complexes (Bateman, 1951; Cawthorn, 1996; Force, 1991; Lister, 1966; Ram Mohan et al., 2012). However, the mechanisms by which millions of tonnes of Fe, Ti and V become concentrated to form massive Fe–Ti

⁎ Corresponding author. Tel.: +86 10 82322195; fax: +86 10 82322176. E-mail address: [email protected] (Z. Zhang). 0169-1368/$ – see front matter © 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.oregeorev.2012.09.004

oxide deposits remain poorly understood. For example, although stratiform Fe–Ti oxide ores, such as those of the uppermost part of the Bushveld Complex in South Africa and the Bjerkreim–Sokndal in Norway are generally thought to have formed as a result of magma mixing and/or crustal contamination in a dynamic layered intrusion (Cawthorn, 1996), alternative models proposed include gravitational differentiation (Charlier et al., 2006, 2009; Wager and Brown, 1968), increasing oxygen fugacity (Botcharnikov et al., 2008; Toplis and Carroll, 1995) and periodic pressure fluctuation (e.g. Cawthorn and Ashwal, 2009). However, the precise mechanisms by which the oxides crystallized and accumulated are poorly known.

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The Panxi (short for Panzhihua-Xichang) district in southwest China is the most important Fe–Ti oxide ore district in China (e.g. Zhang et al., 2009; Zhong et al., 2002, 2003, 2004, 2009, 2011). Here, several mafic– ultramafic intrusions host some of the world's largest Fe–Ti oxide deposits and several Cu-Ni-(PGE) sulfide deposits (Song et al., 2003, 2009; Tao et al., 2007, 2008). The Fe–Ti oxide ore deposits include Panzhihua, Hongge, Baima, and Taihe, which have been a major source of V, Ti, and Fe since the 1960s, making China a major producer of these metals (Ma et al., 2003). Until now, the majority of previous studies have focused on the Panzhihua intrusion, which is a typical layered mafic intrusion hosting Fe–Ti oxide ores in highly differentiated rocks. Field observations have shown that in the Panzhihua deposit, part of disseminated ores typically occur as conformable layers within layered intrusions, or occur in oxide-rich gabbro, whereas the massive bodies always exhibit sharp contacts with the surrounding gabbro. The origin of massive ores and oxide-rich gabbro is not straightforward, although it is generally held that a magmatic origin is likely. Two main mechanisms of Fe–Ti–V-enrichment have been proposed for the Panzhihua intrusion: crystal accumulation processes (e.g. Zhang et al., 2009), and liquid immiscibility (Zhou et al., 2005). However, as yet, there is no direct evidence from experimental work to support the mechanism of liquid immiscibility. Crystal accumulation, as evidenced by bulk-rock chemistry and cumulus textures in a late-stage gabbroic magma residue after Fe–Ti-poor phase assemblage crystallization has been suggested as the main mechanism leading to the formation of the Fe–Ti oxide ores in the Panxi district (e.g. Ganino et al., 2008; Pang et al., 2008). Such a mechanism requires a fractionated liquid enriched in Fe and Ti. Although the concentration of Ti could be increased by the crystallization of mafic minerals, such as olivine and clinopyroxene, the enrichment of Fe in the residual liquid requires considerable fractionation of Fe-poor minerals such as plagioclase. However, no cumulus plagioclase or anorthositic rocks during the early stages of crystallization have been recognized in the deposits of the Panxi district. This raises an important question: can the model that was conducted from the studies of the Panzhihua deposit be applied to the Taihe deposit and other deposits in the Panxi district, or to the other magmatic Fe–Ti oxide deposits in the world? The Taihe Fe–Ti oxide deposit is the largest one in the northernmost part of the Panxi district (Fig. 1) and was not mined prior to 2008. Due to poor exposure, only few investigations were performed (Shellnutt and Zhou, 2007). Currently, the deposit is being mined by Chongqing Iron and Steel Company, with the mining activities revealing new exposures and providing an opportunity to carry out systematic investigations and sampling for petrological and geochemical studies and to constrain the petrogenesis and metallogenesis of the deposit. In this paper, we present representative mineral chemistry, major and trace element geochemistry on 17 gabbroic samples Sr–Nd isotopic data on 10 samples from the Taihe deposit, aiming to constrain the petrogenesis of the mafic rocks and the mechanism of Fe–Ti oxide ore formation and its relationship with the host gabbroic intrusion. We also address the question of why so many Fe–Ti oxide deposits are clustered in the Panxi district and discuss the geologic setting, source composition, depth of melting, and magma chamber processes leading to the ore formation. 2. Geological setting 2.1. Regional geology The Yangtze Block in South China consists of Mesoproterozoic granitic gneisses and metasedimentary rocks, which have been intruded by the Neoproterozoic (~ 800 Ma) Kangdian granites (Zhou et al., 2002b), and overlain by marine and terrestrial strata of late Neoproterozoic (~ 600 Ma) to Late Permian age (Wang et al., 2012a; Yan et al., 2003). Permian rocks include limestone and the Emeishan continental flood basalts (Emeishan Large Igneous Province; ELIP; Ali

et al., 2010; Shellnutt and Jahn, 2009; Wu and Zhang, 2012). The Triassic strata include both continental and marine sedimentary rocks, whereas the Jurassic to Cretaceous strata are entirely terrestrial. Neoproterozoic arc plutonic-metamorphic assemblages occur along the western and northern margin of the Yangtze Block (e.g., Wang et al., 2012b), which are correlated to subduction of Rodinian oceanic lithosphere toward the Yangtze Block during the period from 860 to 760 million years (Zhou et al., 2002a). A late Paleozoic to early Mesozoic (~280–230 Ma) rifting event has also been recognized (Cong, 1988). The rocks were further deformed during the Paleogene India-Asia collision (Yin and Harrison, 2000). 2.2. Emeishan province and associated Fe–Ti oxide deposits General descriptions of the Emeishan province have been presented by several workers (e.g., Chung and Jahn, 1995; Xu et al., 2001; Xiao et al., 2004; Zhang et al., 2006; Wang et al., 2008; Shellnutt and Jahn, 2009). Here we present only a summary. The province is dominated by the Emeishan flood basalts, which ranges in thickness from a few hundred meters to a maximum of ~ 5 km. Most of the province lies within the broad region of Cenozoic uplift caused by the India–Eurasian collision (Chung and Jahn, 1995). As a consequence, the lava pile is deeply dissected, especially in the western parts of the province. Erosional remnants of the flood basalts are distributed over an area of at least 2.5 × 10 5 km 2. The flood basalts were likely emplaced at or close to sea level (Ukstins Peate and Bryan, 2009). In contrast to the Siberian flood basalts, which formed at a relatively high northern latitude, the emplacement of the Emeishan flood basalts occurred near the Equator (Enkin et al., 1992). Overall, the province appears to be slightly older than the ~ 251 Ma Siberian Traps, with published 40Ar– 39Ar ages of 254 ± 5 Ma (Boven et al., 2002), 255 Ma, and 251–253 Ma (Lo et al., 2002). More recent SHRIMP zircon U–Pb dating on mafic intrusions, dykes and volcanic rocks of the ELIP have indicated emplacement between 257 and 263 Ma (e.g. He et al., 2007). From precise isotope dating of zircons, Shellnut et al. (2012) demonstrated that the Emeishan magmatism was short-lived, with a rapid (≤2 Ma) emplacement. The Panxi district lies in the central-western part of the ELIP where flood basalts are variably deformed, uplifted and eroded due to strong tectonic activity in the Cenozoic because of India-Eurasia collision. Magmatic Fe–Ti oxide deposits are documented in several layered intrusions in the district, the distribution of which can be explained by exhumation along a major N–S faults. The ore deposits and host layered intrusions include Panzhihua, Hongge, Baima, Taihe and Xinjie (Ma et al., 2003). Most of the intrusions show U–Pb zircon magmatic crystallization ages of ∼260 Ma (Fig. 1; Zhang et al., 2009 and references therein). Giant Fe–Ti–V oxide deposits occur in several relatively large layered intrusions which are spatially associated with contemporaneous flood basalts and many granitoids (Fig. 1). The total oxide ore reserves for the four large Fe–Ti–V oxide deposits are estimated to be 7544 million tonnes with an average ore grade of 36 wt.% Fetotal, 0.28 wt.% V2O5 and 12.6 wt.% TiO2 (Ma et al., 2003). 2.3. Taihe intrusion The Taihe gabbroic intrusion is exposed 12 km west of the Xichang city in Sichuan Province. The intrusion is an ovoid-shaped sill-like body extending NW-SE along a strike length of ca. 3 km, dipping 50–60° NE, and covers an area of ∼ 13 km 2 with an estimated thickness of ~ 1.4 km (Fig. 2). The intrusion contains an ore reserve of 810 million tonnes with a mean grade of ~ 33% total Fe, ~ 12% TiO2 and ~ 0.3% V2O5 (Ma et al., 2003). Published laser-ablation inductively coupled plasma mass spectrometry U–Pb zircon ages by Zhong et al. (2009) indicate a crystallization age of 258.0 ± 1.9 Ma. The eastern part of the ore-bearing gabbroic body is covered by Quaternary alluvial deposits. The roof of the intrusion has been denuded, and

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Fig. 1. Map showing main outcrops of the Emeishan flood basalts and related mafic–ultramafic intrusions in the central part of the Emeishan Large Igneous Province, SW China (modified from Zhang et al., 2009). Ages of the intrusions are compiled from Zhou et al. (2005) and Zhong et al. (2009).

the floor of the intrusion has not been reached by current mining and the deep drill holes. Northwestern and southern parts of the intrusion are intruded by late Permian alkaline granite and syenite (Shellnutt and Zhou, 2007; Shellnutt et al., 2010). The wall rocks include Sinian metamorphosed sandstone, and Triassic shales and coal measures, which are in fault contact with the intrusion. This contact dips NW and is interpreted as a thrust fault. However, the fault between the intrusion and adjacent granite as shown on local geological maps is not visible in the field. Internally, the intrusion is cut by several brittle faults, which have disrupted the continuity of primary igneous structures. The intrusion is cut by diabase and syenitic dykes, ranging from ~ 5 to 20 m in thickness. Granites to the west of the intrusion were dated

at ~ 260 Ma (Luo et al., 2006). The ores occur over an interval of 400 to 700 m of igneous stratigraphy in the lower part of the intrusion. Based on differences in internal structure and the extent of oxide mineralization, domestic geologists (Li et al., 1981) have identified three “zones” in the intrusion (from base to top): a lower zone (LZ), a middle zone (MZ), and an upper zone (UZ). Specifically, the ~ 1.4 km-thick mafic stratigraphy consists of: (1) a ~ 200 m thick LZ comprising coarse-grained gabbro, apatite-bearing gabbro, troctolite and intercalated gabbro, and clinopyroxene-bearing troctolite; (2) a 300 to 350 m thick MZ of gabbro and intercalated clinopyroxenite, plagioclase-bearing clinopyroxenites with major oxide layers (ore bodies) up to 150 m-thick that mainly occur in the lower part (Fig. 3), with the abundance of Fe–Ti oxides in the rocks decreasing upwards;

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oxides with variable amounts of clinopyroxene, plagioclase, and olivine. The sparse silicate minerals occur as interstitial phases (~ 5%) (Fig. 4f). Where they are in contact with oxides, the silicate minerals are always rimmed with brown hornblende (Fig. 4e), implying higher H2O content of the evolved Fe–Ti-rich magma where these minerals crystallized. Moreover, throughout the intrusion, there is no significant reversal of differentiation trends as marked by a gradual or abrupt increase in the An content of plagioclase upsection in the intrusion that would suggest an influx of fresh primitive magma (Li et al., 1981) (Fig. 3).

3. Analytical methods

Fig. 2. Simplified geological map for the Taihe area (Modified from Zhang, et al., 1988).

and (3) the UZ is a 800 m-thick sequence of olivine gabbro, and layered gabbro, including unmineralized leucogabbro and melanogabbro, with some small oxide ore bodies in the lower part. These barren gabbros are relatively rich in apatite (Fig. 3). Small amounts of anorthosite, syenite, granophyre, and felsic pegmatite also occur as dykes or lenses within the intrusion. Both the MZ and lower part of the UZ of the intrusion are well layered with cm to dm scale rhythmic layering, similar to those described in the Fe–Ti oxide ore-bearing portion of the Skaergaard intrusion (e.g. McBirney, 1996) and the Bushveld intrusion (e.g. Cawthorn, 1996; Cawthorn and Spies, 2003). Both modal layering and grain-size layering are present, but the layering varies in form, frequency, and spacing. On the basis of their range of modal mineralogy, the vertical variations of silicate and oxide minerals in each layer are shown in the stratigraphic section presented in Fig. 3. Troctolites occur only in the LZ and consist predominantly of approximately 50% (as modal proportions, hereinafter) euhedral plagioclase and 40% olivine enclosed by anhedral clinopyroxenes with minor hornblende, which display a typical poikilitic texture (Fig. 4a). The gabbros are composed of ~ 50% clinopyroxene and ~ 40% plagioclase, which are locally replaced by tremolite and albite, respectively, but the majority of rocks are unaltered (Fig. 4b). Some euhedral Fe–Ti oxide minerals (~ 5%) occur as inclusions in the olivine grains, whereas most of them occur as anhedral interstitial minerals in clinopyroxene and plagioclase (Fig. 4a). The melanogabbros are coarse-grained dark rocks composed of approximately 40% clinopyroxene, 50% plagioclase, and up to 10% olivine and a few percent each of magnetite and hornblende. Leucogabbros are somewhat finer-grained and characterized by abundant plagioclase crystals. A typical mode consists of 60–70% plagioclase, with typical sizes ranging between 1.5 and 2.5 mm, 25–30% clinopyroxene (0.2–0.5 mm), 2–3% hornblende (~ 0.2 mm), and up to 5% magnetite (~ 0.2 mm). Some samples also contain small amounts of olivine. Clinopyroxenites always exhibit a granular texture, containing ~ 70% anhedral clinopyroxene and ~ 5% plagioclase (Fig. 4c), and considerable amounts of Fe–Ti oxide (~10%). Disseminated ores are generally coarse-grained and consist of ~ 50% Fe–Ti oxide, ~ 20% clinopyroxene, ~ 20% plagioclase, ~ 10% ilmenite, and minor amounts of olivine. These euhedral silicate minerals, such as olivine, clinopyroxene, and plagioclase are enclosed by late-crystallizing Fe–Ti oxides (Fig. 4d). Massive ores typically contain > 80% Fe–Ti

We selected representative gabbroic samples from a surface section in the open pit from the three stratigraphic cases of the intrusion (Fig. 3). After screening under the microscope, relatively fresh samples were selected and sawed into slabs and the central parts were used for preparing thin sections and whole-rock chemical analyses. Polished thin sections were made from individual samples and the least altered representative samples were selected for electron microprobe analysis using a JXA-8200 electron microprobe at Washington University in St. Louis, MO. The “Probe for Windows” was used for data reduction (see http://www.probesoftware.com/). Electron microprobe analyses were performed on olivine, clinopyroxene and plagioclase. The measured data were corrected using CITZAF after Armstrong (1995). Oxide and silicate standards were used for calibration (e.g., Amelia albite for Na, Si; microcline for K; Gates wollastonite for Ca; Alaska Anorthite for Al; synthetic fayalite for Fe; synthetic forsterite for Mg; synthetic TiO2 for Ti; synthetic Mn-olivine for Mn; synthetic Cr2O3 for Cr). Samples were crushed in a steel mortar and ground in a steel mill to provide powders of ~ 200 mesh. Major elements were acquired through the analysis of fused glass discs using a scanning wavelength dispersive X-ray fluorescence (XRF) spectrometer at the Center of Modern Analysis, Nanjing University. The analytical uncertainties are less than 1%, estimated from repeated analyses of two standards (andesite GSR-2 and basalt GSR-3). Loss on ignition was determined gravimetrically after heating the samples to 980 °C for 30 minutes. Trace elements were determined by solution ICP-MS performed at the ICP-MS Laboratory at the National Research Center for Geoanalysis, Beijing. After complete dissolution, ~ 40 mg of powder was dissolved in distilled HF + HClO4 in 15 ml Savillex Teflon screw-cap breakers. Precision for most elements was typically better than 5% RSD (relative standard deviation), and the measured values for Zr, Hf, Nb and Ta were within 10% of the certified values. The detailed sample preparations, instrument operating conditions and calibration procedures follow those established by Qi and Grégoire (2000). Two standards (granite GSR-1, basalt GSR-3) were used to monitor the analytical quality. Rb–Sr and Sm–Nd isotopic compositions were obtained using a VG354 multi-collector mass spectrometer at the Center of Modern Analysis, Nanjing University. The weighed ~ 100 mg samples were spiked before dissolution with mixed isotopic tracers. They were dissolved overnight using HF and HNO3, evaporated to dryness, then followed by oven dissolution in fresh HF and HNO3 for 7 days at 160 °C. Separation of Rb and Sr was carried out with a cationexchange column (packed with AG50Wx8). Sm and Nd were further purified using a second cation-exchange column (packed with AG50Wx12). Total procedural blanks were ~ 200 pg for Sr and ~ 50 pg for Nd. The mass fractionation corrections for Sr and Nd isotopic ratios were based on 88Sr/ 86Sr of 8.37521 and 146Nd/ 144Nd of 0.7219, respectively. During our analyses, measured 87Sr/ 86Sr ratios for standard NBS987 are 0.710242 ± 0.000012, and measured 143 Nd/ 144Nd ratios for the JMC standard are 0.511124 ± 0.000010 (2σ uncertainty for 18 analyses).

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Fig. 3. Stratigraphic section showing the vertical variations in abundances of silicate, oxide and phosphate minerals and mineral compositions and sample localities (modified from Zhang et al., 1988).

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Fig. 4. a) Euhedral crystals of olivine and plagioclase, enclosed by large crystal of clinopyroxene in troctolite, under crossed polars. b) Euhedral crystals of clinopyroxene, surrounded by interstitial plagioclase and clinopyroxene, under crossed polars. c) Granular texture in plagioclase-bearing clinopyroxenite. d) Euhedral crystals of clinopyroxene and olivine, enclosed by opaque Fe–Ti oxide minerals; ilmenite exsolution lamellae is observed in clinopyroxene, under crossed polars. e) Magnetite (black) enclosing clinopyroxene and plagioclase, both of which are rimmed by brown amphibole, suggesting reaction between the hydrous oxide melt and the silicate minerals, under plain-polarized light. f) Oxide ore containing isolated grains of clinopyroxene and plagioclase in an oxide matrix as interstitial phases, under crossed polars. Note: Pl, plagioclase; Cpx, clinopyroxene; Am, amphibole.

4. Results 4.1. Mineral chemistry Representative analyses of olivine from the LZ and MZ of Taihe intrusion have Fo contents ranging from 71 to 75 mol% (Table 1). The maximum Fo content, 75, is lower than that in Panzhihua intrusion (82.3; Pang et al., 2009). No significant zoning (i.e. ~ 1 mol% in Fo content) is observed in individual olivine crystals. The analyzed olivine in the LZ contains moderate Ni content (180–530 ppm). Except one olivine crystal that contains extremely low content of Ni (below detect limit), the rest of the grains have b 460 ppm Ni. However, these differences are insignificant as compared to the intra-sample variations. The analyzed clinopyroxene in LZ and MZ is augite and diopside

(En40–43Fs5–14Wo40–43), with TiO2 contents ranges from 0.11 to 2.05 wt.% and Al2O3 from 2.94 to 5.56 wt.%, similar with those reported from the Panzhihua intrusion (En40–44Fs10–12Wo43–47 Pang et al., 2009). The analyzed clinopyroxene in UC is augite but contains higher content of ferrosilite (En41.2Fs19.0Wo27.5), TiO2 (5.63 wt.%) and Al2O3 (13.24 wt.%) (Table 2). Representative plagioclase from the Taihe intrusion has An contents ranging from 60 to 76 mol% (Table 3), overlapping with those of the Panzhihua intrusion (Pang et al., 2009). 4.2. Major and trace elements All of the analyzed gabbroic samples are relatively fresh as observed under the microscope and as indicated by their small loss-on-ignition

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Table 1 Representative composition of olivine in LZ and MZ of Taihe intrusion. Sample

th03-ol1

Rock

Olivine gabbro

Unit

LZ

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO NiO CaO Na2O K2O Total

38.69 0.02 0.00 0.03 22.55 0.38 37.88 0.03 0.14 0.19 0.06 99.97

th03-ol2

th03-ol3

th03-ol4

th03-ol5

th03-ol6

th04-ol1

th04-ol2

th04-ol3

th04-ol4

th04-ol5

th04-ol6

37.91 0.00 0.00 0.01 24.99 0.45 35.97 0.04 0.09 0.03 0.03 99.51

37.78 0.01 0.01 0.06 24.69 0.44 35.44 0.02 0.17 0.19 0.05 98.85

39.18 0.03 0.00 0.04 24.37 0.44 36.55 0.05 0.09 0.01 0.02 100.75

40.55 0.02 0.00 0.06 24.21 0.37 37.37 0.00 0.16 0.18 0.04 102.97

1.006 0 0 0 0.554 0.01 1.423 0.001 0.002 2.997 71.59 27.90

1.011 0 0 0.001 0.552 0.01 1.413 0.001 0.005 2.993 71.54 27.95

Olivine gabbro MZ

38.83 0.05 0.00 0.01 22.28 0.37 37.98 0.03 0.05 0.05 0.04 99.69

Apfu (O = 4) Si 1.011 Ti 0 Al 0 Cr 0.001 Fe2+ 0.493 Mn 0.008 Mg 1.475 Ni 0.001 Ca 0.004 Total 2.992 Fo 74.65 Fa 24.93

1.013 0.001 0 0 0.486 0.008 1.477 0.001 0.002 2.987 74.93 24.65

39.33 0.00 0.03 0.04 24.03 0.34 36.80 0.03 0.14 0.09 0.02 100.83

1.021 0 0.001 0.001 0.522 0.007 1.425 0.001 0.004 2.982 72.92 26.70

39.37 0.00 0.00 0.04 24.20 0.37 36.87 0.02 0.09 0.05 0.00 101.00

38.94 0.02 0.00 0.06 24.66 0.43 37.20 0.05 0.10 0.14 0.04 101.63

1.02 0 0 0.001 0.524 0.008 1.425 0 0.002 2.981 72.79 26.80

35.65 0.05 0.00 0.06 23.02 0.38 36.37 0.03 0.16 0.03 0.01 95.77

1.008 0 0 0.001 0.534 0.01 1.436 0.001 0.003 2.993 72.55 26.97

0.982 0.001 0 0.001 0.53 0.009 1.493 0.001 0.005 3.021 73.48 26.08

38.60 0.02 0.00 0.00 24.59 0.41 36.53 0.05 0.03 0.02 0.00 100.25

1.012 0 0 0 0.539 0.009 1.427 0.001 0.001 2.989 72.26 27.28

38.89 0.04 0.00 0.00 25.49 0.46 36.27 b.d.l. 0.05 0.02 0.01 101.23

1.013 0.001 0 0 0.555 0.01 1.408 0 0.001 2.988 71.36 28.13

1.02 0.001 0 0.001 0.53 0.01 1.418 0.001 0.002 2.982 72.43 27.09

1.03 0 0 0.001 0.514 0.008 1.415 0 0.004 2.973 73.04 26.54

Table 2 Representative composition of clinopyroxene from Taihe intrusion. Sample

th03-03

th04-01

Rock

Olivine gabbro

Olivine gabbro

Unit

LZ

MZ

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O NiO Total Mg#

52.71 1.15 3.58 0.02 6.57 0.16 14.58 22.63 0.56 0.01 0.00 101.96 80

49.45 1.96 5.10 0.13 6.70 0.19 13.76 22.41 0.47 0.02 0.01 100.20 79

Apfu (O = 6) Si Ti Al Cr Fe2+ Mn Mg Ca Na K Total Fe2+ Fe3+ Fe2+/(Fe2+ + Fe3+) Fe3+/(Fe2+ + Fe3+) Wo En Fs Ac

1.909 0.031 0.153 0 0.199 0.005 0.787 0.878 0.039 0 4.003 0.191 0.008 0.961 0.039 46.0 41.3 10.7 2.1

1.834 0.055 0.223 0.004 0.208 0.006 0.761 0.89 0.034 0.001 4.015 0.162 0.045 0.784 0.216 46.9 40.1 11.2 1.8

th04-02

th04-03

th04-04

th04-05

th05-01

th05-02

th05-03

th05-04

th05-05

Olivine gabbro

th08-04 Oxide-bearing gabbro UZ

49.59 1.93 5.48 0.01 6.12 0.19 14.05 22.76 0.42 0.01 0.04 100.61 81

1.828 0.054 0.238 0 0.189 0.006 0.772 0.899 0.03 0 4.015 0.143 0.045 0.762 0.238 47.4 40.7 10.2 1.6

50.44 1.75 5.02 0.17 7.04 0.20 13.78 21.98 0.50 0.00 0.00 100.88 78

1.854 0.048 0.218 0.005 0.217 0.006 0.755 0.866 0.036 0 4.004 0.204 0.012 0.943 0.057 46.1 40.2 11.8 1.9

50.43 1.75 4.98 0.02 7.12 0.18 13.94 22.07 0.51 0.00 0.01 101.00 78

1.852 0.048 0.215 0.001 0.219 0.006 0.763 0.868 0.037 0 4.01 0.189 0.029 0.869 0.131 45.9 40.4 11.8 1.9

50.77 1.33 3.98 0.15 7.15 0.21 14.11 22.23 0.45 0.02 0.04 100.44 78

1.878 0.037 0.174 0.004 0.221 0.007 0.778 0.881 0.033 0.001 4.013 0.182 0.038 0.827 0.173 45.9 40.5 11.8 1.7

49.66 2.05 5.56 0.16 7.15 0.14 13.63 21.96 0.51 0.00 0.01 100.81 77

1.83 0.057 0.241 0.005 0.22 0.004 0.748 0.867 0.037 0 4.009 0.193 0.027 0.879 0.121 46.2 39.9 11.9 2.0

49.51 1.73 4.89 0.21 6.26 0.17 13.85 22.58 0.45 0.01 0.00 99.67 80

48.74 1.87 4.75 0.04 7.43 0.15 13.69 21.97 0.53 0.01 0.00 99.17 77

1.843 0.048 0.215 0.006 0.195 0.005 0.769 0.901 0.032 0.001 4.015 0.151 0.044 0.776 0.224 47.4 40.4 10.5 1.7

1.833 0.053 0.21 0.001 0.234 0.005 0.768 0.885 0.039 0 4.028 0.148 0.083 0.64 0.36 45.9 39.8 12.3 2.0

49.90 1.79 4.83 0.13 6.88 0.19 14.09 22.14 0.56 0.01 0.00 100.51 79

1.844 0.05 0.211 0.004 0.213 0.006 0.776 0.877 0.04 0 4.019 0.154 0.058 0.726 0.274 45.9 40.6 11.4 2.1

51.99 1.11 2.94 0.16 7.28 0.18 15.16 21.29 0.48 0.00 0.00 100.58 79

1.912 0.031 0.127 0.005 0.224 0.006 0.831 0.839 0.034 0 4.009 0.198 0.026 0.885 0.115 43.4 43.0 11.8 1.8

41.36 5.63 13.24 0.02 10.48 0.18 12.66 11.77 2.90 0.78 0.02 99.04 69

1.567 0.161 0.591 0.001 0.332 0.006 0.715 0.478 0.213 0.038 4.102 0.027 0.297 0.083 0.917 27.5 41.2 19.0 12.3

116

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Table 3 Representative composition of plagioclase from Taihe intrusion. Sample

th03-pl01

Rock

Olivine gabbro

Unit

LZ

SiO2 Al2O3 FeO CaO Na2O K2O TiO2 Total

52.45 31.66 0.16 13.76 3.82 0.09 0.08 102.01

Apfu (O = 8) Si 9.35 Al 6.652 Fe2+ 0.024 Ca 2.628 Na 1.32 K 0.02 Total 19.994 An 66.23 Ab 33.26 Or 0.50

th03-pl02

th03-pl03

th03-pl04

th03-pl05

th03-pl06

th04-pl03

th04-pl04

th04-pl05

Olivine gabbro MZ 51.01 31.88 0.16 13.83 3.64 0.06 0.12 100.70

9.227 6.797 0.024 2.681 1.278 0.015 20.021 67.47 32.16 0.37

53.91 30.65 0.34 12.49 4.57 0.07 0.13 102.16

9.576 6.417 0.05 2.376 1.574 0.017 20.011 59.90 39.69 0.42

52.47 31.14 0.23 12.85 4.23 0.09 0.08 101.09

9.427 6.594 0.034 2.474 1.475 0.02 20.023 62.33 37.17 0.50

49.95 30.01 0.29 15.41 3.23 0.60 0.04 99.52

9.228 6.535 0.045 3.05 1.155 0.141 20.154 70.17 26.58 3.25

(LOI) values (Tables 4 and 5). As expected from their range of modal mineralogy, the rocks exhibit significant compositional variations. For example, SiO2 contents range from 40 to 45 wt.% and Al2O3 contents from 12 to 26 wt.% (Table 4). Na2O and K2O range from 1.1 to 2.6 wt.% and from 0.09 to 0.99 wt.%, respectively, whereas CaO shows a relatively narrow range between 13 and 15 wt.%. The gabbros are relatively rich in TiO2 and Fe2O3, ranging between 0.75 and 3.71 wt.% and 3.4 and 14 wt.%, respectively. There is no systematic variation of major oxides through the intrusion, although there is a slight overall decrease in MgO and Fe2O3 (as total iron) and an increase in SiO2, Na2O, Al2O3, and CaO upward (Fig. 5). TiO2 and vanadium contents are highest in the upper part of the MZ and decrease upward through the intrusion. In the gabbros, CaO, Al2O3 and total alkalis (Na2O+K2O) increase with increasing SiO2, whereas P2O5, MgO, Fe2O3, and TiO2 decrease (Fig. 6). V and Fe2O3 clearly show positive correlation with the TiO2 contents (Fig. 7). On an AFM diagram, the samples show a tholeiitic Fe-enrichment trend (Fig. 8). They are richer in Fe and Ti, but poorer in MgO than normal gabbroic rocks on an MgO-(Al2O3 +CaO)–(FeOtotal +TiO2) ternary diagram (Fig. 9). Whole-rock trace element data of gabbros and ores are listed in Tables 4 and 5. All samples, including the ores from the Taihe intrusion are enriched in light rare earth elements (LREE) relative to heavy rare earth elements (HREE) with pronounced positive Eu anomalies. Chondrite-normalized REE patterns of the samples are overall similar to ocean island basalts (OIB) (Fig. 10). Generally, all of the rocks in the Taihe intrusion have significant negative Zr and Hf, and moderate negative Nb and Ta anomalies (Fig. 10), and variable Ti concentrations corresponding to the variable content of Fe–Ti oxides in the samples. The Cr and Ni contents of the gabbros correlate positively with the MgO contents (Fig. 7c and d) while the Cr shows considerable variation with FeO (not shown), implying that Cr-spinel and olivine, respectively, are the dominant host phases, although the fractionation of clinopyroxene may also influence the chromium contents of the residual melt (Kornprobst et al., 1981). The oxide ores exhibit positive Nb and Ta anomalies (Fig. 10c and d). In contrast, U, Th, Zr, and Hf are significantly depleted relative to elements with similar compatibilities. 4.3. Sr and Nd isotopes The Taihe gabbros define a small range of age-corrected εNd(t) (t = 260 Ma) from − 0.6 to 0.7 and ( 87Sr/ 86Sr)t values ranging from 0.7040 to 0.7050 (Table 6). These data overlap with the range defined

53.42 31.00 0.09 12.68 4.41 0.10 0.08 101.78

9.515 6.508 0.013 2.42 1.524 0.023 20.004 61.00 38.42 0.58

49.80 32.41 0.45 14.54 3.18 0.04 0.09 100.51

9.057 6.946 0.068 2.833 1.122 0.009 20.036 71.46 28.30 0.24

th08-pl01

th08-pl02

Ore-bearing gabbro UZ 50.35 33.01 0.14 15.00 3.04 0.06 0.11 101.70

9.041 6.984 0.022 2.884 1.059 0.013 20.003 72.91 26.77 0.32

51.80 32.15 0.17 13.94 3.58 0.08 0.05 101.77

9.259 6.772 0.025 2.669 1.241 0.018 19.985 67.94 31.60 0.46

46.66 29.92 0.61 18.12 3.18 0.04 0.07 98.60

47.49 31.02 2.86 13.03 2.98 0.11 0.06 97.55

8.822 6.668 0.096 3.67 1.166 0.01 20.431 75.75 24.06 0.20

8.986 6.917 0.452 2.641 1.094 0.026 20.116 70.21 29.10 0.70

by the Panzhihua intrusion (e.g. Zhang et al., 2009) and the Lijiang picrite (Zhang et al., 2006), and show similar ( 87Sr/ 86Sr)t, but lower εNd(t) values as compared to those of the Baima Fe–Ti oxide ore-bearing intrusion (Pang et al., 2010). Compared with the Xinjie and Hongge intrusions, the Taihe gabbros exhibit lower ( 87Sr/ 86Sr)t values (Fig. 11a). 5. Discussion 5.1. Magma recharge and continental crustal contamination Many classic layered mafic intrusions, such as those of Skaergaard and Kiglapait, have long been thought to represent products of closed magmatic system that involved differentiation of a single injection of magma (Morse, 1980; Wager and Brown, 1968). In other words, any minor replenishment of magma in these intrusions did not significantly change the course of crystallization in the main magma body. However, cyclic layered units have been attributed to repeated influxes of new magma into the chamber, such as the Muskox, Stillwater, Great Dyke, Bushveld, Rum, and Jimberlana (e.g. Naslund and McBirney, 1996, and references therein). On the basis of a systematic study of mineral compositions carried out on stratigraphic sections through the Panzhihua intrusion, Pang et al. (2009) proposed that the intrusion was recharged at least twice by more primitive magmas. However, as indicated in Fig. 3, there is no significant reversal of compositional trends in the Taihe intrusion as marked by gradual or abrupt increases in the An content of plagioclase through the intrusion. Thus, there is no evidence for the recharge and mixing of more primitive (i.e. Fe–Ti-undepleted) magmas in the exposed sequence. However, we do not preclude the possibility that the intrusion may have been recharged at a deeper level since the floor of the intrusion has not yet been exposed. Crustal contamination has been proposed by many researchers as a crucial factor in the genesis of stratiform chromitites, such as those of the Bushveld and Stillwater complexes (Campbell, 1977; Irvine, 1977; Kinnaird et al., 2002). Ganino et al. (2008) concluded that the Fe–Ti oxides had crystallized at an early stage of the solidification of Panzhihua intrusion, because of oxidizing conditions in response to an interaction of the relatively evolved basaltic magma with sedimentary wall rocks where the Panzhihua intrusion was emplaced. At its present level of exposure, the wall rocks of Taihe intrusion include late Permian alkaline granites and syenites, whereas the Sinian meta-sandstones, and

Table 4 Major (wt.%) and trace elements (ppm) compositions of Taihe samples. Sample th01

th02

th03

Rock

Gabbro

Olivine gabbro Olivine gabbro Olivine gabbro Ore-bearing Ore-bearing Gabbro

43.51 44.08 42.59 2.02 2.16 2.79 20.81 19.89 12.63 8.07 8.97 12.06 0.02 0.03 0.05 0.12 0.13 0.16 5.07 5.98 10.25 14.44 14.87 15.72 0.99 0.09 0.19 2.03 1.97 1.23 0.25 0.26 0.06 2.67 1.53 2.42 100 99.96 100.15 56 57 63 7.9 9.97 5.78 7.17 9.14 10.3 16.3 21.3 20.6 2.4 2.89 2.58 12.1 14.9 11.4 3.26 3.52 2.01 1.71 1.89 1.75 3.12 3.6 1.49 0.4 0.46 0.25 1.86 2.56 1.28 0.33 0.41 0.19 0.81 1.07 0.57 0.1 0.1 0.05 0.48 0.63 0.37 0.06 0.08 0.06 45.8 31.1 90.1 1763 1688 2129 0.28 0.2 0.27 202 159 413 0.32 0.33 0.55 0.21 0.4 0.17 4.59 2.31 2.25 17.3 23.5 14.5 0.63 0.8 0.45 2.87 3.4 2.92 0.26 0.27 0.24 113 123 2.16 29.5 35.5 9.25 74.2 90.4 15.1 19.8 23.5 3.21 10840 13140 2975 211 243 46.5 54.4 26.7 2806 17.4 20.3 16.2

43.54 2.29 12.25 10.98 0.07 0.16 11.23 15.62 0.18 1.2 0.04 2.43 99.99 67 9.71 2.79 8.97 1.67 10.7 3.26 1.6 3.52 0.47 2.34 0.38 1.02 0.1 0.61 0.07 6.5 961 0.25 48.3 0.05 0.05 2.12 16.8 0.78 1.45 0.17 371 56.5 176 37.4 15000 320 74.8 15.1

th05

42.62 2.38 12.73 11.78 0.06 0.15 10.41 15.87 0.17 1.15 0.05 2.39 99.76 64 10 2.66 9.22 1.73 10.6 3.15 1.67 3.26 0.48 2.35 0.39 0.93 0.11 0.6 0.08 7.41 1097 0.25 51.7 0.05 0.05 2.18 16.3 0.75 1.02 0.11 407 56.6 179 41.8 15290 352 66 16.4

th052

th053

Gabbro

Gabbro

33.45 7.27 19.51 22.97 0 0.14 2.59 8.65 0.55 2.69 0.11 1.66 99.59 18 4.77 4.74 10.9 1.5 7.23 1.52 1.03 1.35 0.2 1.01 0.18 0.46 0.05 0.24 0.05 17.8 1821 0.8 201 0.25 0.08 2.09 23 0.74 6.14 0.64 4.3 99.8 61.5 10.6 44070 793 204 26.2

35.15 9.35 15.82 13.18 0.08 0.15 3.64 14.46 0.09 3.7 0.12 4.35 100.09 36 5.4 6.6 14.9 1.97 9.81 1.99 1.25 1.85 0.27 1.24 0.19 0.58 0.05 0.28 0.05 22.7 1835 1.66 278 0.26 0.08 1.19 24.7 0.85 7.24 0.79 9.63 91.8 61.9 13.6 47080 604 187 22.5

th06

th061

th07

th08

th09

Ore-bearing Ore-bearing Ore-bearing Gabbro Gabbro

44.24 27.84 1.72 9.96 15.98 17.96 9 25.13 0.06 0 0.12 0.15 8.78 3.73 15.43 10.68 0.59 0.3 1.4 1.02 0.04 0.2 2.4 2.8 99.76 99.77 66 23 21 4.61 13.8 4.77 34.2 11 4.42 1.56 20.5 7.69 4.87 1.63 1.96 1 4.72 1.5 0.76 0.21 4.28 1.03 0.84 0.16 2.47 0.39 0.31 0.05 2.24 0.29 0.39 0.05 26.9 9.72 1253 1773 0.29 0.62 136 176 4.13 0.22 0.9 0.05 1.95 1.44 16.5 23.1 0.77 0.67 8.49 6.53 0.65 0.68 482 4.08 43.3 103 141 78.1 36.4 12.5 9913 53440 245 803 115 192 18.6 23.1

Gabbro

Gabbro

39.88 3.71 17.93 14.23 0.01 0.13 5.24 13.63 0.6 1.91 0.49 2.55 100.31 42 11.9 10.4 25.9 3.83 19.8 4.26 2.34 4.57 0.62 3.11 0.48 1.24 0.13 0.74 0.11 28 1690 0.31 194 0.49 0.22 2.99 28.7 0.95 5.02 0.46 9.66 50.5 65.7 22.6 20910 446 127 22

40.15 6.35 18.07 14.09 0.01 0.16 3.61 13.54 0.59 1.01 0.25 2.52 100.35 34 5.74 6.82 16.3 2.25 10.3 2.29 1.18 2.07 0.27 1.34 0.22 0.6 0.07 0.37 0.05 21.8 1354 0.67 195 0.37 0.09 1.32 18.2 0.65 4.11 0.42 32.4 128 132 11.5 40870 1123 128 28.5

th11

th12

Olivine gabbro Gabbro

40.01 40.26 3.64 1.96 18.12 14.71 13.95 12.48 0.3 0.07 0.14 0.15 5.31 11.19 13.46 13.32 0.64 0.1 1.82 1.1 0.46 0.29 2.53 4.32 100.38 99.95 43 64 14.6 9.68 9.8 4.02 26.5 11.8 4.17 2 21.8 11.6 5.64 3.28 2.28 1.69 5.43 3.52 0.79 0.49 3.61 2.42 0.59 0.42 1.5 1.03 0.15 0.11 0.81 0.66 0.12 0.1 12 4.11 1005 1212 0.44 0.27 218 28.6 0.39 0.09 0.09 0.09 3.21 2.88 34.3 10.9 1.27 0.61 6.45 1.71 0.58 0.14 3.34 420 89.4 71.3 130 253 36.3 28.1 35950 11240 807 220 274 170 25.1 16.2

th13

th14

th15

Gabbro

Gabbro

Gabbro

39.88 42.75 45.29 44.86 2.03 1.81 0.75 0.77 14.87 12.56 25.7 25.81 22.91 12.03 3.44 3.59 0 0.05 0 0 0.16 0.13 0.04 0.03 3.78 9.87 1.41 1.68 12.24 15.69 15.46 15.63 0.31 0.18 0.93 0.91 1.08 2.45 2.54 2.62 0.21 0.05 0.18 0.19 2.77 2.41 3.91 3.88 100.24 99.98 99.65 99.97 25 62 45 48 8.63 5.78 8.97 17.5 4.03 10.3 11.1 17.9 11.2 20.6 25 44.5 1.88 2.58 3.04 6.07 10.8 11.4 14.3 29.4 2.87 2.01 2.98 6.2 1.67 1.75 1.86 2.63 3.22 1.49 2.45 5.61 0.42 0.25 0.38 0.8 2.09 1.28 1.88 4.03 0.33 0.19 0.33 0.68 0.86 0.57 0.93 1.74 0.1 0.05 0.11 0.2 0.47 0.37 0.59 1.35 0.07 0.06 0.09 0.17 45.7 90.1 47.1 42 1378 2129 2046 1941 0.27 0.27 0.24 0.76 217 413 379 270 0.22 0.55 1.34 1.35 0.05 0.17 0.36 0.47 2.26 2.25 2.13 3.2 15.9 14.5 35.1 39.9 0.77 0.45 1.04 1.2 1.57 2.92 4.26 14.6 0.16 0.24 0.41 1.21 235 2.16 7.58 5.94 40.8 9.25 14.7 36.5 115 15.1 23.9 42.1 29.9 3.21 5.11 7.42 11150 2975 4216 15960 254 46.5 77.7 375 49.5 2806 10.5 112 17.5 16.2 19.4 24.1

T. Hou et al. / Ore Geology Reviews 49 (2012) 109–127

SiO2 TiO2 Al2O3 TFe2O3 Cr2O3 MnO MgO CaO K2O Na2O P2O5 LOI Total Mg# Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Rb Sr Cs Ba Th U Pb Zr Hf Nb Ta Cr Co Ni Sc Ti V Cu Ga

Gabbro

th04

LOI = Loss on ignition; Mg# = Mg/(Mg + Fe2+) with Fe3+/Fe2+ = 0.10. 117

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T. Hou et al. / Ore Geology Reviews 49 (2012) 109–127

Table 5 Selected major (wt.%) and trace elements (ppm) compositions of Taihe Fe–Ti ores. Sample

THO04

THO06

THO08

THO10

THO12

THO14

THO16

THO18

THO21

SiO2 TiO2 Al2O3 Fe2O3 Cr2O3 MnO MgO CaO K2O Na2O P2O5 LOI Total V Cr Co Ni

0.94 15.75 4.58 69.82 0.03 0.34 5.2 0.68 0.03 0.25 0 2.1 99.72 16182.30 455.70 493.90 476.10

0.96 15.58 4.63 70.06 0.02 0.33 5.24 0.67 0.04 0.26 0 2.15 99.94 15326.80 463.20 501.70 489.60

4.52 14.9 5.45 67.53 0.03 0.34 5.63 0.33 0.03 0.25 0 0.3 99.31 14780.80 325.70 453.10 500.60

4.61 15.01 5.58 67.69 0.02 0.35 5.76 0.32 0.04 0.27 0 0.29 99.94 14967.80 336.70 463.10 479.20

9.18 22.78 7.84 51.66 0 0.3 4.68 2.5 0.07 0.11 0.02 1.25 100.39 7434.90 39.00 299.70 187.20

9.23 22.89 7.79 50.95 0 0.31 4.72 2.48 0.08 0.12 0.03 1.24 99.84 15782.90 40.50 301.20 266.40

1.54 15.92 5.03 69.23 0.03 0.36 3.59 2.09 0.03 0.28 0 2.17 100.27 15433.90 378.70 461.50 255.40

1.59 15.01 5.46 68.18 0.03 0.37 3.65 2.52 0.97 0.29 0 2.26 100.33 15324.80 368.40 456.20 278.60

1.66 15.24 5.3 69.62 0.02 0.38 3.6 2.44 0.02 0.09 0 2.42 100.79 14802.30 302.30 519.80 334.00

Triassic shales and coal measures shown in the geological map (Fig. 2), are not the wall rocks where the intrusion was emplaced. Thus, the original wall rock(s) into which the Taihe intrusion was emplaced is not known. Notably, all the Taihe samples have relatively constant wholerock Sr–Nd isotopic compositions and do not show significant crustal contamination (Fig. 11a). Using both the Yangtze lower and upper crust as the basement through which the magmas passed, modeling by element and Nd isotopic compositions shows that the samples with low εNd(t) values may have undergone very low degrees (up to 5%) of bulk-crustal contamination (Fig. 11b). Therefore, a substantial portion of the geochemical signatures such as the variable depletion of Nb, Ta, Zr and Hf contents, and the enriched Sr–Nd isotopic signature of the Taihe intrusion, probably reflect the influence of the magma source region at depth rather than processes in the magma chamber at the site of emplacement. 5.2. Plume-derived magma The data from the Taihe intrusion show minor crustal contamination and display OIB-like trace element patterns with enrichment of LILE, HFSE, and LREE, similar to the features displayed by the Emeishan basalts and Lijiang picrites (Fig. 10). Their low (87Sr/86Sr)t ratios (0.7040 to 0.7050) and higher εNd(t) values (−0.6 to +0.7) indicate a mantle source that was enriched to some degree, relative to the bulk Earth

(Sun and McDonough, 1989). This type of mantle source is similar to that of many primary oceanic and continental alkaline suites that possess positive εNd(t) in association with LREE enrichment, and provide evidence for an OIB-like, asthenospheric mantle source. It is widely recognized that the other Fe–Ti oxide ore-bearing intrusions in the ELIP with the OIB-like signatures appear to have been derived from a slightly enriched asthenospheric melt at ~260 Ma, presumably generated by the same mantle plume that formed the Emeishan flood basalts (Zhang et al., 2009). Previous studies showed that the Taihe intrusion crystallized at 258.0 ± 1.9 Ma (Zhong et al., 2009), coeval with the large-scale eruption of Emeishan continental flood basalts, which has been considered to be genetically related to a mantle plume (Ali et al., 2010; Hanski et al., 2010). Thus, these Fe–Ti oxide ore-bearing intrusions, including Taihe, appear to be part of the Emeishan Large Igneous Province (Pang et al., 2010). Such an inference is supported by the similarities of isotopic compositions and trace element patterns between the Taihe intrusion and the Lijiang picrite, which is one of the most primitive and first eruptive products of the Emeishan plume (Zhang et al., 2006). 5.3. Parental magma composition The Taihe intrusion displays an evolved signature, such as the relatively low Fo olivine (Fo71–75, Table 1), and a strong Fe enrichment

Fig. 5. Chemical variations of major element oxides and trace elements [SiO2, TiO2, Fe2O3 (as total iron), V, Al2O3, MgO, CaO, Na2O + K2O, and P2O5] from the bottom to the top of the sampled section of the Taihe intrusion, SW China. Note that the stratigraphic position of each sample is relative and that the four cases are not defined petrologically but based on mining criteria.

T. Hou et al. / Ore Geology Reviews 49 (2012) 109–127

119

Fig. 6. SiO2 vs TiO2, Fe2O3 (as total iron), Al2O3, CaO, P2O5, MgO, Mg# and Na2O + K2O for rocks of the Taihe intrusion, SW China.

trend compared with typical tholeiitic suites, although we realize that the compositions of the analyzed Taihe samples may not represent liquid compositions (Figs. 8 and 9). The composition of the parental magma of the Taihe intrusion is difficult to estimate, because the

intrusion does not have obvious chilled margins that can be used for this purpose, and there is clear evidence of crystal accumulation in the intrusion. The bulk composition of the intrusion, excluding the Fe–Ti orebodies, calculated by weighting the average chemical

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T. Hou et al. / Ore Geology Reviews 49 (2012) 109–127

Fig. 7. (a) TiO2 vs Fe2O3 (as total iron) and (b) V in the samples, (c) MgO vs Ni and (d) Cr in the gabbros from Taihe intrusion, SW China. Symbols are the same as in Fig. 6.

compositions of each case according to its outcrop area, shows markedly Fe and Ti rich and Si poor nature than those of normal tholeiitic magmas (Table 7). If the oxide ore bodies are included, the entire intrusion would be even richer in Fe and Ti and more depleted in Si. Furthermore, the presence of ilmenite exsolution lamellae in

clinopyroxene in the Taihe intrusion (Fig. 4d) indicates that the pyroxene crystallized from a Ti-rich melt. The enrichment of Fe and Ti in all of the rocks of the Taihe intrusion and the presence of Ti-rich clinopyroxene are indicative of Fe- and Ti-rich parental magmas that generated the Taihe intrusion.

Fig. 8. AFM diagram showing geochemical variations in the Taihe intrusion, SW China. The tholeiitic and calc-alkaline trends are after Wilson (1989). Symbols are the same as in Fig. 6.

Fig. 9. (CaO + Al2O3)–(FeOtotal + TiO2)–MgO plot of the Taihe intrusion, SW China. Mineral phases plagioclase (Pl), clinopyroxene (Cpx), orthopyroxene (Opx), amphibole (Amp) and oxide are also plotted in this diagram. The arrow indicates the Fe enrichment of rocks from Taihe in comparison with the normal gabbros, which should follow the line between plagioclase and clinopyroxene. Symbols are the same as in Fig. 6.

T. Hou et al. / Ore Geology Reviews 49 (2012) 109–127

121

Fig. 10. (a) and (c): Chondrite-normalized REE patterns of the gabbros and massive Fe–Ti oxide ores from the Taihe deposit. Primitive-mantle normalizing values and average ocean-island basalt (OIB) pattern are from Sun and McDonough (1989). b) Primitive-mantle normalized trace element diagrams of gabbros and Fe–Ti oxide ores from the Taihe deposit. Data for the ores are from Pang et al. (2010).

Various possibilities can be considered to generate such a parent magma including: (1) as a melt that was already rich in Fe and Ti when it originated from the mantle, (2) from a normal mantle-derived magma that became enriched in Fe and Ti as a result of differentiation prior to final emplacement, or (3) by a combination of both processes. Although ferrobasalts are commonly rich in Fe and Ti, enrichment of these elements plus depletion in SiO2 is rare among terrestrial basalts (Kushiro, 1979; Yoder and Tilley, 1962). Therefore, the low SiO2 in our samples cannot be readily explained by low-pressure differentiation of tholeiites as in the Skaergaard trend (McBirney and Naslund, 1990; Morse, 1990; Tegner, 1997; Wiebe, 1997). Experiments by Whitaker et al. (2007) successfully produced liquids with Fe–Ti-rich and SiO2-depleted affinity by crystallizing an olivine tholeiite at ~9.3 kb under anhydrous conditions (~0.4 wt.% H2O). However, there is no evidence to suggest that the Taihe intrusion crystallized at a pressure above 5 kb. Whereas Fe–Ti-rich liquids can form through low-pressure differentiation of tholeiites

(McBirney and Naslund, 1990; Morse, 1990; Tegner, 1997; Wiebe, 1997), enrichments in Fe and Ti combined with SiO2 depletion appear to require high pressure (Pang et al., 2008). Moreover, assuming the parental magma of Taihe intrusion is a normal tholeiite and had evolved in a deeper-level magma chamber (corresponding to the high pressure), the early-stage fractional crystallization of mafic minerals, i.e. olivine and/or Cr-spinel, which contain high Fe content (e.g., FeO content in the most Mg-rich olivine in Taihe gabbro is ~17 wt.%; Li et al., 1981), would lead to Fe-poor affinity of the residual melt. Hence, the Fe-rich parental magma of the Taihe intrusion is unlikely to be the result of early-stage crystal fractionation of olivine (±Cr-spinel) within a deeper-level magma chamber. Even in such a case, the Fe–Ti enrichment combined with SiO2 depletion in the residual melt requires fractionating a large amount of Si-rich and Fe-poor mineral, i.e. plagioclase (e.g. Morse, 1980). However, the remarkable positive Eu anomalies in the chondritic REE patterns of the samples indicate that significant plagioclase accumulated in the magma chamber rather than fractionated at depth. Therefore, the Fe–Ti enrichment combined

Table 6 Sr and Nd isotopic compositions of the Taihe samples. Sample

th01

th02

th04

th05

th06

th061

th08

th11

th12

th14

Rb Sr 87 Rb/86Sr 87 Sr/86Sr 2δ Sm Nd 147 Sm/144Nd 143 Nd/144Nd 2δ (143Nd/144Nd)i εNd(t) (87Sr/86Sr)i

41.87 1794 0.0672 0.705036 0.000008 3.523 14.61 0.1468 0.512522 0.000007 0.512272 −0.6 0.70479

36.65 1773 0.0615 0.705092 0.000009 3.776 15.54 0.1437 0.512526 0.000008 0.512281 −0.4 0.70486

20.19 1073 0.0582 0.705071 0.000009 3.014 11.36 0.1754 0.512601 0.000008 0.512302 0.0 0.70486

3.506 1012 0.0014 0.704985 0.000006 1.657 8.895 0.1112 0.512528 0.000006 0.512339 0.7 0.70498

22.92 1176 0.0607 0.705084 0.000007 2.968 10.15 0.1728 0.512608 0.000007 0.512314 0.2 0.70486

3.804 1682 0.0102 0.704882 0.00001 2.745 10.13 0.1617 0.512543 0.000007 0.512268 −0.7 0.70484

1.856 1203 0.0051 0.705029 0.000007 1.899 9.227 0.1196 0.512517 0.000005 0.512313 0.2 0.70501

1.295 1162 0.0046 0.704874 0.000009 1.813 8.938 0.1175 0.512536 0.000009 0.512336 0.6 0.70486

23.79 1675 0.0426 0.704898 0.000007 1.957 9.643 0.1272 0.512544 0.000007 0.512328 0.5 0.70474

13.88 1779 0.0234 0.705118 0.000008 1.755 9.042 0.1198 0.512531 0.000009 0.512327 0.5 0.70503

Note: Chondrite uniform reservoir (CHUR) values ((143Sm/144Nd) 0CHUR = 0.512638, (143Nd/144Nd) 0CHUR = 0.1967) are used for the calculation. λRb = 1.42 × 10−11/year (Steiger and Jäger, 1977), λSm = 6.5 × 10−12/year (Lugmair and Harti, 1978). (87Sr/86Sr)t and εNd(t) were calculated at 260 Ma.

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Fig. 11. a) Whole-rock initial εNd(t) versus initial (87Sr/86Sr)t ratios of the Taihe intrusion, calculated as t = 260 Ma. Different fields denote Sr–Nd isotopic compositions of ore-bearing intrusions and picrite in the Emeishan LIP after Pang et al. (2010) and Zhang et al. (2006). b) Plot of Sm/Nd versus εNd(t) to show minor crustal contamination, using sample TH05 (Sm/Nd = 0.19,εNd(t) = 0.7) as the original magma and the Yangtze upper crust (Sm/Nd = 2.0,εNd(t) = −10) and Yangtze lower crust (Sm/Nd = 0.2, −30) as contaminants (Jahn et al., 1999). c) Diagram of (Th/Ta)PM vs. (La/Nb)PM (after Neal et al., 2002). d) (Sm/Yb)PM vs (Th/Nb)PM of the Taihe gabbros. The field of the Emeishan continental flood basalts (ECFB) is from Xu et al. (2001), Xiao et al. (2004), Wang et al. (2007), Fan et al. (2008), Qi et al. (2008), Qi and Zhou (2008), Song et al. (2008). PM = Primitive mantle normalized to values of Sun and McDonough (1989). EMI range is taken from Zindler and Hart (1986) and Cliff et al. (1991). GLOSS (global subducting sediment) values taken from Plank and Langmuir (1998) and Chauvel et al. (2008). UC = upper crust values taken from Ruddnick and Gao (2003).

with SiO2 depletion cannot be attained solely by fractionation of mafic minerals prior to final emplacement. However, we still suggest that the formation of some plagioclase-rich cumulates in the LZ, such as the troctolite, increased the Fe and Ti content in the MZ of the intrusion to a limited extent, consistent with the chemical variation as seen in the stratigraphic section

(Fig. 3). Additionally, the fractionation of olivine and plagioclase could account for the low Ni (15–250 ppm), Cr (2.16–420 ppm) and Co (9.25–89.4 ppm) contents in our samples which have a very low Mg-number (Table 4). However, we cannot rule out the possibility that the fractional crystallization in the deeper part of the intrusion could also be responsible for the low content

Table 7 Comparison of the estimated parental magma and other high-Fe mafic magmas.

SiO2a TiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O P2O5 Mg#b (1) (2) (3) (4) (5) (6) a b

Estimated parental magma of Taihe intrusion

Lijinag picrite (1)

East Greenland lava (2)

Galapagos (3)

Pechenga (4)

Fe-rich tholeiite (5)

Ferrodiorite (6)

43.1 4.1 15.5 15.7 5.4 12.2 2.6 0.5 0.5 40

44.47 2.35 11.15 12.33 14.78 11.38 1.89 0.08 0.34 69

47.9 4.4 12.5 15.9 5.49 10 2.72 0.64 0.45 41

51.2 3.97 10.1 18.4 3.92 8.74 2.63 0.29 0.47 30

46.5 2.29 10.1 15.6 14.8 8.62 0.4 1.03 0.21 66

49.79 0.82 15.82 13.01 6.14 10.93 2.97 0.25 0.07 48

49.78 3.02 14.37 15.99 3.39 7.99 3.29 1.20 0.74 32

Picrite (sample DJ-2), Lijiang area, ELIP (Zhang et al., 2006). Tertiary ferrotholeiite from East Greenland (Larsen et al., 1989). Galapagos, microprobe analyses of glass (Fornari et al., 1983). Pechenga, average of ferropicrites in Pechenga, Finland (Hanski, 1992). Proposed parental magma to Upper Zone, Bushveld Complex (Davies and Cawthorn, 1984). Average of 17 ferrodiorites, Harp Lake massif, Labrador (Ashwal, 1993). Major and minor element oxides are in weight percent and are calculated to 100% total on a volatile-free basis. Mg# = [molar 100 × Mg/(Mg + Fe2+)], assuming 10% of total iron oxide is ferric.

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of transition metal elements (e.g. Ni, Cr, Co) since the floor of the intrusion has not been observed. Moreover, undoubtedly, the precipitation of copious amounts of Fe–Ti oxides led to the depletion of Fe and Ti in the UZ of the intrusion.

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The trend of isotopic compositions for Taihe gabbros (Fig. 11a) toward lower εNd(t) and higher (87Sr/86Sr)t could indicate the involvement of low-εNd materials in the source mantle or minor contamination of magmas with Mesoproterozoic or late Paleoproterozoic lithosphere (crust or mantle) overlying the high-εNd source region reflected in the Lijiang picrites. Since the possibility of considerable crustal contamination has been excluded, they can be interpreted in terms of interaction between plume-derived magmas and the subcontinental lithosphere (SCLM; Davies et al., 1989; Hofmann, 1997; Green and Falloon, 1998, 2005), similar to the petrogenesis of many Emeishan basalts proposed by other researchers (e.g., Xiao et al., 2004; Xu et al., 2001). Hence, we propose that the parental melt of the Taihe intrusion could have been generated by plume-lithosphere interaction or ascending plumederived melts contaminated by the lithospheric mantle, a model that is consistent with a previous study on the Panzhihua intrusion (Zhang et al., 2009). This model has been advocated to account for the petrogenesis of basalts and picrites in the ELIP by many workers (e.g. Chung and Jahn, 1995; Xu et al., 2001; Zhang et al., 2006). However, if the lithospheric mantle did not experience large-scale impregnation by fluids, it is generally recognized as refractory mantle with a high viscosity because it is cold and nearly anhydrous. Partial melting of this type of mantle hardly occurs, even when heated by a plume (Rudnick et al., 2004). Thus, the lithospheric mantle was probably metasomatized by fluids/ melts and enriched (Guo et al., 2003). This inference is consistent with the radiogenic isotopic composition of the Taihe intrusion. Unfortunately, due to the lack of peridotite xenoliths, we have no information about the composition of the SCLM (subcontinental lithospheric mantle) directly under the study area to effectively test these ideas.

source region that has incorporated subduction-related materials. Moreover, both the Taihe gabbros and Lijiang picrites could be derived from the same plume as indicated by their Sr–Nd isotopic compositions. Hence, these characteristics of the Taihe gabbros suggest that the subduction-related materials were stored in the lithospheric mantle, and added into the Emeishan plume-derived melts by plumelithosphere interaction. The Taihe intrusion data on a (Th/Nb)PM vs (Sm/Yb)PM diagram plot exclusively near the EMI area further supporting the addition of EMI-components which were possibly formed via subduction (Weaver et al., 1986), and were incorporated via plume-lithosphere interactions (Fig. 11d). Similar characteristics, such as higher (Sm/Yb)PM and Nb-Ta-Zr-Hf depletion have been recognized in other Fe–Ti oxide ore-bearing intrusions such as Panzhihua, Baima and Hongge (Pang et al., 2010), which could be attributed to subduction-related metasomatism in the source. Our previous noble gas isotope studies on olivine and clinopyroxene separated from the Panzhihua, Baima, Hongge and Taihe gabbros show that the source regions of Fe–Ti oxide deposits in the Panxi area were affected by subduction-related melts/fluids (Hou et al., 2011). Such a mantle has a substantially lower solidus temperature than the refractory SCLM (Leeman et al., 2009). Indeed, during the Neoproterozoic, the western margin of the Yangtze Block was an active margin with subduction zone, where the Panxi area is located. The subducted oceanic crust likely underwent eclogite-facies metamorphism and generated slab melts. This scenario is supported by the presence of widelydistributed adakitic plutons, such as those of Datian and Dajianshan felsic plutons in the Panxi area. Furthermore, these plutonic rocks are inferred to be the products of melting of a subducted oceanic slab, in the presence of garnet as a residue in the source, i.e., in the form of eclogite (e.g., Zhao et al., 2008; Zhou et al., 2002a). On the basis of trace element and Pb–Sr–Nd isotopic investigations, Song et al. (2004) and Xiao et al. (2004) also proposed that the primitive magmas of the ELIP were produced by partial melting of a rising mantle plume that interacted with the lithospheric mantle that has been previously modified and enriched by pelagic sediments during Neoproterozoic subduction. The subduction and metasomatic events affecting lithospheric mantle during the Neoproterozoic are also suggested by the Re-Os isotopic compositions of the mantle-derived peridotite xenoliths from the Maguan district, Yunnan province, which is not far from Panxi district (see inset of Fig. 1; Huang et al., 2011). Eclogitic material in the source has been also suggested by modeling La/Yb and Dy/Yb ratios (Fig. 12). The Taihe data lie closer to a mixing line between amphibole-bearing lherzolite and amphibole-bearing eclogite and plot within the field defined by data of the Panzhihua gabbros. Although the exact position of the mixing lines is clearly model-dependent, the inference that much of the melting occurred in garnet-facies peridotite mantle coexisting with eclogitic material is logical.

5.5. The role of subduction-related materials

5.6. Constraints on mineralization

As stated in the previous section, the Taihe gabbros show minor crustal contamination, and these signatures probably reflect the influence of the magma source region at depth rather than processes in the magma chamber at the site of emplacement. The Taihe gabbros show enrichment of the LILE (e.g. Rb, K, Th, and U) and moderate depletion of HFSE (e.g. Nb, Ta, Zr, and Hf), and the HREE (Fig. 10), which suggest that they originated from a mantle source metasomatized by the slab-derived fluid/melt (Wilson, 1989). During the subduction processes, Zr and Hf are immobile in fluid-dominated regimes (Münker et al., 2004). This implies that the lithospheric mantle beneath the Panxi region has likely been metasomatized by ancient slab-derived melts (Hou et al., 2011). Depletion of Zr and Hf by zircon fractionation is unlikely due to the low SiO2 content of the magmas involved. Hence, considering the typical features of magmatism subduction zones, the Taihe intrusion was probably derived from a

Many researchers have suggested that phosphorus plays an important role in the formation of magmatic Fe–Ti oxide ore deposits (e.g. Thompson et al., 2007). However, the absence of apatite in the Fe–Ti oxide ores (Table 5), combined with its presence in the silicate rocks in the upper part of the Taihe intrusion that is barren of Fe–Ti oxide ores, suggests that phosphorus is not a key factor that controls ore formation in the Taihe deposit. Thus, we suggest that the presence of apatite in some Fe–Ti oxide layers in the upper part maybe fortuitous; if apatite crystallized at the same time as the oxides, it simply accumulated along with them (e.g. Pang et al., 2008). The relative enrichment of apatite in the upper part of the intrusion is probably a result of crystallization of other silicate-rich and phosphorous-poor minerals in the lower part. Some experiments indicate that carbon and its oxides might play a role in the formation of Fe–Ti oxide ores (e.g., Lindsley et al., 1999;

Accordingly, we suggest that the melt that evolved into the Taihe intrusion was already rich in Fe and Ti when it originated from the mantle, and was probably ferropicritic. Relatively primitive, high-Fe mantle-derived melts are represented by ferropicrites that occur as isolated lava flows in the Precambrian or at the base of continental flood basalt provinces in the Phanerozoic (Gibson et al., 2000). Although rare, the occurrence of picrites with high concentrations of Fe and Ti has been documented from the ELIP (e.g. Chung and Jahn, 1995; Zhang et al., 2006). A ferropicritic parental magma composition has also been proposed for the Fe–Ti oxide ore-bearing Panzhihua intrusion by Zhang et al. (2009) and the Hongge intrusion by Bai et al. (2012a, b). 5.4. Nature of the mantle source

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of relatively large amounts of Fe–Ti oxide required for the development of ore-rich layers. In addition, variations in the relative proportions of coexisting oxide and silicate phases may be related to differences in concentrations and diffusion rates of ions in the liquid immediately above the crystallizing layer (Reynolds, 1985). Precipitation of abundant magnetite lowered the density of the stagnant layer until it reached that of the overlying magma. The two liquids then mixed, thus terminating the cycle of Fe–Ti oxide layer formation and marking a return to silicate-dominated fractionation. 5.7. Implications for the formation of large Fe–Ti–V ore clusters

Fig. 12. Variation of La/Yb vs. Dy/Yb for the Taihe gabbros compared with modal partial melting curves for amphibole-bearing garnet lherzolite (Amp Gt lh) and amphibolebearing eclogite (Amp Ec). D values are from http://earthref.org/cgi-bin/er.cgi? s=kdd-s0-main.cgi. Source mineralogy is 50% olivine, 20% orthopyroxene, 20% clinopyroxene, 5% garnet, 5% amphibole for amphibole-bearing garnet lherzolite; 52% garnet, 43% clinopyroxene, 5% amphibole for the inferred eclogitic material stored in the lithospheric mantle. Data for Panzhihua gabbros are from Zhang et al. (2009) and Zhou et al. (2005).

Zhou et al., 2005). Accordingly, Ganino et al. (2008) proposed that interaction of the relatively evolved basaltic magma with sedimentary wall rocks of the Panzhihua intrusion, i.e. carbonate rocks, was crucial for the formation of the Fe–Ti oxide ores in the intrusion. The footwall of the Taihe intrusion is unknown. This, combined with the geochemical characteristics which indicate only minor crustal contamination, implies that carbon is not essential for producing Fe–Ti oxide ores in the present case. Experiments at low pressures indicate that crystallization of Fe–Ti oxides from parental ferropicrite is favored under relatively oxidizing conditions (Juster et al., 1989; Toplis and Carroll, 1995). Periodic fluctuation of fO2 in the magma has been postulated as the mechanism for the formation of titanomagnetite layers in the Upper Zone of the Bushveld Complex (Klemm et al., 1985). Extensive troctolitic and gabbroic rocks that occur in the LZ of the Taihe intrusion (Fig. 3) implies a protracted period of fractional crystallization that resulted in the concentration of large amounts of Fe, Ti, and V in the residual magma, as also suggested by petrographic observations (Fig. 4). This process would have created favorable conditions for the precipitation of large quantities of Fe–Ti oxides (e.g. Reynolds, 1985). Thick Fe–Ti oxide layers are mainly observed in the lower part of the MZ. This is indicative of a dense Fe–Ti– (V)-enriched liquid that did not mix with the overlying magma and formed a stagnant layer from which large amounts of Fe–Ti oxide may have crystallized. Crystallization of Fe–Ti oxides is controlled by the Fe2O3/FeO ratio of the liquid and is a function of fO2, temperature, and the fH2O/fH2 ratio (Kress and Carmichael, 1991). Modeling of estimated parental magma composition (Table 7) using the MELTS program successfully simulated cotectic crystallization of olivine (~ Fo82; Li et al., 1981) and Fe–Ti spinel at 1,154 °C, 2 kb, the FMQ buffer, and an initial H2O content of 1.5 wt%. The modeling also showed that under these relatively oxidizing conditions, Fe–Ti oxides formed early in the crystallization sequence. The Fe2O3 content of the magma in the Taihe intrusion would also be elevated with the crystallization of ilmenite and Fe +2-bearing silicates (clinopyroxene and olivine). This is evidenced by the higher Fe 3+/(Fe 2+ + Fe 3+) ratio in ore-bearing gabbro of UZ, compared with those in the gabbro of LZ and MZ (Table 2). The H2O content of the residual liquid will also be elevated by the fractionation of anhydrous silicate phases as reflected by considerable amounts of hornblende in the rocks (Fig. 4e). Therefore, it is reasonable to infer that a complex interplay of these factors resulted in the precipitation

As discussed above, during the Neoproterozoic, subducted oceanic crust beneath the Panxi area likely underwent eclogite-facies metamorphism and partial melting. The infiltration of subduction-related melts and fluids into the lithospheric mantle led to the formation of metasomatic minerals, such as pyroxene, phlogopite, and/or amphibole. Considerable amounts of melts were produced by partial melting of eclogite-facies oceanic crust as reflected by the Neoproterozoic adakitic plutons in the Panxi area (e.g., Zhao et al., 2008). Such melts would have moved upwards through the lithospheric mantle and reacted with olivine in the surrounding peridotite to form orthopyroxene and garnet (Gibson, 2002). Re-melting of residual eclogite could also produce andesitic partial melts that react with the surrounding peridotite and enrich it in garnet and clinopyroxene (Yaxley, 2000), forming the equivalent of eclogite. Melting of this relatively anhydrous, re-fertilized peridotite produces Fe-rich melts that are enriched in incompatible trace elements (Yaxley and Green, 1998). Such a re-fertilized peridotite source has a lower solidus than normal peridotite (Yaxley, 2000). During the late Permian, when the Emeishan plume ascended from the lower mantle (Zhang et al., 2008), it is likely that it encountered the enriched lithospheric mantle containing the eclogitic material. Thus, the primary magmas may have been produced by partial melting of both the plume and the eclogitic material. The magmas were parental to the ore-bearing intrusions, carrying a relatively enriched isotope signature, derived from a hybridized mantle source which contained normal garnet-zone lherzolitic components (Fig. 12) and a mixture of re-fertilized peridotite and recycled material. As stated above, we have inferred that the Fe-rich parental magma was derived from a Fe-rich mantle source. Tuff et al. (2005) demonstrated that ferropicritic magmas can be generated by the interaction between melts of garnet pyroxenite and the surrounding peridotites under high pressure (~5 GPa) and temperature (~1550 °C). This is consistent with the conclusions drawn from previous experiments that the iron content of primary magmas increases with both increasing mean degree of partial melting and increasing mean temperature at a given pressure (Langmuir and Hanson, 1980). In the Panxi district, the Fe-rich magma could have been generated by this process, i.e., partial melting of a mixture of an upwelling mantle plume and the subcontinental lithospheric mantle that comprises the equivalent of eclogite and other recycled materials due to the ancient subduction-related enrichment. This type of melting involving an eclogite or pyroxenite component is also capable of explaining the significant volumes of melts seen in the ELIP and other LIPs (e.g., Campbell, 1998; Cordery et al., 1997; Leitch and Davies, 2001; Takahashi et al., 1998). In summary, the evidence supports the model that the isotopic and elemental composition of the Taihe gabbroic intrusion can be attributed to the involvement of subducted crustal components, and that the parental magma of the Taihe intrusion and other ore-bearing intrusions was ferropicritic derived from possible eclogitic materials of former oceanic crust. Thus, we propose a model where the interaction of these subduction-generated lithospheric mantle sources and the Emeishan plume led to the production of many world-class Fe–Ti oxide deposits in the Panxi area.

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6. Conclusions The ore bodies of the Taihe intrusion probably formed by latestage crystallization of Fe–Ti–V rich melt under a high oxygen fugacity and volatile-rich environment. The rocks of the Taihe intrusion were derived from a ferropicritic melt and are more evolved in chemistry upward and follow a tholeiitic differentiation trend with enrichment in Fe, Ti, and V. The ferropicritic parental magma was formed by the interaction of Permian Emeishan plume and the lithospheric mantle, with the latter enriched in eclogitic component related to Neoproterozoic subduction. The interaction of these subductiongenerated lithospheric mantle sources and the Emeishan plume led to the production of many world-class Fe–Ti oxide deposits in the Panxi area. Acknowledgements Financial support for this work was supported by Projects 2012CB416805 and 2007CB411405 of the State Key Fundamental Program (973), “The Fundamental Research Funds for the Central Universities”, Key Project of Chinese Ministry of Education (no. B07039), the National Natural Science Foundation of China (nos. 40925006 and 40821061), the 111 Project (B07011), and PCSIRT. References Ali, J.R., Fitton, J.G., Herzberg, C., 2010. Emeishan large igneous province (SW China) and the mantle plume up-doming hypothesis. J. Geol. Soc. London 167, 953–959. Armstrong, J.T., 1995. CITZAF-a package of correction programs for the quantitative electron microbeam X-ray analysis of thick polished materials, thin-films and particles. Microbeam Anal. 4, 177–200. Ashwal, L.D., 1993. Anorthosites. Springer-Verlag, Berlin, Heidelberg, p. 148. Bai, Z.J., Zhong, H., Li, C.S., Zhu, W.G., Xu, G.W., 2012a. Platinum-group elements in the oxide layers of the Hongge mafic–ultramafic intrusion, Emeishan Large Igneous Province. SW China. Ore Geol. Rev. 46, 149–161. Bai, Z.J., Zhong, H., Naldrett, A.J., Zhu, W.G., Xu, G.W., 2012b. Whole rock and mineral composition constraints on the genesis of the giant Hongge Fe–Ti–V oxide deposit in the Emeishan Large Igneous Province, SW China. Econ. Geol. 107, 507–524. Bateman, A.M., 1951. The formation of late magmatic oxide ores. Econ. Geol. 46, 404–426. Botcharnikov, R.E., Almeev, R.R., Koepke, J., Holtz, F., 2008. Phase relations and liquid lines of descent in hydrous ferrobasalt—implications for the Skaergaard intrusion and Columbia River Basalts. J. Petrol. 49, 1687–1727. Boven, A., Pasteels, P., Punzalan, L.E., Liu, J., Luo, X., Zhang, W., Guo, Z., Hertogen, J., 2002. 40Ar/39Ar geochronological constraints on the age and evolution of the Permo-Triassic Emeishan volcanic province, southwest China. J. Asian Earth Sci. 20, 157–175. Campbell, I.H., 1977. Study of macro-rhythmic layering and cumulate processes in Jimberlana intrusion, Western Australia. 1. Upper layered series. J. Petrol. 18, 183–215. Campbell, I.H., 1998. The Mantle's Chemical Structure: Insights from the Melting Products of Mantle Plumes. In: Jackson, I. (Ed.), Earth's Mantle: Composition, Structure and Evolution. Cambridge University Press, Cambridge, pp. 259–310. Cawthorn, R.G., 1996. Layered Intrusions. Elsevier, Amsterdam . (531 pp.). Cawthorn, R.G., Ashwal, L.D., 2009. Origin of anorthosite and magnetitite layers in the Bushveld Complex, constrained by major element compositions of plagioclase. J. Petrol. 50, 1607–1637. Cawthorn, R.G., Spies, L., 2003. Plagioclase content of cyclic units in the Bushveld Complex, South Africa. Contrib. Mineral. Petrol. 145, 47–60. Charlier, B., Duchesne, J.C., Auwera, J.V., 2006. Magma chamber processes in the Tellnes ilmenite deposit (Rogaland Anorthosite Province, SW Norway) and the formation of Fe–Ti ores in massif-type anorthosites. Chem. Geol. 234, 264–290. Charlier, B., Namur, O., Duchesne, J.C., Wiszniewska, J., Parecki, A., Auwera, J.V., 2009. Cumulate origin and polybaric crystallization of Fe–Ti oxide ores in the Suwalki anorthosite, Northeastern Poland. Econ. Geol. 104, 205–221. Chauvel, C., Lewin, E., Carpenter, M., Arndt, N.T., Marini, J.C., 2008. Role of the recycled oceanic basalt and sediment in generating the Hf–Nd mantle array. Nat. Geosci. 1, 64–67. Chung, S.L., Jahn, B.M., 1995. Plume-lithosphere interaction in generation of the Emeishan flood basalts at the Permian–Triassic boundary. Geology 23, 889–892. Cliff, R.A., Baker, P.E., Mateer, N.J., 1991. Geochemistry of inaccessible island volcanics. Chem. Geol. 92, 251–260. Cong, B.L., 1988. Formation and evolution of the Pan-Xi paleorift (in Chinese). Science Press, Beijing, p. 424 (in Chinese). Cordery, M.J., Davies, G.F., Campbell, I.H., 1997. Genesis of flood basalts from eclogitebearing mantle plumes. J. Geophys. Res. 102, 20179–20197. Davies, G., Cawthorn, R.G., 1984. Mineralogical data on amultiple intrusion in the Rustenburg Layered Suite of the Bushveld Complex. Mineral. Mag. 48, 469–480.

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